SEDIMENTARY BASINS OF THE WORLD An Introduction to the Series
Etymology reveals much about the essence of a word. Science in German, Wissenschaft, is the art of observing whereas science in Chinese, koxue, is the study of classifying. Scientific observations, with the help of modern equipments, have made leaps and bounds in our century, but taxonomy seems irrelevant. Classification can be science. The rise of the natural sciences in Europe could be traced back to Carl Linnaeus in 1750 when he used criteria of mutual exclusiveness to establish the taxonomy of living organisms. Unfortunately, this prerequisite in dividing and subdividing is not always appreciated, and a common practice in geology has been to "classify" basins through reference to incidental attributes. So, we have coastal basins, back-arc basins, extensionally rifted basins, successor basins, deep-sea basins, flysch basins, etc. These qualifications describe the geography, tectonic setting, principal stresses, orogenic chronology, depositional environment or sedimentary association of a basin, but these all could be different aspects of one and the same. "A basin is a basin is a basin", paraphrasing Gertrude Stein. Even though no two basins are exactly alike, subsidence is the common denominator of all. A systematic classification of basins depends upon recognition of the mutually exclusive causes of subsidence. Forty years ago, when I first went to study in the United States, I was fascinated by the debate whether basin subsidence is isostatically induced by sedimentary load. Later, in the 1950's, I began to realize that depressions on Earth are underlain by thin crust and came to the conclusion that subsidence could be the surface manifestation of an endogenetic process of crustal thinning, in response to the Airy isostasy. Still later, in the early 1960's, after geophysical studies had revealed the heterogeneity of the Earth's mantle, subsidence could be related to mantle cooling and corresponding density change, in response to the Pratt isostasy. Meanwhile, we all concede that the weight of a sedimentary pile filling a subsiding depression will induce isostatic subsidence. The three different mechanisms of isostatic adjustment are, however, not mutually exclusive, and they may have operated concurrently. To classify basins on the basis of the various modes of isostasy can thus not be systematic. But, as indicated by gravity studies, not all sedimentary basins are isostatically adjusted, and not all subsidence is isostatic. For a first-order division we could thus recognize two classes of basins, those which have subsided isostatically and those which have not. Crustal thinning is a prerequisite to initiate Airy isostasy. Thinning can be induced by two different systems of principal stresses, with the least principal stress either horizontal or vertical. The former leads to the genesis of rifted basins under horizontal extension, and the latter results in pull-apart basins in transform or strike-slip fault zones. The orientation of principal stresses could thus be the criterion for second-order distinctions. Rift basins are commonly present in the continental interior. After the appearance of oceanic crust between separated continents, the rifts become narrow oceanic gulfs (like the Red Sea). Eventually the loci of subsidence are shifted to passive margins where coastal plains are underlain by thick basinal sediments. Rifted basins may also form on an active margin where a segment of continent is torn apart from the mainland to form islands arcs; those are back-arc basins. The position with respect to plate margins can thus be used as the criterion for third-order distinctions. Subsidence induced by horizontal compression, the third of the three possible configurations of principal stress, is not isostatic. Plate-tectonics theory relates the origin of trenches to the underthrusting of ocean lithosphere on an active plate margin and the origin of foreland basins to the underthrusting within the continental lithosphere. These two major basins of compressional origin are thus also distinguished by the third-order criterion concerning their position with respect to plate margins.
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SEDIMENTARY BASINS OF THE WORLD m AN INTRODUCTION TO THE SERIES
When we first planned the series of the Sedimentary Basins of the World, we intended to adopt a genetic classification. Basins are subdivided into the three sets of criteria discussed above: 1. Isostasticallyadjusted basins 1.1. Extensionalbasins 1.1.1. Rift basins in the continental interior 1.1.2. Narrowoceanic gulf basins 1.1.3. Basins of deposition on passive margins 1.1.4. Rifted basins on active margins or back-arc basins 1.2. Transcurrent(transform or strike-slip) pull-apart basins 2. Isostasticallynot adjusted basins 2.1. Compressionalbasins 2.1.1. Forelandbasins (in a continental interior) 2.1.2. Oceanic trenches (on a continental margin) With this scheme in mind, the editors of Elsevier and I made up a list of the volumes for the projected series, and we started our search for volume editors. Our first priority, taking into account current demand, was to bring out a volume on China. We were concerned that the Chinese basins could not be fitted into our scheme, because we were told that they represent a special group of unclassifiable basins on "paraplatforms". However, as I became personally involved in researches on the geology of China, I came to the conclusion that the Chinese basins were not "unclassifiable". Rifted basins, back-arc basins, pull-apart basins, foreland basins, etc., exist in China, as they are present elsewhere in the world. The Chinese basins of different origins do share a common history in geologic evolution, and they are united by their geography. It would be illogical to discuss the various Chinese basins in separate volumes. Yet if we are to include all of them in one, we have to designate that volume by their unifying geography. When we started to work on our second volume on rifts, we were still trying to keep our genetic scheme, although we were resigned to make a single exception for the Chinese opus. We were thinking of the East African Rift Valleys, and the emphasis was on Africa. As chance would have it, I just happened to accept a consulting contract on Africa. After a year of working on the assignment, I realized that basins on that continent are as diverse in origin as those in China; there are foreland basins, pull-apart basins, as well as rift basins in Africa. About this time, our choice of the editor for a volume on rifted basins, Professor R.C. Selley, brought up another issue. He pointed out to me the impracticability of putting out volumes on the sole basis of their postulated origins. The purpose of the series is to provide information on the geology of sedimentary basins of the world in order to help a novice to start a project. Commonly, the one who seeks information knows the geographic extent of his interests, but not necessarily the genesis of his targets. Taking, for example, the case of a person who is to start an exploration venture in some region, how should he know if he is to study a monograph on pull-apart basins or one on foreland basins. He knows, of course, if the location is in China or in Africa, and could consult an opus on Sedimentary Basins of China, or that on Africa accordingly. The arguments by Professor Selley finally convinced me to change our scheme. The criterion of dividing the volumes will have to be geographical. In addition to those on China, Africa, and the Caribbean, volumes on sedimentary basins of Australia, South America, and the Soviet Union (Russian Platform and Siberia) are planned. Geographical groupings are satisfactory if the basins of various origins in a region share some common heritage, but to throw all heterogeneous entities into one big pot could be disconcerting. To produce, for example, a volume on the Sedimentary Basins of Europe to include all those in the Russian Platform, under the North Sea, and in the Prealps may make a good lexicon, but not an opus harmonized by a unifying theme. This consideration led us to the decision to place priority on certain natural boundaries, so that each volume would sustain a certain coherence in geology as well as in geography. The Cenozoic basins in the Tethyan orogenic belt, for example, are to be grouped under the title of Foreland Basins of the Alpine-Mediterranean Region. The pull-apart and back-arc basins on the shores of the Pacific will be included in two or more volumes of the
Sedimentary Basins of the Circum-Pacific Region. There are, of course, other sedimentary basins of the Earth, especially those of the Near East, North America and Antarctic, which should be included in the series if the coverage is to be complete. On the other hand, we shall also evaluate the demand of the profession for such volumes as the initial volumes of the series successively appear during the next decade.
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It is my hope that the series of volumes would not be compared to philately albums; there should be unity in style and in substance. Yet the accumulation of geological information has reached such immense proportions since Eduard Suess wrote his monumental work Das Antlitz der Erde that no single person could ever hope to master the geology of the world. The series of our volumes will, therefore, have to be collective efforts. Coordinations by volume editors are indispensable. My job, as the series editor, is to further enhance the unity and harmony of the whole. We have, however, to accept the fact that each article of a volume may "speak" a different dialect, and each volume of the series may "speak" a different language. Perfect consistency can only be achieved if a person has the time or the capacity to translate all those hundreds of articles in more than a dozen volumes into one universal script. This is not possible, and the practical alternative to the ideal is, therefore, to leave each author or group of co-authors a maximum freedom in their style of presentation and in their interpretations of geology. The articles are to be accepted as expressions of the present state of understanding of an area by leading geologists working in that area. They may or may not represent the understanding of the volume editor, or that of the series editor. Through my experiences in editing the first volume on Chinese basins, I appreciate the potential dangers of such freedom of expressions; a lack of precision in semantics could lead to grave misunderstandings. I felt impotent when I saw basic terms, like orogeny, platform, shelf sediments, intracratonic basins, etc., defined, in certain communities of our profession, on a basis distinctly different from that adopted by modem students of geology. The misconceptions in some instances are so deeply rooted, that nothing short of a rewrite could save the situation. Yet neither the volume nor the series editor could completely revise all the articles. To avoid complete chaos, I plan, therefore, to write a summary, at the end of each volume (or at least some), in my style, and to interpret the geology on the basis of my understanding. Such summaries may contain an overdose of personal opinions and may involve interpretative errors by a single geologist, but they should, at least, be consistent, and may eliminate misconceptions caused by the divergent meaning of the same words as they are used in various "dialects" and "languages". The preceding pages were written in January, 1988 for the first volume of the series on Chinese basins. The reviewers of the Chinese volume gave me encouragement that I, as the series editor, should continue my role as an interpreter of different cultures. I have, therefore, taken the initiative to write the last chapter of the second volume m A Distant View of the South Pacific Geology. The volume edited by Peter Ballance and my summary are both written in English. There is little need for translation. Nevertheless, New Zealanders speak English with their local accent, and the same "words" are pronounced differently by one who speaks English with a strong Chinese and Swiss accent. It is not surprising that Peter, a dear old friend, found it difficult to "agree entirely" with me. On the other hand, he was tolerant enough not to protest too strongly, and I could have my Distant View for the reference of other distant readers. In reviewing the articles in the second volume of our series, I became more convinced than ever of the wisdom of Selley's advice that the basins should be grouped regionally. I have devoted most of my professional career with a process-oriented approach. When I edited a book on Mountain-Building Processes, the emphasis was on processes. Instead of regional syntheses, I adopted an analytical approach to look into the different processes involved in mountain-building, sedimentary, magmatic, deformational, and metamorphic. In editing this series on Sedimentary Basins, there is no better alternative than regional synthesis. Not only basins in China have diverse origins, those in the South Pacific are equally diverse. Yet they all seem to belong to the same set of diverse basins. If the geology of the South Pacific seems very different from that of China, the apparent distinction can be attributed to the fact that they have advanced to different stages of tectonic evolution. Four years have gone by since the publication of the second volume of the series of The Sedimentary Basins of the World. On the eve of the publication of the third volume, I am encouraged to find that the series will not be philately albums; they will not be collections of random observations. I saw the parallelism in the pattern of orogenic deformations in China and in South Pacific, and I could see the same pattern in Africa. There is a difference in the stage of the tectonic evolution. China, as a part of Eurasia has undergone a billion years of amalgamation. The South Pacific is still in an earlier phase of accretion. Africa has gone a long away since its separation of Pangaea, and it is being pushed toward Eurasia to its ultimate destiny of a place in a supercontinent. The figures and the colorations of the mosaic pieces are different, but they are all to be pieced together for a unifying theory of global tectonics.
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S E D I M E N T A RBASINS Y OF THE WORLD - - AN INTRODUCTION TO THE SERIES
With the publication of the Caribbean volume, the Sedimentary Basins of the World series has come of age. The "philately albums" have accumulated such a wealth of information, and the series is on its way to becoming an encyclopaedia of sedimentary basins. More significant is the fact that geology could shed its label of provincialism. In addition to basins formed by ocean-continent interactions, the Caribbean volume provided a basis for interpreting geology on the basis of assuming ocean-ocean plate-interactions. The perspectives have caused me to think again about the genesis of some Chinese basins. Were not the Tarim and Qaidam basins fragments of an oceanic plateau, like the Colombia and Venezuela basins, that survived as rigid blocks during the Phanerozoic deformations? We originally hoped that the series would be completed in 10 or 15 years. With the publication of this fourth volume, we see the light at the end of the tunnel, even though another decade will have to elapse before we complete our task. The two volumes on sedimentary basins of the former Soviet Union are being edited. We are pleased to persuade Andrew Miall to edit the volume of North American basins. Other volumes being planned are South American basins, Mediterranean and Tethyan basins, Southwestern Asian basins, West Pacific basins, Australian and Antarctic basins. In order to fulfill our dream of producing an opus comparable to Suess's The Face of the Earth, two more series are required: The Mountains of the World and The Oceans of the World. Elsevier might see the wisdom to initiate these ambitious projects, even though I myself may be too ancient to see their completion. Zurich, Switzerland, January, 1999
KENNETH J. HSU Series Editor Sedimentary Basins of the World
Dedication This volume on Caribbean sedimentary basins is dedicated to Professor Willem A. van den Bold, a pioneer in geologic and paleontologic studies of the circum-Caribbean basins. Early years. 'Wim' van den Bold was born in 1917 and raised in Amsterdam. In 1942, he began his Ph.D. studies at the Geological Institute of the University of the Utrecht under the supervision of Professor L.M.R. Rutten, a well known structural geologist and tectonicist. The proposed topic of his Ph.D. study was a study of Cretaceous and Tertiary ostracodes contained within samples from Cuba that had been collected by Rutten and other geologists from the University of Utrecht as part of their ongoing study of the regional geology and tectonics of that island. Within one year of starting this study, Wim was forced into hiding in Utrecht to avoid conscription by Nazi authorities for forced labor in the German war effort. When it appeared that the Nazis had given up their search for him, Wim retrieved his microscope and literature from the university and resumed his Ph.D. study in his house. To compensate for the loss of electricity in 1944, Wim was able to illuminate his microscope by devising a set of mirrors and lenses to direct sunlight through a small opening in his attic. His persistence paid off and in 1945 Wim completed his Ph.D. study, 'Contribution to the Study of Ostracoda, with Special Emphasis to the Tertiary and Cretaceous Microfauna of the Caribbean Region'. In addition to his original Cuban samples from Rutten, his study was expanded to include sample material from Bonaire, British Honduras (now Belize) and Guatemala. His dissertation results showed that the vertical ranges of many species of ostracodes in various sedimentary basins were limited enough for stratigraphic correlations and that the ability of many ostracoda to adapt to
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changing water conditions made them ideal indicators for reconstructing marine vs. brackish-water environments. Shell years. Wim joined Royal Dutch Shell in 1946 to participate in the post-war resurgence of international oil exploration. His first assignments took him to Venezuela, Colombia and Trinidad where he contributed to a wider industry-sponsored stratigraphic correlation effort that led to the modem Tertiary biostratigraphic age correlation schemes in use today. The tectonic inversion of deep- and shallow-water basins related late Tertiary migration of the Caribbean plate provided Wim and his colleagues thick onland exposures for identifying and correlating planktic and benthic formaninifera. Wells drilled during oil exploration led to rapid increases in the understanding of the exposed geology of the basin edges and the subsurface geology of basins like the Magdalena basin of Colombia, the Maracaibo and Eastern Venezuela basins of Venezuela, and the Southern basin of Trinidad. In his initial assignments for Shell in Venezuela and Colombia, Wim further developed the use of ostracodes as stratigraphic markers in exploration wells, but, after his transfer in 1950 to Trinidad, he participated with Hans Bolli of Trinidad Leaseholds in development of the first planktic formaniniferal zonation of that island. During this period, Hans G. Kugler of Trinidad Leaseholds and later the University of Basel provided him access to ostracode material from the tectonically active Northern and Southern basins of Trinidad and underlying 'passive margin' section that provided him the basis of many significant papers on the Cenozoic biostratigraphy and regional geology of Trinidad in the late 1950's and early 1960's. Years at Louisiana State University. In 1958, Wim was invited to join the Department of Geology of Louisiana State University (LSU) by Henry V. Howe, a leading ostracode researcher and founding chairman of that department. His hiring of Wire insured the continuation of ostracode studies at LSU along with the continued growth of the geology department. During his early years at LSU, Wim continued his studies of the biostratigraphy of the Trinidad area with seminal papers on the ages and environments of the Brasso Formation (1958), the Eocene and Oligocene section (1960), and the Upper Miocene and Pliocene section (1963) of Trinidad. His teaching duties in the small department included structural geology and igneous and metamorphic petrology along, with a long-running tenure as instructor for the department's summer field camp in the Front Range of Colorado. Students, impressed by his ability to toil in the field for long days without water or rest, nicknamed him 'the Camel'. During the mid-to-late 1960's, Wire expanded his scale of observation of Caribbean basins to include field-based studies in Guatemala, Panama, Jamaica, Puerto Rico, Antigua, Anguilla, St. Martin, St. Croix, Costa Rica, Nicaragua, the Lesser Antilles, and Mexico. He systematically described the ostracode faunas in these areas and, where possible, related these faunas to better known biostratigraphic frameworks based on planktic foraminifera. The resulting regional correlations laid the framework for looking at the Caribbean region as a large-scale tectonic and sedimentary system rather than as a mosaic of individual and unrelated basins. In 1971, Wire returned to Cuba, the main topic of his Ph.D. study more than 30 years before. Through the invitation of Cuban geologists, Wim began a field-based study of the Neogene basins of the Matanzas and Santiago areas. In returning to Cuba at the height of the Cold War and the nadir of Cuba-US relations, Wim became one of the first US-based earth scientists to 'break the ice' and begin collaborative work with Cuban geologists. From the late 1960's to the present, graduate student enrollment at LSU increased and Wim supervised the research of many masters and doctoral students working on ostracodes and foraminifera. These student projects adapted Wire's tenets of establishing the basic geology and stratigraphy, determining the microfauna in a systematic way and cross-comparing ostracode and planktic foraminiferal faunas, and using ostracodes for stratigraphic correlation and assessment of marine vs. brackish conditions. Some of these students have gone on to become influential researchers and teachers in industry and academia, where they have in turn passed his time-proven methods of basinal studies on to a third generation of researchers. Follow-up studies to some of his work in this volume include Gill et al. for St. Croix, Babb and Mann for Trinidad, and Mann et al. for the Dominican Republic. Retirement years. In 1990, Wim retired with his wife Nettie to Baarn, Holland, near Utrecht. His children are now living both in the U.S. and Holland. He is currently a Professor Emeritus
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at his alma mater, the University of Utrecht, where he is completing a monograph on Caribbean ostracodes. His mailing address there is: Dr. Willem A. van den Bold Rijksuniversiteit te Utrecht Instituut voor Aardwetenschappen Budapestlaan 4, Postbus 80.021 3508 TA Utrecht, Netherlands
Dedication. We dedicate this volume to Wim van den Bold in acknowledgement of his tireless pursuit of his research goals and for his unselfish sharing of his knowledge of Caribbean basins and microfauna with us and his other students and colleagues. We dedicate this volume to him as not only a mark of respect, but also one of affection. Acknowledgement. We thank Hans van den Bold for the photograph and some of the information in this dedication. This dedication was submitted by Peter E McLaughlin, Jr. (Exxon Exploration Company) and Barun K. Sen Gupta (Louisiana State University). Selected dissertations supervised by Willem A. van den Bold at Louisiana State University (1971-1989) George Esker, 1968, Biostratigraphy of the Cretaceous-Tertiary Boundary in the East Texas Embayment Based on Planktonic Foraminifera. Paul Steineck, 1973, Paleoecologic and Systematic Analysis of Foraminifera from the EoceneMiocene Montpelier and Lower Coastal Groups, Jamaica, West Indies. Burton Bordine, 1974, Neogene Biostratigraphy and Paleoenvironments, Lower Magdalena Basin, Colombia. Walter Kessinger, 1974, Stratigraphic Distribution of the Ostracoda of Comanche (Cretaceous) Series of North Texas. Gilbert Taylor, 1975, The Geology of the Limon Basin of Costa Rica. Patricia Fithian, 1980, Distribution and Taxonomy of the Ostracoda of the Paria/Trinidad/Orinoco Shelf. Maria-Luisa Machain Castillo, 1985, Pliocene Ostracoda of the Saline Basin, Veracruz, Mexico. Peter E McLaughlin, 1989, Neogene Basin Evolution in the Southwestern Dominican Republic: A Foraminiferal Study (co-supervised by B.K. Sen Gupta).
SELECTED PUBLICATION LIST W.A. VAN DEN BOLD Bold, W.A. van den, 1946. Contribution to the study of Ostracoda with special reference to the Tertiary and Cretaceous microfauna of the Caribbean region. Doctoral thesis, Utrecht Univ., 175 pp. Reprint Antiq. Junk, 1970, with addition of a list of changes in genetic and specific assignments of species, and collection numbers of holo- and paratypes. Bold, W.A. van den, 1950. Miocene Ostracoda from Venezuela. Journal of Paleontology 24 (1), 76-88. Bold, W.A. van den, 1957. Ostracoda from the Paleocene of Trinidad. Micropaleontology 3 (1), 1-18. Bold, W.A. van den, 1957. Oligo-Miocene Ostracoda from southern Trinidad. Micropaleontology 3 (3), 231-254. Bold, W.A. van den, 1958. Ambocythere, a new genus of Ostracoda. Annals and Magazine of Natural History, ser. 12, 10, 801-813. Bold, W.A. van den, 1958. Ostracoda of the Brasso Formation of Trinidad. Micropaleontology 4 (4), 391-418. Bold, W.A. van den, 1960. Eocene and Oligocene ostracoda of Trinidad. Micropaleontology 6 (2), 145-196. Bold, W.A. van den, 1960. The genus Eucyitheridea Bronstein (Crustacea, Ostracoda) with a redescription of the type species. Annals and Magazine of Natural History, ser. 13, 4 (41), 283-303. Bold, W.A. van den, 1961. Some new Ostracoda of the Caribbean Tertiary. Koninkl. Nederlandse
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Akad. Wetensch. Proc., ser. B., 64 (5), 627-639. Bold, W.A. van den, 1963. Anomalous hinge structure in a new species of Cytherelloidea. Micropaleontology 9 (1), 75-78. Bold, W.A. van den, 1963. Ostracods and Tertiary stratigraphy of Guatemala. Bulletin of the American Association of Petroleum Geologists 47 (4), 696-698. Bold, W.A. van den, 1963. The ostracode genus Orionina and its species. Journal of Paleontology 37 (1), 33-50. Bold, W.A. van den, 1963. Upper Miocene and Pliocene Ostracoda of Trinidad. Micropaleontology 9 (4), 361-424. Bold, W.A. van den, 1964. Ostracoden aus der Oberkreide von Abu Rawash, Aegypten. Palaeontographica, Abteilung A: Palaeozoologie-Stratigraphie 123; 111-135. Bold, W.A. van den, 1965. Middle Tertiary Ostracoda from northwestern Puerto Rico. Micropaleontology 11 (4), 381-414. Bold, W.A. van den, 1965. Pseudoceratina, a new genus of Ostracoda from the Caribbean. Koninkl. Nederlandse Akad. Wetensch. Proc., ser. B, 68 (3), 160-164. Bold, W.A. van den, 1965. New species of the ostracod genus Ambocythere. Annals and Magazine of Natural History, ser. 13, 8 (1), 1-18. Bold, W.A. van den, 1966. Les ostracodes du Neogene du Gabon. Revue de l'Institut Frawais du Petrole 21 (2), 155-176. Bold, W.A. van den, 1966. Miocene and Pliocene Ostracoda from northeastern Venezuela. Koninkl. Nederlandse Akad. Wetensch. Verh., Affi. Natuurk., ser. 1, 23 (3), 43 pp. Bold, W.A. van den, 1966. Ostracoda from Col6n Harbour, Panama. Caribbean Journal of Science 6, 43-64. Bold, W.A. van den, 1966. Ostracoda from the Antigua Formation (Oligocene, Lesser Antilles). Journal of Paleontology 40 (5), 1233-1236. Bold, W.A. van den, 1966. Ostracoda of the Poz6n section, Falc6n, Venezuela. Journal of Paleontology 40 (1), 177-185. Bold, W.A. van den, 1966. Ostracode zones in Caribbean Miocene. Bulletin of the American Association of Petroleum Geologists 50 (5), 1029-1031. Bold, W.A. van den, 1966. Upper Miocene Ostracoda from the Tubara Formation (northern Colombia). Micropaleontology 12 (3), 360-364. Bold, W.A. van den, 1967. Miocene Ostracoda from Costa Rica. Micropaleontology 13 (1), 7586. Bold, W.A. van den, 1967. Ostracoda of the Gatun Formation, Panama. Micropaleontology 13 (3), 306-318. Bold, W.A. van den, 1968. Ostracoda of the Yaque Group (Neogene) of the northern Dominican Republic. Bulletins of American Paleontology 54 (239), 106 pp. Bold, W.A. van den, 1969. Messinella, a new genus of Ostracoda in the Caribbean Cenozoic. Micropaleontology 15 (4), 397-400. Bold, W.A. van den, 1969. Neogene Ostracoda from southern Puerto Rico. Caribbean Journal of Science 9 (3-4), 117-125. Bold, W.A. van den, 1970. Ostracoda of the lower and middle Miocene of St. Croix, St. Martin and Anguilla. Caribbean Journal of Science 10 (1-2), 35-61. Bold, W.A. van den, 1970. The genus Costa (Ostracoda) in the upper Cenozoic of the Caribbean region. Micropaleontology 16 (1), 61-75. Bold, W.A. van den, 1971. Ostracoda of the Coastal Group of formations of Jamaica. Transactions Gulf Coast Association of Geological Societies 21,325-348. Bold, W.A. van den, 1971. Ostracode associations, salinity and depth of deposition in the Neogene of the Caribbean region. In: Paleoecology of ostracodes. Centre de Recherches de Pau (Societe Nationale des Petroles d'Aquitaine), Bulletin 5 (5), 449-460. Bold, W.A. van den, 1972. Ostracodos del post-Eoceno de Venezuela y regiones vecinas. In: Congreso Geologico Venezolano, 4th, Memoria, Vol. 2. Boletfn de Geologfa Publicaci6n Especial 5, 999-1063. Bold, W.A. van den, 1973. Distribution of Ostracoda in the Oligocene and Lower and Middle Miocene of Cuba. Caribbean Journal of Science 13 (34), 145-159. Bold, W.A. van den, 1973. Nota geologica; notas sobre los ostracodos de la Formaci6n Punta Gavilan. Boletfn de Geologfa (Caracas) 12 (22), 333-335. Bold, W.A. van den, 1973. Ostracoda of the La Boca Formation, Panama Canal Zone. Micropale-
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ontology 18 (4), 410-442. Bold, W.A. van den, 1974. Neogene of Central Haiti. AAPG Bulletin 58 (3), 533-539. Bold, W.A. van den, 1974. Ornate Bairdidae in the Caribbean. In: Bold, W.A. van den (Ed.), Geoscience and Man 6, Ostracoda, the Henry V. Howe memorial volume. Louisiana State University, School of Geoscience, Baton Rouge pp. 29-40. Bold, W.A. van den, 1974. Taxonomic status of Cardobairdia (Van den Bold, 1960) and Abyssocypris n. gen.; two deepwater ostracode genera of the Caribbean Tertiary. In: Bold, W.A. van den (Ed.), Geoscience and Man 6, Ostracoda, the Henry V. Howe memorial volume. Louisiana State University, School of Geoscience, Baton Rouge pp. 65-79. Bold, W.A. van den, 1975. Distribution of the Radimella confragosa group (Ostracoda, Hemicytherinae) in the late Neogene of the Caribbean. Journal of Paleontology 49 (4), 692-701. Bold, W.A. van den, 1975. Neogene biostratigraphy (Ostracoda) of southern Hispaniola. Bulletins of American Paleontology 66 (286), 549-639. Bold, W.A. van den, 1975. Ostracodes from the late Neogene of Cuba. Bulletins of American Paleontology 68 (289), 119-167. Bold, W.A. van den, 1975. Remarks on ostracode-biostratigraphy of the late and middle Tertiary of southwest Puerto Rico. Caribbean Journal of Science 15 (1-2), 31-40. Bold, W.A. van den, 1976. Distribution of species of the tribe Cyprideidini (Ostracoda, Cytherideidae) in the Neogene of the Caribbean. Micropaleontology 22 (1), 1-43. Bold, W.A. van den, 1976. Ostracode correlation of brackish-water beds in the Caribbean Neogene. Transactions of the Caribbean Geological Conference VII, Saint Francois, Guadeloupe, June 30-July 12, 1974, pp. 169-175. Bold, W.A. van den, 1980. Notes on the distribution of some mid-Tertiary ostracodes of Puerto Rico. In: Llinas, C.R., Gil, N., Seaward, M., Tavares, I., and Snow, W. (Eds.), Transactions of the 9th Caribbean geological conference, Santo Domingo, Dominican Republic, Aug. 16-20, 1980, Vol. 1, pp. 225-230. Bold, W.A. van den, 1981. Distribution of Ostracoda in the Neogene of central Haiti. Bulletins of American Paleontology 79 (312), 136 pp. Bold, W.A. van den, 1983. Shallow-marine biostratigraphic zonation in the Caribbean post Eocene. Proc. 8th Int. Symp. Ostracoda, 400-416. Bold, W.A. van den, 1985. Heinia, a new genus of Ostracoda from the Gulf of Mexico and the Caribbean. Journal of Paleontology 59 (1), 1-7. Bold, W.A. van den, 1986. Distribution of Ostracoda at the Eocene-Oligocene boundary in deep (Barbados) and shallow-marine environment (Gulf of Mexico). In: Pomerol, C., and Premoli-Silva, I. (Eds.), Terminal Eocene events. Elsevier, Amsterdam, pp. 259-263. Bold, W.A. van den, 1986. Fresh and brackish water Ostracoda from the Neogene of northern Venezuela. Tulane Studies in Geology and Paleontology 19 (34), 141-157. Bold, W.A. van den, 1988. Neogene paleontology in the northern Dominican Republic; 7, The subclass Ostracoda (Arthropoda). Bulletins of American Paleontology 94 (329), 105 pp. Bold, W.A. van den, 1988. Ostracoda of Alacran Reef, Campeche Shelf, Mexico. Tulane Studies in Geology and Paleontology 21 (34), 143-155. Bold, W.A. van den, 1989. Ostracoda of the Montezuma Formation, Pliocene, Nicoya Peninsula, Costa Rica. Tulane Studies in Geology and Paleontology 22 (1-2), 61-64. Bold, W.A. van den, 1990. Late Holocene Ostracoda in and around Lake Enriquillo, Dominican Republic. In: Larue, D.K., and Draper, G. (Eds.), Transactions of the 12th Caribbean geological conference, St. Croix, United States Virgin Islands, Aug. 7-11, 1989, pp. 163-189 Bold, W.A. van den, 1990. Short review of the Ostracoda of the Montpelier and Coastal Groups of Jamaica. In: Larue, D.K., and Draper, G. (Eds.), Transactions of the 12th Caribbean geological conference, St. Croix, United States Virgin Islands, Aug. 7-11, 1989, pp. 99-103. Bold, W.A. van den, 1990. Stratigraphical distribution of fresh and brackish water Ostracoda in the late Neogene of Hispaniola. In: Whatley, R.C., and Maybury, C. (Eds.), Ostracoda and global events m Tenth International Symposium on Ostracoda. Chapman and Hall, pp. 221-232. Bold, W.A. van den, 1990. Note on the stratigraphy of the Bellad6re and Comendador structures, Neogene of Haiti and the Dominican Republic. Transactions of the 10th Caribbean geological conference (Cartagena, 1983), pp. 238-242. McLaughlin, EE, Bold, W.A. van den, and Mann, E, 1991. Geology of the Azua and Enriquillo basins, Dominican Republic; 1, Neogene lithofacies, biostratigraphy, biofacies, and paleogeogra-
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phy. In: Mann, E, Draper, G., and Lewis, L.E (Eds.), Geologic and tectonic development of the North America-Caribbean Plate boundary in Hispaniola. Geological Society of America Special Paper 262, 337-366. McLaughlin, EE, Gill, I.E, and Bold, W.A. van den, 1995. Biostratigraphy, paleoenvironments and stratigraphic evolution of the Neogene of St. Croix, U.S. Virgin Islands. Micropaleontology 41 (4), pp. 293-320.
CARIBBEAN SEDIMENTARY BASINS Preface The purpose of this volume is to provide a collection of original studies of Caribbean basins that were mainly carried out in the early and mid-1990's. It is hoped that this collection of studies will become a valuable source of information for future students of Caribbean basins and will help them better focus their research. I solicited original contributions crafted with a regional tectonic scope rather than more localized geologic reviews mainly because there have been several recent and comprehensive volumes on Caribbean geology. These volumes include: 9 Geological Society of America, Decade of North American Geology, volume H, on the geology and geophysics of the Caribbean region, edited by Gabriel Dengo and James E. Case (1990). This benchmark volume contains a wealth of information on Caribbean geology and basins arranged in a comprehensive area by area manner. 9 Geological Society of America Special Paper 295 on the on- and offshore geology and geophysics of southern Central America in Panama and Costa Rica, edited by myself (1995) (fig. 1B). This volume contains a complete summary of basin studies in southwestern Caribbean through the early 1990's. 9 American Association of Petroleum Geologists, Petroleum Basins of South America, edited by A.J. Tankard, R. Su~irez and H.J. Welsink (1995). This award-winning volume provides comprehensive coverage of basins in the northern and northwestern areas of South America. Given these objectives and existing Caribbean volumes, the scope and emphases of this volume include the following: 9 A focus on sedimentary basins in the region of the Caribbean Sea and its margins. The common tectonic events discussed in the chapters of this volume include the rifting and passive margin history of North and South America that led to the formation of the Caribbean region, the entry of an exotic, Pacific-derived Great Arc of the Caribbean at the leading edge of the Caribbean oceanic plateau, the terminal collision of the arc and plateau with the passive margins fringing North and South America, and subsequent strike-slip and accretionary tectonics that affected the arc-continent collision zones (Fig. 1). 9 An emphasis on new subsurface and potential field data (well logs, shallow and deep penetration seismic reflection, ship-based and satellite-based gravity and magnetics, and aeromagnetics). These data are leading to rapid breakthroughs in our mainly outcrop-based understanding of Caribbean tectonic history and Caribbean tectonic processes. The papers in this volume attempt to synthesize these subsurface and potential field data at a regional or basin-wide scale to facilitate tectonic interpretations. 9 Inclusion of new biostratigraphic and structural data on the highly deformed and sometimes metamorphosed 'bits and pieces' of Caribbean sedimentary basins that are important for tectonic reconstructions. These types of data, which vary from biostratigraphic dating studies to structural and isotopic dating studies, are needed in this region where many of the major tectonic events occurred in the Cenozoic and have strongly overprinted the existing Cretaceous and older rocks. 9 A regional plate tectonic introduction to Caribbean basins in the form of two introductory chapters by myself and Mtiller et al. These chapters attempt to place Caribbean basins into a wider regional context using revised quantitative plate reconstructions. This regional context helps to better understand their often complex tectonic settings, short-lived subsidence mechanisms, and later inversion histories. The last systematic effort at a quantitative Caribbean plate model was by Ross and Scotese (1988) and was not able to incorporate the recent advances in seafloor mapping using satellite-based gravity measurements that have now become available (Sandwell and Smith, 1997). The volume is divided into six parts. Part 1 consists of the two tectonic setting papers by myself and Mtiller et al.
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Fig. 1. (A) Political subdivision of the Caribbean region. (B) Locations of chapters in this volume (numbers refer to chapter number). Chapters 1 and 2 cover the entire area shown and Chapter 3 focusses on the central part of Mexico that is not shown on this map. Diagonally lined area represents the Pacific-derived Caribbean oceanic plateau province, Great Arc of the Caribbean and area of back-arc basins formed during eastward and northeastward migration of the Great Arc.
In Chapter 1, I use Geosat gravity data from Sandwell and Smith (1997) to illustrate the large-scale tectonic features of the Caribbean with emphasis on Cenozoic sedimentary basins visible on the gravity maps. I classify these basins using a simplified basin classification scheme. I present a series of thirteen plate reconstructions based on the North A m e r i c a - S o u t h America motions used with permission from the Mtiller et al. study and use these reconstructions to place the major basin-forming events into a plate tectonic context. These tectonic events support the idea of Pindell and Barrett (1990) and previous workers that the Caribbean is a Pacific-derived oceanic plateau that diachronously collided with the passive margins of North and South America during
SEDIMENTARY BASINS OF THE CARIBBEAN m PREFACE
XVII
Cenozoic time. To conclude this overview, I suggest possible avenues of future research based on this compilation and the new results presented in the other chapters of this volume. In Chapter 2, Mfiller et al. compute plate motions and uncertainties of motions for the Mesozoic-Cenozoic interactions of North and South America using magnetic anomaly and satellite gravity-based fracture zone identifications from the Atlantic Ocean. This refined data set supports previous plate models for the Caribbean by showing the two Americas drifted apart until about 71 Ma (Campanian) and then underwent a steady convergence across the area of the present-day Caribbean Sea from about 55 Ma (Eocene) to the present. The authors propose that their much better defined Cenozoic convergent phase is responsible for shaping many of the convergent zones and basins along the east-west-trending northern and southern margins of the Caribbean plate. Part 2 consists of five chapters mainly focussed on basins overlying the North America plate and recording its rifting from South America in Late Jurassic to Cretaceous time. Because the structures related to this rift plane in Caribbean tectonic history are largely covered by water and Cretaceous and younger carbonate platform rocks, three of these chapters rely heavily on either well and seismic reflection data or on outcrop data from deformed sections now exposed onland. Chapter 3 by Marton and Buffler uses seismic reflection data and the results of DSDP Leg 77 to define major tectono-stratigraphic sequences in the area of the southeastern Gulf of Mexico (Fig. 1). This area is critical to understanding the early rift history of the Gulf of Mexico and North and South America opening because it is adjacent to, but not overprinted by, the Paleogene thrust event of Cuba. These data indicate that the Yucatan Peninsula underwent a major counter-clockwise rotation from Oxfordian to late Berriasian time which produced a series of rifts in the southeastern Gulf of Mexico. As rifting ceased, and the rift blocks subsided Early Cretaceous carbonate platforms formed atop the rift blocks. These workers attempt to correlate the Mesozoic offshore stratigraphic record of the southeastern Gulf of Mexico with onshore data described by Pszcz6~kowski in chapter 4 from western Cuba. Chapter 4 by Pszcz6~kowski chronicles the major stratigraphic and tectonic events known from the rifted and passive margin of North America now exposed by Paleogene thrusting and subsequent Cenozoic faulting in western Cuba (Fig. 1). The major tectono-stratigraphic events included a rift event (Lower Jurassic-?Callovian-early Oxfordian), a passive margin subsidence event (?Callovian/middle Oxfordian-Santonian) and the beginning of the arc-continental collision phase (Campanian-Paleocene). Chapter 5 by Pessagno et al. present new stratigraphic, age, and paleobathymetric data from the area of Mexico southwest of the northwest-striking Walper megashear, a major terrane boundary that bisects Mexico and thought to have been active during the early rift period between North and South America. Faunal (radiolaria and megafossils) and paleomagnetic data indicate that the area southwest of the megashear was transported from higher paleolatitudes during the Oxfordian and reached lower paleolatitudes by the Berriasian. The authors propose that terranes southwest of the megashear share similar stratigraphic signatures with the Jurassic and Early Cretaceous rocks of western Cuba that are described in detail in Chapter 4 by Pszcz6~kowski (Fig. 1). The authors argue that the Mexican and Cuban rocks may have formed a once continuous San Pedro del Gallo terrane prior to disruption of the Cuban part of the terrane by Paleogene collision of the Caribbean arc. Chapter 6 by Scott and Finch presents new stratigraphic and paleontologic data from a Berriasian-Albian carbonate platform formed above the Chortis block in Honduras (Fig. 1). The Chortis block provides a critical for clues to the paleoposition of the Caribbean because the Chortis block is the only subaerial and continental part of the present-day Caribbean plate. At the time of Mesozoic rifting between North and South America, the Chortis block probably formed a southern extension of the continental area of southern Mexico. The ages and faunas reported by Scott and Finch support the idea that by Aptian and Albian time the Chortis block had established a biogeographic connection to the Caribbean. In Chapter 7, Masaferro and Eberli use multi-channel seismic reflection lines to reveal the nature and evolution of the Great Bahama bank, a Late Jurassic-Recent carbonate platform formed during the early rifting between North and South America (Fig. 1). Their data show Early Jurassic rifts and extensional structures overlain by a ca. 5-kin-thick carbonate platform. The carbonate platform shows strong fault control that is interpreted as an effect of the collision between the Caribbean arc and the Great Bahama bank in Late Cretaceous-Middle Eocene time. Fault activity largely ended with the collision in Late Eocene time and fault scarps and tectonic depressions were masked by infilling of the highly productive carbonate platform.
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Part 3 consists of six chapters that focus on smaller, usually heavily faulted and onshore Cenozoic basins of the northern Caribbean that formed in response to arc, collisional, and strike-slip activity between the evolving North America-Caribbean plate boundary. In Chapter 8, Av6 Lallemant and Gordon present new structural and isotopic data from Roatan Island, which is one of the Bay Islands off the coast of northern Honduras (Fig. 1). Structures in metamorphic and sedimentary rocks exposed on these islands include older, ductile structures formed at metamorphic conditions during the Late Cretaceous to early Tertiary left-oblique collision between the Chortis and Maya (southern Mexico) blocks. Most brittle faults are younger and formed after the Late Eocene or Early Oligocene exhumation of the metamorphic rocks. Younger structures indicate transtension related to North America-Caribbean strike-slip motion along the Cayman trough. The Bay Islands form the northern flank of the offshore Tela sedimentary basin. These data indicate a transtensional origin for this basin in late Neogene time. In Chapter 9, Manton and Manton present new stratigraphic and age data from Miocene marine sedimentary rocks from the southern margin of the Tela basin in the fault-bounded, coastal range of northern Honduras (Fig. 1). These data suggest that this range is a deformed turbiditic basin of Miocene age that has been inverted along strike-slip faults related to the North America-Caribbean plate boundary. Deformed sedimentary rocks of this inverted range may have once been continuous with undeformed sedimentary rocks of the Tela basin to the north. In Chapter 10, Montgomery and Pessagno present new identifications of radiolaria and foraminifera from deformed, deep basinal rocks faulted against Caribbean Cretaceous arc rocks in Jamaica and Hispaniola (Dominican Republic) (Fig. 1). The authors interpret these rocks as accreted crust of the proto-Caribbean and Atlantic seafloor that was subducted to the west and southwest beneath the Caribbean arc. Pacific-derived fragments subducted to the east and northeast beneath the Caribbean are also present in these areas and suggest a complex subduction polarity reversal in the Caribbean arc sometime during the Cretaceous. In Chapter 11, de Zoeten and Mann present new stratigraphic, paleocurrent, and sandstone petrographic data from Paleocene to Early Pliocene basinal and carbonate platform rocks of northern Hispaniola (Dominican Republic) (Fig. 1). We use this sedimentary record to interpret three major tectonic phases that affected this segment of the North America-Caribbean plate boundary: (1) an Early to Middle Eocene collisional event between the Caribbean arc and the Bahama carbonate platform; (2) a period of strike-slip faulting and basin formation from Late Eocene to Early Miocene; and (3) a period of transpressional uplift and folding from Late Miocene to Early Pliocene. In Chapter 12, Mann et al. present new outcrop, well, and seismic reflection data from the Enriquillo basin of Hispaniola (Dominican Republic) and document the existence of two distinct facies of Early Pliocene evaporites which formed in the center and edges of the tectonically active ramp basin within the North America-Caribbean strike-slip zone. Analysis of seismic reflection data tied to an exploration well in the center of the basin showed that the basin center deposit is a ca. 1500-m-thick halite and gypsum deposit whose depocenter was controlled by reverse faults that crosscut the center of the basin. Three occurrences of basin-edge, tidal flat gypsum deposits are interpreted as having been deposited during Early Pliocene eustatic falls in sea level. In Chapter 13, Gill et al. use stratigraphic and paleontologic data from both wells and outcrops to document the evolution of the Neogene Kingshill basin on St. Croix (U.S. Virgin Islands) within the North America-Caribbean strike-slip plate boundary. These data reveal a deep-water, pre-basinal section of Oligocene to early Middle Miocene age which underwent transtensional faulting no later than late Middle Miocene to form the rift-like basin. They relate this transtensional event to oblique, left-lateral opening between the islands of St. Croix and Puerto Rico. Near the end of the Middle Miocene, the occurrence of shallow-water limestone indicates an uplift of the basin-bounding blocks. This uplift continued through the Pliocene when St. Croix acquired its approximate present-day land area. Part 4 consists of two chapters that focus on Cenozoic basins related to the Lesser Antilles arc system of the eastern Caribbean. In Chapter 14, Huyghe et al. use seismic reflection profiles and sidescan sonar images to review the tectonic control and sedimentary patterns of late Neogene piggyback basins of the accretionary wedge formed in Cretaceous?-Cenozoic time in front of the east-facing Lesser Antilles island arc. These basins provide important modem examples of piggyback basins that are now preserved in older more deformed areas of the circum-Caribbean like Cuba and Venezuela. Their data show two
SEDIMENTARY BASINS OF THE CARIBBEAN - - PREFACE
XIX
main stages in the history of these piggyback basins: (1) rapid tilting of the basin in response to the growth of a fault-bend fold at one edge of the basin; (2) reactivation of the bounding thrust surface by folding or sediment diapirism. In Chapter 15, Bird et al. integrate gravity, seismic reflection and refraction data with their previously published magnetic data from the Grenada back-arc basin to produce a model for its east-west opening during Paleogene time. Disruption of basin gravity and magnetic trends in the northern part of the basin is attributed to localized tectonic shortening effects. Part 5 consists of three chapters on the Jurassic-Recent sedimentary basins of the eastern Venezuela and Trinidad area of the southeastern Caribbean. These basins reflect both the JurassicCretaceous rifting and passive margin history of separation between the North and South America plates as well as a much younger phase of Oligocene to Recent transpression between the eastward-migrating Lesser Antilles arc and accretionary wedge and the South American continent. In Chapter 16, di Croce et al. use seismic reflection data tied to wells in the eastern Venezuela and Trinidad areas to propose four distinct tectono-stratigraphic phases for this part of the South America margin: (1) an ill-defined Paleozoic/pre-Jurassic pre-rift phase; (2) a Jurassic syn-rift phase related to the separation of North and South America; (3) a Cretaceous to Oligocene passive margin phase related to the thermal subsidence of the South American rifted margin; and (4) a Neogene foredeep phase related to the oblique collision between the Lesser Antilles arc and accretionary prism and the South American continent. The passive margin period of phase 2 from 131 Ma to 30 Ma is subdivided into five second-order transgressive-regressive cycles interpreted as eustatic sea-level fluctuations superimposed on the thermally subsiding margin. The Oligocene and younger history of the margin and its major sequences is dominated by the formation of a collisional foredeep basin produced by the oblique collisional event and rapid basin infilling in response to uplift and erosion of the Andes. In Chapter 17, Flinch et al. use seismic reflection data and wells from eastern Venezuela, the Gulf of Paria and Trinidad to reveal the presence of a submarine late Neogene pull-apart basin beneath the shallow Gulf of Paria west of Trinidad (Fig. 1). The pull-apart formed at a stepover between the fight-lateral Casanay-Arima fault zone to the north and the Warm Springs fault zone to the south. These data also reveal northward thrusting coeval with the pull-apart formation that is interpreted as a passive roof duplex formed above deeper southward-directed thrusts. In Chapter 18, Babb and Mann integrate seismic reflection and well data from the Northern and Gulf of Paria basins of Trinidad with existing outcrop and map information to identify three Neogene deformational phases related to initial Late Miocene-Early Pliocene movement along the E1 Pilar fault zone, formation of the Gulf of Paria pull-apart basin between the E1 Pilar and Warm Springs-Central Range fault zones, and increasing movement on the Warm Springs-Central Range fault zones. We relate the southward expansion of the late Neogene deformational zone in Trinidad to the presence of oceanic crust southeast of Trinidad that allows for the southeastward migration of strike-slip fault-bounded blocks in that direction. Part 6 consists of three chapters containing revealing new information on the Cretaceous Caribbean igneous plateau that forms the basement of the central Caribbean Sea beneath the Venezuelan and Colombian basins (Fig. 1). All three chapters make use of deep penetrating multi-channel seismic data to show the stratigraphic and structural features of this oceanic plateau province which was previously known from isolated outcrops in the circum-Caribbean and from widely spaced DSDP sites in the Venezuelan and Colombian basins. In Chapter 19, Diebold and Driscoll use these deep-penetrating seismic lines to reveal previously unseen structures within the entire thickness of the Cretaceous Caribbean oceanic plateau basaltic crust east of the Beata Ridge. The lower basaltic section consists of submarine basaltic volcanoes with dipping flanks that appear to have maintained their primary sense of dip. Flows can be followed for distances of 20 to 100 km and appear to have been erupted on an existing section of rough ocean floor of normal thickness. The upper basaltic sequence, whose top defines the well known B" reflector, also contains widespread flows but is more heterogeneous than the lower unit and fills morphological and extensional lows in the top of the lower sequence. Rifting and flexural uplift have faulted both sequences along the Beata Ridge and large escarpments. The age relations between the lower sequence and the upper sequence are not known since only the top of the upper sequence (B" reflector) has been drilled. In Chapter 20, Driscoll and Diebold use the same seismic reflection data set to document the history of rift features in the upper basaltic section and the sedimentary and tectonic history of the
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section overlying the basaltic basement described by them in Chapter 19. Extensional deformation accompanying the formation of the rifted upper basaltic section of the Cretaceous plateau led to the formation of the large escarpments along the Hess escarpment and Beata Ridge. The sedimentary section between the B" and A" reflectors was probably derived from the uplift of ranges in northern South America, is correlated to DSDP and ODP sites, and is constrained to be younger than Senonian (~88 Ma) and older than Middle Eocene (~50 Ma). This section exhibits only minor Neogene faulting along the Beata Ridge but does exhibit reworking and drifts related to bottom currents. Chapter 21 by Mauffret and Leroy uses an extensive seismic reflection, gravity, and magnetic data set that is centered on the Beata Ridge. Regional mapping reveals post-Early Miocene strike-slip and compressional features along the eastern edge of the Beata Ridge that formed by a northeast-southwest-shortening event. These young deformational features increase from south to north along the trend of the Beata Ridge. The authors propose that these faults are related to shortening between the subduction zone along the northern margin of South America and the subduction and strike-slip zone in Hispaniola and the Muertos trench. Additional color illustrations are available at the Sedimentary Basins of the Caribbean site on the Internet ~. The supplementary material includes color versions of selected figures from chapters 1, 2, 11, 12, 14 and 17 and additional figures for chapters 5, 7 and 15. P. MANN (Editor)
REFERENCES
Dengo, G. and Case, J.E. (Eds.), 1990. The Caribbean Region. The Geology of North America, Vol. H, Geological Society of America, Boulder, Colorado, 528 pp. with attached volume of map enclosures. Mann, E (Ed.), 1995. Geologic and Tectonic Development of the Caribbean Plate Boundary in Southern Central America. Geological Society of America Special Paper 295, 349 pp. Pindell, J.L. and Barrett, S.E, 1990. Geologic evolution of the Caribbean region. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Vol. H, Geological Society of America, Boulder, Colorado, pp. 405-432. Ross, M.I. and Scotese, C.R., 1988. A hierarchical tectonic model of the Gulf of Mexico and Caribbean region. Tectonophysics, 155: 139-168. Sandwell, D.T. and Smith, W.H.E, 1997. Marine gravity anomaly from Geostat and ERS-1 satellite altimetry. Journal of Geophysical Research, 102: 10,039-10,054. Tankard, A.J., Sufirez Soruco, R. and Welsink, H.J., 1995. Petroleum Basins of South America. American Association of Petroleum Geologists Memoir 62, 792 pp.
1http://www.elsevier.nl/locate/caribas/with mirror sites: http://www.elsevier.com/locate/caribas/and http://www.elsevier.jp/locate/caribas/
Reviewers I would like to thank the following people for volunteering their time to review the papers in this volume. Lewis Abrams James A. Austin, Jr. Hans G. Av6 Lallemant Nathan Bangs Charles D. B lome Burke Burkart Kevin C. Burke Eric Calais Millard E Coffin Daniel M. Davis James E Dolan Thomas W. Donnelly Grenville Draper Robert A. Duncan Paul A. Dunn Peter Emmet Lee Gerhard Jan Golonka Mark B. Gordon Nancy R. Grindlay Chistoph E. Heubeck Albert C. Hine Troy L. Holcombe
Frederick Hutson Keith James James N. Kellogg John E Lewis Fernando Martinez Gyorgy L. Marton Kristian Meisling Henry E. Mullins Ian Norton Steven Pierce James L. Pindell Walter Pitman Andrzej Pszcz6tkowski Edward Robinson Robert Rogers Amos Salvador Kathryn M. Scanlon Charlotte B. Schreiber Robert E. Sheridan Norm Silberling Carol Telemaque Elazar Uchupi R MANN (Editor)
List of Contributors* *
W. ALl 17 Trinmar Point Fortin Trinidad and Tobago
J.E CASEY 15 Department of Geosciences University of Houston Houston, TX 77204-5503, USA
G. AVI~ LALLEMANT 8 Department of Geology and Geophysics, MS-126 Rice University Houston, TX 77005-1892, USA
Y. DENIAUD 14 Ddpartement de G6ologie et Oc6anographie m URA 197 Avenue des Facult6s 33405 Talance, France
S.E. BABB 18
Institute for Geophysics University of Texas at Austin 4412 Spicewood Springs Road, Bldg. 600 Austin, TX 78759, USA A.W. BALLY 16 Department of Geology and Geophysics, MS-126 Rice University Houston, TX 77005-1892, USA D.E. BIRD 15 Bird Geophysical 16903 Clan Macintosh Houston, TX 77084, USA
[email protected] R.T. BUFFLER 3 Institute for Geophysics University of Texas at Austin 4412 Spicewood Springs Road Bldg. 600 Austin, TX 78754, USA
[email protected] S.C. CANDE 2 Scripps Institution of Oceanography, UCSD 9500 Gilman Drive La Jolla, CA 92093-0215, USA cande @gauss.ucsd.edu A. CANTU-CHAPA 5 Department of Geology Florida International University University Park Miama, FL 33199, USA
* Superior ciphers refer to the chapter number. t E-mail address of first authors included when available.
V. DE LISA 17 Petroleos de Venezuela Edificio PDVE&P La Estancia Chauo, Caracas, Venezuela R. DE ZOETEN 11 Unocal Thailand Central Plaza Office Bldg. 2993 Phaholyothin Road Bangkok 10990, Thailand rdezoeten @unocal.com J. D I CROCE 16 J. DIEBOLD 19,20 Lamont-Doherty Earth Observatory of Columbia University Palisades, NY 10964-8000, USA johnd @lamont.ldgo.columbia.edu N. DRISCOLL 19,20 Geology and Geophysics Woods Hole Oceanographic Institution Woods Hole, MA 02543, USA ndriscoll @whoi.edu G.R EBERLI 7 Comparative Sedimentology Laboratory Rosentiel School of Marine and Atmospheric Science 4600 Rickenbacker Causeway Miami, FL 33149, USA gerberli @rsmas.miami.edu
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LIST OF CONTRIBUTORS
J-C. FAUGERES 14 D6partement de G6ologie et Oc6anographie URA 197 Avenue des Facult6s 33405 Talance, France
G. HERNANDEZ a7 Petroleos de Venezuela Edificio PDVE&P La Estancia Chauo, Caracas, Venezuela
R.C. FINCH 6 Department of Earth Sciences Box 5062 Tennessee Technological University Cookeville, TN 38505, USA rcf7332.tntech.edu
K.J. HSO 22 Geologisches Institut ETH-Ziirich ZUrich, Switzerland
J.E FLINCH 7 Departamento de Geologia Lagoven Caracas, Venezuela Present address:
Exploration and Production Total Paris la Defense, France joan.flinch @total.com I. GILL 13 Department of Geology University of Puerto Rico, Mayaguez E O. Box 5000 Mayaguez, Puerto Rico 00681-5000 i-gill @rumac.upr.clu.edu E. GONTHIER 14 D6partement de G6ologie et Oc6anographie m URA 197 Avenue des Facult6s 33405 Talance, France M.B. GORDON 8 Department of Geology and Geophysics MS-126, Rice University Houston, TX 77005-1892, USA Present address:
GX Technology 5847 San Filipe, Suite 3500 Houston, TX 77057, USA mgordon@ gtx.com R. GRIBOULARD 14 Ddpartement de G6ologie et Oc6anographie URA 197 Avenue des Facult6s 33405 Talance, France S.A. HALL 15 Department of Geosciences University of Houston Houston, TX 77204-5503, USA
D.K. HUBBARD 13 Virgin Islands Marine Advisors 5046 Cotton Valley Rd Christiansted, St. Croix 00820 D.M. HULL 5 Programs in Geosciences University of Texas at Dallas EO. Box 830688 Richardson, TX 77083, USA E HUYGHE 14 D6partement de G6ologie et Oc6anographie m URA 197 Avenue des Facultds 33405 Talance, France Present address:
Laboratoire de G6odynamique des Chaines Alpines et VPRESA, 15 rue Maurice Gignoux 38031 Grenoble cedex, France huyghe @uj f-grenoble.fr M. KELLDORF s Programs in Geosciences University of Texas at Dallas EO. Box 830688 Richardson, TX 77083, USA M.E. LAMAR 12 Department of Geological Sciences University of Texas at Austin Austin, TX 78713, USA Present address:
Scott and White Memorial Hospital Department of Obstetrics and Gynecology 2401 South 31st Street Temple, TX 76508, USA S.R. LAWRENCE 12 Exploration Consultants Ltd. Highlands Farm, Greys Road Henley-on-Thames, Oxon, RG9 4PR UK s.lawrence @ecgc.com
Chapter 1
Caribbean Sedimentary Basins" Classification and Tectonic Setting from Jurassic to Present
P. MANN
INTRODUCTION
Plate rates
The purpose of this introductory chapter is to describe the active tectonic setting of the Caribbean, its major crustal provinces, and to provide a simple classification for sedimentary basins in the Caribbean region. In addition to this background information on Caribbean basins, I provide a series of thirteen quantitative plate reconstructions based on the revised plate model of Mtiller et al. (Chapter 2). These reconstructions serve to place individual basins into a better tectonic framework.
Rates of relative plate motion as predicted by the Nuvel-lA plate motion model of DeMets et al. (1994) are relatively slow (11-13 mm/year) between the Americas and the Caribbean plate but much faster (59-74 mm/year) between the Cocos, Nazca and Caribbean plates (Fig. 1). Recent GPS-based studies of the relative motion between the North America and Caribbean plates in the northeastern Caribbean by Dixon et al. (1998) have shown that the actual North America-Caribbean rate of east-west strike-slip motion may be twice as fast as predicted by the Nuvel-lA plate motion model.
ACTIVE TECTONIC SETTING OF CARIBBEAN SEDIMENTARY BASINS
Earthquake studies Major plates The distribution of recorded earthquakes, active calc-alkaline volcanoes, and spreading ridges defines five rigid plates in the Caribbean region: North America, South America, Caribbean, and Nazca (Molnar and Sykes, 1969; Mann et al., 1990) (Fig. 1). Geologic and seismic studies indicate that the Caribbean plate is moving eastward relative to the Americas, and this movement is accommodated by left-lateral strike-slip faults along its boundary with the North America plate, and right-lateral strike-slip faults along its boundary with the South America plate. Oceanic lithosphere of the North and South America plates is consumed along the eastern edge of the Caribbean at the Lesser Antilles subduction zone. Oceanic lithosphere of the Cocos and Nazca plates is consumed along the western and southwestern edge of the plate at the Middle America subduction zone (Fig. 1).
There have been many first-motion studies on large Caribbean earthquakes over the past 25 years and I have compiled a representative group from the Harvard focal mechanism catalogue on Fig. 1. Because large earthquakes occur more frequently in subduction zone settings, many more focal mechanisms are available from the Lesser Antilles and Middle America arcs than for the northern and southern dominantly strike-slip plate boundaries. In general, subduction zone earthquakes are characterized by shallow thrust faulting with the auxiliary plane striking approximately parallel to the trend of the trench and dipping steeply away from the arc and with the fault plane striking subparallel to the arc trend and dipping gently beneath the arc. First-motion studies from the strike-slip plate boundaries are consistent with shallow focus leftlateral fault displacements along the northern plate boundary zone and right-lateral displacements along the southern plate boundary zone (Fig. 1).
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 3-31. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
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E MANN
Fig. 1. Earthquake focal mechanisms and plate motions relative to a fixed Caribbean plate based on the NUVEL-1A global plate motion model of DeMets et al. (1994). Numbers give rate of plate motion in mm/year and arrows give directions of plate motion. Focal mechanisms are color-coded according to depth: red mechanisms are from earthquakes from 0 to 75 km in depth; blue mechanisms are from earthquake from 75 to 150 km in depth; and green mechanisms are > 150 km in depth. Basemap is the Geosat gravity map of the Caribbean compiled by Sandwell and Smith (1997). In accordance with plate motion model predictions, earthquake focal mechanisms show predominantly left-lateral motion along the northern edge of the Caribbean plate, predominantly right-lateral motion along the southeastern edge of the plate, predominantly thrust motion at the eastern and western ends of the plate, and predominately northeast-directed right-lateral motion associated with the displacement of the Maracaibo block in the southwestern Caribbean. The Maracaibo block is a continental-arc fragment of northwestern South America that is escaping to the north and northeast as a response to the late Neogene collision of the Panama arc and oblique subduction of the northern Nazca plate.
First-motion studies along with geologic and GPS-based geodetic studies have shown that the northwestern corner of South America (Maracaibo block of Mann and Burke, 1990) is being displaced northward and northwestward along the B o c o n 6 eastern Andean right-lateral strike-slip fault system in Colombia and western Venezuela (McCann and Pennington, 1990; Kellogg and Vega, 1995) (Fig. 1). This displacement appears to be a consequence of late Neogene collision of the Panama arc with northwestern South America (Mann and Burke, 1990).
MAJOR CRUSTAL PROVINCES OF THE CARIBBEAN REGION
The Caribbean region consists of a rim of Cretaceous-Recent arc terranes and associated back-arc basins molded about a sub-circular core consisting of a continental fragment in the western Caribbean (Chortfs block) and an oceanic plateau province beneath the central and eastern Caribbean (cf. Case et al., 1990, for a comprehensive review of all Caribbean crustal provinces) (Fig. 2). Chortis block
S u b d u c t e d slabs
Inclined subducted slabs extend to depths of 150 km under most o f the land areas adjacent to the Lesser Antilles and Middle America arcs (McCann and Pennington, 1990; Dewey and Sufirez, 1991) (Fig. 1). Subducted slabs are also present beneath much of the northwestern comer of South America (van der Hilst and Mann, 1994).
The Chortfs block of northern Central America consists of well-dated Mesozoic and Cenozoic formations which unconformably overlie a continental basement of poorly dated, metamorphic rocks of Paleozoic and possible Precambrian age (Gordon, 1991) (Fig. 2). Seismic velocities confirm that the Chortfs block is continental (Case et al., 1990) but the isotopic ages of metamorphic protolith rocks
C A R I B B E A N S E D I M E N T A R Y BASINS" C L A S S I F I C A T I O N A N D T E C T O N I C S E T T I N G
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Fig. 2. Major crustal provinces of the Caribbean: Precambrian-Paleozoic Chortfs block; Cretaceous oceanic plateau of the central Caribbean; Early Cretaceous-Recent Great Arc of the Caribbean; and passive margins of North and South America. Red lines indicate active plate boundaries and white lines are magnetic anomaly and fracture zone trends from Coffin et al. (1992). Key to abbreviations: YB -----Yucatan basin; GB = Grenada basin.
exposed in Honduras have not been established. Limited isotopic age dates indicate a late Paleozoic metamorphic event around 300 Ma and show that these rocks are pre-Mesozoic (Gordon, 1991). The dominant basement rock types are phyllitic and graphitic schists which can contain interlayers of metaconglomerate and quartzite. The overlying Mesozoic stratigraphy consists of Middle Jurassic through Early Cretaceous clastic rocks overlain by Aptian-Albian shallow marine limestone (Scott and Finch, Chapter 6). The stratigraphic record indicates that the Chortfs block was a region of minimal tectonic activity during the Mesozoic but was subject to regional folding and faulting events in the Late Cretaceous and Cenozoic (Av6 Lallemant and Gordon, Chapter 8; Manton and Manton, Chapter 9).
Caribbean oceanic plateau The central and eastern Caribbean is underlain by an oceanic plateau with a 12-15 km thickness that is intermediate between oceans and continents over most of its area (Case et al., 1990; Diebold and Driscoll, Chapter 19) (Fig. 2). Caribbean ocean floor made in Cretaceous or earlier times would normally have subsided as it aged to depths of between 5
and 6 km below sea level, and where it is overlain by 2 km of sediments, it might be expected to lie about 1 km deeper. The shallow depth of most of the Caribbean ocean floor, averaging 1-2 km less than predicted, is commonly attributed to the rapid and widespread emplacement of basaltic flows and sills to form an' immense oceanic plateau during Santonian time (~88 Ma) (Burke, 1988; Donnelly et al., 1990; Sinton et al., 1997). Deformation of the Caribbean plate edges has led to exposure of the edges of the Caribbean oceanic plateau in Costa Rica (Sinton et al., 1997), Panama (Bowland and Rosencrantz, 1988), southern Hispaniola (Sen et al., 1988), Colombia, and northern Venezuela (Kerr et al., 1997). The thickness and geochemistry of the Caribbean oceanic plateau is similar to that of western Pacific oceanic plateaus including the Manihiki and Ontong Java (Bowland and Rosencrantz, 1988; Kerr et al., 1997). Diebold and Driscoll (Chapter 19) and Driscoll and Diebold (Chapter 20) show that Caribbean oceanic plateau volcanism was a two-phase event. They suggest that the Santonian age sampled from circum-Caribbean outcrops and from DSDP and ODP cores in the Colombian and Venezuelan basins (Donnelly et al., 1990) dates only the second smaller phase of plateau formation.
6 Great Arc of the Caribbean
Arc rocks of a Cretaceous-Eocene island-arc chain are found in a semi-continuous belt from Cuba to the north coast of South America (Fig. 2). The northern, or Greater Antilles, segment of the arc from Cuba to the Virgin Islands east of Puerto Rico has been inactive since its collision with the Bahamas Platform in Late Paleocene to earliest Oligocene time. Major pulses of collision were broadly diachronous and occurred in Late Paleocene/earliest Eocene time in western Cuba (Bralower and Iturralde-Vinent, 1997; Gordon et al., 1997), Early to Middle Eocene in central Cuba (Hempton and Barros, 1993), Middle Eocene to Recent in Hispaniola (Mann et al., 1991) and Late Eocene to Early Oligocene in Puerto Rico (Dolan et al., 1991). Arc rocks of Early to Late Cretaceous age are found in a continuous belt along the Aves Ridge, the remnant arc produced by rifting of the Grenada back-arc basin, and the Leeward Antilles along the northern coast of South America (Av6 Lallemant, 1997) (Fig. 2). Although the lack of reliable isotopic ages does not allow recognition of di~chroneity in the age of the arc, the arc is adjacent to a west to east-younging fold-thrust belt along the northern margin of South America (Av6 Lallemant, 1997). Because all arc segments initiated during the Early Cretaceous and exhibit lithologic and geochemical similarities, several groups of workers have interpreted circum-Caribbean island-arc rocks as a continuous volcanic arc chain that ringed the Cretaceous Caribbean oceanic plateau (Malfait and Dinkelman, 1972; Pindell and Barrett, 1990) (Fig. 2). This apparently continuous volcanic chain has been called the 'Great Arc of the Caribbean' (Burke, 1988), the 'Mesozoic Caribbean Arc' (Bouysse, 1988), and the 'Proto-Antillean Arc' (Donnelly et al., 1990). In this chapter, I adopt the term 'Great Arc' for three reasons: (1) brevity; (2) the age of the arc in the Greater Antilles ranges into the Cenozoic and therefore is not always restricted to the Mesozoic; and (3) the extent of island-arc rocks related to the arc may extend far beyond the present geographic area of the Greater and Lesser Antilles. Back-arc basins associated with the Great Arc
Paleogene back-arc basins are present along most of the length of the Great Arc of the Caribbean (Fig. 2). Marine heat-flow measurements and depth to basement calculations using marine seismic profiles from the Yucatan basin suggest that it formed during a brief period of northeasterly extension between Paleocene and Early Eocene time (Rosencrantz, 1990). Similar calculations in the Grenada back-arc basin indicate that it formed during the Paleocene hiatus in arc activity along the Lesser
R MANN Antilles island arc (Bouysse, 1988). Bird et al. (Chapter 15) show that the direction of opening of the Grenada basin was approximately east-west rather than north-south as proposed by Pindell and Barrett (1990). Heubeck et al. (1991) proposed that a narrow belt of now-inverted, Paleogene basinal rocks exposed on the island of Hispaniola may represent the continuation of the Grenada and Yucatfin back-arc basins in this area (Fig. 2). Deformation of Caribbean crust and basin formation
Plate tectonics within collages of continental and island-arc lithosphere like the Caribbean is well recognized to be more complicated than that in the oceans because of the existence of many older faults which act as lines of weakness and because silicarich and feldspar-rich rocks of continents and island arcs deform more easily at low temperatures than do oceanic basalts (Fig. 2). These facts explain the broad (~200-250 km) zones of plate-edge seismicity and late Neogene deformation along all margins of the Caribbean plate as well as the diffuse zones of seismicity and active faulting within the Chortfs block suggestive of large-scale, internal plate deformation (Manton, 1987; Gordon and Muehlberger, 1994) (Fig. 1). This complex crustal and active plate setting leads to basin subsidence in response to a variety of subsidence mechanisms as well as basin inversions related to abruptly changing tectonic settings.
PREVIOUS CLASSIFICATIONS AND REGIONAL STUDIES OF CARIBBEAN SEDIMENTARY BASINS
There have been many previous classifications and regional studies of Caribbean sedimentary basins within a plate-tectonic framework. Gonzalez de Juana et al. (1980) carried out a thorough compilation on sedimentary basins in Venezuela for the Venezuelan oil industry. Burke et al. (1984) compiled information on Mesozoic rifts in the Caribbean and Gulf of Mexico related to the breakup of North and South America along with post-Eocene strikeslip basins from the southern and northern margins of the Caribbean. Ladd and Buffler (1985) compiled data on trench and forearc basins of the Middle America arc and Speed and Westbrook (1984) compiled data on forearc and intra-arc basins of the Lesser Antilles arc as part of data syntheses sponsored by the Ocean Drilling Program. Burkart and Self (1985) and Manton (1987) compiled data on active rift basins of the Chortfs block. Eva et al. (1989) compiled information from mainly onland rift, arc and strike-slip basins in Venezuela, Trinidad and the Leeward Antilles. Holcombe et al. (1990)
CARIBBEAN SEDIMENTARY BASINS: CLASSIFICATION AND TECTONIC SETTING and Ladd et al. (1990) compiled information on submarine basins from the plate interior and active plate margins, respectively, as part of the GSA Decade of North American Geology volume on the Caribbean. St6phan et al. (1990) presented fourteen reconstructions of the Caribbean with superimposed paleogeographic information for the period from the Jurassic to the present-day. Dolan et al. (1991) compiled information and presented a tectonic synthesis of onland Paleogene sedimentary basins of Hispaniola and Puerto Rico. Pindell (1995) compiled information on rifts related to North America-South America breakup and along with foreland basins related to the diachronous collision of the Great Arc of the Caribbean and the passive margins of North and South America.
BASIN CLASSIFICATION USED IN THIS OVERVIEW
Fig. 3 summarizes the nomenclature I use in this chapter to classify Caribbean sedimentary basins. Four main types of basins are recognized that are associated with strike-slip, island-arc, collisional and rift environments. Strike-slip basins
Using nomenclature developed by geologists in California and New Zealand, I classify Caribbean strike-slip basins into five basin types based on their bounding fault structure: (1) pull-apart basins produced by extension at a discontinuity or 'step' along a section of a strike-slip fault; (2) fault-wedge basins occurring at intersections of bifurcating strike-slip
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faults; (3) fault-angle depressions parallel to a single strike-slip fault trace; (4) fault-flank depressions between transverse secondary folds or normal faults; and (5) ramp or 'push-down' basins between reverse or thrust faults related to strike-slip movement (Cobbold et al., 1993) (Fig. 3A). All of these late Neogene basin types typically mark zones of strike-sliprelated tectonic subsidence and are found offshore or in topographically low onland depressions or valleys. Areas of most rapid tectonic uplift in both the northern and southern Caribbean region are often localized on restraining bend strike-slip fault segments or 'push-ups' related to shortening at a discontinuity or 'step' along a throughgoing strike-slip fault. Because these bends are usually the sites of rapid, long-term (>5 m.y.) uplift, the bends typically form deeply eroded mountainous areas that are typically structural domes exposing Cretaceous and older basement rocks. Island-arc basins
Island-arc basins include trench-fill basins, forearc basins, intra-arc basins bounded by highs or volcanoes within the volcanic arc, and back-arc basins (Fig. 3B). The most prominent examples of island-arc basins in the Caribbean include back-arc basins of the active Middle America arc (MedianNicaraguan, Mann et al., 1990), the extinct Greater Antilles segment of the Great Arc (Yucat~in basin, Rosencrantz, 1990), and the active Lesser Antilles arc (Grenada basin, Bird et al., Chapter 15). Island-arc basins of the Great Arc are commonly deformed and require careful mapping to delineate their extent and internal facies.
Fig. 3. Basin classification nomenclature used in this chapter to classify Caribbean sedimentary basins shown on Figs. 5-10. (A) Strike-slip basin types. (B) Island-arc basin types. (C) Collisional basin types. (D) Rift basin types.
8 Collisional basins
This basin type includes foreland or foredeep basins, which are by far the most extensive and thickest of all the basin types shown on Fig. 3. In the Caribbean, these basins mark the flexure of the continental or thinned crust of the North and South America plates beneath the overriding thrust sheets of the Great Arc of the Caribbean. Piggyback basins which can also form in non-collisional accretionary prism settings like the Barbados Ridge complex in front of the Lesser Antilles arc (Huyghe et al., Chapter 14) m form and are filled while being carried on moving thrust sheets (Ori and Friend, 1984). Rift basins
This basin type includes full-grabens or rifts and half-grabens or rifts. Full-graben means a graben bounded on both sides by normal faults while halfgraben means a graben bounded only on one side by a normal fault. The full and half types can occur singly or together and be linked by transverse strikeslip faults called transfer faults. The best examples of uninverted rifts in the Caribbean are Jurassic rifts related to the breakup of North and South America found in the southeastern Gulf of Mexico (Marton and Buffler, Chapter 3) and Paleogene rifts related to the early formation of the Cayman trough pull-apart basin (Leroy et al., 1996). Inverted basins
Structural inversion of basins means that the basin-controlling extensional faults reversed their movement because of convergent tectonics and the basin was turned inside out to form a present-day mountain range (Williams et al., 1989). Because of the continuing activity along its margin there are many examples of inverted Caribbean sedimentary basins that include inverted Jurassic rifts in northwestern South America (Lugo and Mann, 1995), inverted intra-arc basins of the Greater Antilles segment of the Great Arc (Mann and Burke, 1990; Dolan et al., 1991), and inverted forearc basins of the Middle America arc (Kolarsky et al., 1995a). Inverted basins provide valuable insights into the early stratigraphic history of basins provided that their sediments can be well dated and their structural overprint can be removed.
R MANN (1997) make an excellent tool for the study and classification of Caribbean sedimentary basins. On these maps, free-air gravity highs marked by the yellow and orange colors correspond to seafloor highs that include active volcanic arcs, remnant volcanic arcs, uplifted oceanic plateau crust, peripheral bulges in flexed Mesozoic oceanic crust of the Atlantic Ocean and Cenozoic crust of the Pacific Ocean and carbonate platforms and isolated banks. Fracture zones in oceanic crust are expressed as fine lineaments traceable over distances up to several hundred kilometers. Trenches at subduction zones, sedimentary accretionary wedges, and major sedimentary basins of the types shown on Fig. 3 are marked by large free-air gravity lows. I have divided the Caribbean gravity data set into six sub-areas that allow better resolution of basins in the individual areas (Fig. 4). These areas include: (1) basins associated with the Middle America trench, arc, and back-arc in the western Caribbean (Central America) (Fig. 5); (2) basins associated with the North America-Caribbean plate boundary in the northern Caribbean (Fig. 6); (3) basins associated with the Lesser Antilles trench, arc, and back-arc in the eastern Caribbean (Fig. 7); (4) basins associated with the South America-Caribbean plate boundary in the southern Caribbean (Fig. 8); (5) basins associated with the Panama arc-South America collisional zone in the southwestern Caribbean (Panama and Costa Rica) (Fig. 9); and (6) basins associated with the central Caribbean plate (Nicaraguan Rise, Colombian basin, and Venezuelan basin) (Fig. 10). Because free-air gravity is a close approximation of seafloor bathymetry, only those submarine Cenozoic basins with prominent morphologic expression can be distinguished on these maps. Older deformed basins from previous tectonic phases might be expressed as a gravity high or intermediate gravity value. For this reason, the offshore basins classified in this study are generally Cenozoic basins generated during the more recent phases of Caribbean strike-slip and subduction tectonics. For this reason, these maps should not be considered as complete compilations of all Caribbean sedimentary basins.
BASINS AND MAJOR TECTONIC FEATURES ASSOCIATED WITH THE MIDDLE AMERICA TRENCH, ARC, AND BACK-ARC (WESTERN CARIBBEAN) Rifts associated with the f r a g m e n t a t i o n of the western Chortis block
USE OF GRAVITY MAPS TO ILLUSTRATE CARIBBEAN SEDIMENTARY BASINS
Gridded, 2-min, satellite-derived free-air gravity data compiled and described by Sandwell and Smith
Seven approximately north-south-striking Neogene rifts and half-rifts are present in the northwestern corner of the Caribbean plate (Chortfs block) between the Median back-arc basin and the North
CARIBBEAN SEDIMENTARY BASINS: CLASSIFICATION AND TECTONIC SETTING
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Fig. 4. Map showing locations of regional gravity maps of Sandwell and Smith (1997) used in this chapter in Figs. 5-10 to illustrate regional tectonic features. America-Caribbean strike-slip zone in Guatemala, Honduras and E1 Salvador (Nos. 1-7 in Fig. 5). Little is known about of the exact age for the onset of rifting and the thickness of the sedimentary fill of these rifts. Several authors have attributed the transverse nature of the faulting to internal deformation of the Caribbean plate as it moves eastward relative to the North America Plate along concave southward, left-lateral strike-slip faults of the MotaguaPolochfc system (No. 8 in Fig. 5) (Plafker, 1976; Burkart and Self, 1985).
Inverted forearc basin rocks adjacent to the subducting Cocos Ridge Shallow subduction of the Cocos Ridge in Late Miocene to Recent time has inverted a marine forearc basin of Oligocene and Miocene age between the arc and trench (Kolarsky et al., 1995a) (No. 13) along with trench-slope facies on the Pacific peninsulas of Panama and Costa Rica (Corrigan et al., 1990; Collins et al., 1995) (No. 14). This inverted basin is collinear with the undisturbed, offshore Sandino forearc basin to the north along the margin of Nicaragua and E1 Salvador (No. 16).
Median-Nicaraguan back-arc basin This late Neogene back-arc basin forms a prominent, 800-km-long structural depression parallel to the Middle America volcanic arc and trench (No. 9). This basin is most prominent in Nicaragua where it is occupied by two large lakes (Managua and Nicaragua) (No. 10). Late Quaternary arc volcanoes occur at the edges, in the center and adjacent to the basin. Back-arc basin sedimentary rocks of Oligocene to Neogene age are inverted along the southeastern extension of the back-arc basin in Costa Rica (No. 11). This localized back-arc basin inversion is Late Miocene to Recent in age and is related to the shallow subduction of the Cocos Ridge (Kolarsky et al., 1995a) (No. 15). The Median back-arc basin to the north (No. 12) is a less distinctive and linear basinal feature than the Nicaraguan basin to the south.
Cocos Ridge The Cocos Ridge (No. 15) is a hotspot trace of the Galapagos hotspot that stands 2 to 2.5 km higher than the surrounding seafloor and is presently subducting beneath the southern Middle America trench (Kolarsky et al., 1995a). Collins et al. (1995) propose on the basis of detailed biostratigraphic work that the Cocos Ridge contacted the Middle AmeriCa trench about 3.6 Ma and inverted the back-arc area of Costa Rica by 1.6 Ma.
Panama fracture zone This fault (No. 17) is a right-lateral transform fault that separates oceanic crust of the Cocos and Nazca plates.
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P. M A N N
Fig. 5. Geosat free-air gravity map of marine areas associated with the Middle America trench, arc and back-arc in the western Caribbean (Central America). Gravity highs are shown by darker colors and lows are shown by lighter colors. Numbers identify Cenozoic basins that are described in the text.
Fig. 6. Geosat free-air gravity map of marine areas associated with the North America-Caribbean plate boundary in the northern Caribbean. Gravity highs are shown by darker colors and lows are shown by lighter colors. Numbers identify Cenozoic basins that are described in the text.
C A R I B B E A N S E D I M E N T A R Y BASINS" C L A S S I F I C A T I O N A N D T E C T O N I C S E T T I N G
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Fig. 7. Geosat free-air gravity map of marine areas associated with the Lesser Antilles trench, arc and back-arc in the eastern Caribbean. Gravity highs are shown by darker colors and lows are shown by lighter colors. Numbers identify Cenozoic basins that are described in the text.
Fig. 8. Geosat free-air gravity map of marine areas associated with the South America-Caribbean plate boundary in the southern Caribbean. Gravity highs are shown by darker colors and lows are shown by lighter colors. Numbers identify Cenozoic basins that are described in the text.
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E MANN north from rough Cocos lithosphere created at the Galapagos rift system to the south (Protti et al., 1995). Forearc morphology adjacent to the Galapagos seafloor and the Cocos Ridge is tectonically eroded by the higher standing and rougher seafloor southeast of the rough-smooth boundary (von Huene et al., 1995; Kolarsky et al., 1995a).
BASINS AND MAJOR TECTONIC FEATURES ASSOCIATED WITH THE NORTH AMERICA-CARIBBEAN PLATE BOUNDARY (NORTHERN CARIBBEAN) North America-Caribbean foreland basin and active plate boundary in Central America, Cuba, Hispaniola, and the Puerto Rico trench
Fig. 9. Geosat free-air gravity map of marine areas associated with the Panama arc-South America collisional zone in the southwestern western Caribbean (Panama and Costa Rica). Gravity highs are shown by darker colors and lows are shown by lighter colors. Numbers identify Cenozoic basins that are described in the text.
Rough-smooth boundary of the Cocos plate This boundary (No. 18) separates a smooth Cocos lithosphere created at the East Pacific Rise to the
A semi-continuous foreland basin recording the collision between the Great Arc of the Caribbean and the passive margin of North America can be traced from the Sepur foreland basin of northern Central America (No. 1 in Fig. 6), along the eastern edge of the Yucatan Peninsula (Rosencrantz, 1990; Lara, 1993; No. 2), along the northern (Denny et al., 1994; No. 3) and northeastern coasts of Cuba (Ball et al., 1985; No. 4), along the northwestern (Dillon et al., 1992; No. 5) and northeastern (Dolan et al., 1998; No. 6) coasts of Hispaniola, and in the Puerto Rico trench (Masson and Scanlon, 1991; Grindlay et al., 1997; No. 7). The age of the foreland basin is diachronous with Late Cretaceous thrust-related subsidence in northern Central America (Rosenfeld, 1990), Paleocene-Early Eocene subsidence in western Cuba (Bralower and Iturralde-Vinent, 1997), Early to Middle Eocene in central Cuba (Hempton and Barros, 1993), Middle Eocene to Recent in Hispaniola (Mann et al., 1991) and Late Eocene to Early Oligocene in Puerto Rico (Dolan et al., 1991).
Fig. 10. Geosat free-air gravity map of marine areas associated with the central Caribbean plate (Nicaraguan Rise, Colombian basin, Beata Ridge, Venezuelan basin). Gravity highs are shown by darker colors and lows are shown by lighter colors. Numbers identify Cenozoic basins that are described in the text.
CARIBBEAN SEDIMENTARY BASINS: CLASSIFICATION AND TECTONIC SETTING
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Yucat~in back-arc basin
Muertos trough and 'forearc' basin
The Yucatfin basin formed in Paleocene time behind the Cuban segment of the Great Arc of the Caribbean as it moved to the north and northeast prior to its collision with the Bahama Platform (Rosencrantz, 1990). The basin exhibits three sub-basins. The West Yucatan basin (No. 8) is an oceanic-floored pull-apart basin formed along leftlateral faults bounding the Yucatan Peninsula. The Central Yucatan basin (No. 9) and Cayman Rise (No. 10) are basins formed on stretched arc or continental crust thinned in a back-arc setting. The Cayman Ridge (No. 11) south of the Cayman Rise could be considered a remnant arc although it has been strongly overprinted by strike-slip faulting related to the active plate boundary in the Cayman trough (No. 12).
The Muertos trench (No. 18) accommodates northward underthrusting of the Caribbean oceanic plateau of the Venezuelan basin beneath Hispaniola (Ladd et al., 1990; Dolan et al., 1998). A forearctype basin has formed on the overriding plate south of Hispaniola (Ladd et al., 1990; No. 19) but is not associated with a volcanic arc probably because the angle of subduction of the Caribbean plate is too low to generate wellling.
Cayman trough The Cayman trough (No. 12) is an 1100-km-long pull-apart basin that began its protracted history of oceanic spreading at a 100-kin-long spreading ridge during the Early Eocene (Rosencrantz et al., 1988). The eastern end of the trough (No. 14) is marked by half-grabens of Paleocene-Eocene age (Leroy et al., 1996) that may be coeval with an inverted PaleoceneEocene graben in Jamaica (Mann and Burke, 1990).
Basins and inverted basins associated with the southern edge of the Cayman trough Late Neogene basin formation, restraining bend uplifts, and inverted Paleogene rifts in this area are linked to left-lateral strike-slip movements along the Enriquillo-Plantain Garden fault zone (No. 13). Some basins like the Tela of northern Honduras (No. 14) (Av6 Lallemant and Gordon, Chapter 8; Manton and Manton, Chapter 9) are not clearly linked to a specific fault zone and instead appear to be part of broad structural borderland within the broad strike-slip plate boundary zone.
Basins associated with the Anegada fault zone This fault (No. 20) has been interpreted by Jany et al. (1990) as an active right-lateral fault bounding the eastern edge of a Puerto Rico-Hispaniola microplate (Jany et al., 1990). Two right-steps along the fault are interpreted as pull-apart basins that formed by right-lateral motion in Late Miocene to Recent time. Gill et al. (Chapter 13) propose that motion along the Anegada fault zone is left-lateral, rather than right-lateral on the basis of detailed stratigraphic studies on St. Croix (U.S. Virgin Islands) to the south of the fault.
BASINS AND MAJOR TECTONIC FEATURES ASSOCIATED WITH THE LESSER ANTILLES TRENCH, ARC, AND BACK-ARC(EASTERN CARIBBEAN) Aves Ridge remnant arc The Aves Ridge (No. 1 in Fig. 7) is a remnant arc formed when Paleogene east-west opening of the Grenada back-arc basin separated the ridge from the Lesser Antilles arc (Bouysse, 1988; Bird et al., Chapter 15). Dredge hauls and marine geophysics indicate that the ridge formed part of the Late Cretaceous Great Arc that ceased activity by the time of back-arc opening of the Grenada basin in Early Paleogene time.
Grenada back-arc basin Convergent strike-slip basins of the Hispaniola restraining bend Convergence of the eastward-moving Caribbean plate relative to the southeastern extension of the Bahama Platform has led to localized convergence and topographic uplift in Hispaniola (Mann et al., 1995). Three late Neogene basins in Hispaniola (Nos. 15, 16, 17) are the thrust-bounded ramp type (Mann et al., Chapter 12) (Fig. 3A).
The Grenada back-arc basin (No. 2) with an average water depth of 2-3 km, contains 2 km (north) to 9 km (south) of Cenozoic sediment derived from both the erosion of South America and the Lesser Antilles arc (Bouysse, 1988). Opening of the basin was in an east-west direction (Bird et al., 1993). The gravity low of the Grenada basin can be traced along much of the margin of northern South America where it is oriented east-west, is narrower, and parallel to collisional structures of the margin.
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Lesser Antilles volcanic arc The Lesser Antilles volcanic arc (No. 3) initiated in Early Cretaceous and has remained active to the present (Speed and Westbrook, 1984; Bouysse, 1988). Its position at the leading edge of the eastward-moving Caribbean plate means that it may be far-traveled and probably originated somewhere in the eastern Pacific Ocean (Pindell and Barrett, 1990). Underthrusting of Jurassic-Cretaceous age oceanic crust of the Atlantic Ocean beneath the Lesser Antilles results in line of active calc-alkaline volcanoes forming the volcanic arc. The gravity high of the volcanic arc can be traced to the southwest along the northern margin of South America and to the northeast through the Virgin Islands and Puerto Rico.
Kallinago basin This basin (No. 4) forms an intra-arc basin with the northern Lesser Antilles volcanic arc and formed in the Late Miocene by the westward migration of the volcanic line from the islands on its eastern flank (Limestone Caribbees). The Kallinago basin is not a typical rift basin related to back-arc spreading because the volcanic line jumped westward away from the trench and not toward the trench as found in most back-arc basins. McCann and Pennington (1990) attributed this unusual behavior related to the subduction of the Barracuda fracture zone ridge (No. 9) which lowered the angle of subduction and caused the volcanic arc to migrate westward.
Tobago trough This 10-km-thick, Miocene to Recent basin (No. 5) is bounded on its eastern edge by back-thrusts within the accretionary wedge of the Lesser Antilles arc (Barbados Ridge complex, No. 7) (Speed et al., 1989). The Tobago trough extends to the west along the northern margin of South America.
Lithospheric trace of the Lesser Antilles subduction zone The contact or lithospheric trace between crystalline rocks of the Lesser Antilles arc and downgoing oceanic crust of the Atlantic Ocean (No. 6) lies at a depth of about 20 km north of Trinidad but is marked by the line of demarcation between faint northeast gravity trends on strike with Atlantic fracture zone trends and prominent north-south trends of the Lesser Antilles arc. The island of Tobago with a Cretaceous to Eocene record of arc activity is located just to the west of this contact and is therefore the most eastward outcrop of arc rocks of the
MANN Caribbean plate. The north-south lithospheric trace and the east-west E1 Pilar fault zone of Trinidad and northern Venezuela (No. 13) can be seen to form a continuous and curving lineament.
Barbados Ridge accretionary complex and deformation front This accretionary wedge (No. 7) between 100 and 300 km wide and from 0.5 (toe of slope) to 20 km (above lithospheric trace) thick consists of the mainly clastic fluvial and pelagic sediments offscraped from the downgoing Atlantic ocean floor (Ladd et al., 1990). Southward widening of the complex reflects the fluvial addition of material from the Orinoco delta area south of Trinidad (No. 6). The deformation front is marked by the most eastward thrust fault juxtaposing the Barbados Ridge accretionary complex with undeformed seafloor of the Atlantic Ocean. This front is roughly east-west and irregular in the area to the east of Trinidad. Piggyback basins (Fig. 3C) and shale diapirs derived from muds in the prodelta area of the Orinoco River are common in this area (Huyghe et al., Chapter 14).
BASINS AND MAJOR TECTONIC FEATURES ASSOCIATED WITH THE SOUTH AMERICA-CARIBBEAN PLATE BOUNDARY (SOUTHERN CARIBBEAN)
The Aves Ridge (No. 1 in Fig. 8), the Grenada back-arc basin (No. 2), the Lesser Antilles volcanic arc (No. 3), and the Tobago trough (No. 4), which extend into this area, are described above and are also shown on Fig. 7.
Guyana passive margin of South America This Cretaceous-Recent margin (No. 5 in Fig. 8) formed by rifting and strike-slip of the Africa plate past the South America Plate in earliest Cretaceous time (Pindell and Barrett, 1990). The margin projects into a buried passive margin buried beneath the Gulf of Paria west of Trinidad and may control the location of strike-slip faults in this area (Babb and Mann, Chapter 18).
Orinoco delta The Orinoco River drains a large area of the northeastern South American continent and forms one of the major shelf-margin deltas in the world (No. 6).
CARIBBEAN SEDIMENTARY BASINS: CLASSIFICATION AND TECTONIC SETTING Eastern Venezuelan and Maracaibo basins
This basin is a major foreland basin marked by a 150 mGal gravity low formed by oblique convergence between the South American continent (Guyana Shield) and the Caribbean arc system in Oligocene and Miocene time (di Croce et al., Chapter 16; Flinch et al., Chapter 17). The basin is subdivided into two sub-basins which increase in age from east (Maturfn, Oligocene to Recent, No. 7) to west (Gu~irico, Eocene to Pliocene, No. 8). The western extension of the foreland basins is represented by the Maracaibo basin (Late PaleoceneEocene, No. 9) (Lugo and Mann, 1995). The change in strike of the older, western part of the foreland basin (Maracaibo, No. 9) is related to late Neogene northward displacement of the Maracaibo block along the fight-lateral Bocon6 fault (No. 13) and the left-lateral Santa Marta-Bucaramanga fault (No. 14).
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BASINS AND MAJOR TECTONIC FEATURES ASSOCIATED WITH THE PANAMA ARC-SOUTHWESTERN CARIBBEAN COLLISIONAL ZONE (SOUTHWESTERNCARIBBEAN) C6baco basin complex
This submarine basin complex in a shelf setting (No. 1 in Fig. 9) is an active pull-apart formed at a left-step in the Azuero-Son~i fault zone of western Panama (Kolarsky et al., 1995b). Tonosi basin
This now folded and uplifted Oligocene-Miocene turbiditic basin (No. 2) appears to have been a forearc basin that has now become inverted as a result of strike-slip movements along the obliqueslip margin of southwestern Panama (Kolarsky et al., 1995b). Riffs of the Canal area
Accreted rocks of the South America passive margin and Great Arc of the Caribbean
These fold-thrust belts (No. 10) contain mixtures of arc and margin-related lithologies formed during the oblique Cenozoic collision of the arc and the passive margin (Av6 Lallemant, 1997). Cariaco pull-apart basin on the El Pilar fault zone
The Cariaco pull-apart basin (No. l l) is a 1400-m-deep, closed depression in the Venezuelan shelf at a 35-km south-step between the right-lateral E1 Mor6n and E1 Pilar strike-slip faults. Schubert (1984) estimated that 70 km of right-lateral motion on the faults was necessary to produce the basin over the last 2 million years. South Caribbean marginal fault
This fault (No. 12) forms the deformation front of a large accretionary wedge formed by the underthrusting of the Caribbean oceanic plateau and its overlying sedimentary cover beneath the South American continent (Ladd et al., 1990).
These late Neogene basins (No. 3) formed as a consequence of diffuse east-west extension within this topographically lowest part of the Panama Isthmus. These basins may have formed as a response to bending of the isthmus of Panama following its collision with northwestern South America in Late Miocene to Early Pliocene time (Mann and Kolarsky, 1995). San Bias 'forearc' basin
This basin (No. 4) has formed in a 'forearc' setting above the accreted North Panama deformed belt in late Neogene times (Reed and Silver, 1995). Bayano-Chucanqu6 basin
This basin (No. 5) has formed in a large syncline formed in response to the bending and strike-slip deformation of the Isthmus of Panama following its collision with the South America margin (Mann and Kolarsky, 1995). Sambfi basin
This basin (No. 6) appears to be a pull-apart basin formed at a left-step in the left-lateral Samb6 fault zone.
Maracaibo block Pearl Islands basin
This triangular-shaped block of continental crust is bounded to the east by the Bocon6 right-lateral fault (No. 13) and to the east by the left-lateral Santa Marta-Bucaramanga fault (Mann et al., 1990) (No. 14).
This Middle Miocene-Pleistocene basin (No. 7) formed as a small foreland basin in front of east-dipping reverse faults of the East Panama deformed belt (Mann and Kolarsky, 1995).
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Colombian accretionary complex and forearc basin
This margin (No. 8) developed in response to eastward subduction of oceanic crust of the Nazca plate beneath Colombia (Westbrook et al., 1995). Atrato-San Juan basin
This basin (No. 9) formed along the approximate suture zone between the Panama arc and the South American continent (Bueno Salazar, 1989; Kellogg and Vega, 1995).
MANN Beata Ridge
The Beata Ridge (No. 4) is marked by a triangular-shaped uplift of oceanic plateau crust at the place where the Caribbean sea is narrowest, between northern South America and Hispaniola. Driscoll and Diebold (Chapter 20) interpret the uplift as a mainly relict extensional fault block related to the formation of the Cretaceous oceanic plateau while Mauffret and Leroy (Chapter 21) emphasize its late Neogene uplift history on thrust faults and its role as major tectonic boundary between the Colombian and Venezuelan basins. Venezuelan basin
BASINS AND MAJOR TECTONIC FEATURES ASSOCIATED WITH THE CENTRAL CARIBBEAN PLATE Basins and carbonate banks of the Nicaraguan Rise
The Nicaraguan Rise is a broad submarine swell underlain by island arc and continental crust of the Chortfs block that extends from northern Central America to Jamaica and is bounded on the north by the Cayman trough and on the south by the Hess Escarpment (No. 1 in Fig. 10). Carbonate banks of Cenozoic age occupy structural highs formed by poorly understood faults with a predominantly northeast strike. Holcombe et al. (1990) interpreted the northeast faults as a set of left-lateral strike-slip faults bounding a set or more northward-striking rift basins associated in one case (San Andres trough) with Quaternary basaltic volcanism present on the island of San Andres (No. 2). The linear Hess Escarpment (No. 3) has been interpreted by Mann et al. (1990) as a possible Neogene strike-slip feature whereas others like Driscoll and Diebold (Chapter 20) have interpreted it as Cretaceous normal fault linked to the Beata Ridge (No. 4) and to the formation of the Caribbean oceanic plateau (Fig. 2). Colombian basin
The low-relief Colombian basin (No. 5 in Fig. 10) is underlain by the Cretaceous Caribbean oceanic plateau (Bowland and Rosencrantz, 1988) and is bounded to the north by the Hess Escarpment, to the south by the South Caribbean margin fault (Ladd et al., 1990; No. 6), and to the east by the Beata Ridge (No. 4). van der Hilst and Mann (1994) used tomographic data to show that the oceanic plateau crust of the Colombian basin is underthrust at the South Caribbean front fault (No. 6) by a distance of several hundred kilometers beneath the South America margin.
The low-relief Venezuelan basin (No. 7) is a rectangular area of oceanic plateau crust bounded on the north by the Muertos trench, on the south by the South Caribbean marginal fault, on the west by the Beata Ridge, and on the east by the Aves Ridge, the remnant volcanic arc of the Lesser Antilles. A prominent east-west arch trends parallel to the long axis of the basin and may be a regional flexure of the Caribbean plate produced by its ongoing subduction at the Muertos trough to the north and the South Caribbean marginal fault (No. 6) to the south.
TECTONIC EVOLUTION OF THE CARIBBEANPLATE AND ITS SEDIMENTARYBASINS Two models for Caribbean evolution
There are two contrasting models for the platetectonic evolution of the Caribbean. The first model most recently put forward by Frisch et al. (1992) proposes that the Caribbean region formed during the period of 130 Ma to 80 Ma as South America moved southeast away from North America (Fig. l lA). Igneous upwelling in the space that formed between the two continents is thought to have produced the anomalously thick oceanic plateau crust of the Caribbean and Central America (Kerr et al., 1997; Diebold and Driscoll, Chapter 19; Driscoll and Diebold, Chapter 20). This model recognizes some strike-slip motion along the northern and southern margins of the plate but does not view these offsets as large enough to restore the Caribbean to a position in the eastern Pacific. An alternative school of thought and adopted in the reconstructions shown in this review was first formulated by Wilson (1966) and later elaborated by Malfait and Dinkelman (1972), Ross and Scotese (1988), Pindell and Barrett (1990), and others. This mobilistic view is that the Caribbean was originally an area of eastern Pacific Ocean floor and oceanic
CARIBBEAN SEDIMENTARY BASINS: CLASSIFICATION AND TECTONIC SETTING
17
Fig. 11. Two possible origins for the Caribbean. (A) The Caribbean oceanic plateau forms by the separation of North and South America during the period 130 to 80 Ma (Frisch et al., 1992). Crosses indicate areas of continental crust. The numbers give positions of the northern margin of South America according to Pindell and Barrett (1990). (B) The Caribbean oceanic plateau forms as normal Pacific oceanic crust drifts over the Galapagos hotspot, is thickened in the middle and Late Cretaceous, and passes into the gap between North and South America. The numbers give positions of the leading edge of the Caribbean arc system and oceanic plateau through time, according to Pindell and Barrett (1990). Note that the positions of continental masses allow 'no free face' for arc migration. The 'free face' of the Atlantic Ocean acts to channel the arc in an eastward direction.
plateau that has been rafted behind the eastwardmoving Great Arc of the Caribbean of Burke (1988) (Fig. 11B). This area of Pacific normal ocean crust appears to have been modified and thickened into the present-day Caribbean oceanic plateau province in the Cretaceous w h e n the crust drifted over the Galapagos hotspot (Duncan and Hargraves, 1984;
Sinton et al., 1997) (Fig. l i B ) . The passage of this area of crust from a Pacific realm to an Atlantic one is recorded by the diachronous history of collisions b e t w e e n the Great Arc at the leading edge of the plateau and the passive margins of North and South A m e r i c a (Pindell and Barrett, 1990). These collisions c o m m e n c e in Late Creta-
18 ceous time in northern Central America and northwestern South America and continue through to the present-day in the northeastern and southeastern Caribbean (Fig. 11B). Using North A m e r i c a - S o u t h America motion to infer Caribbean tectonics
Because magnetic anomalies and fracture zones nearly as old as the times of separation between North and South America have been mapped in the Atlantic Ocean, the motions of North and South America with respect to Africa can be fully described using the vectorial closure condition required by a three-plate system (Pindell and Barrett, 1990; MUller et al., Chapter 2). Improved maps of Atlantic fracture zones using Geosat gravity data by Mtfller et al. (Chapter 2) has allowed them to more precisely reconstruct the motion history of the two Americas. A summary of the Jurassic to recent path of two points on northern South America relative to a fixed North America using the data of MUller et al. (Chapter 2) is shown on Fig. 12. This diagram illustrates the steadily widening space between the two Americas that was presumably filled by oceanic crust of Jurassic and Early Cretaceous age. This expanse of crust known as the proto-Caribbean Ocean is thought to have been consumed by the Great Arc of the Caribbean from the Late Cretaceous to Recent time. Remnants of the proto-Caribbean Ocean are preserved only as small fragments within rocks of the Great Arc (Montgomery and Pessagno, Chapter 10). The vector diagram in Fig. 12 can be used to make general inferences about the regional deformational style of the intervening Caribbean plate. For example, from the Late Jurassic to Maastrichtian, one would expect a generally divergent tectonic style to pervade much of this region and from Maastrichtian to the Present one would expect a convergent or strike-slip style to be present (Fig. 12). However, this direct dependence of Caribbean deformational style on the relative motion of North and South America assumes that motion is taken up along a single plate boundary, such as during Jurassic separation. Studies including Marton and Buffler (Chapter 3) show that even the young North America-South America Plate boundary during Jurassic time was multi-branched and involved the motion of an intervening microcontinent, the Yucat~in block. As the gap between the Americas widened, later Cretaceous-Cenozoic relative plate motions acted across at least two plate boundaries (northern and southern Caribbean arc or strike-slip boundaries). Therefore, the North America-South America motions shown on Fig. 12 are only indirectly manifested in Caribbean deformation.
R MANN Relative motion path of South America relative to North America
The path shows South America moving away from North America during the Jurassic through the Late Cretaceous, a process that led to the formation of ocean floor on the sites of the Caribbean and Gulf of Mexico. The orientation of Mesozoic graben is generally perpendicular to this direction except in regions affected by the independent rotation of the Yucatan block (Marton and Buffler, Chapter 3). While the age of most circum-Caribbean rifts is confined to the Jurassic, the path shows continued separation of the Americas up through the Maastrichtian. Numerous geological studies such as those by Pessagno et al. (Chapter 5), Marton and Buffler (Chapter 3), Masaferro and Eberli (Chapter 7), Scott and Finch (Chapter 6), and di Croce et al. (Chapter 16) show that the Cretaceous was a time of carbonate passive margin formation atop these early rift structures. These bank margins probably fronted large expanses of Jurassic and Cretaceous ocean crust that formed following the separation of the two plates in Late Jurassic-earliest Cretaceous time. The behavior of the Great Arc of the Caribbean in response to the convergence or strike-slip motion of North and South America during the period of 71 Ma to the present-day cannot be predicted with accuracy given that these larger plate motions will only be indirectly manifested across multiple strikeslip and subduction Caribbean boundaries. Mann et al. (1995) and Gordon et al. (1997) propose that the trend and direction of the Great Arc during the Late Cretaceous and Cenozoic is governed mainly the direction of a free face, or area of subductable oceanic crust in front of the moving arc. For example, the presence of the Bahamas Platform led to arc collision and reorientation of the arc to subduct Atlantic oceanic crust in a more eastward direction. MUller et al. (Chapter 2) propose that 200-300 km of post-Early Miocene north-south convergence across the Caribbean plate may have led to significant underthrusting of the Caribbean plate beneath North and South America and may have modified existing sedimentary basins.
MAIN PHASES OF CARIBBEAN BASIN DEVELOPMENT WITHIN THE FRAMEWORK OF NORTH AMERICA-SOUTH AMERICA RELATIVE MOTION HISTORY
Using the second plate-tectonic model shown in Fig. liB, five main phases of basin evolution can be predicted for the margins of the North and South America plates. These phases include pre-rift phase, Late Jurassic rift phase, Cretaceous passive
CARIBBEAN SEDIMENTARY BASINS: CLASSIFICATION AND TECTONIC SETTING
19
Fig. 12. Relative plate motion vectors of three points of northern South America with respect to a fixed North America based on data presented by M~iller et al. in Chapter 2 of this volume. The position of South America with respect to North America provides a framework in which to base key events in Caribbean evolution such as the entry and diachronous collision of the Great Arc of the Caribbean. Points in millions of years along the vectors correspond to the ages of plate reconstructions give in Figs. 13-25 of this chapter.
margin phase, Late Cretaceous-Recent arc-passive margin collisional phase, and late Cenozoic strikeslip phase. I subdivide thirteen plate reconstructions of the Caribbean based on the plate parameters of Miiller et al. (Chapter 2) into these four phases. Several Cretaceous and Cenozoic structural phases characterize the overriding Great Arc and the adjacent oceanic plateau and Chortfs block. For example, the oceanic plateau undergoes a two-phase Cretaceous volcanic and stretching event (Driscoll and Diebold, Chapter 20), the Great Arc may have experienced a subduction polarity reversal (L6bron and Perfit, 1994; Draper et al., 1996; Kerr et al., 1997; Montgomery and Pessagno, Chapter 10) and the Chortfs block experienced a Late Cretaceous folding and faulting event (Scott and Finch, Chapter 6).
Pre-rift phase Plate reconstructions such as those by Pindell and Barrett (1990), Marton and Buffler (Chapter 3) and Pszcz6tkowski (Chapter 4) leave no space for the Caribbean-Gulf of Mexico region in its present position when South America is closed up against North America to reform western Pangea in pre-Late Jurassic time (Fig. 13). Prior to rifting of the Americas in the Middle Jurassic, three crustal age provinces are present in the future area of rifting between North and
South America shown in Fig. 13. These provinces include: (1) Pan-African crustal age province of Africa and Brazil; (2) Grenville crustal age province of North and South America that includes a possible continuation through the Oaxaca area of southern Mexico and the Chortfs block (Renne et al., 1989; Hutson et al., 1998); and (3) Guyana Shield of pre-Grenville age (> 1.2 Ga) of northern South America. The Yucat~in block fills the central part of the Gulf of Mexico and is presumably underlain by a prong of the Appalachian-Marathon-Ouachita orogenic belt (Marton and Buffler, Chapter 3).
Late Jurassic rift phase Rifts of Late Jurassic age in the Caribbean form part of a band of rifts associated with the early opening of the central and northern Atlantic that crudely follow orogenic grains from the North Atlantic to Guyana. By Oxfordian time, rifts are active along the northern Gulf of Mexico, the southeastern Gulf of Mexico (Marton and Buffler, Chapter 3), the Bahama Platform (Masaferro and Eberli, Chapter 7), the northeastern margin of South America (di Croce et al., Chapter 16), and the northwestern margin of South America (Eva et al., 1989; Lugo and Mann, 1995) (Fig. 14). Rifts which extend southward along the western margin of South America may be related
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P. M A N N
Fig. 13. Reconstruction of the Caribbean region at 180 Ma (Bajocian). Key to abbreviations: M S M = Mohave-Sonora megashear; TMVB -- Trans-Mexican volcanic belt; E A F Z -- eastern Andean fault zone.
Fig. 14. Reconstruction of the Caribbean region at 156 Ma (Oxfordian, magnetic anomaly M29). Gray areas represent oceanic crust of normal thickness. Stippled areas indicate rifted areas. Key to abbreviations: P C - proto-Caribbean oceanic crust (dark line represents speculative position of spreading ridge)" N B F Z -- northern Bahamas fracture zone; M S M -- Mohave-Sonora megashear; T M V B = Trans-Mexican volcanic belt; E A F Z = eastern Andean fault zone.
CARIBBEAN SEDIMENTARY BASINS: CLASSIFICATION AND TECTONIC SETTING
21
Fig. 15. Reconstruction of the Caribbean region at 145 Ma (Tithonian, magnetic anomaly M19). Dots in Atlantic Ocean represent magnetic anomaly and fracture zone picks by Mt~ller et al. (Chapter 2) based on interpretation of Geosat gravity images. Key to abbreviations: PC = proto-Caribbean oceanic crust (dark line represents speculative position of spreading ridge); M S M = MohaveSonora megashear; TMVB = Trans-Mexican volcanic belt; EAFZ = eastern Andean fault zone. to back-arc rifting produced by subduction at that margin. The widening gap between North and South America was presumably occupied by oceanic crust generated at a proto-Caribbean spreading ridge. This early oceanic corridor between the Atlantic and Pacific widens through continued rifting and oceanic spreading into the Tithonian (Fig. 15).
the Chortfs block occupied a southern extension of Precambrian and Paleozoic orogenic belts in Mexico and was subsequently displaced eastwards by strike-slip faults in the Cenozoic (Av6 Lallemant and Gordon, Chapter 8; Manton and Manton, Chapter 9) (Fig. 16).
Cretaceous passive margin phase
Late Cretaceous-Recent arc-passive margin collisional phase
By earliest Cretaceous time, rifting had ceased, the Yucatan block had rotated to its present-day position, and a post-rift passive margin section composed mainly of carbonate rocks had blanketed the rift topography in the southeastern Gulf of Mexico (Marton and Buffler, Chapter 3), the Bahamas Platform (Masaferro and Eberli, Chapter 7), the northeastern margin of South America (di Croce et al., Chapter 16; Babb and Mann, Chapter 18), and the northwestern margin of South America (Lugo and Mann, 1995). These passive margins enjoyed open ocean circulation and probably fronted a protoCaribbean oceanic basin that was several hundred kilometers wide (Fig. 16). It is interesting to note that the Chortfs block experienced a similar rift and passive margin history to the above intra-Caribbean margins (Scott and Finch, Chapter 6). Presumably
By Late Cretaceous time, the Great Arc of the Caribbean and its adjacent oceanic plateau province was colliding with the passive margin of northwestern South America and the southern margin of northern Central America (Pindell and Barrett, 1990; Kerr et al., 1997) (Fig. 17). In northwestern South America, extensive areas of the plateau and arc rocks accreted to the continental cratonic rocks of northwestern South America (Kerr et al., 1997). In these areas, ages of rocks of the Great Arc extend back to the Early Cretaceous but ages of the oceanic plateau are generally confined to the Santonian (Kerr et al., 1997; Sinton et al., 1997). Diebold and Driscoll (Chapter 19) and Driscoll and Diebold (Chapter 20) present evidence that the oceanic plateau eruption event was a two-phase event with the Santonian event probably corresponding to the younger event.
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P. M A N N
Fig. 16. Reconstruction of the Caribbean region at 118 Ma (Aptian, magnetic anomaly C34n). Key to abbreviations: M S M = MohaveSonora megashear; TMVB = Trans-Mexican volcanic belt; EAFZ -- eastern Andean fault zone.
Fig. 17. Reconstruction of the Caribbean region at 83 Ma (Campanian, magnetic anomaly C34n). Key to abbreviation: EAFZ = eastern Andean fault zone.
CARIBBEAN SEDIMENTARY BASINS: CLASSIFICATION AND TECTONIC SETTING By Maastrichtian time, subsidence of the Sepur foreland basin was ending as the Chortfs block was sutured to the area of southern Mexico and the Yucatan block (Fig. 18) and the Great Arc was migrating to the northeast towards its eventual collision with the Bahama Platform in the Late Paleocene and Eocene (Fig. 19). In Paleocene time, the end of the arc moved along a complex strike-slip zone at the eastern edge of the Yucatan Peninsula (Lara, 1993) and opened the Yucatan back-arc basin in its wake (Rosencrantz, 1990) (Fig. 19). In northwestern South America, a Maastrichtian foreland basin associated with the accretion of oceanic plateau material widened and began to affect the area of western Venezuela (Pindell and Barrett, 1990) (Fig. 18). These foreland basin deposits will later become overprinted by the effects of the late Neogene collision of the Panama arc with northwestern South America. By Early Eocene, arc-continent collision was complete in western Cuba and collision proceeded in a diachronous manner along the edge of the Bahamas Platform (Gordon et al., 1997; Masaferro and Eberli, Chapter 7) (Figs. 19 and 20). This diachronous collision accompanied transfer of microplates from the Caribbean plate to the North America Plate in a clockwise fashion as forward progress of the Great Arc was halted by its collision with the Bahamas Platform (Mann et al., 1995). A thin foreland basin formed between the collision zone and the Bahamas carbonate platform (Hempton and Barros, 1993). Similarly, in northern South America, collision ended in Eocene time in the Lake Maracaibo area of western Venezuela and proceeded in a diachronous manner eastward along the northern margin of South America in Middle to Late Eocene time (Fig. 20). Initiation of oceanic spreading in the Cayman trough in Middle Eocene time may be the result of a change in the direction of the Great Arc from a northeastward to an eastward direction to move around the salient formed by the southeastern Bahama Platform (Mann et al., 1995). By Late Oligocene, the zone of active collision is in the present-day area of Puerto Rico on the northern plate boundary and eastern Venezuela on the southern boundary (Fig. 21). Late Cenozoic strike-slip phase
The Miocene to Recent period of Caribbean history corresponds to its strike-slip phase since by this time the arc-continent collisional zones have lengthened and converted into long strike-slip faults along the northern and southern edges of the Caribbean plate. During the Middle Miocene, the Cocos and Nazca plates ruptured along the Galapagos rift probably as a response to simultaneous subduction in two directions beneath the Middle America arc to the
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north and the Colombian trench to the south (Wortel and Cloetingh, 1981) (Figs. 22 and 23). By Late Miocene, localized convergence between the eastward-moving Caribbean plate and the southeastern extension of the Bahama Platform led to thrusting and topographic uplift in Hispaniola (Mann et al., Chapter 12; de Zoeten et al., Chapter 11) (Fig. 24). Along the southeastern margin of the plate, tectonic activity involved the Trinidad area (Babb and Mann, Chapter 18; di Croce et al., Chapter 16; Flinch et al., Chapter 17) (Fig. 24). By Late Pliocene, the margins of the Caribbean had reached its present-day configuration (Fig. 25). Two important tectonic and paleoceanographic events of this time was closure of the Panama seaway by collision of the Panama arc against northwestern South America (Kellogg and Vega, 1995: Mann et al., 1995) (Fig. 25) and the collision of the Cocos Ridge with southern Central America by about 1.6 Ma. The Panama arc collisional event may have accelerated the northwestward expulsion of the Maracaibo block into the Caribbean.
FUTURE WORK ON CARIBBEAN SEDIMENTARY BASINS
To conclude this review, I would like to leave the reader with some large-scale Caribbean tectonic problems that could be addressed by future studies of Caribbean sedimentary basins. Pacific vs. in situ origin of the Caribbean
The problem of the origin of the Caribbean is by no means solved despite the Pacific-origin approach that I have followed in this introduction and in the tectonic reconstructions. There are several problem areas for the Caribbean origin problem. First, North America-South America relative plate motion history (Mtiller et al., Chapter 2) (Fig. 12) only indirectly bears on the position of this intervening Caribbean plate and Great Arc through time. Second, paleomagnetic studies in the Caribbean are handicapped by several factors: (1) the problem of distinguishing large-scale plate-tectonic rotation from local structural rotation about vertical axes and apparent tectonic rotation (cf. MacDonald, 1980, for a discussion of paleomagnetic data from the Chortfs block); (2) the inability of paleomagnetism to address longitudinal changes in plate position of the type assumed for an eastward-moving Caribbean Great Arc and oceanic plateau in Cenozoic time; and (3) large error limits on existing data (Gose, 1985). And, third, studies of individual strike-slip offsets are problematic because, as shown on the reconstructions, these faults form somewhat late in the Caribbean tectonic history and therefore represent
24
P. M A N N
Fig. 18. Reconstruction of the Caribbean region at 71 Ma (Maastrichtian, magnetic anomaly C32n.2n). Key to abbreviation: eastern Andean fault zone.
Fig. 19. Reconstruction of the Caribbean region at 55.9 Ma (Early Eocene, magnetic anomaly C25n). Key to abbreviations: back-arc basin; G B -- Grenada back-arc basin; M B = Maracaibo foreland basin.
YB
EAFZ
--
= Yucat~in
CARIBBEAN SEDIMENTARY BASINS" CLASSIFICATION AND TECTONIC SETTING
25
Fig. 20. Reconstruction of the Caribbean region at 41.3 Ma (Middle Eocene, magnetic anomaly C19n). Key to abbreviations: Maracaibo foreland basin; E A F Z = eastern Andean fault zone.
MB
=
Fig. 21. Reconstruction of the Caribbean region at 25.5 Ma (Late Oligocene, magnetic anomaly C7An). Key to abbreviation: Gufirico foreland basin.
GB
=
26
P. M A N N
Fig. 22. Reconstruction of the Caribbean region at 15.1 Ma (Middle Miocene, magnetic anomaly C32n.2n). Key to abbreviations: G B = Gu~rico foreland basin; M B = Maturfn foreland basin.
Fig. 23. Reconstruction of the Caribbean region at 11.5 Ma (latest Middle Miocene, magnetic anomaly C5r.2n). Key to abbreviations: G B -- Gu(trico foreland basin; M B = Maturfn foreland basin.
27
CARIBBEAN SEDIMENTARY BASINS" CLASSIFICATION AND TECTONIC SETTING
Fig. 24. Reconstruction of the Caribbean region at 9.2 Ma (Late Miocene, magnetic anomaly C4Ar.2n). Key to abbreviations: Gufirico foreland basin; M B = Maturfn foreland basin.
GB
--
Fig. 25. Reconstruction of the Caribbean region at 3.1 Ma (Late Pliocene, magnetic anomaly C2An.2n). Key to abbreviation: Maturfn foreland basin.
MB
=
28 only a small part of the total Caribbean displacement. There are several promising new approaches for study of the origin of the Caribbean plate that can augment traditional plate reconstruction and paleomagnetic methods. Paleoenvironmental studies such as those by Pessagno et al. (Chapter 5) attempt to define changes in the paleolatitude of terranes using macro- and micropaleontologic data. Montgomery and Pessagno (Chapter 10) point out key indicator rocks, such as red cherts, that can be used to distinguish a Pacific vs. Atlantic environment of deposition. Finally, geochemical and high-resolution dating of igneous and metamorphic rocks of rocks of the Great Arc and oceanic plateau (e.g., Sinton et al., 1997) or the grains in sedimentary rocks from the allochthonous areas (e.g., Hutson et al., 1998) allows better constraints on plate reconstructions.
Age and environments of the Caribbean oceanic plateau Diebold and Driscoll (Chapter 19) and Driscoll and Diebold (Chapter 20) present data showing a two-stage Cretaceous evolution of the plateau and suggest that existing DSDP and ODP dated drill samples from the plateau may constrain only the later, smaller plateau-building event. Further outcrop studies and deep ocean drilling are needed to constrain this hypothesis.
Polarity reversal of the Caribbean arc Montgomery and Pessagno (Chapter 10) note two types of accreted sedimentary material in the Great Arc of the Caribbean: Pacific-derived material accreted when the arc was west or southwest-facing and Atlantic-type material accreted when the arc was east or northeastward-facing as it is today. Structural and stratigraphic outcrop studies are needed to confirm the existence and age of the proposed Early Cretaceous arc polarity reversal discussed by these authors and previous workers like L6bron and Perfit (1994) and Draper et al. (1996).
Origin of Caribbean ophiolites Extensive ophiolites were obducted during collision of the Great Arc with the passive margins of North and South America in northern Central America, the Greater Antilles, and northern South America. Gealey (1980) proposed that these ophiolites represent the basement of the forearc basement of the Great Arc. Other possible origins for the ophiolites include proto-Caribbean oceanic crust involved in the arc-continent collision and Caribbean oceanic plateau crust (Kerr et al., 1997). The overly-
P. MANN ing and interbedded sedimentary rocks could provide important clues to the origin of the ophiolites and their paleolatitudes through time.
Triggering of back-arc basins formation Depth to basement calculations and heat-flow measurements for the Yucat~in (Rosencrantz, 1990) and Grenada back-arc basins (Bird et al., Chapter 15) indicate that both basins formed rapidly over a short time interval in the Paleogene. Further geophysical work and deep-sea drilling is needed to confirm this history and understand why this basins opened rapidly and then became dormant despite continued subduction beneath the Great Arc.
Diachronous arc-continent collision and termination of arc activity The timing of this event summarized on Fig. 11B could be improved through careful biostratigraphic and stratigraphic studies as done by Bralower and Iturralde-Vinent (1997) in Cuba and several of the papers in this volume.
Amount of allochthoneity of Caribbean arcs The amount of overthrusting of the Great Arc over the passive margins of North and South America is not well understood because deep seismic data has not been attempted over the arc-continent collision zones. If the arc is far-traveled on a predominantly unmetamorphosed passive margin sequence, potential hydrocarbon deposits may exist at depth in areas where crystalline rocks are present at the surface.
Driving forces of Caribbean plate motion Mann et al. (1995) and Mann (1996) proposed that the Caribbean plate is driven as a response to dense oceanic slabs sinking beneath the Great Arc at the leading edge of the plate. The direction of the arc movement is therefore always oriented in the direction of oceanic crust or the 'free face'. However, MUller et al. (Chapter 2) have noted that the Caribbean plate remains fixed in a mantle reference frame since Middle Eocene time. For a stationary Caribbean plate, North America-South America post-Eocene north-south convergence rather than the presence of an oceanic free face may be the dominant plate-driving force affecting the Caribbean. GPS-based geodetic studies spanning the Caribbean plate could be used to test these differing dynamic scenarios.
C A R I B B E A N SEDIMENTARY BASINS: C L A S S I F I C A T I O N AND T E C T O N I C SETTING
Nature and driving forces of internal Caribbean plate deformation Diebold and Driscoll (Chapter 19) and Driscoll and Diebold (Chapter 10) propose that internal deformation of the Caribbean plate in the Colombian and Venezuelan basins and along the Beata Ridge and Hess Escarpment is a response to divergent deformation associated with the formation of the Cretaceous Caribbean oceanic plateau. In their view, modern escarpments on the seafloor are largely relict features that lack significant neotectonic deformation. In contrast, Mauffret and Leroy (Chapter 21) propose that the scarps in this region reflect internal disruption of the Caribbean plate along the line of the Beata Ridge. Faulting reflects mainly shortening in this intra-plate zone of deformation. Continued geophysical studies, reexamination of existing data, and GPS-based geodetic studies spanning the Caribbean plate are needed to distinguish these two ideas.
ACKNOWLEDGEMENTS
I would like to thank Dietmar MUller for providing the plate information to create the reconstructions in Figs. 13-25 and Lisa Gahagan and the UTIG PLATES project for creating the reconstructions. UTIG contribution 1422.
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Chapter 2
New Constraints on the Late Cretaceous/Tertiary Plate Tectonic Evolution of the Caribbean
R. D I E T M A R
MULLER,
JEAN-YVES
ROYER, STEVEN
C. C A N D E , W A L T E R R. R O E S T
a n d S. M A S C H E N K O V
We review the plate tectonic evolution of the Caribbean area based on a revised model for the opening of the central North Atlantic and the South Atlantic, as well as based on an updated model of the motion of the Americas relative to the Atlantic-Indian hotspot reference frame. We focus on post-83 Ma reconstructions, for which we have combined a set of new magnetic anomaly data in the central North Atlantic between the Kane and Atlantis fracture zones with existing magnetic anomaly data in the central North and South Atlantic oceans and fracture zone identifications from a dense gravity grid from satellite altimetry to compute North America-South America plate motions and their uncertainties. Our results suggest that slow sinistral transtension/strike-slip between the two Americas at rates roughly between 3 and 5 mm/year lasted until chron 25 (55.9 Ma). Subsequently, our model results in northeast-southwest-oriented convergence until chron 18 (38.4 Ma) at rates ranging between 3.7 4- 1.3 and 6.5 + 1.5 mm/year from 65~ to 85~ respectively. This first convergent phase correlates with a Paleocene-Lower Eocene calc-alkaline magmatic stage in the West Indies, which is thought to be related to northward subduction of Caribbean crust during this time. Relatively slow convergence until chron 8 at rates from 1.2 + 0.9 to 3.6 + 2.1 mm/year from 65~ to 85~ respectively, is followed by a drastic increase in convergence velocity. After chron 8 (25.8 Ma), probably at the Oligocene-Miocene boundary, this accelerated convergence resulted in 92 • 22 km convergence from chron 8 to 6, 127 + 25 km from chron 6 to 5, and 72 + 17 km from chron 5 to the present measured at 85~ near the North Panama Deformed Belt at convergence rates averaging 9.6 4- 3.1 and 9.6 4- 2.1 mm/year from chron 8 to 6 and chron 6 to 5, respectively, slowing down to 5.2 • 1.3 mm/year after chron 5. Neogene convergence measured at the eastern Muertos Trough, at 17.5~ 65~ is 41 4- 18 km from chron 8 to 6, 58 • 25 km from chron 6 to 5, and 22 4- 17 km from chron 5 to present day, at rates between 4.4 + 1.7 and 1.6 4- 1.0 ram/year. These well-resolved differential plate motions clearly show an east-west gradient in plate convergence in the Neogene, correlating well with geological observations. We suggest that the Early Miocene onset of underthrusting of the Caribbean oceanic crust below the South American borderland in the Colombian and Venezuelan basins, the onset of subduction in the Muertos Trough, and folding and thrust faulting at the Beata Ridge and the Bahamas, and the breakup of the main part of the Caribbean plate into the Venezuelan and Colombian plates, separated by the Beata Ridge acting as a compressional plate boundary (Mauffret and Leroy, Chapter 21) may all be related to the accelerated convergence between the two Americas. The main differences with previous analyses are that (1) our model results in substantial variations in convergence rates between the two Americas after chron 25 (55.9 Ma), (2) we have computed uncertainties for our North America-South America plate flow lines, and (3) we show Tertiary Caribbean plate reconstructions in an Atlantic-Indian hotspot reference system. Our absolute plate motion model suggests that the Caribbean plate has been nearly stationary since chron 18 (38.4 Ma). The east-west gradient in convergence between the Americas in the Neogene has not resulted in substantial eastward motion of the Caribbean plate, but rather contributed to causing its breakup into the Colombian and Venezuelan plates along the Beata Ridge where east-west-oriented compressional stresses are taken up. Our model also suggests that the eastward escape of the Caribbean plate in a mantle reference frame ceased when seafloor spreading started in the Cayman Trough, if the current interpretation of magnetic anomalies in the Cayman Trough is not grossly in error. Our model suggests that the opening of the Cayman Trough was accomplished by westward motion of the North American plate relative to a stationary Caribbean plate in a mantle reference system. This implies that subsequent North America-Caribbean and South America-Caribbean tectonic processes were no longer dominated by Cocos-Caribbean and Nazca-Caribbean plate interactions, as the latter had ceased to drive the Caribbean plate eastwards. We conclude that the west-northwestward motion of South America relative to a trapped, stationary Caribbean plate caused oblique collision along the passive margin of eastern Venezuela in the Neogene.
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsii), pp. 33-59. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
34 INTRODUCTION Many decades of research on deciphering the tectonic and sedimentary history of the Caribbean area (Fig. 1) have resulted in a fairly well understood tectonic framework of its evolution. The plate tectonic history between the two Americas has been reconstructed based on regional geophysical and geological data and its implications for Caribbean geology and have been evaluated in syntheses including Pindell et al. (1988), Ross and Scotese (1988), Pindell and Barrett (1990) and Stdphan et al. (1990). Our analysis builds on the knowledge that
R.D. MfJLLER et al. has accumulated from these and many other regional studies pertaining to Caribbean tectonic history. The purpose of this paper is not a comprehensive review of Caribbean tectonic evolution. Hence the reader will not find tables of syntheses of all tectonic events that may have occurred during Caribbean tectonic history, or all models that have been put forward to explain them. This information has been thoroughly reviewed by Pindell and Barrett (1990). Here we rather focus on extracting information from recently declassified dense satellite altimetry data that allow us to map the structure of the ocean floor in much more detail than previously possible and to
Fig. 1. Seafloor spreading isochrons in the Atlantic and eastern Pacific oceans from M~illeret al. (1997). Light gray shades correspond to young ocean floor ages and dark grays to old ages. The bold frame outlines the Caribbean area shown in Fig. 2.
NEW CONSTRAINTS ON THE PLATE TECTONIC EVOLUTION OF THE CARIBBEAN
35
Fig. 2. Main tectonic elements of the Caribbean area. Isochrons shown in the Cayman Trough are from Rosencrantz et al. (1988). The cross-hatched area centered on the Beata Ridge indicates the plate boundary between the Venezuelan and Colombian microplates. Regional names which are not tectonic elements are shown in italics.
better constrain past plate motions along consuming or transform plate boundaries such as the boundaries between the North American, South American, and Caribbean plates (Fig. 2). We utilize these data jointly with new and existing magnetic anomaly data to investigate the Late Cretaceous and Tertiary plate kinematic framework of the Caribbean region, including the computation of uncertainties for our plate reconstructions. Before the equatorial Atlantic between Africa and South America started opening at about chron M-0 time (120 Ma) (Pindell and Dewey, 1982; Mascle et al., 1988), the Caribbean tectonic framework was largely dependent on North AmericaAfrica plate motions, which were identical to North America-South America motion vectors prior to the opening of the South Atlantic. After initial opening of the equatorial Atlantic, plate boundaries between the two Americas were affected by North America-South America relative plate motions resulting from the difference vectors between North America-Africa and South America-Africa seafloor spreading, as first computed by Ladd (1976). Reconstruction of the Caribbean tectonic frame for this period requires closure of the North AmericaAfrica-South America plate circuit by using magnetic anomaly and fracture zone date sets that are standardized with respect to time. Because of the relatively slow velocity of North America-South
America plate motions after chron 34 (83 Ma) it is particularly important to obtain estimates of the uncertainties for the relative plate motion vectors between the two Americas in order to evaluate the resolution of plate motions models. We have combined a set of new magnetic anomaly data in the central North Atlantic between the Kane and Atlantis fracture zones, the 'CanaryBahamas Transect' (Maschenkov and Pogrebitsky, 1992), with existing magnetic anomaly data in the central North and South Atlantic oceans as well as with Seasat and Geosat satellite altimetry data to create a self-consistent data set for the two ocean basins. The finite motion poles and their uncertainties were estimated for 15 times from chron 34 to the present using an inversion method developed by Chang (1987, 1988), Chang et al. (1990), and Royer and Chang (1991), which allows a simple parameterization of the rotation uncertainties along a plate circuit path. The resolution of the estimates for North America-South America plate motions differs through time and is largely dependent on the velocity of their relative motions, i.e. faster plate motions are better resolved than slower motions. Even though the first-order features of our model are similar to Pindell et al.'s (1988) model, our results differ in detail, especially in the late Tertiary, and we stress the evaluation of uncertainties of plate motion vectors. In particular, our conclusions differ from
36 Pindell et al. (1988) in that we suggest that North America-South America plate motions may have had substantial effects on the structural development of the Caribbean area after chron 34 (83 Ma), especially during a period of rapid plate convergence in the Neogene.
DATA Magnetic anomaly data
The central North Atlantic is probably the ocean basin best covered by magnetic anomaly data. The three most comprehensive individual data sets are the northeast-southwest-trending Kroonvlag data (Collette et al., 1984), the trans-Atlantic Geotraverse (TAG) data set, which comprises a number of long east-west-oriented lines (Rona, 1980), and the Canary-Bahamas Transect (Maschenkov and Pogrebitsky, 1992), which comprises a dense set of survey lines between the Atlantis and Kane fracture zones from the Mid-Atlantic Ridge to about magnetic anomaly 13 (Fig. 3). These data sets are supplemented by a large number of other geophysical surveys (see Klitgord and Schouten, 1986, for previous compilation), resulting in dense data coverage.
R.D. MULLER et al. A comprehensive analysis of magnetic anomaly data in the South Atlantic was carried out by Cande et al. (1988). In order to create a self-consistent set of magnetic anomaly identifications for closing the North America-Africa-South America circuit, we use Cande et al.'s (1988) magnetic anomaly crossings in the South Atlantic (Fig. 3) and identify the same magnetic chrons as Cande et al. (1988) in the central North Atlantic. All magnetic anomaly identifications correspond to the young end of normal-polarity intervals, except for anomaly 33o. The phase shift angles were determined from paleomagnetic poles for North America from Harrison and Lindh (1982) and from the IGRF90 reference field. We use the young end of the following normal-polarity intervals according to the Cande and Kent (1995) magnetic reversal time scale: chron 5 (9.74 Ma), 6 (19.05 Ma), 8 (25.82 Ma), 13 (33.06 Ma), 18 (38.43 Ma), 21 (46.26 Ma), 24 (52.36 Ma), 25 (55.90 Ma), 30 (65.58 Ma), and 32 (71.59 Ma), 33o (79.08 Ma), and 34 (83.00 Ma). Fracture zone data
Fracture zones represent important information constraining plate motions and can be used in concert with magnetic anomaly data for computing
Fig. 3. New dense magnetic anomaly data in the Canary-Bahamas Transect area north of the Kane Fracture Zone (Maschenkov and Pogrebitsky, 1992), combined with data from other sources. Our magnetic anomaly identifications are shown as triangles (C5y, C18y, C24y), squares (C6y), upside-down triangles (C8y, C21y), and circles (C13y).
NEW CONSTRAINTS ON THE PLATE TECTONIC EVOLUTION OF THE CARIBBEAN finite rotations. Geosat, Seasat and ERS-1 altimetry data provide a unique data set to uniformly map the height of the sea surface, whose short-wavelength topography reflects uncompensated basement topography, such as that related to fracture zones. Miiller et al. (1991) demonstrated that there is an excellent correlation between the geoid anomaly and the basement structure of the Kane Fracture Zone in the central North Atlantic. They used geoid data from Geosat and subsatellite basement topography profiles of the Kane Fracture Zone to show that the average horizontal mismatch between geoid low and the axis of the basement trough, as mapped by Tucholke and Schouten (1988), is 5 km. The results of this comparative study represent 'ground truth' for the use of satellite altimetry data for accurately mapping slowly slipping Atlantic-type fracture zones. Following the Kane Fracture Zone study, Mtiller and Roest (1992) identified a number of small- and medium-offset fracture zones from the along-track Geosat and Seasat gravity data by picking the center of the gravity troughs corresponding to the deepest portion of the central fracture valleys. We re-identified these fracture zones from dense satellite-derived
37
gravity data (Sandwell and Smith, 1997). For the central North Atlantic reconstructions we use the Atlantis, Northern and Kane fracture zones (Fig. 4). When fracture zone offsets change from medium to small, as has happened in the case of the Northern Fracture Zone, whose offset diminished from 80 km at chron 25 to 20 km at chron 13, they may become unstable and commence to migrate along the ridge, producing V-shaped patterns. We used only those portions of fracture zones that appear to follow flow lines, i.e. have not migrated along the ridge axis. The location of the Kane Fracture Zone is constrained by Tucholke and Schouten's (1988) compilation of basement structure. Numerous fracture zones from the equatorial Atlantic to the southern South Atlantic record plate flow lines of seafloor spreading in the South Atlantic. Shaw and Cande (1990) pointed out that the northernmost fracture zones in this spreading system, i.e. the Marathon, Mercurius, Doldrums, and Four-North fracture zones in the equatorial Atlantic (Fig. 2), put important constraints on South America-Africa plate motions due to their proximity to the finite rotation poles. However, because no
Fig. 4. Gravity anomalies from satellite altimetry from Sandwell and Smith (1997) and interpreted and rotated magnetic anomaly and fracture zone identifications in the central North Atlantic. The unrotated magnetic and fracture zone identifications are identified by the following symbols: triangle (C5, 9.74 Ma; C18, 38.43 Ma; C30, 65.58 Ma); square (C6, 19.05 Ma, C21; 46.26 Ma; C32, 71.59 Ma), upside down triangle (C8, 19.05 Ma; C24, 52.36 Ma; C33o, 79.08 Ma), circle (C13, 33.06 Ma; C25, 55.90 Ma; C34, 83 Ma). All rotated data points are marked by crosses. Paleoridge or transform segments as defined by magnetic anomaly or fracture zone identifications, approximated as great circles in the inversion method used here, are denoted by alternating small and large symbols.
38
R.D. M U L L E R et al.
Fig. 5. Gravity anomalies from satellite altimetry from Sandwell and Smith (1997) and interpreted and rotated magnetic anomaly and fracture zone identifications in the South Atlantic. The symbols used for plotting magnetic and fracture zone identifications follow the same convention as in Fig. 4.
magnetic anomaly data have been identified in this area, we cannot resolve which portion of a fracture zone is relevant for a particular age, and whether these fracture zones reflect South America-Africa spreading for their entire length. We find that by using fracture zones south of 8~ only, we obtain very similar rotations compared with reconstructions in which equatorial fracture zones are included. Accordingly, our South Atlantic reconstructions are constrained by the Bodo Verde, Martin Vaz, Rfo Grande, and some unnamed fracture zones (Fig. 5). Both in the central North Atlantic and the South Atlantic only those fracture zones were used whose offsets through time are constrained by magnetic anomaly data. This is important, because the use
of incorrect fracture zone segments for constraining a given reconstruction would skew our results. We also avoid large-offset fracture zones, as they are not reliable indicators of plate motion changes over short geological time spans.
RECONSTRUCTION METHOD
The finite plate motion poles and their uncertainties were estimated using an inversion method developed by Chang (1987, 1988), Chang et al. (1990), and Royer and Chang (1991), based on the criterion of fit by Hellinger (1981). The uncertainties of a rotation are expressed as a covariance matrix,
NEW CONSTRAINTS ON THE PLATE TECTONIC EVOLUTION OF THE CARIBBEAN which is conceptually equivalent to the 'partial uncertainty rotations' described by Stock and Molnar (1983). In this method magnetic anomaly and fracture zone data are both regarded as points on two conjugate isochrons, which consist of great circle segments. The best fit reconstruction is computed by minimizing the sum of the misfits of conjugate sets of magnetic anomaly and fracture zone data points with respect to individual great circle segments. Consequently, both the resulting best-fitting rotations, as well as the sum of the misfits, depend critically on correctly identifying conjugate data points that belong to a common isochron segment. In practice, the application of Hellinger's criterion of fit poses no problem, because rotations have been published for most plates describing their Late Cretaceous/Tertiary history of motion. Hence, a starting rotation can be used for an initial reconstruction to identify conjugate isochron segments and data points. A possible disadvantage of applying Hellinger's (1981) criterion of fit to both magnetic and fracture zone data is that fracture zones cannot necessarily be expected to fit as well as magnetic anomaly data. Although all portions of a fracture zone have at some time been the location of a transform fault and part of an isochron, fracture zone morphology becomes overprinted successively at the transform-ridge intersections during changes in spreading direction. The amplitude of this effect is expected to increase as a function of transform length. For this reason we do not use fracture zones with offsets of more than about 150 km. Shaw and Cande (1990) recognized this problem and suggested an inversion method that incorporates fracture zones by minimizing the misfit of fracture zone data with respect to plate flow lines. The benefits of this model were expected to be a utilization of the 'integral constraints' of fracture zones, i.e. their continuity, as well as the possibility of allowing for consistent asymmetries of fracture zone limbs on conjugate plate flanks. Shaw and Cande (1990) implemented this method by minimizing the misfits of fracture zone data to symmetric flow lines. However, forcing fracture zones to fit symmetric flow lines may obscure distinct changes in spreading direction recorded in fracture zones bounded by two asymmetrically spreading corridors. Even though seafloor spreading appears to be remarkably symmetric in the long run, accretion of ocean floor through time periods of short 'stages' resolvable by magnetic anomaly data can be quite asymmetric. A critical aspect in our reconstruction method is the correct assessment of the uncertainties in the location of the data that will propagate into the uncertainties on the rotation parameters (location of the rotation pole and rotation angle). Following
39
a detailed analysis of the dispersion of magnetic anomaly-5 crossings in the Indian Ocean by Royer et al. (1997) we assigned 1 - a nominal uncertainties of 4 km to the magnetic anomaly crossings and of 5 km to fracture zone crossings following Mfiller et al.'s (1991) analysis. The uncertainties assigned to the data (6-) are related to their true unknown estimates (or) by the quality factor: 2 O"
Although ~c is unknown, the method developed by Royer and Chang (1991) allows to estimate s from the misfit, the geometry of the plate boundary and the number of data: N-2s-3
where N is the number of points, s the number of great circle segments, and r the total weighted misfit. Note that N - 2s - 3 corresponds to the number of degrees of freedom. Thus the parameter ~ indicates whether the assigned uncertainties are correct (s ~ 1), underestimated (~ > 1). Our reconstructions indicate that our nominal uncertainties are generally slightly overestimated. For the central Atlantic, ~ ranges from 1.07 to 3.08 and all degrees of freedom are larger than 50. This means that the average true uncertainties range from 2.3 (= 4/3,,/-3-,~.08) to 3.9 (= 4/~/-f.07) km for the magnetic crossings, and 2.8 to 4.8 km for the fracture zone crossings (Table 1). For the South Atlantic, ranges from 0.67 to 1.74 with degrees of freedom between 28 and 42. Thus the uncertainties should lie between 3.0 and 4.9 km for the magnetic data, and 3.8 to 6.1 for the fracture zone data (Table 2). Several factors may contribute to this discrepancy in the dispersion of the magnetic crossings: (i) there are much less data in the South Atlantic, hence we are probably combining data from different spreading corridors into individual segments; (ii) there are much less recent (i.e. GPS-navigated) cruises in the South Atlantic than in the Central Atlantic. Without further investigating this question, we decided to keep an uncertainty of 4 km for the magnetic crossings. Our inversions also suggest that the dispersion of the fracture zone data is generally better than 5 km. However, since fracture zones are not the optimal records for plate reconstructions, and since we do not know how well gravity troughs relate to the actual location of (paleo-)transform segments, we choose to remain conservative and use the Mtiller et al. (1991) result of an uncertainty of 5 km for fracture zones.
40
R.D. M U L L E R et al.
Table 1 North America-Africa finite rotations Chron 5 6 8 13 18 21 24 25 30 32 33o 34
Age (Ma)
Latitude (+~
Longitude (+~
Angle (o)
r (km)
t~ (km -1)
9.7 19.0 25.8 33.1 38.4 46.3 52.4 55.9 65.6 71.6 79.1 83.0
80.98 80.89 79.34 75.99 74.54 74.23 77.34 80.64 82.74 81.35 78.64 76.81
22.82 23.28 28.56 5.98 0.19 -5.01 -1.61 6.57 2.93 -8.32 -18.16 -20.59
2.478 5.244 7.042 9.767 11.918 15.106 16.963 17.895 20.84 22.753 26.981 29.506
76.3 62.3 39.9 79.2 36.9 42.9 17.8 51.7 80.6 66.9 55.2 52.1
2.46 1.96 2.53 1.19 1.65 1.19 3.08 1.26 1.07 1.33 1.85 1.82
df 188 122 101 94 61 51 55 65 86 89 102 95
N
s
271 191 140 159 94 96 94 110 135 142 155 144
40 33 18 31 15 21 18 21 23 25 25 23
6 mag (km)
6 fz
2.5 2.9 2.5 3.7 3.1 3.7 2.3 3.6 3.9 3.5 2.9 3
3.2 3.6 3.1 4.6 3.9 4.6 2.8 4.5 4.8 4.3 3.7 3.7
(km)
Parameters are: r -- sum of misfits; N -- number of data points" s = number of great circle segments; df = degrees of freedom; t~ = df/r (see text for discussion). The parameters ~ and the misfit r in this table are calculated with nominal uncertainties of 4 km for the magnetic anomaly crossings and 5 km for the fracture zone crossings. The 'true' data uncertainties O'mag and 6fz are related to their unknown estimates (o-) (i.e. 4 and 5 km for magnetic and fracture zone data, respectively) by the quality factor x -- (6-/o-) 2. Table 2 South America-Africa finite rotations Chron 5 6 8 13 18 21 24 25 30 32 33o 34
Age (Ma)
Latitude (+~
Longitude (+~
Angle (o)
r (km)
~ (km -1)
df
9.7 19.0 25.8 33.1 38.4 46.3 52.4 55.9 65.6 71.6 79.1 83.0
62.05 58.77 57.59 56.17 57.10 56.95 58.89 61.35 63.88 63.41 62.92 61.88
-40.59 -37.32 -36.27 -33.64 -33.00 -31.15 -31.18 -32.21 -33.61 -33.38 -34.36 -34.26
3.18 7.049 9.962 13.41 15.912 19.107 21.38 22.273 24.755 26.573 30.992 33.512
39.3 26.3 24.6 52.4 18.1 47.1 28.7 27.7 24.2 24.1 20.7 41.2
0.71 1.10 0.85 0.67 1.44 0.89 0.87 1.01 1.74 1.16 1.74 0.97
28 29 21 35 26 42 25 28 42 28 36 40
N
s
57 58 52 64 57 79 52 53 73 53 65 75
13 13 14 13 14 17 12 11 14 11 13 16
6-mag (km)
6 fz
4.7 3.8 4.3 4.9 3.3 4.2 4.3 4.0 3.0 3.7 3.0 4.1
5.9 4.8 5.4 6.1 4.2 5.3 5.4 5.0 3.8 4.6 3.8 5.1
(km)
For explanation of parameters, see Table 1.
RESULTS
Table 3 North America-Africa covariance matrices
North A m e r i c a - A f r i c a and South A m e r i c a - A f r i c a finite rotations F i n i t e r o t a t i o n s for N o r t h A m e r i c a - A f r i c a South America-Africa
and
plate motions were computed
for 12 t i m e s f r o m c h r o n 34 to the p r e s e n t . T h e r e s u l t i n g finite r o t a t i o n s , s t a t i s t i c a l p a r a m e t e r s , a n d c o v a r i a n c e m a t r i c e s are l i s t e d in T a b l e s 1 - 4 . T h e 9 5 % c o n f i d e n c e r e g i o n s are 3 - d i m e n s i o n a l e l l i p s o i d s in latitude, l o n g i t u d e , a n d r o t a t i o n a n g l e s p a c e . In o r d e r to r e p r e s e n t t h e m o n a m a p , the e l l i p s o i d s are p r o j e c t e d o n t o the l a t i t u d e - l o n g i t u d e
s p h e r e (cf. R o y e r a n d
Chron
a
b
c
d
e
f
5 6 8 13 18 21 24 25 30 32 33r 34
4.721 6.134 17.122 8.318 19.815 18.001 26.464 11.855 9.118 12.322 4.085 6.197
-4.047 -4.971 -17.129 -8.466 -20.774 -20.328 -31.305 -13.570 -10.652 -16.202 -4.022 -4.831
2.792 3.561 12.390 5.972 14.248 13.636 20.687 8.641 6.636 9.702 2.339 2.468
4.354 5.202 19.550 10.323 24.125 26.773 40.473 18.742 15.770 24.153 5.730 6.957
-2.897 -3.586 -14.054 -7.144 -16.397 -17.973 -26.665 -11.775 -9.831 -14.573 -3.475 -4.290
2.064 2.679 10.432 5.207 11.551 12.551 17.992 7.871 6.448 9.103 2.382 3.069
C h a n g , 1991). T h e p r o j e c t e d u n c e r t a i n t y e l l i p s e s inc l u d e u n c e r t a i n t i e s in b o t h l a t i t u d e , l o n g i t u d e , a n g l e , b u t o n e m u s t k e e p in m i n d that the true size o f the r o t a t i o n u n c e r t a i n t i e s (i.e. e l l i p s o i d s ) , w h i c h are des c r i b e d b y the c o v a r i a n c e m a t r i c e s , m i g h t n o t b e re-
Covariance matrices can be reconstructed in the following way:
NEW CONSTRAINTS ON THE PLATE TECTONIC EVOLUTION OF THE CARIBBEAN
Table 4 South America-Africa covariance matrices Chron
a
b
c
d
e
f
6 5 8 13 18 21 24 25 30 32 33r 34
27.203 20.913 63.296 9.525 31.698 11.165 25.100 25.115 15.017 30.163 9.249 8.711
-7.073 -4.353 -20.659 -2.385 -12.098 -3.955 -11.021 -11.568 -6.988 -13.978 -2.900 -4.332
-20.420 -13.079 -35.876 -5.374 -18.390 -6.408 -11.714 -13.323 -8.168 -16.671 -4.696 -4.971
3.179 1.974 8.134 2.117 5.766 2.546 6.070 7.341 5.207 9.696 3.493 4.014
5.487 2.804 11.603 1.676 7.124 2.525 5.352 6.681 4.322 8.326 2.513 3.230
16.685 9.422 23.341 3.765 12.671 4.334 6.722 8.144 5.358 10.185 3.347 3.649
For reconstruction of covariance matrices see Table 3.
flected well by their 2-D projection onto the sphere. For instance, in the case of the South Atlantic, the 2-D uncertainty in the location of the chron-5 bestfitting pole appears to be at least 10 times larger than the 2-D uncertainty in the location of the chron-21 rotation pole, whereas the volume of the 95% confidence region for chron-5 rotation is only about 3 times larger than for the chron-21 rotation (7838 versus 2773 km 3, respectively). Conversely, the 95% uncertainty volume for chron 8 (19,705 km 3) is 2.5 larger than for chron 5, whereas its 2-D projection is about 3 times smaller than for chron 5. In Fig. 6 we compare our results with the North America-Africa pole path from Klitgord and Schouten (1986) and the South America-Africa pole path from Shaw and Cande (1990). The main difference for the central North Atlantic is that our path shows a sharp bend at anomaly-8 time (25.8 Ma), resulting in a pole path for chron 5 to chron 8 that is distinctly different from the chron-21 to chron-8 path, whereas Klitgord and Schouten's (1986) path is continuous between chrons 21 and 5. The consequences of this difference for Neogene North America-South American plate motions are discussed in subsequent sections. Our South Atlantic pole path is generally similar to Shaw and Cande's (1990) path, but less smooth. The similarity reflects the fact that both models are based on very similar data sets, whereas the differences reflect properties of the different inversion methods used to compute finite rotations as well as new fracture zone identifications. The 'integral constraints' of fracture zone continuity used in Shaw and Cande's (1990) method is probably the main reason for the smoothness of their pole path. However, the use of symmetric flow lines to evaluate the fit of fracture zone data might smooth out real cusps in a finite pole path that is computed from reconstructing data from individual isochrons independently. The main difference between Shaw and
41
Cande's (1990) model and our model for the South Atlantic (Fig. 5) is the anomaly-13 reconstruction, which results in a distinct cusp in our pole path, while there is a less accentuated cusp in their path. One way to qualitatively test whether or not such a cusp is supported by the data is to construct plate flow lines and evaluate their fit to fracture zones. For South Atlantic spreading this is done best in the equatorial Atlantic, because the fracture zones here are the best recorders of changes in spreading direction due to their proximity to the finite motion poles. However, plotting plate flow lines in this area raises the problem of knowing where the South America-North American plate was located through time. Plotting flow lines both for North America-Africa and South America-Africa for the same fracture zone allows us to address this problem. We plot South Atlantic plate flow lines derived from our model in Fig. 7a with the gridded gravity anomaly field computed from Seasat, Geosat, and ERS-1 data (Sandwell and Smith, 1997). All flow lines are constructed using both ridge-transform intersections as seed points. The area encompassed by the resulting dual flow lines approximates the amount of transpression or transtension that occurred during changes in spreading direction, assuming symmetric plate accretion. The Marathon and Mercurius fracture zones do not fit the computed flow lines well. However, the Four-North and Doldrums fracture zones show a good fit to our flow lines. The latter observation confirms the validity of our South Atlantic plate motion model, whereas the former indicates that the North-South American plate boundary may be located in the vicinity of the Marathon and Mercurius fracture zones, resulting in deviations from South America-Africa flow lines. In comparison, we show central North Atlantic plate flow lines in the equatorial Atlantic Ocean in Fig. 7b. The Marathon Fracture Zone flow line clearly does not match the gravity anomaly expression of this fracture zone. The post-chron 6 flow line of the Mercurius Fracture Zone fits its eastern limb better than the South America-Africa flow line (compare with Fig. 7a), but not its western limb. Its is clear that this fracture zone would not be useful to constrain either North America-Africa nor South AmericaAfrica plate motions, since it appears to have been affected by plate boundary processes. We can draw the conclusion that applying Hellinger's criterion of fit jointly to the magnetic anomaly data and fracture zone data from continuous fracture zones, which follow plate flow lines, and reconstructing these data independently for each isochron, results in plate models that produce smooth continuous flow lines, even though this property is not utilized or imposed by the model inversion, as it is in Shaw and Cande's (1990) method.
R.D. MULLER et al.
Fig. 6. Finite rotation poles and 95% confidence ellipses for North America-Africa (upper right) and South America-Africa (lower left) plate motions for 12 reconstruction times from chron 34 (83 Ma) to the present. As a comparison we show the North America-Africa pole path from Klitgord and Schouten (1986) and the South America-Africa pole path from Shaw and Cande (1990) (gray triangles). Note the sharp cusp in our North America-Africa pole path at chron 8 (25.8 Ma), resulting in a pole path for chron 5 to chron 8 that is distinctly different from Klitgord and Schouten's (1986) path. The differences between ours and both Klitgord and Schouten's (1986) and Shaw and Cande's (1990) models reflect a much improved accuracy in locating fracture zones based on a dense gravity anomaly grid from satellite altimetry (Sandwell and Smith, 1997), and some new magnetic anomaly data.
North America-South America finite rotations Tables 5 - 8 list the rotation parameters for the North A m e r i c a - S o u t h America relative motions, for each chron resulting from the product of the South A m e r i c a - A f r i c a rotation with the corresponding Africa-North America rotation (Table 9). The resulting pole path (Fig. 8) shows: (i) very stable poles of motion from chron 25 to 21, and from chron
18 to 8; (ii) an important northward migration of the rotation poles from chron 34 to chron 18; (iii) followed by a southward migration until chron 6. For this reason we computed only 8 stage rotations (Table 6) to describe the main episodes of relative motion between the two American plates. The North A m e r i c a - S o u t h America rotation stage poles always lie outside the Caribbean plate (Table 6); for the ages younger than chron 25 (55.9 Ma), they lie within or
NEW CONSTRAINTS ON THE PLATE TECTONIC EVOLUTION OF THE CARIBBEAN
43
Fig. 7. (a) Gravity anomalies and South Atlantic symmetric plate flow lines in the Equatorial Atlantic Ocean. Flow lines are composed of segments for stages bounded by our reconstruction times (see Fig. 4). All flow lines have been constructed using both ridge-transform intersections as seed points. The area encompassed by the resulting dual flow lines approximates the amount of transpression or transtension that occurred during changes in spreading direction, assuming symmetric plate accretion. The Marathon and Mercurius fracture zones do not fit the computed flow lines well. However, the Four-North and Doldrums fracture zones show a good fit to our flow lines. The latter observation confirms the validity of our South Atlantic plate motion model, whereas the former indicates that the North-South American plate boundary may be located in the vicinity of the Marathon and Mercurius fracture zones, resulting in deviations from South America-Africa flow lines. (b) Gravity anomalies and central North Atlantic symmetric plate flow lines in the Equatorial Atlantic Ocean. The Marathon Fracture Zone flow line clearly does not match the gravity anomaly expression of this fracture zone. The post-chron-6 flow line of the Mercurius Fracture Zone fits its eastern limb better than the South America-Africa flow line (compare with Fig. 7a), but not its western limb. It is is clear that this fracture zone would not be useful to constrain either North America-Africa nor South America-Africa plate motions, since it appears to have been influenced by plate boundary processes.
in the vicinity of the North A m e r i c a - S o u t h A m e r i c a plate boundary, implying a different sense of motion along strike of this plate boundary.
North America-South America plate motions in the Caribbean area Ideally we would like to be able to compute North A m e r i c a - C a r i b b e a n and South A m e r i c a - C a r i b b e a n plate motions and compare the m o d e l e d motion vectors with m a p p e d plate boundary deformation. Accurate estimates for North A m e r i c a - C a r i b b e a n plate motions are available for times after chron 6 (19.0 Ma), as recorded by seafloor spreading in the
C a y m a n Trough (Rosencrantz et al., 1988; Mauffret and Leroy, Chapter 21). The magnetic anomalies on older portions of the C a y m a n Trough are less straightforward to identify, and the chronology of both pre- and post-chron 6 seafloor spreading here is still the subject of debate. For these reasons we first analyze the North A m e r i c a - S o u t h America plate kinematic framework, as it is independent on knowledge of the spreading history in the C a y m a n Trough, and compare the results with geological and geophysical data from the northern and southern Caribbean plate boundaries. Fig. 9 shows North A m e r i c a - S o u t h A m e r i c a n plate motion through time, illustrated by the suc-
44
R.D. MULLER et al.
Table 5 North America-South America finite rotations Chron
Age (Ma)
Latitude (+~
Longitude (+~
Angle (o)
df
5 6 8 13 18 21 24 25 30 32 33o 34
9.7 19.0 25.8 33.1 38.4 46.3 52.4 55.9 65.6 71.6 73.6 83.0
15.06 14.76 17.52 16.99 18.12 13.61 13.95 13.87 11.25 9.58 9.03 7.31
-56.59 -52.43 -53.52 -53.13 -54.60 -53.09 -52.33 -53.04 -53.89 -53.74 -56.57 -58.00
1.408 3.433 5.146 6.067 6.498 7.11 8.188 8.777 9.009 8.947 9.12 9.332
216 151 122 129 87 93 80 93 128 117 138 135
Table 6 North America-South America stage rotations (North America reference frame) Stage (Chron a-b)
Age (Ma)
Latitude Longitude Angle (+~ (+~ (o)
0- 5 5- 6 6- 8 8-18 18-25 25-30 30-33o 33o-34
9.7 9.3 6.8 12.6 17.5 9.7 8.0 9.4
15.06 14.42 23.08 20.57 1.76 -43.69 -31.75 -32.74
-56.59 -49.56 -55.54 -58.63 -49.85 -80.57 -124.85 -102.67
1.408 2.03 1.727 1.358 2.355 0.484 0.557 0.417
cessive motion of three points attached to the North American plate with respect to the South American plate for eleven stages from chron 34 (83 Ma) to the present. The paths of these points through time as shown on Fig. 9 correspond to conventional plate flow lines illustrating relative North A m e r i c a - S o u t h America plate motion. The only difference here is that for each stage rotation vector a simultaneous
95% confidence region is plotted about the young ends of the relative motion vectors, i.e. the ellipse about points at chron 33o reflect the uncertainty for the chron 3 4 - 3 3 o stage rotation. A simultaneous confidence region represents the area (on the surface of the Earth) in which a point on a given plate may have been located with equal likelihood for a particular reconstruction time, with 95% confidence. The simultaneous 95% confidence regions increase in size with increasing distance from the stage pole of motion; in the Caribbean area they increase in size correspondingly from east to west (Fig. 9). The reader may note that even though the projected confidence ellipsoids of our finite rotation poles for North A m e r i c a - S o u t h American plate motion for chrons 18-6 show large overlaps (Fig. 8), the simultaneous 95% confidence regions for the three rotated points shown show only little overlap. This may appear puzzling, but only reflects the imperfect nature of the 3-dimensional finite rotation pole error ellipsoid projections onto a spherical surface on Fig. 8, as discussed before. In other words, this result shows that the covariance matrices and uncertainty ellipsoids for these reconstructions are different enough to distinguish different phases of convergence in the late Tertiary outside of the 95% error bounds. We cannot resolve very slow relative motions between chrons 25 and 24; some overlap between the confidence regions for chrons 32 and 30 and chrons 18 and 13 is also observed. However, all other stage vectors are well resolved. Consequently, we compute a new set of 8 stage rotation poles (Table 6) and corresponding relative motion vectors, including only stages for which we can resolve relative North A m e r i c a - S o u t h America plate motions (Fig. 10). The relative motion vectors in this set of stages can be divided into four age groups: (1) slow sinistral strike-slip/transtension between chrons 34 and 25 (~2.8 • 0.8-4.8 • 1.1 m m / y e a r at 85~ (2) northeast-southwest-oriented conver-
Table 7 Covariance matrices for North America-South America finite rotations Chron
a
b
c
d
e
f
5 6 8 13 18 21 24 25 30 32 330 34
25.693 33.509 81.193 18.284 52.534 30.632 54.006 38.048 25.161 44.118 13.746 15.362
-8.424 -12.099 -38.062 -11.008 -33.276 -24.772 -43.018 -25.517 -18.055 -30.783 -7.130 -9.517
-10.306 -16.901 -23.479 0.625 -4.194 7.601 9.682 -4.414 -1.233 -6.327 -2.336 -2.616
6.305 8.283 27.207 12.113 29.357 27.959 44.281 25.021 19.886 32.094 8.844 10.472
-0.060 1.988 -2.057 -5.265 -8.715 -14.851 -20.378 -4.705 -5.158 -5.741 -0.753 -0.732
11.451 19.289 33.476 8.859 23.734 16.779 24.534 16.000 11.870 19.408 5.696 6.765
See parameter legend in Table 3.
45
N E W C O N S T R A I N T S ON THE PLATE T E C T O N I C E V O L U T I O N OF THE C A R I B B E A N
Fig. 8. Finite rotation poles for 12 reconstruction times for North America-South America relative plate motions since chron 34 (83 Ma). These poles have been computed by combining the finite rotations and their 95% confidence regions shown in Fig. 6. Note the large uncertainty for the chron-5 (9.7 Ma) reconstruction. This partly reflects that the South Atlantic reconstruction for this time is not well constrained. It also reflects that the uncertainty of finite rotations increases with decreasing finite rotation angle.
Table 8 Volumes of uncertainty ellipsoids for North-South America finite rotation poles Chron
Latitude
Longitude
Angle
df
5 6 8 13 18 21 24 25 30 33o 34
15.06 14.76 17.52 16.99 18.12 13.61 13.95 13.87 11.25 9.03 7.31
-56.59 -52.43 -53.52 -53.13 -54.6 -53.09 -52.33 -53.04 -53.89 -56.57 -58
1.408 3.433 5.146 6.067 6.498 7.11 8.188 8.777 9.009 9.12 9.332
216 151 122 129 87 93 80 93 128 138 135
Vol. 95%
Vol. 1 - o-
14,548 22,448 62,456 16,470 46,085 28,396 48,212 38,850 26,246 13,898 14,696
650 993 2741 724 1988 1230 2069 1682 1154 613 647
46
R.D. MOLLER et al.
Fig. 9. North America-South America plate motion through time, illustrated by the successive motion of three points attached to the North American plate with respect to the South American plate for eleven stages from chron 34 (83 Ma) to the present. The paths of these points through time as shown on Fig. 9 correspond to conventional plate flow lines illustrating relative North America-South America plate motion. The only difference here is that for each stage rotation vector a simultaneous 95% confidence region is plotted about the young ends of the relative motion vectors, i.e. the ellipse about points at chron 33o reflect the uncertainty for the chron 34-33o stage rotation. A simultaneous confidence region represents the area (on the surface of the Earth) in which a point on a given plate may have been located with equal likelihood for a particular reconstruction time, with 95% confidence. This figure demonstrates that we cannot resolve very slow relative motions between chrons 25 and 24; some overlap between the confidence regions for chrons 32 and 30 and chrons 18 and 13 is also observed. However, all other stage vectors are well resolved. Consequently, we compute a new set of relative motion vectors, including only stages for which we can resolve relative North America-South America plate motions well, shown in Fig. 10.
gence from chron 25 to 18 (6.5 4- 1.5 m m / y e a r at 85~ (3) slow compressional motion from chron 18 to 8 (3.6 4-2.1 m m / y e a r at 85~ and (4) fast n o r t h - s o u t h - o r i e n t e d convergence from chron 8 to 6 (9.6 + 3.1 m m / y e a r for chrons 8 - 6 and 9.6 4- 2.1 m m / y e a r for chrons 6 - 5 at 85~ followed by a deceleration in n o r t h - s o u t h - o r i e n t e d convergence post-chron 5 (5.2 4- 1.3 m m / y e a r at 85~ In order to directly compare our model with Pindell et al.'s (1988) results, we compute N o r t h - S o u t h A m e r i c a relative motions for the same 7 stages as
Table 9 Volumes of uncertainty ellipsoids for North America-Africa stage rotation poles Stage
Latitude
Longitude
Angle
Vol. 95%
5- 6 6- 8 8-18 18-25 25-30 30-33o 33o-34
14.42 23.08 20.57 1.76 -43.69 -31.75 -32.74
-49.56 -55.54 -58.63 -49.85 -80.57 -124.85 -102.67
2.03 3 6 , 3 1 7 1662 1.727 116,330 5324 1.358 149,039 6821 2 . 3 5 5 117,073 5358 0 . 4 8 4 87,663 4012 0 . 5 5 7 5 6 , 3 6 8 2580 0 . 4 1 7 3 9 , 2 9 4 1798
Pindell et al.'s (1988) (Fig. 11). The general shape of both models is similar, reflecting the similarity of the magnetic anomaly data sets used to constrain both models. The main difference between the models is that Pindell et al.'s (1988) model implies relatively constant convergence between the Americas of rates b e t w e e n 6 and 4 m m / y e a r since chron 21 (46.3 Ma). Our model results in substantial variations in convergence rates since chron 25 (55.9 Ma), as d o c u m e n t e d in Fig. 10 and Table 10. In particular our model resolves the onset of fast convergence after chron 8 at a speed of nearly 10 m m / y e a r , c o m p a r e d with less than 4 m m / y e a r from chrons 18 to 8 measured at 85~ Without identifying magnetic anomaly 8, two stages of slow (chron 13-8) and fast (chron 8 - 6 ) convergence would be averaged.
Vol. 1 - cr
DISCUSSION Lesser Antilles to Mid-Atlantic Ridge The North A m e r i c a - S o u t h A m e r i c a plate boundary east of the Caribbean region is characterized by
47
NEW CONSTRAINTS ON THE PLATE TECTONIC EVOLUTION OF THE CARIBBEAN
Table 10 Uncertainties on motion vectors (3-D 95% confidence limits) Chron
Time span (Ma)
rmaj (km)
rmin (km)
Om~
-2 -34 48 30 -82 -64 -33 -42
27 42 55 54 41 34 27 24
6 10 12 12 12 11 11 12
53 56 57 58 60 62 60 60
10 25 37 38 28 14 8 8
-9 -159 72 44 -102 -67 -35 -45
27 43 56 55 42 35 28 24
7 l0 13 15 13 12 12 12
49 53 53 54 56 58 57 56
4 32 24 11 74 48 61 42
-4- 18 + 23 -4- 21 + 33 • 30 i 14 -t- 9 • 8
-154 - 170 108 84 -119 -71 -37 -48
28 45 58 57 44 36 28 25
7 11 15 17 16 13 13 12
46 49 49 50 51 54 52 52
+ 1.0 -4- 1.7 + 1.4 + 0.9 + 1.3 4- 1.1 i 0.9 + 0.7
22 58 40 22 91 47 59 40
+ 13 4- 22 4- 13 • 15 4-32 • 14 • 10 + 9
172 178 136 138 -139 -76 -39 -52
29 46 60 59 45 38 30 26
8 13 19 22 20 16 14 13
41 44 44 43 44 46 45 45
3.5 7.1 6.8 2.4 5.1 3.4 5.0 2.9
• 1.2 • 1.9 + 2.1 -t- 1.5 -+- 1.4 4- 1.1 + 1.0 + 0.7
47 94 65 42 127 47 57 38
-t- 16 • 25 • 20 -t- 26 4-34 + 15 -+- 11 i 9
174 177 154 161 -151 -85 -45 -60
30 48 63 64 49 41 32 29
l0 15 22 26 23 19 16 14
35 39 37 34 33 35 35 36
5.2 9.6 9.6 3.6 6.5 3.5 4.8 2.8
+ 1.3 • 2.1 + 3.1 + 2.1 + 1.5 • 1.1 -t- 1.1 • 0.8
72 127 92 65 162 47 55 37
-t- 18 • 27 + 30 • 36 + 38 • 15 4- 12 + 10
174 176 161 167 - 159 -93 -51 -70
31 49 66 68 54 45 35 31
11 17 25 29 26 21 17 15
30 34 31 27 24 26 26 28
Rate of motion (mm/year)
Motion vectors (km)
9.7 9.3 6.8 12.6 17.5 9.7 8.0 9.4
1.6 0.6 3.6 1.6 2.4 3.6 5.4 3.3
• 4• 4• • 4•
0.8 0.5 4.0 1.9 0.9 1.0 0.7 0.6
22 7 34 29 60 49 61 43
• • • + • 4• 4-
10 7 38 33 23 14 8 8
9.7 9.3 6.8 12.6 17.5 9.7 8.0 9.4
0.7 1.0 2.7 1.0 2.6 3.5 5.4 3.2
-t- 0.7 • 2.0 • 3.9 • 2.1 • 1.1 • 1.1 • 0.8 • 0.6
9 13 25 18 64 48 61 42
• 4• • • • • •
9.7 9.3 6.8 12.6 17.5 9.7 8.0 9.4
0.3 2.4 2.6 0.7 3.0 3.5 5.3 3.1
• 1.3 -t- 1.8 + 2.2 + 1.8 -t- 1.2 + 1.1 • 0.8 • 0.6
9.7 9.3 6.8 12.6 17.5 9.7 8.0 9.4
1.6 4.4 4.2 1.2 3.7 3.4 5.2 3.0
9.7 9.3 6.8 12.6 17.5 9.7 8.0 9.4 9.7 9.3 6.8 12.6 17.5 9.7 8.0 9.4
Azimuth (o from N)
(o from N)
Longitude: 48~ 0- 5 5- 6 6- 8 8-18 18-25 25-30 30-33o 33o-34
Longitude: 52~ 0- 5 5- 6 6- 8 8-18 18-25 25-30 30-33o 33o-34
Longitude: 58~ 0- 5 5- 6 6- 8 8-18 18-25 25-30 30-33o 33o-34
Longitude: 65~ 0- 5 5- 6 6- 8 8-18 18-25 25-30 30-33o 33o-34
Longitude: 75~ 0- 5 5- 6 6- 8 8-18 18-25 25-30 30-33o 33o-34
Longitude: 85~ 0- 5 5- 6 6- 8 8-18 18-25 25-30 30-33o 33o-34
The parameters rmaj and rmin are the semi-major and semi-minor axes of the ellipse of confidence (95% level). The variable 0maj is the azimuth of the semi-major axis.
48
R.D. MULLER et al.
Fig. 10. Gravity anomalies in the Caribbean area and North America-South America plate motion through time of three points attached to the North American plate with respect to the South American plate for eight stages from chron 34 (83 Ma) to the present. The relative motion vectors can be divided into four groups: (1) slow sinistral strike-slip/transtension between chrons 34 and 25 (55.9 Ma) (~2.8 4-0.8-4.8 + 1.1 mm/year at 85~ (2) northeast-southwest-oriented convergence from chron 25 to 18 (38.4 Ma) (6.5 -+- 1.5 mm/year at 85~ (3) slow motion from chron 18 to 8 (25.8 Ma) (3.6-t-2.1 mm/year at 85~ and (4) fast north-south-oriented convergence from chron 8 to 6 (19.0 Ma) (9.6 4- 3.1 mm/year for chrons 8-6 and 9.6 4- 2.1 mm/year for chrons 6-5 at 85~ followed by a deceleration in north-south-oriented convergence post-chron 5 (9.7 Ma) (5.2 4- 1.3 mm/year at 85~ See Table 6 for a complete list of stage motion vectors. Note acceleration in convergence at chron 8.
a number of anomalous ridges and troughs. The Barracuda and Tiburon ridges east of the Lesser Antilles Arc (Fig. 12) both exhibit unusually large Bouguer gravity anomalies (up to ~135 mGal). The eastern continuation of the plate boundary is expressed in the Researcher Ridge and Royal Trough (Fig. 12). The Royal Trough exhibits en-dchelon-shaped tectonic fabric and fresh basalts on a basement characterized by spreading center type faulting identified from GLORIA data, whereas the Researcher Ridge has a large magnetic anomaly. Both features are interpreted as extensional structures (Collette et al., 1984; Roest and Collette, 1986). For the area at the Tiburon and Barracuda ridges, our plate model results in a total of 151 -+- 31 km left-lateral transtension from chron 34 (83 Ma) to 25 (55.9 Ma). A phase of slow transpression is predicted for chron 25 (55.9 Ma) to 18 (46.3 Ma), which is not well resolved, followed by extremely slow motion between chrons 18 and 6, during which time it is within the computed 95% confidence areas for the entire part of the plate boundary east of the Lesser Antilles subduction zone (Fig. 12). The plate boundary area east of 56~ was located quite close to the stage poles of motion; the resulting vectors of relative motion are so small that they cannot be resolved. Convergence in
the Barracuda-Tiburon ridge area started at chron 6. We model 32 4-23 km of convergence from chron 6 to 5, but virtually no relative motion after chron 5 ( 4 + 18 km), which is much too small to be resolved, by our model. The post-chron 8 relative motion in the Royal Trough area can only be resolved for post-chron 5 time. Here our model implies 22 + 10 km of extension for the last 10 million years. Our plate model supports the suggestion that the present topographic and gravity expression of the Tiburon and Barracuda ridges may have resulted primarily from Neogene plate convergence (cf. Mtiller and Smith, 1993). In particular, the strong Moho uplift at the Tiburon Ridge, where the crust is modeled to be 1.5-2 km thick with a Moho uplifted more than 4 km (Mt~ller and Smith, 1993), is extremely unstable. Without compressive forces, oceanic crust as thin as 1.5-2 km would subside and form a depression, rather than a ridge. The plate model used by Miiller and Smith (1993) did not allow them to discriminate when in the Tertiary North-South America convergence in this area and the uplift of the two ridges might have been initiated. However, they argued that Middle Eocene-Late Oligocene turbidites on the slope of the Tiburon Ridge, which is now located 800 m above the abyssal plain, may
NEW CONSTRAINTS ON THE PLATE TECTONIC EVOLUTION OF THE CARIBBEAN
49
Fig. 11. North America-South America plate motion through time of three points attached to the North American plate with respect to the South American plate computed for the same seven stages as Pindell et al.'s (1988) model, shown by triangles. The general shape of both models is similar, reflecting the similarity of the magnetic anomaly data sets used to constrain both models. The main differences between the models are as follows. (1) Pindell et al.'s (1988) model implies sinistral strike-slip between North and South America from chron 34 to 21, followed by convergence. Our model implies strike-slip until chron 25, followed by convergence. (2) Pindell et al.'s (1988) model implies relative constant convergence between the Americas of rates of 4-6 ram/year since chron 21. Our model results in substantial variations in convergence rates from chron 25, as documented in Fig. 10 and Table 6. In particular we resolve the onset of fast convergence after chron 8 of nearly 10 ram/year, compared with less than 4 ram/year from chrons 18 to 8 measured at 85~ This figure also shows the estimated motion of the Caribbean plate in a hotspot reference frame since chron 18 (38.4 Ma) (open circles); solid circles show the present-day position of the two points attached to the Caribbean plate used as starting points for these flow lines. See text for discussion.
indicate that most of its uplift occurred in postOligocene times, a conclusion strongly supported by the plate model presented here. This idea contrasts the interpretation of Dolan et al. (1989, 1990), who suggested that the present topography of the Tiburon Ridge had existed prior to deposition of the turbidites, which would have been emplaced by upslope deposition from the abyssal plain onto the rise. Dolan et al. (1989, 1990) argue that the Tiburon Rise has existed as a bathymetrically shallow feature since the Late Cretaceous. An extensive discussion of this question can be found in Miiller et al. (1993) and Mtiller and Smith (1993). We find that there is little evidence that would show conclusively that the present bathymetric elevation of the Tiburon Rise has existed since the Late Cretaceous. In contrast, the crustal structural modeling carried out by Mtiller and Smith (1993), together with the plate model presented here, indicates that much of the unusually shallow Moho topography and crustal uplift of the Tiburon Rise is a result of the onset of North A m e r i c a - S o u t h America convergence at chron 6 (19.0 Ma). A Neogene formation of the present topography of the Tiburon Rise would also
alleviate the need for upslope turbidite deposition on the rise in the Middle E o c e n e - L a t e Oligocene, as suggested by Dolan et al. (1989, 1990). Instead, the present elevation of the turbidites on the slope of the rise would reflect post-depositional uplift.
Middle American Trench to Lesser Antilles Our results suggest that slow sinistral transtension/strike-slip between the two Americas lasted from chron 34 (83 Ma) until chron 25 (55.9 Ma) (~2.8 + 0.8-4.8 + 1.1 m m / y e a r at 85~ However, we cannot attribute any particular Caribbean tectonic events to this phase of slow sinistral motion. Our model results in roughly northeastsouthwest-oriented slow convergence from chron 25 (55.9 Ma) to chron 18 (38.4 Ma) (6.5 + 1.5 m m / y e a r at 85~ This time period of convergence includes a calc-alkaline magmatic stage in the West Indies, dated as P a l e o c e n e - L o w e r Eocene (Perfit and Heezen, 1978), which is thought to be related to southward subduction of proto-Caribbean crust during this time. North-South America relative motion was characterized by relatively slow transpression
50
R.D. MULLER et al.
Fig. 12. North America-South America plate motion through time of three points attached to the North American plate with respect to the South American plate for eight stages from chron 34 (83 Ma) to the present and their simultaneous 95% confidence regions. Only the Barracuda Ridge area has been affected by North-South American plate motions as old as chron 34 (83 Ma). The ocean crust to the east becomes successively younger. The oldest relative motion vectors plotted correspond roughly to the age of the ocean crust south of the Fifteen-Twenty Fracture Zone. Relative plate motion in this area is not well resolved, except for oblique transtension in the Barracuda Ridge area from chron 34 to 25, and post-chron-6 north-south compression, post-chron-6 extension in the Royal Trough area. Color version at http://www.elsevier, nl/locate/caribas/ until chron 8 (25.8 Ma) (3.6-+-2.1 mm/year at 85~ This phase of slow relative motion between the two Americas correlates with a period of tectonic quiescence along some parts of the Caribbean margins during the Oligocene-Early Miocene that was characterized by subsiding basins, unconformably overlying the previously deformed belts (Calais et al., 1989). However, at the same time, collision has occurred off Venezuela (Mann et al., 1995). Our model is different from Pindell et al.'s (1988) plate model for North-South America plate motions, in that it results in a drastic increase in convergence velocity subsequent to chron 8 (25.8 Ma) (9.6 -+- 3.1 mm/year for chrons 8-6 and 9.6 + 2.1 mm/year for chrons 6-5 at 85~ compare also Figs. 10 and 11). A slowdown in convergence after chron 5 resulted in a post-chron-5 convergence rate drop to 5.2 4- 1.3 mm/year at 85~ The modeled Neogene convergence results in 92-4-30 km convergence from chron 8 to 6, 127 -1-27 km from chron 6 to 5, and 72 4- 17 km from chron 5 to the present near the North Panama Deformed Belt at 85~ (Fig. 10). The Neogene convergence measured south of the Muertos Trough at 65~ is 40 4- 13 km from chron 8 to 6, 58 4- 22 km from chron 6 to 5, and 22 4- 13 km from chron 5 to present day. Geological data from the circum-Caribbean plate boundaries indicate that the tectonic regime changed
drastically in the Early to Middle Miocene. The eastern portion of the North America-Caribbean plate boundary displays an accretionary wedge along the Muertos Trough (Case et al., 1984). Based on seismological evidence Byrne et al. (1985) showed that the Muertos Trough is an active structure and suggested it to be the site of subduction. North-south convergence is accommodated by oceanic crust underthrusting the Greater Antilles, or by folding and thrust faulting only, where the crust is thicker, as at the Beata Ridge and the Bahamas (Ladd et al., 1990). Correlation of seismic reflection data from the turbidite fill in the Muertos Trough with the Venezuelan Basin indicates a Neogene, or possibly late Neogene age for the initiation of underthrusting (Ladd et al., 1990). In the Early Miocene, the Peralta and Rio Ocoa sediment groups on Hispaniola were deformed in a southwest verging fold-and-thrust belt (Heubeck et al., 1991), supporting evidence for the onset of Early Miocene convergence within this part of the Caribbean-North American plate boundary. Tectonic events at this time have also been mapped in the Dominican Republic, in Haiti, and on Cuba (Calais et al., 1992). The Caribbean-South American plate boundary comprises a wide and complicated plate boundary zone. It starts at the deformation front of the subduction zone, includes various dextral strike-slip faults
NEW CONSTRAINTS ON THE PLATE TECTONIC EVOLUTION OF THE CARIBBEAN of northern South America (e.g. Bocono, Oca, E1 Pilar) and continues to the south as a fold-thrust belt (e.g. Ladd et al., 1984). Biju-Duval et al. (1984) analyzed multichannel seismic reflection data in the Venezuelan Basin and concluded that the present configuration of the margin, i.e. underthrusting of the Caribbean oceanic crust below the South American borderland, developed in the Early or Middle Miocene. They also realized that north-south shortening observed at the North Venezuelan margin may be the result of regional North-South America convergence. Recently, a comprehensive analysis of new and existingseismic data from the Beata Ridge and adjacent areas by Mauffret and Leroy (Chapter 21) has shown that the Beata Ridge, a Cretaceous plateau, is bounded to the east by compressive structures reactivated by right-lateral strike-slip, and by normal faults to the west. Uplift of the ridge increases from south to north, and is estimated to have started in the Early Miocene (23 Ma), resulting in a total shortening between 170 km and 240 km as a function of latitude (Mauffret and Leroy, Chapter 21). They interpret the Beata Ridge as a compressional plate boundary, resulting from overthrusting of the Colombian microplate onto the Venezuelan microplate. The implied clockwise rotation of the Colombian microplate and convergence between the latter and the Venezuelan microplate are consistent with differential convergence between the North and South American plates, increasing from east to west, thereby 'squeezing' the Colombian plate out to the east, as suggested by Burke et al. (1978). Mauffret and Leroy (Chapter 21) suggest that the observed deformation may also be caused by the buoyancy of the Cocos plate, which is subducting under the Caribbean plate (England and Wortel, 1980; Meijer, 1992). However, Central America experiences extension in the back arc in the north where the subducting Cocos plate is older than in the south, whereas a younger Cocos plate in the south causes shortening. Alternatively, compression in Costa Rica, as expressed by the April 22, 1991 Costa Rica earthquake (Plafker and Ward, 1992), has been attributed to the subduction of an aseismic ridge (Adamek et al., 1987). Mann and Burke (1984) suggested that the Beata Ridge may be the consequence of northward motion of the Maracaibo block, a tectonic block of South America. Recent work has confirmed a north- to northeast-directed motion of these blocks relative to the Caribbean plate (Ego et al., 1995). In summary it is unclear what the role of the Cocos plate may be in terms of contributing to compression at the Beata Ridge. An east-west gradient in convergence between the Americas is also supported by a recent analysis of present-day relative plate motions between North and South America based on GPS data (Dixon and
51
Mao, 1997). They found an increase of differential north-south convergence from east to west from about 1 mm/year at the Barracuda ridge to about 9 mm/year at 85~ It is interesting to note that their modeled present-day convergence rate of 9 mm/year at 85~ is similar to the rates we calculate for chron 8-5 time, and significantly faster than our post-chron-5 convergence estimate at 85~ of 5.2 -+1.3 mm/year. It is not clear whether this difference reflects a recent acceleration in North AmericaSouth America plate convergence, or whether it is related to problems in our anomaly-5 reconstruction. The latter may well be the case, since the South Atlantic reconstruction for this time is not well constrained; it is based on a small number of data points only, and Table 2 shows that we have slightly underestimated the uncertainties of both magnetic and fracture zone identifications for this data set. In contrast, the anomaly-5 reconstruction in the central North Atlantic is extremely well constrained. We propose that the post-chron-8 convergence between North and South America has also played a substantial role in the Neogene Panama Arc collision and subsequent arc deformation to an S-shaped pattern. In the collision area, the computed northsouth convergence between the Americas resulted in a total of 291 -+- 75 km of north-south convergence. During the Middle to Late Miocene the onset of the collision of the Costa Rica-Panama Arc with the western Cordillera of northwestern South America (Wadge and Burke, 1983; Eva et al., 1989; Mann et al., 1990) started forming the North Panama Deformed Belt. The tectonic events outlined above are all slightly younger (mostly Early Miocene) than the onset of North-South America convergence predicted by our model (~26 Ma, Late Oligocene). This may reflect an artifact of our model. Our plate model lacks resolution between anomalies 8 (25.8 Ma) and 6 (19.0 Ma), because it is not straightforward to identify the magnetic anomalies between 6 and 8 with confidence in a slowly spreading tectonic regime. It is possible that the Early Miocene tectonic events in the Caribbean correspond to a global change in plate motions. Evidence for this idea comes from a detailed survey of the Pitman Fracture Zone in the South Pacific that shows a distinct change in spreading direction at chron 6c (Cande et al., 1995) which represents the Oligocene-Miocene boundary (23.8 Ma). It is virtually impossible to identify this magnetic anomaly in the slowly spreading central North Atlantic and South Atlantic oceans. Therefore, we consider it possible that convergence started at the Oligocene-Miocene boundary, 2 m.y. later than predicted by our plate model. Pindell et al.'s (1988) model results in an acceleration in convergence at chron 6 (20 Ma in the DNAG timescale
52 (Kent and Gradstein, 1986), 19 Ma in the timescale used here (Cande and Kent, 1995), with extremely slow convergence from chron 13 to chron 6. This demonstrates that a model not constrained by any magnetic anomaly identification between anomaly 6 and 13 (a time interval 14 m.y. long) results in an apparent acceleration in convergence at 19 Ma, about 5 m.y. later than the Oligocene-Miocene boundary. Plate motions relative to the mantle
We put Caribbean plate motions into an absolute hotspot reference flame based on AtlanticIndian ocean hotspot tracks (Mtiller et al., 1993) for understanding cause and effects of plate motions between the Americas and the Caribbean plate(s). Ross and Scotese (1988) used a paleomagnetic reference flame for their model (which cannot resolve longitudinal motions of plates), and correspondingly do not show a geographic flame on their reconstructions. Pindell et al. (1988) used the absolute plate motion model by Engebretson (1982) and Engebretson et al. (1985) for the Pacific based on hotspot tracks to calculate Pacific-Caribbean relative motions, which have likely exerted controls on Caribbean tectonic evolution in the Mesozoic and early Tertiary (Pindell et al., 1988), but not necessarily in the Neogene. In any case, Pacific absolute plate motions are only of limited use to constrain absolute motions of plates bordering the Atlantic Ocean, as our knowledge on closing plate circuits crossing the boundary between East and West Antarctica is still inadequate (Molnar and Stock, 1987; Cande et al., 1995). We use the relative plate motion model for the central North Atlantic and South Atlantic presented here, a revised relative plate motion model for the Caribbean area, largely based on the tectonic elements and plate hierarchy from Ross and Scotese (1988), and the model for motion of plates in the Atlantic and Indian Ocean Hemisphere relative to major hotspots (Mtiller et al., 1993). The combined rotation model has been adapted to the Cande and Kent (1995) timescale for post-chron-34 (83 Ma) times and the Gradstein et al. (1994) timescale for earlier times. The absolute plate motion model by MiJller et al. (1993) is based on jointly fitting dated hotspot tracks on the Australian, Indian, African, and North and South American plates relative to present-day hotspots assumed fixed in the mantle. Therefore this model is better constrained than a model solely based on the hotspot tracks of one plate. Pindell et al. (1988) suggested that most of the total opening by seafloor spreading between the two Americas was accomplished some time between 100 and 90 Ma, when the Caribbean plate started entering from the west. However, the age of the
R.D. MfJLLER et al. initial contact of a Caribbean plate originating from the Pacific has been revised to late CampanianMaastrichtian, when syn-orogenic sedimentation and northward verging folding, thrusting and obduction of ophiolites have occurred at the southern margin of the Yucat~in Peninsula in Guatemala (Rosenfeld, 1990). The arguments in favor of an allochthonous nature of the Caribbean oceanic crust have been reviewed comprehensively by Pindell and Barrett (1990). In the Paleocene the Yucat~in Basin opened along a left-lateral strike-slip fault (Rosencrantz, 1990) and the prograding arc started colliding with the Bahamas Platform in western Cuba (Bralower et al., 1993). For the time after the Middle Eocene we have an estimate for North America-Caribbean plate motions from the spreading history in the Cayman Trough (Rosencrantz et al., 1988), even though spreading here may not reflect the total North America-Caribbean motion (Rosencrantz and Mann, 1991). Burke et al. (1980) suggested that cumulative offsets of strike-slip faults on Jamaica suggest a minimum rate of offset of 4 mm/year, in addition to an average of 16 mm/year of total opening in the Cayman Trough. We implemented this suggestion in our rotation model, similar to Ross and Scotese (1988), by allowing for 4-6 mm/year of left-lateral strike-slip between Jamaica and southern Hispaniola. In Fig. 11 we show the resulting path of two points attached to the Caribbean plate relative to the mantle (without considering relative motion between the Colombian and Venezuelan microplates through time, which we cannot reconstruct). The two absolute plate motion paths in Fig. 11 as well as the plate reconstructions in Fig. 13b-d show that the Caribbean plate has been virtually stationary with respect to the mantle at least since the onset of seafloor spreading in the Cayman Trough. The errors from combining the 'absolute' and relative plate motions models involved in this calculation are probably larger than the total length of the path shown. The Caribbean plate could have only maintained a substantial eastward component of motion if either the Cayman Trough opened much later and faster than presently assumed, and/or if there has been much faster strike-slip between the Caribbean plate and Jamaica than suggested by Burke et al. (1980). North America's and South America's plate motions in the mantle reference flame are both characterized by relatively fast westward motion, with a small component of convergence added at chron 13 due to a clockwise change in South American plate motion, and even faster convergence after chron 8 due to a counterclockwise change in North America absolute plate motions. In contrast, the Caribbean plate appears to have been virtually stationary in a mantle reference flame at least since chron 18. This
NEW CONSTRAINTS ON THE PLATE TECTONIC EVOLUTION OF THE CARIBBEAN result agrees with an idea put forward by Sykes et al. (1982), who noted that only a small fraction of its perimeter is attached to a subducting slab. Even though the Caribbean slab under the South American plate has been shown to be longer than previously thought, the forces assumed to be most important for driving plates, namely ridge push, slab pull and trench suction (this force acts to draw plates together at a trench; Elsasser, 1971) must be relatively small. If they were not, then the Caribbean plate would not rest in a mantle reference frame, as found by our plate kinematic analysis. Our result is also in accordance with the analysis by Gripp and Gordon (1990) of present-day Caribbean plate motions with respect to the hotspots. Their analysis, based on motion of the Pacific plate relative to its underlying hotspots, and the NUVEL-1 relative motion model by DeMets et al. (1990) results in roughly west-southwest-oriented motion of the Caribbean plate. However, their motion vectors do not differ significantly from zero. It must be concluded that tectonic plate boundary processes between the Caribbean plate and the Americas are entirely driven by relatively fast, mostly westward motion of North and South America. The resulting differential motion between North and South America affects a stationary Caribbean plate trapped between two larger plates by edge-driven plate tectonic interactions, equivalent to some small plates in the Middle East (e.g. Arabia/Anatolian plate; McKenzie, 1972). Sykes et al. (1982) recognized this possibility, but suggested alternatively that the Caribbean plate may be forced to move eastward in response to the gradient in convergence rate between North and South America, increasing from east to west, as also found by our analysis. In contrast, the plate motion paths plotted in Fig. 11 suggest that the eastward motion of the Caribbean plate with respect to the two Americas is entirely due to westward motion of the latter two plates with respect to the mantle, and that the east-west convergence gradient quoted by Sykes et al. (1982), which has been constrained to post-chron-8 (25.8 Ma) times by our model (probably post-chron 6c as discussed above), may not have resulted in substantial eastward motion of the Caribbean plate with respect to the mantle. Rather, the east-west gradient in post-chron-6c convergence may have contributed to causing east-west compression at the Beata Ridge, as described in Mauffret and Leroy (Chapter 21). Our combination of relative and absolute plate motions indicates that throughout the Tertiary tectonic processes at the northern and southern boundary of the Caribbean plate were governed by the relatively fast westward motion of both the North and South American plates with respect to a nearly
53
stationary Caribbean plate. The differences between North America-Africa and South America-Africa plate motions, as described here, resulted in changes in relative motion between the two Americas whose effects are clearly seen in the tectonic development along the northern and southern margin of the Caribbean area. Since the Caribbean plate does not appear to have moved substantially relative to the mantle during the Neogene, there are no major tectonic processes which can be attributed to the eastward 'escape' of the Caribbean plate during this time (e.g. Mann, 1997). In particular for the time since chron 6 (19 Ma), for which Caribbean-North America relative motion is better constrained than for earlier times, we find that the Caribbean plate was virtually fixed relative to the mantle. This observation suggests that accelerated convergence post-chron 6c (23.8 Ma) at the OligoceneMiocene boundary reduced the space within the eastward 'escaping' arc could operate such that Caribbean plate motion relative to the mantle ceased. It follows that most deformation at the northern and southern Caribbean plate boundaries in the Miocene and younger was entirely governed by changes in absolute plate motion of the North American and South American plates, and the resulting motion relative to the Caribbean plate. While strike-slip along the northern and southern Caribbean margins continued since the east-west component of absolute plate motion of the Americas was far larger than the northsouth components, the magnitude of the latter increased, probably at the Oligocene-Miocene boundary, resulting in convergence between the two Americas. The rate of convergence increased from east to west, resulting in an eastward-directed 'squeeze' on the Caribbean plate which caused its breakup along the Beata Ridge, where east-west-oriented compressional stresses are absorbed (Mauffret and Leroy, Chapter 21). Mann et al. (1995) show a model for the formation of Caribbean microplates in six stages from the Maastrichtian to present-day in a fixed South American framework. Their figures show that the Caribbean plate has moved eastward by about 800 km since the mid-Oligocene (relative to South America). Mann et al. (1995) reason that collision ceased in the Middle Eocene in central Cuba since the arc could advance no further to the north-northeast above the Bahamas Platform, and that this event rotated Caribbean plate motion clockwise in a more easterly direction. They favor the 'tectonic escape' mechanism proposed by Burke and Seng6r (1986), which results in the motion of a colliding plate towards the remaining 'free face', e.g. an island arc. In case of the Caribbean in Middle Eocene times, the remaining free face would have been towards the east, i.e. the Lesser Antilles Arc. This argument is
54
R.D. MULLER et al. America t:J,'
20 ~
Caribbean plate
15 ~
stationary
% 10 ~
(a)
America
Chron 18, 38.4 Early Eocene
3 crr~/a
America
d"
20 ~
Caribbean plate
15 ~
stationary
10 ~
South America
(b)
2.15 cnga
Chron 8, 25.8 Ma Late Oligocene Platform 20 ~
Gonave Venezuela plate stationary
15 ~
Colombia plate 10 ~
Venezuela South America Cocos
(C)
Plate
1.3 Cm/a
Chron 5, 9.7 Ma Late Miocene 00~ _90 ~
_85 ~
_80 ~
-75 ~
.70 ~
-65*
-60*
NEW CONSTRAINTS ON THE PLATE TECTONIC EVOLUTION OF THE CARIBBEAN plausible. However, if we put Caribbean plate reconstructions in the Atlantic-Indian hotspot framework (Fig. 13), it appears that the eastward motion of the Caribbean plate had ceased at Middle Eocene times. it follows that the apparent continuing apparent 'escape' of an arc system as described by Royden (1993), e.g. for the Scotia Arc between South America and the Antarctic plate, may not necessarily involve the absolute motion of a small plate (e.g. the Scotia Sea plate) relative to the mantle. A retreating subduction boundary may be initiated by the change in polarity of a subduction system, as in the case of the proto-Caribbean, but the Caribbean plate never reached the 'open ocean' as in the case of the Scotia Sea (Royden, 1993), since its eastward migration was inhibited by boundary forces to the north and south, due to progressive convergence between the two Americas. As a result, Caribbean absolute plate motion stopped. Subsequently, all relative motion observed between the Americas and the Caribbean plate and associated tectonic elements has been caused by the absolute plate motion of North America and South America relative to a stationary Caribbean plate.
CONCLUSIONS
New gravity anomaly data from satellite altimetry and new magnetic data allow us to construct a modified plate model for plate motions between the two Americas, and calculate its uncertainties. For the N o r t h - S o u t h America plate boundary area east of the Lesser Antilles Arc our results are in good agreement with the observed strong plate de-
55
formation of the oceanic crust at the Barracuda and Tiburon ridges. MUller and Smith (1993) inverted Bouguer anomalies for crustal layer structure, and found that the Moho is uplifted 2 - 4 km over short wavelengths ( ~ 7 0 km) at the Barracuda and Tiburon ridges, implying large anelastic strains and an unstable density distribution. Together with the plate model presented here, these results indicate that much of the unusually shallow Moho topography and crustal uplift of the Tiburon Rise is a result of North A m e r i c a - S o u t h America convergence after chron 6 (19.0 Ma). This model is in contrast with Dolan et al.'s (1989, 1990) suggestion that the present topography of the Tiburon Rise has existed since the Late Cretaceous. Our results suggest that slow sinistral transtension/strike-slip between the two Americas lasted from chron 34 (83 Ma) until chron 25 (55.9 Ma), followed by roughly northeast-southwest-oriented convergence until chron 18 (38.4 Ma). This first convergent phase correlates with a P a l e o c e n e - E a r l y Eocene calc-alkaline magmatic stage in the Greater Antilles, which is thought to be related to southward subduction of proto-Caribbean crust during this time. Relatively slow transpression until chron 8 is followed by a drastic increase in convergence velocity. Subsequent to chron 8 (25.8 Ma), probably at the O l i g o c e n e - M i o c e n e boundary, fast convergence resulted in 92 4- 22 km convergence from chron 8 to 6, 127 • 25 km from chron 6 to 5, and 72 4- 17 km from chron 5 to the present measured at 11~ 85~ near the North Panama Deformed Belt. The Neogene convergence measured at the eastern Muertos Trough, at 17.5~ 65~ is 41 + 18 km from chron 8 to 6, 58 4- 25 km from chron 6 to 5, and 22 4- 17 km from chron 5 to present day. The modeled conver-
Fig. 13. Plate reconstructions of the Caribbean area in an Atlantic-Indian mantle reference system for chrons 18 (38.4 Ma), 8 (25.8 Ma), and 5 (9.7 Ma). See Fig. 2 for labels of tectonic elements. Bold arrows show North American and South American plate motion relative to the mantle for the time intervals from 38.4 to 25.8 Ma (a), 25.8 to 9.7 Ma (b), and 9.7 Ma to present day (c). The eastward escape of the Caribbean plate had ceased when the opening of the Cayman Trough started (a). This time corresponds to the collision in central Cuba which prevented a further advance of the Caribbean plate to the north-northeast above the Bahamas Platform. Only if the Cayman Trough would have opened later and spread considerably faster than interpreted by Rosencrantz et al. (1988), and/or if contemporaneous strike-slip between Jamaica and the Caribbean plate was considerably faster than suggested by Burke et al. (1980), would the Caribbean plate have maintained any eastward-directed motion relative to the mantle after the Middle Eocene. (b, c) Relatively slow convergence between the Americas from chron 18 (38.4 Ma) to chron 8 (25.8 Ma) was followed by rapid convergence after chron 8, probably starting at the Oligocene-Miocene boundary, averaging 9.6 mm/year until chron 5 (9.7 Ma), slowing down to 5.2 mm/year after chron 5. Accelerated convergence was caused by a counterclockwise change in the absolute plate motion direction of North America. As a result, the total area available for the western Caribbean plate at 85~ was reduced by at least "~230 km in north-south direction in the last 25 m.y. We suggest that about half of the north-south extent (500 km) of the Maracaibo slab (subducted Caribbean oceanic plateau crust) under the South American continent (van der Hilst and Mann, 1994) may have resulted from post-Oligocene South AmericaCaribbean convergence. At least since chron 18 (38.4 Ma) Cocos plate-Caribbean interactions have not resulted in any substantial motion of the Caribbean plate relative to the mantle. Continuing oblique collision along the passive margin of eastern Venezuela (Algar and Pindell, 1993) must be attributed to the west-northwestward motion of South America relative to the mantle and relative to a stationary Venezuelan plate, rather than to continuing eastward movement of the Caribbean plate. Equivalently, the Miocene and younger transpression observed in Hispaniola (Heubeck and Mann, 1991) due to collision of arc rocks with the Bahamas Platform is the result of continuing westward motion of the North American plate (and the Bahamas Platform) relative to a stationary Venezuelan plate in a mantle reference frame, rather than a continuing eastward 'escape' of the Caribbean plate. The Gonave microplate has been transferred from the Venezuelan plate to North America in the Pliocene (Mann et al., 1995), leaving the rest of the Venezuelan plate behind.
56 gence may correspond to the Early Miocene onset of underthrusting of the Caribbean oceanic crust below the South American borderland in the Colombian and Venezuelan basins, the onset of subduction in the Muertos Trough, and folding and thrust faulting at the Beata Ridge and the Bahamas, and the breakup of the main part of the Caribbean plate into the Venezuelan and Colombian plates, separated by the Beata Ridge acting as a convergent plate boundary (Mauffret and Leroy, Chapter 21). The east-west shortening between the latter two plates may reflect the differential convergence between the two Americas, increasing from east to west. The main differences with Pindell et al.'s (1988) model are the following. (1) Pindell et al.'s (1988) model implies relatively constant convergence between the Americas of rates of about 5 mm/year or less since chron 21, with the exception of faster convergence between chron 6 and 5. Our model results in substantial variations in convergence rates from chron 25, as documented in Fig. 10 and Table 10. In particular, we resolve an initial phase of fast convergence between chron 8 (25.8 ma) and chron 6 (19.0 Ma) of nearly 10 mm/year, compared with less than 4 mm/year from chrons 18 to 8 measured at 85~ We suggest that most of this convergence occurred after chron 6c (23.8 Ma), which corresponds to the OligoceneMiocene boundary, a plate reorganization in the South Pacific (Cande et al., 1995), and the formation of the present-day deformed belts north and south of the Caribbean area. Without identifying magnetic anomaly 8, two stages of slow (chrons 13-8) and fast (chrons 8-6) convergence are averaged. (2) We have computed uncertainties for our North America-South American plate flow lines. Uncertainty ellipses for rotated data points are especially helpful to evaluate whether or not we can resolve relatively slow phases of relative motion. (3) We have put Caribbean plate reconstructions into the Atlantic-Indian hotspot reference system from Mtiller et al. (1993). This allows us to evaluate causes and effects of relative plate motions in the Caribbean area. St6phan et al. (1986) put forward the hypothesis that both the northern and the southern Caribbean deformed belts are the result of the bending of the Caribbean continental frame related to eastwest shortening. East-west shortening appears to have a variety of different causes. In the Panama area, east-west shortening is related to Nazca-South America convergence in that direction and collision of an east-west-oriented arc with a northsouth-oriented margin (Mann and Corrigan, 1990; Wadge and Burke, 1983). In Hispaniola, northeastsouthwest shortening is related to the interaction of the westward-moving North American plate rel-
R.D. MCILLER et al. ative to a stationary Caribbean plate as shown before. Our plate model shows well resolved north-south convergence between the Americas during the Neogene, and we argue that this plate convergence is likely the main cause for the formation of those deformed belts which cannot be attributed to east-west convergence as described above. We also suggest that the fast Neogene plate convergence between North and South America contributed to the Late Miocene onset of the collision of the Costa RicaPanama Arc with the western Cordillera of South America (Wadge and Burke, 1983; Eva et al., 1989; Mann et al., 1990; Mann and Corrigan, 1990). One of the main tectonic events affecting the Caribbean plate in the Neogene has been its breakup into the Venezuelan and Colombian plates (see Mauffret and Leroy, Chapter 21). The breakup may have been caused the observed east-west gradient in convergence between the Americas, the subduction of the buoyant Cocos plate under the Caribbean plate (Mauffret and Leroy, Chapter 21), or the northnortheastward motion of the Maracaibo block, which in turn may be related to differential North-South America convergence. By combining Atlantic-Indian hotspots as a reference frame with revised North America-African and South America-African relative plate motions, and with a revised plate model for the Caribbean area we are able to show the following. (1) The eastward escape of the Caribbean plate appears to have ceased when the opening of the Cayman Trough started. This time corresponds to the collision in central Cuba which prevented a further advance of the Caribbean plate to the northnortheast above the Bahamas Platform. Only if the Cayman Trough would have opened later and spread considerably faster than interpreted by Rosencrantz et al. (1988), and/or if contemporaneous strikeslip between Jamaica and the Caribbean plate was considerably faster than suggested by Burke et al. (1980) would the Caribbean plate have maintained any eastward-directed motion relative to the mantle after the Middle Eocene. If we rely only on the interpreted post-chron-6 (19 Ma) spreading history of the Cayman Trough by Rosencrantz et al. (1988), which is better constrained than its previous opening, then the Caribbean plate is still found to have been without any substantial motion relative to the mantle subsequent to chron 6 within the errors of absolute plate motion models. (2) It is not the case that North America-South America plate motions had only minor effects on the development of the Caribbean region after the Campanian, as suggested by Pindell et al. (1988). After the breakup of the Caribbean plate into the Venezuelan and Colombian plates, the eastward driving force
NEW CONSTRAINTS ON THE PLATE TECTONIC EVOLUTION OF THE CARIBBEAN of the latter plate may have still been derived from interactions with the Cocos plate. However, even if this is so, our model suggests that post-chron-8 (25.8 Ma) differential motion between the Americas has resulted in a total of 291 + 64 km in convergence at 85~ near the North Panama Deformed Belt. In other words, the total area available for the Colombian plate at 85~ was reduced by at least ~230 km in north-south direction in the last 25 m.y. Surely this reduction in space had a profound influence on the Colombian plate margin, and has contributed to convergence along the North Panama Deformed Belt. van der Hilst and Mann (1994) show that the subducted Maracaibo slab underlying northwestern South America extends up to 500 km from the Caribbean-South America boundary to the south. The Maracaibo slab corresponds to subducted Caribbean oceanic plateau crust. Our results suggest that about half of the north-south extent of the Maracaibo slab under the South American continent may have resulted from Miocene and later South America-Caribbean convergence, if most of the North America-South America convergence was taken up at this boundary. (3) The Caribbean plate has been trapped between two larger plates, and has been subject to edge-driven plate tectonic interactions since then. It follows that the main control on North AmericaCaribbean and South America-Caribbean plate interactions has not originated from Cocos plateCaribbean interactions, as these interactions have not resulted in any substantial motion of the Caribbean plate relative to the mantle. Continuing oblique collision along the passive margin of eastern Venezuela as reported by Algar and Pindell (1993) must be attributed to the west-northwestward motion of South America relative to the mantle and relative to a stationary Venezuelan plate, rather than to continuing eastward movement of the Caribbean plate. Equivalently, the Miocene and younger transpression observed in Hispaniola (Heubeck and Mann, 1991) due to collision of arc rocks with the Bahamas Platform is the result of continuing westward motion of the North American plate (and the Bahamas Platform) relative to an approximately stationary Venezuelan plate in a mantle reference frame, rather than a continuing eastward 'escape' of the Caribbean plate. By the same token, the Gonave microplate, which has been detached from the Caribbean microplate in the Pliocene and accreted to North America has not been 'left behind' (Mann et al., 1995), but has rather been transferred from a stationary Venezuelan plate to North America, leaving the rest of the Venezuelan plate behind. (4) The Tertiary tectonic history of the Caribbean plate can be described by 'tectonic escape' up to the Middle Eocene. Subsequently the Caribbean
57
plate came to a halt in the Atlantic-Indian mantle reference system due to a progressive reduction in space between the two Americas for the arc to move eastward and due to its collision with the Bahamas Platform. We show that the most severe reduction in space started at the Oligocene-Miocene boundary, resulting in the gradual formation of many of today's tectonic elements and sedimentary basins in the Caribbean area.
ACKNOWLEDGEMENTS
The contents of this paper have been clarified substantially by Jim Pindell's comprehensive review of an early draft. Given that we are no experts on Caribbean geology, our discussion and the interpretation of our results has benefited from many discussions with Eric Calais and Alain Mauffret, as well as from thorough reviews by Jan Golonka, Mark Gordon, and Ian Norton. We thank Paul Mann for his encouragement to make a contribution to this volume. JYR acknowledges support from the Centre Nationale de la Recherche Scientifique (CNRS) that enabled his visit at the University of Sydney, and from the Plan Nationale de Teledetection Spatiale. UMR 6526 G6osciences Azur contribution 139, Geological Survey of Canada contribution 1997027.
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Rosenfeld, J.H., 1990. Sedimentary rocks of the Santa Cruz o p h i o l i t e - a proto-Caribbean history. Trans. 12th Caribbean Geol. Conf., U.S. Virgin Islands, pp. 513-519. Ross, M.I. and Scotese, C.R., 1988. A hierarchical tectonic model of the Gulf of Mexico and Caribbean region. Tectonophysics, 155: 139-168. Royden, L.H., 1993. Evolution of retreating subduction boundaries formed during continental collision. Tectonics, 12: 629638. Royer, J.-Y. and Chang, T., 1991. Evidence for relative motions between the Indian and Australian plates during the last 20 Myr from plate tectonic reconstructions: implications for the deformation of the Indo-Australian plate. J. Geophys. Res., 96: 11,779-11,802. Royer, J.-Y., Gordon, R.G., DeMets, C. and Vogt, ER., 1997. New limits on the motion between India and Australia since chron 5 (11 Ma) and implications for lithospheric deformation in the equatorial Indian Ocean. Geophys. J. Int., 129: 41-74. Sandwell, D.T. and Smith, W.H.E, 1997. Marine gravity anomaly from Geosat and ERS-1 satellite altimetry. J. Geophys. Res., 102:10,039-10,054. Shaw, ER. and Cande, S.C., 1990. High-resolution inversion for South Atlantic plate kinematics using joint altimeter and magnetic anomaly data. J. Geophys. Res., 95: 2625-2644. Stdphan, J.F., Blanchet, R. and Mercier de L6pinay, B., 1986. Northern and southern Caribbean festoons (Panama, Colombian-Venezuela and Hispaniola-Puerto Rico), interpreted as pseudosubdivisions induced by the east-west shortening of the peri-Caribbean continental frame. In: WezelForese, C. (Ed.), The Origin of Arcs. Developments in Geotectonics, vol. 21, Elsevier, Amsterdam, pp. 401-422. Stdphan, J.F., Mercier De Ldpinay, B., Calais, E., Tardy, M., Beck, Ch., Carfantan, J.-Ch., Olivet, J.-L., Vila, J.-M., Bouysse, Ph., Mauffret, A., Bourgois, J., Thery, J.-M., Tournon, J., Blanchet, R. and Dercourt, J., 1990. Paleogeodynamic maps of the Caribbean: 14 steps from Lias to Present. Bull. Soc. G6ol. Fr., 8: 915-919. Stock, J. and Molnar, E, 1983. Some geometrical aspects of uncertainties in combined plate reconstructions. Geology, 11: 697-701. Sykes, L.R., McCann, W.R. and Kafka, A.L., 1982. Motion of Caribbean plate during last 7 million years and implications for earlier Cenozoic movements. J. Geophys. Res., 87: 10,656-10,676. Tucholke, B.E. and Schouten, H., 1988. Kane fracture zone. Mar. Geophys. Res., 10: 1-39. Van der Hilst, R. and Mann, E, 1994. Tectonic implications of tomographic images of subducted lithosphere beneath northwestern South America. Geology, 22:451-454. Wadge, G. and Burke, K., 1983. Neogene Caribbean plate rotation and associated tectonic evolution. Tectonics, 2: 633643.
Chapter 3
Jurassic-Early Cretaceous Tectono-Paleogeographic Evolution of the Southeastern Gulf of Mexico Basin
G Y O R G Y L. M A R T O N and R I C H A R D T. B U F F L E R
A new opening model for the Gulf of Mexico basin provides a framework in which the Jurassic-Early Cretaceous tectono-paleogeographic evolution of the southeastern Gulf of Mexico and surrounding regions can be discussed. A detailed analysis of available seismic data and the results of DSDP Leg 77 define four major tectono-stratigraphic sequences bounded by major unconformity surfaces: crystalline basement, Paleozoic(?) pre-rift rocks, a Late Jurassic syn-rift sequence, and an Early Cretaceous post-rift sequence. The pre-rift rocks are interpreted to represent a pre-Mesozoic (Late Paleozoic?) sedimentary cycle. The Late Jurassic syn-rift sequence in the central continental domain of the southeastern Gulf of Mexico occurs in grabens or half-grabens and is interpreted to consist of two units, a lower non-marine unit overlain by a marine carbonate unit consisting of carbonate buildups (platforms) and adjacent deeper marine sediments. Jurassic rocks are absent over high-standing blocks as well as the adjacent Yucatfin and Florida blocks. The Lower Cretaceous post-rift sequence drapes the entire area and consists of deep-water carbonate sediments in the central basin flanked by shallow-water platforms on the adjacent Yucatan and Florida blocks. Six tectono-paleogeographic maps covering the eastern Gulf of Mexico and northwestern Cuba (palinspastically restored to the southeastern margin of Yucatan) document the evolution of the area. During the late Middle Jurassic (Callovian) the southeastern Gulf was a bridge between Yucatan and Florida, separating an area of widespread extension and salt deposition to the north in the Gulf of Mexico from another area of extension and clastic sedimentation to the south between Yucatan and northern South America. By Oxfordian time Yucatfin had rotated 11~ counter-clockwise and major continental rifting and non-marine sedimentation in rift basins had begun all along the southeastern Gulf. To the north salt deposition had ceased and a major marine transgression culminated in deposition of Smackover carbonates. South of Yucatan shallow-water carbonate sedimentation also prevailed. During Kimmeridgian, Tithonian and into earliest Cretaceous time, rifting continued in the southeastern Gulf as Yucatan continued to rotate counter-clockwise. As the basin subsided a marine seaway, characterized by shallow-water carbonate platforms on high-standing blocks, became established, connecting the Gulf of Mexico with the proto-Caribbean. In the northeastern Gulf the mixed clastic/carbonate Haynesville and Cotton Valley sequences were deposited, while to the south of Yucatan shallow-water carbonate sedimentation gave way to deeper water sedimentation as the margin subsided. In late Berriasian spreading ceased in the Gulf of Mexico, Yucatan reached its present-day position, and rifting stopped in the southeastern Gulf. Carbonate platforms atop rift blocks drowned as the basin subsided and sea level rose, and the southeastern Gulf became the deep-water seaway that it is today. The marine transgression reached the Yucatan and Florida blocks, where extensive carbonate platforms became established and flourished throughout the Early Cretaceous. To the south of Yucatan, deep-water pelagic sedimentation continued throughout the Early Cretaceous.
INTRODUCTION T h e s o u t h e a s t e r n G u l f of M e x i c o (Fig. 1) has b e e n a s e a w a y c o n n e c t i n g the G u l f of M e x i c o with the C a r i b b e a n since the L a t e Jurassic. It has had a long and c o m p l e x g e o l o g i c a l history b e g i n n i n g with a Jurassic t h r o u g h E a r l y C r e t a c e o u s passive
It was during the L a t e Jurassic rift stage that the m a i n tectonic e l e m e n t s f o r m e d , w h i c h c o n t r o l l e d s u b s e q u e n t C r e t a c e o u s and C e n o z o i c s e d i m e n t a t i o n patterns and p a l e o c e a n o g r a p h i c conditions. T h e p r i m a r y p u r p o s e of this p a p e r is to s u m m a rize the Jurassic t h r o u g h E a r l y C r e t a c e o u s tectonop a l e o g e o g r a p h i c e v o l u t i o n of the s o u t h e a s t e r n G u l f
m a r g i n e v o l u t i o n that can be tied directly to the b r e a k u p of w e s t e r n P a n g e a . This early e v o l u t i o n is
of M e x i c o b a s e d on an analysis of r e g i o n a l m u l t i f o l d s e i s m i c data and D S D P drilling results. T h e s e data
c h a r a c t e r i z e d by pre-rift, rift, and post-rift stages.
h a v e b e e n u s e d to define four m a j o r tectono-strati-
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 63-91. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
64
G.L. MARTON and R.T. BUFFLER
graphic sequences bounded by major unconformity surfaces: crystalline basement, Paleozoic(?) pre-rift
PHYSIOGRAPHIC/GEOLOGIC SETTING
rocks, a Late Jurassic rift sequence, and an Early Cretaceous post-rift sequence. Emphasis here will be
The southeastern Gulf of Mexico is a deep seaway at the intersection of the northern Yucat~in Straits and the western Straits of Florida (Fig. 1) (Marton and Buffler, 1994; Marton, 1995). It lies just north of Cuba in between the steep Campeche and Florida escarpments, which reflect the paleotopography of Lower Cretaceous carbonate margins. The northern part of the area is extremely fiat, an expression of the distal Mississippi Fan turbidite plain (Florida Plain). Most of this flat area is underlain by Mesozoic oceanic crust (Marton, 1995). To the south the seafloor rises above the turbidite plain to a broad area dissected by several prominent, northwest-trending erosional channels (Fig. 1). The major channel bifurcates to the south, with one arm extending into the Yucat~in Straits and one arm extending into the Straits of Florida. Seismic data provide evidence that this relatively higher standing area is underlain by more or less extended continental crust (Schlager et al., 1984; Marton, 1995). Cenozoic and Upper Cretaceous sediments are thin or absent over the central part of this platform. This overall shallow bathymetric high
on the Late Jurassic rift sequence, as new interpretations of these rocks are presented. This important Late Jurassic sequence has no tectonic equivalent around the Gulf of Mexico and Caribbean basin. The Early Cretaceous rocks have been the subject of earlier studies, and they will be discussed as appropriate to complete the picture. The evolution of the southeastern Gulf of Mexico area is summarized using a set of regional tectono-paleogeographic maps that includes the eastern Gulf of Mexico as well as northwestern Cuba. The details of this history, however, cannot be told without first considering the regional tectonic setting and the tectonic events that have influenced the region. A revised tectonic model for the Gulf of Mexico basin as well as the southeastern Gulf has been set forth in earlier works by the authors, and it is reviewed briefly herein. The reader is referred to these references for additional background and details (Marton and Buffler, 1994; Marton, 1995).
Fig. 1. Map of the southeastern Gulf of Mexico study area showing major physiographic features. Also shown are the tectonostratigraphic provinces in the Cordillera de Guaniguanico of western Cuba (1 = Sierra de los Organos; 2 = Sierra del Rosario meridional; 3 = Sierra del Rosario septentrional). Water depth in meters.
EVOLUTION OF THE SOUTHEASTERN GULF OF MEXICO BASIN and the linear northwest-trending features along the northern margin are believed to be an expression of the underlying shallow basement configuration and Jurassic tectonic elements of the region (Schlager et al., 1984; Marton, 1995). The prominent reentrant in the Campeche Escarpment (Catoche Tongue) also reflects an underlying Jurassic graben structure (Shaub, 1983). The steep margin along the north coast of Cuba is the result of the collision between a Cretaceous volcanic arc system and North America, which culminated in Early to Middle Eocene time (Angstadt, 1983; Angstadt et al., 1985). Conspicuous in the area are several large knolls that probably represent continental basement highs covered with thin sediments (Catoche Knoll) or basement highs capped with Lower Cretaceous platforms or atolls (Jordan and Pinar del Rfo) (Bryant et al., 1969; Schlager et al., 1984; Marton, 1995) (Fig. 1).
65
ton, 1995). Late Jurassic rifling characterized by generally east-west extension occurred contemporaneously with seafloor spreading and ocean crust formation to the north in the central Gulf of Mexico. When oceanic crust formation stopped, i.e., Yucatan reached its present-day position, rifting in the southeastern Gulf of Mexico between Yucatan and Florida also ceased. This important event is marked by a prominent post-rift unconformity in the southeastern Gulf of Mexico. Dating of this unconformity, based on the results of DSDP drilling and interpretation of seismic data, is established as earliest Cretaceous. The regional-scale model, therefore, is vital for explaining the tectonic processes responsible for the formation of the southeastern Gulf of Mexico, and it provides the tectonic framework for discussing the paleogeographic evolution of the region.
REFLECTION SEISMIC DATA REGIONAL TECTONIC FRAMEWORK The authors' view of the Jurassic development of the Gulf of Mexico basin has been discussed at length elsewhere (Marton and Buffler, 1994; Marton, 1995). The preferred opening model is a new two-stage model constrained by a refined definition of oceanic crust and by the known kinematic framework of the large continental blocks surrounding the Gulf of Mexico basin (North American plate and Afro-South American plate) (Fig. 2). During the Late Triassic-Early Jurassic(?) to late Middle Jurassic (Callovian) syn-rift stage, the relatively stable Yucatan block translated southeastward along a major transform zone in eastern Mexico (Fig. 2A,B). This motion accommodated a large amount of extension in the area of the future northern Gulf. At the same time the Florida-Bahamas block extended also in a southeast direction to form a series of basins and arches, partly accommodated by a postulated major shear zone, the Bahamas Fracture Zone (BFZ) (Fig. 2A,B). A rotation pole for the Yucatan block in the southeastern Gulf of Mexico (23.18~ 84.24~ is proposed for the late Middle Jurassic (Callovian) to earliest Cretaceous (Berriasian) drifting stage (Fig. 2B-D). Around this pole the Yucatan block rotated about 42 ~ counter-clockwise out from the northern Gulf to accommodate the newly formed oceanic crust in the basin. The implications of this regional-scale model bear directly on the Jurassic evolution of the southeastern part of the basin. The inferred opening motion resulted in a southward propagating rift/spreading center in the eastern Gulf of Mexico (Fig. 2). Just north of the rotation pole, a tectonic setting occurred, which has no counterpart elsewhere in the Gulf of Mexico (Marton and Buffler, 1994; Mar-
The seismic data for the regional study of the entire eastern Gulf of Mexico consist of 6380 km of 12-24 fold marine seismic lines collected between 1977 and 1983 (Marton, 1995). A map showing these data is presented here as Fig. 3. A more detailed map showing the subset of seismic data used in the southeastern Gulf plus the location of figures is included in Fig. 4. The earliest data set was collected in 1977-78 and was designed to explore the Catoche Tongue as well as produce a regional overview of the southeastern and eastern Gulf (Catoche Grid (CATG) and Gulf Tectonics (GT) lines; Fig. 3). The second period of data acquisition during the first half of 1980 resulted in a higher-quality data set and provided better resolution (Straits of Florida (SF) lines; Fig. 3). These lines were designed to collect a comparatively dense grid over the central part of the area to prepare the drilling sites for DSDP Leg 77, and after the drilling, to provide the data base for further seismic stratigraphic studies and to extrapolate regionally the drilling results (Fig. 4). In a third period of data acquisition in 1983 the Mississippi Canyon (MC) lines were collected, while a fourth set of data (GULFREX-Gulf) were provided by Chevron Corporation as paper copies only. The seismic data coverage is uneven, with the highest concentration being in the southeastern Gulf (Figs. 3 and 4). The University of Texas seismic data base was originally processed at the University of Texas Institute for Geophysics (then Marine Science Institute) immediately following the acquisition of the data, and it is available on field tapes (in demultiplexed form), on final stack tapes and on films at high vertical exaggeration. In this form the data were used to study the Cretaceous and younger sediments, which occur in
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E V O L U T I O N OF THE S O U T H E A S T E R N G U L F OF M E X I C O BASIN
67
Fig. 3. Location of available seismic data in the eastern Gulf of Mexico. M C = Mississippi Canyon; SF = Straits of Florida; G T = Gulf tectonics; CATG -- Catoche Grid; Gulfrex = Chevron data. Also shown are the location of two studies used for construction of tectono-paleogeographic maps presented in Fig. 14 (Dobson, 1990; DeBalko, 1991). the s h a l l o w e r , t e c t o n i c a l l y u n d i s t u r b e d p a r t o f the
T w o d i f f e r e n t a p p r o a c h e s w e r e u s e d as p a r t o f
s e i s m i c s e c t i o n s . H o w e v e r , d u e to the i n c r e a s i n g o v e r -
this s t u d y to i m p r o v e the q u a l i t y o f the a v a i l a b l e
b u r d e n a n d t e c t o n i c c o m p l e x i t i e s , the o l d e r s e c t i o n
s e i s m i c d a t a a n d to i m p r o v e r e s o l u t i o n at d e p t h : (A)
( J u r a s s i c a n d b a s e m e n t ) is n o t s u f f i c i e n t l y r e s o l v e d .
a c o m p l e t e r e p r o c e s s i n g o f 1070 k m o f s e i s m i c d a t a
Fig. 2. Four-stage evolution of the Gulf of Mexico basin and adjacent areas: (A) Early Jurassic; (B) 166 Ma (Callovian)" (C) 160 Ma (Oxfordian); (D) 140 Ma (Berriasian) (see text for discussion, and Marton and Buffler, 1994 and Marton, 1995, for more details). A = Africa; A F B = Appalachian foldbelt; B F -- La Babia fault; B F Z = Bahamas fracture zone; BP = Blake plateau; CA -- Central Atlantic; DP = Demarara plateau; G M -- Gulf of Mexico; GP = Guyana plateau; L U = Llano uplift; M F B -- Marathon folded belt; M g r = Mid-Gulf ridge; M S M = Mojave-Sonora megashear; OFB = Ouachita foldbelt; P C = proto-Caribbean; SA = South America; S M F -San Marcos fault; T M V = Trans-Mexican volcanic belt; URB -- undifferentiated rift basins; YUC -- Yucatan block; W M T = western main transform. Light and dark stippled areas in northern Gulf are large-scale highs and basins, respectively. Solid lines are major faults and foldbelts. Dashed lines outline areas of Late Triassic-Early Jurassic sediments and volcanics. Hachured pattern outlines basement terranes in Mexico. Gray line outlines distribution of late Middle Jurassic salt (Louann and equivalent). P in southeastern Gulf is rotation pole for the Yucatan block with degrees of rotation.
68
G.L. MARTON and R.T. BUFFLER
Fig. 4. Seismic grid used in the study of the southeastern Gulf of Mexico, including locations of Figs. 5 (SF-15), 6 (GT3-75), 7 (SF-2), 8 (SF-6), 9 (SF-7), 10 (SF-13), 11 (SF-17), and 12 (SF-11), and DSDP Leg 77 drill sites 535, 536, 537, 538A, and 540. P is location of pole of rotation for Late Jurassic opening of the Gulf of Mexico basin (after Marton and Buffler, 1994; Marton, 1995).
that represent key NW-SE- and NE-SW-oriented sections and form a comprehensive grid over the southeastern Gulf, and (B) time migration of 3200 km of data, the remaining stacked seismic data in the southeastern Gulf. This reprocessed data set includes all of the CATG, GT and SF lines south of approximately 25~ (Figs. 3 and 4). The remaining 2100 km of seismic data (MC lines and GT lines north of approximately 25~ were not reprocessed or migrated, but they also were used in interpretation and mapping.
The results at each site are discussed in detail in Buffler et al. (1984) and are described briefly below, with emphasis on observations directly bearing on the early Mesozoic evolution of the area. Sites 535/540 (basin sites)
DSDP LEG 77 RESULTS
Sites 535 and 540 were drilled to sample the thick Mesozoic and Cenozoic basin fill sediments in the eastern part of the study area (Fig. 4). The combined holes at these two sites sampled an almost complete Lower Cretaceous deep-water carbonate section, providing the first view of Early Cretaceous sedimentation in the deep part of the Gulf of Mexico basin.
In December/January 1980-1981, the R/V Glomar Challenger drilled at five primary sites in the southeastern Gulf of Mexico (Fig. 4) (Sites 535, 536, 537, 538, 540: Buffler et al., 1984). The locations of the sites were based on an analysis of the seismic data discussed above. Holes 536, 537, 538A were drilled on high-standing basement blocks (basement sites), while Sites 535 and 540 were designed to sample the Mesozoic-Cenozoic basin fill (basin sites). The geologic and tectonic setting of each of these sites are shown schematically on three cross sections based on seismic lines (Figs. 4-7).
Site 535 Site 535 bottomed in Berriasian age rocks, leaving unsampled the thick sedimentary section below observed in the seismic data (Fig. 5). This unsampied sedimentary section is regionally extensive and is interpreted to comprise the Jurassic rocks (J) that form the major focus of this study (Schlager et al., 1984; Marton, 1995). At the bottom of Site 535, 100 m of condensed upper Berriasian to lower Valanginian alternations of massive bioturbated to poorly laminated limestone and marly limestone with numerous hardgrounds were recovered. They
EVOLUTION OF THE SOUTHEASTERN GULF OF MEXICO BASIN were interpreted to be deposited in at least bathyal depths. The lower 25 m section (upper Berriasian) is characterized by a low (5%) porosity, a high (4.7 km/s) sonic velocity, and a 2.57 g/cm 3 bulk density (Buffer et al., 1984). This is in contrast to the 17% porosity, 3.4 km/s sonic velocity, and 2.43 g/cm 3 density measured in the overlying Valanginian limestones. The abrupt change in acoustic properties results in a high-continuity, high-amplitude reflector, which is correlatable throughout most of the mapped area. Because of the limited resolution of the seismic (30-60 m at best) the horizon here is referred to as 'top Berriasian' or TB (Fig. 5). It occurs approximately at the Jurassic-Cretaceous boundary, and it is used as the main mapping horizon in this study. An overlying unit in Hole 535 consists of about 460 m of lower Valanginian to Cenomanian(?) laminated, variably bioturbated, frequently organic-rich limestones and marly limestones, with increasing sand-size, shallow-water skeletal debris towards the top of the sequence (Fig. 5).
Site 540 Site 540 was drilled in a higher physiographic setting (Fig. 5) and penetrated 460 m of Cretaceous (middle Albian to Cenomanian section), mostly pelagic sediments with fine-grained input and skeletal debris from the adjacent shallow platform to the east (Florida Escarpment). Truncating the Lower Cretaceous section is the widespread Middle Cretaceous Sequence Boundary (MCSB), which at Site 540 is characterized by a thin Late Cretaceous (Turonian to lower Paleocene) condensed interval consisting of carbonate debris flows and turbidites. Sites 537/536/538 (basement sites) Site 537 Site was drilled on the crest of a small knoll, about 25 km north of the Campeche Escarpment at the entrance of the Catoche Tongue (Figs. 4 and 6). At total depth 19 m of phyllite was recovered with an age of about 500 Ma (Pan-African), based on 4~ measurements (Dallmeyer, 1984). Above the metamorphic basement, quartz porphyry fragments were cored, indicating acid or silicic igneous activity in the area prior to and perhaps during early Berriasian time. These fragments were interpreted to represent an in-situ igneous body (sill or dike), or clasts from a nearby source (Buffler et al., 1984). Above the porphyry is a 50-m-thick, earlymiddle Berriasian transgressive sequence, starting with coarse-grained, poorly sorted arkosic sandstone, which was interpreted to be deposited in a fluvialalluvial or arid estuarine setting (Buffler et al., 1984). Thin bentonite beds represent contemporaneous volcanism. The overlying unit, consisting of mixed
69
dolomitic marls, muddy dolomites, and arkosic sand represent a transition to marine conditions. Brackish-water ostracods, plant and coal fragments in this unit were interpreted as deposition in a shallow-marine or estuarine environment. The lower to middle Berriasian transgressive sequence is followed by a 58 m upper Berriasian to Valanginian skeletal limestone, grainstone, and wackestone sequence with abundant shallow-water fragments. These rocks were interpreted as in-situ deposits, related to a deep carbonate platform. Alternatively, these sediments may had been reworked from a shallow-water setting and redeposited in a deep-water environment (Buffler et al., 1984; Schlager et al., 1984). This sequence is capped by a thin sequence of Cretaceous (Aptian) to Cenozoic (Lower Pliocene) nannofossil ooze with interbedded chert, mud, and volcanic ash, representing a typical pelagic environment (Fig. 6). Based on the recovered abbreviated section at Site 537, it can be concluded that higher standing parts of the Yucatan terrace, and probably the eastern Yucatfin block, were not transgressed until the very end of the Jurassic-earliest Cretaceous. Fault movement separating this block from the sediment source areas of the Berriasian fluvial deposit, cannot be older than late Berriasian to early Valanginian. Ongoing tectonic activity (rifting) at this time is compatible with a model which suggests that in the southeastern Gulf of Mexico, rifting occurred contemporaneously with Late Jurassic-earliest Cretaceous oceanic crust formation in the central Gulf of Mexico.
Site 536 Site 536 was drilled at the base of the Campeche Escarpment on the northern end of the Yucatfin terrace (Figs. 4 and 6). At the base of the hole, a 25-m-thick dolomite sequence of unknown age was drilled. The depositional setting of this section was established as a shallow-water platform, based on observed algal laminations and possible desiccation cracks. This dolomite is devoid of diagnostic fossils and is characterized by a very tight fabric (1% porosity) and a very high sonic velocity (over 6 kin/s). A 87Sr/86Sr age for this rock can be interpreted as either Middle Jurassic-Early Cretaceous or late Paleozoic (Permian) (Testarmata and Gose, 1984). However, the accompanying paleomagnetic study from the same work supported the Middle JurassicEarly Cretaceous age (Testarmata and Gose, 1984). Based on the physical characteristics of the rocks and their tectonic setting, an interpretation that the enigmatic dolomite is pre-Middle Jurassic (possibly late Paleozoic) in age is preferred here. Seismic evidence suggest that the Upper Jurassic section pinches out and is not present on the Yucatfin terrace (Figs. 6 and 8), which eliminates the Upper Jurassic-Early Cretaceous ages. Marine post-Per-
Fig. 5. Location of DSDP Sites 535 and 540 along seismic line SF-15. See Fig. 4 for location. (A) Uninterpreted seismic line. (B) Depth converted line drawing interpretation of seismic line with simplified stratigraphic columns of Sites 535 and 540. J -- Jurassic; TB -- top Berriasian post-rift unconformity; L K = Lower Cretaceous" M C S B = Mid-Cretaceous sequence boundary; U K - C = Upper Cretaceous-Cenozoic" VE = vertical exaggeration. (C) Stratigraphic summary of earliest Cretaceous rocks at DSDP Site 535 showing location of TB. Env. = depositional environment. Modified from Buffler et al. (1984). 7~
9 ;Z
O
=
Fig. 6. Location of DSDP Sites 536 and 537 along seismic line GT3-75C. See Fig. 4 for location. (A) Uninterpreted seismic line. (B) Depth converted line drawing interpretation of seismic line with simplified stratigraphic columns of Sites 536 and 537. J Jurassic" TB -- top Berriasian post-rift unconformity; L K -- Lower Cretaceous; M C S B = Mid-Cretaceous sequence boundary; U K - C = Upper Cretaceous-Cenozoic. (C) Stratigraphic summary of earliest Cretaceous rocks at DSDP Sites 536 and 537 overlying basement rocks. Env. = depositional environment. Modified from Buffler et al. (1993).
:Z
r.~
>.
0
X
0
2: D
,.H
0 ~H U: >.
Ul
0 Z 0 ~H
,< O
Fig. 7. Location of DSDP Site 538 along seismic line SF-2. See Fig. 4 for location. (A) Uninterpreted seismic line. (B) Depth converted line drawing interpretation of seismic line with simplified stratigraphic columns of Site 538. J = Jurassic; TB -- top Berriasian post-rift unconformity; L K = Lower Cretaceous; M C S B -- Mid-Cretaceous sequence boundary; U K - C = Upper Cretaceous-Cenozoic. (C) Stratigraphic summary of earliest Cretaceous rocks at DSDP Site 538 overlying basement rocks. Env. = depositional environment. Modified from Buffler et al. (1993). 7z
7z
9 ~Z
--.3 t',3
EVOLUTION OF THE SOUTHEASTERN GULF OF MEXICO BASIN
Fig. 8. Seismic line SF-6B along northern edge of Yucatfin terrace. See Fig. 4 for location. (A) Uninterpreted. (B) Interpreted. J Jurassic; T B = top Berriasian post-rift unconformity; L K = Lower Cretaceous; M C S B ---=Mid-Cretaceous sequence boundary; U K - C Upper Cretaceous-Cenozoic. Vertical exaggeration about 10 at sea bottom.
mian to pre-uppermost Middle Jurassic sediments are unknown from the Gulf of Mexico proper (Salvador, 1991), also making it improbable that the age of the dolomite is Early or Middle Jurassic. The possible equivalent of this early Mesozoic(?)/Paleozoic(?) unit extends southward beneath the Yucatfin terrace and may attain several kilometers in thickness (Figs. 6 and 8). Equivalent rocks in the deep central domain (e.g., near Site 535/540, Fig. 5)
73
=
=
have been interpreted from seismic as a pre-rift section in the basins and sometimes on the top of tilted blocks (Fig. 6), indicating that deposition of this unit predates faulting in the southeastern Gulf of Mexico. The dolomite sequence is overlain by 108 m of Aptian to Albian porous, coarse-grained, skeletal, lime-packstone-grainstone with neritic material and fine-grained radiolarian intercalations (Fig. 6). This Lower Cretaceous sequence was interpreted as talus
74
G.L. MARTON and R.T. BUFFLER
deposits at the foot of the Campeche Bank(Fig. 4). In this association the shallow-water component represents redeposited platform sediments shed into a deep-water environment, and the fine-grained component represents ongoing pelagic sedimentation. Site 5 3 8 a
Site 538A was drilled on the top of a large knoll (Catoche Knoll) that rises high above the Florida Plain (Figs. 4 and 7). In Hole 538A at total depth, 64 m of basement rocks were recovered, which consist of mylonitic gneiss and amphibolite. 4~ measurements again suggest a 500 Ma (Pan-African) age (Dallmeyer, 1984). These crystalline rocks are intruded by several generations of diabase dikes with 190-160 Ma (Lower-Middle Jurassic) 4~ intrusion ages. This rock assemblage was interpreted as rifted continental crust or 'transitional' crust (Dallmeyer, 1984; Schlager et al., 1984), an interpretation strongly supported by the general high-standing, block-faulted tectonic setting (Marton, 1995) (Fig. 7). Basement is covered by 67 m of Early Cretaceous (lower(?)-upper Berriasian to Valanginian) skeletal-oolitic-oncolitic limestones. There is a gradational change between the more oncolitic (deeper-water) facies at the base towards the more oolitic (shallow-water) facies at the top. Based on similarities between Sites 537 and 538A, Schlager et al. (1984) suggested that the limestones may have been deposited in-place on an extensive carbonate platform, rather than on an isolated block. Alternatively, it is possible that gradual shallowing of this earliest Cretaceous section records uplift(?) of the block (Catoche Knoll) from 50-200 m to as little as a few meters (Marton, 1995). Such uplift could be explained by fault-block rotation (footwall uplift) during tectonic activity (rifting) in the southeastern Gulf of Mexico during the earliest Cretaceous final phase of rifting and seafloor spreading. This latter interpretation would imply deposition on an already isolated block. The Berriasian to Valanginian limestones are capped by 210 m of A1bian to Upper Pliocene pelagic sediments, including chalk, foraminiferal-nannofossil ooze, radiolarian mudstone, and chert, indicating low sedimentation rates, frequent non-deposition/erosion on the finally submerged basement high (Fig. 7).
PREVIOUS SEISMIC STRATIGRAPHIC STUDIES
Following Leg 77, several seismic stratigraphic studies were undertaken in the study area, designed to integrate the results of the drilling with existing seismic data. Phair and Buffler (1983) and Phair (1984) described in detail the Lower Cretaceous section by subdividing it into four main depositional se-
quences and tying them into the drill sites. Angstadt (1983), Angstadt et al. (1983) and (1985) conducted a detailed study of the overlying Late CretaceousCenozoic rocks, mapped four main units, and documented development of an early Cenozoic foredeep along the north flank of the Cuban orogen, followed by current-controlled late Cenozoic pelagic sedimentation in the western Straits of Florida. Other seismic studies in the area include a description of the Catoche Tongue (Shaub, 1983). The studies by Angstadt (1983) and Phair (1984) were incorporated into a paper included in the DSDP Leg 77 volume that made a preliminary interpretation of the seismic stratigraphy, structural setting, and geologic history of the entire area (Schlager et al., 1984). Schlager et al. (1984) established a seismic stratigraphic framework, which included an inferred Late Triassic-Early Jurassic unit (TJ), two inferred Jurassic units (J1 and J2), as well as four Early Cretaceous units (EK 1-4).
TECTONO-STRATIGRAPHIC FRAMEWORK
In the present study of the southeastern Gulf of Mexico, the main purpose of seismic interpretation is to delineate the distribution and reveal the tectonic and depositional setting of the Jurassic sediments. The Lower Cretaceous sequences were studied earlier, and they are only discussed briefly to complete the picture. Major bounding surfaces (unconformities) previously defined in earlier studies have been re-correlated, including the Mid-Cretaceous Sequence Boundary (MCSB, formerly MCU) and the late (top) Berriasian horizon (here designated TB) (Marton, 1995). In addition, two other horizons, the top of the pre-rift section and the top of acoustic basement (crystalline basement), have been defined (Marton, 1995). All four of these surfaces have been interpreted where possible and mapped across the entire eastern Gulf of Mexico as well as in the southeastern Gulf (Marton, 1995). These four regional unconformity surfaces have been used to define four major tectono-stratigraphic sequences in the study area: (1) crystalline basement (rifted continental crust and oceanic crust); (2) Paleozoic(?) pre-rift rocks; (3) a Late Jurassic syn-rift sequence (J); and (4) a Lower Cretaceous post-rift sequence (LK). Each of these units are shown on the regional sections presented above (Figs. 5-8). The seismic character and additional details and interpretations of each unit, along with their relationships to tectonic development of the south-central part of the study area, are illustrated using four additional seismic lines oriented east-west, starting in the north and proceeding to the south (Figs. 9-12). The location of these lines is shown in Fig. 4 and also on a structure
EVOLUTION OF THE SOUTHEASTERN GULF OF MEXICO BASIN
75
Fig. 9. Uninterpreted and interpreted seismic line SF-7. J = Jurassic; TB = top Berriasian post-rift unconformity; LK = Lower Cretaceous; MCSB = Mid-Cretaceous sequence boundary; U K - C = Upper Cretaceous-Cenozoic. Vertical exaggeration about 10 at sea bottom. See Fig. 4 for location.
map of the pre-Jurassic surfaces (top pre-rift or top acoustic (crystalline) basement) in this area (Fig. 13). These sections and the map show (1) how faulting and
depositional style change in a north to south direction, (2) major rift-related features, such as half-grabens, grabens, major faults, fault blocks, and possible ac-
76
G.L. MARTON and R.T. BUFFLER
Fig. 10. Uninterpreted and interpreted seismic line SF-13. J Jurassic; T B -- top Berriasian post-rift unconformity; L K - - Lower Cretaceous; M C S B -- Mid-Cretaceous sequence boundary; U K - C = Upper Cretaceous-Cenozoic. Vertical exaggeration about 10 at sea bottom. See Fig. 4 for location. -
commodation zones, as well as (3) major depositional features, such as non-marine alluvial fans/fan deltas, deep marine beds and carbonate buildups. Each of the four tectono-stratigraphic sequences and their bounding surfaces are discussed in more detail below.
Crystalline basement Rifted continental crystalline basement (transitional crust), probably composed mainly of crystalline metamorphic rocks intruded by mafic rocks,
EVOLUTION OF THE SOUTHEASTERN GULF OF MEXICO BASIN
77
Fig. l l. Uninterpreted and interpreted seismic line SF-17. J = Jurassic; TB = top Berriasian post-rift unconformity; LK = Lower Cretaceous; MCSB = Mid-Cretaceous sequence boundary; U K - C -- Upper Cretaceous-Cenozoic. Vertical exaggeration about 10 at sea bottom. See Fig. 4 for location.
as drilled at Sites 537 and 538, is inferred to underlie the entire higher-standing central and southern parts of the study area (Schlager et al., 1984; Marton and Buffer, 1994; Marton, 1995) (Figs. 6 and 7). To the north beneath the Florida Plain, basement is believed to be oceanic crust (Marton, 1995) (Fig. 6). The boundary between oceanic and continental crust has been well defined
based on gravity, magnetic, and seismic data (Hall and Najmuddin, 1994; Marton and Buffler, 1994; Marton, 1995). The top of crystalline basement throughout the study area is interpreted to be the top acoustic basement on the seismic data. Acoustic basement is characterized by chaotic, incoherent reflections, which distinguishes it from the overlying sedimen-
78
G.L. MARTON and R.T. BUFFLER
Fig. 12. Uninterpreted and interpreted seismic line SF-11. J Jurassic; TB : top Berriasian post-rift unconformity; LK -- Lower Cretaceous; MCSB -- Mid-Cretaceous sequence boundary; U K - C = Upper Cretaceous-Cenozoic. Vertical exaggeration about 10 at sea bottom. See Fig. 4 for location. -
-
tary sections characterized by more organized reflections (Figs. 5-12). Most of the seismic lines show good examples of this boundary. However, the exact location of this boundary on some of the seismic data is not well imaged, due to greater depth and/or lack of acoustic impedance change along this boundary. The boundaries illustrated on all the seismic examples shown here, therefore, represent only the best interpretations possible and are not unequivocal everywhere.
Paleozoic(?) pre-rift rocks The pre-rift rocks are defined on the seismic data as the section overlying acoustic basement and underlying interpreted Late Jurassic syn-rift rocks filling grabens and half-grabens, or in some cases younger Lower Cretaceous rocks filling sag basins where the Jurassic syn-rift section is missing (Figs. 5-12). The top of this pre-rift section (or the top of basement when this section is missing),
EVOLUTION OF THE SOUTHEASTERN GULF OF MEXICO BASIN
79
Fig. 13. Structure map of the pre-Jurassic surfaces (top Paleozoic(?) pre-rift section or top acoustic (crystalline) basement) (pre-rift surface) in the southeastern Gulf of Mexico, showing: (a) distribution of Upper Jurassic carbonate platforms (PL 1-3), (b) major faults (heavy black lines), and (c) basement highs and lows. Graben morphologyin the central part of the area is shown for lines corresponding to Figs. 5 (SF-15), 9 (SF-7), 10 (SF-13), 11 (SF-17), 12 (SF-11). Contours are in meters. White areas indicate where Jurassic inferred to be thin or absent. HG = half grabens; AZ = possible accommodationzone.
which also defines the base of the syn-rift section, is defined here as the pre-rift surface. A structure map of this surface in the central study area shows the regional rift morphology (Fig. 13). The pre-rift section is characterized by organized but generally low-amplitude reflections. It is interpreted here to be an extensive sedimentary unit that occurs throughout most of the area with a fairly uniform thickness (Figs. 5-12). In half-grabens (e.g., Line SF 7, Fig. 9) the pre-rift section is observed as uniform packages of sub-parallel reflections on top of rotated fault blocks. Here this unit is distinct from the reflection-free basement below and from the more diverging and onlapping Late Jurassic synrift section above. These relationships indicate that these rocks have rotated with the underlying fault blocks and were deposited prior to the Late Jurassic rifting. In some places, the unit apparently is absent over high-standing fault blocks. This can be seen on
the block between CDP's 1300 and 1400 on Line SF 7 (Fig. 9), suggesting long-term uplift and erosion of the block. The pre-rift section on Line SF-13 (Fig. 10) has been offset along a large west-dipping normal fault bounding a half-graben on the east. On Line SF-17 (Fig. 11) it is down-dropped into a full graben. The section is thinner and truncated on the eastern up-thrown block, possibly due to erosion during or following rifting. Beneath the shallow Yucatan terrace area to the southwest where the interpreted Jurassic rocks are absent (Fig. 13), the pre-rift section is flat-lying and is very thick (estimated over 2 km) (Figs. 6, 8 and 12). Along the northern edge of the terrace, the unit is composed of two separate sub-units separated by a prominent angular unconformity, indicating a complex tectonic history (Fig. 6). On Line SF-11 (Fig. 12) the thick pre-rift section along the eastern edge of the Yucatfin terrace is down-dropped to the
80 east along a half-graben, while the western margin shows a series of smaller antithetic faults. As argued above under the discussion of Site 536, a tentative correlation of these rocks with the dolomite drilled at the bottom of the hole (Fig. 6) suggests a possible late Paleozoic age for these pre-rift rocks. The nearest Paleozoic sedimentary rocks of late Paleozoic age occur only in outcrop and in the subsurface to the south in Belize and Guatemala (Santa Rosa group) (e.g., Bateson, 1972) and in the Appalachian foldbelt to the north (e.g., Thomas et al., 1989).
Late Jurassic syn-rift sequence Distribution A more detailed study of the Late Jurassic syn-rift sequence (J) throughout the study area was one of the primary focuses of this study, as these rocks have not been examined in any detail. The distribution and tectonic setting of the Late Jurassic rocks (J) is illustrated on all the seismic lines (Figs. 512). A tectonic control on sedimentation is indicated, with presumed Jurassic sediments confined to depocenters in half-graben or graben settings. The characteristic wedge shape of the Jurassic sequence, i.e., the thickening against the border faults and the progressive onlap onto the pre-rift section away from the fault, is diagnostic of syn-sedimentary faulting (syn-rift setting). On some high standing blocks the sequence onlaps and is very thin or absent (e.g., Figs. 7-9 and 13), suggesting that many of the blocks remained high through the end of the Jurassic. The generalized structure map on the pre-rift surface (Fig. 13) also shows the tectonic setting of the Jurassic rift section in the south-central part of the study area, including major basement blocks, major faults and graben morphology. The map shows the areas where Jurassic rocks are inferred to be missing. This includes the extensive lack of Jurassic over the Yucat~in terrace to the southwest (Fig. 13), as well as at Sites 536, 537 and 538, which suggests that Jurassic sediments were not deposited across the Yucat~in block to the west. The Jurassic section is separated from the overlying Lower Cretaceous section by the prominent Late Berriasian unconformity (TB). This horizon was drilled at Site 535 (Fig. 5; see also earlier discussion of Site 535). At the drilling site it is seismically a strong event which can be correlated throughout the area. It is interpreted to separate Jurassic sediments, characterized by rapidly changing thicknesses, from the overlying Cretaceous sediments, characterized by more uniform distribution (Figs. 512). This horizon is commonly onlapped by Lower Cretaceous sediments, indicating its unconformable nature (Figs. 5-12).
G.L. MARTON and R.T. BUFFLER The Late Jurassic syn-rift section in the subbasins often can be divided into two units based on a distinctive vertical change in internal reflection character or seismic facies. These units are interpreted to be a non-marine unit below and a marine unit above.
Nonmarine(?) unit The lower unit occurs in a stratigraphically deeper position lying directly above the pre-rift surface. It represents the initial rocks deposited in the rift basins. The seismic response of this unit is more discontinuous, and of more variable amplitude, compared with that of the overlying younger Jurassic sediments. This is best illustrated on Lines SF-7 and SF-13 (Figs. 9 and 10). The wedge-shaped character of this lower unit in half-grabens adjacent to fault scarps indicates its syn-rift origin. This vertical change in facies is also observed in the eastern part of the graben shown on Line SF 17 (Fig. 11). The lower unit, however, changes seismic facies laterally to a more continuous facies to the west, suggesting a lateral change in depositional setting. The subdivision into the two units is not as obvious on Line SF-11 (Fig. 12), although the section does become more discontinuous along the faulted flanks of the half-graben. Based on the seismic character and the tectonic setting, these lower rocks are tentatively interpreted to have been deposited in a non-marine, rift-basin setting, possibly as footwall-sourced alluvial fan deposits (Figs. 9-11) and/or hanging wall alluvial cones or dip slope fans (Fig. 12). Alternatively, they may represent fan-deltas deposited in lakes (i.e., represented by the seismic facies change from discontinuous to continuous on Line SF-17; Fig. 11). The overall depositional setting envisaged here is a series of non-marine rift basins with interior drainage systems. Marine unit The upper unit contains two laterally equivalent seismic facies. Areas in the rift basin are characterized by more continuous reflections (Figs. 9-12), while adjacent basement highs are often capped by chaotic, discontinuous, or reflection-free seismic facies (Figs. 10 and 11). The continuous seismic facies in the basins is almost identical in seismic character with that of the overlying deep marine Lower Cretaceous section (Figs. 9-12), which has been sampled at DSDP Sites 535 and 540 (Fig. 5). These rocks, therefore, are tentatively interpreted to also represent sedimentation in a marine setting. The adjacent reflection-free facies overlying lower-standing basement highs is characteristic of and is interpreted here to represent contemporaneous carbonate platforms or buildups, possibly with central reef cores (Figs. 10 and 11). Above the reflectionfree zones drape can be observed, which represents
EVOLUTION OF THE SOUTHEASTERN GULF OF MEXICO BASIN differential compaction above the reef tops (Figs. 10 and 11). Further, between the reef more continuous reflections are interpreted as the acoustic response of lagoonal sediments (CDP 550, Fig. 10). The more continuous reflectors adjacent to the banks are interpreted as off-bank, more uniformly deposited, deeper marine sediments (Figs. 10 and 11). They probably consist of deeper-water shales and/or pelagic-hemipelagic carbonates. The distribution of the interpreted carbonate buildups on basement highs is shown on the tectonic map (Fig. 13). One large interpreted reef complex occurs on a northwest-trending elongated basement high (horst), which is bounded by a major fault on the southwest side (Figs. 10 and 13). The area of this bank appears to have become more restricted, with the margin stepping back through time (Fig. 10). The bank appears to have finally drowned out near the end of the Jurassic (TB; Fig. 10). Another Jurassic carbonate buildup was interpreted on the northeast side of a tilted footwall block shown on Line SF-17 at about CDP 500 (Figs. 11 and 13). The location of this carbonate edifice is somewhat different from the previously discussed bank, since it was established on a tectonically controlled east-dipping slope, rather than on the top of a block. Syn-rift tilting of the block resulted in a buildup which faces to the northeast. The margin of the platform is characterized by prograding clinoforms (CDP 550). Accelerated tilting and/or rising sea level then resulted in the drowning of the platform toward the end of the Berriasian (TB). Lower Cretaceous onlap onto the tilted syn-rift section is particularly well expressed on this section (Fig. 11). Based on the occurrence of the interpreted carbonate buildups, and based on the widespread occurrence of laterally equivalent reflections with marine characteristics, it can be postulated that the upper part of the Jurassic section was deposited in a shallow- to moderately deep-marine environment with a dominantly carbonate facies. The age of the underlying transition from non-marine to marine and the timing of the establishment of a marine seaway through the southeastern Gulf is not known, but it is inferred to have taken place during the middle part of the Late Jurassic. A Kimmeridgian age is suggested, which corresponds to the age of a possible similar shallow- to deep-water setting that occurs to the south within the Mesozoic section of western Cuba (Iturralde-Vinent, 1994). This is also a time of worldwide rise in sea level (Haq et al., 1988).
Tectonic setting Interpreted faults indicate that deposition during the Jurassic was controlled by faulting, which resulted in typical graben and half-graben filling geometries. Line SF-7 (Fig. 9) shows impressive
81
horst and features, related to significant N E - S W extension. The dominant sense of faulting is to the southwest, except for a pair of northeastward-dipping large faults on the northeastern end of the line (CDP 1400). On Line SF-13 (Fig. 10) only one half-graben can be identified with a southwestward-dipping basin bounding fault. This fault was reactivated in the Early Cretaceous, probably due to differential compaction and/or due to mild tectonic rejuvenations. Jurassic faulting is also observed around CDP 1000. These faults became inactive by the end of the Jurassic, as they do not cut the Top Berriasian surface (TB). Further south on Line SF-17 (Fig. 11) syn-rift Jurassic sedimentation in the basin southwest of CDP 1000 was controlled by two faults dipping in opposite directions, resulting in a full-graben setting. Thickness of the Jurassic section exceeds 0.5 s (1200 m). On the upthrown footwall, just northeast of CDP 1000, the Jurassic is thin or absent. This indicates that even in the deep central part of the seaway large fault blocks may have remained at or above sea level throughout the Jurassic rifting. The southernmost line shown here (SF-11, Fig. 12) shows a pair of southwestward-dipping major faults controlling Jurassic sedimentation in the half-graben between CDPs 500 and 1000. A thick, tectonically disrupted pre-rift section is particularly well defined with numerous, relatively small-scale faults interpreted around CDP 500. These faults apparently were active only in the early phase of rifting as they cut only the pre-rift section and the lowermost part of the Jurassic section. It can be observed that Jurassic sediments which fill the half-graben thin toward the southwest and pinch out against the Yucatan terrace (Fig. 12). Southwestward thinning and pinchout of the Jurassic section against a tilted block is also observed on the northeastern part of the section (CDP 1000-1200). The Jurassic is thin or absent over the tilted block, and either was never deposited or was later eroded. Between CDP 1000 and 1500 very well-defined onlap of Cretaceous beds defines the angular unconformity between the tilted Jurassic section and the more horizontal overlying Lower Cretaceous section (TB). Faults which extend into the Lower Cretaceous section (about CDP 1200) again represent minor tectonic rejuvenations and/or differential compaction. It has been recognized that half-grabens are major tectonic elements in continental rifts (e.g., Rosendahl, 1987). In the seismic examples shown here it was documented that Jurassic sediments were deposited in both half-graben and full-graben settings in the central part of a presumed Jurassic seaway (Figs. 9-13). In these examples the halfgrabens change their sense of asymmetry along the northwest-trending axis of the rift system (Fig. 13).
82 The transitional zones or overlapping zones between contrasting polarity rift segments are known as transfer or accommodation zones (e.g., Morley, 1988). A possible example of an accommodation zone may occur in the vicinity of Line SF-17 (Figs. 11 and 13), where there is an overlap between the main faults forming the opposing half-grabens. Here the Jurassic occurs in a graben setting and thickening of the Jurassic against major faults can not be observed (Fig. 11). In contrast, to the north and south on Lines SF-15 (Fig. 5) and SF-11 (Fig. 12), the Jurassic section thickens against major faults, halfgraben morphology and syn-depositional faulting is interpreted (Fig. 13). Another accommodation zone probably occurs at the overlap of fault zones between Lines SF 13 and SF-15 (Fig. 13). The tectonic setting described above is generally characteristic of the syn-rift period in the evolution of passive margins. This setting in the southeastem Gulf, however, is different, in that the rifting is contemporaneous with seafloor spreading in the central Gulf to the north, rather than occurring prior to seafloor spreading (Fig. 2). Faults in the area generally do not cut the late Berriasian surface (TB), indicating the cessation of rifting near the Berriasian-Valanginian boundary. This is used as evidence for the time of abandonment of seafloor spreading and oceanic crust formation in the north, as proposed in the regional model of Gulf evolution by Marton and Buffler (1994) and Marton (1995) (Fig. 2). L o w e r C r e t a c e o u s post-rift s e q u e n c e (LK)
The Lower Cretaceous post-rift sequence (LK) forms a widespread drape over the entire study area. It is relatively thin over the western and central part of the study area, but becomes thicker to the north over oceanic crust and particularly to the east toward the Florida Escarpment (Figs. 5-12). The geology of this sequence has been studied in detail by Phair (1984), who divided the unit into four seismic sequences. He mapped the thickness, seismic facies, and interpreted the depositional history of these sequences using the DSDP sites as control. Results of this study were also reported by Phair and Buffler (1983) and Schlager et al. (1984). These studies concluded that the primary source for the mainly hemipelagic deep-sea carbonates making up these sequences was sediment shed from the Florida Escarpment to the east. These units thin and onlap to the west onto basement highs in the south-central part of the area (e.g., Figs. 5, 6, 11 and 12), and they also thicken and onlap onto the underlying Jurassic section in residual sag basins overlying the Jurassic rift grabens and half-grabens (Figs. 6, 8 and 9). This latter relationship is espe-
G.L. MARTON and R.T. BUFFLER cially well expressed in the sub-basin around CDP 1000 in Fig. 9. Here thickness variations within the Lower Cretaceous section are not directly controlled by the major faults. In general, the Lower Cretaceous section fills the remnants of the rift basins and forms a widespread blanket. This geometry and the lack of major basement-involved faulting are characteristic for sag basins in a post-rift setting. The Lower Cretaceous sequence is capped by the Middle Cretaceous Sequence Boundary (MCSB) (e.g., Schlager et al., 1984; Buffer, 1991). The MCSB always coincides with a high-amplitude reflector or reflection package. Its erosional character is well expressed on many seismic lines, where it truncates underlying Lower Cretaceous sediments (Figs. 5-12). It is onlapped by younger Upper Cretaceous to Cenozoic sediments
TECTONO-PALEOGEOGRAPHIC EVOLUTION OF THE SOUTHEASTERN GULF OF MEXICO BASIN AND SURROUNDING REGIONS
The Jurassic through Early Cretaceous tectonopaleogeographic evolution of the southeastern Gulf of Mexico is outlined in this section. These reconstructions are based on the constraints drawn from the regional scale model (Fig. 2), interpretation of the seismic data (Figs. 3-12), available DSDP well control (Figs. 4-8), plus the additional interpretations presented by Marton and Buffer (1994) and Marton (1995), particularly interpretations of regional geopotential data and structural and isopach maps. To visualize the main events along with the already published and discussed kinematics of the eastern Gulf, a series of tectono-paleogeographic maps are presented (Fig. 14). Although paleogeographic data are extremely scarce, an attempt is made to tie the tectonic events to the seismically identified rocks and infer, at least at first order, tectonic styles, and depositional styles, as well as the paleo-extent of seas. To place the southeastern Gulf of Mexico in a more regional context, the Jurassic-Lower Cretaceous settings of the adjacent northeastern Gulf as well as western Cuba are included on the maps. Data from the northeastern Gulf are taken from seismic stratigraphic studies of Lower Cretaceous rocks of the offshore northeastern Gulf (Corso, 1987), Jurassic rocks of the offshore northeastern Gulf (Dobson, 1990), and Mesozoic rocks along the Florida Escarpment in the deep eastern Gulf (DeBalko, 1991) (Fig. 3). Additional information from these studies are included in Corso et al. (1989), DeBalko and Buffler (1992), Buffer et al. (1993), and Dobson and Buffler (1997). These studies document for the first time the evolution of Jurassic through Early Creta-
EVOLUTION OF THE SOUTHEASTERN GULF OF MEXICO BASIN ceous shelf margins and provide a paleogeographic context for this part of the Gulf basin. Mesozoic rocks also occur throughout western Cuba just to the south of the study area in the Cordillera de Guaniguanico region, and they provide the most comprehensive Mesozoic section of the northern proto-Caribbean passive margin (Fig. 1). They occur as north-northwestward-thrust tectonic slices contained in three main tectonostratigraphic units: (1) the Sierra de los Organos; (2) Sierra del Rosario meridional (southern); (3) Sierra del Rosario septentrional (northern) (Fig. 1) (Pszcz6lkowski, 1987; Iturralde-Vinent, 1994). Although the exact palinspastic reconstruction of these rocks can only be inferred, they have been interpreted by IturraldeVinent (1994) to have originated along the southern continental margin of Yucat~.n, and then later incorporated into the Cuban foldbelt during a Late Cretaceous to early Cenozoic orogenic event. This palinspastic reconstruction is adopted for the purposes of the paleogeographic maps presented here (Fig. 14). There are significant differences among the three regions (northeastern Gulf, southeastern Gulf, and western Cuba) shown on the maps. The northeastern Gulf stratigraphy represents sedimentation along a Late Jurassic to Early Cretaceous passive margin which faces the eastern Gulf of Mexico ocean basin. The southeastern Gulf stratigraphy represents an active Late Jurassic to earliest Cretaceous continental rift, which becomes a subsiding passive margin only in the Early Cretaceous. The western Cuban section is interpreted to represent a portion of a south-facing Mesozoic passive margin of the western proto-Caribbean located further to the south, and it probably only had very broad similarities to the former two areas.
Late Triassic(?) to late Middle Jurassic (Fig. 14A) During the Late Triassic(?) to late Middle Jurassic (Callovian), the Gulf of Mexico region accommodated a large amount of northwest-southeast extension related to the incipient breakup of western Pangea (Salvador, 1991; Marton and Buffler, 1994; Marton, 1995; Fig. 2A,B). In general, this extension occurred above sea-level and was concentrated between the relatively unextended Yucatfin block and the southern margin of North America. Another locus of extension was located between Yucatfin and the northern margin of South America (Fig. 2A,B). Marine rocks of this age have not been reported from the Gulf of Mexico basin proper, which indicate a long period of extension in a continental rift setting (Salvador, 1991). The end of this rifting phase is best characterized by subsidence below sea level in the Callovian, resulting in the widespread deposi-
83
tion of evaporites (mainly salt) in a restricted Gulf of Mexico basin (Louann-Campeche salt) (Salvador, 1991) (Fig. 14A). An important result of the regional model (Fig. 2) (Marton and Buffler, 1994; Marton, 1995) was that oceanic crust formation began in the Gulf of Mexico in late Middle Jurassic (Callovian) time. In the northeastern Gulf of Mexico salt occurs above an interpreted 'breakup unconformity', which indicates that although salt deposition may have started during the very end of the rifting stage in the central Gulf, it probably continued into the early spreading stage around the periphery of the basin and was deposited on the initial oceanic crust in the center of the basin (Marton and Buffler, 1994; Marton, 1995). The southeastern Gulf of Mexico formed a quasicontinuous bridge between Yucatfin and Florida separating the extending areas to the north (Gulf of Mexico) and to the south (proto-Caribbean) (Fig. 14A). Although Lower to Middle Jurassic dikes in DSDP Hole 538A provide some evidence for the onset of extension, rifting during this period in the southeastern Gulf was subordinate to rifting in the Late Jurassic. The only well-defined east-northeast-trending extensional feature (corresponding to the general northwest-southeast extension) is the Catoche Tongue (Fig. 14A). This half-graben, based on its orientation, may represent an Early to Middle Jurassic rifting episode (Shaub, 1983). The amount of early to late Middle Jurassic extension in the southeastern Gulf was possibly equivalent to the adjacent Yucatfin block (fl = 1.25) or Florida block (fl = 1.4) as discussed by Marton (1995). The emergent position of the future seaway (continental bridge) is supported by the lack of preserved rocks of this age in the DSDP wells and in the area as inferred from the seismic data. Further, distribution of thick Callovian to Oxfordian salt also suggests that marine waters did not reach the southeastern Gulf of Mexico area (Fig. 14A). During the Late Triassic(?) to late Middle Jurassic period north of 27~ in the northeastern Gulf largescale basins and structural highs formed, including the Apalachicola basin, Southern platform, and the Tampa embayment (Fig. 14A). Distribution of the late Middle Jurassic salt reflects the areas of more extreme extension, which subsided below sea level before the termination of rifting. Farther south, along the southern margin of Yucatfin, the Lower/Middle Jurassic(?) to Oxfordian terrigenous San Cayetano Formation was deposited (Fig. 14A) (Pszcz6tkowski, 1987). Maximum thickness of the formation reaches and may exceed 3000 m in the Sierra de los Organos of present-day western Cuba (Fig. 1). The Yucatfin block has been interpreted as the source area of these clastic sediments (Pszcz6tkowski, 1987). The well-bedded shales and
Fig. 14. Maps showing interpreted Late Jurassic tectono-paleogeographic reconstructions of the eastern Gulf of Mexico and northwestern Cuba (palinspastically restored to southeastern margin of Yucat~in block) for: (A) Callovian, (B) Oxfordian, (C) Kimmeridgian, (D) Tithonian, (E) Berriasian-Valanginian, and (F) Early Cretaceous. Gray line around Yucat~in and Florida block is present location of
9 Z
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Fig. 14 (continued). Lower Cretaceous platform margin for reference. Dashed heavy lines indicate inferred landward limit of seas. Pole of rotation in southeastern Gulf is shown with degrees of rotation of the Yucatfin block. Other lines and patterns on maps are as annotated. See text for discussion of each map.
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86
G.L. MARTON and R.T. BUFFLER
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EVOLUTION OF THE SOUTHEASTERN GULF OF MEXICO BASIN sandstones of the San Cayetano sediments represent deposition in a deltaic and local shallow-marine setting (e.g., Haczewski, 1987). In the Sierra del Rosario (present-day western Cuba, Fig. 1) thickness of this unit never exceeds 1000 m, and facies represent a more distal setting for this segment of the Cordillera de Guaniguanico (Haczewski, 1987). The San Cayetano sediments may represent deposition in an extensional rift basin, which formed as Yucatfin and South America rifted apart. This is corroborated by the occurrences of basic dikes and sills in the San Cayetano Formation (Iturralde-Vinent, 1994, 1996).
87
South of Yucatfin in western Cuba the middle Oxfordian shallow-water sediments of the Jagua Formation (in the Sierra de los Organos, Fig. 1) and the Francisco Formation (in the Sierra del Rosario, Fig. 1) consist of shales and limestones similar to the Smackover Formation. A thick unit of Oxfordian to early Kimmeridgian tholeiitic basalts, interbedded with shale and limestones, in the northern Sierra del Rosario has been interpreted as the result of continental margin volcanism (Iturralde-Vinent, 1994, 1996).
Kimmeridgian (Fig. 14C) Oxfordian (Fig. 14B) By Oxfordian time the Yucatan block had rotated 11~ counter-clockwise, a narrow oceanic tract had propagated into the northernmost segment of the southeastern Gulf of Mexico, and salt deposition was completed in the Gulf of Mexico basin (Fig. 2C, Fig. 14B). The identified salt in the southeastern Gulf extends only slightly southward of the tip of the oceanic crust and is confined more to the east, suggesting asymmetrical rifting and subsidence along the Florida margin. Major continental rifting had begun all along the southeastern Gulf, but the area remained at or just above sea level, preventing deposition of thick salt. The older part of the wedge-shaped Jurassic syn-rift units, which occur along the marginal faults bounding syn-rift basins (e.g., Figs. 9 and 10), probably represent this continental (non-marine) syn-rift period, with the local deposition of alluvial fans or fan deltas building into small bodies of fresh water (Fig. 14B). In the northeastern Gulf of Mexico, above the Louann salt, the early to middle Oxfordian basal clastics of the Norphlet Formation were deposited, which include an alluvial-fluvial-eolian system bordering the Oxfordian sea (Salvador, 1987; Dobson, 1990; Dobson and Buffler, 1997). Pronounced transgression of the Oxfordian sea is documented by the deposition of the carbonate-dominated Smackover Formation (Fig. 14B). Shallow-water carbonates were deposited on a monocline or in a ramp setting, which deepened towards the rapidly subsiding basin center. On local highs basin margin carbonate buildups (composed of corals, sponges and stromatolites) were established (Dobson, 1990) (Fig. 14B). These sediments can be regarded as the foundation of the newly established passive margin facing the spreading Gulf of Mexico. West of the Tampa embayment and the Sarasota arch, in a deepmarine (slope) setting, DeBalko (1991) has identified an Oxfordian section more than 1000 m thick. This may represent deep-sea fan sediments deposited in a rapidly subsiding basin next to the spreading center.
During Kimmeridgian time the southeastern Gulf of Mexico seaway became wider as seafloor spreading continued, the Yucatfin block rotated 18~ counter-clockwise, and the tip of the spreading center propagated farther south (Fig. 14C). It is possible that by this time a marine connection between the Gulf of Mexico and the proto-Caribbean had become established, and the continental bridge had subsided below paleosea-level. Although drilling data are not available, it is postulated that the seismically identified carbonate buildups in the central part of the seaway (Figs. 10 and 11) started to grow during this time. Several of the large blocks in the central part of the seaway probably remained above sea-level, as well as the whole Yucatfin block and the Yucatfin terrace (Figs. 13 and 14C). In the northeastern Gulf of Mexico the carbonate-dominated Smackover Formation gave way to the mixed clastic/carbonate sequence of the Kimmeridgian Haynesville Formation (Salvador, 1987) (Fig. 14C). Based on seismic stratigraphy and limited (up-dip) well information, Dobson (1990) interpreted clastic(?) shelf margin progradation in the Tampa embayment and a pronounced carbonate shelf margin in the outer Apalachicola basin (Gilmer equivalent) (Fig. 14C). Deep sea fans continued to be deposited in the deep basin adjacent to the Tampa embayment/Sarasota arch (DeBalko, 1991). In the late Oxfordian to Kimmeridgian in the Sierra de los Organos segment of western Cuba (Fig. 1) extensive carbonate platform(s) evolved (Pszcz6lkowski, 1987) (Fig. 14C). The thickness of the thick-bedded shallow-water platform sediments represented by the San Vicente Member of the Guasasa Formation reaches 650 m. In the Sierra del Rosario segment (Fig. 1) the equivalent section reaches only a few tens of meters. Kimmeridgian sediments in the Sierra del Rosario represent relatively slow sedimentation, but they do not have characteristics of typical pelagic deposits. They may instead represent a transition zone between the shallow-water platform of the Sierra de los Organos and
88 a deeper basinal area to the south (Iturralde-Vinent, 1994) (Fig. 14C).
Tithonian (Fig. 14D) By Tithonian time the tip of oceanic crust had propagated into the northern segment of the southeastern Gulf of Mexico, leaving behind rapidly subsiding passive margins of increasing length on both sides of the seaway (Fig. 14D). Rifting had reached a more advanced stage in the southern segment, as the Yucat~in block rotated 25 ~ counter-clockwise around the rotation pole. By Kimmeridgian or Tithonian time a full-fledged seaway may have evolved between the opening Gulf of Mexico and the proto-Caribbean. This is based on the seismically identified carbonate buildups that are present in the axial part of the seaway and the equivalent off-bank reflectors with marine characteristics (Figs. 10 and 11). Although no direct evidence is available, it is possible that drowning and step-back of these small carbonate platforms began in the Tithonian (Figs. 10 and 11). Several of the large blocks still remained above sea-level, including the Yucat~in block and the Yucat~in terrace (Figs. 13 and 14D). Tithonian sedimentation in the northeastern Gulf of Mexico is characterized by deposition of the thick fluvial-deltaic sequence of the Cotton Valley clastics that prograded across the Apalachicola basin and into the Tampa embayment (Dobson, 1990) (Fig. 14D). Deep-sea fans influenced by southward currents were interpreted in the deep basin adjacent to the Tampa embayment/Sarasota arch (DeBalko, 1991). In western Cuba shallow-water sedimentation ceased in the Tithonian (Pszcz6tkowski, 1987). Slow sedimentation and deepening environments characterizes the E1 Americano Member of the Guasasa Formation in the Sierra de los Organos (Fig. 1). Similarities between this formation and the La Zarza Member (upper part) of the Artemisa Formation in the Sierra del Rosario (Fig. 1) indicate uniform pelagic deposition throughout the entire Cordillera de Guaniguanico region (Pszcz6ikowski, 1987) (Fig. 14D). This deep-water environment in the Guaniguanico suggests a transgression towards the north onto the southern margin of the Yucat~in block (Fig. 14D). Continued extension along this margin is suggested by coeval basaltic sills and dikes in the Sierra del Rosario sections (Iturralde-Vinent, 1996).
Berriasian-Valanginian (Fig. 14E) In the late Berriasian spreading ceased in the Gulf of Mexico and the Yucat~in block reached its present position (42~ rotation) (Fig. 2D, Fig. 14E). Consequently, rifting also ceased throughout the southeastern Gulf of Mexico and rapid (thermal) subsidence
G.L. MARTON and R.T. BUFFLER commenced in the entire area. Late Jurassic carbonate buildups in the central part of the seaway were drowned and marine transgression submerged previously emergent blocks. Early to middle Berriasian transgression in the tectonically active area near block 'BL' (Fig. 14E) is well documented at DSDP Site 537 (Fig. 6). Fault activity at this time or slightly later finally separated the block from Yucat~in, which had to be the source area for the drilled fluvial to nearshore sediments. The Berriasian to Valanginian transgression also probably reached the area of the Yucat~in terrace and possibly the entire Yucat~in block. Ephemeral shallow-water carbonate platforms were established on the Catoche Knoll and on 'BL', represented by recovered late Berriasian to early Valanginian shallow-water limestones at Site 537 (Fig. 6). Deep-sea conditions in the late Berriasian in the central part of the seaway was documented at DSDP Site 535 (Fig. 5). In the northeastern Gulf of Mexico, Cotton Valley sedimentation continued into the Berriasian (Fig. l lE). Close to the future Lower Cretaceous platform margin, the seismically and lithologically well defined Knowles ramp was mapped by Dobson (1990) based on drilling and seismic stratigraphy. The exact age of the transgression onto the higher standing parts of the Sarasota arch and south Florida platform is not well established. In south Florida, in the type well (Bass Collier Country 12-2), Applegate et al. (1981) described the latest Jurassic(?) to earliest Cretaceous Wood River Formation overlying Early Jurassic (189 Ma radiometric age) rhyolite porphyry. Based on palynological studies Salvador (1987) stated that only the basal sandstone- and shale-dominated 30-50 m of the section is Tithonian. Consequently, the upper (600 m) dolomite and evaporite section of the Wood River Formation represents earliest Cretaceous (and younger) platform deposits. In western Cuba in the Berriasian and Valanginian, pelagic sedimentation prevailed in the entire Cordillera de Guaniguanico region (Fig. 1), including radiolarian ooze deposits and cherts (Pszcz6ikowski, 1987) (Fig. 14E). Dramatically reduced sedimentation rates and very deep-water facies in western Cuba indicate gradual deepening and starved basin conditions south of the southern Yucatfin margin in the proto-Caribbean. In contrast, limestones of the Tumbadero and Tumbitas Members of the Guasasa Formation in the Sierra de los Organos, and the Sumidero Member of the Artemisa Formation in the Sierra del Rosario, do not contain any terrigenous material (Fig. 14E).
Early Cretaceous (Aptian-Albian) (Fig. 14F) During Early Cretaceous time in the southeastern Gulf of Mexico passive margin conditions prevailed.
EVOLUTION OF THE SOUTHEASTERN GULF OF MEXICO BASIN This period is represented by the deposition of thick sections of platform carbonates on the Yucatfin and Florida blocks (Fig. 14F). Steep margins formed which separated the platform areas from the deep basins. Present-day relief along the Campeche Florida escarpments reaches and exceeds 3000 m. In the axial trough deep-sea pelagic and hemipelagic sedimentation occurred, as interpreted at Sites 535 and 540. Exceptions are Pinar del Rfo Knoll and Jordan Knoll, on top of which shallow-water carbonate platforms were able to keep up with the fast subsidence (Fig. 14F). In Albian-Cenomanian time these carbonate platforms were progressively drowned, and the platform margins stepped back to shallower positions towards Yucatfin and Florida. The southeastern Gulf of Mexico became a sediment-starved deep seaway, the floor of which was scoured at various times by deep-sea currents during the mid- and Late Cretaceous.
CONCLUSIONS The regional opening model (Fig. 2) (Marton and Buffler, 1994; Marton, 1995) provides a framework in which the Mesozoic tectono-stratigraphic evolution of the southeastern Gulf of Mexico can be discussed. This involved a southward-propagating spreading center/rift as Yucatfin rotated counter-clockwise relative to Florida during the Late Jurassic about a nearby pole in the southern part of the study area. The continental domain of the southeastern Gulf of Mexico, based on interpretation of reflection seismic data, is characterized by variable, 3500 to 7000 m basement depths, and by a continental rift morphology. The dominant tectonic elements in this area are northwestward-trending horst and features, corresponding to the main northeast-southwest maximum extensional strain. A detailed analysis of the available geophysical data as well as the review of the DSDP Leg 77 results in the southeastern Gulf of Mexico provide important information to decipher details for the development of this Late Jurassic-Early Cretaceous seaway. Based on this analysis, four major tectono-stratigraphic sequences bounded by major unconformity surfaces have been defined: crystalline basement, Paleozoic( ?) pre-rift rocks, a Late Jurassic rift sequence, and an Early Cretaceous post-rift sequence. Major conclusions of the tectono-paleogeographic evolution of the area are summarized below. (1) The oldest sedimentary unit in the southeastern Gulf of Mexico has been delineated based on the primary criterion that its thickness variations are not correlated with offset along Jurassic extensional faults. Because deposition of this unit (up to 2000
89
m thick) is not related to subsequent extension, it is referred to as the 'pre-rift' section. Its age is not clearly documented in the available DSDP wells. Based on the interpreted structural setting and possible correlation to DSDP Site 536, it is suggested that this pre-rift unit represents a Pre-mesozoic (late Paleozoic?) sedimentary cycle. (2) Based on the available seismic data, the Late Jurassic syn-rift section shows a complex basin-fill geometry in the continental domain of the southeastern Gulf of Mexico. The syn-rift nature of the Jurassic section is interpreted based on its characteristic wedge-shaped occurrence in half-grabens and its progressive onlap onto the pre-rift section. In general, Jurassic is interpreted to be missing on the high-standing Yucatfin and Florida blocks, as well as on large-scale basement highs in the central part of the southeastern Gulf of Mexico and the Yucatfin terrace to the southwest. (3) Syn-rift sedimentation has been interpreted to evolve from a non-marine to a normal marine setting during the latest Middle Jurassic to earliest Cretaceous period. Lack of salt deposition and the non-marine character of seismic reflectors in the deep half-grabens indicate that in the Callovian to Oxfordian(?) in the southern and central segments of the southeastern Gulf of Mexico rifting occurred in a continental setting. The southeastern Gulf of Mexico during this time formed a continental bridge between Yucatfin and Florida. Seismically identified carbonate buildups and the more marine character of seismic reflectors in the upper part of the Jurassic section suggest that the later rifting-stage (Kimmeridgian(?) to late Berriasian) in the southeastern Gulf of Mexico occurred in a marine setting. Although no drilling data are available, it is suggested that in Kimmeridgian to Tithonian time a marine seaway opened between the Gulf of Mexico and the proto-Caribbean as the southeastern Gulf of Mexico subsided below sea level. In late Berriasian time spreading in the Gulf of Mexico ceased and the southeastern Gulf of Mexico rift-system aborted. (4) Onset of rapid thermal subsidence in the earliest Cretaceous resulted in the final submergence of several of the large fault blocks as well as the Yucatfin and Florida platforms. Long-lasting carbonate platforms became established on rift shoulders (Yucatfin and Florida blocks), while deep-sea sedimentation prevailed in the axial part of the southeastern Gulf of Mexico, as indicated by the results of the DSDP Leg 77. Along the constructive carbonate margins of Yucatan and Florida, thicknesses of the Lower Cretaceous platform carbonates reached and exceeded 1500 m. In the adjacent basinal setting, depending of the availability of off-platform material, a 500-m to 1500-m-thick carbonate-dominated section accumulated in the Early Cretaceous.
90
G.L. MARTON and R.T. BUFFLER
ACKNOWLEDGEMENTS
S u p p o r t for this study was in part p r o v i d e d by the N a t i o n a l S c i e n c e F o u n d a t i o n grant O C E - 9 0 2 0 6 7 3 , Texas H i g h e r E d u c a t i o n C o o r d i n a t i n g B o a r d Advanced Research program, generous fellowships g r a n t e d by T h e U n i v e r s i t y of Texas D e p a r t m e n t of G e o l o g y and Institute for G e o p h y s i c s and the Society of E x p l o r a t i o n G e o p h y s i c i s t s . T h e authors are p a r t i c u l a r l y thankful to Paul M a n n for his support, s u g g e s t i o n s and e n c o u r a g e m e n t during the preparation of the m a n u s c r i p t . R e v i e w o f this p a p e r by J a m i e Austin, A1 H i n e and A n d r z e j P s z c z 6 l k o w s k i is d e e p l y a p p r e c i a t e d . T h e i r c o m m e n t s and professional insights significantly i m p r o v e d the quality of the m a n u s c r i p t . This is T h e U n i v e r s i t y of Texas Institute for G e o p h y s i c s C o n t r i b u t i o n 1339.
REFERENCES
Angstadt, D.M., 1983. Seismic Stratigraphy and Geologic History of the Southeastern Gulf of Mexico/southwestern Straits of Florida. M.A. Thesis, The University of Texas at Austin, Austin, Texas, 206 pp. Angstadt, D.M., Austin, J.A., Jr. and Buffler, R.T., 1983. Deepsea erosional unconformity in the southeastern Gulf of Mexico. Geology, 11: 215-218. Angstadt, D.M., Austin, J.A., Jr. and Buffler, R.T., 1985. Seismic stratigraphy and geologic history of the southeastern Gulf of Mexico-southwestern Straits of Florida. Am. Assoc. Pet. Geol. Bull., 69: 977-995. Applegate, A.V., Winston, G.O. and Palacas, J.G., 1981. Subdivision and regional stratigraphy of the pre-Punta Gorda rocks (lowermost Cretaceous-Jurassic?) in south Florida. Gulf Coast Assoc. Geol. Soc. Suppl. Trans., 31: 447-453. Bateson, J.H., 1972. New interpretation of geology of Maya Mountains, British Honduras. Am. Assoc. Pet. Geol. Bull., 56: 956-963. Bryant, W., Meyerhoff, A.A., Brown, N., Furrer, M., Pyle, T. and Antoine, J., 1969, Escarpments, reef trends, and diapiric structures, eastern Gulf of Mexico. Am. Assoc. Pet. Geol. Bull., 53: 2506-2542. Buffler, R.T., 1991. Seismic stratigraphy of the deep Gulf of Mexico basin and adjacent margins. In: A. Salvador (Editor), The Gulf of Mexico Basin. The Geology of North America, J, Geological Society of America, Boulder, Colo., pp. 353-387. Buffer, R.T., Schlager, W. and Shipboard Scientific Party, 1984. Initial Reports of the Deep Sea Drilling Project, 77. U.S. Government Printing Office, Washington, D.C., 747 pp. Buffer, R.T., Dobson, L.M. and DeBalko, D.A., 1993. Middle Jurassic through Early Cretaceous evolution of the northeastern Gulf of Mexico basin. In: J.L. Pindell and B.B. Perkins (Editors), Mesozoic and Early Cenozoic Development of the Gulf of Mexico and Caribbean Region. Gulf Coast Section, Society of Economic Paleontologists and Mineralogists, pp. 33-50. Corso, W., 1987. Development of the Early Cretaceous Northwest Florida Carbonate Platform. Ph.D. dissertation, The University of Texas at Austin, Austin, Texas, 136 pp. Corso, W., Buffer, R.T. and Austin, J.A., 1989. Erosion of the southern Florida escarpment. In: A. Bally (Editor), Atlas of Seismic Stratigraphy. Am. Assoc. Pet. Geol. Stud. Geol., 27 (2): 149-157.
Dallmeyer, R.D., 1984. Ar40/Ar39 ages from a pre-Mesozoic crystalline basement penetrated at holes 537 and 538A of the Deep Sea Drilling Project Leg 77, southeastern Gulf of Mexico: tectonic implications. Init. Rep. DSDP, 77: 497-506. DeBalko, D.A., 1991. Seismic Stratigraphy and Geologic History of Upper Middle Jurassic through Lower Cretaceous Rocks, Deep Eastern Gulf of Mexico. M.A. Thesis, The University of Texas at Austin, Austin, Texas, 143 pp. DeBalko, D.A. and Buffer, R.T., 1992. Seismic stratigraphy and geologic history of Middle Jurassic through Lower Cretaceous rocks, deep eastern Gulf of Mexico. Trans. Gulf Coast Assoc. Geol. Soc., 42: 89-105. Dobson, L.M., 1990. Seismic Stratigraphy and Geologic History of Jurassic Rocks, Northeastern Gulf of Mexico. M.S. Thesis, The University of Texas at Austin, Austin, Texas, 165 pp. Dobson, L.M. and Buffler, R.T., 1997. Seismic stratigraphy and geological history of Jurassic rocks, northeastern Gulf of Mexico. Am. Assoc. Pet. Geol. Bull., 81: 100-120. Haczewski, G., 1987. Sedimentological reconnaissance of the San Cayetano Formation: an accumulative continental margin in the Jurassic of western Cuba. In: A. Pszcz6tkowski, K. Piotrowski, A. De la Torre, R. Myczynski and G. Haczewski (Editors), Contribucion a la Geologia de las Provincia Pinar del Rio (Contributions to the Geology of Pinar del Rio Province). Editorial Cientffco-T6cnica, Ciudad de la Habana, La Habana, pp. 228-247. Hall, S.A. and Najmuddin, I.J., 1994, Constraints on the tectonic development of the eastern Gulf of Mexico provided by magnetic anomaly data. J. Geophys. Res., 99: 7161-7175. Haq, B.U., Hardenbol, J. and Vail, ER., 1988, Mesozoic and Cenozoic chronostratigraphy and cycles of sea-level change. In: C.K. Wilgus, B.S. Hastings, C.G.St. Kendall, H.W. Posamentier, C.A. Ross and J.C. Van Wagoner (Editors), Sea-Level Changes: An Integrated Approach. Soc. Econ. Paleontol. Mineral. Spec. Pap., 42: 40-45. Iturralde-Vinent, M.A., 1994. Cuban geology: a new plate-tectonic synthesis. J. Pet. Geol., 17: 39-70. Iturralde-Vinent, M.A. (Editor), 1996. Ofiolitas y Arcos Volcanicos de Cuba. Project 364, Caribbean Ophiolites and Volcanic Arcs, Special Contribution No. 1, Miami, Fla., 254 pp. Marton, G.L., 1995. Jurassic Evolution of the Southeastern Gulf of Mexico. Ph.D. Dissertation, The University of Texas at Austin, Austin, Texas, 276 pp. Marton, G. and Buffer, R.T., 1994. Jurassic reconstruction of the Gulf of Mexico basin. Int. Geol. Rev., 36: 545-586. Morley, C.K., 1988. Variable extension in Lake Tanganyika. Tectonics, 7: 785-801. Phair, R.L., 1984. Seismic Stratigraphy of the Lower Cretaceous Rocks in the Southwestern Florida Straits, Southeastern Gulf of Mexico. M.A. Thesis, The University of Texas at Austin, Austin, Texas, 319 pp. Phair, R.L. and Buffler, R.T., 1983, Pre-Middle Cretaceous geologic history of the deep southeastern Gulf of Mexico. In: A.W. Bally (Editor), Seismic Expression of Structural Styles m A Picture and Work Atlas. Am. Assoc. Pet. Geol., Stud. Geol., 15 (2): 2.2.3-141. Pszcz6tkowski, A., 1987. Paleogeography and paleotectonic evolution of Cuba and adjoining areas during the Jurassic-Early Cretaceous. Ann. Soc. Geol. Pol., 57: 127-142. Rosendahl, B.R., 1987. Architecture of continental rifts with special reference to east Africa. Annu. Rev. Earth Planet. Sci., 15: 445-503. Salvador, A., 1987. Late Triassic-Jurassic paleogeography and origin of Gulf of Mexico Basin. Am. Assoc. Pet. Geol. Bull., 71: 419-451. Salvador, A., 1991. Triassic-Jurassic. In: A. Salvador (Editor), The Gulf of Mexico Basin. The Geology of North America J, Geological Society of America, Boulder, Colo., 131-180.
E V O L U T I O N OF THE S O U T H E A S T E R N G U L F OF MEXICO BASIN Schlager, W., Buffler, R.T., Angstadt, D., Phair, R.L., 1984. Geologic history of the southeastern Gulf of Mexico. Init. Rep. DSDP, 77: 715-738. Shaub, EJ., 1983. Origin of the Catoche Tongue. In: A.W. Bally (Editor), Seismic Expression of Structural Styles - - A Picture and Work Atlas. Am. Assoc. Pet. Geol., Stud. Geol., 15 (2): 2.2.3-129. Testarmata, M.M. and Gose, W.A., 1984. A paleomagnetic evaluation of the age of the dolomite from site 536, Leg 77,
91
southeastern Gulf of Mexico. Init. Rep. DSDE 77: 525-530. Thomas, W.A., Chowns, T.M., Daniels, D.L., Neathery, T.L., Glover, L. and Gleason, R.J., 1989. The subsurface Appalachian-Ouachita orogen beneath the Atlantic and Gulf Coastal Plains. In: Hatcher, R.D. Jr., Thomas, W.A. and Viele, G.W. (Editors), The Appalachian-Ouachita Orogen in the United States. The geology of North America F-2, Geological Society of America, Boulder, Colo., pp. 445-458.
Chapter 4
The Exposed Passive Margin of North America in Western Cuba
ANDRZEJ PSZCZOLKOWSKI
The Mesozoic successions of western Cuba, now exposed in the Guaniguanico terrane, were deposited to the east of the present NE Yucatan coast. The evolution of these passive margin successions encompasses the syn-rift stage (Early Jurassic?Callovian/early Oxfordian), drift stage (?Callovian/middle Oxfordian-Santonian), and the beginning of the active margin stage (Campanian-Paleocene). Prior to the middle Oxfordian, the San Cayetano basin was located in an originally narrow rift zone formed between Yucatan and South America. The onset of shallow-water carbonate sedimentation in the Sierra de los Organos and Cangre belts occurred in the late Oxfordian or earliest Kimmeridgian. Drowning of a carbonate bank, or platform, in the early Tithonian resulted in a considerable uniformity of facies in all belts of the Guaniguanico terrane, expressed by widespread occurrence of ammonite-bearing limestones and radiolarian microfacies, especially in the upper Tithonian deposits. Pelagic limestones accumulated during the Berriasian and Valanginian, while siliciturbidites occurred in the Northern Rosario, La Esperanza and Placetas belts of western and central Cuba during the Valanginian-Barremian. These belts belonged to a deep-water sector of the basin that extended between the Yucatan and Bahamas platforms. During the Aptian-Albian, siliceous deposition extended across the entire deeper part of the northwestern proto-Caribbean basin. Pelagic carbonate sedimentation resumed in the Cenomanian. Origin of the regional late Turonian (or Coniacian)-Santonian hiatus in the deep-water, pelagic sequence of the northwestern proto-Caribbean basin was probably related to paleoceanographic conditions that existed during Late Cretaceous times. These conditions were associated with paleogeographic changes in the southern part of the proto-Caribbean basin, when the Nicaraguan Rise-Greater Antilles Arc partially closed the connection with the Pacific. During the Campanian, abundant volcaniclastic detritus appeared in the upper Moreno Formation of the Northern Rosario belt. The Bahfa Honda segment of the volcanic arc was located east of the Yucatan block margin and south of the Moreno depocenter. This arc could be the westernmost part of the Greater Antilles Arc (GAA). Unlike previous interpretations, at the end of the Cretaceous a more southerly position of the extinct volcanic arc is inferred from the paleotectonic reconstruction and lithology of the late Maastrichtian deposits. During the Late Paleocene, clastic deposition occurred in a foreland basin setting, in front of a thrust belt along the southern side of the remnant proto-Caribbean Sea.
INTRODUCTION
The Jurassic to Paleocene sedimentary successions of the passive margins of North America are exposed in Cuba (Fig. 1). The Mesozoic platform and/or slope deposits crop out in the northern part of central Cuba. These deposits, traditionally linked to the Bahamas platform (Meyerhoff and Hatten, 1974; Pardo, 1975) occur also in the Matanzas Province (Pszcz6~kowski, 1986b) and in eastern Cuba (Iturralde-Vinent, 1996). In western Cuba (Fig. 2) the Jurassic to Paleocene rocks occur in the Guaniguanico tectonostratigraphic unit (terrane). These rocks are considered to belong originally to the eastern margin and slope of the Yucatan platform (Iturralde-Vinent, 1994, 1996). Metamorphic rocks exposed in the Isla de la Juventud (Isle of Pines)
and in the Sierra de Escambray (Fig. 1) are similar to Mesozoic successions of the Guaniguanico terrane (Khudoley and Meyerhoff, 1971; Mill~n and Myczyfiski, 1978). Stratigraphic and lithologic similarities existing between the Guaniguanico, Pinos and Escambray terranes (Fig. 1) clearly suggest their paleogeographic proximity prior to the Late Cretaceous and Paleogene tectonic events (Pszcz6ikowski, 1981; Iturralde-Vinent, 1994). Studies of the Pinar del Rio geology were carried out by oil companies before 1959, which resulted in many advances in understanding of stratigraphy and tectonics of this area (Hatten, 1957, 1967; RigassiStuder, 1963; Meyerhoff, in Khudoley and Meyerhoff, 1971; Pardo, 1975). Mapping and research carried out in western Cuba during the past 26 years has resulted in publication of many papers, includ-
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 93-121. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
94
A
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Fig. 1. (A) Map of Cuba showing the location of selected geological structures and deep wells. 1 = terranes of passive margin origin exposed in western and south-central Cuba: GU -- Guaniguanico (stratigraphic terrane), P = Pinos (metamorphic terrane in the Isla de la Juventud, or Isle of Pines), E = Escambray (metamorphic terrane in the Sierra de Escambray); 2 - the Placetas and Camajuanf belts in north-central Cuba (Kimmeridgian?/Tithonian to Maastrichtian slope successions) and the Asunci6n metamorphic massif (AN) in eastern Cuba; 3 - the Remedios belt (Cretaceous platform succession)" 4 = well location sites (shown as encircled numbers: 1 Martfn Mesa 1 (situated in the Martfn Mesa tectonic window), 2 = Pinar 1, 3 = Guanahacabibes, 4 = Los Arroyos 1), BH = Bahfa Honda terrane (ophiolite and Cretaceous volcanic arc), Gh = Guanahacabibes Peninsula. (B) Schematic map showing the location of the main terranes and belts in western and central Cuba and adjacent areas (partly after Rosencrantz, 1996 and Case et al., 1984, 1990): a = Yucat;in platform; b = Florida and Bahamas platforms; c = Camajuanf and Placetas belts (undivided) in north-central Cuba (CA & PS)" d -- Yucatfin basin; e = Cayman ridge; f = Camagtiey trench; GU = Guaniguanico terrane in western Cuba; BH -- Bahia Honda terrane (ophiolite and Cretaceous volcanic arc); P -- Pinos terrane; E = Escambray terrane; RS & CC = Remedios and Cayo Coco belts (undivided) in north-central Cuba (shallow-water and pelagic carbonates); SGM = southeastern Gulf of Mexico. Arrows indicate relative movement along faults and barbed continuous lines denote major thrusts.
ing overviews of Cuban geology (Lewis and Draper, 1990; Iturralde-Vinent, 1994, 1996), and geological maps. The present paper focuses on the evolution of the Jurassic to Early Paleocene passive margin successions now exposed in the Guaniguanico terrane of western Cuba, before their Paleogene tectonic deformation.
TECTONIC SETTING
In this paper, a tectonostratigraphic terrane is defined following the criteria of Howell et al. (1985), adopted in some recent Caribbean geological studies (for example, Mann et al., 1991). Also in Cuba some geological structures have been characterized as 'terranes' (Lewis and Draper, 1990; Pszcz6tkowski, 1990; Piotrowska, 1993; Iturralde-Vinent, 1994, 1996; etc.), although the overall, generally accepted scheme of Cuban tectonostratigraphic terranes is still to be achieved. It was not the aim of this paper to propose such a scheme. Rather, the concept of terranes is merely used herein to explain a Mesozoic evolution of a passive margin successions exposed in
western Cuba. Two types of terranes may be distinguished in western and south-central Cuba, namely the stratigraphic and metamorphic terranes. However, in western Cuba both types of terranes are not completely separated, as the Guaniguanico terrane includes also metamorphic rocks (the Cangre belt). The term 'Guaniguanico terrane' was introduced by Iturralde-Vinent (1994). This author (IturraldeVinent, 1994, 1996) proposed a generalized tectonic scheme of the Pinar del Rio Province and placed the Guaniguanico terrane among the southwestern Cuban terranes (with the Pinos and Escambray terranes). In his opinion, the Guaniguanico terrane is composed of five juxtaposed belts: Los Organos, Rosario South, Rosario North, Quifiones and Felicidades. The Guaniguanico terrane (Figs. 1 and 2) is located mainly in the Pinar del Rio Province but its eastern extremity reaches the Havana Province. The Pinar fault forms the southern boundary of this tectonostratigraphic unit. The eastern part of the northern boundary of the Guanigianico terrane is defined by the tectonic contact with the Bahfa Honda terrane and the Guajaibdn-Sierra Azul unit (Fig. 2B and C). To the west, the northern
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA 83o30 '
84000 '
I
,co
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o APS
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,
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1 cm) derived from the extinct volcanic arc, even in the Cangre and Sierra Chiquita tectonic units. The thickest sections of the Cacarajfcara megabed clastic deposits (450 m) were measured in these two tectonic units (Fig. 13), and this fact shows that their position was well to the south (Fig. 20), not at the entrance of the Gulf of Mexico between Florida and Yucatfin, as proposed by Iturralde-Vinent (1992). During the Paleocene, the convergence of the extinct GAA segment with the Bahamas platform margin continued (Fig. 21). The position of the extinct, westernmost GAA segment (occurring now in the Bahia Honda terrane) at the Yucatfin margin as indicated in Fig. 21, results from a tectonic and lithostratigraphic analysis of the Lower Paleocene deposits in the Guaniguanico terrane. In the Northern Rosario belt, the Manacas Formation shales overlying the Cacarajfcara Formation are evidence of a major change from pelagic conditions during the Cretaceous to the Paleogene foreland basin environment. The narrow northwestern sector of the proto-Caribbean basin was either a peripheral or a retroarc foreland basin, located in front of the thrust belt along the southern side of the remnant proto-Caribbean Sea. The discrimination of ancient peripheral and retroarc foreland basins is difficult (Ingersoll, 1988). Within the plate tectonic model accepted herein (Figs. 19-21), the Paleocene basin of western and central Cuba was a peripheral foreland basin. However, according to an alternative plate-tectonic model (Iturralde-Vinent, 1994, 1996) an retroarc foreland basin formed in western and central Cuba during the Paleocene. The material derived from the volcanic suites and ophiolite contributed to the foreland basin deposits
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
117
CONCLUSIONS
Fig. 21. Simplified paleogeographic reconstruction of the northern Caribbean region for the Early Paleocene: Vb -- Vfbora basin (western Cuba), Cf = Cienfuegos basin (south-central Cuba), T.f = La Trocha fault in central Cuba (see Hatten, 1967); 1 -- Caribbean plate oceanic crust, 2 = proto-Caribbean and Atlantic oceanic crust, 3 = pelagic biomicrites and breccias of the Anc6n Formation (western Cuba), 4 = shales and claystones (Manacas and Vega Alta formations), 5 - breccias and calcirudites of the Vega Formation (Camajuanf belt of the Bahamas platform margin), 6 = Remedios belt in north-central Cuba, 7 = syn-sedimentary normal faults. Other symbols as in Fig. 20.
in western Cuba. Pelagic limestones prevailed in the Southern Rosario belt (Anc6n Formation), although with a clear influence of the arc-originated detritus in the Cinco Pesos tectonic unit. Pelagic limestones and breccias are widespread in the Sierra de los Organos belt. The limestone and chert breccias were formed as a result of a considerable erosion of the underlying Cretaceous limestones (in places also Tithonian), along syn-sedimentary fault escarpments (Pszczdtkowski, 1978). These faults, schematically shown in Fig. 21, could originate in a zone of extension induced by bending of the underthrusting plate during arc-passive margin collision (Bradley and Kidd, 1991). The Paleocene-Middle Eocene limestone breccias with a considerable thickness are also known in the Camajuanf belt of west-central and central Cuba (Pszcz6tkowski, 1983). The JurassicCretaceous sedimentary successions deposited on (and along) the passive margin of Yucatfin, now exposed in western Cuba, formed the foreland basin substrate during the Paleocene-Early Eocene. In the Sierra de los Organos belt, a change from the passive margin to a foreland basin occurred during the Early to Late Paleocene. The terminal collision of the extinct volcanic arc with the passive margin occurred in the Late Paleocene to Early Eocene in western Cuba (Bralower et al., 1993; Bralower and Iturralde-Vinent, 1997; Gordon et al., 1997).
The Mesozoic successions of western Cuba, now exposed in the Guaniguanico terrane, were deposited more than 100 km to the east of the present northeast Yucat~in coast. The evolution of these Yucat~in passive margin successions encompasses the synrift stage (Lower Jurassic-?Callovian/early Oxfordian), drift stage (?Callovian/middle OxfordianSantonian), and the beginning of the active margin stage (Campanian-Paleocene). Prior to the middle Oxfordian, the San Cayetano basin was located in a narrow, but steadily widening, rift zone formed between Yucatfin and South America. The advance of the middle Oxfordian transgression resulted in a major facies change, when the San Cayetano deltaic sediments were replaced by shallow-water limestones with bivalves and/or by deeper ammonitebearing deposits. The restoration of the carbonate shallow-water sedimentation in the Sierra de los Organos and Cangre belts and its onset in the Rosario belts occurred in the late Oxfordian or earliest Kimmeridgian. Subsidence kept pace with the relatively high rate of sedimentation in the Sierra de los Organos belt; about 400 to 650 m of shallowwater limestones and dolomitic limestones formed during the Kimmeridgian. Drowning of the shallow-water carbonates in the early Tithonian resulted in a considerable uniformity of pelagic facies in all belts of the Guaniguanico terrane. The Hauterivian-Barremian siliciturbidites that occur in the Northern Rosario, La Esperanza and Placetas belts of western and central Cuba are interpreted to have a common source for the clastic material. These belts probably belonged to the deep-water sector of the basin, which extended between the Yucat~in and Bahamas passive margins. During the Aptian-Albian, the siliceous deposition extended across the entire northwestern, deeper part of the northwestern proto-Caribbean basin. In the Cenomanian, pelagic carbonate sedimentation was restored in the Rosario and Placetas belts. The late Turonian (or Coniacian)-Santonian deposits are very scarce, or even entirely missing in western Cuba, due to non-deposition and the Late Cretaceous and/or Paleocene erosion. During the Turonian-Santonian, the Nicaraguan Rise-Greater Antilles Arc partially closed the connection of the proto-Caribbean basin with the Pacific. Among other factors, this paleogeographic change could create specific paleoceanographic conditions in the northwestern part of the proto-Caribbean basin. The eastern passive margin of Yucatfin was affected by approaching arc terrane in the Campanian. In the Northern Rosario belt of western Cuba, the Moreno Formation contains abundant volcaniclastic material, mainly in its upper part. During the
118 C a m p a n i a n , the volcanic arc was located east of the Yucatfin b l o c k m a r g i n and south of the M o r e n o depocenter. This arc could be the w e s t e r n m o s t part of the G r e a t e r Antilles Arc ( G A A ) , as p r o p o s e d by Pindell and D e w e y (1982) and Pindell et al. (1988). T h e late M a a s t r i c h t i a n deposits of the passive margin of Yucatfin are r e p r e s e n t e d by the Cacarajfcara F o r m a t i o n in the Rosario belts. A m o r e southerly position of the extinct volcanic arc at the end of the C r e t a c e o u s is inferred from the p a l e o t e c t o n i c rec o n s t r u c t i o n and lithology of the late M a a s t r i c h t i a n deposits. D u r i n g the P a l e o c e n e , the s e d i m e n t a r y successions of the G u a n i g u a n i c o terrane, originally deposited on (and along) the passive m a r g i n of Yucatfin, f o r m e d the foreland basin substrate. This foreland basin was located in front of a thrust belt along the southern side of the r e m n a n t p r o t o - C a r i b b e a n Sea.
ACKNOWLEDGEMENTS
T h e author is grateful to Richard T. Buffler, T h o m a s W. Donnelly, John E L e w i s and G y 6 r g y M a r t o n for their review of the manuscript, and to Paul M a n n for his useful c o m m e n t s on the text and figures. T h e discussions with R y s z a r d Myczyfiski and M a n u e l Iturralde-Vinent on s o m e p r o b l e m s concerning the g e o l o g y of Cuba are appreciated.
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THE E X P O S E D PASSIVE M A R G I N OF N O R T H A M E R I C A IN W E S T E R N C U B A Organos, Cuba. Arch. Sci. Soc. Phys. Hist. Nat. G6n~ve, 16: 339-350. Rodriguez, E, 1987. Divisi6n estratigrfifica de la Formaci6n Esperanza y comparaci6n de los cortes de las subzonas Esperanza y Sierra del Rosario. In: Memorias del III Encuentro Cientffico-t6cnico de Geologfa en Pinar del Rfo. Soc. Cubana de Geologfa, pp. 46-50. Rosencrantz, E., 1990. Structure and tectonics of the Yucatan basin, Caribbean Sea, as determined from seismic reflection studies. Tectonics, 9: 1037-1059. Rosencrantz, E., 1996. Basement structure and tectonics in the Yucatan basin. In: M.A. Iturralde-Vinent (Editor), Ofiolitas y arcos volcanicos de Cuba (Cuban ophiolites and volcanics arcs). IUGS/UNESCO Project 364, Contrib. 1, pp. 36-47. Ross, M.I. and Scotese, C.R., 1988. A hierarchical tectonic
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Chapter 5
Stratigraphic Evidence for Northwest to Southeast Tectonic Transport of Jurassic Terranes in Central Mexico and the Caribbean (Western Cuba)
EMILE A. PESSAGNO, JR., ABELARDO CANTIJ-CHAPA, DONNA M. HULL, MICHAEL KELLDORF, JOSE E LONGORIA, CHRISTOPHER MARTIN, XIANGYING MENG, HOMER MONTGOMERY, JAIME URRUTIA FUCUGAUCHI and JAMES G. OGG
Jurassic and Early Cretaceous stratigraphic data from terranes in Central Mexico situated southwest of the Walper Megashear demonstrate similar records of paleobathymetry and tectonic transport. In general, each of these terranes shows the same paleobathymetric fingerprint: (1) marine deposition at inner neritic depths during the Callovian to early Oxfordian (Middle to Late Jurassic); (2) marine deposition at outer neritic depths during the late Oxfordian (Late Jurassic); (3) sudden deepening to bathyal or upper abyssal depths (ACD = aragonite compensation level) from the early Kimmeridgian (Late Jurassic) until the end of the Cretaceous. This paleobathymetric fingerprint differs markedly from that occurring to the east-northeast of the Walper Megashear in the Coahuiltecano terrane (emended herein: ~Sierra Madre Oriental terrane). In the Coahuiltecano terrane (e.g., Peregrina Canyon near C. Victoria, Tamps.), no Mesozoic marine deposits older than late Oxfordian occur. The paleobathymetric fingerprint of this terrane was (1) inner neritic during the late Oxfordian (Late Jurassic) to ~Barremian (Early Cretaceous) and (2) bathyal to abyssal during the remainder of the Cretaceous (Aptian to Maastrichtian). Though varying in detail, each succession that has been examined in the mosaic of suspect terranes to the southwest of the Walper Megashear shows evidence of tectonic transport from higher latitudes to lower latitudes during the late Middle Jurassic, the Late Jurassic, and the Early Cretaceous. For example, the paleolatitudinal signature of the San Pedro del Gallo terrane (Durango) supplied by faunal data (radiolarians and megafossils) and preliminary paleomagnetic data indicates that this terrane was transported tectonically from higher paleolatitudes (Southern Boreal Province: ~40~ during the Late Jurassic (Oxfordian) to lower paleolatitudes (Tethyan Realm: Northern Tethyan Province) by the Early Cretaceous (Berriasian). The Jurassic and Lower Cretaceous successions at Mazapil (Zacatecas), Sierra de la Caja (Zacatecas), Sierra de Zuloaga (Zacatecas), Symon (Durango), and Sierra de Catorce (San Luis Potosi) are all genetically related to that at San Pedro del Gallo. They are regarded as representing dismembered remnants of the San Pedro del Gallo terrane. Faunal data (radiolarians and megafossils) from the Mazapil succession (Sierra Santa Rosa) indicate that this remnant of the San Pedro del Gallo terrane was situated at Southern Boreal paleolatitudes (>30~ during the Oxfordian and Kimmeridgian and at Northern Tethyan paleolatitudes (22 to 29~ during the Tithonian and Berriasian. Preliminary paleomagnetic data from the upper Tithonian to Berriasian part of the Mazapil succession indicates 25~ Farther to the southeast (San Luis Potosi, Hidalgo, Veracruz, Puebla) in the Huayacocotla segment of the Sierra Madre Oriental, previous investigations indicate tectonic transport from Southern Boreal paleolatitudes (>30~ during the Callovian to Northern Tethyan paleolatitudes (22 ~ to 29~ during the Kimmeridgian and Tithonian and to Central Tethyan paleolatitudes (30~ to lower Tethyan paleolatitudes during the Middle Jurassic, Late Jurassic, and Early Cretaceous.
IMPORTANCE
OF FAUNALAND FLORAL DATAIN
PALEOGEOGRAPHIC RECONSTRUCTIONS
Much of Mexico west of the Walper Megashear of Longoria (1985a,b, 1986, 1987, 1994) consists of suspect terranes or displaced terranes (Fig. 1). Previous studies by Taylor et al. (1984), Pessagno and Blome (1986), Pessagno et al. (1986, 1993a,b), and Montgomery et al. (1992, 1994a,b) have established the importance of faunal and floral data in paleogeographic reconstructions in North America as well as in the Caribbean. Recognition of displaced tectonstratigraphic terranes depends primarily on paleolatitudinal data derived from paleontology and paleomagnetism. Faunal and floral data can be used to constrain existing paleomagnetic data and in some cases can also help determine whether tectonostratigraphic terranes originated in the Northern or Southern Hemisphere or in the Eastern or Western Pacific. For example, paleomagnetic data presented by Jones et al. (1977) indicate that the Wrangellia terrane originated 15~ north or south of the Triassic paleoequator. During the Late Triassic and Early Jurassic both the molluscan and radiolarian assemblages tend to be predominantly Tethyan in origin (Tipper, 1981; Taylor et al., 1984; Pessagno and Blome, 1986). However, the discovery by Taylor et al. (1984, pp. 128, 135) of very rare Boreal ammonites (amaltheids) in the upper Pliensbachian por-
Fig. 1. Map showingapproximateposition of area of displaced and suspect terranes west of WalperMegashear of Longoria.
N W TO SE T E C T O N I C T R A N S P O R T
OF JURASSIC TERRANES
IN MEXICO AND THE CARIBBEAN
125
Fig. 2. Index map showing important Jurassic localities in Mexico and Cuba. The most important localities for this report are 3-18, 24. Key to localities: 1 = Tlaxiaco: Sierra Madre del Sur, Oaxaca. 2 = Pletalcingo: Sierra Madre del Sur, Puebla. 3 = Huayacocotla Anticlinorium: Taman-Tamazunchale, San Luis Potosi; Huayacocotla, Veracruz; Huachinango, Puebla. 4 -- Sierra Catorce, San Luis Potosi. 5 = Sierra Santa Rosa, Zacatecas. 6 = Sierra de la Caja, Zacatecas. 7 = Sierra Cadnelaria, Zacatecas. 8 = Sierra Sombretillo and Sierra Zuloaga, Zacatecas. 9 = Sierra de Ramirez, Zacatecas-Durango. 10 - Sierra de Chivo, Durango. 11 -- Sierra de Palotes, Durango. 12 --- San Pedro del Gallo, Durango. 13 = Santa Maria del Oro, Sierra de la Zarca, Durango. 14 = Sierra Vieja-Arroyo Doctor, Tamaulipas. 15 = Huizachal Anticlinorium, Tamaulipas. 16 =- Sierra Galeana-Iturbide, Nuevo Leon. 17 = Sierra de Parras, Coahuila. 18 - Sierra de Jimulco, Coahuila. 19 = Sierra Menchaca, Cohuila. 20 = Sierra Plomosas-Place de Guadalupe, Chihuahua. 21 -- Sierra E1 Cuchillo Parado, Chihuahua. 22 = Sierra de Samalayuca, Chihuahua. 23 = Sierra de Cucurpe, Sonora. 24 = Cordillera de Guaniguanico, Cuba. Base map partly derived from that in Salvador et al. (1992).
-30 ~
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r
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- 7
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30 ~ '~'~.
"31
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if,, g '/~\
9
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~.
e
Fig. 3. Megashear map of Longoria (1994).
\
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126 tion of the Maude Formation of the Queen Charlotte Islands (B.C.) indicates that Wrangellia was situated in the Northern Hemisphere. Moreover, the presence of the pectenacid Weyla in Lower Jurassic strata in the Queen Charlotte and other remnants of the Wrangellia terrane indicates that this terrane originated in the Eastern Pacific (Smith, 1980; Taylor et al., 1984; Pessagno and Blome, 1986; Pessagno et al., 1986). Although the focus of the present report is on geologic terranes immediately southwest of and adjacent to the Walper Megashear, it is worth noting that faunal data from a variety of sources suggest either northwest to southeast movement or southeast to northwest movement along other possible megashears (Fig. 3). In the states of Oaxaca and Guerrero Burckhardt (1927, 1930) was the first to record the presence of a Middle Jurassic (Bajocian to Callovian) ammonite assemblage which strongly resembles that in the Andes, specifically Argentina (Figs. 1 and 2). Subsequently, the strong Andean affiliation of the Middle Jurassic ammonite assemblage of Oaxaca and Guerrero has been noted by Arkell (1956), Imlay (1980), Sandoval and Westermann (1988), Sandoval et al. (1990), and von Hillebrandt et al. (1992). The fact that no Middle Jurassic ammonite faunas with strong Andean affiliation are known from elsewhere in Mexico and North America suggests that the Middle Jurassic (Bajocian to Callovian) succession in these states has undergone southeast to northwest tectonic transport along a more outboard megashear paralleling the Walper Megashear (i.e., the Cserna Megashear of Longoria, 1994; Fig. 3). In the Vizcaino Peninsula (Baja California Sur; Fig. 1), radiolarians, though abundant and well-preserved in strata of Early Jurassic (Pliensbachian) to Late Cretaceous (Cenomanian) age, are representative of a poorly diversified Boreal assemblage (Whalen and Pessagno, 1984; Whalen, 1985; DavilaAlcocer, 1986; Pessagno et al., 1986). Foraminifers occurring in the Late Jurassic to Late Cretaceous strata of the Vizcaino Peninsula are likewise Boreal in nature and show strong affinity to those of the California Coast Ranges. In fact, studies made by Longoria (unpublished PEMEX report) indicate that the planktonic foraminiferal assemblage occurring in the Upper Cretaceous Valle Formation is like that described by Douglas (1969) from the Great Valley Supergroup (California Coast Ranges). This fauna is not well-developed south of the latitude of Bakersfield in California and occurs northward to Alaska and to Japan. In addition to the foraminifers and radiolarians the Late Jurassic (Tithonian) to Early Cretaceous (Valanginian) strata of the Eugenia Formation contain a Boreal bivalve assemblage characterized by species of Buchia (cf. Fig. 5).
E.A. PESSAGNO et al. RADIOLARIAN PALEOLATITUDINAL MODEL
It is now apparent from our analyses of radiolarian faunal data from North America and elsewhere in the world that radiolarians can be utilized in paleobiogeographic investigations and to monitor the tectonic transport of terranes both in the Northern and Southern Hemispheres. Much of the Circum-Pacific margin is comprised of a collage of tectonostratigraphic terranes. Many of these terranes have been displaced paleolatitudinally for hundreds, or in some cases, possibly thousands of kilometers. Circum-Pacific Jurassic paleogeography, as a result, is difficult to discuss in simplistic terms and must be viewed through this complex mosaic. In the Northern Hemisphere Pessagno and Blome (1986) and Pessagno et al. (1986, 1987, 1993a,b) divided the Tethyan Realm into a Central Tethyan Province characterized by a radiolarian assemblage with high pantanelliid abundance and diversity and the absence of Parvicingula/Praeparvicingula and into a Northern Tethyan Province with high pantanelliid abundance and diversity and common Parvicingula/Praeparvicingula (Fig. 4). The Boreal Realm was subdivided into a Southern Boreal Province and a Northern Boreal Province. The Southern Boreal Province is characterized by a sharp decline in pantanelliid abundance and diversity and by the abundance and diversity of species of Parvicingula/Praeparvicingula; the Northern Boreal radiolarian assemblage is distinguished by abundant Parvicingula/Praeparvicingula and by its total lack of pantanelliids. In Fig. 4 the boundary between the Tethyan Realm and Boreal Realm is placed at ~30~ the boundary between the Central Tethyan Province and the Northern Tethyan Province is established by associated paleomagnetic data at ,~22~ (Pessagno et al., 1987; Yeh and Cheng, 1996). Pessagno and Blome (1986) originally proposed that the model for the Southern Hemisphere is the mirror image of that in the Northern Hemisphere. Subsequently, new data from the Southern Hemisphere has substantiated this model from Argentina (Pujana, 1989, 1991, 1993, 1996), New Zealand (Aita and Grant-Mackie, 1992), Antarctica (Kiessling, 1995; Kiessling and Scasso, 1996), and the Sula Islands (Pessagno and Hull, in prep.).
PALEoLATITUDINAL RECONSTRUCTIONS USING MULTIPLE CRITERIA
Although radiolarian paleobiogeographic reconstructions are useful and can stand alone, they are far more effective when combined with information derived from paleomagnetism, analysis of the total faunal and floral assemblage, and other criteria
NW TO SE TECTONIC TRANSPORT OF JURASSIC TERRANES IN MEXICO AND THE CARIBBEAN
127
Fig. 4. Paleolatitudinal model based on distribution of selected radiolarians from the Jurassic and Lower Cretaceous.
having paleolatitudinal or paleolongitudinal significance (see Fig. 5). The tenet stressed herein is that paleogeographic reconstructions should use all criteria available, where possible, and should not focus on any one facet (e.g., analysis of only the ammonite assemblage). Nevertheless, even in eugeoclinal terranes, where other fossils are absent, it is often possible to determine the relative paleolati-
TETHYAN REALM I [ ! i I,
CENTRAL TETHYAN PROVINCE
~--
-
NORTHERN TETHYAN PROVINCE
TETHYAN AMMONITES
I
i
'~..
.t,t"
i~...........7;,.... I
E
,'---.-.~ ....
/ ..........i.;'A../
g
]
~.1
' ~,
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,,~
k
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i
Interbedded phosphatic limestone, black chert and red calcareous mudstone. Abundant Radiolaria and ammonites. Thickness = 13.3 m
o:I
"UNIT E"
,
Upper 5.64 m consisting of medium-bedded light gray micrite and late Tithonian Substeueosceras spp. ~ thin-beddded black chert, Remainder of unit consisting of thin-bedded, ' i Paradontoceras aft. i buff-weathering calcareous siltstone and black chert (2). Abundant Radiolaria upper part may be Berriasian and ammonites. callistoides, Durangites sp. Thickness = 23.5 m
O
Faunal Realm/Province
Paleobathymetry
Calpionellids calcified Radiolaria
'
-~
Diagnostic Faunal Elements
Thurmannites spp., Asteria aft. psilostoma .- fide Burckhardt (1930).
c-
Fig.
,
.
eL..
~c'~
o c-
Outer Neritic
o CL. c~ cD
i middle Oxfordian
o Z
Dichotomosphinctes
__, middle Oxfordian Nerinea,bivalves, corals, or and sponge spicules older See Burckhardt (1910, 1930)
Inner Neritic
O
eL.. cD t---5 o
Stratigraphic summary for Mazapil remnant.
and steep continental slopes at sites of upwelling of nutrient-rich waters. Whether this scenario could exist in the distal backarc setting characterizing all San Pedro del Gallo remnants is questionable. Phosphatic limestones and shales were recorded by Burckhardt (1930) in the Sierra Santa Rosa, Sierra de la Caja, and Sierra de Zuloaga (Fig. 2.: Locs. 5, 6, 8). They are not known from San Pedro del Gallo, Sierra Catorce, the Huayacocotla Anticlinorium, or from western Cuba (Fig. 2: Locs. 3, 4, 12, 24). An alternative to the upwelling model may be a large kill of fish and other organisms by a red tide, producing an abundance of phosphatized bones and other material at lower bathyal or upper abyssal depths. The upper Oxfordian to lower upper Tithonian part of the La Caja Unit E (Fig. 7) contains Buchia, Tethyan ammonites, abundant Parvicingula/Praeparvicingula, and poorly diversified pantanelliids indicative of Southern Boreal paleolatitudes. The remainder of Unit E and all of Units D, C, B, and A are assignable to the Northern Tethyan Province based on the presence of abundant, diversified pantanelliids, abundant to common Parvicingula/Praeparvicingula, and the presence of calpionellids (Units A and B, Fig. 5). These data indicate that the Mazapil remnant of the SPG terrane has been transported from Southern Boreal paleolatitudes (>30~ to Northern Tethyan paleolatitudes
(22~ during the early Kimmeridgian to late Tithonian interval. It should be noted that Ogg's preliminary paleomagnetic data for the late Tithonian indicate 25~ Sierra de Catorce remnant (Fig. 2" Loc. 4; Figs. 9, 11, 14) Erben (1956, p. 46), Carrillo-Bravo (1961, p. 42), and Imlay (1980) noted the presence of red phylitic shale with Sinemurian ammonites in the Sierra de Catorce. These strata are the chronostratigraphic equivalent (part) of the Huayacocotla Formation of east-central Mexico. The Upper Jurassic succession at the Sierra de Catorce is much like that of other remnants in terms of its lithostratigraphic and paleobathymetric signatures. The Sierra de Catorce succession begins with the massively bedded, micritic, inner neritic Zuloaga Limestone which is identical to that at San Pedro del Gallo and at Mazapil (Web Fig. 5.10). Verma and Westermann (1973) divided the La Caja Formation in the Sierra de Catorce into two members: a lower E1 Pastor Member and upper E1 Verde Member. The E1 Pastor Member consists of massively bedded darkto medium-gray micritic limestone with interbedded black chert, siltstone, sandstone, and shale. The E1 Verde Member includes thin-bedded micritic lime-
NW TO SE TECTONIC TRANSPORT OF JURASSIC TERRANES IN MEXICO AND THE CARIBBEAN Litho Unit
Description
unnamed limestone unit
Correlated incorrectly by Verma and Westermann (1973) and Imlay (1980) with the La Taraises Formation which has its type area in Sierra de La Parra. Probably the Chapulhuacan Limestone. Thickness is unknown.
Light gray, massively bedded micrite with yellow, oval chert concretions.
Age
Diagnostic Faunal Elements
139 Faunal
Paleobathymetry Realm/Province
Valanginian
Olcostephanus sp p.
O3 c-
Berriasian?
o m
El Verde Member
Thin-bedded medium gray micrite, black chert, light gray siltstone and sandstone. Chert and micrite with abundant Radiolaria. Siltstone and sandstone are carbonate wacke (turbidite) frequently showing graded bedding withgrading of molluscan fragments. Chert and micrite with abundant Radiolaria. Abundant ammonites occur in micrite, Thickness = - 25 m,
late Tithonian
Kossmatiaspp, Durangitesspp. Corongocerasspp, Substeuerocerasspp., Berriasellaspp.
oo oo
c~ ..K:=
cD c-
E
.+_,
o u_
late early
-~, ,
c--
El Pastor
Member
Zuloaga
Limestone
Cahuasas
Formation
Kimmeridgian Massively bedded medium to dark gray micrite interbedded with thin-bedded to black radiolarian chert and siltstone or sandstone. Siltstone and sandstone early Tithonian; representing carbonate wacke occurs in layers up to 0.4 m thick and shows late Kimmeridgian grading of molluscan fragments. Massive micrite beds averaging about is missing; early 0.8. Closely resemble micrites of "lower massively bedded limestone Tithonian strata member"of the Taman Formation in Huayacocotla remnant. Micrite and assumedly overlie black chert with abundant Radiolaria Micrite with abundant ammonites. early Kimmeridgian Buchia noted by Burckhardt (1930). Thickness = - 28 m. strata unconformably
Massively bedded micritic limestone with nodules of black chert.. According to Imlay (1980, p. 45) Zuloaga rests unconformably on the continental red beds of the Cahuasas Formation. Thickness = 200 m fideVerman and Westermann (1973).
o
Mazapilites mexicanus, Pseudolissoceras zitteli
Abundant Radiolaria
Idocerasspp, in lower part together with Buchia concentrica
middle Oxfordian Nefinea, bivalves, corals, or and sponge spicules older See Burckhardt (1910, 1930)
Continental red beds. Red shale, sandstone, and siltstone. Probably Congomerate at base. Overlies Lower Jurassic (Sinemurian) strata of Bathonian to Huayacocotla Group. The latter unit contains Sinemurian ammonites. This unit Bajocian based on was assigned to the La Joya Formation by Imlay. However, because it overlies age in Huayacocotla the Huayacocotla Group, it is assigned here to the Cahuasas. See Figure 13. remnant. Thickness = 120 to 160 m.
z
Unfossiliferous
Inner Neritic
NON MARINE
Fig. 14. Stratigraphic summary for Sierra Catorce remnant.
stone, black thin-bedded radiolarian chert, thin-bedded siltstone and sandstone, and shale. The siltstone and sandstone in both members often display graded bedding with graded molluscan fragments and are interpreted as turbidites (Web Figs. 5.10 and 5.11). These strata were deposited at upper abyssal depths rather than inner neritic depths as incorrectly suggested by Verma and Westermann (1973) and Salvador et al. (1992). Most stratigraphers in the past have totally ignored the far more abundant microfauna. Because they failed to examine the rocks in thin-section, they reached erroneous conclusions concerning water depth. Burckhardt (1930) recorded the presence of several species of Buchia from the Kimmeridgian portion of La Caja Formation at Sierra de Catorce. These Boreal bivalves are associated with Tethyan ammonites and Parvicingula/Praeparvicingula (thin-section analysis). Hence, we conclude that the Sierra de Catorce remnant was situated in the Southern Boreal Province during the Oxfordian to Kimmeridgian interval. Tithonian La Caja strata contain Tethyan ammonites, Parvicingula/Praeparvicingula and lack Buchia. Hence, during the Tithonian the Sierra de Catorce remnant was probably situated in the Northern Tethyan Province. Our investigations of the strata in this area are not, however, as complete as they are in the San Pedro
del Gallo remnant, the Mazapil remnant, and the Huayacocotla remnant. Huayacocotla remnant (Figs. 9, 11, 15) The Mesozoic succession begins in this area with the deposition of the Lower Jurassic (Sinemurian to lower Pliensbachian) Huayacocotla Formation which consists of black shale, mudstone, and graywacke. A rich ammonite assemblage occurs in the lower and middle parts of the unit (Burckhardt, 1930; Erben, 1956; Imlay, 1980). Bivalves and land plants have been recorded from the upper part. The Huayacocotla Formation overlies the Mississippian to Lower Permian strata of the Guacamaya Formation with marked hiatus associated with an angular unconformity. According to Nestell (1979) the Guacamaya Formation contains fusulinids with South American affinities. This likewise seems tobe true of fusulinids occurring in the Guacamaya Formation at Peregrina Canyon near Ciudad Victoria (Tamps.). The thickness of the unit varies from 560 to 1200 m. As far as can be determined, references to the presence of the Upper Triassic Huizachal Formation (continental red beds) in this area are erroneous. The Huizachal has largely been confused with the Middle Jurassic Cahuasas Formation which also consists of continental red beds (cf. Imlay, 1980).
140
E.A. PESSAGNO et al. Litho Unit
Chapulhuacan
Limestone
Pimienta Formation
Description
Age
Medium to massively bedded very fine-grained cream to light gray micrite with abundant Radiolaria, calpionellids, and rare ammonites. Micrite with black chert nodules and lenses. Thickness = - 30 m.
Berriasian
Thin-bedded cream colored to light gray micrite interbedded with dark gray shale, common black radiolarian chert, and light green vitric tuff. Micrite with abundant Radiolariaand calpionellids together with rare ammonites and common sponge spicules. Thickness = - 200 m.
to Valanginian.
late Tithonian
Diagnostic Faunal Faunal Elements Paleobathymetry Realrn~rovince Subthurmannia sp. i Neolissoceras sp. Spiticeras sp. Thurmanniceras sp, :Paradontocerasaff, callistoides Durangites, Substeueroceras j +
r"c~ t"
I c,o
SubzoneZletRadiolaria at some localitiesto south
,
>, '-'
c 0
!
d o
oi
"upper thin-bedded
~ limestonemember" II b-c---,
Thin-bedded dark gray to medium gray micrite with thick interbedds of dark to medium gray shale. Shale layers with abundant micrite nodules. Micrite nodules and beds with abundant Subzone 4~ Radiolariasiliceous sponge spicules and commonammonites.
Durangites, Kossmatia, Salinites grossicostatum
r t-"
+
late Tithonian
very abundant Subzone
O
4~ Radiolaria.
a.. t'-
Ataxicoceras, Idoceras, Massively bedded dark gray to medium gray micrite interbedded with thin"Glochiceras fialar" bedded to medium bedded shale. Upper part of unit with numerous dark gray ~ "lowermassively micrite nodules. Massive micrites and micrite nodules with abundant Radiolaria early Kimmeridgian Mazapilites, to "c-- beddedlimestone rare to common ammonites and common pectenacids (Aulacomyella). Hyaline Virgatosphinctes + Radiolaria early Tithonian assignable to Subzone 2ctl, =~ member" calpionellids occur in basal Zone 4, Subzone 4[3 strata at same horizon as lower Tithonian ammonite Mazapilites. Zone 3, & Subzone 4l]. i
c-, t-.,
t--el3 t"-
i
Outer Neritic Santiago
Formation
Silty black shale, mudstone, micrte. Containing ammonites. Radiolaria occurring in upper-most part. Thickness = N169 m.
early Callovian to late Oxfordian
c-"
?
Reineckeia, Dichotomosphinctes, Discosphinctes, Ochetoceras.
I !
Reineckeia Neuqueniceras
!
O Z
I Tepexic Limestone
Calcarenite containing ammonites and bivalves. Thicknes -- - 39 m.
late early Callovian t
Cahuasas
Formation
Continental red beds. Dominantly red shale, siltstone, sandstone, and conglomerate. Commonly cross-bedded. Overlies Lower Jurassic (Sinemurian) strata of HuayacocotlaGroup.The latter unit contains Sinemurian ammonites. Thickness = 40-1200 m. (1)
Bathonian to Bajocian.
Fossil plants.
NON MARINE
Fig. 15. Stratigraphic summary for H u a y a c o c c o t l a remnant.
The Cahuasas Formation consists of 40 to 1200 m of red arkosic sandstone, conglomerate, and shale that rest with angular unconformity on the Huayacocotla Formation (Imlay, 1980, p. 49). Imlay indicates that the Cahuasas must be older than Callovian because where it crops out on the surface it lies disconformably below marine beds of early to middle Callovian age. In the subsurface, however, it underlies latest Bathonian to early Callovian marine shale of the Palo Blanco Formation (see Palo Blanco below). Imlay also indicates that the Cahuasas must be younger than Toarcian (Early Jurassic) in that it passes downward into plant-bearing beds which are early Middle Jurassic. The Cahuasas Formation is overlain disconformably in surface outcrops by the inner neritic early Callovian Tepexic Limestone. The Tepexic is a calcarenite containing common to abundant Liogryphaea nebrascaensis and ammonites such as Neuquenisceras neogaeum and Reineckeia (Cantd-Chapa, 1969, p. 19; Imlay, 1980, p. 50). In the subsurface and at some surface localities a inner neritic black shale unit, the Palo Blanco Formation (Cantd-Chapa, 1969, p. 5; Imlay, 1980, p. 49), underlies the Tepexic Limestone and rests disconformably on the Cahuasas. The Palo B lanco Formation contains the late Bathonian to early Callovian ammonite Kepplerites (Cantd-Chapa, 1969, p. 5; Imlay, 1980).
The Tepexic Limestone is overlain conformably by silty black shale, siltstone, and silty micritic limestone constituting the Santiago Formation (middle Callovian to upper Oxfordian). The lower and middle parts of the Santiago contain bivalves (e.g., small Ostrea, senior author's observations) and ammonites; microfacies analysis by Longoria (1984) indicates that most of this unit was deposited at inner neritic depths. The uppermost (upper Oxfordian) part of the Santiago Formation (e.g., at Taman, S.L.E) contains common radiolarians as well as ammonites (Pessagno et al., 1987). These Santiago strata reflect the same sudden change in water depths from inner neritic to outermost neritic during the late Oxfordian that was noted in the San Pedro del Gallo and Mazapil remnants. The Santiago Formation is overlain conformably by the Taman Formation (sensu Pessagno et al., 1984, 1987). The Taman Formation (thickness about 30-60 m) consists of two informal units (Web Figs. 5.12 and 5.13): (1) a massively bedded to medium-bedded micritic limestone member (lower Kimmeridgian to upper Tithonian), and (2) a thin-bedded micritic limestone member (upper Tithonian) (Pessagno et al., 1984, 1987). Both members of the Taman Formation contain profusely abundant radiolarians, rare foraminifera (chiefly Textulariina), common siliceous sponge spicules, ammonite aptychi, and
NW TO SE TECTONIC TRANSPORT OF JURASSIC TERRANES IN MEXICO AND THE CARIBBEAN occasional ammonites. The abundance of radiolarians together with the sparse benthonic foraminiferal assemblage and the rarity of ammonites suggests that Taman strata were deposited at upper abyssal depths at or somewhat below the ACD (aragonite compensation level) (see microfacies analysis in Longoria, 1984). The Taman is overlain conformably by the latest Tithonian (Late Jurassic) to Berriasian (Early Cretaceous) Pimienta Formation (sensu Pessagno et al., 1984, 1987) and overlain conformably by the Chapulhuac~in Limestone Berriasian to Valanginian). The Pimienta Formation includes 200-400 m of light-gray thin-bedded micritic limestone with thick shale intervals, thin-bedded black radiolarian chert, and light green vitric tuff; it contains abundant radiolarians, calpionellids, siliceous sponge spicules, and common ammonites (Fig. 10). Pimienta deposition likewise took place at upper abyssal depths somewhat above the ACD (compensation level of aragonite). The Chapulhuac~in Limestones consists of about 30% of medium to massively bedded, very finegrained, cream to light-gray micrite with abundant radiolarians, calpionellids, nannoconids, and planktonic foraminifera, and rare ammonites at most localities (senior author's observations and those of Longoria, 1984, p. 69); Chapulhuac~in strata were also deposited at upper abyssal depths somewhat above the ACD. Deposition continued at these depths during the remainder of the Cretaceous. The upper Bathonian (Middle Jurassic) to upper Oxfordian (Upper Jurassic) part of the succession contains Boreal megafossils such as the ammonite Kepplerites in the Palo B lanco Formation (Cantf-Chapa, 1969, p. 5; Imlay, 1980, p. 50). Elsewhere in western North America this ammonite is known from Middle Jurassic strata in the Sierra Nevada, from the upper Bathonian part of the Snowshoe Formation, Izee terrane (east-central Oregon), and from Boreal Middle Jurassic strata as far north as Alaska (Imlay, 1980; Pessagno and Blome, 1986; Pessagno et al., 1986, 1987). The lower Kimmeridgian to upper Tithonian (Upper Jurassic) part of the succession (Taman Formation sensu Pessagno et al., 1984, 1987) contains a rich Northern Tethyan radiolarian assemblage characterized by the abundance and diversity of pantanelliids and by the presence of common to abundant Parvicingula and Praeparvicingula. Calpionellids (Tethyan) occur in the upper Tithonian part of the Taman Formation. Moreover, the megafossil assemblage is Tethyan in aspect (Figs. 4 and 5) (see Imlay, 1980; Cantf-Chapa, 1989). The Pimienta Formation as well as the overlying Chapulhuac~in Limestone is characterized by a Tethyan ammonite assemblage and by a microfossil assemblage includ-
141
ing abundant calpionellids and nannoconids lacking
Parvicingula/Praeparvicingula. This association of faunal elements is indicative of the Central Tethyan Provinces (Figs. 4 and 5). These data indicate that the Huayacocotla remnant of the SPG terrane underwent tectonic transport from Southern Boreal paleolatitudes (> 30~ during the late Bathonian (Middle Jurassic) to Northern Tethyan paleolatitudes by the early Kimmeridgian (Late Jurassic) to Central Tethyan paleolatitudes (t "I
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382 In the arcward areas of the wedge, the influence of mud diapirism increases and tectonic disturbances are induced by fault reactivation and secondary faulting. The superimposition of clay diapiric structures on deformations linked to tectonic accretion and oceanic basement ridges leads to complex basin morphologies bounded by steep topographic features (Fig. 16). The progressive tectonic evolution from frontal to arcward basins is associated to a change in the regional topography with a slope angle varying from about 1.5 ~ in the zone of initial accretion to about 0.25-0.4 ~ in the arcward zone. This flattening of the topography is evidenced in more than 100-kmlong profiles (Fig. 16) at a larger scale than the convex shape induced by compaction and cohesion increase in the very frontal part (Zhao et al., 1986).
Mechanics of the accretionary wedge In active margins, wedges develop and deform until a critical taper is attained. They then slide stably, continuing to grow self-similarly as additional material is accreted at the toe. A critically tapered wedge that is accreting fresh material deforms internally while sliding in order to accommodate the
E HUYGHE et al. influx and to maintain a constant mean slope (Davis et al., 1983). Assuming a constant subduction rate and a constant basin sedimentary column, Le Pichon et al. (1990) have shown that the horizontal accretion velocity dramatically decreases through time whereas 'internal thickening' increases. This 'internal thickening' is mainly related to motion along thrusts (Mulugetta, 1988) and out-off-sequence reactivation (Chalaron et al., 1995). The development of prominent relief on the sea floor at the hangingwall of the faults and the sedimentary pattern of the SBRC piggyback basins attest for such reactivations. The source of the extensive clay diapirism of the SBRC is provided by subcretion of muddy Miocene sediments occurring at the footwall of a major ramp of the decollement (Brown and Westbrook, 1988; Dia et al., 1997). These muds are overpressured and induce ductile deformation in the lowermost part of the wedge beneath the arcward basins. Adaptations of the critical wedge model (Dahlen, 1990; Williams et al., 1994) suggest that such a ductile deformation in the lowermost part of a wedge would induce the flattening of its topography. We suggest that the brittle-ductile wedge model applies to the SBRC. The frontal areas are char-
Fig. 15. (a) SAR image of the gravity structures on the flank of mud domes. 1 - scarps; 2 -- compressional ripples; (b) -- track of seismic profile shown in (b). (b) 3.5-kHz seismic profile 9031 showing the slumps and the associated erosion that affected the departure zone of the features. Horizontal scales are the same for (a) and (b). (c) Line drawing of SAR image shown in (a). The SAR interpretation is drawn on the bathymetric map (1 = slumps; 2 = compressional ripples). N-S grey line outlines the axis of the mud volcanoes alignment. (d) Line drawing of (b). Horizontal scales are the same for (c) and (d).
LATE NEOGENE PIGGYBACK BASINS ON THE BARBADOS RIDGE COMPLEX
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acterized by a brittle behavior of the wedge and mean slope angle of about 1.4 ~ (Fig. 16g) and the major principal stress direction (cq in Fig. 16g) is not horizontal, instead making an angle relative to the basal decollement (Davis et al., 1983). This angle is nonetheless small as fluids pressure along the decollement zone (Moore, 1989; Labaume et al., 1995; Bangs et al., 1996) reduce the shear stress along the decollement. The high traction along the decollement near the toe favors tectonic activity on forward verging thrusts (Davis and Lillie, 1994) rather than the initiation of steeper backthrusts. In the arcward zone, the flattening of the regional topography reflects the ductile deformation of the lowermost part of the wedge. The major principal stress direction becomes nearly horizontal above the ductile zone of deformation. As a result, backthrusts and pre-existing thrusts may equally be reactivated (Davis and Lillie, 1994), as evidenced by the opposite wedge-shaped geometries of the sedimentary bodies developed back of both thrusts and back-thrusts (Fig. 11). Therefore distinction between arcward and frontal basins is based on the occurrence of abundant mud diapirism and on the regional flattening of the topography that are both related to changes in the rheology of the deep material of the wedge.
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Sedimentary environment E Piggyback basins have been initially defined in shallow marine conditions (Fig. 1, adapted from Ori and Friend, 1984) inducing temporary erosions under subaerial conditions. Piggyback basins also developed under continental conditions as in the Subandean Belt of Bolivia (Baby et al., 1995), where the greatest piggyback basins of the world occur. The late Neogene piggyback basins of the SBRC formed under deep marine conditions. As a result, the sediment supply is small compared to other piggyback basins (see Table 1). The size and sedimentation rate of the SBRC piggyback basins are similar to those of the Sub-Himalayan piggyback basins. The reasons for such a small supply are quite different: in the Sub-Himalayan Belt (Chalaron et al., 1995; Bilham et al., 1997) the sedimentation rate is controlled by the base-level of the rivers, whereas in the SBRC most of the terrigenous sediment supply transits from the Guyana Margin to the abyssal plain through canyons (Damuth, 1975) and a small sedimentation occurs in the piggyback basins (0.12 mm/yr for the late Quaternary at 11~ latitude, from
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TECTONIC EVOLUTION OF THE GRENADA BASIN Table 2 Back-arc basins of the world Basin
Subduction zone
Aegean Andaman Banda Bismark Fiji Plateau Grenada Havre Kurile Lau Mariana Okinawa Parece-Vela Sea of Japan Shikoku South Aleutian South Fiji South Sandwich Tyrrhenian Yucatan
Hellenic Andaman Java New Britian New Hebrides Lesser Antilles Kermadec Kurile Tonga Mariana Ryuku Mariana Japan Bonin Aleutian Kermadec, Tonga South Sandwich Hellenic Cuban
Bandy and Hilde, 1983; Brooks et al., 1984; McCabe et al., 1985, 1986). Some back-arc basins have anomaly patterns characterized by weak or subtle trends including the Grenada, Sea of Japan, PareceVela, and West Fiji basins (Weissel, 1981; Brooks et al., 1984; Speed and Westbrook, 1984; Tamaki, 1985; Bird et al., 1993). Finally, some back-arc basins have no predominant magnetic pattern over them such as the Kurile, Okinawa, and Mariana basins (Lee et al., 1980; Weissel, 1981; Brooks et al., 1984; Okuma et al., 1990). Magnetic anomaly patterns produced by back-arc spreading may be related to the orientations of back-arc extension, which in turn may be related to the interactions between overtiding plates and their respective subducting slabs. Four scenarios that lead to a lack of coherent magnetic lineations over back-arc basins are: (1) complex rifting and segmentation of spreading centers producing incoherent anomaly trends (Tamaki, 1985); (2) young basins, such as the Mariana and Okinawa, not sufficiently developed to produce well defined trends; (3) basins formed by e a s t - w e s t extension near the magnetic equator (the Grenada basin is a probable candidate for this scenario); and (4) a basin formed during a time when there were no geomagnetic reversals, such as in Mid-Cretaceous time (about 118-84 Ma). Refraction data from the Andaman, Banda, Celebes, Grenada, Havre, Mariana, Okinawa, Parece-Vela, Sea of Japan, Shikoku, and Sulu basins reveal that the crustal structure of back-arc basins is similar to 'typical' oceanic crust described by Ludwig et al. (1971). The nature of back-arc basin crust, however, is more variable. Six back-arc basins, Grenada, Mariana, Okinawa, Parece-Vela, Sea of
397 Table 3 Means and standard deviations for the velocity structure of selected back-arc basins (regarding transition layers, means and standard deviations are only calculated for the total of all transition layers) Back-arc basin Layer2
Sea of Japan Okinawa Shikoku Parece-Vela Andaman Sulu Celebes Banda Grenada Mariana Havre Means
Layer3
M
STD M
STD
5.5 5.1 4.9 5.3 5.4 5.4 5.2 5.1 4.9 5.1 4.4 5.1
0.2 0.3 0.6 0.5 0.3 0.3 0.2 0.5 0.6 0.3
0.4 0.2 0.3 0.1 0.3 0.3 0.3 0.2 0.4 0.3
6.4 6.0 6.7 6.9 6.2 6.4 6.7 6.6 6.2 6.2 6.6 6.4
Transition Mantle
7.4, 7.5 7.1, 7.4 7.7 7.2 7.4, 7.4 7.2, 7.4 7.3, 0.2
M
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0.2 0.3 0.2 0.3 0.4 0.2
Japan, and Sulu basins (Ewing et al., 1957; Karig, 1971; Hayes et al., 1978; Bibee et al., 1980; Hussong and Uyeda, 1981; Curray et al., 1979), appear to contain an additional layer exhibiting intermediate velocities between layer 3 and mantle velocities (defined as the transitional layer for this study). Table 3 displays results from the eleven back-arc basins studied by analyses of refraction velocities. Means and standard deviations of the data are displayed for each basin as well as for the entire data set. The standard deviations illustrate the variability of the data, while the means support the concept of layered crusts, similar to normal oceanic.
REGIONAL SETTING
Geology The Grenada basin is bounded to the north by the Saba Bank at the junction of the Greater and Lesser Antilles, and to the south by the continental rise of northern Venezuela (Fig. 1). The Aves Swell (or Ridge) and the Lesser Antilles arc form the western and eastern limits of the basin, respectively. The shape of the basin is arcuate with approximate dimensions of 640 km (north-south) by 140 km (east-west), its water depth ranges from about 2 to 3 km. Sediment thickness ranges from 2 km in the north to 9 km in the south (Bouysse, 1988). Morphologically, the ocean floor of the Grenada basin can be divided into northern and southern parts at about 15.5~ The bathymetry of the northern part is rugged while the southern part of the basin is characterized by a near-horizontal, smooth seafloor. The nature of the sediments in the Grenada basin is not known; however, refraction data indicate that
398 sediments of the Aves Ridge extend and thicken into the basin (Westbrook, 1975). The Aves Swell, an extinct island arc (Bouysse, 1984, 1988), is oriented north-south and dips steeply into the Venezuela basin to the west. Its eastern edge is arcuate, similar to the Grenada basin, and descends in steps into the Grenada basin. Fox and Heezen (1975) report volcanic rocks consisting of andesites, basalts, dacites and volcanic breccias recovered from pedestals and scarps of the Aves Ridge. They also report Middle Eocene fossiliferous limestones, marls and chert recovered from dredge hauls. Late Cretaceous to Paleocene granodiorites, diabases and basalts, dredged from the southern part of the Aves Ridge, may be part of the Aves Ridge or the northern edge of the South American platform (Fox and Heezen, 1975). Like the eastern edge of the Aves Swell and the Grenada basin, the general shape of the Lesser Antilles island chain is arcuate. It bifurcates at about 15~ with a maximum separation of about 50 km at its northern limit. The outer arc is older (generally middle and late Paleogene to early Neogene) and inactive, while the inner arc is younger (generally Neogene to Quaternary) and presently active (Fox and Heezen, 1975). The western limit of the arc is marked by steep bathymetric gradients into the Grenada basin. Bouysse (1988) reports two episodes of volcanism: the first in Late Cretaceous time prior to back-arc spreading (84-66.4 Ma), and the second from Eocene to present. Warner (1991) reports that Early Paleocene to Early Oligocene volcanic rocks have been recovered from the Saba Bank. The episodic nature of volcanism is clear because volcanic rocks younger than Early Paleocene to Early Oligocene (Warner, 1991) have not been collected from the Aves Swell and volcanic rocks older than Middle Eocene (with one exception) have not been collected from the Lesser Antilles (Bouysse, 1988). This age difference indicates that the Grenada basin probably formed in early Tertiary time. The duration of back-arc spreading was restricted, as found in other back-arc basins. Jurassic age basalts on the island of la D6sirade are thought to represent obducted oceanic crust (Fink, 1968, 1970).
Seismic refraction The thickness of the crust increases beneath the Aves Ridge and Lesser Antilles and decreases beneath the Grenada basin (Kearey, 1974; Kearey et al., 1975; Westbrook, 1975). Modeling also suggests that the base of the crust generally mirrors the topography and bathymetry (Boynton et al., 1979; Bird, 1991). Oceanic crustal layers 2 (4.9-5.3 km/s) and 3 (6.2-6.4 km/s) are present in the Grenada basin south of about 15~ (Speed and Westbrook, 1984).
D.E. BIRD et al. Boynton et al. (1979) reports that the Grenada basin crust is approximately 14 km thick near 14~ with an average velocity of 4.8 km/s. Descriptions of the velocity structure vary; however, general structure and probable rock types follow: (1) 2.2 and 3.7 km/s m unconsolidated sediments and partially lithified sediments (Boynton et al., 1979); (2) 5.3 km/s m (upper oceanic crust), basaltic flows and dikes (observed only in the southern part of the basin (Speed and Westbrook, 1984); (3) 6.2 km/s (lower oceanic crust), gabbros; and (4) 7.4 to 7.5 km/s lower crust to mantle velocity transition zone. Upper and lower crusts of the Lesser Antilles arc are defined as 6.3 km/s and 6.9 km/s layers totaling about 35 km in thickness (Westbrook, 1975; Boynton et al., 1979). Rocks of the lower crust may be basic in composition while the rocks overlying the upper crust (3.4-4.5 km/s layer) may be composed of limestones, pyroclasts and sediments (Boynton et al., 1979). Officer et al. (1957) describe the seismic structure of the Aves Ridge as similar to that of the Lesser Antilles.
Seismic reflection Neither of the prominent Caribbean reflectors, A' or B', can be traced across the Aves Ridge into the Grenada basin; however "... a lower or middle Miocene horizon can be followed throughout the Grenada basin and west to the crest of the Aves Ridge..." (Speed and Westbrook, 1984). Two outer ridges of the Aves Swell enclose horizontal layers of sediment which are thickest near 13.4~ (Kearey, 1974). Reflection data also suggest that the Aves Ridge continues northward with its topographic expression buried by layers of sediment (Kearey, 1974). For the purpose of this work, the most important aspect of the seismic reflection data is the character of sediments overlying deep structures in the basin. The rugged bathymetry of the northern part of the basin is produced by deep structures in several locations. Conversely, reflection horizons of the southern part of the basin are smooth and relatively undisturbed (Fig. 7). Sediment thickness increases to the south from approximately 2 to 9 km.
Magnetics Magnetic anomalies over the Grenada basin (Fig. 9 and Web-Figs. 15.5,6) have been carefully examined by Bird et al. (1993). Anomaly amplitudes of hundreds of nanoteslas and wavelengths ranging from 10 to over 50 km are observed. Anomalies over the southern part of the basin display longer wavelengths and smaller amplitudes than those over
TECTONIC EVOLUTION OF THE GRENADA BASIN the northern part. Similarly, shapes and trends of anomalies change from north to south. The shape of the anomalies over the northern part are typically oblong with an east-west trend degrading to patchy and more disorganized farther south. The magnetic anomalies over the Aves Swell are similar to those over the northern part of the basin except that they are oriented north-south with larger amplitudes. The magnetic anomalies over the Lesser Antilles range in amplitude from 150 to 600 nT with wavelengths from 5 to 40 km. Short-wavelength anomalies (20 to 50 km) clearly delineate the island chain.
Gravity Several high-amplitude (approximately 80 to 150 mGal) free-air gravity anomalies, parallel to and just east of the island arc, are shown in Fig. 4 and WebFigs. 15.1,2. The wavelengths of these anomalies increase from about 20 km in the south to 50 km in the north. Similarly, several north-south-trending free-air gravity anomalies are observed over the Aves Ridge reflecting bathymetric variations. These anomalies display wavelengths and amplitudes of about 20 km and 50 mGal, respectively. North of the Aves Ridge, the free-air gravity field is subdued, displaying a broad positive northeast gradient (0.6 mGal/km) over the area. The average free-air gravity value ranges from about 0 to - 2 0 mGal over the northern half of the basin, then gradually decreases to about - 8 0 mGal over the southern half. Airy isostatic reduction was performed by Kearey (1974). Isostatic anomalies over the Aves Ridges are generally negative ( - 1 5 mGal) with some positive values to the south. Isostatic anomalies over the Grenada basin decrease from +10 to - 3 0 mGal, north to south. Kearey (1974) calculated 50 reGal positive anomalies over the Lesser Antilles.
INTERPRETATION The northern part of the Aves Ridge (north of 15~ appears to be displaced to the west and may be related to the late Tertiary tectonic event which resulted in the westward shift of the Lesser Antilles (Figs. 4 and 5 and Web-Figs. 15.1-4). Therefore, the data are inspected for features which support this hypothesis. Subtle, curvilinear 'discontinuities' are interpreted for total intensity magnetic anomaly, free-air gravity anomaly, and bathymetry data sets. Since these data sets are physically different and related to physically different rock properties, trends are not coincident between the data sets. Discontinuities consist of connected gradients, highs and lows, and/or connected truncations of gradients, highs and lows. Subduction of the aseismic Barracuda Ridge
403 (McCann and Sykes, 1984), or differential motion between the North American and South American plates (Bougault et al., 1988), in the late Tertiary has caused the subducting slab(s) to shoal under the northern part of the overriding Caribbean plate. McCann and Sykes (1984) have further suggested that the Barracuda and Main Ridges are continuous and have been recently overridden by the relative eastward motion of the Caribbean plate. In either case this shoaling has resulted in a westward shift of the center of volcanism beneath the northern part of the Lesser Antilles island arc. In addition to tectonic compression, this shoaling of the subducting slab(s) may have caused sections of the overriding plate (i.e., northern Aves Ridge, Grenada basin, and Lesser Antilles Arc) to be displaced and/or disrupted. Therefore curvilinear discontinuities are interpreted to be related to disruptions within the crust of the basin, which are in turn interpreted to be related to the tectonic event responsible for the bifurcation of the arc.
Gravity High-amplitude free-air gravity anomalies near the Lesser Antilles arc are interpreted to be produced by a combination of shallow bathymetry as well as dense volcanic rocks. Hence the contrast between the magmatic arc with the surrounding water and sediments results in positive gravity anomalies as seen in the results of 2-D modeling. The long-wavelength low over the southernmost part of the basin is also interpreted to be produced by a combination of effects. Sediment thickness increases to about 10 kin, causing the crust to warp down into the mantle. Kearey (1974) reports that negative isostatic anomalies (less than - 3 0 mGal) over the southern part of the Grenada basin may be related to downward flexing of the crust into the mantle. Bouguer gravity anomalies were calculated in order to compensate for the water bottom interface and allow for the interpretation of crustal variations. Bouguer gravity anomalies were calculated for two cases: (1) substituting rock with a density of 2.0 g/cm 3 for the water layer, and (2) substituting rock with a density of 2.67 g/cm 3 for the water layer (Fig. 10 and Web-Figs. 15.7,8). In both cases a broad high (i.e., positive anomaly) is located over the center of the basin, with a north-south elongation, from about 13~ to 15~ This high is slightly displaced to the south as the density used to calculate the Bouguer anomalies is increased. A large, triangular-shaped residual low is produced by the Bouguer calculations over the southernmost part of the Grenada basin and adds confidence to the notion that the crust is downwarped into the mantle. In general, the only differences between utilizing a rock
404
D.E. BIRD et al.
Fig. 9. Total intensity magnetic anomalies over the study area. Heavy solid and dashed lines indicate interpreted curvilinear zones of disruption (dashed lines indicate reduced confidence). Contour interval is 50 nT. density of 2.0 or 2.67 g/cm 3 in the Bouguer calculation are the amplitudes of the resultant anomalies. Anomaly shapes and wavelengths are essentially unaffected by the calculations. The large Bouguer gravity high over the center of the basin is interpreted to be caused by crustal thinning and the formation of the Grenada basin. This gravity high migrates southward as the rock density which replaces the water layer is increased, because as the water column increases, the increased density assigned to it produces increased anomaly amplitudes. Small wavelength anomalies over the Lesser Antilles and Aves Ridge are relatively unchanged by the calculation. If the Bouguer gravity
is compensating for the effect of the water bottom, then cross-cutting northeast-trending discontinuities also seen in these maps are interpreted to be related to deep-seated, crustal features
Two-dimensional models Two-dimensional models were constructed using profile gravity, bathymetry and seismic (both reflection and refraction) data. Basement horizons, interpreted from multiple channel seismic reflection lines, were converted to depth prior to being incorporated in models. Oceanic crustal layers 2 and 3, and a probable transition from lower crust to man-
TECTONIC EVOLUTION OF THE GRENADA BASIN Table 4 Densities used in 2-D and 3-D forward models Layer
Density (g/cm 3)
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1.03 2.30 2.57
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tle velocity, were incorporated utilizing refraction information. Densities were interpolated from the Nafe-Drake curve (Ludwig et al., 1971). In order to maintain consistency throughout the study area, density variations were not introduced within individual rock layers of the models. Table 4 displays horizons and densities used in modeling. Depth conversion of two-way seismic reflection time is achieved by utilizing a velocity function derived from velocity analyses of seismic line C1904 (LDGO). This velocity function is developed in a two-step process. First, the velocity analyses in the basin are converted to depth using the Dix formula (Dix, 1955). Time versus depth curves are then overlain and a 'best fit' curve is determined. Second, the interpreted velocity curve is modified to satisfy seismic refraction data. Reversed seismic refraction line 29 (Officer et al., 1957) indicates that the top of oceanic seismic layer 2 is at about 8 km below sea level near LDGO line 15, cruise RC1904. The lower limit of the velocity function was set at this point. The use of a single velocity function for basinwide time-to-depth conversions is easily implemented, but there are inherent pitfalls. The sediment and water layer thicknesses vary from 0 to 9 km and 0 to 3 km, respectively. The effect of an average velocity function tends to make shallow horizons too deep, and deep horizons too shallow. A high degree of accuracy with respect to shallow depth conversions is not critical for the purpose of this study. Fortunately the velocity analyses from cruise RC1904 are tied to seismic refraction velocities, hence the accuracy of depth conversions for deep horizons is good (:k: 0.5 km) when considering the overall dimensions of the basin (640 by 140 km). Connecting models (A-B, B-C, and C - D - E ) over the Grenada basin, oriented north-south and concentric with the Lesser Antilles Arc, reveals that the depth to the base of oceanic crustal layer 3 (Ludwig et al., 1971) decreases from about 20 to 18 km in C - D - E and remains nearly constant at about 15 km thereafter (Fig. 11). The combined thickness
405
of oceanic crustal layers 2 and 3 decreases from about 15 km in the south, to 8 km in the center of B-C (near 14.5~ and then increases to 15 km to the north. In general, the thickness of the transitional layer is modeled to thicken away from the central portion of the basin. East-west-oriented models (FF' and G-G') reveal that the minimum depth to the base of oceanic crustal layer 3 ranges from 13 to 15 km (Fig. 12). These minima are located at about 62.5~ and 61.5~ for models F-F' and G-G', respectively. The combined thickness of layers 2 and 3 ranges from about 17 km (Aves Ridge) to 7 km (Grenada basin) to 28 km (Lesser Antilles arc). For the two-dimensional model F-F', the depth to the base of layer 3 decreases to about 15 km at 15.75~ 62.5~ and increases to about 17 km to the east (Fig. 12A). The minimum combined thickness of combined layers 2 and 3 is 11 km. Dramatic variations in the combined thickness of layers 2 and 3 modeled in the eastern part of F-F' are interpreted to be related to the discontinuities. Other dramatic variations in the combined thickness of layers 2 and 3 are observed in the western part of F-F'. These variations are interpreted to represent density variations within the Aves Ridge. In model G-G', the shape of the base of layer 3 is asymmetric and shallowest (about 14 km) in the eastern part of the model (14.25~ 61.5~ From here the depth increases east and west to about 18 km (Fig. 12B). Overall, the two-dimensional models display good correlations between calculated gravity and free-air gravity profiles. Long-wavelength gravity anomalies over this area are considered to be two-dimensional because of the overall north-south trough-like shape of the gravity field over the study area. The excellent correlation of gravity profiles adds confidence to the overall concept of the basin's crustal structure. That is, the crust of the Grenada basin thins toward the center and thickens under the Aves Ridge and Lesser Antilles island arc. The observed gravity data are consistent with such a model. Three-dimensional models
A three-dimensional forward model was constructed to help interpret gridded gravity data. Gridded bathymetry data and an interpreted depth to acoustic basement surface (Speed and Westbrook, 1984) were incorporated into a two-and-one-half layer model representing the water column, sediments, and basement respectively. The depth to acoustic basement surface was interpreted from twoway travel time using many of the multi-channel seismic lines utilized in this study (Speed and Westbrook, 1984). Fig. 13, and Web-Figs. 15.9,10, show the depth-to-acoustic basement surface used in 3-D
Fig. 10. Bouguer gravity anomalies over the study area calculated by substituting rock density of (A) 2.0 g/cm 3 and (B) 2.67 g / c m 3 for the water layer. Contour interval is 10 mGal.
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modeling. Speed and Westbrook (1984) interpreted this basement surface in time (Web-Fig. 15.11) utilizing extensive multiple- and single-channel seismic reflection data sets. This surface was converted to depth utilizing the velocity function described above. The calculated gravity field is produced by the water layer, sediment layer, and the upper surface of layer 2. The observed free-air gravity field is produced by these same layers as well as layer 3, the transitional layer, and deeper sources. The observed free-air gravity field also reflects density variations within the sediments and basement which are unknown, as well as variations in crustal thickness. In general, density variations within the sediments and shallow parts of the basement should produce
short-wavelength anomalies. With respect to longwavelength anomalies, the calculated gravity field (Fig. 14A and Web-Fig. 15.12) correlates well with the observed free-air gravity field; however, there are several important differences between the calculated and observed fields. The location of the gravity minima over the Grenada basin is displaced to the north for the calculated field. A north-south-trending, subtle, broad high is superimposed along the center of the basin for the observed free-air gravity. The area of the continental shelf of Venezuela exhibits an anomaly high in the calculated gravity. The observed free-air gravity field has high frequency anomalies, particularly over the Aves Ridge and Lesser Antilles, which are not observed in the calculated field.
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To increase confidence in the 3-D model, the free-air gravity field is filtered (high cut: 60 km) to remove short-wavelength anomalies, then the calculated gravity is subtracted from the filtered free-air
gravity (Web-Figs. 15.13,14). This operation isolates anomalies which are produced by deep crustal and upper mantle variations. Residual gravity anomalies over the basin are similar to Bouguer gravity anoma-
TECTONIC EVOLUTION OF THE GRENADA BASIN 64"W
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lies and are defined here as residual Bouguer anomalies. A broad high exists over the basin and lows over the Aves Ridge and Lesser Antilles arc. Also, the triangular-shaped anomaly over the southernmost part of the basin correlates well with Bouguer anomalies.
DISCUSSION A key to understanding the Grenada basin is an understanding of the differences and similarities between the northern and southern parts of the basin. The differences are dramatic for bathymetry, seismic reflection, and magnetic data. However, free-air and Bouguer gravity data suggest that the crustal
structure of the basin is broadly similar from north to south. Although seismic refraction coverage is sparse in the north, when utilized as control for twodimensional and three-dimensional modeling, these data also indicate similarities from north to south. The only mantle velocity (8.2 km/s) recorded for the basin is Lesser Antilles Seismic Project (LASP) Profile B (Boynton et al., 1979) and is oriented approximately east-west near 13~ (Fig. 15). Two-D model G-G' is crossed by the north-northeast-oriented reversed Profile 29 (Ewing et al., 1957). There is no doubt that the crust in this part of the basin is similar to typical oceanic layering. Average velocities of back-arc basins are 5.1, 6.4, 7.3, and 8.2 km/s for layer 2, layer 3, transition, and mantle,
Fig. 14. (A) Calculated gravity anomalies (free-air) over the study area. (B) Filtered free-air minus calculated gravity anomalies = residual Bouguer anomalies. Contour interval is 10 mGal.
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by a broadly arcuate pattern gently dipping northnortheast with a hinge-zone marked at 5 s (twt) that turns into a southeasterly direction along the offshore Guyana passive margin. The deepest and northernmost well to reach basement in the Matuffn Basin is Soledad-1 or p o i n t ' S ' in Fig. 12 and Fig. 13b. The well encountered an undated granite at a depth of 4267 m (Feo-Codecido et al., 1984). Note that the roughly east-west-trending basement strike to the north parallels the fracture zones in the adjacent At-
lantic Ocean suggesting that the eastern Venezuela passive margin may have originated as a transform margin as also suggested by George and Sams (1993).
P R E - C R E T A C E O U S S T R A T I G R A P H Y ( P R E - R I F T AND R I F T PHASE)
Onshore a few seismic profiles suggest the presence of pre-Cretaceous sediments. Occasionally in
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN the northern Maturfn Basin a strong regional reflector can be observed about 6 s below the base of the Cretaceous. Together with other profiles it can be shown that this reflector has an overall northwesterly dip. The top of the crystalline Precambrian on these sections has been designated as SB-1 and the base of the Cretaceous as SB-2 (Di Croce, 1995). Farther east, a few seismic sections both onshore and offshore show locally poorly defined diverging reflectors (e.g. western third of regional profile Fig. 9a). It is useful to differentiate the deep regional pre-Cretaceous reflector from other more local reflectors that obviously diverge. The more local diverging reflectors may correspond to possible Jurassic half-grabens and the deeper reflector may represent the base of the lower Paleozoic cover of the Guyana Craton. For convenience, the lower reflector was designated as SB-1 and the base of the more local diverging reflectors was designated as SB- 1.1. The relationship between SB- 1 and SB- 1.1 cannot be observed on any single profile and the proposed interpretation by Di Croce (1995) is rather vague. Pre-Cretaceous sediments have not been drilled in the study area but are known farther to the west. Much of the relevant information was summarized by Feo-Codecido et al. (1984) and was incorporated in Fig. 12. Accordingly a zone underlain by an undeformed wedge of lower Paleozoic sediments has been penetrated by a number of wells. According to Feo-Codecido et al. (1984) the Carrizal-2X well penetrated 1827 m of Carrizal clastics. We believe that the above-mentioned northwesterly dipping reflector may correspond to the base of the Cambrian Carrizal and Hato Viejo formations. Note that to the north of the Espino Graben, the Paleozoic is deformed and slightly metamorphosed and the Apure thrust fault separates the deformed Paleozoic deformed belt to the north from the Guyana Craton and its subsurface Paleozoic cover to the south. The Jurassic Espino Graben is filled with an unfossiliferous red-bed section with intercalated basalt flows (162 Ma, see Feo-Codecido et al., 1984). These flows are underlain by some Carboniferous clastics and the Cambrian Carrizal Formation. By analogy to sections farther west, the Jurassic is assigned to the 'La Quinta' Formation of the Lake Maracaibo area (Gonzalez de Juana et al., 1980). Although the evidence for extension shown on our seismic profiles is far from satisfactory, we speculate that a subdued Jurassic rifting event may affect much of the Eastern Venezuelan Basin. Large Jurassic rifts have been reported from offshore French Guyana (Gouyet et al., 1992). In offshore Guyana (southeast of the study area) none of the wells drilled penetrated this sequence. However, Upper Jurassic samples (angular pebble-sized fragments of light-green,
441
consolidated, sorted, medium- to coarse-grained calcarenaceous orthoquartzite which is composed of rounded quartz, shell debris, nonskeletal granules with minor amounts of glauconite) were dredged from a 4400 m deep scarp at the northern edge of the Demerara plateau (Fox et al., 1970; Gouyet et al., 1992). The possible presence of Jurassic rifts is further supported by observations from the Tacutu rift system of northeastern Brazil which parallels the Espino Graben. According to Eiras and Kinoshita (1989) the Tacutu Basin is bottomed by extensive Middle to Upper Jurassic basaltic volcanics (150 Ma to 180 Ma), overlain by about 1250 m of Jurassic clastics and evaporites and 4900 m of possibly Lower Cretaceous clastics. These observations are relevant because they support an opening of the Atlantic facing the Orinoco Delta during the Late Jurassic (Heezen and Freeman-Lynde, 1976). To sum up, we suggest that the basement reflector SB-1 at the western end of the area is probably overlain by a thick wedge of lower Paleozoic, mostly Cambrian sediments. These could possibly correlate with the Paleozoic sequences of the Bove Basin of southern Senegal (Villeneuve et al., 1989; Villeneuve and Komara, 1991) and its western extension in Florida which on many Pangea reconstructions were adjacent to the north of our study area (e.g. Dercourt et al., 1993). Farther east in the Eastern Venezuelan Basin, sporadic diverging reflectors suggest the occurrence of limited rifting which may be coeval with the rifting events of the Espino Graben to the west and the Tacutu Graben of northeastern Brazil. The sequence boundary SB-I.1 separates the underlying Precambrian craton from the overlying 'Jurassic' graben fill. We conclude that a rifting phase preceded the Upper Jurassic opening of the Venezuelan Atlantic margin. However, over most of the area the pre-Cretaceous basement is the Precambrian craton and therefore we refer to the amalgamated SB-1/SB-2 simply as the pre-Cretaceous basement top.
THE CRETACEOUS-PALEOGENE (PASSIVE MARGIN PHASE)
During the Cretaceous and Paleogene siliciclastic sequences were deposited along the passive margin of Venezuela responding to tectonic subsidence and worldwide eustatic sea-level changes. A seaward thickening wedge of sediments represents this tectono-stratigraphic megasequence. The chronostratigraphic calibration of the offshore seismic profiles is mostly based on offshore wells A and B (Fig. 14a). The proposed stratigraphic correlation with the outcrops of the Serranfa del Interior (Interior Range) is shown in Fig. 14b, which also
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SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN shows the relation of the formations known from the Serranfa del Interior to the offshore Orinoco platform. Calibration of the seismic profiles permits subdivision of the passive margin Cretaceous into three stratigraphic units, i.e. from oldest to youngest: Unit II (Cretaceous), Unit III (Paleocene-Eocene) and Unit IV (Oligocene) (Fig. 6b). The age attribution in Ma to the proposed sequence boundaries and maximum flooding surfaces is merely an inference because, in fact, the available paleontological control is not adequate to fully support the assignment of specific ages. In other words, the proposed subdivision is a working hypothesis that needs to be amplified by future paleontological studies from both old and new wells. The description of Unit II (Cretaceous) and its subdivision will be discussed separately for the offshore and the onshore areas, respectively. Unit III and Unit IV are mostly present in the offshore area because in the onshore area (i.e. south of the deformation front) a hiatus corresponds to much of the Paleogene. This hiatus may be due to a combination of nondeposition and erosion preceding the deposition of the onlapping Early Miocene foredeep.
Unit II (Cretaceous) offshore On reflection profiles it is observed that the offshore stratigraphic Unit II (Cretaceous) onlaps on the basement and locally on some Jurassic halfgrabens corresponding to sequence boundary SB-2 (i.e. the breakup unconformity of some authors) which merges with sequence boundary SB-1 to form a major regional stratigraphic break, i.e. the pre-Cretaceous unconformity. At its top Unit II is bounded by sequence boundary SB-2.1 (58.5 Ma or the basal Late Paleocene). Unit II is an overall seawardthickening wedge (Fig. 7b) trending north-northeast where its thickness in the best preserved portion reaches almost 3.0 s (about 12,500 m). A time-structure map of the top of the Cretaceous (Fig. 15a) is characterized by a monotonous seaward-dipping attitude with widely separated contours to the south (0.0 to 4.5 s), becoming closer-spaced to the north-northeast (5.0 to 7.0 s), reflecting the gradient change from platform to slope. In fact, the 5-s contour approximates the seismic shelf edge of the Cretaceous platform (Fig. 7a). Based on internal stratal configurations a tentative subdivision into five distinct packages is made. These packages amount to transgressive-regressive (T/R) cycles bounded by four major continuous high-amplitude reflectors, labeled K1, K2, K3, K4, and SB-2.1 (Fig. 16). The K1-K4 reflectors are interpreted as downlap surfaces that represent maximum flooding events corresponding to the lower Aptian (111 Ma), the uppermost Albian (98.5 Ma),
443
middle Cenomanian (95.75 Ma) and middle Turonian (91.5 Ma). From bottom to top the characteristic of each subunit is as follows. (1) The basal subunit II-A (inferred to be 'prelower Aptian') is a wedge-shaped section that pinches out in an updip direction, due to the progressive onlap on the crystalline basement (the SB-1/SB-2 or pre-Cretaceous unconformity) of divergent, discontinuous and variable-amplitude seismic reflectors (Fig. 16a). (2) The overlying subunit II-B (inferred to be 'Aptian-upper Albian') exhibits an aggradational stacking pattern with a nearly uniform thickness. The high-amplitude seismic reflectors are subparallel, continuous and can be traced throughout the area (Fig. 16a). (3) The following subunit II-C (inferred to be 'upper Albian to lower Cenomanian') displays an aggradational stacking pattern and pinches out in a downdip direction (Fig. 16a). The seismic reflectors are subparallel but discontinuous and have low to medium amplitudes. (4) The subunit II-D (inferred to be 'lower Cenomanian to middle Turonian') is a thin package corresponding to two or three reflectors (Fig. 16b) with suggestions of a downlap surface at its base. (5) The subunit II-E (inferred to be 'middle Turonian-Senonian') is characterized by a thin prograding wedge of sediments with a pronounced sigmoidal stratal pattern that downlaps on the K4 surface and thins basinward. The low- to mediumamplitude reflectors are fairly continuous but locally chaotic (Fig. 16b). Fig. 17a shows the relation of well 'A' to our proposed seismic stratigraphic subdivision of the offshore Cretaceous Unit II, which includes the Lower and Upper Cretaceous. Two of the abovementioned seismic-stratigraphic packages, i.e. 'prelower Aptian' and the 'Aptian to upper Albian', were penetrated by well A where the 'Lower Cretaceous' consists of a thick unit of multicolored and mottled shales with sandstones and silty shale. The biostratigraphic characteristics of this unit are mainly represented by palynomorphs. The occurrence of species such as Cellassopollis, Ephedrapites, Liliacidites, Cicatricosisporites, among others, suggests that the Aptian to Albian was penetrated (Furrer, 1979). Palynological and sedimentological analyses indicate that these sediments were deposited in a continental environment. The Upper Cretaceous subunits II-D and II-E inferred to be 'Cenomanian' and 'Turonian-Senonian' consist of thinly bedded light-colored algal limestone, rich in forams and pelecypods, interbedded with micritic limestone, glauconitic shale and sandstone. The study of its fossils (Furrer, 1979) has given the following ages: (1) Cenomanian-Turonian,
444
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Fig. 15. Structure map and isopachs of the Eastern Venezuelan Basin. (a) Time-structure map of the top of Cretaceous. (b) Time-isopach map of the Lower Miocene (Unit V). Note the isopach maximum to the west (1.25 s). (c) Time-isopach map of the Pliocene (Unit VIII). Compared with B the depocenter migrated eastward with the highest isopach values (5.0 s) located to the southwest of Trinidad.
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN
445
Fig. 16. Seismic expression of Cretaceous sequences (see Fig. 3c for location). (a) Uninterpreted and interpreted dip-oriented segment of a seismic profile showing the subdivision of Unit II (Cretaceous). (b) Uninterpreted and interpreted dip-oriented segment of a seismic profile (see Fig. 3c for location) showing details of the forestepping prograding pattern of the Upper Cretaceous Unit II-E. The downlap surface (K4) marks the inception of the overall regression from the Upper Cretaceous to the Tertiary. characterized by the presence of lnoceramus, Nerinea sp., miliolids and Acicularia; (2) Campanian, characterized by the presence of Orthokarstenia bramletti cretacea, Orthokarstenia cretacea and Orthokarstenia parva, Gaudryina palmula, and Globotruncana fornicata; (3) Maastrichtian, characterized by the presence of a planktonic assemblage
of Globotruncana, Pseudotextularia, Heteroelix and Rugoglobigerina. Based on sedimentological and paleontological analyses, the thin carbonate-shale facies of the Cenomanian, the Turonian, and the Senonian is interpreted to be deposited on a stable shallow-water platform with sequence boundaries which based on the limited paleodata are
446
J. DI CROCE et al.
Fig. 17. Regional Cretaceous and Paleogene correlation from Guyana to the Eastern Venezuelan Basin. (a) Passive margin stratigraphy based on correlation of selected key wells across the Guyana Basin, the Eastern Venezuelan Basin and outcrop sections of the Serranfa del Interior. Stratigraphic columns simplified after Petroconsultants (1989), unpublished Lagoven reports and Chevalier (1993). (b) Onshore Eastern Venezuelan Basin: stratigraphic correlation of Unit II (Cretaceous) penetrated by selected key wells.
compatible with the Haq et al. (1987) sea-level curve. Wells A and B penetrated a dominantly siliciclastic Cretaceous that was derived from the Guyana Shield. Only a few carbonate layers were encountered. It is not demonstrated that the eroded steep
seismic shelf margin observed on the northeastern segment of Fig. 7a consists of carbonates as suggested by the overall similarity of the seismic image with other eroded carbonate margins from the Blake Plateau (Dillon et al., 1985, 1988) or west Florida (Wu et al., 1990).
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN Unit II (Cretaceous) onshore
Onshore Unit II consists of a northward-thickening wedge which rests unconformably and onlaps the Precambrian crystalline basement (Fig. 8a). At its base, Unit II is underlain by SB-1/SB-2 pre-Cretaceous unconformity. At its top Unit II is bounded by SB-3 which corresponds to a major regional unconformity, which has already been mentioned as the 'basal foredeep unconformity'. Unit II also includes a couple of strong, continuous and parallel reflectors. These reflectors are probably due to an acoustic impedance contrast of clastics and carbonates. In general, Unit II is uniformly thick along strike but onlaps the basement on dip profiles. The sequence is truncated at the top. Based on well information Unit II is equivalent to the Aptian-Santonian Temblador Group (well 'H' in Fig. 17b). Unit II reaches 600 m to the north. The same well also permits subdivision of this unit into two subunits (Fig. 17b). The lower subunit correlates with the middle Aptian to Albian Canoa Formation which consists of mottled coarsegrained sandstones interbedded with siltstones that were deposited in a continental environment. The upper subunit correlates with the Cenomanian to Campanian Tigre Formation. This upper subunit is composed of two lithofacies. A basal lithofacies consists of sandstones interbedded with shales. The upper lithofacies consists of dolomitic limestones and glauconitic shales with Exogira, Lingula, Plicatula sp., Astarte sp. (Gonzalez de Juana et al., 1980). Sedimentological and biostratigraphic analyses indicate that this subunit was deposited in an environment ranging from lagoonal/near shore/marginal marine at its base to the outer-shelf towards the top. Unit II: onshore-offshore correlation
A correlation between the onshore and the offshore Cretaceous of the Eastern Venezuelan Basin with the stratigraphy of the Serranfa del Interior and the wells of the Guyana offshore is illustrated in Fig. 17. Based on seismic correlation and age, onshore Unit II appears to be equivalent to most of the offshore Unit II. Onshore, the top of Unit II with its omission of documented Maastrichtian may be more eroded. Toward the southeast, the continuation of the study area is the Guyana Basin, a portion of the passive margin basins of northeast and east South America (Jankowsky and Schlapak, 1983; Veeken, 1983). The Guyana offshore basin shows stratigraphic characteristics similar to the Orinoco platform. Thus, the Lower Cretaceous unit consists of mostly AptianAlbian shallow-water clastics and carbonates. However, in the DSDP Site 144 on the northern edge of the Demerara Plateau (east of the Guyana Basin) a
447
silty shale of Barremian to Aptian age was recovered by piston core and the oldest sedimentary unit so far (Upper Jurassic) (Fox et al., 1970) was dredged near DSDP Site 144. The Upper Cretaceous of the Guyana Basin consists of Cenomanian deep-water shales and marls and of Turonian-Maastrichtian platform-type clastics with carbonates at the periphery of the basin and deep water clastics with minor carbonates beyond that margin (Petroconsultants, 1989). By far the most complete Cretaceous section of the former passive margin is outcropping in the Serranfa del Interior (e.g. Rossi et al., 1985; Vivas, 1986; Chevalier, 1993; Erikson and Pindell, 1993), where the lowermost subdivision of the Barranquin Formation and its four cycles cannot be easily correlated with the subsurface data of the Eastern Venezuelan Basin. However, a comparison of the outcrop section with wells A and B from the offshore (Fig. 14b) supports the plausibility of the inferred assignment of the offshore downlap surfaces, i.e. K1 (111 Ma) is equivalent to the Aptian Garcfa Member, K2 (98.5 Ma) to the upper Albian Chimana Formation, K3 and K4 (95.75 Ma and 91.5 Ma, respectively) to the Cenomanian-Turonian Querecual Formation. In conclusion, the Cretaceous is onlapping on a crystalline basement over most of the area and only locally overlies the inferred Jurassic rift Unit I. The Cretaceous Unit II is best subdivided into the following five transgressive-regressive (T-S) cycles (Figs. 14a and 17a). (1) A wedge-like clastic pre-lower Aptian II-A cycle (SB-2 to K1, i.e. from ?132 to 111 Ma) corresponds roughly to the siliciclastic upper Barranquin section of the Serranfa del Interior. (2) A clastic inferred 'Aptian-upper Albian' II-B cycle (K1 to K2, i.e. 111-98.25 Ma) with a basal condensed downlap sequence may be equivalent to the Garcfa Member of the Serranfa del Interior. In the Serranfa these clastics are replaced by the carbonates of the E1 Cantil Formation. (3) A clastic inferred 'upper Albian to lower Cenomanian' II-C cycle (K2 to K3, i.e. 98.5-95.75 Ma) with a basal condensed downlap sequence is equivalent to the Chimana Formation of the Serranfa del Interior. For a similar correlation see also Galea-Alvarez et al. (1996). (4) A thin carbonate shale 'lower Cenomanian to middle Turonian' II-D cycle (K3-K4, i.e. 95.7591.5 Ma) with an ill-defined downlap surface at its base is equivalent to the lower Querecual Formation of the Serranfa del Interior. The Turonian marine flooding surface is also documented paleontologically in Trinidad by Huang (1996) who correlates the surface with a worldwide Turonian marine flooding event.
448 (5) A thin wedge of prograding clastics and carbonates of the 'middle Turonian-Senonian' II-E cycle (K4-SB-2.1, i.e. 91.5 to ?66.5 Ma) with a pronounced downlap surface at its base is probably equivalent to the combined upper Querecual, San Antonio and San Juan formations of the Serranfa del Interior. Significant parts of the II-E cycle may be eroded in much of the onshore of the Eastern Venezuelan Basin. Note that much of the Cretaceous in the Maturfn sub-basin and of the Orinoco shelf is dominated by shallow-water deposits, while the corresponding deeper-water lowstand sediments appear on outcrops and in wells in the Southern Basin of Trinidad (Sprague et al., 1996).
Unit III (Late Paleocene-Eocene) offshore and onshore Unit III is a thin and condensed section defined by two strong reflectors with onlap on the lower sequence boundary SB-2.1 (58.5 Ma) (Fig. 18a). The top is the SB-2.2 downlap surface. Offshore, the thickness of this sequence is nearly constant (i.e. less than 100 ms) and can be followed throughout the offshore area. On the other hand, onshore Unit III is mostly eroded or never was deposited. However, only in onshore w e l l ' S ' over 450 m of Middle and Upper Eocene pelagic shales and sandstones have been reported. Unit III can be subdivided into two depositional subunits. The data from offshore wells A and B (Fig. 14a) show that a lower subunit III-A (90 m thick) consists of two distinctive lithofacies. A basal lithofacies is characterized by shallow-water limestone with algae, ostracods and abundant planktonic foraminifera. Based on the occurrence of Globorotalia pseudomenardii, this basal lithofacies appears to be Late Paleocene (early Thanetian) in age. The upper lithofacies consists mostly of silty shale and dark shale with abundant benthic and planktonic foraminifera. The assemblage includes Bathysiphon, Glomospira,
Cyclammina, Globorotalia aegua, Globorotalia angulata, Globorotalia acuta and Globorotalia velascoensis. These paleontological data (Furrer, 1979) suggest deep (bathyal) water conditions for this lithofacies which is dated as Upper Paleocene (Thanetian). There is a pronounced hiatus separating Unit III from the underlying Unit II. This hiatus omits the Lower Paleocene and locally the Maastrichtian. Thus, the base of Unit III could be considered to be equivalent to the 58.5 Ma sequence boundary of Haq et al. (1987) and the base of the upper facies could correspond to the 56.5 Ma maximum flooding event of these authors. The upper subunit III-B (115 m thick) consists of glauconitic shale interbedded with thin-beds of siltyshale and fine-grained sandstone, that grades upward
J. DI CROCE et al. to reworked glauconitic limestone. The faunas again (Furrer, 1979) permit subdivision of this subunit into a Lower Eocene unit, based on the occurrence of
Globorotalia formosa, Globorotalia broennimanni and Globorotalia palmerae, a Middle Eocene unit, based on the occurrence of Globorotalia spinuloinflata, Globorotalia centralis and Truncatulinoides rohri, and Upper Eocene unit (only in the well B) based on the occurrence of Globorotalia cerroazulensis cerroazulensis (Furrer, 1979). The same author interpreted an Oligocene unit in well A resting on top of the Middle Eocene. However, the very poor state of preservation of the planktonic foraminifera did not permit a clear definition of the age of this interval. Sedimentological analysis and the occurrence of benthic foraminifera, such as Cyclammina, suggest a bathyal (500 m) depositional environment for these sediments. The presence of reworked limestone that includes shallow-water forms such as Lepidocyclina, Nummulites and Discocyclina, suggests that these sediments were transported and reworked in a deep marine setting, possibly by gravity flows. Note that onshore w e l l ' S ' reports Middle to Upper Eocene pelagic faunas from a sandstone-shale sequence. This suggests that subunit III-B extends well into the onshore area. Fig. 17a illustrates the regional correlation of the Paleocene and the Eocene. To sum up, in the offshore the Late Paleocene overlies the SB-2.1. Much of the Paleocene, the Danian and in places the Maastrichtian appears to be missing. A similar unconformity in the Guyana offshore is less well constrained (Petroconsultants, 1989). The lower Unit III-A corresponds to a middle Thanetian sequence, presumably sandwiched between the 58.5 Ma and the 56.5 Ma or 54.2 Ma sequence boundary of Haq et al. (1987). The upper subunit III-B includes most of the Eocene, with pelagic faunas of the Upper, Middle and Lower Eocene. Specific correlations with the Guyana offshore are difficult to make, but it appears reasonable that all units reported from the Venezuelan offshore area are also present in the Guyana offshore. The offshore 'Upper Paleocene' subunit III-A should be correlated with the top of the Vidofio Formation of the Serranfa del Interior, because this formation shows a very similar paleontological assemblage and lithological features. Subunit III-B appears to be equivalent to the Caratas Formation of the Serranfa, but the lithofacies of these units appears to be somewhat different. Note also that in the Serranfa (stratigraphic column J, Fig. 14b) a significant unconformity causes the omission of much of the Upper Eocene which is reported in well A (Fig. 14a). Subunit III-B is also represented in the Eastern Venezuelan Basin and has been penetrated by well 'S'.
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN
449
Fig. 18. Seismic expression of Paleogene sequences (see Fig. 3c for location of this section which is very close to well A shown in Fig. 14). (a) Uninterpreted and interpreted dip-oriented segment of seismic profile showing Unit III (Paleocene-Eocene). (b) Uninterpreted and interpreted dip-oriented segment of seismic profile showing Unit IV (Oligocene).
Unit IV (Oligocene) offshore and onshore Offshore Unit IV is a sedimentary wedge downlapping onto sequence boundary SB-2.2 (36 Ma of Haq et al., 1987). The top of Unit IV is the sequence boundary SB-3 (probably 25.5 Ma corresponding to the lowermost Miocene) which is characterized by truncated reflectors and overlain by onlapping reflectors (Figs. 6a and 18b). This upper boundary of
sequence IV is correlated with a regional unconformity described as the 'basal foredeep unconformity' that separates the overlying foredeep tectono-stratigraphic unit from the underlying passive margin unit. In the offshore area Unit IV occurs mostly to the south and no Oligocene has been reported in other wells of the offshore. Therefore, it is suggested that in the downdip offshore the section is eroded or never was deposited. In the onshore area except for well
450 S no Oligocene fossils have been reported from the Maturfn Basin, but significant siliciclastic sections are reported from the oil fields of the folded belt. Offshore Unit IV can be subdivided into two subunits separated by a high-amplitude reflector, represented by sequence boundary SB-2.3 (inferred to be 30 Ma by Haq et al., 1987). The lower subunit IV-A consists of subtle sigmoid seismic reflection patterns that pinch out in a downdip direction. The upper subunit IV-B displays a progradational configuration and shows a pronounced downlap geometry (Fig. 18b). Based on well 'A' (Fig. 14a) Unit IV (max. 200 m) consists of two lithofacies. A basal lithofacies is composed of glauconitic shale with abundant fossils (gastropods and foraminifers) that grade up into skeletal and glauconitic limestone (stromatoporoids and foraminifers) interbedded with fine- to medium-grained calcareous white sandstone. This lower lithofacies is considered to be Early Oligocene (Rupelian) in age based on the occurrence of Bulimina sculptiles and poorly preserved specimens of Globigerina euapertura (Furrer, 1979). The upper lithofacies consists of a coarsening-upward sequence of poorly consolidated medium- to coarse-grained white sandstone and appears to be Late Oligocene (Chattian) in age, based on the occurrence of Siphogenerina senni, Cassigerinella chipolensis and possible specimens of Globigerina ciperoensis (Furrer, 1979). In addition, the sandiest portion contained the macrofossils of Heterostegina antillea, Lepidocyclina and Nummulites. These sedimentological and paleontological analyses indicate that both units were deposited in shallow-water conditions closely associated with a carbonate platform. Based on correlation of the seismic facies and sedimentological characteristics, Unit IV is subdivided into: (1) a lower subunit IV-A (equivalent to Lower Oligocene) that consists of very condensed transgressive and highstand deposits which contain the maximum flooding surface 35 Ma and overlie sequence boundary SB-2.2 (36 Ma); these two surfaces merge updip in a single seismic reflector; (2) an upper subunit IV-B (equivalent to the Upper Oligocene) bounded at its base by sequence boundary SB-2.3 (30 Ma). Seismic data show that a lowstand deposit (mostly sandstone) associated with this sequence boundary pinches out updip near the relict offlap break of the previous highstand deposit and becomes thinner basinward (Fig. 18b). The only onshore record of the Oligocene of the Eastern Venezuelan Basin is reported from well 'S', where some 325 m of sandstone and pelagic shales are reported to contain Oligocene faunas including Globigerina
ciperoensis. Fig. 14b shows the correlation of Oligocene Unit IV between the composite stratigraphic column of
J. DI CROCE et al. the Serranfa del Interior with the Venezuela offshore. Fig. 17a ties the Venezuelan onshore-offshore to the Guyana Basin. It is important to note that in Fig. 14a the faunally documented Oligocene in well A is absent in well B, suggesting a significant hiatus, as already suggested by Prieto (1987). An alternative is that the Oligocene downdip of the progradational wedge shown on the seismic (Fig. 18b) is so condensed that its detection was not possible in well B. Onshore well 'S', which did encounter the Oligocene, places the zero edge of the formation to the south of the well. At this time it is not known whether the Oligocene has been encountered in other wells that were recently drilled farther north in the basin. The twofold division observed in offshore well A and on adjacent seismic lines can be reasonably correlated with wells in the E1 Furrial-Carito trend and from there with a section in the Serranfa del Interior (Fig. 17a). Many authors have accepted a threefold subdivision in the Serranfa, i.e. from bottom to top Los Jabillos clastics, the Areo shales and the Naricual clastics (e.g. Gonzalez de Juana et al., 1980; Sams, 1995). Conventionally, these three formations were all included in the Merecure Group (e.g. Gonzalez de Juana et al., 1980), which has its type locality in the Santa Ana field of the Anaco trend. Because no fossils are reported from the Merecure type section, the overall correlation with the sections of the Serranfa may be debatable. In conclusion, it appears that the Oligocene faunas reported from wells A and S, the faunas reported from E1 Furrial area, and the faunas reported from the Areo Formation are all roughly correlatable. Additional paleontological details from the Orocual field were reported by Giffuni and Castro-Mora (1996). Based on offshore data, the Oligocene second-order cycle may be split into two third-order cycles bounded respectively by SB-2.2 (36 Ma), SB-2.3 (30 Ma) and SB-3 (25.5 Ma). Throughout the area in the updip direction and toward the craton, the Oligocene is absent due to Lower Miocene to Upper Miocene erosion. Downdip of the progradational wedge shown near well A, the Oligocene may be either absent because of sediment starvation or so thin and condensed that the interval escaped detection in well B. The paleogeography of the Oligocene will be discussed below.
THE NEOGENE (FOREDEEP PHASE)
From uppermost Oligocene to Early Miocene a major change of tectonic subsidence regimes occurred in the Eastern Venezuelan Basin. In the context of the oblique convergence between the Caribbean and South American plates (see Fig. 4)
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SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN the east-west-trending Cretaceous-Paleogene passive margin of South America collided with the transpressional front of the Caribbean Plate. Thus, the northern Venezuelan transpressional folded belts became associated with a foredeep and its depocenters that migrated from west to east. Lugo (1991) illustrated the Paleogene foredeep of the Lake Maracaibo area and Audemard (1991), Lugo and Mann (1995) and Audemard and Lugo (1996) sketched how this foredeep migrated farther east into the area of this study. Three directions of sediment transport characterize the foredeep basin fill as follows: (1) an important longitudinal east-west-directed transport; (2) a southerly sediment source from the adjacent Guyana Shield; and (3) a north-northwesterly source of sediments from the emergent fold belt of the Serranfa del Interior which reworked the Cretaceouslower Tertiary units that outcrop to the north. Note that a significant amount of Plio-Pleistocene sediments derived from the Serranfa is also trapped in compressional satellite ('piggyback') basins that overlie the Miocene accretionary wedge of the Monagas foothills and do not reach the foreland basin itself (e.g. Parnaud et al., 1995).
Well calibration and key chronostratigraphic seismic horizons of foredeep stratigraphy As will be shown later the estimated age attributions for the Neogene stages differ and there is no agreement among various authors on the chronostratigraphic timing of stratigraphic stages and series. For the purpose of this paper the assigned ages in Ma of the chronostratigraphic horizons on figures and seismic profiles have been made in accordance with the global cycle chart of Haq et al. (1987). Because changes in the time scale are anticipated, sequence boundaries are numbered and the inferred age is given in brackets. Fig. 19a is a simplified summary showing key onshore chronostratigraphic horizons interpreted on well logs and seismic lines. It also shows major facies cycle subdivisions of the stratigraphic column. The foredeep stratigraphy is characterized by two overall transgressiveregressive cycle wedges, subdivided into four major units: the lower Unit V, corresponding to the Lower Miocene; Unit VI, Middle Miocene; Unit VII, Upper Miocene; and Unit VIII, Plio-Pleistocene to the present. In the following sections a more detailed description of these units is given, including comments about their overall geometry, seismic facies, stratigraphy and depositional environment.
Unit V (Lower Miocene) Seismostratigraphic Unit V is bounded at its base by sequence boundary SB-3 (i.e. the basal foredeep
455
unconformity) and at its top by sequence boundary SB-3.3 (16.5 Ma). The basal Lower Miocene (Aquitanian) sequence boundary SB-3 is a composite surface characterized on the regional scale by three main features: (1) onlap that recorded the deepening of the basin at the beginning of the foredeep phase (Fig. 14a and Fig. 20a); (2) on the distal stable platform a set of prograding sequences; and (3) on the onshore proximal stable platform an erosional surface defined by truncation of Cretaceous strata associated with the southward-migrating peripheral bulge (Fig. 20b). Fig. 15b shows an isopach in seismic twt-time of the Lower Miocene-Middle Miocene, which is characterized by its overall wedge-shaped geometry with values ranging from less than 0.5 s to over 2.5 s. On the longitudinal profile (Fig. 9a), which obliquely crosses the axis of the foredeep basin, Unit V (i.e. the SB-3-SB-3.3 interval) is relatively thick (3000 m) to the west (near the Anaco trend) and gradually thins towards the SSE. On NW-SE-oriented profiles (Fig. 8) Unit V pinches out towards the south as it onlaps the truncated Cretaceous (Unit II). The seismic and sedimentological characteristics differ between the onshore and offshore areas (see Fig. 20a). Therefore in the following sections the onshore and offshore areas will be described separately.
Unit V onshore On west-east-oriented seismic profiles (see Fig. 21a) the stacking pattern of key wells permits the subdivision of Unit V into at least three backstepping depositional sequences (i.e. third-order sequences) which are bounded by the sequence boundaries SB-3 (lowermost Miocene, approx. 25.5 Ma), SB-3.1 (21 Ma), SB-3.2 (17.5 Ma) and SB-3.3 (16.5 Ma). These boundaries are characterized by local truncation of the underlying reflectors and by onlapping reflectors. Internally, these depositional sequences are characterized by highstand and transgressive system tracts. However, the transgressive system tracts and maximum flooding surfaces are poorly developed. The following stratigraphic summary of Unit V is derived from a sedimentological report of well L (see Fig. 3a) shown in a generalized manner in Fig. 19a and located in the axis of the western portion of the basin. From bottom to top Unit V consists of two major lithofacies: (1) lithofacies A, overlying, but not reaching the SB-3 (i.e. basal foredeep unconformity), a section of blocky (15-25 m thick) and massive, medium- to coarse-grained sandstone interbedded with thin (2-4 m thick) layers of shale and occasional lignite which is bounded at its top by sequence boundary SB-3.1; this lithofacies was deposited in a fluvial environment; and (2) lithofacies B, bounded by SB-3.1,
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Fig. 20. Details of profiles showing the basal foredeep unconformity. (a) Uninterpreted and interpreted dip-oriented segment of seismic profile showing the configuration of the Neogene foredeep in the offshore area. For location see Fig. 3c. Note in the downdip direction the onlap of deep water facies on relatively thick Cretaceous passive margin sequence. Abbreviations: SB-3 = basal foredeep unconformity; SB-3.10 -- Upper Miocene (5.5 Ma) sequence boundary. See Fig. 3c for location. (b) Uninterpreted and interpreted dip-oriented segment of seismic profile showing the configuration of the Neogene foredeep in the onshore area. See Fig. 3c for location. Here a Neogene shallow-water facies onlaps SB-3 which overlies a thin truncated proximal passive margin Cretaceous sequence. SB-3.2 and SB-3.3, and consisting of two prograding packages of coarsening-upward sequences which are characterized by basal shales grading upward into alternating facies of siltstone and sandstone
and ending with fine- to m e d i u m - c o a r s e sandstone. These packages were deposited in a littoral to shallow marine environment as coastal bars prograding progressively landward.
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN Unit V offshore The stratigraphic configuration of Unit V in the offshore area differs considerably from the onshore but is still characterized by a northward-thickening wedge (Fig. 6a). In this part of the basin Unit V corresponds to the deep-water phase which illustrates the inception of the foredeep (Fig. 20a). Like in the onshore area, offshore Unit V is bounded at its base by SB-3 (i.e. the basal foredeep unconformity) with its deep-water Lower Miocene sediments onlapping onto thin Oligocene or probably older sediments. This interpretation is particularly important because a similar situation can be observed in wells of the E1 Furrial-Carito oil fields, i.e. within the folded belt to the north of the study area. There, the well data (Fig. 19b) show that a substantial deepening of the basin occurred during the Early Miocene as recorded by an important unconformity separating underlying Oligocene neritic sediments (lower Merecure Formation) from the overlying middle to upper bathyal shales and thin turbidites of the Lower Miocene (lower Carapita Formation). In the offshore to the north and northeast, Unit V and upper units have been deeply eroded by submarine currents during the Upper Miocene (Fig. 20a). Consequently, part of Unit V is directly overlain by sequence boundary SB-3.10 (5.5 Ma). Farther southeast still in the offshore but in the passive margin domain, Unit V exhibits a uniform thickness (approx. 110 ms) along strike profiles and thins updip and downdip (i.e. toward the south-southwest and toward the northnortheast). On seismic dip profiles Unit V can be subdivided into two depositional sequences which consist of a backstepping mound-shaped configuration, each of them bounded by onlap surfaces and underlain by truncated reflectors (Fig. 21b). Unit V is characterized by low- to medium-amplitude and moderately continuous reflectors which display bi-directional terminations (i.e. onlap updip and downlap basinward). Based on seismic and well correlations (Fig. 14a) in this portion of the basin, Unit V is composed of two lithofacies. A first lithofacies present to the south-southwest consists of gently dipping shallow-water carbonates that overlie Unit IV (Oligocene). On seismic this facies is not visible, and only a single high-amplitude reflector with moderate continuity is present. The carbonate section consists of white to beige skeletal coral micritic limestones that contain macroforaminifera such as Miogypsina, Lepidocyclina and, Nummulites, gastropods, echinoderms, bryozoa and algae (Halimeda and Lithothamniun). Sedimentological analysis and the faunas suggest the development of a coral reef complex during the Early Miocene (Furrer, 1979). The carbonate section shows cave and dissolution porosity suggesting subaerial exposure that perhaps has been related to a sea-level fall that occurred
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during the latest Early Miocene. This corresponds with SB-3.3 or the 16.5 Ma sequence boundary of Haq et al. (1987). Laterally and upward the second lithofacies of Unit V occurs basinward to the northeast and consists of mainly brown-olive silty shale. Sedimentological and paleontological analyses indicate that these sediments were deposited in an upper bathyal to outer shelf water depth and are considered to be Early Miocene (Burdigalian) based on the occurrence of Globigerinatella insueta. In light of the above description and patterns observed on seismic profiles it is suggested that the upper lithofacies of Unit V is represented by lowstand deposits characterized by muddy slope fans (Fig. 2 l b). The sequence boundaries that bound each of these lowstand deposits are from base to top as follows: SB-3 dated 25.5 Ma, SB-3.1 dated 21.0 Ma and SB-3.3 dated 16.5 Ma. Summary and comments Unit V has an overall transgressive character and is bounded at its base by SB-3 dated approximately 25.5 Ma (i.e. basal foredeep unconformity) and its top by SB-3.3 dated 16.5 Ma. Regional sequence stratigraphy allowed the subdivision of Unit V (Lower Miocene) as follows: in the onshore, Unit V consists of three depositional sequences, characterized mostly by highstand and transgressive deposits. This unit includes part of the Merecure Formation and the basal portion of the Oficina Formation. In the offshore Unit V consists of a thin carbonate platform located to the south and two depositional sequences characterized by lowstand deposits. Proceeding from west to east along the axis (i.e. from onshore to offshore) of the basin, the seismic reflection patterns of these depositional units change. In the western area, the seismic reflection patterns are represented by medium- to high-amplitude continuous reflectors and the sequences boundaries are defined by onlap and local truncation. The configuration changes laterally and basinward, to low-medium-amplitude discontinuous reflectors. The sequence boundaries are difficult to follow and some may merge into a single surface or seismic reflector. The seismic facies of the west and southwest are interpreted as coastal to platformal facies, gradually deepening to bathyal water depth to the north and northeast. This facies marks the inception of the foredeep phase in the basin. Unit VI (uppermost Lower Miocene-Middle Miocene)
Unit VI is bounded at its base by sequence boundary SB-3.3 (16.5 Ma) and at its top by sequence boundary SB-3.7 (10.5 Ma). On dip-oriented seismic profiles (Fig. 9a) Unit VI shows an overall
Fig. 21. Interpreted segments of seismic profiles illustrating Miocene sequences. (a) Interpreted segment of an onshore seismic profile showing depositional sequences and sequence boundaries of Unit V (Lower Miocene) and Unit VI (Upper Miocene) For location see Fig. 3c. (b) Interpreted segment of offshore seismic profile in the offshore area showing the configuration of Unit V (Middle Miocene). See Fig. 3c for location. (c) Interpreted segment of seismic profile in the offshore area showing the configuration of Unit VI (Middle Miocene) and Unit VII (Upper Miocene). See Fig. 3c for location.
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wedge shape, trending west-east, very similar in geometry to the underlying Unit V. In the western and best preserved portion of the basin the unit is over 2 s thick and thins toward the east. On N-S-oriented profiles (Fig. 8), Unit VI thickens toward the north and pinches out towards the south as it onlaps the truncated Cretaceous (Unit II) and the crystalline basement. Onshore, two different seismic facies characterize Unit VI. The western portion exhibits a pattern with high- to medium-amplitude reflectors with excellent continuity. Toward the east Unit VI is formed by reflectors with low- to medium-amplitude and disrupted/irregular reflection configurations. On the other hand, in the offshore the seismic facies exhibits a pattern with low amplitude and moderate continuity.
of siltstone and sandstone and ending with fineto medium-coarse sandstone. These sediments were deposited in a litoral to shallow marine environment and consist of coastal bars which prograded progressively landward. This section is bounded by SB-3.4 (15.5 Ma) and SB-3.5 (13.5 Ma). (3) Abruptly, the previously described lithofacies changes to a massive thick section (approx. 600 m) of silty shales with abundant planktonic and benthic faunas. The planktonic assemblage includes
Unit VI onshore Based on internal configuration and well data, Unit VI consists of four depositional sequences (third order) defined from bottom to top by sequence boundaries SB-3.3 (16.5 Ma), SB-3.4 (15.5 Ma), SB-3.5 (13.8 Ma), SB-3.6 (12.5 Ma) and SB-3.7 (10.5 Ma) (Fig. 21a). These sequence boundaries show onlap reflection terminations and SB-3.7 is a well-defined seismic horizon (particularly in the western portion of the study area) characterized by an erosive surface showing underlying truncated reflectors and overlying onlapping reflectors (Figs. 2 l a and 22a). Unit VI is characterized by two different trends. The lower two depositional sequences exhibit a backstepping behavior, building up from prograding higher-order sequences (Fig. 21a). The upper two depositional sequences are formed by subtle progradational forestepping but an overall aggradational configuration (Fig. 22b). Note also that some of these depositional sequences are composed of lowstand deposits. The ages of the sequence boundaries and of the maximum flooding are well constrained by paleontological data contained in the following stratigraphic summary of Unit VI which is derived from sedimentological reports of key wells and particularly of well L (Figs. 3a and 19a). From bottom to top the lithofacies that characterize Unit VI consist of the following. (1) Gray-brownish shale with occasional thin layers of sandstone and profuse pellets of glauconite throughout the interval yielded abundant and diverse faunas of an latest Early Miocene age. This age is mainly based on the occurrence of Globigerinoides sicanus which is restricted to Zones N7-N8 (Blow, 1969). The section was deposited in a middle to outer shelf environment. It includes the maximum flooding surface of 16 Ma of Haq et al. (1987) and is bounded by SB-3.3 (16.5 Ma) and SB-3.4 (15.5 Ma). (2) Coarsening-upward stacking patterns with basal shales grading upward into variable facies
indicating that these sediments were deposited in middle to upper bathyal water depths and mark the maximum deepening of the basin during the Middle Miocene (Serravallian). This section is bounded by SB-3.5 (13.5 Ma) and SB-3.6 (12.8 Ma) and contains the maximum flooding surface of 13.4 Ma of Haq et al. (1987) which is shown in Fig. 19a. (4) A thick (between 450 and 750 m) section of gray-olive shale, locally interbedded with a few thin layers of fine-grained sandstone. Benthic foraminifera such as Lenticulina, Bathysiphon, Haplophragmoides, Textularia, Miliammina suggest that this section was deposited in water depths ranging from upper bathyal/outer shelf to inner neritic. The planktonic foraminifera within this section are abundant and diversified. However, only one distinct marker was found, corresponding to the Globorotalia aft. siakensis, Zones N13-N14 (Blow, 1969), i.e. the upper part of the Middle Miocene. This zone is close to SB 3.7 which correlates with the sequence boundary of 10.5 Ma of Haq et al. (1987).
Globorotalia praemenardii, Globorotalia peripheroacuta, Globorotalia prefohsi, Globorotalia fohsi lobata, and Globorotalia obesa, which corresponds to Zones N10-N12 (Blow, 1969). The benthic faunas consist of Cyclammina cancellata, Bathysiphon,
Lenticulina subpapillosa, Valvulina flexilis, Bolivina imporcata, Globulina ovata, among many others,
Unit VI offshore The seismic stratigraphic configuration of Unit VI of the offshore passive margin domain differs from the onshore. Offshore Unit VI is a single depositional unit bounded at its base by SB-3.3 (16.5 Ma) and at its top by an onlap surface, corresponding to sequence boundary SB-3.7 (10.5 Ma). To the north in the offshore foredeep domain Unit VI is thin as a result of widespread deep erosion during the Early Pliocene. In the updip area Unit VI exhibits a very subtle sigmoidal pattern with moderate progradational offlap, characterized by low-amplitude and moderate continuous reflectors with northeast-dipping downlap terminations (Fig. 21c). Unit V (Lower Miocene) and Unit VI are separated by a thin continuous drape sheet (approx. 50 ms) which represents the downlap surface. Correlation with nearby wells suggests that this surface corresponds
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Fig. 22. Uninterpreted and interpreted segment of onshore seismic profiles showing sequence boundaries and relations of Middle Miocene to Pliocene. For location see Fig. 3c. (a) Onshore, the Unit VII (Upper Miocene) is characterized by three deeply incised unconformities (i.e. SB-3.7, SB-3.9 and SB-3.10). (b) This is another view of the relationship of Unit VI and Unit VII showing a more subdued onlap on SB-3.7.
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN to the 16 Ma maximum flooding surface (M1) of Haq et al. (1987). In offshore wells (Fig. 14a), Unit VI consists of thick (~ 180 m) mostly fossiliferous gray-olive shale with a few layers of sandstone and an increased glauconite content towards the lower part. Sedimentological and paleontological analyses indicate that these sediments were deposited in an outer-shelf setting. Based on the occurrence of planktonic foraminifers, such as (from bottom to top) Globorotalia fohsi fohsi, Globorotalia fohsi lobata and Globorotalia siakensis, which correspond to Zones N10 through N14 (Blow, 1969), the whole Middle Miocene is represented.
Summary To sum up, Unit VI together with the underlying Unit V appear to form a complete second-order transgressive-regressive cycle wedge in the onshore area (Fig. 19a) as well as in the offshore. The transgressive phase comprises the Lower Miocene to the middle portion of the Middle Miocene. During the transgressive phase two flooding events are interpreted. The older event is the 16-Ma maximum flooding surface of Haq et al. (1987) which can be correlated along the basin. The second flooding event, which culminates with a major paleobathymetric deepening, is interpreted to mark the peak transgression and is recorded as the 13.4-Ma (Serravallian) maximum flooding surface of Haq et al. (1987). Following that maximum transgression the overall regression began. The regressive phase comprises the upper portion of the Middle Miocene and SB-3.7 (10.5 Ma) is interpreted as the peak regression for this cycle. The Middle Miocene (Unit VI) in the onshore corresponds to the Freites Formation (e.g. Gonzalez de Juana et al., 1980). Unit VII (Upper Miocene) Unit VII on seismic profiles is best defined in the western onshore where three depositional sequences are observed (Fig. 22a). Overall aggradational to subtle onlapping patterns predominate and involve from bottom to top the sequence boundaries SB-3.7 (10.5 Ma), SB-3.8 (8.2 Ma), SB-3.9 (6.3 Ma) and SB-3.10 (5.5 Ma). Seismic data in the west show that these sequence boundaries are characterized by erosional surfaces and medium- to high-amplitude reflectors with moderate continuity. Moving to the east, along the axis of the basin, Unit VII shows seismic facies characterized by low-amplitude, discontinuous reflectors. While the sequence boundaries are reasonably well defined in the western part of the area they are difficult to follow eastward because they are intercepted and offset by a complex system of listric faults (Figs. 9a and 10; see also Lilliu,
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1990; Daza and Prieto, 1990). In the offshore, in some areas mostly to the south, Unit VII shows a prograding sigmoidal configuration with a northeastdipping attitude (Fig. 21c). Note also that most of the sigmoid strata and the successive offlap breaks were partially eroded by canyons, thus displaying only remnant slope clinoforms truncated at the top (Fig. 23a).
Unit VII onshore Overlying the SB-3.7 (10.5 Ma), a fining-upward stacking pattern on logs and reported paleowaterdepths suggest that the lower portion of Unit VII represents an overall transgressive phase (Fig. 19c). Two major lithofacies are differentiated in Unit VII, as shown in wells located in the axis of the basin, which penetrated a 1200-m-thick section. The lower portion consists of siltstone and shale interbedded with white fine- to medium-grained poorly sorted sandstone. Plant remains and occasional lignite layers are present. Sedimentological analysis suggests that these sediments were deposited in environments ranging from continental to coastal plains. This portion is bounded at the base by SB-3.7 (10.5 Ma) and at its top by SB-3.8 (8.2 Ma). Its depositional sequence consists of transgressive and highstand deposits. The upper portions contain predominantly gray-olive shale with abundant benthic foraminifera, siltstone and poorly sorted sandstone. Sedimentological and paleontological analyses suggest that this upper portion was deposited in environments ranging from shallow marine to outer-shelf/upper bathyal. Two depositional sequences have been interpreted within this upper portion, bounded successively by SB-3.8 (8.2 Ma), SB-3.9 (6.3 Ma) and at its top by SB-3.10 (5.5 Ma). These depositional sequences are characterized mostly by lowstand deposits (Fig. 19c). A second deepening event is observed within this upper portion. However, the peak transgression is paleontologically poorly constrained due to the lack of reliable biostratigraphic successions that would permit a more precise dating of sequence boundaries and maximum flooding surfaces. Based on seismic interpretation and well data (Figs. 14a and 19c) the peak transgression occurred between sequence boundaries SB-3.8 (8.2 Ma) and SB-3.9 (6.3 Ma) and, consequently, the maximum transgression M2 is tentatively correlated with the 7-Ma maximum flooding event of Haq et al. (1987). Unit VII offshore The stratigraphic configuration of Unit VII in the offshore differs from its onshore equivalents. In the offshore passive margin domain, Unit VII is characterized by a single depositional sequence with a lowstand system tract of slope fan and lowstand prograding wedge deposits (Figs. 2 l c and 23a). Unit
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Fig. 23. Canyon systems associated with SB-3.10 (5.5. Ma). (a) Uninterpreted and interpreted segment of an offshore seismic profile. The SB-3.10 (5.5 Ma) sequence boundary marks a major erosional event. The line drawing also shows Unit VII (Upper Miocene) characterized by a prograding sigmoidal configuration (see Fig. 3c for location). (b) The detailed time-structure map of the top Upper Miocene shows the geometry and orientation of the canyons in the offshore area. VII is bounded at its base by SB-3.7 (10.5 Ma) and at its top by SB-3.10 (5.5 Ma) with its deeply incised valleys, which farther basinward (i.e. northeast) develop into large-scale submarine canyons. In some areas only isolated remnants of Unit VII can be observed on seismic profiles (Figs. 6a, 7b and 19d). To the north, where Unit VII has been penetrated by wells (Fig. 14a), mostly gray-greenish shale interbedded with siltstone and fine-grained sandstone has been reported. The presence of benthic foraminifera such as Cyclammina, Haplophrag-
moides, Alveo-Valvulinella suteri, Uvigerina, Lenticulina, Bolivina, Recurvoides, Buliminella, among many others, indicates that these sediments were deposited in upper to middle bathyal water depths characteristic for lowstand deposits. The limited occurrence of planktonic species of Globigerina nephenthes and Sphaeroidinellopsis paenedehiscens
gives an Late Miocene to earliest Pliocene age for this unit.
Summary Unit VII is overall transgressive and bounded by two major regional unconformities, i.e. at its base by SB-3.7 (10.5 Ma) and at its top by SB-3.10 dated 5.5 Ma. Regional sequence stratigraphy allowed the subdivision of Unit VII (Upper Miocene). In the onshore, Unit VII consists of three depositional sequences: the two lower depositional sequences, represented mostly by transgressive and highstand deposits, and the upper depositional sequence, characterized by lowstand deposits. The second peak transgression is interpreted to have occurred during the Late Miocene and is correlated with the 7 Ma of maximum flooding surface of Haq et al. (1987). Unit VII includes the La Pica Formation in the onshore
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN portion of the basin (e.g. Gonzalez de Juana et al., 1980). In the offshore area Unit VII consists of a single depositional sequence characterized by lowstand deposits. SB-3.10 (5.5 Ma), a Messinian erosional event
Unit VIII, the youngest part of the stratigraphic column, is separated from Unit VII by the sequence boundary SB-3.10, which no doubt is the most conspicuous feature of the basin. On seismic profiles, this unconformity is easily identified, showing large-scale erosion with deep-water channels and submarine canyons that create an irregular scoured surface with truncations of the parallel-planar reflectors of the subcropping depositional units. Fig. 23b shows a time-structure sketch map of the main canyon systems and the physiographic features of these northeast-trending canyons characterized by closer contours and landward V-shaped valleys. The canyons are better visible on shelf-parallel-oriented profiles which generally show the U-shaped canyon flanks. On perpendicular, i.e. dip-oriented profiles, the truncation appears minimal because the canyon surfaces often mimic the underlying slope clinoforms (Fig. 21c). The seismic facies of the canyon fill consists of low-amplitude, discontinuous to chaotic and irregular reflectors which often pinch out on the slope. Upward and on dip profiles a series of prograding clinoforms plunge from the wall canyon head and pinch out within the upper slope (Fig. 21c). The lower portion of the canyon fill consists of slumps and slope deposits, and the upper portion exhibits isolated lowstand prograding wedges. Well data suggest that the canyon fill is overwhelmingly dominated by shales which are characterized by monotonous high values on gamma-ray logs and low resistivity values occasionally interrupted by sandstone spikes (Fig. 19d). When the shaley infill is superposed on slope and basin shale, the well-log response of canyon fill and outside-canyon shales is virtually the same, i.e. without showing any abrupt changes in the log patterns. Based on well reports Prieto (1987) suggests that the canyons developed between 7.0 Ma and 4.5 Ma. The approximate age of the canyon formation can be reasonably related to a major sea-level fall that occurred during the Messinian, which corresponds to a 5.5 Ma sequence boundary of Haq et al. (1987). Note particularly in Figs. 6a and 20a that in the downdip direction the SB-3.10 surface is overlain by a pronounced regional deep-water onlap that mimics the underlying basal foredeep unconformity. This suggests that the deep-water canyon cutting event is strongly enhanced by the differential subsidence of the foredeep and therefore leads to the formation of a 'secondary basal foredeep unconformity'.
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Unit VIII (Plio-Pleistocene)
The overall regional pattern of Unit VIII (i.e. both west to east and north to south) is characterized by its well-defined wedge-shaped configuration (Figs. 8 and 9a). Fig. 15c is a time-isopach (two-waytime) of the unit. A monoclinal trough characterizes the isopachs. Considerable thickness variations distinguish this unit which pinches out to the south increasing to thicknesses in excess of 7000 m (~5 s) in the offshore area up to the shelf-break, where the sequence is involved in a major growth fault zone (Figs. 9a and l la). The maximum depth of more than 5 s in its northwestern part reflects the subsidence and infill of the foredeep by the eastward advance of the Neogene depocenter.On reflection seismic profiles of the western part of the basin Unit VIII consists of parallel to divergent mediumto high-amplitude reflectors with moderate continuity. Moving to the east (i.e. offshore) forestepping progradational sigmoidal configurations dominate (Figs. 24 and 25). Unit VIII onshore Fig. 19c is a composite log summary of well K that defines the chronostratigraphic horizons interpreted on seismic profiles and is tied to other wells. Unit VIII represents an overall regressive regime showing a distinctive change in log-facies from a fining-upward trend in the underlying Unit VII to a coarsening-upward trend (Fig. 19c). Unit VIII shows great lateral variation of seismic facies which is related to the progressive eastward movement of the Caribbean Plate relative to the South American Plate during the Neogene (Fig. 4c and d). Unit VIII is bounded at its base by SB-3.10 (5.5 Ma) and at its top by the actual topography. At least three depositional sequences have been interpreted within Unit VIII. These are bounded by SB-3.10 (5.5 Ma), SB-3.11 (4.2 Ma), SB-3.12 (3.8 Ma) and SB-3.13 (3.0 Ma) which are defined seismically by regional onlap and local truncation and tied to well logs. These depositional sequences are dominated by transgressive and highstand system tracts. The ages of the sequence boundaries are poorly constrained. Based on well reports, the lithofacies of Unit VIII consists of stacked patterns of coarsening-upward sequences of lignitic shale, red-brown micaceous siltstone and shale, poorly sorted fine- to coarse-grained sandstone, and conglomeratic sandstone. Faunas are very poor and marine fossils are lacking within this unit although fossil turtles, Corbicula sp. and fish teeth (Funkhouser et al., 1948), sparse species of Miliammina fusca (Hedberg et al., 1947) and the fluvial mollusk of Hyria trinitaria (Palmer, 1945) all suggest a Pliocene age. Unit VIII was deposited in a littoral/marginal marine to mostly continental setting.
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Fig. 24. Seismic expression of Plio-Pleistocene (Unit VIII) prograding sequences on the offshore Orinoco platform. For location see Fig. 3c. SB-3.15 is the inferred top of the Pliocene. SB-3.10 is the Messinian canyon-forming event. Unit VIII offshore The stratigraphy of Unit VIII in well A (Fig. 19d) shows an overall coarsening- and thickening-upward trend. On seismic profiles, this unit is represented by a thick (~7000 m) progradational sigmoidal pattern consisting of multiple offlapping depositional sequences that represent high-frequency Plio-Pleistocene glacio-eustatic sea-level fluctuations (Fig. 24). Reliable paleontological data are not available but the seismic resolution is good and several sequence boundaries were interpreted. The dating was inferred by comparing the interpretation with the idealized Neogene model of Vail et al. (1991). Seven Plio-Pleistocene sequences are recognized. These sequences are mainly third-order, but some may be fourth-order sequences (Vail and Wornardt, 1990; Mitchum and Van Wagoner, 1991; Vail et al., 1991). The inferred dates can be obtained by combining Fig. 25b and Fig. 26. In general, the depositional sequences bounded by these sequence boundaries show a well-developed lowstand system tract with slope fan and prograding wedge deposits, a very thin transgressive system tract and variable thickness highstand system tracts. Fig. 25b
shows the geographic position of the shelf edge of each sequence. The depositional framework was controlled by the advance of the Orinoco Delta since the Late Miocene. The shelf edges display a continued shift from the southwest to the northeast, which is due to the high rate of sediment supply from the west. Furthermore, the progressive development of a growth fault system was established with a base that is controlled by the top of the underlying passive margin (Fig. 9b). Unit VIII consists of a typical deltaic progradation composed of the following. (1) A lower portion characterized by gray-olive shale with abundant foraminifers such as Orbulina universa, Globigerinoides obliquus, Lenticulina sp., Bathysiphon sp., and Cyclammina interbedded with a few thin layers of turbidites and fine-grained sandstone. Sedimentological and paleontological analyses indicate that these sediments were deposited in an upper- to middle-bathyal water depth. (2) A middle portion consisting of fine- to medium-grained sandstone interbedded with gray siltstone and shale. The presence of Sphaeroidinella seminulina marks the top of the Pliocene. In addi-
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN
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Fig. 25. Offshore Orinoco platform, prograding systems of Unit VIII (Pliocene). (a) Line drawing of an offshore W-E profile adjacent to the Orinoco Delta showing the sequence stratigraphy of Unit VIII (Plio-Pleistocene). The heavy dashed line marks the interpreted top of the Pliocene, i.e. SB-3.15. (b) Map showing the position of the prograding shelf-margin from Late Miocene to Present. The growth fault system to the east is in essence a Quaternary growth system. tion these shales contain pelecipods, gastropods and foraminifera such as Ammonia, Elphidium, Florilus, Haplophragmoides, Eponides, Globigerina and Globigerinoides. Sedimentological and paleontological analyses indicate that these sediments were deposited in a middle to outer shelf environment that probably represents a prodeltaic facies. (3) An upper portion that consists of coarsening- and thickening-upward strata of fine- to coarsegrained sandstone with frequent microconglomerates and abundant fossil and wood fragments, chert, pellets and pyrite nodules. Interbedded with this facies, thin layers of gray siltstone and shale are present, with planktonic foraminifera such as Globorotalia truncatulinoides, suggesting a Pleistocene to Recent age. Sedimentological and paleontological analyses indicate that these sediments were deposited in a shallow-water depth. They represent the delta front and delta plain facies of the Orinoco Delta. Based on stacking patterns in well logs, the paleo-water depth and seismic configuration, Unit VIII represents the overall regressive Pliocene to Present phase of the Eastern Venezuelan Basin.
TERTIARY PALEOGEOGRAPHY OF EASTERN VENEZUELA
The map of Fig. 27a shows the Oligocene facies distribution, its inferred updip zero edge, and the starved downdip continuation of that formation. Because of the limited data the Oligocene paleogeography is difficult to reconstruct. The Oligocene zero edge is due to erosional truncation. Downdip from the lowstand deltas of well A, the Oligocene is absent in well B probably due to sediment starvation. To the northwest of the starved area more complete pelagic sediments are reported from the Serranfa del Interior and from Trinidad. Combined with the plate tectonic reconstruction shown in Fig. 4b this paleogeographic map suggests that the area now occupied by the Serranfa del Interior and the Venezuelan offshore to the north was still an intact passive margin during the Oligocene. The paleogeographic map of Fig. 27b depicts the main depositional regimes of the Early Miocene (25 Ma), at the onset of the foredeep phase. To the west and south of the basin, continental to coastal
466
J. DI CROCE et al. ii ii ii iiiiii
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]500 km for the Caribbean plate north of the Mor6n-E1 Pilar fault system. Observed offsets are much smaller but may represent minimum offsets as some major strike-slip faults like the E1 Coche-North Coast fault zone are not well mapped. The amount of shortening in the Serranfa del Interior fold-thrust belt has been estimated at 45-90 km based on balanced cross-section estimates of Parnaud et al. (1995). Right-lateral strike-slip faults that strike east-southeast to southeast have well-mapped offsets ranging from 10 to 40 km. (B) Cenozoic sedimentary basins in northern South America. Major basins are formed either by strike-slip faults of the right-lateral Mor6n-E1 Pilar fault system or as foredeeps along major thrust faults at the southern limits of the fold-thrust belts. The Morochito basin studied by Parnaud et al. (1995) is a piggyback basin formed by thrusting within the Serranfa del Interior thrust belt. (C) Lithologic belts of northern South America formed as the Caribbean arc obliquely collided with the passive margin of northern South America. The age of deformation youngs from an Eocene-Oligocene event that formed the Foothills fold-thrust belt and obducted crystalline rocks of the Villa de Cura belt, to an Oligocene-Pleistocene event that formed the Serranfa del Interior fold-thrust belt, to the Trinidad belt where deformation began definitely in the Late Miocene and is continuing to the present.
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY original trench setting off northwestern South America while the low-grade parts of the belt experienced a similar style of deformation at a shallower depth in the trench (Fig. 4C). Structures mapped by Av6 Lallemant (1997) indicate a protracted history of northwest-dipping (present geometry) thrusts. The syn-metamorphic thrusting and folding event affecting passive margin sedimentary rocks of the Northern Range of Trinidad has an opposite sense of dip and vergence to the Araya Peninsula and may indicate deformation in a different tectonic setting from its present geographic location (Algar and Pindell, 1993). Caucagua-Paracotos-Villa de Cura belt These rocks are mainly metasedimentary passive margin rocks overthrust by the large, ophiolitic Villa de Cura nappe (Fig. 4C). Foothills belt This belt forms a narrow fold-thrust belt in Cretaceous and Paleogene sedimentary rocks south of the Caucagua-Paracotos-Villa de Cura belt (Fig. 4C). Figueroa de Sanchez and Hern~indez (1990) document late Eocene-Oligocene north to south overthrusting into the Gu~irico sub-basin adjacent to this belt (Fig. 4B) using evidence from an exploration well. Serrania del Interior belt This east-northeast-trending fold-thrust belt falls on strike with the deformed zone in Trinidad and has been well studied and dated during petroleum exploration of the Eastern Venezuelan basin (Erlich and Barrett, 1992; di Croce et al., Chapter 16; Flinch et al., Chapter 17) (Fig. 4C). Thrusting is south- to southeast-vergent, involves unmetamorphosed Cretaceous to Pliocene lithologies, and occurred in Oligocene through Pliocene time (Parnaud et al., 1995; Roure et al., 1995). Based on balanced crosssections, the total north-south shortening across the belt is about 45-90 km (Roure et al., 1995). Guyana shield The Guyana shield is a pre-Grenville age Precambrian craton that dips northward beneath the Foothills and Serranfa del Interior fold-thrust belts (Fig. 4C). Trinidad belt The Trinidad belt is a zone of strike-slip and thrust deformation of Middle Miocene to Recent age in Trinidad, the Gulf of Paria, and the eastern shelf of Trinidad. This deformation forms the topic of this paper and is also discussed by Flinch et al. (Chapter 17) and di Croce et al. (Chapter 16) (Fig. 4C).
505
SUMMARY OF PREVIOUS TECTONIC MODELS FOR TRINIDAD Introduction
The tectonic interpretation of complex structural relationships in Trinidad spans the period since systematic and detailed outcrop mapping of the entire island was completed at a scale of 1 : 100,000 by Kugler (1959). Onland mapping in Trinidad combined with the larger-scale recognition of the eastward relative motion of the Caribbean plate (Molnar and Sykes, 1969; Fig. 2B) has led to a variety of predictive tectonic models for the Trinidad region. Most of the models agree that the MesozoicPaleogene passive margin of northeastern South America converted to the zone of presently active deformation sometime during the Neogene and that this event was linked to the interaction between the margin and the eastward or southeastward-moving Caribbean plate (e.g., Pindell and Barrett, 1990). Differences in models center on three fundamental questions concerning the interaction of the exotic Caribbean arc and the passive margin of northeastern South America: (1) When did the Trinidad area of the northeastern passive margin of South America convert from a passive to active margin? (2) What type of active margin was initiated (strike-slip vs. thrust) by the oblique passage of the Caribbean arc along the passive margin? (3) What were the main fault controls (normal vs. strike-slip vs. thrust) on active margin sedimentation? Sampling of tectonic models for Trinidad
A sampling of four of the more recent and widely cited tectonic models are illustrated in Fig. 5 and briefly summarized here. Perez and Aggarwal (1981) They compile earthquake epicenter locations and focal mechanisms of larger events to propose a shift in the location of strike-slip faulting from the east-west-striking, right-lateral E1 Pilar fault zone to the northwest-striking, right-lateral Los Bajos and E1 Soldado fault zones in the Gulf of Paria, which they inferred to have a cumulative right-lateral offset of 22 km (Fig. 5A). In this model, these surficial strike-slip faults act as transform or tear faults separating the subducted oceanic part of the South America plate from the unsubducted continental part of the plate (Fig. 2B). These authors proposed that the north-to-south shift in plate motion from the E1 Pilar fault zone to the Los Bajos-E1 Soldado system (Wilson, 1940) was a Pleistocene response to
506
S. BABB and R MANN Trinidad from a relatively minor offset (~30 km) strike-slip fault on the Paria Peninsula (Fig. 4A).
Robertson and Burke (1989) They present a regional strike-slip model based on a grid of seismic lines tied to wells in the North Coast basin off the north coast of Trinidad and field results from Late Pleistocene outcrops of the Northern basin south of the E1 Pilar fault zone (Fig. 5C). They conclude that Trinidad and the North Coast basin was a diffuse and active zone of fightlateral faulting related to the eastward motion of the Caribbean plate past South America (Fig. 4A). They proposed that the stratigraphy of Trinidad was the result of the strike-slip juxtaposition of several blocks with distinct depositional histories that are now expressed morphologically as the Northern, Central and Southern Ranges and the Northern and Southern basins. Observed deformation patterns in Late Pleistocene rocks near the E1 Pilar fault zone were attributed to a right-lateral simple shear mechanism.
Erlich and Barrett (1992)
Fig. 5. Summary of previous tectonic models proposed for Trinidad and the easternmost Eastern Venezuelan basin. (A) Perez and Aggarwal (1981) emphasize strike-slip tectonics along the southern limit of the Benioff zone (cf. Fig. 2B). (B) Speed (1985) and Russo and Speed (1992)emphasize a southeastwardto south-southeastward-verging fold-thrust belt with little or no associated strike-slip tectonics in Trinidad. (C) Robertson and Burke (1989) emphasize a diffuse zone of east-west-striking right-lateral strike-slip faults. (D) Erlich and Barrett (1992) emphasize two, east-west-striking strike-slip faults and intervening normal faults bounding a triangular-shaped rift zone in northern Trinidad. See text for further discussion of each model.
They followed the idea of a predominantly strikeslip boundary but proposed normal throws on southeast and northeast faults bounding the rifted margins of the triangular Gulf of Paria-Northern basin (Fig. 5D). The mechanism for rifting in the Gulf of Paria-Northern basin was attributed to simultaneous fight-lateral motion on the E1 Pilar and Los Bajos fault zones and transfer of slip from the E1 Pilar to Los Bajos. A similar rift-related model for the Gulf of Paria-Northern basin was proposed by Salvador and Stainforth (1968). A Miocene age of movement was inferred from previously published structural and stratigraphic data along with some new industry data presented in their paper. The Warm Springs fault zone was interpreted as a post-Miocene horsetail splay fault of the Los Bajos fault zone.
Previous published studies using industry data from Trinidad the southward widening of the Benioff zone formed by the subducted Atlantic slab beneath Trinidad (Fig. 2B).
Speed (1985) and Russo and Speed (1992) Using outcrop, gravity and regional seismicity data, these workers proposed that Trinidad and eastern Venezuela is a south to southeast-verging foldthrust belt with an associated foreland basin formed by thrust loading (Fig. 5B). In their model, the E1 Pilar fault in Trinidad is regarded as a north-dipping thrust along which the Northern Range has been uplifted. In this model, the main strike-slip plate boundary fault has not yet propagated eastwards into
Several of the above tectonic models were handicapped by the lack of subsurface data needed to constrain the character of fault zones like the controversial E1 Pilar fault zone and the overall structural style and age of structures in the Gulf of Paria and Northern basins (Fig. 4B). While limited industry seismic and well data had been published from the Gulf of Paria (Eva et al., 1989), offshore Los Bajos fault zone (Tyson, 1989) and the Gulf of PariaNorthern basin (Payne, 1991), most studies in the Southern basin-Gulf of Paria-Northern basin region were local in scale, did not present a large amount of systematic mapping using seismic grids tied to
D E V E L O P M E N T OF A N E O G E N E T R A N S P R E S S I O N A L PLATE B O U N D A R Y
wells, and consequently did not attempt a tectonic synthesis for the entire area of northern Trinidad. However, in the area of the Caribbean Sea north of Trinidad, Robertson and Burke (1989) synthesized a large industry data set consisting of eleven wells and 5000 km of seismic profiles. Payne (1991) proposed large, late Neogene normal movement on the E1 Pilar fault zone using seismic reflection data that were reinterpreted by Algar and Pindell (1993) who reached the same interpretation of normal faulting. Payne (1991) also proposed that the Gulf of Paria is a typical rift basin characterized by a horst and graben structure. He proposed two types of faults depending on their strike: eastwest or northeast-southwest-striking faults are transpressional while northwest-southeast-striking faults are transtensional to extensional.
DATA USED IN THIS STUDY OF THE GULF OF PARIA AND NORTHERN BASIN, TRINIDAD
A map showing new seismic reflection and well data used in this paper is shown in Fig. 6A. The data set includes approximately 1600 km of migrated seismic reflection data. Some of these data were reprocessed by the Petroleum Company of Trinidad and Tobago (Petrotrin). These data are tied to 28 wells with electric logs and well completion reports. The seismic reflection data used in this study are concentrated in the Gulf of Paria while the well data are concentrated in the Northern basin and its nearshore area in the Gulf of Paria (Fig. 6A). Only one reflection line used in this study (Fig. 36) has been previously published but is differently interpreted by Payne (1991) and Algar and Pindell (1993).
MAJOR SUB-BASINS AND HIGHS OF THE GULF OF PARIA AND WESTERN PART OF THE NORTHERN BASIN
Tectonic map and regional cross-section A tectonic map of Trinidad showing the major onshore basins and ranges was adapted and simplified from the more detailed regional maps of Kugler (1959) and Persad (1984) (Fig. 6A). This map is keyed to a regional cross-section A-A' adapted from section D-D' that accompanied the regional geologic map of Persad (1984) (Fig. 7). Formation names shown on the cross-section are from Persad (1984) and are also shown in tabular form in Fig. 8. Fig. 6B provides an index map of more detailed maps of the Gulf of Paria and Northern basins that
507
are derived from well and seismic reflection data and presented later in this paper.
Basin and sub-basin nomenclature used in this study Basins For this study, the term 'Gulf of Paria basin' is used to describe all sedimentary rocks in the northern offshore Gulf of Paria area that overlie latest Early Cretaceous rocks of the passive margin section of northern South America that is horizontally ruled on Figs. 7 and 9. The 'Northern basin' or 'Caroni basin' (cf. Robertson and Burke, 1989 for this second usage) refers to the similar succession of sedimentary rocks within this onshore area of Trinidad (Fig. 6A). Data coverage in the Gulf of Paria basin extends only to the Trinidad-Venezuela boundary that bisects the Gulf of Paria in a roughly north-south direction (Fig. 6A). Data from both the Venezuelan and Trinidadian part of the Gulf of Paria basin are described by Flinch et al. (Chapter 17) and Di Croce et al. (Chapter 16). Sub-basins Mapping of seismic sequences on seismic reflection lines tied to wells reveals the presence of four sub-basins within the Gulf of Paria basin and the western part of the Northern basin (Fig. 9A). These basins, formally named in this paper for the first time, include the following. (1) E1 Pilaf sub-basin with a total Late MioceneRecent thickness in two-way time of about 3 s. This semicircular basin is asymmetrical with its steep northern side bounded by the sub-vertical plane of a southern strand of the E1 Pilar fault zone and its more gently sloping, southwestern flank bounded by the Gulf high, a remnant Cretaceous passive margin carbonate bank (Fig. 9A). (2) Puerto Grande sub-basin with a total Late Miocene-Recent thickness in two-way time of about 2.3 s. This rectangular basin is also asymmetrical with its steep northern side bounded by an unnamed fault south and parallel to the E1 Pilar fault zone. Its southwestern end is bounded by the Avocado high, a remnant Cretaceous passive margin carbonate bank (Fig. 9A). (3) Goodrich sub-basin with a total Late Miocene-Recent thickness in two-way time of about 3.3 s. This basin has a more complex and irregular geometry than the E1 Pilar and Puerto Grande sub-basins. Its northern and northwestern edges are bounded by the Galf and Domoil highs (Fig. 9A). Its northeastern and southwestern edges are bounded by large oblique-slip faults between the E1 Pilar and Warm Springs fault zones and its southern boundary is bounded by the Warm Springs fault zone.
508
S. BABB and E MANN
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY PREVIOUS LITHOSTRATIGRAPHIC STUDIES OF THE GULF OF PARIA AND NORTHERN BASIN Introduction Previous studies subdividing the lithostratigraphy of outcrops and well data from the Northern basin and Gulf of Paria area include Kugler (1953), Barr and Saunders (1968), Higgins and Saunders (1968), and Payne (1991). Payne (1991) made the first attempt to subdivide the post-Middle Miocene, subsurface sequences for the Northern basin using wells and seismic reflection surveys. He defined a depositional sequence framework using differences in seismic responses, provided relative ages of the sequences, and discussed stratigraphic correlation across the Northern basin. This study attempts to build on Payne's work by integrating the subsurface geology of the Gulf of Paria and Northern basins and the outcrop geology of the Northern and Central Ranges into a single stratigraphic framework (Fig. 8). Outcrop stratigraphy of the Northern and Central Ranges adjacent to the Gulf of Paria-Northern basins Structure of Jurassic-Cretaceous rocks in the Northern and Central Ranges Late Jurassic-Early Cretaceous carbonate, siliciclastic, and volcanic rocks in Trinidad crop out as a variety of mapped formations of low-grade metasedimentary rocks that have been studied in detail in the Northern Range (Trechmann, 1935; Kugler, 1959; Barr and Saunders, 1968; Furrer, 1968; Saunders, 1972; Potter, 1976; Wadge and Macdonald, 1985; Frey et al., 1988; Algar and Pindell, 1993) (Fig. 10). Unmetamorphosed Barremian to AptianAlbian black-gray shale (Cuche Formation) has been described from deep wells drilled in the onshore Northern basin and offshore Gulf of Paria basin as well as from outcrops within the Mount Harris push-up block along the Central Range right-lateral strike-slip fault (Barr, 1952; Bartenstein et al., 1957, 1966; Koutsoukos and Merrick, 1985; Algar and Pindell, 1993) (Fig. 6A).
509
The Mesozoic section in the Northern Range exhibits large-scale northward-verging folds formed either by large-scale syn-sedimentary slumping or tectonic deformation along with four phases of subsequent brittle faulting (Algar and Pindell, 1993) (Figs. 7 and 10A, B). The major structure of the Northern Range is a large, overturned antiform whose gently curved, east-west-trending axial trace roughly parallels the 500- to 925-m-high, topographic crest of the range (Fig. 10A). The Northern Range is a direct along-strike continuation of the Coastal Fringe/Margarita belt of the Araya-Paria Peninsula of Venezuela (Gonzalez de Juana et al., 1968; Av6 Lallemant, 1997; Fig. 4C) which exhibits many of the same Mesozoic rock types and styles of deformation (Fig. 10A). Fission track ages by Algar and Pindell (1993) from Mesozoic metamorphic rocks of the Northern Range indicate that the main phase of shallow unroofing and erosion through a temperature of 200o-250 ~ began early Late Miocene about 11 m.y. ago. This age of unroofing and initial erosion of Mesozoic rocks in the Northern Range is consistent with the presence and composition of the siliciclastic wedge of Late Miocene age (Cunapo Formation) shed into the Northern basin (Fig. 8). The isolated outcrop in the Lower Cretaceous Cuche Formation in the Central Range does not exhibit either large-scale folding or metamorphism but does show evidence of brittle faulting probably related to its position and late Neogene uplift history within the Mount Harris push-up block of the Central Range fault zone (Fig. 6A). Jurassic-Cretaceous outcrop stratigraphy in the Northern and Central Ranges Although the lithostratigraphy of the Northern Range is well mapped, the formation names and structural interpretations by Kugler (1959) were questioned and subsequently modified by Algar and Pindell (1993). We have maintained the Kugler (1959) interpretations as outlined in Fig. 10B because they are consistent with the main features of the Kugler (1959) geologic map that are schematically shown in Fig. 10A. The Mesozoic section of the Northern Range con-
Fig. 6. (A) Tectonic map of Trinidad showing major fault systems and associated sedimentary basins. Inset shows direction and rate of South America plate relative to the Caribbean plate from DeMets et al. (1994). Thin lines in the Gulf of Paria and Northern basins are multi-channel seismic reflection lines provided by the Southern Basin Consortium and Petrotrin for use in this study. Wells are identified by capital letters. Well C corresponds to the Couva Marine-1 well of Bray and Eva (1983). Transtensional strike-slip basins including the E1 Pilar, Goodrich and North Soldado sub-basins of the Gulf of Paria are found mainly offshore, while basins affected by transpressional tectonics like the Northern basin and Southern basin are found onshore. Outcrops of metamorphosed Mesozoic passive margin rocks are restricted to the Northern Range and Paria Peninusla of Venezuela. One isolated unmetamorphosed outcrop of lower Cretaceous passive margin rocks (Cuche Fm.) is found in the Mount Harris area. (B) Index map showing locations of more detailed maps presented in this paper and line of cross-section A-A' shown in Fig. 7. Regional cross-section A-A', modified from Persad (1984) and Robertson and Burke (1989), is shown in Fig. 7.
510
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sists of a latest Jurassic through Barremian section of low-grade metamorphic rocks, mainly of a mixed siliciclastic and carbonate origin. These metamorphic rocks are conformably overlain by metamorphosed sedimentary rocks of Late Cretaceous age (Fig. 10B). Most workers agree that carbonate and minor evaporitic Mesozoic rocks of the Northern Range and in the Couva Marine-1 well of the Gulf of Paria basin (Bray and Eva, 1983; Eva et al., 1989) record an early passive margin phase in the geologic history of Trinidad as shown on the chart in Fig. 8. Tertiary passive margin stratigraphy In general, the Cenozoic stratigraphic succession of the Northern Basin-Gulf of Paria is indicative of a progressive shallowing and infill from Paleogene-Early Miocene, deeper-water, finergrained passive margin or 'flysch-type sediments' of the Chaudiere, Pointe-a-Pierre, and Brasso Formations to the Early Miocene-Late Miocene, shallower-water and coarser-grained active margin or 'molasse-type' sedimentary rocks of the Cunapo and Manzanilla Formations (Tyson and Ali, 1990) (Fig. 8). The Paleogene-Early Miocene deeperwater rocks are unevenly distributed and generally thin as one moves from south to north across the Central Range as seen on the regional cross-section in Fig. 7. For this reason, Paleogene-Early Miocene deeper-water rocks are thin in the subsurface of the Northern basin and are commonly difficult to distinguish on seismic reflection lines from the overlying, Neogene shallower-water units.
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DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
511
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Fig. 8. Nomenclature, age, lithology, and depositional environments of formations in the Northern basin and Gulf of Paria and their correlation to sequences defined in this study. Data are compiled from Barr and Saunders (1968) and this study. Lined pattern indicates passive margin lithologies and is keyed to the formations with a lined pattern shown on the cross-section in Fig. 7. Tectonic interpretations in right column are proposed in this study. (1993) have inferred that the shallow-water conditions over the Central Range represent the surface expression of large-scale folds forming along the plate boundary zone. Termination of carbonate sedimentation of the Tamana Formation correlates with the rapid influx of shallow water siliciclastic sedimentary rocks during the Late Miocene (Tyson et al., 1991; Erlich et al., 1993) (Fig. 8).
Stratigraphic nomenclature and age control of Payne (1991) for the Northern basin and Gulf of Paria Payne (1991) subdivided the Late MioceneRecent stratigraphic succession of the Northern Basin-Gulf of Paria into two sequences, A and B. Sequence A consists of the massive, Late Miocene Cunapo conglomerate, which aggrades and progrades southward from the southern flank of the Northern Range and interfingers with sandstone of the Manzanilla Formation of the Northern and
Gulf of Paria basins (Fig. 8). The maximum drilled thickness of the Manzanilla Formation approximates 6500 feet (1950 m) near its depocenter in the E1 Pilar sub-basin south of the E1 Pilar fault zone (Fig. 9A). Sequence B consists of sandstone and clay of the Plio-Pleistocene Springvale and Talparo Formations (Fig. 8). The thickness of the Springvale Formation varies widely and suggests changing tectonic and sedimentary conditions during its deposition. For age control of his sequences A and B, Payne (1991) uses the palynological zonation of the Northern Basin developed by E. Gonzales (contract palynologist for the Petroleum Company of Trinidad and Tobago). The zonation is based on statistical pollen counts and the first appearances of distinctive marker palynomorphs. The specific species upon which this zonation was based were not provided by Payne (1991) or in the internal reports available for this study. The pollen zones identified by Gonzales include: UM I, UM II pollen zones Upper Miocene; PI, PII, PIII pollen zones - - Pliocene; and
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S. B ABB and E MANN
Fig. 9. (A) Major tectonic elements and total isochron map for Late Miocene through Pleistocene sequences 3, 4, and 5 in the Gulf of Paria (contour interval is 1000 ms). The depocenters are mainly formed by transtensional fault movements associated with or between the E1 Pilar and Warm Springs fault zones. The Domoil and Gulf highs are remnant, constructional carbonate topography inherited from a Late Jurassic-Early Cretaceous carbonate passive margin. Well locations are given by letters. (B) Location map of regional seismic sections and geologic sections based on well logs shown in Figs. 11-20.
a Pleistocene zone. These zones and their correlation to the lithostratigraphic subdivision of the Northern Basin by Payne (1991) are used in this study and are s u m m a r i z e d on the stratigraphic column in Fig. 8.
A correlation b e t w e e n these pollen zones with the planktonic biochronozones shown on the Haq et al. (1987) sea level chart was proposed by R. Liska (pers. commun., 1995): U M I is equivalent to lower
D E V E L O P M E N T OF A N E O G E N E T R A N S P R E S S I O N A L PLATE B O U N D A R Y
N 16; UM II is equivalent to N 17 and the upper part of N 16; the P zones are probably equivalent to N18 and part of N19 (3.5-5.2 Ma).
SEISMIC REFLECTION DEFINITION AND G E O L O G I C CORRELATION OF LATE M I O C E N E TO PLEISTOCENE SEQUENCES 3-5 IN THE GULF OF PARIA AND NORTHERN BASINS
Criteria for defining seismic sequences in the subsurface of the Gulf of Paria and Northern basin The definition of seismic sequences proposed in this study was based on the recognition of onlap, downlap, toplap or truncation surfaces on seismic reflection lines. The sequence boundaries were correlated on all seismic reflection lines and wells in the seismic grid shown in Fig. 6A. Using these criteria, we identify five Late Jurassic-late Neogene sequences in the subsurface of the Gulf of Paria and Northern basins and correlate them with previously proposed lithologic formations in Fig. 8 and on the cross-section of Fig. 7. The five proposed sequences include the following. Sequence 1: Late Jurassic-Early Valanginian carbonate megaplatform consisting of carbonateevaporitic lithologic cycles (Fig. 8). This sequence is described in detail by Babb (1997) and is not discussed in this paper. Sequence 2: Middle Valanginian-middle Aptian restricted carbonate banks developed on the megaplatform (Fig. 8). This sequence consists of evaporite-sand-shale cycles formed in a restricted bank setting. This sequence is described in detail by Babb (1997) and is also not discussed in this paper. Sequence 3: Late Miocene-Early Pliocene shallow marine to brackish water conglomerate and sandstone. These rocks are equivalent in part to the Cunapo and Manzanilla Formations known from wells along the northern flank of the Northern basin (Persad, 1984, 1985) (Fig. 8). The Cunapo Formation represents the first influx of coarse sedimentary rocks from the Northern Range into the Northern basin and interfingers southwards with sandstone of the Manzanilla Formation (Fig. 8). Sequence 4: Early to middle Pliocene inner neritic to shallow marine conglomerate, sandstone, silt and clay. These rocks are equivalent to the Springvale Formation and the upper part of the Manzanilla Formation known from outcrops in the uplifted flanks of the Northern basin (Persad, 1984, 1985) (Fig. 8). Water depths increased slightly during deposition of sequence 4 (Springvale Formation) as shown by the retrogradational nature of the well logs from this unit discussed later in this paper.
513
Sequence 5: Late Pliocene to Pleistocene marine to brackish-water sand, silt, clay, and minor conglomerate. These rocks are equivalent to the Talparo and Cedros Formations known from outcrops in the uplifted northern and southern flanks of the Northern basin (Persad, 1984, 1985) (Fig. 8). These two formations exhibit increasingly brackish-water environments probably related to the regional uplift and shallowing of the basin floor during this period and the constriction of the basin caused by uplift of the Northern and Central Ranges.
Seismic character and definition of sequence 3: Late Miocene Manzanilla and Cunapo Formations Distribution of sequence 3 Using the seismic lines and wells shown in Fig. 6A, sequence 3 was mapped in the E1 Pilar and Goodrich sub-basins of the Gulf of Paria basin. Sequence 3 is also known from wells and seismic lines to extend eastward and underlie much of the Northern basin. Sequence 3 is thickest in the E1 Pilar sub-basin adjacent to the E1 Pilar right-lateral strike-slip fault zone and thins to the south in the Goodrich sub-basin (Fig. 9A).
Base reflector of sequence 3 The base of sequence 3 on a north-south line crossing the Gulf high between the E1 Pilar and Goodrich sub-basins onlaps the top of sequence 2 formed by remnant topography of the Cretaceous carbonate bank in the center of the Gulf of Paria basin (Fig. 11). On a north-south line crossing the Avocado high, another Cretaceous carbonate high to the east of the Gulf high, sequence 3 can also be observed onlapping sequence 2 (Fig. 12). In the south of the Gulf of Paria basin, sequence 3 rests unconformably on a northward-thinning packet of reflectors that is known from wells C in Fig. 12 and V in Fig. 13 to be deep-water sandstone and shale of the Middle Miocene Brasso Formation (Fig. 8). The Brasso Formation along with the rest of the underlying, deep-water early Tertiary 'passive margin' section of Trinidad characteristically thins northward onto Early Cretaceous carbonate bank rocks of sequence 2 (Persad, 1984) (Fig. 7).
Top of sequence 3 The top of sequence 3 is defined on the regional northwest-southeast line in the Goodrich sub-basin by onlap, toplap, and truncation reflection terminations as seen on the seismic lines shown in Figs. 1113.
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Fig. 11. North-south seismic section showing representative example of seismic sequences 2, 3, 4, and 5 with sequence 3 onlapping the Lower Cretaceous carbonate bank and the top of sequence 3 defined by toplap and truncation (line of section shown on map in Fig. 9B). The Gulf high is inferred to be constructional carbonate morphology that is elongate in a northwest-southwest direction and developed on relatively flat megacarbonate platform of sequence 1 (see Fig. 9A for map view of Gulf high). Steep sides of platform led to gravitational slumping of sequence 3 at its edges. On the southern margin of the bank, low-angle normal faults sole out at the base of sequence 3. At the base of the northern margin of the bank, the basal, mounded part of sequence 3 suggests gravity flows derived from the Gulf high.
Seismic facies within sequence 3 A variety of seismic facies are present within sequence 3 that include the following types: mounded to chaotic, wedge, hummocky, sub-parallel to hummocky, and prograding. Mounded and chaotic facies are present in depositional lows at the base of the Gulf high in the E1 Pilar sub-basin (Fig. 11). Low-angle faulting is associated with this facies on the southern flank of the Gulf high carbonate bank (Fig. 9A). Near the northern margin of the E1 Pilar subbasin near the E1 Pilar strike-slip fault zone, reflections within sequence 3 consist of predominantly high-amplitude, discontinuous reflectors that have a wedge-shaped external form which tapers in a
northward direction (Fig. 12). On the southern flank of the Avocado high (Fig. 9A), sequence 3 exhibits variable-amplitude, continuous-discontinuous reflectors that show medium frequency and are subparallel to h u m m o c k y (Fig. 12). This seismic section shows reflections south of well C that are of lower frequency and slightly higher continuity than the equivalent section north of well C. In the eastern part of the Goodrich sub-basin, sequence 3 exhibits sigmoidal-oblique, high-lowamplitude reflection packets whose sense of progradation is from northwest to southeast (Fig. 13). Well V confirms that this facies consists of sandstone and shale.
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Fig. 12. North-south section showing sequence 3 overlying the Lower Cretaceous section to the north and the Lower Miocene Brasso Formation to the south (line of section shown on map in Fig. 9B). Logs of wells B and C provide lithologic control on seismic facies. Well B shows that sequences 3 and 4 are conglomeratic and well C shows that the equivalent section is more distal sandstone and shale. This pattern of north-to-south fining is observed throughout the Gulf of Paria and Northern basins and is interpreted as coarse sedimentation shed off the east-west-striking E1 Pilar fault zone.
Correlation of seismic facies of sequence 3 with lithologies from electric logs and cores from wells We have made stratigraphic cross-sections based on aligning wells in the Goodrich sub-basin and Avocado high to constrain the lithology and sedimentary facies of seismic sequences 3, 4 and 5. Cross-sections were made in both an approximately north-south (Fig. 14) and in an approximately eastwest direction (Figs. 15 and 16) to provide a better understanding of how seismic/sedimentary facies vary on a basin-wide scale (locations of aligned wells shown in Fig. 9B). Spacing between wells along these sections varies from 5 to 16 km. Lithologic variations of sequence 3 in a north-south direction Sequence 3 can be divided into three lithologic units in wells B, C1, C2, and V which vary in character in a north-south direction (Fig. 14). Well B at the western end of the Puerto Grande sub-basin adjacent to the E1 Pilar strike-slip fault contains
mainly conglomeratic rocks of the Cunapo Formation of Late Miocene to Early Pliocene age (Fig. 8). Unit 1 of this conglomeratic section is an 80-m-thick aggradational stack of massive conglomerate with thin shale interbeds increasing towards the top of the unit (Fig. 14). Unit 1 is bounded at its base by an unconformity surface above the Early Cretaceous carbonate platform section of the northern Avocado high (Figs. 9, 12). The top of unit 1 is bounded by a 10.5-m-thick shale interval. Unit I appears to pinch out to the south. The character of this pinch out is demonstrated in the seismic line through sequence 3 in Fig. 12. A conglomeratic section very similar to unit I is present in well F within the E1 Pilar fault zone and on the north flank of the Puerto Grande sub-basin (Fig. 9A). Unit II in well B begins with sandstone and conglomerate units which display a blocky-shaped electric log signature and are possibly channel features (Fig. 14). The electric log profile suggests aggradation of units I and II. At wells C 1 and C2, 10 km to the southwest of well B, and well V 18 km to the southwest of B, unit II shows a more distal fa-
D E V E L O P M E N T OF A N E O G E N E T R A N S P R E S S I O N A L PLATE B O U N D A R Y
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Fig. 13. Northwest-southeast seismic section showing the three sequences mapped and defined in this study (line of section shown in Fig. 9B). Sequence 3 is a southeastwardly prograding unit defined at its upper boundary by onlap, toplap, and minor truncation. During the Late Miocene-Early Pliocene, a reversal in progradation direction is indicated by the northwestwardly downlap of sequence 4, which thins to the north by onlap. Within the upper part of sequence 4, a lapout surface is interpreted as a maximum flooding surface. The boundary between sequences 4 and 5 appears concordant except in the very southeasterly part of the line, where possible truncation of sequence 4 occurs. Sequence 5 is characterized by distinct, high-frequency shingling of northwestwardly prograding units. Well V (cf. log in Fig. 14) is projected onto this section and provides age and lithologic control.
cies of sandstone and shale with minor conglomerate (Fig. 14). The sand is predominantly grayish white, very fine-grained to medium-grained, fair to poorly sorted, and contains minor amounts of carbonaceous material. The shale is gray, commonly sericitic, and locally carbonaceous. Unit III in well B consists of an interbedded conglomerate-shale interval which exhibits several 9- to 12-m-thick prograding units (Fig. 14). At wells C1 and C2, unit III consists of interbedded sandstone and shale. The sandstone is fine- to coarse-grained, generally poorly sorted, rarely pebbly, and contains subangular to subrounded grains. The sand contains
varying amounts of white, metamorphic vein quartz and disseminated carbonaceous material. At well V, unit III consists of thinly bedded sandstone and shale. The sand is medium-grained with rare coarse grains and pebbles. Geologic interpretation of sequence 3 in a north-south direction In the seismic section of Fig. 12, sequence 3 contains a southward-prograding conglomerate facies to the north adjacent to the E1 Pilar strike-slip fault. This unit would correspond to the conglomeratic facies of units I and II in well B that is closest to the
518
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Fig. 14. North-south stratigraphic correlation between electric logs of wells V and B showing the correlation of sequences 3, 4 and 5. The datum is the base of the Durham sand at the base of sequence 5. The lithologies of units II, III, IV, and V are described in detail in the text. SB denotes sequence boundary.
fault, and is interpreted as a fanglomerate deposited within a in a shallow marine to brackish setting. Coarse grain sizes of the conglomerate are attributed to local source areas in the Northern Range and steep topographic slopes formed by movement along the E1 Pilar strike-slip fault zone (Fig. 9). Mounded reflectors at the base of sequence 3 on the line in Fig. 11 suggest the occurrence of gravity flow processes in more distal areas of the E1 Pilar sub-basin. A shallow-water depositional setting is inferred for the reflectors north of well C in Fig. 12 because of their position on the Avocado high (Fig. 9A) and because the type of subparallel to hummocky reflectors
seen south of well C generally indicate sediments prograding into shallow water. Clinoform geometries of reflectors south of well B in Fig. 12 indicate a southward increase in water depth. A general increase in the sorting of sand in units II and III of sequence 3 suggests a southward increase in the degree of marine reworking southward along this inferred north-to-south paleoslope. Clinoform geometries in sequence 3 south of well C (Fig. 12) and south of well V (Fig. 13) indicate a similar pattern of southward increase in water depth and related increase in the amount of marine reworking.
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14 km) beneath the Curacao Ridge. Given the eastern migration of the Caribbean Plate with respect to the South American Plate, during the Late Cretaceous and early Tertiary the depocenter would have been farther west, located roughly adjacent to the northern extension of Lake Maracaibo (see Fig. 12). The contour current appears to flow along and parallel to the 4000 m isobath. Current activity began in the Middle Eocene, above reflector A", and the current intensity started to wane toward the Early-Middle Miocene.
in Early Miocene time; the only major river system still disgorging sediment from South America to the Caribbean Plate is the Magdalena River. Note that, like the previous drainage system, the Magdalena River is an axial-parallel river system with one branch being sourced from the basin between the Western and Central Cordillera and the other sourced from the basin separating the Central from the Eastern Cordillera (Hoorn et al., 1995; Fig. 14A). This change in sediment source region is consistent with the inferred sediment provenance change from the Paleocene-Eocene and the Miocene and Pleistocene recorded at ODP Site 1001 (Sigurdsson et al., 1997). Onshore sediment provenance studies indicate that the early Tertiary sediments sampled in Barbados and Venezuela were sourced from the Cordillera system and Guyana Shield (Kasper and Lame, 1986; Baldwin et al., 1986). Furthermore, fission-track analysis of these samples yields age ranges from 80 to 15 Ma, which is consistent with the erosion of the high Cordillera mountain system beginning in the Late Cretaceous (Baldwin et al., 1986). As a result of the northward damming and/or diverting of the rivers and the renewed Oligocene
and Miocene Andean tectonic activity (Fig. 13), sufficient sediment was supplied to the developing foreland basin to infill the paleo-structural lows east of the north-south-trending foreland basin (e.g., the Marfijo and Solim6es basins; Fig. 14). As a consequence of this regrading, fluvial systems flowed to the east across the South American continent and delivered sediments to the Atlantic Ocean that led to the development of the Amazon delta and fan. On the basis of limited biostratigraphy, it is postulated that most of the Amazon fan was deposited after the Middle Miocene in response to Andean uplift (Damuth and Flood, 1984). Nevertheless, because the Andes were high-standing regions since the Late Cretaceous, the initiation of the Amazon fan is not simply a consequence of Middle Miocene uplift. We propose that two important events were responsible for modifying the drainage of South America and forming the presentday Amazon and Orinoco drainage systems. First, and most important, incipient uplift and deformation of the northern Andes (Central Cordillera, Eastern Cordillera, Santander Massif, Santa Marta Massif, Sierra de Perija), Venezuelan Andes (Mdrida An-
622 des), and the Venezuelan Caribbean Mountains in Late Oligocene to Early Miocene time dammed or diverted the existing axial-parallel drainage system of the Maracaibo-Peruvian foreland basin (Figs. 1214A). Either damming or increasing the length of the fluvial system would raise base level and cause deposition along the upper reaches of the fluvial profile (i.e., the equilibrium profile; Schumm, 1991). Renewed tectonic activity in the Central and Eastern Cordillera caused additional uplift, but also increased tectonic subsidence in the foreland basin east of the Cordillera due to flexure loading of the footwall block. Second, through time, the evolving drainage systems infilled the paleo-structural lows because the axial-parallel drainage systems were forced east by the tectonic dam to the north. Upon infilling the structural lows and regrading of the fluvial profiles by erosion, deposition, and isostasy, the fluvial systems were able to bypass the structural barriers across the Marfijo and Solim6es basins and deliver sediments to the Atlantic Ocean. The dramatic change in sediment accumulation and distribution that we observe in the MCS data from the Venezuelan Basin indicates that axial-parallel drainage networks delivered sediment from the Late Cretaceous to Early Miocene. Our results are consistent with early Tertiary paleo-flow directions based on sediment provenance studies in Barbados and Venezuela (Kasper and Lame, 1986). These studies indicated that the major drainage of the South American Craton was towards the north in the Late Cretaceous and early Tertiary. The present Orinoco and Amazon drainage systems developed as a result of tectonic deformation along the northern South American margin (e.g., Venezuelan Caribbean Mountains; Fig. 14). Even though the Late Cretaceous to Early Miocene sediments in the Venezuelan Basin appear to be derived from South America, the deposition was predominantly by gravity flows prior to the Middle Eocene and by current-control after the Middle Eocene (i.e., above and below A"). A change in climatic conditions during the Late Eocene and the onset of abyssal currents, together with the acoustic character of the sediments, suggest that the deposition of the A" to eM sequence was controlled by abyssal current circulation (Figs. 4 and 6; McCave and Tucholke, 1986; Westall et al., 1993; Driscoll and Laine, 1996).
Late Eocene to Early Miocene current-controlled deposits Similar to the B" to A" succession, the sediment drift sequence above A" thickens dramatically toward the southeast displaying a pronounced increase in thickness across the rough-smooth boundary (Fig. 4). The depocenter of this deposit is lo-
N.W. DRISCOLL and J.B. DIEBOLD calized within the central Venezuelan Basin. The hemipelagic-pelagic sediments which comprise the post-A" sequence thin markedly towards the northwest away from the depocenter predominantly by onlap, suggesting that the hiatus results from nondeposition rather than erosion. However, near DSDP Site 150 (Fig. 8), the sedimentary section thins by truncation suggesting that, in part, the hiatus also is due to erosion. Farther north away from the abyssal current axis, DSDP Site 146 sampled a more complete sedimentary section (Fig. 9). We attribute this pattern of thickening, thinning, and truncation to the flow of geostrophic bottom currents around the Venezuelan Basin. Geostrophic flows usually parallel the contours because the pressure gradient is balanced by the coriolis force. The core of bottom water flow appears to have been concentrated near the 4000 m isobath and given the topography and the northern latitude, the circulation pattern is counterclockwise. The onset of this current-controlled deposition above reflector A" began in the Middle Eocene and waned in the Early Miocene (~20 Ma; Fig. 13). This resulted in more subdued current-controlled features in post-Early Miocene time. Even though current intensity has waned the circulation pattern appears to have been maintained to the present (Fig. 4). Analysis of 3.5 kHz echograms suggests that the finer portions of the turbidites entering the Venezuelan abyssal plain in recent times are entrained by the bottom currents and deposited in blankets of sediment waves mantling the underlying deposits. Given that the acoustic character and distribution of the A" to eM is diagnostic of current-controlled deposition, the cogent question becomes, what is the source for the bottom water and through which gateways did it gain access to the Caribbean. There are no present-day gateways through the northern and eastern Caribbean structural barrier (Greater and Lesser Antilles) that are deep enough to allow lower North Atlantic Deep Water (NADW) and/or Antarctic Bottom Water (AABW) access to the Caribbean. The sill depth of the deepest gateway along the Antilles, the Anegada Passage, is 1960 m (Worthington, 1966) and it only allows Antarctic Intermediate Water (AAIW) and upper NADW access to the Caribbean (Worthington and Wright, 1970; Wright and Worthington, 1970; Wunsch and Grant, 1982). Even so, the configuration of the eastern Caribbean structural barrier might have been quite different during the Middle Eocene to Early Miocene. Towards understanding the paleo-gateways to the Caribbean, Donnelly (1990) analyzed pelagic sediments collected by DSDP drilling in the Atlantic, Caribbean, and Pacific. Geochemical analysis of the pelagic sediments, used to infer intermediate and bottom water chemistry, indicated that a barrier
TECTONIC AND STRATIGRAPHIC DEVELOPMENT OF THE EASTERN CARIBBEAN between the Caribbean and western North Atlantic was established at least by Middle to Late Eocene time (,~40 Ma). If the Caribbean intermediate and deep waters were in communication with the western Atlantic, then the excess silica in the sediments of both basins should vary in concert. However, the relatively precipitous Late Eocene disappearance of excess silica in the western Atlantic was not observed in either the Caribbean or Pacific samples. Siliceous sedimentation in the Caribbean continued until Early Miocene time suggesting that a structural barrier isolating the Caribbean from the Atlantic existed since at ~40 Ma. The barrier is presumed to be the Aves Ridge, or the Lesser Antilles, or both. The reduction of excess silica in Caribbean sediments in the Early Miocene is interpreted to record the gradual shoaling of the Panamanian Isthmus (Donnelly, 1990). Consequently, we propose that the gateway for the deep Caribbean circulation is the Panamanian Isthmus between Central and South America, which allowed abyssal water (AABW??) in the eastern Pacific access into the Caribbean. The gradual shoaling of the Central American Isthmus closed the gateway in Late Oligocene-Early Miocene time, consistent with the diminished occurrence of current-controlled features since the Early Miocene observed in the seismic sections (i.e., the more uniform thickness off the sequence overlying horizon eM). We interpret the truncation and non-deposition between A" and eM horizons observed in the seismic reflection data to record the counterclockwise flow of bottom currents from the eastern Pacific around the Venezuelan Basin (Figs. 4 and 6-8). A core of bottom water flow appears to be concentrated near the present-day 4000 m isobath (Figs. 13 and 15). This ribbon of flow would explain the erosion and non-deposition observed along line 1320 (Fig. 8). DSDP Site 150 is located in an area where the maximum hiatus occurs due to erosion and non-deposition. Farther north, DSDP Site 146 (Fig. 9) sampled the sedimentary section farther away from the ribbon of swift current and recovered a more complete sedimentary section. The onset of current-controlled deposition above horizon A" began in the Middle Eocene and waned in the Early Miocene (~-,20 Ma). Even though the dominant current-controlled features diminished during the Early Miocene, current-controlled deposition has continued to the Holocene. The present-day circulation could be the result of AAIW or upper NADW entering the basin through the Anegada and Cayman passages and causing a sluggish cyclonic circulation around the basin (~5 cm/s). An alternative hypothesis, is that Guyana surface current entering the Caribbean and flowing across the Caribbean into the Gulf of Mexico excites a deep water circulation in the Venezuelan Basin (Fig. 15). If this alternative
623
hypothesis is correct, then surface circulation could have been more vigorous when the Panamanian Isthmus was open than it is today. The increased intensity of the surface current might explain the different styles of the current-controlled deposition before and after the Early Miocene. The current-controlled deposit observed in EW9501 seismic data appears to be a detached drift with a morphology similar to the Greater Antilles, Blake, and Eirik drifts (McCave and Tucholke, 1986). Within the A" to eM sediment interval, the laminated character systematically decreases westward away from line 1293 (Figs. 4 and 6). In addition, the acoustic character diminishes slightly toward the east away from line 1293 (Fig. 2). Using the hummocky acoustic character to define the region where the swiftest-flowing ribbons of current occurred, has allowed us to reconstruct the paleo-circulation (Figs. 13 and 15). Given that much of the structuring of the present-day Caribbean Plate had occurred by Eocene time (i.e., Venezuelan Basin, Beata Ridge, and Hess Escarpment) or was well underway (Caribbean-South American Plate boundary), we propose that the increase in hummocky character within A" to eM toward the west records a velocity increase associated with the basin constriction around Beata Ridge (Fig. 15). Likewise, the regions with hummocky acoustic character immediately above eM are structurally shallower and are closer to the axis of the current at the time of deposition (Figs. 3 and 4). Conversely, the regions with well laminated sequences overlying eM are structurally deeper and basinward of the main current axis at the time of deposition (Figs. 6 and 7).
CONCLUSIONS The newly acquired high-resolution MCS data during EW9501 clearly imaged the crustal structure, crustal thickness, and the character of the overlying sedimentary successions. These data coupled with previous collected onshore and offshore data, have allowed us to develop the following model for the geologic evolution of the Venezuelan Basin. (1) The proto-Caribbean crust was formed by seafloor spreading in Late Jurassic-Early Cretaceous time. We have assumed normal oceanic crust thicknesses of approximately 6 km for the protoCaribbean crust because the seismic reflection data indicates that faulting and extension were concomitant with, and subsequent to, the magmatic activity. (2) Prior to the Senonian, widespread and rapid eruption of basaltic flows began in concert with extension and thinning of the 'old' plate. We are not able to determine whether extensional deformation pre-dated or post-dated the onset of magmatic activity.
624 (3) The plate was thickened by at least two if not more magmatic events. The excess thickness was generated by the extrusion of lava flows over a vast area, intrusion of dikes and sills, and underplating by residual mantle from the melting event. The roughsmooth B" boundary imaged in seismic profiles is the edge of the basalt province created by these magmatic events over the older deforming plate. (4) The large divergent volcanic wedges observed along the rough-smooth B" boundary and the abrupt shoaling of Moho appear to be controlled by a northwestward-dipping fault system. The dip of the reflectors that comprise the divergent wedge diminish upsection and resemble the stratal geometry observed in rift basins. In contrast to the common model cited for the formation of these volcanic wedges (seaward-dipping reflectors, SDRs), we propose that the source of the basalt flows is either along strike or toward the northwest, structurally updip, and that they progressively infill the deforming basin. (5) Extension continued after magmatic thickening of the Venezuelan crust as evidenced by faulted and rotated tongues of smooth basement toward the east. The location of the major extensional deformation migrated through time from the Venezuelan Basin to the western flank of the Beata Ridge. The extensional unloading of the footwall caused uplift and rotation of the Beata Ridge and collapse of the hangingwall (i.e., Hess Escarpment). The sediment thickness and stratal geometry of the overlying sedimentary successions across the Venezuelan Basin and Beata Ridge suggest that the majority of the observed deformation in this region occurred soon after the emplacement of the volcanics. Minor fault reactivation in the Neogene along the eastern flank of the Beata Ridge is associated with an accommodation zone (i.e., tear fault) that records a change in the deformation style from subduction of the Caribbean Plate along the Muertos Trough to obduction of the Caribbean Plate along Hispaniola. (6) Coincident with the rough-smooth boundary is a marked increase in the thickness of the A" to B" sediment interval. Correlation with DSDP and ODP sites in the Caribbean constrains the age of this interval to be younger than Senonian (,~88 Ma) and older than Middle Eocene (~50 Ma). We conclude that the Late Cretaceous to Early Miocene sediments observed in the Venezuelan Basin are terrigenous deposits transported northward by an axial-parallel fluvial system draining the MaracaiboPeruvian foreland basin of South America. (7) The initiation of the present-day South American drainage patterns is not simply a consequence of Middle Miocene uplift of the Andes, as the Andes are the result of several tectonic episodes beginning in the Late Cretaceous. We propose that during the Late Cretaceous to Early Miocene the dominant
N.W. DRISCOLL and J.B. DIEBOLD drainage was a north-flowing axial-parallel system that supplied sediment to the Venezuelan Basin. The middle Tertiary uplift and deformation along the northern South America Plate boundary blocked the axial-parallel fluvial networks. As a result of the northward damming of the rivers and the renewed Oligocene and Miocene Andean tectonic activity, abundant sediment was supplied to the developing foreland basin east of the Andes that infilled the paleo-structural lows in the Mar~jo and Solim6es basins. The consequent regrading of the fluvial systems allowed drainage systems to flow east across the South American continent (e.g., Amazon and Orinoco) and deliver sediment to the Atlantic Ocean. (8) Finally, the seismic reflection data also imaged an Eocene-Early Miocene current-controlled drift deposit which might reflect the movement of eastern Pacific bottom currents into the Caribbean during this period. The gradual shoaling of the Central American Isthmus in Late Oligocene-Early Miocene time closed the gateway. Such closure is consistent with the diminished occurrence of current-controlled features since the Early Miocene observed in the seismic sections and Caribbean paleo-deep and intermediate water geochemistry of the sediments.
ACKNOWLEDGEMENTS
We thank the crew and scientific staff of the R/V
Ewing for making cruise EW9501 in the Caribbean such a great success. This manuscript benefited from numerous discussions with Lewis Abrams, Peter Buhl, Thomas Donnelly, Bob Duncan, Edward Laine, Sylvie Leroy, Garry Karner and Elazar Uchupi. James Kellogg, Paul Mann, Walter Pitman, James Pindell, and Elazar Uchupi critically read this manuscript and their comments are greatly appreciated. Garry Karner compiled the topographic data shown in Fig. 14. Support for this research was provided by the National Science Foundation grant OCE-93-02578. Woods Hole Oceanographic Institution contribution #9503. Lamont-Doherty Earth Observatory #5684.
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C h a p t e r 21
Neogene Intraplate Deformation of the Caribbean Plate at the Beata Ridge
A L A I N M A U F F R E T and S Y L V I E L E R O Y
We have studied all the seismic profiles existing in the Beata ridge area, in addition to Seabeam maps, to determine the tectonics of this structure and its relationships with the adjacent areas. The basis of this work is the multichannel Casis seismic survey carried out by the R/V Nadir in 1992. These seismic lines are migrated and clearly show evidence of compression and transpression. The Presqu'~le du Sud d'Hispaniola is an uplifted part of the volcanic igneous province that formed the Caribbean plate during Cretaceous time. This region was initially a part of the thick Beata volcanic plateau that collided with the central part of Hispaniola. A Seabeam map and single-channel seismic lines from the Seacarib 1 cruise show the active collision between the northeastern tip of Beata ridge and the western termination of the Muertos trough. Structural analysis indicates that the Muertos trough is an Eocene feature that has been reactivated in Recent times. The progressive emergence of the Muertos prism and its onland extension results from vertical stacking by thrusting of different parts of the prism. Several compressional structures can be seen on the eastern flank of the Beata ridge. The importance of these structures decreases systematically towards the south. One of these structures, the Taino ridge, was surveyed in detail during the Casis cruise. We describe reverse faults, pop-up and strike-slip faults. These tectonic features are compatible with a NE-SW compressive stress. A detailed site survey of the Aruba Gap was also performed during the Casis cruise. We show again compressional and wrench faults and an increase of the deformation towards the north. In contrast, we found no evidence of any compressional deformation on the western side of the Beata ridge. Here, steep NE-SW scarps shown by Seabeam maps, are predominant. We conclude from this tectonic framework that the Beata ridge has been deformed by compression and strike-slip faulting since the Early Miocene (23 Ma) by a NE-SW-oriented compressive stress. The Beata ridge is progressively uplifted from the south to the north up to the emergence of the Presqu'~le du Sud. However, the Beata ridge is a Cretaceous plateau and initial topography must be taken into account. The Beata ridge is placed in the regional tectonic framework and we show that the compression of the ridge is probably connected to the Sinu subduction zone in Colombia. We distinguish between the Colombian and the Venezuelan microplates separated by the Beata compressional zone. The former drifts towards the northeast faster than the latter. From our structural analysis we deduce 0.9 cm/yr of relative motion between the two plates.
INTRODUCTION The Caribbean plate is a large volcanic province probably formed, during the Cretaceous, in the Pacific O c e a n (Duncan and Hargraves, 1984). This plate is m o v i n g eastwards relative to the North A m e r i c a n ( N O A M ) and South A m e r i c a n ( S O A M ) plates (Pindell and Barrett, 1990). The C a r i b b e a n plate is d e l i m i t e d (Fig. 1) to the north by a left-lateral strike-slip fault zone ( C a y m a n - P u e r t o Rico fault system) and to the south by a c o m p l e x set of rightlateral faults. In addition, the N O A M and S O A M plates are slowly converging with a present pole of rotation located near the M i d - A t l a n t i c ridge (Pindell and Barrett, 1990; Mtiller and Smith, 1993; Mtiller et al., 1996). S e d i m e n t a r y d e f o r m e d belts (Ladd
and Watkins, 1978; L a d d et al., 1981, 1984, 1990) north of South A m e r i c a ( C o l o m b i a n and Venezuelan d e f o r m e d belts, Fig. 1) and south of Puerto Rico (Muertos trench, Fig. 1) m a y result from this c o m p r e s s i o n which increases towards the west. The D S D P results of Leg 15 (Edgar et al., 1973b) and the O D P results of L e g 165 (Scientific Party, Leg 165, 1996) indicated that the volcanic b a s e m e n t of the C a r i b b e a n plate was f o r m e d during a short period of the Cretaceous time (late Turonian to Campanian, 8 8 - 7 4 Ma). The same volcanic rocks outcrop in the Presqu'~le du Sud and B a h o r u c o Peninsula of Hispaniola (Maurasse et al., 1979) and Curaqao Island (Klaver, 1987) and these areas are considered to be uplifted pieces of the C a r i b b e a n plateau. In the central part of the C a r i b b e a n Sea, the Beata ridge,
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 627-669. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
628
A. MAUFFRET and S. LEROY
Fig. 1. Plate tectonic frameworkof the Caribbean plate. The Gonave microplate (Rosencrantz and Mann, 1991) is indicated.
2-4 km deep, has a triangular shape and lies (Fig. 2) between the Colombian basin to the west and the Venezuelan basin to the east. The northern parts of the Colombian and Venezuelan basins have been named the Haiti and Dominican sub-basins, respectively, and we gave new names (from Indian tribes) to the main ridges that composed the Beata ridge. The Bahoruco Peninsula is the northern prolongation of the Beata ridge, whereas this ridge is separated from the deformed margin of the South American plate by the Aruba Gap. Many speculative models have been proposed for the formation of the Beata ridge: normal faulting (Fox et al., 1970; Fox and Heezen, 1975; Holcombe et al., 1990); a buoyant thick oceanic plateau which resists subduction (Burke et al., 1978) or reverse faulting related to a transpressive motion (Vitali, 1985; Mauffret et al., 1994). New multichannel seismic profiles, acquired during the Casis cruise performed in 1992 on the R/V Nadir in the Caribbean Sea (Fig. 2), demonstrate strong transpressive tectonics of a former volcanic plateau (Mauffret et al., 1994; Leroy, 1995; Leroy and Mauffret, 1996). The Beata ridge was studied in the framework of a comprehensive work on the geophysical and geological data of the Caribbean Sea (Leroy, 1995) to promote an ODP Leg on the deep structure of the Caribbean igneous province.
BATHYMETRY OF THE CENTRAL CARIBBEAN BASIN
Since 1980 (Case and Holcombe, 1980) no new bathymetric map of this region has been published. In 1985 a Seabeam survey was performed during the French Seacarib cruise (1985, R/V Charcot). Some data were published (Mercier de Lepinay et al., 1988; Jany et al., 1990; Mauffret and Jany, 1990) but the surveys of the Beata ridge as a whole are presented for the first time in this paper (Fig. 3). These data are integrated in the new map presented here (Fig. 3) that includes a compilation of previous data provided by the Marine Geophysical Data Center (Leroy, 1995). The Muertos trough lies in the northern part of the central Caribbean region between 5 and 4 km deep. This depression is delimited towards the north by a deformed slope. Towards the west the Muertos trough undergoes a prominent bend then disappears near Hispaniola. The Bahoruco Peninsula is bounded towards the east by a steep scarp from the coast to 3 km deep. The deep part of the scarp and the continental rise, below 3 km deep, are interrupted by several seamounts that trend north-south. The southern tip of the Bahoruco Peninsula extends southward by a spur from the coast to 2 km depth. The Beata ridge is bounded towards the northwest by a steep scarp that trends NE-SW. At the foot of
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
629
uertos
Tairona Ridge,
Dominicansub basin
31 TainoRidge DSDP 151 153 Warao Rise
ta Platea~''~
Colombia Basin
Fig. 2. The northern parts of the Colombian and Venezuelan basins are renamed the Haiti and Dominican sub-basins, respectively. The Warao rise is a buried feature that bounds the Haiti sub-basin to the south. The backbone of the Beata ridge is formed by the DSDP 151 and Tairona ridges. Warao, Taino and Tairona are the names of Indian tribes. The position of the DSDP Sites (Legs 4 and 13) and ODP Sites (Leg 165) are indicated. The seismic tracks of the Casis cruise is also shown.
the scarp lies the Haiti sub-basin that is surrounded by the 4.2 km bathymetric contour. This basin is delimited towards the west by the Hess escarpment and towards the north by the steep continental slope of Haiti and the Haiti plateau that trends N W - S E . The eastern boundary of the Beata ridge with the Venezuelan basin is subdued, but two prominent features, the Taino ridge 3 to 4 km deep, and the Beata plateau surrounded by the 4-km-bathymetric contour outline the southeastern limit of the ridge. The central part of the Beata ridge, delineated by the 3-km-bathymetric contour, is formed by two prominent features: Tairona and DSDP 151 ridges that trend north-south. The Aruba Gap, 4 km deep, is located between the Beata ridge and the South
American deformed belt and forms a sill between the Colombian and Venezuelan basins.
SEISMIC CONTROLS In addition to the Casis cruise we examined all the MCS (multichannel seismic profiles) and SCS (single-channel seismic profiles): from the University of Texas at Austin, Shell, Institut Franqais du Pdtrole for the MCS profiles; from the Lamont Doherty Earth Sciences Observatory (Vema, R / V Conrad data and the monitors of the Ewing 9501 cruise); Texas A & M (Alamino cruise) and the Seacarib 1 cruise (SCS profiles) (Fig. 4). All these
630
A. MAUFFRET and S. LEROY
Fig. 3. New bathymetric map of the central part of the Caribbean Sea (from Leroy, 1995, modified). The Seacarib 1 Seabeam surveys (position indicated) were incorporated to the conventionalbathymetric data base (NGDC).
profiles were digitized and converted in depth with a velocity curve derived from DSDP results and velocity analysis of the Casis data.
SEISMIC STRATIGRAPHY
The seismic stratigraphy of the central Venezuelan basin was defined in the early work of Ladd and Watkins (1980). The DSDP results reveal the presence of four main seismic intervals (Fig. 5). An upper seismic interval in Hole 31 (Bader et al., 1970) and Holes 146 and 153 (Edgar et al., 1973b), corresponds to lithic unit 1, a chalk marl ooze and clay of Early Miocene to Recent age (eM, 23 Ma; Fig. 5). The second seismic interval corresponds to a Middle Eocene to Early Miocene radiolarian chalk and radiolarian chalk unit. Horizon A" (Fig. 5) forms the
base of this interval. The third seismic interval corresponds to lithified chalks, cherts, limestones and black shales which rest upon Santonian to Coniacian basalts. Horizon B" correlates with these basalts. In addition, we have identified in the Aruba Gap area a lower sedimentary unit between an equivalent of B" and a deeper horizon (V, Fig. 5B). In the volcanic crust several reflectors have been identified (sub-B" R reflectors and Moho, Fig. 5A), but a study of the deep crust is beyond the scope of this paper (Mauffret and Leroy, 1997). The upper seismic interval has a variable thickness, relatively thin in the flank of the Beata ridge, and increasing to 2 km thick in the southern part of the Venezuelan and Colombian basins where it fills a trench related to the South American deformed belt (Talwani et al., 1977; Biju Duval et al., 1982b). In the Colombian basin, the layer of Early Miocene to
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
631
Fig. 4. Seismic tracks used in this study.
recent sediments is thick in relation with the Magdalena deep-sea fan (Kolla et al., 1984). The DSDP Sites and the seismic profiles indicate that this interval consists of turbiditic sediments in basins, but a pelagic composition is inferred on the Beata ridge. The second seismic interval corresponds to a pelagic unit. The upper part, and in some places the entire interval (Fig. 5B), is chaotic. This layer is current-controlled, as confirmed by the presence
of several hiatuses in the DSDP holes that have been related to strong Early Miocene currents which were active as the Caribbean Sea opened towards the Pacific Ocean (Edgar et al., 1973a; Holcombe and Moore, 1977). The hummocky aspect of the Early Miocene reflector is widespread and the unit has been described as a prominent horizon by Houtz and Ludwig (1977). The Early Miocene to Middle Eocene seismic interval dips toward the South Amer-
Fig. 5. Seismic stratigraphy tied with the DSDP results. Except for (B) all the seismic profiles cross exactly the DSDP Site area. The main reflectors are Early Miocene (eM), Eocene (A ~I) and Santonian to Coniacian (B") in age, respectively. Note the thickening of the A"-B 'I interval from west (A, B) to east (C, D). The deep reflectors (V, sub-B", R and Moho) are described in another paper (Mauffret and Leroy, 1997). The location of profiles is indicated in inset.
t7z
-]
>,
t,~
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE ican deformed belt and this disposition indicates that compressional deformation along South America occurred since the Early Miocene epoch. At this time the floor of the Caribbean Sea was completely deformed, with a general tilting towards the south related to the formation of trenches (Biju Duval et al., 1982b). The structure of the young accretionary prism indicates a large amount of offscraping (Ladd et al., 1984) and tomography data indicate a long slab extending beneath northwest South America (Hilst and Mann, 1994). However, the south dip of the Venezuelan basin crust is partly the result of original construction of the Cretaceous volcanic plateau that thins from north to south (Diebold et al., 1999). The current-controlled layer cannot be identified on the top of the Beata ridge and the transparent facies of the layer overlying the top of the ridge suggests a pelagic environment (Fig. 5E). The A"-B" interval is evident in the Venezuelan basin and in the western part of the Colombian basin (Ladd and Watkins, 1980; Bowland, 1993). However, horizon A", which correlates with Middle Eocene chert, is not clearly identified in the eastern Colombian basin and on the Beata ridge (Fig. 5E). The A " - B " interval is thick in the Venezuelan (Fig. 5D) and Colombian basins and also on the eastern flank of the Beata ridge (Fig. 5C), but thin-
633
ner in the Aruba Gap area where it was penetrated during drilling at DSDP Site 153 (Fig. 5A and B). A prominent hard ground is correlated with a late Maastrichtian hiatus at the DSDP Site 153. A similar hard ground was described at DSDP Site 151, but here the Paleocene directly overlies the Santonian sediments and basalts (Edgar et al., 1973b). In a basin located to the north of the DSDP Site 153, the A " - B " interval correlates in thickness and seismic facies with units encountered at the DSDP Site 153 (Fig. 5B). However, the acoustic basement (V, Fig. 5) is very different from the typical smooth B" reflector drilled at DSDP Site 153, and is overlain by a sedimentary layer with a compressional velocity of 3.9 k m / s (Fig. 5B).
NORTHERN BOUNDARY OF THE CENTRAL CARIBBEAN REGION
The present-day northern limit of the Caribbean plate is the Oriente fault Fig. 1), the extension of the Oriente fault in Hispaniola that lies in the Cibao valley (Fig. 6; Calais et al., 1992; Russo and Villasenor, 1995) and the Puerto Rico trench (Masson and Scanlon, 1991). A strain partitioning may occur and a component of compression may be absorbed in the
Fig. 6. Tectonic framework of Hispaniola. The main boundary between the Caribbean plate and the North American plate is located along the Oriente fault and the Cibao valley. However, a strain partitioning with a compressional boundary is possible north of Hispaniola. The central part of Hispaniola is occupied by island arc crust (Central Cordillera), deformed sedimentary belts (Peralta and Neiba) and basins (San Juan and Enriquillo). The Enriquillo basin bounds the Presqu'~le du Sud that is formed by Cretaceous volcanic rocks identical to that of the Caribbean basement (Maurasse et al., 1979). The Presqu'~le du Sud is split into two parts by the left-lateral strike-slip Enriquillo fault. This fault is related to the Navassa pull-apart basin (Mann et al., 1995). The southern offshore part of the Presqu'~le du Sud is severely deformed by compression and transpression (Bien-Aime Momplaisir, 1986). The Haiti sub-basin has probably a quasi oceanic crust. East of the Beata ridge the Seacarib Seabeam survey is indicated. The Muertos prism can be divided into three parts by recent reverse faults. The location of Figs. 7 and 8 is shown.
Fig. 7. The Haiti sub-basin has a thin crust as shown by the refraction data (refraction line 36W, Ewing et al., 1960). The sedimentary layers onlap a wedge at the base of the slope indicating old deformation. The bottom of the wedge is indicated by the white line and arrows (flat reflector). In contrast the upper slope is actively deformed with the Ile-a-Vache anticline and a deep syncline bounded by reverse faults (Bien-Aime Momplaisir, 1986). The location of the seismic profile is indicated in Fig. 6.
t-
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NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE sedimentary deformed belt that lies north of Hispaniola (Fig. 6; Austin, 1984; Dillon et al., 1992). The Caribbean plate boundary has been located along the north of Hispaniola since the Miocene and before this boundary was probably placed along the Oriente fault, the Central Cordillera and Muertos trough (Mercier de Lepinay, 1987; Pindell and Barrett, 1990). Hispaniola can be divided into four main blocks. A Septentrional block separated from the Eastern and Central Cordilleras by the Cibao valley. These cordilleras underwent a Mesozoic island arc tectonism (Mercier de Lepinay, 1987; Lebron and Perfit, 1994). A central area, including the Peralta flysch, the San Juan basin, the Sierra de Neiba that is an accretionary prism since the Eocene up to the Pliocene (Mann and Lawrence, 1991; Mann et al., 1991a,b, 1995). This third block is delimited by the Enriquillo depression (Fig. 6). The southern block is formed by the Presqu'ile du Sud d'Haiti that is an uplifted portion of the Cretaceous Caribbean igneous province (Maurasse et al., 1979). This block collided with the northern block recently (Mercier de Lepinay et al., 1988). The Early Pliocene strata are folded in the western part of the Enriquillo basin and the collisional process is probably presently active Vila et al., 1990). The Presqu'ile du Sud is cut by the Enriquillo fault that extends towards the west to Jamaica and the Cayman spreading center (Sykes et al., 1982; Rosencrantz and Mann, 1991). The Navassa trough (Fig. 6) is a pull-apart basin located along the left-lateral Enriquillo fault (Mann et al., 1995). The margins of the Presqu'ile du Sud were studied in detail by Bien-Aime Momplaisir (1986) and she showed that the northern and southern margins are affected by a compressional or/and transpressional tectonics governed by a N40-45 compressive stress. The Presqu'*le du Sud can be assimilated to a giant positive flower structure. A seismic profile (Fig. 7) shows the deformation of the upper margin with a broad syncline and a thrust of the sedimentary cover of this basin on the Tle-~-Vache structure. A narrow anticline is located at the top of this feature. At the base of the steep slope lies the Haiti sub-basin. The correlation of the seismic profile with refraction results (Ewing et al., 1960) indicates that this basin has a thin oceanic crust. The Haiti plateau is a thick block related to the Cretaceous igneous province, whereas the thin crust of the Haiti sub-basin seems to be not affected by the Cretaceous volcanic event. However, the presence of some intra basement reflectors (sub-B"?; Fig. 7) may suggest a weak volcanic contamination of the thin crust in the Haiti sub-basin. A small sedimentary wedge lies above a flat reflector (d6collement?) at the base of the slope. However, the sedimentary layers of the basin onlap this wedge. Consequently
635
this wedge was formed during old compressional tectonics and the present deformation is restricted to the upper margin. Nevertheless the sedimentary wedge may also be formed by slope breccias at the base of the scarp and the chaotic aspect of the upper sedimentary layers suggests slumped sediments and erosional products. The presence of dipping reflectors (Fig. 8) into the acoustic basement of the Beata plateau indicates the volcanic formation of this feature. The deepest sedimentary unit onlaps the basement and this seismic configuration suggests that the basement relief is old. This basement and the lower sedimentary layers are tilted towards the NNE, whereas the thickness of the recent layers increases in the same direction (Fig. 8). A wedge of deformed sediments is evident at the base of the western slope of the Sierra de Bahoruco. Thickening of the recent sedimentary layers and tilting of the lower layer in the basin, reverse fault and d6collement suggest a compressional origin for the wedge. This compression is recent but probably inactive at the present day as shown by other seismic profiles crossing the southern part of the deformed wedge (see later). Consequently the Haiti plateau and the Haiti sub-basin had an eastwards motion relative to the Presqu'ile du Sud, but this motion is presently nonexistent. The Muertos trough has been described several times (Ladd and Watkins, 1978; Biju Duval et al., 1982a; Ladd et al., 1981, 1990). The seismicity shows a steep Benioff zone dipping to 125 km depth (Bryne et al., 1985; Russo and Villasenor, 1997). A Seabeam survey (Mercier de Lepinay et al., 1988; Ja W, 1989; Mauffret and Jany, 1990; Vila et al., 1990) of the western part of the Muertos trough shows the relationships between the front of the accretionary prism and the Sierra de Martin Garcia (Fig. 9A). However, this front is connected to the Peralta flysch belt in some publications (Mann and Lawrence, 1991; Mann et al., 199 la,b, 1995). In the Ocoa Bay (Fig. 9B) the interpreted boundary of the Muertos accretionary prism is located between a (piggyback?) basin and a deformed zone (written communication of S.R. Lawrence, co-author of Chapter 12) that in fact is located in the inner part of the Muertos prism. Moreover the 50-km contour of the Muertos Benioff zone is located near Ocoa Bay (Russo and Villasenor, 1997). The Muertos accretionary prism is divided into three parts: the San Pedro basin, an upper prism and a lower prism (Fig. 9A). Two out of sequence thrust zones separate the three units. A detailed study (Biju Duval et al., 1982a) of the San Cristobal basin, an onshore extension of the San Pedro basin, indicates that the Muertos trough is active since the Eocene, but the uplift of the San Cristobal basin is related to a Late Miocene-Recent compressional event. The Eocene
636
A. M A U F F R E T and S. L E R O Y
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NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE paleoprism (Peralta belt and Sierra de Neiba) has been shortened and uplifted by this recent event that is related to the collision of the Beata ridge with the northern block of Hispaniola (Biju Duval et al., 1982a). The uplift is presently occurring by the thrusting of the San Pedro basin over the upper prism that is underthrust by the lower prism (Fig. 10A, Biju Duval et al., 1982a). North of 18~ the lower prism disappears and the upper prism is directly adjacent to the undeformed basin (Figs. 9A, 10C). The slope off Bahoruco Peninsula is very steep from the coast to 2 km depth; then several hills, that trend northsouth, form an extension of the Beata ridge. The magnetic map contoured during the Seacarib cruise (Fig. 9C) shows that some of these hills are made of volcanic rocks. These seamounts are separated from the prism by the narrow Muertos trough (Fig. 10A, B). However, the Beata seamounts are never in contact with the prism and the easternmost seamount disappears abruptly (Fig. 10C). On the other hand the magnetic map (Fig. 9C) shows that N E - S W magnetic lineations of the Caribbean basement are visible beneath the prism and a positive anomaly (431 nanoteslas, Fig. 9C) may be correlated with a buried seamount. The northern tip of the Beata ridge may be cut by a right-lateral strike-slip fault (Fig. 9C). The Enriquillo basin that is the onshore extension of the Muertos trough is a ramp valley (Mann et al., 1991 a,b) bounded by two facing thrusts (Fig. 6). It is probable that the Enriquillo depression is floored by the Cretaceous volcanic basement of the Presqu'~le du Sud because volcanic Cretaceous basalts have been observed northwest of the Enriquillo basin (Pierre Payen anticline, Fig. 6; (Vila et al., 1988). Moreover, the magnetic map (Fig. 9C) shows the Caribbean basement below the Muertos prism. Consequently, the Cretaceous volcanic basement may collapse by normal faulting produced by flexural effects found in the bulge related to underthrusting and this basement finally subducts beneath the Neiba and Muertos prisms. On land these normal faults have been recently described along the northeastern flank of the Sierra de Bahoruco (Pubellier et al., 1999). The Seacarib profiles (Fig. 10) are not migrated and we cannot define if the faults that bound the Beata seamounts are reverse or normal. However, the southwards extension of these faults is clearly reverse (see next paragraph). In conclusion, the northern boundary of the Presqu'~le du SudBeata ridge may be a reverse fault, but the subduction of this block implies a final normal faulting. These normal faults and the strike-slip faults transfer large fragments of the Cretaceous volcanic plateau to the Gonave microplate (Fig. 1). This microplate was defined between the Cayman spreading center, the Oriente fault, the Plantain Garden-Enriquillo fault and the western coast of Hispaniola (Rosen-
637
crantz and Mann, 1991; Mann et al., 1995). The Enriquillo depression connects this microplate to the Dominican sub-basin. We propose that the Gonave microplate belongs to the Venezuelan plate separated from the Colombian plate by the Enriquillo fault and the Beata fault system (see later).
EASTERN BOUNDARY OF THE BEATA RIDGE
South of the Muertos-Beata collision zone previously described several hills trend north-south (Fig. 9A). On a seismic profile (Fig. 11A), already published in Ladd et al. (1981), the two western seamounts are conical but the third seamount is asymmetrical. The thinning of the A"-B" interval (old sedimentary fill, Fig. 11) away from structures suggest that these features are contemporaneous with the formation of the volcanic plateau. The Beata ridge has a thick crust that has been underplated (Mauffret and Leroy, 1997). This underplating induced an uplift and a coeval rifting (Diebold et al., 1999; Driscoll and Diebold, 1999). However, the sedimentary cover of the central basin is upturned and is now on the top of the structure (Fig. l lA). This configuration is abnormal for a tilted block resulting from a normal faulting. The configuration of the sedimentary layers suggests an old feature reactivated and recently uplifted (1 km, Fig. 11A). Although this profile is not migrated we suggest that a transpression is the cause of the uplift. The structure previously described forms a north-south ridge that is offset to the south (Fig. 11D). In the second seismic profile (Fig. 11B) an asymmetrical structure separates two regions of the Dominican sub-basin with a different seismic stratigraphy. The western region shows several sub-B" reflectors, whereas the basement of the eastern region is void of these reflections. Such a prominent contrast between the eastern and the western part suggests a lateral displacement. The western region seems to be transported from the south where the sub-B" reflections are also prominent (Fig. 11C), implying a right-lateral component of a transpressional fault. The vertical uplift of the fault is estimated to be 0.6 km. In addition this profile shows an inversion of basin related to a compressional or transpressional event. The third profile is a regional seismic section (Ewing 9501 cruise, profile 1321; Fig. 11C). A pop-up with a probable reverse fault dipping towards the NNW is identified. Moreover, westwards steep dipping reflectors can be seen in the crust. Although the vertical throw of the pop-up structure is only 0.25 km the probable reverse fault affects the whole crust. These transpressional features were also seen on several single-channel seismic lines. In conclusion, we identify a transpressional zone with a strong component
Fig. 9. Relationship between the Muertos prism and the geologic features of Hispaniola. (A) Seacarib Seabeam survey completed by conventional bathymetric data. The lower prism of Muertos is formed by elongated anticlines. The toe of the prism turns abruptly towards the north when the first seamounts of the Beata ridge appear and the Muertos trough is very narrow. Then the lower prism disappears as well as the Beata seamounts. The slope of the prism is steep and the toe is related to the recent deformed Sierra de Martin Garcia. The location of the seismic profiles illustrated in Figs. 10 and 11A is shown. (B) The toe of the Muertos prism cannot be located in the Ocoa Bay as presumed by Mann et al. (1991a,b, 1995), Mann and Lawrence (1991). The Muertos prism and its forearc (San Pedro and San Cristobal) were formed during the Eocene (Biju Duval et al., 1982a) but were reactivated recently by the collision of the Beata ridge. The San Pedro basin overthrusts the upper prism and the lower prism disappears beneath the upper prism. This successive stacking generates the uplift and the emergence of the prism. (C) Magnetic map performed during the Seacarib 1 cruise. The magnetic basement of the Caribbean basin can be seen beneath the Muertos prism. The southern Beata seamount shows a correlation with a positive magnetic anomaly, but the northern seamount is orthogonal to the magnetic grain. We cannot exclude a rotation of this seamount that is parallel to the Muertos trough (see A). A prominent magnetic anomaly (431 nanoteslas) is located below a reentrant of the prism. This anomaly maybe correlated with a seamount that subducts. A strike-slip fault may transport the tip of the Beata ridge beneath the Muertos prism.
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
639
Fig. 10. Single-channel seismic profile from the Seacarib 1 cruise. (A) The upper prism overthrusts the lower prism, the Muertos trough is narrow and deformed and is bounded by a Beata seamount. (B) The folds of the lower prism are particularly evident. (C) The lower prism disappears as well as the easternmost Beata seamount. This seamount may have subducted; see the magnetic map in Fig. 9C. The locations of the seismic profiles are indicated in Fig. 9A).
of compression. This transpression is presently active. The height of the d e f o r m e d features increases towards the north where the collision of the Beata ridge is in process.
BEATA PLATEAU
This 0 . 8 - k m - h i g h structure, n a m e d by Case and H o l c o m b e (1980), trends N N W - S S E and is isolated
640
A. MAUFFRET and S. LEROY
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
641
Fig. 12. Seismic profile crossing the Beata Plateau. The Eocene A" horizon has been piston-cored (Talwani et al., 1966; Edgar et al., 1971). Although this profile is unmigrated the reverse motion of the faults is clear. See the inset map for profile location.
f r o m the B e a t a ridge (Fig. 12B). A s e i s m i c profile shows two steps that are p r o b a b l y reverse. H o r i z o n A" is e x p o s e d on the h i g h e s t scarp. E a r l y to Middle E o c e n e was i n d e e d piston c o r e d on this scarp (Talwani et al., 1966; E d g a r et al., 1971). 11 k m of s h o r t e n i n g (20%) is e s t i m a t e d for this 4 8 - k m - w i d e structure if w e a s s u m e that the B" reflector was initially flat.
TAINO RIDGE
T h e Taino ridge is l o c a t e d on the e a s t e r n flank of the B e a t a ridge (Fig. 13). T h e D S D P Site 31 ( B a d e r et al., 1970) allows us to calibrate the s e i s m i c profiles and w e c o r r e l a t e the M i o c e n e - O l i g o c e n e b o u n d a r y with a p r o m i n e n t reflector (eM, 23 Ma; Fig. 5C and Fig. 14). A d e t a i l e d survey was perf o r m e d d u r i n g the Casis cruise. T h e s e i s m i c profiles
Fig. 11. (A) This seismic profile, already published by Ladd et al. (1981), shows two conical seamounts and a third western seamount that is asymmetrical. The sedimentary cover of the adjacent basin are tilted and uplifted on the top of the seamount that indicates a transpression. (B) UTIG (University of Texas at Austin) profile cruise courteously provided by J. Austin. The two basins separated by an asymmetrical seamount show a different seismic stratigraphy. The eastern basin is underlain by a basement with several internal reflections (sub-B"). If we compare this profile with the seismic profile illustrated in (C), this basin may be transported from the south along a right-lateral strike-slip fault. (C) Seismic profile shot during the Ewing 9501 cruise courteously provided by J. Diebold. The complete regional seismic profile is also shown in Driscoll and Diebold (1999). Pop-up related to a transpression. Observe the steep dipping reflector that may correspond to a reverse fault at a crustal scale. The Moho reflection is shown at 10 s two-way travel time. The thinning of the old sedimentary fill away from the structures (A, B) indicates that these features were formed during the construction of the Cretaceous volcanic plateau, but a recent reactivation is evident. The uplift of the hills can be estimated to be 0.25 km in the south and 1 km in the north. See the inset map (D) for profile location. This depth to basement (B") map shows the eastern flank of the Beata ridge. The north-south trend of the structures is evidenced by the seismic profiles presented and several other multichannel and single-channel seismic lines. These structures are offset in a dextral sense between the seismic profiles presented in (B) and (A), respectively.
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Fig. 13. Depth to basement (Bf~) in the Taino ridge area. The Casis survey and the Figs. 14-20 are indicated. Reverse faults, pop-up and strike-slip faults are compatible with a NE-SW compressive stress. DSDP Site 31 results (Bader et al., 1970) help to calibrate the seismic profiles. previously presented were unmigrated, whereas all the Casis profiles are migrated. The southern part of the Taino ridge is relatively large (Fig. 13). It is delimited on the western and eastern side by reverse faults with opposite vergence. On the western flank a clear reverse fault (see the seismic trace, Fig. 14) offsets the B fl, A I~ Early Miocene reflectors and the sea floor. The identical thickness west and east of
the reverse fault indicates recent activity of this fault. The throw of this fault is evaluated to 0.6 km. The second part of the Casis B01 (Fig. 15) seismic profile shows a flower structure and reverse faults with an eastern vergence that bound the Taino ridge to the east. The strike-slip fault probably delimits to the north the large plateau that forms the southern part of the Taino ridge (Fig. 13). A new flower structure
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
643
Fig. 14. Reverse fault in the southern part of the Taino ridge. The trace of the fault is visible on the profile. The throw of the fault is estimated to be 0.6 km. See the inset map (depth to basement: B") and Fig. 13 for profile location.
is identified on the next profile (Fig. 16). This structure is visible very deep in the crust and offsets a 10-km-deep horizon (R, Fig. 16). We propose that this major strike-slip fault connects the Taino ridge to the transpressive structure described in the Dominican sub-basin (Fig. 11). To the north the Taino ridge is narrow and is divided into two hills (Fig. 17) that are offset (Fig. 13). The transverse valley is formed by a strike-slip fault (Fig. 5C). The northern part of the Taino ridge is a typical pop-up delimited by reverse faults with opposed vergence (Fig. 18A). A 15-km-deep ddcollement merges with the western fault (Fig. 18B). The B"-A" and the A " - e M intervals are thicker on the west part of the structure than on the top. This difference in thickness indicates that this part of the Taino ridge was an old seamount that
was reactivated by recent faulting. The thickening of the crust below the Taino ridge is estimated to be 0.5 km. To the north the Taino ridge loses its elevation. A strike-slip fault may be responsible for the tilt of a block (Fig. 19). A positive magnetic anomaly is related to the block that could be a Cretaceous volcano. Another volcanic ridge lies west of this block (Fig. 20). We failed to detect any recent deformation of this structure that is probably an initial volcanic feature of the Beata igneous province. The shortening of the northern Taino ridge can be estimated to be 4.6 km for a 21.5-km-wide structure (20%; from 8100 to 8500, Fig. 18A). If we assume that the northern Taino ridge and the southern Taino ridge (Fig. 13) had the same width initially, we estimate the shortening to be 39 km per degree. The
Fig. 15. A flower structure indicates the presence of a strike-slip fault. See the inset map (depth to basement: B ' ) and Fig. 13 for profile location.
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NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
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NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
657
Fig. 28. (A) Seacarib 1 seismic profile (see also Fig. 5E) showing the location of the DSDP Site 151 where an hard ground separates Paleocene sediments from the Santonian sandstone marl and chalk (Edgar et al., 1973b). (B) Seacarib 1 seismic profile showing the steep scarp that trends NE-SW. For location see Fig. 27. entation (N45 ~ of compression has been described in the Presqu'~le du Sud d'Hispaniola (Bien-Aime Momplaisir, 1986). This stress increases towards the north as shown by the structures described in the Venezuelan basin (Fig. 11) along the Taino ridge
(Fig. 13) and in Aruba Gap (Fig. 24). The shortening of the Beata ridge induces a shallowing towards the north. It is evident that the present topography results from an uplift linked to the compression, but it is difficult to evaluate the initial topography. We know
658
A. M A U F F R E T and S. LEROY 72~ t
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Fig. 29. (A) Depth to basement map (B"). (B) Seabeam map of the Seacarib 1 cruise showing the steep scarp that forms the boundary between the Beata ridge and the Haiti sub-basin. The position of the survey is indicated in Fig. 3. The position of the seismic profile illustrated in Fig. 30 is shown.
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
659
Fig. 30. Seacarib 1 seismic profile showing the steep scarp that bounds the Beata ridge. This is not the steepest portion of the scarp and the profile is not normal to the structure. The upper slope shows evidence of slumping and the basement outcrops in the lower slope as shown by a diving survey (Mauffret et al., in prep.). A recent deformation can be noted in the Haiti sub-basin (see also Fig. 8). For location see Fig. 29. that the Beata ridge is a 20-km-thick Cretaceous volcanic plateau. This great thickness is mainly caused by a volcanic underplating that initiated the uplift and rifting of the Beata ridge. Although a thickening by compression may have occurred during the Miocene, we think that the initial Cretaceous thickness was important. The Colombian and Venezuelan basins subduct in a normal way below the South American deformed belt. In contrast the northern part of the Beata ridge collides with the central part of Hispaniola in relation with the eastwards drift of the Caribbean plate relative to Hispaniola. This collision results from the buoyancy of thick volcanic crust that resists subduction beneath Hispaniola (Burke et al., 1978; Mercier de Lepinay et al., 1988). The seismic profile displayed in Fig. 27 indicates that a non-reactivated part of the DSDP 151 ridge is buried in the Colombian basin. This basin and the Haiti sub-basin have a thin crust and were consequently much deeper than the initial Beata volcanic plateau. The Colombian basin is depressed by the loading of the thick sedimentary cover and the subduction below the South American deformed belt (7.4 km, Fig. 32G). The 1.3-km-high step shown in Fig. 32G represents the boundary between the Beata ridge and the Colombian basin. This boundary extends from the Pecos fault zone to the southern tip of the DSDP 151 ridge (Fig. 31). 0.5 km of initial
topography and 0.5 km of uplift was estimated for the Pecos fault zone (Leroy and Mauffret, 1996). We suppose that the initial topography was the same and the reactivation is consequently 0.8 km high. This estimation is close to the 1-km-high regional uplift deduced from the DSDP results (Benson et al., 1970). The buffed part of the DSDP 151 ridge (Fig. 32A) was initially deeper than the northern part but the shape of the ridge is conserved (1.9 km high relative to the adjacent basement) except for the seamount located on the southern tip of the ridge (2.2 km, Fig. 32B). This seamount was probably uplifted (0.3 km) by the N E - S W fault that delimits this feature to the south (Fig. 26). The DSDP 151 ridge-Taino ridge area do not show any appreciable topographic step, but the region located between the DSDP 151 and Tairona ridges is 1.8 km uplifted (from 4 km to 2.2 km, Fig. 32G). The northern part of the Beata ridge, near the Bahoruco Peninsula, is 1 km uplifted relative to the south. The northernmost uplift is actually onland where the Sierra de Bahoruco is 2 km high. We conclude that 6.8 km of uplift may occur between the southern tip of the Beata ridge and Hispaniola. However, the results of a submersible survey (Mauffret et al., in prep.) indicate that the Beata ridge was shallow during the Late Cretaceous after the volcanic event and the 6.8 km uplift results from the constructional history of the
660
A. MAUFFRET and S. LEROY
Fig. 31. Depth to basement (B") of the central part of the Caribbean Sea. The eastern part of the Beata ridge is characterized by reverse faulting offset by some NE-SW faults. N-S highs dominate in the central part of the structure. The NE-SW faults are predominant in the western part of the Beata ridge.
volcanic plateau and a recent compressional event from the Miocene to the Present. We estimated about 40 km of shortening per degree of longitude. Thus, 170 km of shortening may have occurred at 18~ (Bahoruco Peninsula). If we assume that the Sylvie ridge was in strike with
the Marie Aimee ridge and the Bahoruco Peninsula, the Presqu'~le du Sud should be separated from the central Hispaniola by a 240-km gap (Fig. 33). The boomerang shape of the Presqu'~le du Sud-Beata ridge may be due to the progressive collision of the Beata plateau and the collage of fragments of
661
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
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Fig. 32. East-west cross-sections (A to F) and north-south cross section (G) of the Beata ridge. The depth to basement (B') is relative to the sea level. Except for section (B) (2.2 km) the DSDP 151 ridge has the same height relative to the adjacent basement (1.7 to 19 km), but we can see the rising towards the north of the ridge (G). For location see the inset map (depth to basement: B').
662
A. MAUFFRET and S. LEROY
Fig. 33. The DSDP 151 ridge the Tairona ridge and the Presqu'~le du Sud have been displaced to form a north-south structure. This reconstruction shows a 240-km-wide gap between the Presqu'~le du Sud and the central part of Hispaniola. The Presqu'ile du Sud may have underwent a counter-clockwise rotation (Mercier de Lepinay et al., 1988; Van Fossen and Channell, 1988; Jany, 1989). this plateau (Mercier de Lepinay et al., 1988; Jany, 1989). Therefore, the present east-west orientation of the Presqu'~le du Sud may be a neotectonic feature and a substantial counter clockwise-rotation of the Presqu'~le may have occurred (Van Fossen and Channell, 1988). If we assume that the compression occurred since the Early Miocene (23 Ma) then the motion rate is between 0.74 cm/yr (170 km) and 1.04 cm/yr (240 km). The eastern boundary is diffuse with an intraplate deformation, particularly along the Taino ridge. However, we do not have good seismic profiles to illustrate the easternmost deformation (Fig. l lA, B) and a deep thrust, dipping to the west, is suggested (Fig. 11C). The several thrusts observed on the eastern side of the Beata ridge imply a deep d6collement level. Fig. 18C shows that this d6collement is deeper than 15 km. It is clear that the Beata ridge forms a boundary between a Venezuelan microplate and a Colombian
microplate and the latter overthrusts the former. We suggest that the eastern boundary of the Colombian microplate lies along the easternmost thrust (Fig. 11) and the Beata plateau. The southern boundary of this microplate is the Colombian deformed belt. The northern part of this belt trends east-west and is narrow (Vitali, 1985; Vitali et al., 1985). North of the Santa-Marta Massif the deformed belt is segmented with N-S and E N E - W S W orientation of the toe of the prism (Fig. 34). After a new prominent bend (Vernette et al., 1992) the Colombian prism merges with the onshore Sinu belt (Duque-Caro, 1979, 1984; Vitali, 1985; Vitali et al., 1985; Toto and Kellogg, 1992) that is 10 km thick (Lehner et al., 1984). This prism was formed in the Early Miocene (Duque-Caro, 1979) along the buried Sinu trench that trends north-south. This active margin is confirmed by the seismicity that defines a 180-kmdeep Benioff zone. The length of the subducted crust is 350 km (Malav6 and Suarez, 1995). This Benioff
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE zone belongs to the Bucaramanga segment (Pennington, 1981) and cannot be related to the Nazca Plate as suggested by Van der Hilst and Mann (1994). We suggest that the Sinu subduction zone is related to the Beata deformed zone by fight-lateral strike-slip faults and short north-south segments of subduction zones (Fig. 34). The northern part of the Sinu belt is clearly (Vernette et al., 1992) offset by a strikeslip fault zone and the South Caribbean marginal fault illustrated by fig. 13 in Kellogg and Bonini (1982) is probably a transpressive feature. Westwards, the Colombian microplate is delimited by the Panamanian subduction zone (Adamek et al., 1988; Protti and Schwartz, 1994). This Panamanian block is delimited to the north by a strike-slip fault that is connected to the Middle American trench south of Nicoya Peninsula (Fisher et al., 1994; Marshall and Anderson, 1995). This peninsula may belong to the Colombian microplate and a fault north of the peninsula may connect the Middle American trench and the Hess escarpment (Dengo, 1985). The western part of the Hess escarpment is active (Bowland, 1993), but the eastern part of this feature does not show clear evidence of recent faulting. On the contrary, the Pedro escarpment, that delimits the upper Nicaragua rise, shows much evidence of recent faulting (Mascle et al., 1985; Holcombe et al., 1990). Moreover, the lower part of the Nicaragua rise has the same volcanic basement as the Colombian basin (Mauffret and Leroy, 1997) and we placed the northwestern boundary of the Colombian microplate along the Pedro escarpment (Fig. 34). The northern limit of the microplate is the Enriquillo strike-slip fault and the Bahoruco thrust. The Venezuelan microplate is delimited to the east by the Lesser Antilles subduction zone and to the south by the Venezuelan deformed belt. The northern boundary lies along the Anegada Passage (Jany et al., 1990) and Muertos trough. A Gonave microplate was defined (Rosencrantz and Mann, 1991; Mann et al., 1995) between the Cayman spreading center, the Oriente fault to the north, the Walton-Plantain Garden-Enriquillo fault to the south and the western coast of Hispaniola. We suggest that this microplate is related to the Venezuelan microplate through the narrow Enriquillo basin. A recent recompilation of the magnetic data (Leroy, 1995) in the Cayman trough determined 0.9 cm/yr of half-rate spreading between chron 8 (26 Ma) and chron 1 (Fig. 35A). The motion along the Walton-Plantain Garden fault is evaluated in Jamaica Island to 0.4 cm/yr during the Miocene to Quaternary (Rosencrantz and Mann, 1991). However, the slip rate of the Enriquillo fault in Hispaniola is evaluated to 0.8 cm/yr (Mocquet and Aggarwal, 1983) and we suggest a 0.9 cm/yr motion towards the northwest of the Colombian microplate (Fig. 35).
663
Therefore, a 0.5 cm/yr of additional motion can be estimated along the Pedro fault. The velocity vector diagram (Fig. 35A) suggests a velocity as high as 2.6 cm/yr of the Colombian microplate relative the North American plate. GPS measurements in the Presqu'~le du Sud determine 2.3 cm/yr for the relative velocity between this area and the North American plate (Farina et al., 1995). The relative motion between the central block of Hispaniola and the North American plate is 1.5 cm/yr according to the previous study. Therefore, the relative motion between the central block of Hispaniola and the Presqu'~le du Sud is 0.8 cm/yr, a value very close to our estimation (0.9 cm/yr). The plate motion of the southern Caribbean region is documented by the GPS studies (Freymueller et al., 1993; Drewes et al., 1995; Kellogg and Vega, 1995). The Colombian microplate subducts beneath the Panama prism with a rate of 1 cm/yr and the rate of convergence between the Colombian microplate and the North Andes block (Sinu trench) can be evaluated to be 1.7 cm/yr (Kellogg and Bonini, 1982). The differential motion between the two plates may result from the greater convergence of the NOAM-SOAM in the western part of the Caribbean zone than in the eastern part (Mtiller et al., 1996) or from the influence of the large convergent motion of the Cocos plate. In the eastern Colombian basin intraplate deformation has been observed northeast of the Panamanian prism (Bowland, 1993) and the results of the DSDP Site 154 (Edgar et al., 1973b) indicate that the uplift of the deformed structure is related to an Early Pliocene reverse fault with a southwest vergence clearly displayed in the seismic profiles shown by Bowland (1993). The negative buoyancy of the young Cocos oceanic crust may induce a compressional stress in the overriding Caribbean plate (England and Wortel, 1980; Meijer, 1992). We cannot decipher the dominant process (Atlantic or Pacific) to produce a higher displacement of the Colombian microplate than the Venezuelan plate. We observe that the chrons 7 (25 Ma) and 8 (26 Ma) correspond to an increase of convergence between the Nazca plate and the South American plate (Pargo-Casas and Molnar, 1987) and between the North and South American plates (Mtiller et al., 1996), respectively. At this time the Cayman spreading center was also reorganized (Leroy, 1995) and Hispaniola began to drift away from Cuba. A global reorganization of plate motion in the Pacific and Atlantic Oceans may have influenced the internal stress of the Caribbean plate. There is much controversy about the pole position of the North America-Caribbean plate either to the north (Sykes et al., 1982; Deng and Sykes, 1995) or the south (Stein et al., 1988; Calais and Mercier de Lepinay, 1993; Lundgren and Russo, 1996). We agree with Heubeck and Mann (1991), who placed
Fig. 34. (A) The Colombian and Venezuela-Gonave microplates have been differentiated with two patterns. (B) Sketch that shows the relationship between the Sinu trench and the compression along the eastern flank of the Beata ridge. Refer to the text for explanation.
7~ -]
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
665
Fig. 35. (A) Early Miocene reconstruction. Refer to the text for explanation. (B) Sketch (from Leroy and Mauffret, in prep., modified) inspired from Heubeck and Mann (1991). We suggest two different poles for the North American plate-Venezuelan microplate and the North American plate-Colombian microplate motions. A pole for Venezuelan-Colombian microplates motion is proposed near the southern tip of the Beata ridge where the compressional deformation is weak.
a pole for the NOAM-Colombian microplate in the south and a pole for the NOAM-Venezuelan microplate in the north (Fig. 35B). The pole for the Venezuelan-Colombian microplate must be very close to the southern tip of the Beata ridge (Fig. 35B) where the compressional deformation is weak.
CONCLUSIONS
This study together with to investigate ticularly true
shows how detailed seismic surveys, a regional coverage, can be useful a complex tectonic area. This is parfor the Taino ridge and Pecos fault
666 zone where we showed reverse faults and pop-up offset by strike-slip faults. The tectonic framework of these areas is coherent with a N E - S W stress that increases towards the north. The orientation of this stress is incompatible with a simple squeeze of the Caribbean plate between the North and South American plates (Burke et al., 1978) although a westwards increase of convergence between the two major plates may have influenced the faster migration of the Colombian microplate towards the east than the Venezuelan microplate. A clear boundary between the Colombian and Venezuelan microplates cannot be traced, although we do not have enough seismic information to exclude a major structure along the easternmost thrust. However, the deformation seem to be diffuse and localized along several structures. These structures are probably pre-existent but it is very difficult to differentiate between the initial topography and the observed topography that results from the Miocene to Recent compression. Our shortening estimations, that are as high as 20%, can be biased by this poor knowledge of the pre-existent topography. The displacement of the Colombian microplate relative to the Gonave-Venezuelan microplate is evaluated between 170 and 240 km since the Early Miocene. This is much greater than the 50 km of offset along the Enriquillo fault in Presqu'~le du Sud (Mann et al., 1995), although an additional motion may occur south of the Presqu'~le du Sud. However, our rate of 0.9 cm/yr is compatible with the recent and preliminary GPS results (Farina et al., 1995). The present Caribbean plate is a composite, partly formed by continental crust (upper Nicaragua rise), island arc crust (Jamaica and north Hispaniola), but the bulk of this plate is the Cretaceous igneous province. The boundary migrates to the south and fragments of the former plate are abandoned and integrated into the North American plate (Mann et al., 1995). In addition we propose internal boundaries that divide the Caribbean plate in several blocks: upper Nicaragua block, lower Nicaragua block, Colombian and Venezuelan microplates, northern Hispaniola. In such a complex tectonic framework it is impossible to study just the main boundary, i.e. the Cayman spreading centerPuerto Rico trench-Lesser Antilles subduction zone, to deduce the motion between North America and the Caribbean plate and we suggest that the pole for the North America-Venezuelan microplate motion is different from that of the North America-Colombian microplate motion. The South American plate undergoes an internal stress due to the rapid subduction of the young oceanic crust of the Nazca plate (England and Wortel, 1980; Meijer, 1992). In the northern part of South America, the North Andes block undergoes an E - W stress and migrates towards the north relative to the South American main plate. A similar
A. MAUFFRET and S. LEROY stress can be applied to the Caribbean plate by the Nazca and the Cocos plates.
ACKNOWLEDGEMENTS
This work was supported by grants INSU ATP 733 and 780. We thank the officers and crew of the R / V Nadir for assistance in this project. We are especially indebted to the technical crew of GENAVIR who greatly helped in the acquisition of multichannel data. These data were processed at Institut de Physique du Globe de Strasbourg and we thank R. Schlich and M. Schaming who have facilitated access to the processing center and helped us to use the Geovector software. We thank J. Diebold from Lamont-Doherty Earth Sciences Observatory, E Mann and J. Austin from the University of Texas at Austin, A. Mascle from the Institut du P6trole, and E Lehner from Shell to have provided several seismic lines that completed our seismic coverage. We thank E Mann, T. Holcombe, N. Donnelly and E. Calais for their constructive reviews and helpful suggestions. Contribution of URA 1759.
REFERENCES
Adamek, S., Frohlich, C. and Pennington, W., 1988. Seismicity of the Caribbean-Nazca boundary: constraints on micro-plate tectonics of the Panama region. J. Geophys. Res., 93, 20532075. Austin, J.A., 1984. Overthrusting in a deep-water carbonate terrane. In: A.W. Bally (Editor), Seismic Expression of Structural Styles. Am. Assoc. Pet. Geol., Stud. Geol., 15: 3.4.2-1673.4.2-172. Bader, R.G., Gerard, R.D., et al. (Editors), 1970. Initial Reports of the Deep Sea Drilling Project. U.S. Government Printing Office, Washington, D.C., Leg 4. Benson, W.E., Gerard, R.D. and Hay, W.W., 1970. Summary and conclusions. In: R.G. Bader, R.D. Gerard et al. (Editors), Initial Reports of the Deep Sea Drilling Project. U.S. Government Printing Office, Washington, D.C., Leg 4, pp. 659-673. Bien-Aime Momplaisir, R., 1986. Contribution ~ l'6tude g6ologique de la partie orientale du massif de la Hotte (Presqu'~le du sud d'Ha'fti). Synthbse structurale des marges de la Presqu'~le h partir de donn6es sismiques. Thbse de l'Universit6 Paris 6, 223 pp. Biju Duval, B., Bison, G., Mascle, A. and Muller, C., 1982a. Active margin processes: field observations in Southern Hispaniola. In: J. Watkins and C.L. Drake (Editors), Studies in Continental Margins. Am. Assoc. Pet. Geol. Mem., 34: 325344. Biju Duval, B., Mascle, A., Rosales, H. and Young, G., 1982b. Episutural Oligo-Miocene basins along the north Venezuela margin. In: J.S. Watkins and C.L. Drake (Editors), Studies in Continental Margins. Am. Assoc. Pet. Geol. Bull., 34: 347358. Bowland, C.L., 1993. Depositional history Ofthe western Colombian basin, Caribbean Sea, revealed by seismic stratigraphy. Geol. Soc. Am. Bull., 105: 1321-1345. Bowland, C.L. and Rosencrantz, E., 1988. Upper crustal structure
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C h a p t e r 22
Caribbean Modelled After Archipelago Orogenesis" Coda in Esperanto by the Series Editor
KENNETH J. HSU
A colleague recently declined our invitation to serve as a candidate for a volume editorship of the Sedimentary Basins of the World, because he would not accept the agreement between the publisher and myself on the role of the Series Editor. I sensed the resentment of one or two previous volume editors that their hard work should be concluded by a poorly informed maverick with his wise cracks, but they were too much of a gentleman to complain. The most recent episode compels me, however, to give an explanation on the rationale of the agreement. Science is a language. As I indicated in the Introduction to the Series, each article of a volume may "speak" a different dialect, and each volume of the series may "speak" a different language. The series editor cannot possibly request revisions of individual contributions, this is the task of the volume editor. However, the series editor has to bring a certain uniformity to the structuring of the various volumes. There are the problems of different approaches of writing, of different working hypotheses for interpretation, of different standards of judgment, etc. In recommending the volume editors to the publisher, I have expressed my full confidence in their ability to do their best. The past volume editors have not disappointed me, and the editor of this Caribbean volume has done a great job, far beyond my original expectations. I am most impressed by the assembly of the many contributions, especially those of his own. The duty of the series editor is to impart a continuity to the series, by expressing his understanding of the volume in a common language. This is the explanation as to why the subtitle of this article is called Coda in Esperanto. English gave us Shakespeare, German gave us Goethe, and Chinese gave us Li Bei. Esperanto is not a literary language, but is an attempt to use one and the same language.
Spoken languages evolved naturally, but Esperanto is an arbitrary, logical construction. Spoken languages have their beauty. The one merit of Esperanto is its uniformity. The apparent differences between the expressions in different languages are minimized. The volumes of this series have been collective efforts, but the series does not consist of philately albums. The series will not be merely collections of random observations, interpreted idiosyncratically by individual prima donnas. There is the attempt to see a parallelism in the evolution of sedimentary basins, in China, in South Pacific, in Africa, and in the Caribbean, and this attempt is expressed by this Coda in Esperanto. In my Introduction to the Series, I emphasized the indispensability of classification, and the need of having all-inclusive and mutually exclusive categories. I suggested two sets of criteria: isostasy, and orientations of the principal stresses. The classification of basins adopted by the volume editor for the Caribbean are: Strike-slip basins Island-arc basins Collisional basins Rift basins Inverted basins The first four are recognized by the classification recommended for the series. They are, respectively, transcurrent (transform or strike-slip) pullapart basins (1.2), rifted basins on active margins or back-arc basins (1.1.4), compressional basins (2.1), and rift basins in continental interior (1.1.1). The fifth category is a combination of two simple types, such as one on passive margin (1.1.3) converted later into a compressional (2.1) and/or transcurrent basin (1.2). This Caribbean volume "speaks" English with a strong North American accent, and the vocabulary is easily translated into the Esperanto.
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 673-676. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
674 The vocabulary on the sections on tectonics has more of a local flavor. Having been a geologist on three continents, an Esperanto in tectonics is more necessary than one in sedimentology. I have developed during the last 20 years the concept of Tectonic Facies. I shall attempt in this short contribution to express my understanding of the Caribbean geology with an Esperanto vocabulary constructed on the basis of the tectonic facies concept.
MAJOR TECTONIC FEATURES
The tectonic setting of the Caribbean has to be described in terms of plate-tectonics. The volume editor proposed the major crustal provinces as follows: Chortfs block Great Arc of the Caribbean Back-arc basins associated with the Great Arc Caribbean ocean plateau How are these tectonic units compared to those of China, of the South Pacific, or of Africa? In a common language to describe the tectonics of all regions of the world, I suggested all-inclusive and mutually exclusive categories of mountains. They are classified on the basis of the orientations of the orientations of the principal stresses: (1) Germanic Type: extensional, e.g. Black Forest and Vosges on the side of Rhine (2) Alpine Type: compressional, e.g. the Alps (3) Californian Type: transpressional, e.g. Cenozoic California Coast Ranges The tectonics of the Chortfs block belongs to the Californian Type. The other three can be interpreted in terms of the tectonic facies of the Alpine Type of mountains (Hsti, 1994, 1999). The three Alpine tectonic facies are: (a) Rhaetide: overriding block in a collision, commonly underlain by rigid basement. (b) Celtide: collision zone of subduction deformation and metamorphism. (c) Alamanide: Decollement folding and thrust of the thin-skinned cover of the subducted block in a collision. These three tectonic facies have been recognized in the Caribbean region.
THE CARIBBEAN ARCHIPELAGO
If James Hall had devoted, like Willem van den Bold, his life to a study of the Caribbean geology, he would not have come up with a theory of geosyncline. Where are the geosynclinal precursors of the Caribbean mountains? Only the theory of plate-tectonics provides us with a more comfortable answer. An archipelago of islands, marine carbonate banks,
K.J. HSU and oceanic basins came into existence in Jurassic after Africa and South America were split off from North America (Chapter 1). The fragmented terranes underlain by continental crust include the Bahamas, Yucatan/W Cuba, the Maya and Chortfs blocks. The deep-sea basins of the archipelago were the incipient Gulf of Mexico and the proto-Caribbean. Thus portrayed, the Mesozoic Caribbean differed from the Indonesian Archipelago. The first small oceanic basins of the Caribbean owed their origin from regional extension, related to the fragmentation of Pangea and displacement of Africa and South America. The Indonesian Archipelago, on the other hand, resulted from back-arc seafloor spreading; that archipelago is bounded by an island-arc on the active plate margin. There was, however, arc-volcanism in the Trans-Mexican Volcanic belt. Could the Gulf of Mexico and the proto-Caribbean have been basins formed by back-arc seafloor spreading? Could Cuba, Yucatan, Maya and Chortfs block have been remnant arcs behind the frontal arc of the Trans-Mexican Volcanic Belt? The volume editor and his authors presented good evidence against such an alternative interpretation. The series editor accepts their working hypothesis that the Mesozoic Caribbean Archipelago owed its origin to the displacement of Africa and South America from North America.
THE RHAETIDES, CELTIDES, AND ALAMANIDES OF CUBA, CHORTiS, AND HISPANIOLA
The sequence of events indicated by the geology of Central America and the Greater Antilles included Mesozoic rifting and subsidence, followed by Cretaceous and/or early Tertiary collision. The Jurassic and Lower Cretaceous strata in areas underlain by continental crust are mainly carbonate strata, and they have been considered passive margin deposits (Chapters 4, 5, 6 and 7). The presence of arc-volcanics and foredeep sediments indicates, however, the change from a passive margin to a magmatic arc during late Cretaceous and/or early Tertiary time. The geology of Cuba gives evidence of a collision between the Greater Antilles Arc and the North American margin (Chapter 4). In western Cuba, the sedimentary cover of the subducted arc, has been stripped off to form the southerly vergent foreland fold-and-thrust belt (alemanide facies). The overriding block was the Bahamas Platform. The suture zone of the collision is manifested by the ophiolite-melange (celtide facies) of Cuba. The collision structures at the western end of the Great Arc of the Caribbean are northerly vergent. The Chortfs basement (rhaetide) was thrust northward upon the Maya block during late Cretaceous and early Tertiary (Chapter 8). The suture zone is
CARIBBEAN MODELLED AFTER ARCHIPELAGO OROGENESIS
marked by a zone of serpentinites and metamorphic rocks (celtide facies) of Roatfin and Barbareta islands. The collision-structures are cut by the Tertiary transcurrent faulting. Whereas the Yucatan Basin owed its origin to back-arc seafloor spreading, the Cayman trough is an 1100-km long and narrow pullapart basin of Cenozoic age (Chapters 1 and 2).The northerly vergent structure south of the Cayman Trough continued eastward to northern Hispaniola, where the collision between the North American and the Caribbean Plates is well documented (Chapters 9, 10 and 11). The active margin of the Caribbean Plate has been defined by the Hispaniola volcanic arc, which ceased to be active during the Eocene. The subduction complex under the arc consists of serpentinite, gabbros, blue schists, and submarine volcanics (celtide facies). The overriding rhaetide is a granodiorite, which was the root of the magmatic arc (rhaetide facies). The collision-structures on Hispaniola and nearby islands are also modified by post-Eocene strike-slip faulting, related to the eastward movement of the Caribbean Plate (Chapters 11, 12 and 13).
675
AN EXOTIC PACIFIC COMPONENT IN THE CARIBBEAN
Discovery of thick Cretaceous basalt by DSDP Leg XV in 1970 gave the first indications of the unusual nature of the Caribbean basins. The Venezuela Basin and the Colombia Basin are underlain by a Mesozoic ocean crust and a thick sequence of basalt of an oceanic plateau. The basins were not formed by Cenozoic seafloor spreading behind a magmatic arc. The Caribbean Plate, according to the mobilitic working hypothesis adopted by the authors of this volume, "was originally an area of eastern Pacific Ocean floor and oceanic plateau" (Chapter 1). The Mesozoic ocean has become a Cenozoic basin, after an island arc was constructed on the active plate margin. This conclusion is now verified by a wealth of geological, geophysical, and ocean-drilling evidences (Chapters 19 and 20). The Beata Ridge separating the Venezuela and Colombia basins is not a remnant arc such as those of the Philippine Sea. The ridge was a part of the Mesozoic oceanic-plateau and has been deformed by compression during the Neogene (Chapter 21).
THE ISLAND-ARC MARGIN OF THE LESSER
ANTILLES
CARIBBEAN MODEL OF ARCHIPELAGO OROGENESIS
The Cenozoic tectonic evolution of the Caribbean island-arc margin has been very well investigated by geophysics, oceanography, and deep-sea drilling (Chapters 14 and 15). There is the accretionary complex, the oldest of which has been uplifted and is exposed on Barbados. There is the volcanic arc of the Lesser Antilles. There is the Grenada Basin, formed by back-arc seafloor spreading. And there is the Aves Ridge, a remnant arc. Those tectonic units serve as analogues of deformed arcs and basins in orogenic belts.
The geologic history of the Indonesian Archipelago has inspired me to formulate a model of archipelago orogenesis (Hsti, 1994, 1999). My reading of the Caribbean volume has been a revelation. I have come to realize that not all archipelago basins owe their origin to seafloor spreading behind the island arc of an active margin. Nor are all the mountains on the islands related to back-arc basin collapses. The Wilsonian processes of rifting, subduction, and collision are still the paradigm. The geology of the Indonesian Archipelago illustrates, however, only a continent-ocean interaction. The history of the Caribbean Archipelago started out as an ocean-ocean interaction, before the collisions of the Caribbean ocean plate with the continents. I was uncomfortable with my interpretation of the late Precambrian and early Paleozoic geology of China (Hsti, 1999). The North and South China have the sizes of microcontinents, and I had difficulty in considering them remnant arcs, split off during the back-arc seafloor-spreading processes. The interpretation of the Tarim Basin as a relict back-arc basin also invited criticism. After my reading of the Mann Volume, I see the need of revising my Geologic Atlas of China even before it is published. North and South China were more likely microcontinents, formed when the late Precambrian Pangea was fragmented. The Tarim and Qaidam basins were more likely fragments of an oceanic plateau in a
AN ON-GOING A R C - C O N T I N E N T COLLISION
The tectonic evolution of the Trinidad/eastern Venezuela area is distinguished by a sequence of four phases: (1) pre-Jurassic, pre-rift phase, (2) a Jurassic rift phase related to the separation of North and South Americas, (3) a Cretaceous to Oligocene thermal subsidence of the South American rifted margin, and (4) a Neogene foredeep phase related to an arc-continent collision. Superposed on the collision-deformation is a Neogene tectonics dominated by strike-slip faulting, related to the eastward movement of the Caribbean Plate relative to South America. The Chapters, 16, 17 and 18 are written in the standard "language" of plate-tectonics and require no translation into the "Esperanto."
676 proto-Pacific Ocean. They have survived, like the Colombia and Venezuela basins, as rigid blocks during the Phanerozoic deformations. The fate of the Caribbean basins is predictable. With the sediments brought down by the Mississippi, the Magdalena, and other rivers of North and South America, the Caribbean Sea is destined to be converted, in 100 or 200 million years of time, into a desert like the Taklimakan.
K.J. HSU REFERENCES
Hsti, K.J., 1994. The Geology of Switzerland: an Introduction to Tectonic Facies. Princeton Univ. Press, Princeton, NJ, 250 pp. Hsti, K.J., 1999. The Geologic Atlas of China. Elsevier, Amsterdam (in press).
Author Index *
Abbot, EL. 282, 286 Aboud, N. 474 Abrams, L.J. 575, 586, 587, 597, 624 Acton, G.D. 626 Adamek, S. 51, 57, 663, 666 Adams, A.I. 216 Adatte, T. 136, 148 Aden, L.J. 425,473, 476 Aggarwal, Y.E 493, 494, 499, 506, 556, 668 Agterberg, EE 58, 118 Aiello, I.W. 105, 118 Aita, Y. 126, 148 Alba, J.A. 162, 164 Alberding, H. 476 Aldrich, J. 217 Aldrich, M.J., Jr. 215, 216 Algar, S. 478, 479, 489, 492, 494 Algar, S.T. 55, 57, 57, 505, 507, 509, 533, 539, 553,555 Ali, W. 510, 557 Allenbach, E 668 Ambeh, W.A. 388 Amilcar, H. 669 Amilcar, H.C. 669 Andersen, B. 192 Anderson, D.L. 584, 587, 613, 624 Anderson, R.S. 663, 668 Anderson, T.H. 108, 112, 118, 124, 148, 152, 163, 164, 164 Andreieff, E 349, 351-353, 355, 364 Angelier, J. 208, 211,216, 217 Angstadt, D. 91 Angstadt, D.M. 65, 74, 90 Anselmetti, F. 191 Anselmetti, ES. 171,191 Antoine, J. 90, 588, 667 Antoine, J.W. 120, 191,625 Aponte, A. 474 Applegate, A.V. 88, 90 Applin, E.R. 159, 164 Arden, D.D. 197, 199, 217 Areces, A. and 118 Argus, D. 29, 218, 236, 669 Argus, D.E 58, 217, 235, 285, 341,387, 556 Arkell, W.J. 126, 141,148 Arnstein, R. 423,473 Arozena, J. 30, 475 Artoni, A. 385, 386, 386 Atwood, M.G. 155, 164, 165, 217 Aubry, M.-E 341 * Page references to text are in Roman type, to bibliography in italics.
Audemard, E 474, 476 Audemard, EE. 421,423,435, 437, 455, 473 Austin, J.A. 90, 191,635, 666, 667 Austin, J.A., Jr. 29, 90, 167, 168, 191-193, 364 Av6 Lallemant, H.G. 5, 6, 13, 15, 21, 29, 206, 211, 213, 217, 435, 474, 503,505, 509, 541,553, 555 Avedik, F. 387 Aves, H.S. 217, 223, 235 Azavache, A. 424, 474 Azema, J. 152, 164 Azpiritxaga, I. 473 Babb, S. 14, 21, 23, 425, 474, 513, 526, 539, 546-548, 555, 557 Baby, E 385, 386, 386, 388 Bader, R.G. 630, 641,642, 666 Bajo, Bakker, G. 668 Bakker, J.G.M. 626 Baldwin, S.L. 621,625 Bale, E 388 Ball, M. 193 Ball, M.M. 12, 29, 167, 169, 171, 191, 478, 494 Bally, A.W. 435,474, 476 Bandy, W.L. 394, 415 Bangs, N. 371,385, 386, 388 Banks, C.J. 489, 494 Banks, L.M. 437, 474 Banks, N.G. 222, 235 Banks, EO. 217 Banner, ET. 159, 164 Bargar, K.E. 217 Barker, E 394, 415 Barr, K.W. 425, 474, 509, 511,556 Barrett, D.L. 582, 588 Barrett, S.E 6, 14, 16-19, 21, 23, 31, 34, 52, 59, 108, 111, 115, 116, 120, 152, 163, 165, 168, 169, 188, 192, 219, 235, 389-391, 411, 412, 414, 415, 416, 423, 474, 475, 498, 501, 503, 505, 506, 539, 545, 546, 551, 556, 557, 591, 613, 615, 616, 618, 619, 626, 627, 635, 668 Barros, J.A. 6, 12, 23, 30, 168, 169, 191, 192 Bartenstein, H. 509, 556 Bartok, E 423, 474 Barton, J. 286 Barton, R. 152, 157, 165, 198, 218 Bass, M.N. 200-203, 205, 206, 208, 213,218, 219, 223, 235
Bassoullet, J.E 161,164 Bateson, J.H. 80, 90 Bauman, E 473 Baumgartner, EO. 113, 118, 128, 130, 148 Bayes, J. 365 Bazhenov, M.L. 109, 115, 118 Beach, D.K. 171,191 Beall, R. 253, 285 Beard, L.S. 285 Beaudouin, T. 118 Beaumont, C. 435, 474 Beck, C. 31, 57-59, 388, 476 Beck, M.E., Jr. 213,217 Beckmann, J.E 474 Beebe, W. 148 Beets, D. 30 Beets, D.J. 562, 587 Behrens, G.K. 351,353,364 Bejarano, C. 420, 424, 474 Belderson, B. 669 Belderson, R.H. 372, 387 Bellizia, G.A. 435,474 Bellizzia, A. 477, 494, 617, 618, 625 Benedetti, M. 387 Benjamini, C. 112, 118 Benson, W.E. 649, 659, 666 Bercovici, D. 587 Berggren, W.A. 305, 341 Bermudez, EJ. 253, 285 Bernier, E 164 Bernoulli, D. 191 Berryman, K.R. 217 Berthon, J.L. 388 Best, D.M. 415 B6thoux, N. 58 Bettenstaedt, F. 556 Beunk, EE 587 Bibee, L.D. 397, 415, 416 Bickle, M.J. 588 Bien-Aime Momplaisir, R. 633-635, 657, 666 Bierley, R.E. 425,473 Biju-Duval, B. 51, 57, 164, 370-372, 378, 379, 387, 388, 476, 563-566, 581, 582, 587, 593, 597, 615, 625, 630, 633, 635, 637, 638, 651,666 Bilham, R. 385,387 Bird, D.E. 6, 7, 13, 28, 29, 389, 390, 392, 397, 398, 415,415 Birsh, ES. 371,387 Bitri, A. 668 Bizon, G. 625, 666 Bizon, J.J. 164 Blanc G. 387
678 Blanchet, R. 31, 57, 59, 365, 388, 476 Blarez, E. 58 Blome, C . D . 1 2 4 , 126-130, 141, 148-150 Blondeau, A. 285 Blow, W.H. 233, 235, 459, 461,474 Bobier, C. 387, 669 Bock, W.D. 29, 191, 494 Boisseau, M. 243, 245 Boisson, D. 669 Bold, W.A. 292, 297, 304, 305, 307, 309, 342 Bolli, H.M. 159, 164, 346, 364, 466, 474, 556 Bonet, E 154, 159, 164 Bonini, W.E. 663, 667 Borgois, J. 476 Bosch, M. 618, 625 Bosselini, A. 333, 341 Bott, M.H.E 415 Bougault, H. 403,415 Boulbgue, J. 387 Bourgeois, E 192 Bourgois, J. 31, 57, 59, 242, 246, 258, 285, 388 Bouwman, S.A. 476 Bouysse, E 6, 13, 14, 29, 31, 57, 59, 364, 365, 370, 387, 388, 389-391, 397, 398, 411-415, 415, 438, 474, 476, 668 Bowin, C. 617, 625 Bowin, C.O. 247, 249, 279, 285, 292, 296, 341 Bowland, C. 625 Bowland, C.L. 5, 16, 29, 30, 564, 565, 581, 587, 597, 617, 625, 633, 650, 663, 666, 667 Bowles, R.M. 29, 191 Boynton, ~.H. 398, 409, 415 Boynton, W.V. 119 Bradley, D.C. 117, 118, 235, 365 Brakenridge, G.R. 285 Bralower, T.J. 6, 12, 28, 29, 52, 57, 97, 108, 117, 118 Bray, R. 509, 510, 548, 556 Breen, N. 192 Breen, N.A. 388 Breyer, J.A. 235 Briceno, L. 669 Britton, J.C. 235 Brodholt, J. 218, 236, 669 Br6nnimann, E 141,148 Brooks, D.A. 394, 397, 414, 415, 415 Brown, D.J. 589 Brown, K.M. 371,372, 378, 382, 387 Brown, N. 90 Brown, N.K. 191 Browning, J.M. 617, 618, 625 Bryan, G.M. 193 Bryant, W. 65, 90 Bryant, W.R. 167, 191 Bryne, D. 635, 667 Bueno Salazar, R. 16, 29 Buffler, R.T. 6, 8, 18, 19, 21, 29, 30, 64, 65, 67-72, 74, 82, 87, 89, 90, 91, 108-110, 112, 119, 124, 148, 168, 183, 191, 289, 342, 474, 588, 626, 668
AUTHOR INDEX Buhl, E 58, 388, 587, 589, 625, 626, 668, 669 Bullard, E.C. 124, 148 Bullard, T.E 286 Burbank, D.W. 375, 387 Burckhardt, C. 126, 127, 130, 133-135, 137-139, 146, 147, 148 Burkart, B. 6, 9, 29, 152, 164, 199, 200, 213,217 Burke, K. 4-8, 13, 17, 29, 30, 51-53, 55, 56, 57-59, 168, 169, 188, 191, 192, 200, 201, 217, 218, 235, 289, 341, 342, 344, 360, 363, 364, 365, 388, 477, 478, 494, 498, 500, 501, 503, 506, 507, 509, 510, 531, 540-542, 545, 546, 551, 553, 556, 557, 591, 625, 628, 651, 659, 666, 667 Burr, G. 286 Butterlin, J. 285, 667, 669 Byme, D.B. 50, 57 Cabrera, E. 473 Cabrera, S. 474 Caceres Avila, E 215, 217, 219, 223, 235 C~iceres, D. 30, 118, 148, 556 Calais, E. 29, 31, 50, 57-59, 199, 217, 248, 249, 253, 257, 279, 281-284, 285, 341, 364, 388, 476, 556, 633, 663, 667 Calassou, S. 371,387 Calkins, EC. 286 Callomon, J.H. 150 Camargo, Z.A. 119 Campa, M.E 130, 145, 148 Campan, A. 668 Cande, S.C. 36, 37, 39, 41, 42, 51, 52, 56, 58, 59, 120, 218, 416, 668 Cantfi-Chapa, A. 120, 135, 140, 141, 143, 148 Carey, S.W. 124, 148 Carfantan, J.C. 31, 57, 59, 164, 388, 476 Carlson, R.L. 415 Carnevali, J.O. 423,474, 476, 557 Carpenter, G. 387 Carpenter, R.H. 153, 155, 164 Carr, M.J. 30 Carr-Brown, B. 425, 474, 510, 519, 523, 526, 556 Carrillo-Bravo, J. 138, 148 Case, J.E. 4, 5, 29, 50, 58, 94, 118, 197-199, 201, 217, 360, 364, 435, 474, 477, 494, 561, 563, 584, 587, 593, 617, 625, 628, 639, 667, 669 Casero, E 385, 386, 386, 388 Casey, J.E 29, 415, 416, 588 Cashman, S.M. 213, 217 Castrec, M. 387 Castro-Mora, M. 450, 474 Caus, E. 165 Cavanaugh, T. 577, 589 Cazes, M. 668 Cederstrom, D.J. 345, 347, 349, 354, 358, 364 Chalaron, E. 379, 382, 385, 387 Chang, T. 35, 38-40, 58, 59 Channel, J.E. 475
Channell, J.E.T. 662, 669 Chauvin, A. 115, 118 Chen, Y.J. 565, 587 Cheng, Y. 126, 128, 130, 150 Chennouf, T. 387 Chermak, A. 282, 286 Chevalier, Y. 423, 435, 439, 442, 446, 447, 474 Chiari, M. 105, 118 Childs, J.R. 364 Chin, A. 244, 246 Chou, G.T. 388 Chou, T.-A. 217 Chowns, T.M. 91 Christie-Blick, N. 192, 588, 625 Christofferson, E. 561,587 Cita, M.B. 338, 341,342 Clague, D.A. 586, 587 Clark, G.S. 217, 235 Clark, T.E 415 Cloetingh, S. 23, 31 Cluff, L.S. 218 Coates, A.G. 29 Cobbold, ER. 7, 29 Cobiella-Reguera, J.L. 97, 100, 118 Coffin, M.E 5, 29, 583-585, 587, 588, 613, 615, 625 Cole, J.T. 414, 416 Coleman, M.L. 111, 113, 118 Colletta B. 386, 388 Collette, B.J. 36, 48, 58, 59 Collins, J.A. 588 Collins, L.S. 9, 29 Colwell, J.B. 583,587, 588 Condit, D.D. 286 Coney, EJ. 124, 130, 145, 148 Conkin, B.M. 157, 164 Conkin, J.E. 157, 164 Conners, C. 388 Conrad, M.A. 164 Contreras-Montero, B. 136, 148 Coogan, A.H. 159, 162, 164 Cook, H.E. 282, 285 Cooke, W. 286 Cooper, C. 29, 191,217, 364, 494, 625 Cooper, J.C. 317, 341,342, 668 Cooper, M.A. 31 Coriano, M. 474 Corrigan, J. 56, 58 Corrigan, J.D. 9, 29 Corso, W. 82, 90, 192 Cosgrove, E 503, 538-540, 547, 556 Covey, M. 108, 118 Coward, E.L. 29 Coward, L. 191 Cowper, S. 365 Cox, A. 58 Cox, EG. 475 Craig, L.E. 475 Cramez, C. 476 Cramez, C.D. 471,474 Crosby, J.T. 193 Cross, T.A. 394, 415 Crowell, J.C. 185, 191 Crux, J. 424, 474 Curray, J.R. 397, 415 Curry, R.E 164 Curth, EJ. 364
AUTHOR INDEX Cushman, J.A. 159, 164, 347, 364 Cutten, H.N.C. 217 Dahlen, E 388 Dahlen EA. 382, 387, 388 Dallmeyer, R.D. 69, 74, 90 Dallmus, K.E 476 Damond, E 588, 668 Damuth, J.E. 371,385, 387, 625 Daniels, D.L. 91 Danilewski, D. 120 Davies, H.L. 587, 588 Davila-Alcocer, V. 126, 148 Davis, D. 382, 385, 387 Davis, D.M. 388 Davy, E 29 Daza, J. 436, 437, 461,474 de Albear, J.E 96, 98, 100, 119, 120 de Graziansky, EC. 475 de la Torre, A. 98, 105, 106, 108, 118 De Leon, R. 292, 296, 341 De Lepinay, B. 388 De Lepinay, B.M. 476 de Urreiztieta, M. 29 de Zoeten, R. 23, 29, 248, 249, 251, 253-255, 257-260, 266, 267, 273, 283-285, 285, 341 Dean, B.W. 200, 217 DeBalko, D.A. 67, 82, 87, 88, 90 DeCelles, EG. 313,341 Decima, A. 342 Deloffre, R. 164 DeMets, C. 3, 4, 29, 53, 58, 59, 197, 198, 217, 218, 220, 223, 235, 236, 285, 288, 341, 364, 370, 371, 380, 385, 387, 497, 498, 500, 509, 556, 667, 669 Deng, J. 200, 217, 663, 667 Dengo, G. 152, 163, 164, 197, 199, 217, 477, 494, 617-619, 625, 663, 667 Denny, W. 169, 171,191 Denny, W.M., III 12, 29 Denyer, E 31 Dercourt, J. 31, 58, 59, 388, 423, 441, 474, 476 Detrick, R.S. 587, 588 Dewey, J.E 29, 35, 59, 108, 115, 118, 120, 168, 191, 192, 217, 218, 364, 394, 415, 415, 416, 423, 475, 494, 591,592, 625, 626 Dewey, J.W. 4, 29 Di Croce, J. 15, 18, 19, 21, 23,420, 441, 474, 477, 494 Di Giacomo, E. 474 Dia, N.A. 382, 387 Diallo, M.C. 476 Dfaz, M.L. 118 Dickinson, W.R. 124, 148, 275, 277, 279, 282, 285 Diebold, J.B. 5, 16, 19, 21, 28, 29, 563, 565, 566, 577, 580, 581, 583, 587, 593, 597, 613-615, 617, 625, 633, 637, 641,649, 649, 653, 667 Dieni, I. 162, 164 Dietz, R.S. 124, 148 Dill, R.E 357, 364 Dillon, ES. 168, 191 Dillon, W. 12, 29
679 Dillon, W.E 363, 364, 437, 446, 474, 635, 667 Dinkelman, M.G. 6, 16, 30, 168, 192, 591,626 Dix, C.H. 405, 415 Dixon, T. 285, 288, 289, 297, 302, 304, 341,667 Dixon, T.D. 388 Dixon, T.H. 29, 51, 58, 362, 364, 497, 498, 556 Dmitriev, L. 415 Dobson, L.M. 67, 82, 87, 88, 90 Dodd, J.E. 29, 191 Dohm, C.E 253, 285 Dolan, J. 30, 58, 192, 251, 285, 286, 341,365 Dolan, J.E 6-8, 12, 13, 29, 30, 49, 55, 58, 248, 249, 255, 274, 275, 279, 281,283,285, 290, 291,295,341 Donnelly, T.W. 5, 6, 30, 112, 115, 118, 152, 153, 155, 164, 197, 199, 200, 213, 217, 218, 234, 235, 358, 364, 561-563, 573, 577, 585, 587, 588, 591, 593, 597, 613, 622, 623, 625, 653,667 Donovan, S.K. 526, 556 Dooley, T. 492, 493, 494 Doppelhammer, S.H. 120 Douglas, R.G. 126, 148 Doust, H. 668 Drake, C.L. 200, 217, 416 Draper, G. 19, 28, 30, 31, 94, 119, 135, 141, 149, 169, 191, 239, 241-243, 246, 248, 249, 255, 257, 272, 273, 279, 281-283, 285, 286, 342, 668 Drewes, H.D. 663,667 Driscoll, N.W. 5, 16, 19, 21, 28, 29, 580, 581, 588, 593, 597, 611, 613, 614, 617, 622, 625, 637, 641, 653, 667 Driver, E.S. 437, 474 Drobne, K. 165 Duffield, W. 217 Duncan, R.A. 17, 30, 31, 389, 415, 584, 587, 588, 591, 592, 597, 600, 613, 625, 626, 627, 667 Duque-Caro, H. 662, 667 Dyer, B. 557 Dyer, B.L. 503, 538-540, 547, 556 Dziewonski, A.M. 217 Eberle, W. 249, 253,255, 258, 259, 272, 273, 285 Eberli, G. 168, 171,188-190, 191,192 Eberli, G.P. 18, 19, 21, 23, 191 Echeverria, L.M. 562, 588 Edgar, N.T. 29, 30, 58, 241, 246, 364, 415, 564, 575, 588, 597, 600, 602, 617, 621,625, 667, 668 Edgar, T. 667 Edgar, T.N. 627, 630, 631, 641, 650, 663, 667 Edwards, L. 58, 286 Edwards, R.L. 31,192, 342 Edwards, R.S. 415, 416, 626 EEZ Scan Scientific Staff 360, 364 Ego, E 51, 58 Eiras, J.E 441,474
Eisner, EN. 476 Eldholm, O. 583-586, 587, 588, 613, 615, 625 Elsasser, W.M. 53, 58 Embley, R.W. 387 Emmel, EJ. 415 Emmet, EA. 153-155, 157, 164 Endignoux, L. 373, 374, 385, 386, 387, 388 Engebretson, D. 52, 58 Engebretson, D.C. 52, 58 Engelen, J.E 218 Engeln, J.E 236, 388, 669 England, E 51, 58, 663, 665, 667 Eppler, D. 217 Erben, H.K. 138, 139, 148 Erdman, C.E 217 Erikson, J.E 423, 424, 447, 474 Erjavec, J.L. 285 Erlich, R.N. 423, 474, 501, 503, 505, 506, 510, 511,545, 546, 551,556 Escalante, G. 617, 625 Escandon, M. 475 Escobar, C. 216, 217 Espinosa, A. 589 Estrada, J.J. 588 Eugster, H.E 334, 341,342 Eva, A. 509, 510, 548, 556 Eva, A.N. 19, 30, 51, 56, 58, 506, 556 Evans, C.C. 288, 234, 335, 340, 341 Evans, R. 341 Everett, J.E. 148 Everett, J.R. 157, 164 EW-9501 Science Team 667 Ewing, J. 387, 564, 588, 597, 625, 634, 635, 667, 669 Ewing, J.I. 397, 409, 415, 416, 588, 625, 626 Ewing, M. 365, 371,387, 588, 667, 669 Fahlquist, D.A. 120 Falconer, R.K.H. 588 Farfan, EE 556 Farina, E 29, 285, 341, 364, 556, 665, 667 Faug6res, J.C. 370, 372, 385, 386, 387 Faulkner, B. 476 Feinberg, H. 285, 669 Feo-Codecido, G. 439, 441,474 Feray, D.E. 165, 218 Ferguson, R.C. 285 Fernandez, E 475, 556 Fern~indez Carmona, J. 118, 119 Fern~indez, J. 98, 101, 112, 118, 121 Fern~indez Rodriguez, G. 119 Figueroa de Sanchez, L. 505, 556 Finch, R.C. 5, 15, 18, 19, 21, 118, 152-157, 162, 163, 164, 165, 199, 217, 235 Fink, L.K. 398, 416 Finnemore, S. 588 Fisher, D.M. 663, 667 Fitch, T.J. 213, 217 Flemings, EB. 435,474 Flinch, J. 474 Flinch, J.E 23,479, 494 Flood, R.D. 625 Flores, G. 474
680 Flores, R. 30, 118, 148, 556 Flores, W. 217 FRiegel, C.V. 192 Fontas, E 371,384, 387 Ford, R.L. 286 Foreman, H.E 241,246 Foucher, J.E 387 Fourcade, E. 152, 164 Fox, E 57, 667 Fox, EJ. 29, 118, 217, 398, 416, 441, 447, 474, 494, 592, 625, 628, 667 Foye, W.G. 219, 235 Frampton, J. 425, 474, 510, 519, 523, 526, 556 Franke, M. 494 Frankel, A. 360, 364 Freeman-Lynde, R.E 441,475 French, R.B. 150 Frey, M. 509, 556 Freymueller J. 387 Freymueller, J.T. 663, 667 Friedman, G.M. 288, 291,336, 341,342 Friend, EE 8, 31, 369, 370, 385, 388, 489, 494 Frisch, W. 16, 17, 30, 218, 562, 588 Frohlich, C. 666 Frost, S.H. 165, 351-353, 356, 364, 365 Fucugauchi, J.U. 120 Fundora Granda, M. 120 Funes, D. 474 Funkhouser, H.J. 421, 437, 463, 474, 475
Furrazola-Bermtidez, G. 98, 103, 118, 119
Furrer, M. 90 Furrer, M.A. 191, 443, 448, 450, 457, 474, 509, 556 Gahagan, L.M. 29, 59, 192 Gajardo, E. 494 Galea-Alvarez, E 474 Gallango, O. 423,474, 475, 556 Gallo, J. 154, 164 Gapais, D. 29 Garcia, A. 120 Garda, S. 291,292, 341 Gardner, T.W. 667 Garman, K. 168, 191 Gayet, J. 669 Gaylord, M. 192 Gealey, W.K. 28, 30, 168, 191 Gefell, M.J. 208, 218 Geist, E.L. 357, 364 George, R.E, Jr. 440, 474 Georges, G. 588, 668 Gerard, R.D. 666 Gerhard, L.C. 345, 347, 349, 351-355, 364, 365
Ghosh, N. 389, 416, 561,588 Gibbs, A.D. 185, 191 Giegengack, R.E 625 Giffuni, G. 450, 474 Giffuni, R. 474 Gilbert, L. 29, 191 Gilbert, L.E. 474 Gill, I.E 13, 344, 345, 348, 350, 351, 353, 354, 357, 365 Gillett, M. 217, 235
AUTHOR INDEX Ginsburg, R.N. 167, 168, 171, 188-190, 191,192 Gla~on, G. 285 Gleason, R.J. 91 Glover, L. 91 Goff, E 217
Gomberg, D.M. 199, 213, 217 Gomez-Luna, M. 148 Gonthier, E. 387, 387 Gonzales, G. 424, 474, 475 Gonzales-Leon, C. 153, 157, 159, 165 Gonzalez de Juana, C. 6, 30, 419-421, 421, 423, 439, 441, 447, 450, 461, 463, 475, 478, 494, 503, 509, 548, 556
Goodell, H.G. 168, 191 Gordon, M. 249, 286 Gordon, M.B. 4-6, 13, 18, 21, 23, 30, 117, 118, 144, 147, 148, 150, 152, 153, 155, 164, 198, 199, 201, 211, 217, 219, 223, 225, 234, 235, 552, 556
Gordon, R. 29 Gordon, R.G. 53, 58, 59, 217, 218, 235, 236, 285, 341,387, 556, 669
Gorini, C. 668 Gose, W.A. 23, 30, 69, 91, 152, 156, 164, 165, 198, 217 Gou, Y. 475, 476, 556, 557 Gouyet, S. 441,475 Gradstein, EM. 51, 52, 58, 118 Graham, E.A. 365 Graham, R.H. 342 Graham, S.A. 275,285 Grant, A.C. 588 Grant, B. 622, 626 Grant-Mackie, J.A. 126, 148 Green, C. 668 Griboulard, R. 370, 371,379, 380, 384, 387
Grimm, J.E 286 Grindlay, N. 248, 286 Grindlay, N.R. 12, 30, 365 Gripp, A.E. 53, 58 Grippi, J. 57 Grodzicki, J. 120 Groetsch, G.J. 284, 286 Groschel, H. 588 Grue, K. 584, 588 Gruszczyfiski, M. 118 Guellec, S. 388 Gtiendel, E 31 Gueneau, J. 668 Guerra Pena, E 289, 341 Guerrero, J. 626 Gursky, H.J. 150 Guth, L.R. 213,217 Guti6rrez, G. 30, 285 Haczewski, G. 87, 90, 98, 99, 111, 112, 118, 120, 135, 141,148 Hagen, R.A. 584, 588 Hagstrum, J.T. 127, 128, 148 Hall, R. 150 Hall, S.A. 29, 77, 90, 415, 416, 588 Hallam, J.M. 471,475 Hallock, E 556 Hallot, E. 668
Hamoui, M. 162, 165 Hampton, M.A. 274, 286 Hancock, J.M. 475 Haq, B.U. 81, 90, 163, 165, 284, 286, 334, 335, 341, 346, 356, 365, 446, 448-450, 455, 457, 459, 461-463, 466, 471-473,475, 512, 556 Hardenbol, J. 58, 90, 118, 165, 286, 341, 365, 475, 556
Hardie, L.A. 334, 341 Harding, T.E 185, 191,192, 217, 366 Hardy, N.C. 31 Hargraves, R.B. 17, 30, 120, 389, 415, 584, 587, 588, 591, 592, 597, 600, 613,625, 627, 667 Harkrider, D.G. 416, 626 Harland, W.B. 466, 475 Harms, EJ. 317, 341 Harrison, C.G.A. 36, 58, 494 Harrison, T.M. 213, 218, 625 Harry, D.L. 415 Hatten, C.W. 31, 93, 95, 97, 98, 101, 106, 107, 112, 115, 117, 118, 119, 143, 148, 167, 168, 192 Haxby, W. 59, 120, 218, 416 Haxby, W.E 58, 587 Hay, W.W. 666 Hayes, D.E. 397, 416 Hayward, A.B. 342 Heath, R.E 31 Hedberg, H.D. 421,463,474, 475 Heezen, B.C. 59, 398, 416, 441, 474, 475, 625, 628, 667 Heiken, G. 199, 217 Hellinger, S.J. 38, 39, 58 Hempton, M.R. 6, 12, 23, 30, 168, 192, 235, 365
Hennion, J. 588, 625, 667 Hennion, J.E 416, 626 Henry, M. 415 Henry, E 372, 387 Hernandez, E. 668 Hernandez, G. 476 Hernandez, M. 249, 258, 286 Hernandez, N. 667 Hern~indez, L. 556 Herrera, N.H. 96, 97, 119 Herrera, N.M. 143, 148 Herring, J.R. 246 Hess, H.H. 360, 365 Heubeck, C. 6, 29, 30, 50, 55, 57, 58, 214, 217, 220, 225, 235, 285, 295, 341,342, 615, 625, 663, 665, 667 Heubeck, C.E. 284, 286 Hickey-Vargas, R. 31 Higgins, G.E. 371,378, 387, 509, 556 Hilde, T.W.C. 394, 415, 416 Hildebrand, A.R. 107, 119 Hill, I.A. 394, 415 Hill, EJ. 587, 588 Hillhouse, J. 149 Hilst, R. 633, 667 Hine, A.C. 171,192 Hinz, K. 565, 577, 583, 588, 600, 625 Hirdes, W. 285 Hiscott, R. 286 Hobart, M.A. 387 Hogg, J.R. 588, 625
681
AUTHOR INDEX Holcombe, R.T. 587 Holcombe, T. 625, 667 Holcombe, T.H. 631,667 Holcombe, T.L. 6, 16, 30, 58, 118, 197-201, 217, 218, 357, 364, 365, 474, 561, 563, 584, 587, 588, 593, 597, 602, 625, 628, 639, 653, 663, 667, 668 Holden, R.C. 124, 148 Hon, K. 588 Hook, S.C. 388 Hoorn, C. 617, 618, 621,626 Hopkins, H.R. 563, 583, 588, 649, 650, 667 Hopson, C.A. 127, 128, 130, 134, 148, 149 Hor~i6ek, J. 120 Horne, G.S. 118, 155, 164, 165, 198, 199, 217, 219, 224, 235 Hottinger, L. 159, 165 Houlgatte, E. 344, 362, 365 Hou~a, V. 101, 103, 105, 119 Houtz, R. 58, 668 Houtz, R.E. 416, 564, 588, 589, 597, 626, 669 Houtz, R.Z. 631,667 Howell, D. 94, 119 Howell, D.G. 144, 148, 149, 388 Hoyer, M. 667 Hsti, K.J. 289, 338, 341,675, 676 Huang, T. 475 Huang, Z. 58, 118 Hubbard, D. 365 Hubbard, D.K. 345, 351,354, 357, 365 Huggett, Q. 474 Hugh, K.E. 151,165, 218 Hull, D. 150 Hull, D.M. 120, 127-129, 134, 148-150 Husler, G. 668 Husler, J. 588 Hussong, D.M. 397, 416 Hutson, E 19, 28, 30, 57, 111, 112, 118, 119 Huyghe, E 8, 14, 370, 371, 377-379, 387 Imlay, R.W. 98, 103, 111,119, 126, 127, 130, 133-137, 139-141, 143, 146, 147, 149 Ingersoll, R.V. 116, 119, 275, 285, 286 Ingle, J.C., Jr. 29 Inman, K.E 285 Iqbal, J. 542, 557 Irving, E.M. 155, 165, 219, 235 Isea, A. 473 Iturralde de Arozena, J.M. 556 Iturralde, J. 494 Iturralde-Vinent, M. 57, 118, 119, 168, 169, 192 Iturralde-Vinent, M.A. 6, 12, 28, 29, 81, 83, 87, 88, 90, 93, 94, 96, 97, 100, 104-106, 108-110, 112, 116, 117, 118, 119, 144, 147, 149, 277, 286 Ivey, M.L. 219, 223, 235 Jackson, H.R. 565, 582, 588 Jackson, J.B.C. 29 Jackson, T. 30
Jacobsen, S.B. 119 Jaffrezo, M. 164 Jankowsky, W.J. 447, 475, 551,556 Jansma, E 29, 285, 341, 362, 364, 556, 667 Jany, I. 13, 30, 361-363, 365, 628, 635, 662, 663, 667-669 Jaquin, T. 475 Jarrard, R.D. 416 Jean-Poix, C. 669 Johnson, G.L. 474 Johnson, H.R. 415, 416, 626 Johnson, W. 667 Jones, D. 119 Jones, D.L. 124, 148, 149 Jones, EC. 120 Joran, J.L. 415 Jordan, T.E. 435, 474, 475 Jordan, T.H. 220, 235, 370, 387 Joyce, J. 248, 249, 286, 360, 365 Judoley, K.M. 98, 103, 119 Jurgens, A. 192 Kafka, A.L. 59, 236, 286, 388, 651,654, 667, 669 Kahle, H.G. 667 Kanamori, H. 200, 217, 394, 416 Kaneps, A.G. 246 Kaniuth, K. 667 Karig, D.E. 394, 397, 416 Karner, G.D. 394, 415, 416, 588, 625 Karson, J.A. 565,588 Kasper, D.C. 618, 621,622, 626 Kay, R. 587 Kearey, E 398, 403,413,416 Keen, C.E. 565, 582, 583, 588, 589 Keleba, E 476 Kelldorf, M. 120 Kelldorf, M.E. 120, 150 Kellogg, J.N. 4, 16, 23, 30, 617, 626, 662, 663, 667-669 Kelsey, H.M. 217 Kendall, A.C. 289, 291, 334, 335, 337, 338, 341 Kent, D.V. 36, 51, 52, 58, 341 Kent, G.M. 565,588 Kenyon, N. 669 Kenyon, N.H. 387 Kerr, A.C. 5, 16, 19, 21, 28, 30 Kesler, S. 279, 286 Keszthelyi, L. 588 Khain, V.E. 121 Khudoley, K.M. 93, 97, 106, 112, 119, 141,149, 358, 365 Kidd, W.S.E 117, 118 Kieckhefer, R. 415 Kieft, C. 587 Kiessling, W. 126, 149 King, A.E 165, 217 Kinoshita, E.M. 441,474 Kirkland, D.W. 288, 341 Klaus, A. 58 Klaver, G. 30, 587 Klaver, G.T. 562, 588, 627, 668 Klenk, C.D. 217, 235 Klitgord, K. 591,626 Klitgord, K.D. 36, 41, 42, 58, 108, 112, 119, 168, 183, 187, 188, 191,192
Knepp, R.A. 285 Kolarsky, R.A. 8, 9, 12, 15, 30 Kolla, V. 565, 581, 588, 597, 617, 626, 631,668 Komara, S. 476 Korgen, B.J. 415 Kouroma, S. 476 Koutsoukos, E.A.M. 509, 556 Kozuch, M.J. 155, 165, 198, 199, 201, 214, 217, 222, 235 Krantz, R.W. 185, 188, 192 Krijnen, J. 244, 246 Kring, D.A. 119 Kroenke, L. 589 Kroenke, L.W. 588 Kroonenberg, S.B. 617, 626 Krop~i6ek, V. 120 Kruse, S.E. 385, 386, 387 Ku, T. 58, 192, 286 Ku, T.L. 31,342, 668 Kugler, H.G. 425, 475, 483, 488, 494, 501, 505, 507, 509, 514, 526, 535-537, 539, 544, 556 Kulstad, R. 342 Kutek, J. 98, 101, 103, 119, 141, 143, 149 Labaume, E 385, 387 LaBrecque, J. 59, 218, 416, LaBrecque, J.L. 58, 120, 475 Ladd, J. 193, 625, 668 Ladd, J.W. 6, 7, 13-16, 30, 35, 50, 51, 58, 167, 168, 171, 183, 188, 192, 388, 476, 563, 577, 583, 588, 593, 626, 627, 630, 633, 635, 641, 667, 668 Laine, E.E 611, 613, 622, 625 Lallemant, S. 387 Lallemant, S.J.C. 372, 387 Lamar, M.E. 325, 330, 341 Lancelot, Y. 587, 624 Land, L. 217 Land, L.S. 365 Langford, R.E 341 Langseth, M. 387 Langseth, M.G. 371,372, 387 Lara, M.E. 12, 23, 30 Larroque, C. 387 Larson K. 387 Larson, R.L. 168, 192, 587, 624 Larue, D. 476 Larue, D.K. 360, 362, 365, 618, 621, 622, 626 Latreille, M. 473 Lawrence, D.E 213, 217 Lawrence Edwards, R. 668 Lawrence, S.R. 635, 638, 668 Lawver, L.A. 29, 59, 415 Le Pichon, X. 380, 382, 387 Le Quellec, P. 387, 388, 668 Lebr6n, M. 342 L6bron, M.C. 19, 28, 30, 635, 668 Leckie, M. 589 Leckie, R.M. 626 Ledgerwood, R.K. 589 Lee, C.S. 397, 414, 416 Lee, T.-Y. 29 Lehner, E 662, 668
682 Leonard, R. 371, 387, 425, 475, 503, 542, 556 Lepvrier, C. 669 Leroy, S. 8, 13, 16, 29, 30, 43, 51, 53, 56, 199, 218, 565, 580, 588, 592, 615, 625, 626, 628, 630, 632, 637, 649-651, 653, 654, 659, 663, 665, 667-669 Letouzey, E 219, 235, 435, 475 Leturmy, E 386, 387, 388 Lewis, J. 30, 286, 588 Lewis, J.A. 149, 246 Lewis, J.E 30, 31, 94, 119, 135, 141, 149, 168, 169, 192, 246, 285, 342, 668 Liddle, R.A. 421,475 Lidz, B.H. 345, 347, 349-354, 356, 357, 365 Lillie, J.R. 385, 387 Lilliu, A.G. 435, 437, 439, 461,475 Linares Cala, E. 120 Linares, E. 31 Lindberg, EA. 285 Lindh, T. 58 Link, M.H. 283, 286, 493,494 Lithgow, C. 360, 363, 365 Llinas, R. 246 Llinas, R.A. 289, 292, 296, 311,341 Logan, B.W. 289, 333, 337, 341 Long, E 588 Long, R.E. 415 Longoria, J.E 120, 124, 126, 140, 141, 145, 149, 150 L6pez Quintero, J.O. 119 L6pez Rivera, J.G. 97, 98, 119 L6pez-Casillas, A. 162 L6pez-Ramos, E. 109, 111, 118, 119, 146, 149, 164, 217, 235 Lorente, M.A. 626 Lowenstein, T.K. 289, 337, 341 Lowrie, W. 475 Lozej, G.E 153, 154, 159, 165 Lu, R.S. 415, 416, 565, 581, 588, 597, 611,617, 626 Ludwig, W.J. 397, 405, 407, 411, 416, 564, 588, 597, 626, 631,667 Lugo, J. 8, 15, 19,21,30, 115,119,423, 455,473, 475, 552, 556, 618, 626 Lundgen, P. 669 Lundgren, E 218, 236 Lundgren, ER. 663, 668 Luyendyk, B.E 148 Lynch, L.L. 388 Lynts, G.W. 167, 192 Maaskant, R 587 Macdonald, R. 509, 557 MacDonald, K.C. 200, 218 MacDonald, W.D. 23, 29, 30, 118, 217, 220, 235, 494, 625 Macellari, C. 477, 494 Macelli, C.E. 617, 626 MacGillavry, H.J. 669 Mackenzie, G. 192 MacLeod, N. 150 MacPhee, R.D.E. 277, 286 Mahoney, J. 587 Mahoney, J.J. 589
AUTHOR INDEX Makino, M. 416 Malav6, G. 662, 668 Malavielle, J. 387 Malfait, B.T. 6, 16, 30, 168, 192, 591, 626 Maloney, N.J. 494 Mandl, G. 185, 192 Mann, J.E 625 Mann, E 3, 4, 6-8, 12-16, 18, 19, 21, 23, 28, 29-31, 50-53, 55-57, 57-59, 94, 115, 118, 119, 148, 170, 188, 189, 192, 200, 201, 214, 217, 218, 219, 220, 222, 225, 234, 235, 241, 245, 246, 247-249, 251, 253-255, 257, 259, 267, 273, 279, 281-284, 285, 286, 289-297, 299-302, 307, 311, 317, 320, 341, 342, 360, 362, 363, 364, 365, 423, 455, 474, 475, 477, 494, 498, 499, 552, 556, 615, 618, 625, 626, 628, 633, 635, 637, 638, 663, 665, 666, 667-669 Manton, R.S. 224, 235 Manton, W.I. 5, 6, 8, 13, 18, 19, 21, 31, 198, 199, 201, 218, 219, 222-226, 235 Mao, A. 51, 58 Mariner, R.H. 342 Marinho, M. 58 Mariotti, A. 387 Marriner, G.E 30 Marshall, J.S. 663,667, 668 Marshall, M.C. 150 Martin, C. 120 Martin, C.B. 133, 135, 137, 149 Martin, R.G. 29, 58, 118, 191,364, 474 Martfn-Bellizia, C. 439, 475 Martinez, R. 474 Martfnez, D. 95, 96, 119 Martinez-Cortez, A. 148 Marton, G. 64, 67, 68, 77, 82, 83, 89, 90, 108-110, 112, 119, 668 Marton, G.L. 65, 67, 68, 74, 77, 82, 83, 89, 90 Masaferro, J. 18, 19, 21, 23, 168, 188-190, 192 Maschenkov, S. 35, 36, 58 Mascle, A. 57, 364, 365, 370-372, 374, 377, 387, 388, 435, 475, 476, 587, 625, 626, 663, 666, 668 Mascle, G. 387 Mascle, J. 35, 58 Masclr, A.J. 164 Masschenkov, S. 668 Mass6, L. 387 Masson, D. 360-362, 365 Masson, D.G. 12, 31,357, 365, 633,668 Masvall, J. 476 Mata, S. 474, 475 Mathieu, Y. 364 Matta, S. 474 Matthews, J.E. 217, 586, 588 Mattinson, J.M. 31,148 Mattson, EH. 168, 192 Mauffret, A. 16, 29, 30, 31, 43, 51, 53, 56, 57-59, 218, 362, 365, 388, 476, 589, 592, 615, 626, 628, 630, 632, 635, 637, 649-651, 653, 654, 659, 663, 665, 667-669
Mauk, EJ. 150 Maurasse, E 627, 633, 635, 668 Maurrasse, E 31,576, 588 Maury, R. 30 Mayer, L.A. 193, 588, 589 Maync, W. 165 McArthur, J.M. 233,235 McBirney, A.R. 157, 165, 199-203, 205, 206, 208, 213, 217, 218, 219, 223, 235,236 McCabe, R.J. 394, 414, 416 McCaffrey, R. 213,218 McCann, W. 236 McCann, W.R. 4, 14, 31, 57, 59, 286, 360, 363, 364, 365, 388, 403, 413, 416, 667, 669 McCarthy, J. 565,588 McCave, I.N. 622, 623, 626 McClay, K. 379, 385, 386, 388, 492, 493,494 McCrevey, J.A. 120 McDougall, I. 213, 218 McGrew, EO. 199, 218 McKenzie, D. 53, 58, 583, 584, 589, 613, 626 McLaughlin, EP. 345-347, 350, 352, 353, 355, 356, 365, 668 McLaughlin, EE, Jr. 288, 289, 291,295, 297, 301, 304, 305, 307, 311, 317, 320, 330, 334, 335, 338, 342 McMillen, K.J. 565, 581,588, 597, 611, 617, 626 McNally, K. 31 McNulty, C.L. 192 McWilliams, M.O. 149 Meijer, ET. 51, 58, 663, 665, 668 Melia, EJ. 415 Melillo, A. 192 Melson, W. 587 Mendoza, V. 439, 475 Menendez V., A. 118 Meng, X. 120, 133, 134, 137, 149 Mercier de L6pinay, B. 30, 31, 57, 58, 59, 199, 218, 217, 248, 279, 281, 283, 285, 365, 628, 635, 659, 662, 663, 668, 667, 668 Mercier, J. 388 Merrick, K.A. 509, 556 Meschede, M. 213, 218, 588 Meschede, W. 30 Meyerhoff, A.A. 90, 93, 97, 98, 106, 112, 119, 120, 141, 149, 167, 168, 191,192, 365 Meyerhoff, H.A. 357, 358, 365 Meyers, J. 588 Miall, A.D. 435,475 Michaud, F. 152, 164 Michelson, J.E. 503,556 Middleton, G.V. 274, 286 Miles, ER. 589 Millan, G. 31 Milhin, E. 121 Milhin, G. 93, 111, 119 Millegan, ES. 29, 415 Miller, D.E. 416, 626 Mills, R.A. 151-153, 157, 165, 198, 199, 218 Minshull, T.A. 588
AUTHOR INDEX Minster, J.B. 220, 235 Mitacchione, V. 474 Mitchum, R.M. 464, 475, 476 Mocquet, A. 668 Moiola, R.J. 267, 282, 286 Molnar, E 3, 31, 38, 52, 58, 59, 200, 218, 477, 494, 499, 505, 556, 651, 668
Monechi, S. 30, 58, 285, 286, 341 Montadert, L. 587, 625 Montero, W. 30, 667 Montgomery, H. 120, 149, 241-243, 245, 246 Montgomery, H.A. 18, 19, 28, 124, 128, 144, 147, 149 Moore, C. 356, 365 Moore, C.H. 356, 365 Moore, C.H., Jr. 365 Moore, D.G. 415 Moore, G.F 386 Moore, J.C. 371,385,387, 476 Moore, R.E 415 Moore, W.S. 631,667 Morijini, R. 416 Morin, K.M. 241,246 Morley, C.K. 82, 90 Morton, J.L. 588 Mosher, D.C. 588 Mossakovskiy, A. 96, 119 Mountain, G.S. 171,192 Mrozowski, C.L. 416 Muehlberger, W.R. 30, 152, 164, 201, 217, 219, 223, 225, 235 Muff, R. 249, 258, 285, 286 Mugnier, J.L. 373, 374, 379, 385, 386, 387, 388
Muir, J.M. 146, 149 Mukhopadhyay, M. 390, 414, 416 Muller, C. 364, 625, 666 Muller, M.R. 565, 588 Miiller, R.D. 3, 18, 19, 21, 23, 37, 39, 48, 49, 52, 55, 56, 59, 627, 663, 668 Mullins, H. 193, 249, 285, 286, 341 Mullins, H.T. 29, 167, 168, 183, 188, 190, 192, 282, 285 Multer, H.G. 345, 347, 353, 354, 365 Mulugetta, G. 382, 388 Mufioz, I. 149 Mufioz, I.M. 148, 149, 246 Mufioz, M.I. 494 Munoz, N.G. 556 Munro, S.E. 478, 479, 492, 494, 501, 556
Murany, E.E. 437, 475 Murchey, B.L. 127, 128, 148 Murchison, R.R. 217 Murphy, A.J. 364 Murphy, M.T. 588 Murray, G.E. 146, 149 Musgrave, J. 217 Musgrave, R. 589 Mutter, J.C. 565, 577, 583, 585, 585, 586, 588, 600, 615, 626 Mutti, E. 267, 282, 286 Myczyfiski, R. 93, 98-103, 105, 111, 113, 114, 118-120, 135, 141, 143, 144, 149 Myers, C.W. 589
683 Nafe, J.E. 416 Nagle, E 242, 243, 246, 248, 249, 255, 272, 273, 279, 281,285, 286 Najmuddin, I.J. 77, 90 Nakatsuka, T. 416 Neathery, T.L. 91 Needham, H.D. 415 Nely, G. 388 Nemec, M.C. 358, 365 Nestell, M.K. 139, 149 Nettleton, L.L. 175, 192 Neumann, A.C. 171,192 Neumann, M. 159, 162, 165 Newport, R.L. 128, 149 Newton, M.S. 494 Nilsen, T.H. 282, 286 Nivia, A. 30 Norconsult 289, 296 Norell, M.A. 100, 119 Normark, W.R. 267, 286 O'Nions, R.K. 589 Obando, J.A. 29 Odehnal, M. 475 Officer, C.B. 398, 405, 412, 415, 416, 593, 626 Ogawa, Y. 58 Ogden, J.C. 365 Ogg, J.G. 58, 118, 120, 130, 149 Okal, E.A. 388, 557 Okuma, S. 397, 416 Olivet, J.-L. 31, 57, 59, 388, 476 O16ritz, F. 150 Olson, E.C. 218 Onstott, T.C. 31 Orange, D. 192 Ori, G.G. 8, 31,369, 370, 385,388, 489, 494
Orrego, A. 589 Ortega-Guti6rrez, F. 150 Osburn, W.L. 183, 192 Osiecki, ES. 201,218 Palacas, J.G. 90 Palmer, H.C. 248, 281,284, 286 Palmer, V.W. 463,475 Pardo, G. 93, 96, 120, 169, 192 Pargo-Casas, E 668 Paris, E 476 Parnaud, F. 423, 455, 474, 475, 501, 504, 505, 546, 549, 551,556 Parnell, J. 313,342 Parra, M. 387 Parson, L.M. 29, 364, 667 Pascal, J.C. 475, 556 Passalacqua, H. 423,435,439, 475, 546, 556
Patterson, J.M. 476 Paul, C.K. 474 Paulus, EJ. 167, 192 Payne, N. 506, 507, 509, 511, 512, 519, 524, 531,556 Pelaez, M. 285 Pelton, C.D. 387 Pena, L. 286 Penfield, G.T. 119 Pennington, W. 666 Pennington, W.D. 4, 14, 31,663,668
Peralta-Villar, J. 249, 272, 286 Perch-Nielsen, K. 164 Pereira, J.G. 474 P6rez Lazo, J. 109, 110, 115, 120 P6rez, O. 493, 494 Perez, O.J. 499, 506, 556 Perez-Cruz, G. 476 Perfit, M.R. 19, 28, 30, 59, 200, 218, 635, 668 Persad, K.M. 425, 475, 498, 507, 509, 513, 538, 546, 549, 556, 557 Pessagno, E.A., Jr. 18, 19, 28, 103, 109, 111, 120, 124, 126-130, 133-135, 137, 140, 141, 144, 148-150, 245, 246
Peter, G. 371,379, 388, 416 Petit, J.E 208, 218 Petroconsultants 440, 446, 447, 475 Petruccione, L. 192 Phair, R.L. 74, 82, 90, 91 Pia, J. 159, 165 Picard, X. 494, 556 Picard-Cadillat, X. 30, 475 Pickering, K. 256, 257, 264, 266, 286 Pieri, M. 386, 388 Pigram, C.J. 587, 588 Pilger, R.H., Jr. 394, 415 Pilkington, M. 119 Pimentel, M.N. 474 Pindell, J. 505, 507, 509, 533, 539, 553, 555, 625
Pindell, J.L. 6, 7, 14, 16-19, 21, 23, 29, 31, 34, 35, 46, 49-52, 55, 56, 57, 59, 108, 111, 115, 116, 118, 120, 124, 149, 150, 152, 163, 165, 168, 169, 188, 191, 192, 197, 200, 217, 218, 219, 235, 239, 241-243, 246, 248, 249, 255, 257, 273, 279, 281-283, 286, 364, 389-391, 412, 411, 414, 415, 416, 423, 424, 447, 474, 475, 494, 498, 503, 539, 546, 557, 591, 592, 613, 615-618, 625, 626, 627, 635, 668 Pinet, ER. 198, 199, 201,215, 218, 219, 222, 223,234, 235 Piotrowska, K. 94-97, 120 Piotrowski, J. 112, 120 Pitman, W.C. 218 Pitman, W.C., III 59, 120, 416 Plafker, G. 9, 31, 51, 59, 218 Planke, S. 573, 583, 588 Plumley, EW. 365 Poehls, K.A. 394, 416 Pogrebitsky, Y. 35, 36, 58 Pons, J.C. 387 Popenoe, E 119, 192 Potter, D.E 583, 588 Potter, H. 509, 557 Powell, C.M. 31 Powers, S. 219, 235 Prasetyo, H. 388 Prentice, C. 286 Price, E. 388 Priest, S. 217 Prieto, R. 425, 436, 437, 450, 461,463, 474, 476, 503, 557 Project Idylhim members 387
AUTHOR INDEX
684 Protti, M. 12, 3I Protti, S. 663, 669 Prud'Homme, R. 387 Pszcz6tkowski, A. 19, 83, 87, 88, 90, 93-108, 111-117, 120118-120, 135, 141,143, 147, 149, 150 Pubellier, M. 30, 218, 637, 669 Pugaczewska, H. 99, 100, 120 Pujana, I. 126, 150 Pujol, C. 387 Puscharovsky, Yu. 96, 120 Pushcar, P. 217 Pushkar, P. 217, 235 Pyle, T. 90 Pyle, T.E. 109, 120, 191 Pyre, A. 421,475 Racheboeuf, P.R. 476 Radovsky, B. 542, 557 Raineri, R. 159, 165 Raitt, R.W. 415 Rambaran, V. 479, 494 Ramos, N. 217 Ramsay, D.C. 577, 583, 588 Ratschbacher, L. 218 Rayhorn, J.E. 415 Raymond, C.A. 58 Redmond, B.T. 253, 266, 277, 286 Reed, D.L. 15, 31 Reed, K.M. 128, 129, 148 Reid, H. 360, 365 Reid, I. 565,588 Reid, J.A. 365 Remane, J. 148 Renard, V. 365, 387, 668, 669 Renne, P. 19, 30, 31, 57, 115, 118, 119 Renne, P.R. 111,120 Renz, H.H. 421,476 Renz, 0. 192 Reyes, J.R. 668 Ricchi Lucchi, F. 330, 335, 336, 342 Ricci Lucchi, E 257, 282, 286 Richard, P. 185, 188, 192 Richards, M.L. 222, 235 Ricou, L.E. 474 Rigassi-Studer, D. 93, 95, 96, 120 Ritchie, A. 217 Ritchie, A.W. 155, 164, 165 Roberts, M.T. 494 Roberts, R.J. 155, 165, 219, 235 Robertson, P. 371, 388, 478, 494, 500, 501, 506, 507, 510, 531, 540, 541, 545, 546, 551,553, 557 Robertson, R.P. 388 Robinson, C.J. 588 Robir, R.M. 476 Rodriguez, I. 618, 625 Rodriguez, J. 342 Rodriguez, P. 95, 121 Roest, W.R. 37, 48, 58, 59, 589, 668 Rogers, J.J.W. 587 Rogers, R.D. 152, 153, 155, 157, 161-163, 165 Rohr, G.M. 503, 557 Rona, P.A. 36, 59 Roque-Marrero, E 168, 192 Rosales, H. 57, 666 Rosencrantz, E. 5-7, 12, 13, 16, 23, 28,
29, 31, 35, 43, 52, 55, 56, 59, 94, 95, 109, 111, 121, 198, 200, 211, 218, 219, 222, 234, 235, 248, 283, 286, 360, 363, 365, 499, 557, 564, 565, 581, 587, 617, 625, 628, 635, 637, 650, 663, 666, 669 Rosendahl, B.R. 81, 90, 565, 566, 581, 582, 588 Rosenfeld, J.H. 12, 31, 52, 59 Ross, C.P. 286 Ross, M. 168, 183, 188, 189, 192 Ross, M.I. 16, 31, 34, 52, 59, 108, 111, 115, 116, 121, 218, 235, 286, 389, 416, 423,476 Rossello, E.A. 29 Rossi, S. 626 Rossi, T. 435, 447, 476 Rotstein, Y. 588, 626 Roure, E 369, 388, 420, 435, 475, 476, 501,503, 505, 546, 551,556, 557 Rowe, D.W. 557 Rowett, C.L. 124, 150 Rowley, D.B. 59, 120, 218, 416 Rowley, K.C. 388, 557 Royden, L.H. 55, 59, 385, 386, 387 Royer, J.-Y. 35, 38-40, 59 Royer, J.Y. 668 Ruddiman, W.E 625, 667 Russo, R. 388 Russo, R.M. 371, 380, 388, 499, 506, 545, 557, 633, 635, 663, 668, 669 Russomano, F. 473 Ryabukhin, A.G. 112, 121 Ryan, W.B.E 338,341,342, 625, 667 Ryberg, P.T. 285 Sadybakasov, E. 29 Sager, W. 416 Saint-Marc, P. 162, 165, 285, 667 Saito, T. 669 Saleh, J. 388 Salvador, A. 73, 83, 87, 88, 90, 124, 139, 150, 421, 476, 478, 479, 494, 506, 557 Sams, R.H. 440, 450, 474, 476 Sanchez, H. 473 Sandoval, J. 126, 150 Sandwell, D.T. 4, 8, 9, 31, 37, 38, 41, 42, 59 Santiago, N. 474, 475 Sapper, K. 219, 235 Sares, S.W. 286 Sarg, J.E 333, 342 Sarmiento, G.A. 626 Sarzalejo, S. 474 Sass, L.C. 475 Sass, L.S. 474 Sassi, W. 369, 388 Saunders, A.D. 30 Saunders, J. 556 Saunders, J.B. 164, 346, 364, 371, 378, 387, 425, 474, 476, 509, 511, 519, 526, 556, 557, 625, 667 Scanlon, K.M. 12, 29-31,357, 360-363, 364, 365, 474, 633, 667, 668 Scanlon, N. 191 Scasso, R. 126, 149 Schaaf, A. 285, 667 Schaefer, C.T. 307,342
Schaming, M. 588, 626 Schellekens, H. 30 Schellekens, J.H. 149, 246, 365 Schermer, E. 119 Schermer, E.R. 148 Schilling, J.G. 415 Schlager, W. 64, 65, 68, 69, 74, 77, 82, 90, 91, 167, 168, 188, 190, 192, 282, 286 Schlapak, G. 447,475, 551,556 Schlich, R. 583, 586, 588, 600, 626 Schluter, H.-U. 588 Schmidt, V.A. 108, 112, 118, 124, 148, 152, 163, 164, 164 Schmitt, R. 588, 668 Scholl, D.W. 364 Schoomaker, J.E. 476 Schouten, H. 36, 37, 41, 42, 58, 59, 119, 591,626 Schreiber, B.C. 289, 291,311,330, 333, 336-338, 342 Schreiber, E. 342 Schroeder, R. 159, 162, 165 Schubert, C. 30, 31, 58, 477, 494, 499, 556, 557 Schumm, S.A. 618, 622, 626 Schwander, H. 556 Schwartz, D.P. 200, 218 Schwartz, R.K. 341 Schwartz, Y. 663, 669 Scientist Party of Leg ODP 387 Sclater, J.G. 31, 59, 200, 218, 235, 286 Scotese, C. 192 Scotese, C.R. 31, 34, 52, 59, 108, 111, 115, 116, 121, 168, 183, 188, 189, 192, 389, 416, 423,476 Scott, D. 588 Scott, G.R. 120 Scott, R.W. 5, 18, 19, 21, 155, 157, 159, 162, 163, 165 Sebrier, M. 58 Sedlock, R.L. 131,144, 145, 150 Seely, D.R. 366 Segura Soto, R. 96, 121 Self, S. 6, 9, 29, 586, 588 Sen, G. 5, 31 Sen Gupta, B.K. 305, 307, 342 Seng6r, A.M.C. 53, 57, 289, 341,667 Seng6r, C. 57 Shaffer, B.L. 165 Shafiquallah, M. 235 Shafiqullah, M. 217 Shagam, R. 118, 625 Shanmugan, G. 267, 282, 286 Sharp, W.D. 557 Shaub, F.J. 65, 74, 83, 91,148 Shaw, P.R. 37, 39, 41, 42, 59 Shepard, F.P. 476 Shepherd, J. 557 Shepherd, J.B. 499, 500, 557 Sheridan, R. 190, 193 Sheridan, R.E. 167, 171, 183, 188, 192, 193 Shih, T.C. 668 Shih, T.T. 371,388 Shipboard Scientific Party 90 Shipley, T.H. 171, 193, 386, 587, 588, 624
AUTHOR INDEX Shipunov, S.V. 118 Shiroma, J. 29, 285, 341 Shor, G. 416 Shor, G.G. 415 Shouten, H. 192 Shurbet, G.L. 345, 347, 365 Sick, M. 30, 588 Sigurdsson, H. 600, 602, 613, 621,626 Silberling, N.J. 149 Silver, E.A. 15, 31,372, 388, 653,669 Simmons, M.D. 164 Simonson, B.M. 155, 165 Sinton, C. 667 Sinton, C.W. 5, 17, 21, 28, 31,562, 588, 597, 626 Sinton, J.M. 587 Six, W.M. 150 Six, W.M., Jr. 150 Skogseid, J. 586, 588 Skupinski, A. 120 Sleep, N.H. 394, 416, 588 Sliter, W.V. 57, 118, 589 Slootweg, A.E 58 Smith, A.G. 148, 475 Smith, A.L. 30 Smith, D.G. 475 Smith, ED. 474 Smith, ED., Jr. 478, 479, 492, 494, 501, 556 Smith, J. 192 Smith, M.J. 372, 388 Smith, E 150 Smith, EL. 31, 126, 150 Smith, W.H.F. 8, 9, 31, 37, 38, 41, 42, 48, 49, 55, 59, 627, 668 Snelson, S. 435,474 Snoke, A.W. 503, 557 Sobolev, A. 415 Somin, M. 31 Soruco, R.S. 626 Soulas, J.E 477, 494 Southernwood, R. 157, 165, 226, 236 Southernwood, S. 198, 199, 218 Spadea, E 562, 589 Spano, E 435,474 Speck, R. 286 Speed, R.C. 6, 14, 31, 150, 345, 360, 362, 365, 370, 380, 388, 394, 397, 398, 405, 409, 412, 413, 416, 435, 438, 476, 478, 494, 499, 506, 535, 536, 539, 545, 553, 557 Sprague, A.R. 448, 476 Srivastava, S.E 565, 581,588, 589 Stagg, H.M.J. 587, 588 Stainforth, R.M. 421, 476, 478, 479, 494, 506, 557 Stber, K. 667 Stein, C. 218, 236, 669 Stein, S. 29, 58, 217, 218, 220, 235, 236, 285, 341, 370, 387, 388, 476, 556, 663, 669 Steiner, M.B. 199, 218 Steinmetz, J.C. 192 Stephan, E 365 Stephan, J.E 360-363, 365, 370, 388, 423, 424, 476, 668 St6phan, J.E 7, 31, 34, 56, 57, St6phan, J.E 59
685 Stewart, G.S. 200, 217 Stinnesbeck, W. 148 Stock, J. 38, 52, 58, 59 Stoffa, E 588, 650, 651,668, 669 Stoffa, EL. 58, 193, 563, 583, 587, 593, 597, 625, 626, 669 Storey, M. 588, 589, 626 Storti, F. 379, 385, 386, 388 Stow, D. 286 Straub, C. 667 Stride, A.H. 387 Suarez, G. 57, 662, 667, 668 Su~ez, G. 4, 29 Suayah, I.B. 165 Subieta, T. 476, 557 Suchanek, T.H. 365 Suczek, C.A. 279, 282, 285 Supko, ER. 494 Suppe, J. 375, 387, 388, 388 Surdam, R.C. 313,342 Sutter, J. 286 Swart, E 191 Swift, S.A. 588 Swolfs, H.C. 165, 218 Sykes, L. 3, 31,668 Sykes, L.R. 53, 59, 200, 217, 218, 236, 247, 279, 282, 286, 371, 403, 413, 416, 477, 494, 499, 556, 635, 651,663, 667, 669 Sykes, M.A. 476 Sylvester, R.E. 29, 191
589,
Toto, A. 662, 669 Tournon, J. 31, 57, 59, 388, 476 Trechmann, C.T. 509, 557 Tremel, H. 667 Truchan, M. 58, 587, 589, 625, 626, 668, 669 Truskowski, I. 475, 556 Tsai, C.J. 668 Tubb, S.G. 415 Tucholke, B.E. 37, 59, 171, 192, 622, 623, 626 Turn~ek, D. 162, 164 Twiss, R.J. 208, 218 Tyburski, S.A. 199, 201,218, 235 Tyson, L. 501,506, 510, 511,519, 531, 542, 544, 557 Umpleby, D. 588 Unternehr, E 475 Uyeda, S. 394, 397, 416
220, 388, 505,
Tabbutt, K.D. 616-619, 626 Tabor, S. 360, 365 Tajima, E 57 Talwani, M. 58, 563-565, 580, 588, 589, 602, 626, 630, 641,668, 669 Tamaki, K. 392, 394, 397, 416 Tanimoto, T. 587, 624 Tankard, A.J. 435,476, 617, 626 Tappmeyer, D.M. 217, 235 Tarduno, J.A. 584, 589 Tardy, M. 31, 57, 59, 164, 388, 476 Tarney, J. 30 Tavares, I. 246, 285 Taylor, B. 394, 415,416 Taylor, D. 29 Taylor, D.E. 191 Taylor, D.G. 124, 126, 150 Taylor, F. 58, 286 Taylor, F.W. 31, 192, 289, 292, 299, 315,342, 668 Tchekhovich, V.D. 121 Testarmata, M.M. 69, 91 Thery, J.M. 31, 57, 59, 388, 476 Thierry, J. 58, 118, 475 Thomas, J.C. 29 Thomas, W.A. 80, 91, 168, 183, 191 Thompson, G.A. 588 Thompson, S., III 476 Thordarson, T. 588 Tiley, G.J. 388 Tipper, H.W. 124, 150 Tomblin, J.E 389-392, 416 Tondji Biyo, J.J. 29 Torresan, M.E. 587 Torrini, R., Jr. 31 Toskoz, M.N. 394, 416
Vai, G.B. 330, 335, 336, 342 Vail, ER. 90, 165, 286, 341, 365, 464, 466, 471,474-476, 556 Valastro, S., Jr. 342 Valentine, EG. 474 Valery, E 379, 387, 388 van de Wiel, M. 626 van den Berghe, B. 295, 342 van den Bold, W.A. 341,347, 350, 358, 365, 364 Van Buren, H.M. 167, 190, 192 Van Fossen, M.C. 662, 669 van Gestel, J.P. 363, 365 Van der Hilst, R. 55, 57, 59, 663, 669 Van der Hilst, R.D. 4, 16, 31 Van der Voo, R. 124, 150 van Veen, E 58, 118 Van Wagoner, J.C. 154, 164, 464, 475 Vann, I.R. 296, 342 Vaughan, A.E 168, 193 Vaughan, T.W. 253, 286 V~izquez, M. 95, 96, 119 Veeken, EC.H. 447, 476 Vega, V. 4, 16, 23, 30, 617, 626, 663, 667, 668 Verg6s, J. 375, 387 Verhoef, J. 58 Verma, H.M. 135, 139, 150 Vernette, G. 387, 662, 663, 669 Vierbuchen, R.C. 192, 478, 494, 500, 557 Vignali, M. 556 Vila, J. 285, 669 Vila, J.M. 31, 57, 59, 246, 285, 388, 476, 635, 637, 668, 669 Villaroel, V. 420, 437, 476 Villasenor, A. 633, 635, 669 Villasefior, A. 493,494 Villeneuve, M. 441,476 Viniegra, O.E 109, 121 Vinour, E 388 Vitali, C. 628, 662, 669 Vivas, V. 447, 476 Vogt, E 387 Vogt, ER. 59, 217 Von der Hoya, H.A. 199, 201, 214, 215, 218, 223,224, 236
686 von Hillebrandt, A. 126, 150 Von Huene, R. 12, 31 Vrielynck, B. 474 Wadge, G. 30, 51, 56, 58, 59, 201,218, 307, 342, 499, 500, 509, 556, 557 Waggoner, D.G. 31 Walcott, R.I. 213, 218 Wald, D. 285, 341 Wald, D.J. 29 Walker, G.EL. 588 Walker, J.D. 199, 218 Walker, R.G. 282, 286 Walles, EE. 168, 169, 171,174, 193 Walper, J. 124, 150 Walper, J.L. 617, 618, 625 Wanneson, J. 587, 625 Warburton, J. 489, 494 Ward, S.N. 51, 59 Warner, A.J. 358, 366 Warner, A.J., Jr. 398, 416 Warren, J.K. 288, 337, 338,342 Watanabe, T. 416 Waters, A.C. 573,589 Watkins, J. 589 Watkins, J.S. 148, 563, 577, 583, 588, 593, 626, 627, 630, 633, 635, 668 Watson, M. 286 Weaver, C.E. 155, 165 Webb, E 388 Weber, J. 388 Weber, J.C. 371,388 Weibord, N.E. 476 Weidmann, J. 562, 589 Weidner, D.J. 651,654, 667 Weiland, T.J. 151,155, 157, 161,165 Weissel, J.K. 394, 397, 415,416
AUTHOR INDEX Wellner, R. 192 Welsink, H.J. 626 Westall, E 622, 626 Westbrook, G. 58, 625, 667, 668 Westbrook, G.K. 6, 14, 16, 30, 31, 58, 370-372, 378, 379, 382, 386-388, 394, 397, 398, 405, 409, 412, 413, 415, 416, 476, 668 Westercamp, D. 30, 364, 370, 387, 438, 474 Westerfield, J.R. 365 Westermann, G.E.G. 126, 135, 139, 150 Whalen, EA. 126, 150 Whetten, J.T. 344, 358, 360, 366 White, A.H. 313, 342 White, R.A. 200, 201,218 White, R.J. 164 White, R.S. 565, 583, 584, 588, 589, 613, 626 Whitmarsh, R.B. 565, 566, 589 Whittaker, J.E. 164 Wiens, D. 218 Wiens, D.A. 57, 236, 388, 669 Wierzbowski, A. 98, 100, 101, 111, 119, 121, 143, 149, 150 Wilcox, R.E. 357, 366 Wildberg, H. 562, 589 Wildermann, E. 667 Williams, C.A. 382, 388 Williams, G.D. 8, 31 Williams, H. 157, 165, 199, 218, 219, 236 Wilson, C.C. 489, 494, 501, 505, 544, 557 Wilson, H.H. 213, 218 Wilson, J.T. 16, 31, 591,626 Windisch, C. 387, 589, 626
Windisch, C.C. 669 Winker, C.D. 289, 342 Winston, G.O. 90 Winterer, E.L. 193, 588, 589 Wohletz, K. 217 Wooden, J.L. 201,218, 307, 342 Woodhouse, J.H. 217 Woodring, W.P. 286 Woods, D.E 218, 236, 669 Woodside, J. 588 Wornardt, W.W. 464, 466, 476 Wortel, R. 23, 31, 51, 58, 663, 665, 667 Worthington, L.B. 626 Worthington, L.V. 622, 626 Worzel, J.L. 148, 365 Wright, A. 371,388 Wright, W.R. 622, 626 Wu, S. 437, 446, 476 Wunsch, C. 622, 626 Yang, Q. 150 Yeh, K. 126, 128, 130, 150 Yeh, K.Y. 150 Yepes, H. 58 Young, G. 57, 666 Young, K. 153, 155, 164 Youngs, B.C. 313,342 Yule, J.D. 557 Zehnder Mutter, C. 565, 589 Zhang, Y.-S. 587, 624 Zhao, W.L. 382, 388 Zoetemeijer, R. 369, 388 Zonenshein, L.P. 121 Zubieta, D. 386
Subject Index
3-Dimensional ellipsoids 40 3-Dimensional finite rotation 44 40Ar/39Ar measurements 69 At~ B" horizon561 A~t-B~ interval 633 A-subduction zones 435 Absolute plate motion 53 Absolute plate motion model 52 Abuillot Formation 253 Abyssal current circulation 622 Acapulco-Guatemala megashear 164 Active margin stage 93, 117 Africa 18, 19, 35, 183 Africa plate 14 African margin 566 Afro-South America plate 65 Agalteca quadrangle 153, 154 Aggradational configuration 459 Agua Frfa Formation 155 Agu~in 211,220, 222 Agu~in beds 222, 227, 234 Agu~in fault 201, 219 Agu~in Valley 219, 220, 222, 224, 226, 227, 228, 233 Airport/Penitentiary outcrop 352, 356 Airport/Penitentiary roadcut 351 Airport/Penitentiary section 349-352 Alamino cruise 629 Alaska 126, 141 Alluvial-deltaic phase 435 Altamira 253, 257, 259, 260, 273 Altamira block 249, 253,257, 272, 283 Altamira facies 253 Altamira fault zone 251,259, 270 Altamira Formation 247, 251,253, 255, 257, 258, 259, 261,263, 264, 266, 267, 270, 272, 274-277, 279, 281-284 Alturas de Pizarras del Sur 95 Amaime-Chaucha Terrane 617 Amaro Formation 107, 116 Amazon 591,624 Amazon delta 621 Amazon drainage system 621,622 Amazon fan 621 Ammonite-bearing limestones 93 Ammonites 99 Amoco Production Co. 152 Anaco fault 420 Anaco inversion structure 437 Anaco trend 450 Anc6n Formation 107, 108, 117 Andaman basin 390, 391,394, 397, 414 Andean Central Cordillera 618
Andean Cordillera 617 Andes 126, 591 Andros Bank 188 Anegada fault zone 13,362, 363 Anegada Passage 343, 357, 358, 360, 362, 363, 622, 623, 663 Angostura block 301,302 Angostura Formation 302, 304, 305, 315, 336, 340 Anguilla 354, 358 Anguilla/Saba Bank 358 Antarctic Bottom Water 622 Antarctic Intermediate Water 622, 623 Antarctica 52, 111, 126 Antillean island arc 144, 147, 435 Apalachicola basin 83, 88 Apennines 336 Appalachian foldbelt 80 Appalachian-Ouachita orogenic belt 19 Aptian-Albian shallow marine limestone 5 Apure thrust fault 441 Arabian shelf 163 Araya Peninsula 499, 503 Araya-Paria area 500 Araya-Paria Peninsula 503,509, 548, 553 Arc terranes 4 Arctic 582 Areo Formation 450, 473 Areo shales 450 Argentina 126, 126 Arima fault 492 Arroyo Barranca 327, 328 Arroyo Barranca section 333, 336, 337 Arroyo Barrero 325, 328 Arroyo Bermejo 237, 238 Arroyo Bermejo chert 243 Arroyo Berraco Blanco stratigraphic section 260 Arroyo Blanco Formation 287, 296, 300, 305, 317, 333-336, 340 Arroyo Boca de los Guiros 331 Arroyo de la Salvia 323 Arroyo del Pozo 317 Arroyo Guanabano 271 Arroyo Honduras 331 Arroyo Honduras-Facolina 331 Arroyo Las Lavas 267, 269 Arroyo Las Lavas section 271 Arroyo Seco Formation 287, 317, 320, 326, 340 Artemisa Formation 100, 102, 103, 109, 143 Aruba Gap 627-630, 650, 657
Atima Formation 151, 153, 156, 157, 159, 159, 163 Atima limestone 163 Atlantic basin 245 Atlantic floor 391 Atlantic fracture zones 18 Atlantic Ocean 14, 18, 168, 291,471, 499, 591,621,624 Atlantic oceanic crust 18, 439 Atlantic passive margin 437, 439 Atlantic passive margin of South America 435 Atlantic Plate 423 Atlantic-Indian hotspot reference system 33, 56 Atlantic-Indian hotspots 56 Atlantic-Indian mantle reference system 57 Atlantic-type fracture zones 37 Atlantic-type passive margin 420, 424 Atlantis Fracture Zone 33, 35 Atrato-San Juan basin 16 Australia 313, 333 Aves Ridge 6, 13, 14, 16, 389, 391,397, 398, 403, 405,409, 499, 600, 602, 623 Aves Ridge remnant arc 13 Aves Rise 593, 611 Aves Swell 391,397, 398 Avocado high 507, 513, 515, 516, 518, 520, 526, 528 Avocado-Couva High 482, 487, 488, 491,492 Azua basin 295, 296, 305, 317 Azua-Barahona highway 331 Azucar Member 143 Azuero-Son~i fault zone 15 Bt~ horizon 561 Back Rfo Grande 244 Back-arc basins 4, 6, 499 Back-arc rifting 499 Backstepping behavior 459 Backstepping transgressive phase 471 Backstop 551,555 Backstripping 335 Backthrusts 489 Bahfa de Neiba 292, 301 Bahfa de Neiba block 300 Bahfa Honda composite terrane 96 Bahfa Honda quadrangle 241 Bahfa Honda segment of the volcanic arc 93 Bahfa Honda terrane 94, 116 Bahama carbonate platform 247, 249
688 Bahama collision 283 Bahama Platform 13, 22, 52, 241,245, 279, 282-285 Bahamas 50 Bahamas archipelago 167, 168, 169, 188 Bahamas carbonate platform 168, 291 Bahamas Fracture Zone 65, 188 Bahamas passive margin 114, 117 Bahamas Platform 6, 18, 21, 22, 57, 93, 111, 116, 167, 168, 189, 336, 498, 499 Bahamas Platform margin 108 Bahoruco fault 301 Bahoruco fault zone 300, 304 Bahoruco Peninsula 627, 628, 637, 659, 660 Bahoruco thrust 663 Baja California Sur 126 Bakersfield 126 Banda Sea basin 391,394, 397 Bank margins 18 Barahona 312 Barbados 618, 621,622 Barbados accretionary wedge 438, 477, 489 Barbados Ridge accretionary complex 14 Barbados Ridge complex 8 Barbareta 197, 201,203 Barrackpore oil field 538 Barracuda Ridge 14, 48, 51, 55,403 Barranca section 326, 331 Barranquin section 447 Barrero section 320, 328 Basal foredeep 424 Basal foredeep unconformity 419, 435, 438, 447, 449, 457, 466, 471,472 Basement sites 68 Basin classification 7 Basin inversions 6 Basin sites 68 Basin-bounding faults 477 Basin-center evaporite deposits 287 Basin-edge evaporite deposits 287 Bass Collier Country 12-2 88 Bath 243 Bay Islands of Honduras 197, 206, 213, 219, 222 Bayano-Chucanqu6 basin 15 Beata compressional zone 627 Beata deformed zone 663 Beata fault system 637 Beata fault zone 300 Beata igneous province 643 Beata plateau 629, 635, 639, 653, 662 Beata Ridge 16, 28, 29, 33, 50, 51, 53, 561,564, 584, 591,592, 593, 599, 600, 615, 623, 624, 627, 628, 629, 633, 637, 639, 641,649, 650, 651, 657, 659, 665 Beata Ridge flank 580 Beata Ridge footwall 591 Beata seamounts 637 Beata volcanic plateau 627, 649, 659 Bel6n Vigoa tectonic unit 113 Belize 80, 145, 199 Benioff zone 499, 635 Bermeja accretionary complex 245
SUBJECT INDEX Bermeja Complex 128 Best fit reconstruction 39 Bimini Bank 188, 190 Bismark basin 394 Blake Bahama Basin 130, 147 Blake drift 623 Blake Plateau 437, 446 Blessing Formation 343, 344, 353 Blue Mountains 237, 238, 243 Blue Mountains block 243, 244 Blue Mountains inlier 244 Boca Wampti 159 Bocon6 50, 618 Bocon6 right-lateral fault 15 Bocon6-eastern Andean right-lateral strike-slip fault system 4 Bocond-Mordn-E1 Pilar fault systems 618 Bodo Verde 38 Bonacca Ridge 197, 201,206, 215, 222 Bonaire Block 618 Bonito Oriental 220, 231 Boreal ammonites 137 Boreal assemblage 126 Bouguer anomalies 55, 48, 389 Bove Basin 441 Brasso Formation 487, 488, 510, 513 Brasso shales 491 Brazil 19, 441,473 Breakup unconformity 435 Brute stacks 564 Bucaramanga segment 663 Bulk density 69 C. Victoria, Tamps. 123 Cabritos anticline 299 Cabritos- 1 well 302 Cacarajfcara Formation 106, 107, 116, 118 Cahuasas Formation 139, 141 Caigual fault zone 533 Calabaza measured section 259, 264 Calabaza section 260, 264, 266, 269, 282 Calc-alkaline volcanoes 14 California 7, 128 California Coast Ranges 126, 128, 144, 147 Camajuanf belt 111, 116, 117 Camajuanf succession 113 Campana High 487, 491 Campeche 145 Campeche Bank 74 Campeche Escarpment 64, 65, 69, 89 Camfi fault zone 253, 257, 272, 273, 282 Canada Bonita 258, 271 Canada Bonita anticline 263 Canada Bonita Member 258, 259, 260, 261,263, 264, 266 Canary-Bahamas Transect 35, 36 Cangre belt 93, 95, 109, 113, 117 Cangre tectonic unit 116 Cangrejal 224 Cangrejal River 220, 224 Canoa Formation 447 Cantarranas Formation 153, 153 Cation San Matias 135 Canyon fill 463
Canyon San Matias 137 Capas de San Pedro 135, 137 Caracas Group 477 Carapita Formation 424, 489, 491 Carapita overpressure shale belt 489 Caratas Formation 448 Carbonate buildups 63 Carbonate megaplatform 495, 513 Cariaco pull-apart basin 15, 500 Cariaco Trough 477, 478 Caribbean 6, 18, 124, 313, 592 Caribbean arc 503, 539, 552 Caribbean basalt province 561,585, 592, 593, 599, 613 Caribbean crust 561 Caribbean evolution 16 Caribbean large igneous province 564 Caribbean oceanic plateau 5, 13, 15, 498, 561,580, 600, 613 Caribbean oceanic plateau crust 27 Caribbean Plate 3, 4, 14, 28, 29, 35, 115, 144, 145, 147, 152, 189, 211,219, 241,245, 247, 248, 279, 282, 289, 291,304, 360, 362, 363, 391,403, 423,425, 435, 450, 455, 463, 471, 477, 495-498, 505, 552, 580, 586, 591,592, 615, 623, 627, 663 Caribbean Plate south of Hispaniola 591 Caribbean Plateau 583, 586, 600, 627 Caribbean Province 152, 153, 159, 161, 162, 164 Caribbean Sea 241,249, 287, 291,295, 301,498, 564, 628, 631,633 Caribbean slab 53 Caribbean Tertiary reef deposits 353 Caribbean-American collision 188 Caribbean-South American Plate boundary 477, 593, 617, 618, 623 Carmita Formation 106, 115 Caroni Basin 477, 489, 492, 493, 507, 553 Carrizal clastics 441 Carrizal Formation 441 Carrizal-2X well 441 Casanay fault 482, 488, 489, 492 Casanay-Arima 477 Cas-C03 651 Casis cruise 627, 628, 641 Casis data 630 Casis seismic survey 627 Castana No. 1 223 Castilla No. 1 223 Catoche Grid 65 Catoche Knoll 65, 74 Catoche Tongue 65, 69, 74, 83 Caucagua-Paracotos-Villa de Cura belt 505 Cay Sal 189 Cay Sal Bank 168, 190 Cayman passage 623 Cayman pull-apart 477 Cayman Ridge 13, 499 Cayman Rise 13 Cayman spreading center 222, 635, 637, 663 Cayman spreading center-Puerto Rico trench-Lesser Antilles subduction zone 666
689
SUBJECT INDEX Cayman transform 152 Cayman Trough 13, 33, 43, 52, 56, 197, 199, 201,222, 234, 248, 279, 283, 663 Cayman-Puerto Rico fault system 627 Cayo Coco area 168 CCC Test Well 39 349 CCC Test Well 45a 356 CCC Test Well C-26 354 C6baco basin complex 15 Cedros Formation 513, 524, 526 Celebes basin 397, 414 Cenozoic strike-slip phase 19 Central America 12, 112, 249, 623 Central American arc 592 Central American Isthmus 591,623, 624 Central Atlantic 188 Central block of Hispaniola 663 Central Brazilian Shield 617 Central Caribbean Plate 16 Central Cordillera 242, 591, 617, 620, 621,635 Central Cordillera arc of Hispaniola 615 Central Cordillera of Colombia 617 Central Cuba 6, 12, 93 Central fracture valleys 37 Central Mexico 109, 123, 144 Central Range of Trinidad 477, 488, 489, 491,492, 493, 495, 496, 498, 501,506, 509, 510, 523, 526, 528, 531,535, 536, 541,542, 544, 548, 551,554, 555 Central Range fault 493 Central Range fault zone 501,533 Central Range-Caigual fault zone 496, 498, 536, 554 Central Tethyan Province 126, 141 Central Tethyan successions 147 Central Texas 159 Central Venezuelan Basin 622 Central Venezuelan Fault Zone 566, 582 Central Yucat~in basin 13 Cerro de Cabras tectonic unit 95 Cerro de la Virgen Limestone 213 Cerro La Penita 226, 227, 228 Cerro Las Lomas 228 Cerro Panteon quarry unit 2 135, 136, 137 Cerro Wamp6 157 Chapulhuac~in 137 Chapulhuac~in Limestone 136, 137, 141 Charco Largo-1 well 289, 296, 297, 301, 304, 305, 307, 315, 316, 337, 339, 340 Chaudiere Formation 510 Chevron Corporation 65 Chiapas 145 Chiapas-YucaUin area 498 Chicxulub structure of Yucat~in 107 Chimana Formation 447 Choc6 Block 617 Chocoy No. 2 146 Chorotega block 152 Chortfs block 4, 6, 8, 19, 21, 151,155, 159, 197-201,211,213, 215 Cibao 249 Cibao basin 249, 251,288, 291,340 Cibao Valley 253, 334, 633, 635
Cinco de Mayo 145 Cinco Pesos 101 Cinco Pesos area 99 Cinco Pesos tectonic unit 107, 117 Cipero Formation 487, 488 Circum-Pacific margin 126 Ciudad Victoria (Tamps.) 139 Civilian Conservation Corps 345 Classical breakup unconformity 419 Coahuila 146 Coahuiltecana terrane 130, 133, 144 Coahuiltecano terrane 123, 130, 145, 146 Coast of Hispaniola 12 Coast Range Ophiolite 127, 128 Coast Ranges425 Coastal Fringe/Margarita belt 509, 528 Coastal Range/Margarita belt 503 Cochinos Sound 174 Cocos oceanic crust 663 Cocos plate 3, 22, 51, 56, 663, 666 Cocos Ridge 9, 12, 24 Collages 6 Collantes Formation 110 Collision 4 Collisional basins 8 Colombia 4, 16, 498, 617 Colombia River basalt flows 573 Colombian accretionary complex and forearc basin 16 Colombian basin 5, 16, 28, 33, 56, 564, 565, 577, 581,599, 611,617, 650, 651,659, 663 Colombian deformed belt 627, 662 Colombian microplate 51, 52, 627, 662, 663,666 Colombian Plate 33, 56, 592, 615, 637 Colombian Plate margin 57 Colombian prism 662 Colombian trench 22 Columbia University 392 Columbus Basin 425, 503, 542 Columbus channel 437 Columbus foredeep 501 Combined rotation model 52 Common isochron segment 39 Compressional satellite ('piggyback') basins 455 Confluencia Rfos Patuca y Wamp6 quadrangle 155 Conjugate isochrons 39 Conrad 577 Continent-ocean transition 551 Continental bridge 89 Continental encroachment cycle 471 Continental fragment 4 Continental lithospheric A-subduction 435 Continental shelf of Venezuela 407 Continental slope of Venezuela 413 Cooperaci6n Dominicana de Empresas Estatales 296, 312 Cordillera Central 287, 291,292, 295, 296, 301,307, 317 Cordillera de Guaniguanico 83, 87, 88 Cordillera de la Costa belt 503 Cordillera de la Costa Nappe 477 Cordillera Nombre de Dios 220, 223
Cordillera Septentrional 247, 248, 249, 251,267, 273, 279, 282, 283, 284 Cordillera system 621 Coriolis force 622 Corpoven S.A. 420 Costa Rica 5, 9, 599 Costa Rica-Panama Arc 51, 56 Cotton Valley clastics 88 Cotton Valley sequences 63 Couva evaporite 482, 488, 491 Couva Formation 487, 488 Couva Marine- 1 well 510 Covariance matrix 38, 40 Cretaceous Caribbean igneous province 635 Cretaceous Caribbean oceanic plateau 6, 29 Cretaceous Great Arc subduction complexes 245 Cretaceous igneous province 666 Cretaceous passive margin phase 19 Cretaceous volcanic plateau 633 Cruise RC 1904 405 Crustal elements 564 Crustal provinces 3 Crustal structure 564 Crystalline basement 64, 74, 89 Cserna Megashear 126 Cuba 28, 64, 93, 167, 169, 189, 241, 242, 247, 498, 663 Cuba Fracture Zone 188 Cuban arc 168 Cuban orogen 74, 167 Cuban orogenic belt 168 Cuban tectonostratigraphic terranes 94 Cuban-Bahamas collision 189 Cuche Formation 501,509 Cuche shale 488 Cul-de-Sac basin 291,292 Cunapo Conglomerate 482, 487, 488, 492, 493, 511, 531 Cunapo Formation 487, 509, 510, 513, 516, 519, 520, 521,528 Curwao 599, 627 Curwao Ridge 565, 582, 617, 653 Current-controlled drift deposit 591 Cusps 41 Dajabon 237 Dajabon area 237 Dajabon cherts 238, 245 Dajabon rocks 242 Danlf 155 De Purus arch 618 Dead Sea of Israel 337 Decompression melting 584 Deep eastern Gulf 82 Deep-water (flysch) phase 435 Deep-water canyon cutting event 463 Deformation front 14 Degrees of freedom 39 Delaware basin, Texas 337 Demerara Plateau 441,447 Depth to basement calculations 6, 28 Devil's River 238, 243 Difference vectors 35 Dipping reflectors 564 Dispersion 39
690 Dix formula 405 DNAG timescale 51 Doldrums 37 Doldrums fracture zone 41 Dolomita Principale peritidal complex 333 Dolomite Mountains of Italy 333 Dominican Cartographic Institute 296 Dominican Republic 128, 237, 238, 242, 245, 247, 284, 287, 288, 292, 300, 334 Dominican sub-basin 628, 637 Domoil High 477, 482, 487, 492, 526, 528, 533 Domoil-Gupe High 488 Doubloon Saxon 1 well 168, 171, 174 Drift stage 93, 117 Drifting stage 65 DSDP 587 DSDP 151 ridge 629, 651,653, 659 DSDP 151 ridge-Taino ridge area 659 DSDP cores 591 DSDP dated drill samples 27 DSDP drilling 561 DSDP drilling results 63 DSDP Hole 538A 83 DSDP holes 631 DSDP Leg 15 561 DSDP Leg 15 drilling 602 DSDP Leg 15 sites 599 DSDP Leg 15, Site 1001 561 DSDP Leg 77 65 DSDP Site A 241 DSDP Site 4 241 DSDP Site 31 641,649 DSDP Site 95 241 DSDP Site 144 447 DSDP Site 146 241,600, 602, 612, 622, 623 DSDP Site 150 575, 600, 602, 612, 622, 623 DSDP Site 151 633 DSDP Site 152 575, 599, 615 DSDP Site 535 68, 80, 88, 89 DSDP Site 536 68, 89 DSDP Site 537 68, 69, 77, 88 DSDP Site 538 68, 77 DSDP Site 540 68, 69, 80, 89 DSDP sites 82, 599, 615 DSDP well control 82 DSDP wells 89 DSRV Alvin 357 Duarte Complex 128, 242 Duarte highway 240 Durango 130 Durham sand 525 Early Cretaceous post-rift sequence 64, 89 Early shallow carbonate bank phase 546 Earthquakes 3 East End 344 East End Range 353 East Pacific Rise 12 East Panama deformed belt 15 East-west shortening 56 East-central Oregon 128 Eastern Caribbean 413, 585, 591, 611
SUBJECT INDEX Eastern Caribbean structural barrier 622 Eastern Cordillera 591,620, 621,635 Eastern Cuba 93 Eastern Guatemala 112 Eastern Maturfn sub-basin 424 Eastern Mexico 199 Eastern Pacific 14, 503, 584, 591, 613, 623 Eastern Pacific Ocean floor 16 Eastern passive margin of Yucat~in 117 Eastern Venezuela 22, 33, 57, 421,495 Eastern Venezuela foredeep 471 Eastern Venezuela Geosyncline 421 Eastern Venezuela offshore 471 Eastern Venezuelan Basin 15, 419, 421, 423,425,441,491,501,544, 545 Eastern Venezuelan fold belt 491 Eastern Venezuelan fold-and-thrust belt 477 Eastern Venezuelan foreland basin 501 Eastern Yucat~in 123 Eastern Yucat~in block 69 Easternmost Honduras 151 Edge-driven plate tectonic interactions 53 Eirik drift 623 E1 Abra Formation 161' 162 E1 Abra Limestone 146 E1 Aguacate 245 E1 Americano Member 88, 101, 101, 143 E1Cacheal tufts 249, 253, 281,283 E1 Cantil Formation 447 E1 Carbon 220, 226, 227, 231,234 E1 Coche-North Coast 499, 501 E1 Coche-North Coast fault zone 501 E1 Furrial area 450 E1 Furrial fields 423 E1 Furial-Carito oil fields 457, 473 E1 Furrial-Carito trend 450 E1 Granado 321 E1 Granado section 320, 321,325, 326 E1 Limon 267, 270 E1 Limon Member 267, 269-272, 282 E1 Mamey 253, 257, 258, 282 E1 Mamey Formation 253 E1 Mamey Group 247, 255, 257, 273, 274, 275, 277, 282, 283, 284 E1 Mamey measured section 260 E1 Mamey section 259, 260, 267, 282 E1 Mochito 154, 159 E1 Pastor Member 138 E1 Pilar 51 E1Pilar fault 15, 477, 478, 492, 495, 496, 506, 517, 554 E1 Pilaf fault scarp 531 E1 Pilar fault zone 15, 495, 496, 498, 499, 505, 506, 507, 511,515, 518, 528, 539, 541,542, 554, 555 E1 Pilar sub-basin 507, 513, 518, 521, 523, 524, 553 E1 Plan Formation 155 E1 Progreso 220 E1 Sfibalo Formation 99, 100, 109, 111, 112, 113 E1 S~ibalo/Artemisa boundary 100 E1 Salvador 9 E1 Soldado 499
E1 Soldado fault 552 E1 Soldado fault zone 505 E1 Tambor Formation 211 E1 Tambor Group 200, 211 E1 Verde Member 138 Elk Point basin 289 Ellipsoids 40, 44 England 162 Enriquillo 249 Enriquillo basin 287, 288, 289, 291,292, 295, 304, 309, 336, 337, 340, 635, 637 Enriquillo depression 635, 637 Enriquillo fault 635, 637, 663, 666 Enriquillo Valley 287, 296, 313, 315, 334, 337, 340 Enriquillo-Plantain Garden fault zone 13, 291,300, 301,302, 304 Equatorial Atlantic 35 Erin Basin 489, 491 Erin-Siparia syncline 501 ERS- 1 altimetry data 36 ERS- 1 data 41 Escambray terrane 94, 109, 110, 112 Escape of the Caribbean 53 Esperanza belt 103 Espino Graben 441 Esqufas 155 Esqufas Formation 151, 155, 163 Estate Work and Rest 349, 355 Etang Saumatre 292, 292 Eugenia Formation 126 Evans Highway 351,353, 356 Evaporites 83, 287 Ewing 9501 cruise 629, 637 EW-9501 profiles 565 Exogeosyncline 421 Extrusive volcanic mounds 564 Exuma Sound 171 Fairplain 351 Fairplain fault 353, 356 Falcon-Aruba 477 Farallon Plate 592, 108, 168 Fault A 300, 301,302 Fault B 301,302, 304 Fault C 301,301 Fault zone D 301 Fault-angle depressions 7 Fault-wedge basins 7 Felicidades 94 Ferro-di-lancia 336 Fiji Plateau 415 Finite motion poles 35 Finite plate motion poles 38 Finite rotation poles 44 Finite rotations 36, 40 First-motion studies 3 First-order cycle 472 Five Corners 349, 350 Flexural basins 435 Florida 441 Florida block 63, 83 Florida Escarpment 64, 69, 82, 89 Florida Plain 74, 77 Florida platform 89 Florida scarp 437 Florida-Bahamas area 187
SUBJECT INDEX Florida-Bahamas block 65 Flow lines 37, 41, 56 Flysch-type sediments 510 Foothills 503 Foothills belt 505 Foothills fold-thrust belt 499, 501 Foraminifers 99 Forearc basement 27 Foredeep 419, 421,425, 435 Foredeep clastic wedge 435 Foredeep phase 419, 457, 465 Foredeep sequence 435 Foredeep sequence regime 438 Foredeep subsidence 553 Foreland 435 Foreland basin 12, 419, 425 Forestepping regressive phase 471 Four-North fracture zone 37, 41 Fourth-order sequence 464 Fourth-order unconformity 435 Fracture zones 8, 36 Franciscan Complex 128 Francisco Formation 87, 100, 101, 111, 113, 141 Fredensburg Quarry 349 Fredericksburg Group 159 Fredericksted 353 Free-air gravity 8 Free face 28, 551,555 Freites Formation 424, 461 French Seacarib cruise 628 Frontal thrust 618 Full-graben setting 81 Full-grabens 8 Fusulinacea 99 Galapagos hotspot 9, 17, 584, 587, 599, 600, 613 Galapagos hotspot scenario 592 Galapagos mantle plume 592 Galapagos rift 12, 22 Galapagos seafloor 12 Galeota Point thrust fault 501,503 Galice Formation 128 Garcfa Member 447 Gateway 591 Gazzi-Dickinson point-count method 275 Geanticlinal welt 421 Geosat altimetry data 35, 36 Geosat data 41 Geosat gravity data 18 Geostrophic flows 622 Glacio-eustatic sea-level fluctuations 464 Global sea level lowstand 284 GLORIA data 48 GLORIA imagery 357 Gonave microplate 57, 637, 663 Gonave-Venezuelan microplate 666 Gondwana crust 145 Goodrich basin 533, 554, 555 Goodrich pull-apart basin 529, 542 Goodrich sub-basin 495, 496, 507, 513, 515, 516, 520, 521,522, 524, 525, 526, 528, 531,539, 542, 553, 555 Gopa High 477, 482 Gopa-Posa Highs 488 Gorgona 599
691 GPS-based geodetic studies 4, 28, 29 Graben 63, 81, 89 Gran Mangle 253 Gran Mangle series 281 Graphitic schist 5 Gravitational collapse 437 Gravity maps 8 Great Abaco fracture zone 167 Great Arc of the Caribbean 6, 7, 12, 13, 17, 18, 27, 28, 200, 241,245, 498 Great Bahama Bank 167, 168, 174, 187, 188, 191 Great circle segments 39 Great Valley Supergroup 126 Greater Antilles 50, 55, 128, 168, 247, 343, 391,397, 622 Greater Antilles Arc 93, 108, 111, 115, 118, 168, 248 Greater Antilles drift 623 Greenland 109, 144, 147 Grenada basin 6, 7, 13, 28, 389, 392, 397, 398, 403,405, 407, 413,414, 415, 499 Grenville crustal age province 19 Growth fault zone 463 Growth-faults 437 GSA Decade of North American Geology 7 Guacamaya Formation 139 Guachichil 144 Guachichil terrane 145 Guajaib6n Formation 96 Guajaib6n-Sierra Azul belt 95, 96, 114 Guajaib6n-Sierra Azul unit 94 Guanahacabibes Peninsula 95, 109 Guanaja 197, 201,222 Guanaja Island 203, 205, 216 Guananico measured section 260 Guananico section 259, 260, 263, 267, 282 Guane 95, 96 Guaniguanico belts 113 Guaniguanico foreland basin 108 Guaniguanico nappe 97 Guaniguanico tectonostratigraphic belts 96 Guaniguanico tectonostratigraphic unit 93 Guaniguanico terrane 93, 94, 96, 97, 103, 107, 109, 111, 112, 113, 116, 117, 118, 144, 147 Guanoco area 477, 492 Guare Member 155, 162 Gu~irico 15 Gu~irico sub-basin 419, 420, 424, 439, 501,503, 505, 544 Guasasa Formation 101,102, 143, 144, 147 Guatapajaro Anticline 489, 536 Guatemala 9, 80, 110, 145, 153, 199, 200, 222, 477 Guatemala-Belize coastline 234 Guerrero 126 Guerrero block 152, 163 Guiamas River 225 Gulf Coast 154, 157 Gulf Coast region 159 Gulf High 482, 487, 488, 492, 507, 513,
515, 526, 528, 533 Gulf of Guinea 566 Gulf of Honduras 199 Gulf of Maracaibo 617 Gulf of Mexico 6, 18, 124, 144, 145, 153, 159, 168, 188, 289, 291,437, 623 Gulf of Mexico basin 63 Gulf of Paria 14, 419, 477, 478, 487, 488, 489, 491,495 Gulf of Paria basin 495, 498, 507, 524, 539, 553, 554 Gulf of Paria fault zone 501 Gulf of Paria-Northern basin 495, 503 Gulf Tectonics 65 GULFREX-Gulf 65 Guyana 19, 498 Guyana Basin 447 Guyana Craton 441 Guyana margin 551 Guyana offshore 447, 448 Guyana offshore basin 447 Guyana passive margin 14, 440 Guyana Shield 15, 19, 439, 446, 455, 466, 496, 505, 546, 551,617, 621 Haiti 291,292, 599, 629 Haiti Basin 601 Haiti plateau 629, 635 Haiti sub-basin 628, 629, 635, 651,653 Haitian border 309 Half-grabens 8, 13, 63, 437, 443, 546 Harvard focal mechanism catalogue 3 Hato Viejo Formation 441 Havre basin 391,394, 397 Haynesville Formation 87 Haynesville sequences 63 Heat-flow measurements 28 Hess Escarpment 16, 28, 591,600, 613, 615, 623, 624, 629, 653, 663 Hess Oil refinery 353 Hidalgo 123, 147 Higher Boreal paleolatitudes 124 Hildago 137 Himalayas 111 Hispaniola 6, 7, 12, 13, 16, 24, 57, 241, 247, 248, 275, 283, 284, 287, 289, 291,336, 413,498, 624, 627, 635, 637, 659, 663 Hispaniola arc 247, 279, 284 Hispaniola margin 615 Hispaniola restraining bend 13 Hispaniola terranes 241 Hole 146 630 Hole 153 630 Hole 31 630 Honduran shelf break 222 Honduras 5, 9, 110, 154, 163, 197, 199, 200, 219 Honduras and Guatemala 216 Honduras Group 155 Honduras-Facolina section 320 Horizon A" 630 Horizon B" 630 Horst 81, 89 Hotspot trace 9 Huayacocotla Anticlinorium 138, 146 Huayacocotla Formation 138, 139, 140
692 Huayacocotla remnant 129, 139, 141, 143, 144, 145, 147 Huayacocotla segment 123, 145, 147 Huayacocotla terrane 145 Hyde Formation 129 Hydrocarbon deposits 28 Hydrocarbon occurrence 288 Hydrocarbon source rocks 288 Hydrocarbons 498 Iberian margin 565, 566, 581 Ilama Formation 157 ile-fi-Vache structure 635 Imbert 243 Imbert Formation 242, 243, 281 Inclined subducted slabs 4 Indian ocean hotspot tracks 52 Infierno Member 143 Inner arc setting 281 Inner forearc (Los Hidalgos Formation) 283 Institute Fran~ais du P6trole 392, 629 Institute for Geophysics 65 Integral constraints 39, 41 Intermediate unconformity (SB-2) 438 Internal plate deformation 6 Inversion method 35, 38 Inverted basins 8 Inverted Caribbean sedimentary basins 8 Iquitos arch 618 Isla Cabritos 301 Isla de la Juventud 93 Island-arc basins 7 Island-arc terranes 249 Israel 162 Isthmus of Panama 15 Isthmus of Tehuantepec 145 Italy 335 Izee terrane 128, 129, 141 Jacaguas Formation 143 Jagua Clara M61ange 243 Jagua Formation 87, 100, 101, 111, 113, 143 Jagua Vieja Member 100, 143 Jaitique 155 Jaitique Formation 151, 155, 156, 162, 163 Jamaica 13, 16, 52, 56, 153, 237, 243, 245, 247, 477, 635, 663, 666 Japan 126 Jealousy 345 Jealousy Formation 343, 344, 345, 347, 353, 354, 358, 363 Jealousy/Kingshill boundary 345 Jimanf 307 Jimanf Formation 304, 307, 336 Jordan 65 Jordan Knoll 89 Josephine Ophiolite 128, 129 Judith's Fancy 354 Jurassic half-grabens 441 Jurassic rifting event 473 Jurassic-Cretaceous boundary 69 Jutiapa-Trujillo road 220 Kallinago basin 14 Kane Fracture Zone 33, 35, 37
SUBJECT I N D E X Kerguelen 583 Kerguelen Plateau 577, 583, 586 Kingshill basin 343, 346, 353, 354, 363 Kingshill graben 347, 357 Kingshill Limestone 343, 344, 345, 347, 358, 363 Kingshill strata 354 Klamath Mountains 128, 144, 147 Knowles ramp 88 Krause Lagoon 356 Krausirpi beds 151,157, 159, 161,162, 163 Krausirpi quadrangle 159, 162 Kroonvlag data 36 Kurile basin 397 La Caja Formation 128, 135, 137, 139, 146 La Caja Unit E 138 La Casita 135 La Casita Formation 135 La Ceiba 211,220, 222, 223, 224 La Ceiba fault 201,215, 223 La Cumbre Ridge 273 La D6sirade 128, 245, 391,398 La Esperanza 211 La Esperanza belt 93, 95, 105, 112, 114, 117 La Gloria Formation 133 La Gtiira Member 107 La Isla Formation 281 La Legua Member 107 La Palma 96, 112 La Pefia 136 La Pica Formation 424, 462 La Pocilguita del Limon 267 La Pocilguita Member 267, 270, 271, 272 La Quinta Formation 441 La Reine Member 343, 349, 350, 351, 353, 354, 356 La Toca block 249, 251,253, 255, 272, 273, 283 La Toca Formation 242, 251,253, 255, 272-275, 277, 279, 282, 283, 284 La Zarza Member 88, 102, 103, 143 Labrador Sea 565, 581 Lac Assal, Djoubouti, Persian Gulf 337 Lago Enriquillo 292 Lago Enriquillo area 297 Lago Enriquillo block 302 Lagoven S.A. 420 Laguna Rinc6n 299 Lake Macleod, Western Australia 289, 333, 337 Lake Maracaibo 22, 552 Lake Maracaibo region 618 Lake Yojoa 151, 152, 154, 155, 159, 162, 163 Lakes 9 Lamont Doherty Earth Sciences Observatory 629 Lamont-Doherty Geological Observatory 392 Landward-dipping 'D' reflectors 583 Large igenous province 575 Large-scale plate-tectonic rotation 24 Las Lavas Formation 247, 251,255,
257, 260, 264, 267, 269, 270, 271, 272, 274-277, 279, 282, 284 Las Lavas measured section 269 Las Lavas section 267, 271 Las Mangas tonalite 224 Las Piedras Formation 424 Las Salinas 296 Las Salinas fault zone 311 Las Salinas Formation 287, 296, 302, 304, 307, 311,315, 316, 328, 336 Late Berriasian unconformity 80 Late Cretaceous-Recent arc-passive margin collisional phase 19 Late Jurassic rift phase 18 Late Jurassic rift sequence 64, 89 Late Jurassic syn-rift 74, 89 Late Jurassic-Early Cretaceous seaway 89 Later deeper-water passive margin phase 548 Latitude, longitude, and rotation angle space 40 Latitude-longitude sphere 40 Lau basin 391,394, 415 Lebanon 162 Leeward Antilles 6, 391,503 Left-lateral Santa Marta-Bucaramanga fault 15 Leg 77 volume 74 Lesser Antilles 3, 391,392, 397, 398, 407, 499, 622, 623 Lesser Antilles Arc 3, 4, 8, 13, 14, 48, 55, 343, 360, 397, 389, 403,405, 409, 423,539, 555 Lesser Antilles Seismic Project 409 Lesser Antilles subduction zone 14, 48, 498, 499, 500, 663 Limestone Caribbees 14 Lithospheric trace 14 Llanada and Point Sal remnants of the CRO 130 Llanos syncline measured section 266 Llanos syncline section 260, 266, 267, 282 Local structural rotation 24 Loma de Sal y Yeso 287, 289, 296, 302, 307, 309, 312, 315, 340 Loma del Muerto tectonic unit 112 Los Bajos 499 Los Bajos fault 478, 488, 489, 492, 496, 499, 503, 541,552, 555 Los Bajos fault zone 501,505, 506, 555 Los Cayos Member 107 Los Guiros syncline 317, 320 Los Hidalgos Formation 247, 249, 253, 257, 258, 260, 266, 270, 281,284 Los Hidalgos Pass 260 Los Jabillos clastics 450 Los Jabillos Formation 473 Los Organos 94 Louann salt 87, 289 Louann-Campeche salt 83 Lower Carapita Formation 457, 466 Lower Cretaceous carbonate margins 64 Lower Cretaceous post-rift sequence 74 Lower Forest Formation 538, 555 Lower La Pica Formation 487, 488 Lower Merecure Formation 457
SUBJECT INDEX Lower Nicaragua block 666 Lower Oficina Formation 466 Lower Talparo Formation 526 Lower Tethyan paleolatitudes 124 Lower Valle de Angeles Group 151 Lowstand 333 Lowstand fan 284 Lowstand prograding wedge 461 Lucas Formation 105 Luperon facies 253 Luperon Formation 255, 281 Lutgarda Formation 116 Macropaleontologic data 27 Magadi-type chert 313 Magante 240 Magdalena deep-sea fan 631 Magdalena River 621 Magnetic anomaly 35 Magnetic anomaly and fracture zone date sets 35 Magnetic data 55 Magua Formation 242, 281 Maimon Bay 239, 242 Maimon-Amina schists complex 242 Mamey group 242 Manacas Formation 108, 116 Managua Lake 9 Manihiki 5 Mannings Hill 351,352 Mannings Bay Member 343, 349, 351, 352, 356 Mantle plume 583 Mantle plume source 584 Mantua 96, 112 Manzanilla Formation 482, 487, 488, 492, 510, 511,513, 519, 521 Maparito faults 437 Maracaibo 15 Maracaibo basin 15, 419 Maracaibo block 4, 15, 24, 51, 56 Maracaibo-Peruvian foreland basin 591, 617, 618, 624 Maracaibo-Peruvian Trough 617 Marajo basin 618, 621,622, 624 Marathon 37, 41 Margarita 503 Margarita Island 503 Mariana basin 397 Marie Aimee ridge 660 Marine Geophysical Data Center 628 Marine heat-flow measurements 6 Marine seaway 89 Marine seismic profiles 6 Martfn Mesa 1 well 96 Martfn Mesa tectonic window 96 Martin Vaz 38 Matagalpa Formation 157 Matanzas Province 93 Maturfn 15, 440 Maturfn Basin 437, 441,450, 547 Maturfn foredeep 437 Maturfn foreland basin 492, 546 Maturfn sub-basin 420, 424, 448, 501, 503, 544, 545, 546, 551 Maude Formation 126 Maximum flooding 459 Maximum flooding surface 419, 443,
693 459, 461 Maya block 152, 197, 199, 200, 211, 213,215 Maya terrane 144, 145 Mazapil 123, 130, 135, 138, 147, 146 Mazapil remnant 137, 138, 139, 147 Mazapil succession 123, 147 Median back-arc basin 8, 9 Median-Nicaraguan back-arc basin 7, 9 Mediterranean 289, 336, 337 Mella block 301 Mella wells 296, 297, 301 Melvin Evans Highway 349 Mercurius 37 Mercurius fracture zones 41 Merecure Formation 457 Merecure Group 450 Merecure type section 450 M6rida Andes 622 Mesa Formation 424 Mesozoic Caribbean Arc 6 Mesozoic passive margin 472 Messinian 334, 463 Messinian event 356 Messinian sea-level lowering 419 Mestanza tectonic unit 95 Metamorphic protolith rocks 4 Mexico 124, 126, 146, 153, 159, 161, 211 Mexico-Marathon-OuachitaAppalachian structural belt 124 Microcontinent 18 Micropaleontologic data 27 Mid-Atlantic ridge 627 Middle America arc 3, 4, 6 Middle America subduction zone 3,498 Middle America Trench 234, 663 Middle America Trench subduction zone 152 Middle America volcanic arc and trench 9 Middle Cretaceous Sequence Boundary 69, 74, 82 Middle East 159 Middle Fork of Smith River 128 Mid-Las Salinas horizon 304 Mid-Las Salinas reflector 299, 301,304 Minas de Matahambre 112 Minas de Matahambre-La Palma area 96 Misfits 39 Mississippi Canyon 65 Mississippi Fan turbidite plain (Florida Plain) 64 Mobil 296, 297, 304 Mochito area 154 Mochito Mine 154 Mochito shale 151, 154 Mogote zone 95 Moho 55 Moho reflection 565 Moho topography 566, 600 Moho uplift 48 Moho-penetrating reflectors 582 Mojave Sonora Megashear 145 Molasse-type sedimentary rocks 510 Mona Canyon 360 Monagas foothills 423, 424, 455, 466
Montafia de Santa B~irbara 159 Montafias de Col6n 151,153, 155, 157 Montafias de Col6n fold belt 152 Montafias de Col6n region 157 Montafias de Col6n-Rfo Wampti 152 Monte Cristi 267 Montserrat glauconitic sandstone 519 Moreno depocenter 93, 115, 116, 118 Moreno Formation 93, 106, 115, 117 Morningstar sections 349 Morochito piggyback basin 503 Mor6n-E1 Pilar fault zone 499, 500, 501 Mor6n fault 478 Morro area 488 Morro northward-vergent imbricates 489 Motagua 477 Motagua fault 197, 200 Motagua fault zone 197, 199, 200, 211, 213,216 Motagua Valley 211,222 Motagua/Chixoy-Polochic fault zones 152 Motagua-Polochfc system 9 Motion vectors 35 Mount Harris 535 Mount Harris push-up block 509 Mt. Eagle Group 344, 357 Mt. Eagle Series 349 Muertos prism 635, 637 Muertos trench 13, 16, 627, 653 Muertos Trough 33, 50, 55, 56, 360, 363, 591,601,615, 624, 627, 628, 635, 637, 663 Multichannel seismic reflection data 591 Multifold seismic data 63 Nafe-Drake curve 405 Naranjo tectonic unit 113 Naricual clastics 450 Naricual Formation 473 Nariva fold-and-thrust belt of southern Trinidad 489 Nariva turbidites 491 Navarette 258, 267 Navarette measured section 271 Navarette section 267, 271 Navassa trough 635 Nazca plate 3, 16, 22, 663, 666 NE Yucat~in coast 93 Neiba prism 637 Neiba-Plaisance Formation 297 Neogene foredeep 472 Netherlands-Antilles 391 Nevadian island arc 144, 147 New Zealand 7, 111, 126 Nicaragua 9, 152 Nicaragua Lake 9 Nicaraguan Rise 16, 152, 197, 199, 200, 663 Nicaraguan Rise-Greater Antilles Arc 93, 115, 117 Nicely Formation 129 Nicoya Peninsula 663 Normal-polarity intervals 36 Norphlet Formation 87 North America 3, 18, 108, 126, 183, 423, 591
694 North America plate 3, 35, 65, 189, 247, 248, 289, 291,304, 360, 362, 363, 403, 498, 615, 627, 666 North America-Caribbean 43 North America-Caribbean foreland basin 12 North America-Caribbean oblique-slip plate boundary zone 287 North America-Caribbean plate 663 North America-Caribbean plate boundary 247 North America-Caribbean plate boundary zone 197, 279, 284, 343, 345 North America-Caribbean plate motions 43 North America-Caribbean strike-slip zone 9 North America-South America breakup 7 North America-South America motion 18, 33 North America-Venezuelan microplate motion 666 North Andes block 663, 666 North Atlantic 19 North Atlantic Deep Water 622 North Atlantic Province 144 North Atlantic rifted margins 584 North Coast basin 501,506 North Coast fault zone 478 North Coast-E1 Coche fault zone 503 North Fiji basin 391 North Panama deformed belt 15, 33, 50, 51, 55, 57 North Soldado basin 542 North Soldado sub-basin 541,542 North Venezuela margin 51 North-South America plate motions 50 North-South America relative motions 46 Northeast Providence Channel 171 Northeastern coast of Cuba 12 Northeastern Great Bahama Bank 171 Northeastern margin of South America 21 Northeastern Oaxaca 145 Northern Andes 621 Northern Apennines 323 Northern Atlantic Ocean 471 Northern basin of Trinidad 495, 498, 500, 506, 507, 513, 520, 523, 526, 539, 540, 553 Northern basin sidewall 492 Northern Boreal Province 126 Northern Canada Bonita section 259, 260, 263, 266, 282 Northern Central America 12, 16, 18 Northern coast of Cuba 12 Northern Cuba 168 Northern Grenada basin 389 Northern Guarapiche province 477 Northern Guatemala 145 Northern Haiti 281,283 Northern Hispaniola 279, 666 Northern Honduras 13 Northern Italy 336 Northern margin of South America 498
SUBJECT INDEX Northern Mexico 153, 155, 162 Northern province 145 Northern Range of Trinidad 477, 488, 489, 492, 495, 498, 501,503, 505, 506, 509, 510, 518, 523, 526, 531, 533, 541,553, 554 Northern Rosario 93 Northern Rosario belt 93, 95, 99, 100, 102, 103, 105, 106, 108, 112, 113, 114, 116, 117 Northern South America 13, 16, 63, 112, 499 Northern Taino ridge 643 Northern Tethyan Province 109, 126, 137, 138, 139, 144 Northern Transtensional Basin 480, 482, 487, 488, 493 Northern Transtensional Basin fill 480 Northern Trinidad 473, 477 Northern Venezuela 5, 397, 425 Northern Venezuelan transpressional folded belts 455 Northern Yucat~in Straits 64 Northside Range 344, 349, 353, 354, 357 Northwest Bahamas 190 Northwestern margin of South America 21 Northwestern Mexico 157 Northwestern South America 18, 22, 24 Northwestern Venezuela 115 Nova Scotia 583 Nuvel-lA plate 3 NW Atlantic crust 565 Oaxaca 126 Oaxaca area 19 Oca 50, 618 Ocean Drilling Program 6 Oceanic large igneous provinces m Java 580 J Kerguelen 580 Manihiki 580 Ontong 580 Oceanic plateau 5 Oceanic plateau province 4 Oceanic plateau terrane 249 Ocoa Bay 635 Ocoa Group 284 ODP 587, 628 ODP cores 5 ODP dated drill samples 27 ODP drilling 561 ODP Leg 165 561,599 ODP Site 1001 600, 602, 613, 615, 621 ODP Site 642 573 ODP sites 599, 615 Offshore eastern Venezuela 472 Offshore French Guyana 441 Offshore northeastern Gulf 82 Offshore Orinoco platform 443 Oficina Formation 424, 457 Okinawa basin 397 Olanchito 220, 222 Olancho district 159 Old Bahama Channel 174, 175 Oligocene passive margin section 435 Oligocene-Miocene boundary 57 -
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Olistostromes 284 Omoa 220 Ontong Java 5 Ontong Java Plateau 584 Ophiolites 27, 52 Oriente fault 633, 635, 637, 663 Oriente transform fault zone 199 Orinoco 14, 591, 618, 624 Orinoco delta 14, 425, 435,437, 441, 464, 465, 503, 528, 542, 551,554, 555 Orinoco drainage system 621,622 Orinoco Geosyncline 421 Orinoco offshore 425 Orinoco Platform 419, 420, 425, 435, 437, 439, 447 Orinoco River sediments 542 Orinoco shelf 448 Orinoco tar belt 419 Orogenic grains 19 Orthogeosynclinal phase 421 Outer Apalachicola basin 87 Outer arc-trench assemblage 281 Outer forearc-trench assemblage (Imbert Formation) 283 Outer forearc-trench setting 281 Overlaps 44 Pacific Ocean 586, 592, 587, 617, 618, 627, 631 Pacific peninsula of Costa Rica 9 Pacific peninsula of Panama 9 Pacific side of Mexico 163 Pacific source 147 Padre Miguel Group 157 Padre Miguel volcanics 234 Paleogene back-arc basins 6 Paleomagnetic reference frame 52 Paleozoic(?) Pre-rift rocks 64, 74, 89 Palma Picada 258 Palma Picada area 253 Palma Picada Formation 258 Palma Picada intrusions 258, 260 Palma Picada intrusive rocks 249, 284 Palma Picada porphyritic rocks 281 Palma Picada rocks 257 Palmarito Formation 112 Palo Alto-1 well 300 Palo Blanco Formation 140, 141 Pan de Azticar 100 Pan de Azficar Member 100 Pan-African 69 Pan-African crustal age province 19 Panama 5, 241 Panama Arc 4, 16, 22 Panama Arc collision 51 Panama Isthmus 15, 613, 623 Panama prism 663 Panama seaway 24 Panamanian block 663 Panamanian subduction zone 663 Pangea 19, 124, 144, 145, 168, 423,441 Panuco No. 82 146 Parece-Vela basin 391,397 Paria Peninsula 477, 487, 488, 489, 492, 499 Parral 145 Parral terrane 130, 145
695
SUBJECT INDEX Parras 133 Partial uncertainty rotations 38 Passive margin 7, 419, 437, 438 Passive margin of Guyana 420 Passive margin of North America 12 Passive margin phase 419 Passive margin sequence 419, 435 Passive margin successions 94 Patos Island 531 Patuca 153 Peak transgression 461 Pearl Islands basin 15 Pecos fault 651 Pecos fault zone 651,653, 659, 665 Pedernales 491 Pedernales area 489, 492 Pedernales oil field 492 Pedernales region 477 Pedernales shale ridge 489 Pedro escarpment 663 Pedro fault 663 Pedro Garcfa 272, 273 Pedro Garcfa anticline 251 Pedro Garcfa Formation 249, 251,272, 273 Peg-leg and water column multiples 564 Penal oil field 538 Pefialver Formation 107, 116 Pefias Formation 106, 116 Peralta and Rfo Ocoa sediment groups 50 Peralta belt 295, 637 Peralta flysch belt 635 Peregrina Canyon 123, 139, 146 Peripheral bulge 8, 435 Perisutural basins 435 Permian basin 291 Permian Castile Formation 337 Permian Delaware basin 289 Permian Salado Formation 333 Peru 617 Petroleos de Venezuela S.A. 420 Petroleum Company of Trinidad and Tobago 507, 521,524, 526 Petroleum reserves 419 Petrotrin 507 Phase shift angles 36 Philippine plate 414 Philippines 130 Phyllitic schist 5 Pica Pica Member 108 Pico Bonito 223, 224 Piedras Negras 224 Pierre Payen anticline 637 Piggyback basins 8, 14 Pimienta Formation 141 Pimienta Member 100, 101,143 Pinalilla Formation 106, 115 Pinar 1 well 97 Pinar del Rfo 65 Pinar del Rfo geology 93 Pinar del Rio Knoll 89 Pinar del Rio Province 123, 144 Pinar fault 95 Pino Solo tectonic units 95 Pinos terrane 94, 109, 112 Pirital thrust fault 503 Pitman Fracture Zone 51
Placetas belt 93, 111, 114-117 Plantain Garden-Enriquillo fault 637 Plataforma Deltana 420 Plate circuit 35 Plate circuit path 35 Plate hierarchy 52 Plate motions models 35 Plate reconstructions 3, 19 Platform sequence 435 Plume head 613 Point Sal 128 Pointe-a-Pierre Formation 510 Polier Formation 103, 105 Pons 97, 106 Pons Formation 105, 115 Porosity 69 Port-of-Spain 533 Posa area 492 Posa field 488 Posa High 487, 492 Posa oil field 492 Positive inversions 477 Post-Kingshill limestones 351 Post-rift sequence 63 Post-rift stages 63 Post-rift unconformity 65 Precambrian crystalline basement 447 Pre-rift 63, 89 Pre-rift phase 18, 419 Pre-rift rocks 63 Pre-rift section 73, 79 Pre-rift unconformity 435 Presqu'~le du Sud 627, 635, 635, 637, 657, 660, 663, 666 Presqu'~le du Sud-Beata ridge 637, 660 Progradational forestepping 459 Prograding (transitional) phase 435 Promax TM 168 Proto-Antillean Arc 6 Proto-Caribbean 55, 63, 83 Proto-Caribbean basin 93, 108, 110, 111,113-117 Proto-Caribbean crust 109, 168, 591, 613,623 Proto-Caribbean Ocean 18 Proto-Caribbean oceanic crust 27 Proto-Caribbean passive margin 83 Proto-Caribbean Sea 108, 113, 116 Proto-Caribbean seaway 112 Proto-Caribbean spreading ridge 21 Providence Channel 167, 168, 188 Pseudo-oceanic appearance 566 Puebla 123, 147 Puerto Cortez 220 Puerto Grande sub-basin 496, 507, 516, 521,523, 524, 531,553, 554 Puerto Plata 239 Puerto Plata area 242 Puerto Plata basement complex 237, 238, 240, 241,242, 249, 273 Puerto Rico 6, 7, 12, 14, 22, 241,245, 247, 248, 275, 283, 354, 357, 358, 360, 362, 364, 413, 591,627 Puerto Rico microplate 360 Puerto Rico platelet 362 Puerto Rico platform 343, 360 Puerto Rico terrane 360 Puerto Rico trench 12, 199, 248, 663
Puerto Rico-Hispaniola microplate 13 Puerto Rico-Virgin Islands terrane 360 Pull-apart basin 7, 13, 15 Purial area 242 Push-ups 7 Quartzite 5 Quebrada Juana Leandra 224 Queen Charlotte Islands 126 Querecual Formation 447, 448 Quifiones 94 Quifiones Formation 241 Quifiones tectonic unit 106 Quintana Roo 145 Quita Coraza Formation 305, 317, 320, 327, 328, 334, 335, 337 R/V Charcot 628 R/V Conrad 629 R/V Ewing 561,564, 566, 593, 599 R/V Ewing cruise 9501 591 R/V Glomar Challenger 68 R/V Nadir 627, 628 Radiolarian assemblage 127 Radiolarian microfacies 93 Rail Cabin Formation 129 Ramp basin structure 338 Ramp or 'push-down' basins 7 Ranchete 258 Ranchete Member 258, 259, 260, 266, 281 Rattan Hill area 349 Rattan/Belvedere 349 Razorback Ridge 307, 311, 312, 315 Red cherts 27 Red ribbon chert 245 Reflector A" 622 Refraction 564 Regional flooding events 419 Regional opening model 89 Relative motion path 18 Relative plate motion vectors 35 Remnant arc 13 Reprocessing 67 Researcher Ridge 48 Restraining bend 7, 284 Restraining bend tectonics 284, 285 Ridge push 53 Rift 7, 8, 63 Rift basins 8 Rift regime 438 Rifted passive margin 551 Rifts of the Canal area 15 Right-steps 13 Rio Bajabonico fault zone 251 Rio Grande 38, 273 Rio Grande fault 253 Rfo Grande fault zone 249, 257, 258, 269, 272, 273, 274, 283, 284 Rfo Jacagua measured section 271 Rio Jacagua section 267, 269, 271 Rfo Lenin graben 201, 215 Rfo Patuca 152, 153, 155, 157, 159, 162 Rfo Perez measured section 264 Rio Perez section 264, 266, 267, 282 Rfo Perez syncline section 260, 282 Rio San Juan 238, 243 Rfo San Juan area 272
696 Rio San Juan complex 237, 242, 243, 249, 272, 273 Rfo San Juan mudstone 245 Rfo San Juan-Puerto Plata-Pedro Garcfa disrupted terrane 273 Rfo Sutawala 155, 159, 162 Rfo Viejo 201 Rfo Viejo fault 211,223 Rfo Wampfi 152, 153 Rfo Wampti area 155 Rio Wampfi- Rfo Patuca area 157 Rfo Wampfi-Montafias de Col6n 153 Rfo Yaque del Sur 287, 292, 300, 317, 327, 340 Rfo Yaque section 333 Rio Yaroa 273, 274 RoaUin Island 197, 201,203, 205, 208, 210, 213, 214, 215, 222 Roble Member 103 Rogue Formation 128 Romeral Suture 617 Rosario belts 96, 98, 105, 108, 113, 115-118 Rosario North 94 Rosario South 94 Rotation parameters 42 Rotation uncertainties 35, 40 Rough B t~ basement 600 Rough B" crust 565 Rough-smooth B" basement boundary 600, 624 Rough-smooth B" transition 565 Rough-smooth boundary 12 Royal Trough 48 Saba 220, 354 Saba Bank 357, 358, 363, 364, 397, 398, 413 Sag basins 78 Saline giant 287, 289, 291 Salt River Valley 350 Samana Peninsula 249, 284 Sambfi basin 15 Sambfi fault 15 San Andres 16 San Andres Limestone 143, 144, 147 San Andres trough 16 San Antonio Formation 448 San Bias forearc basin 15 San Cayetano basin 93, 99, 111, 117 San Cayetano deltaic sediments 117 San Cayetano Formation 83, 87, 98, 99, 100, 109, 110, 111,112, 141, 143 San Cristobal basin 635 San Esteban 219, 220 San Fernando Bay 488 San Fernando High 488 San Francisco 499 San Francisco fault zone 501 San Francisco-Quiriquire faults 493 San Jose member 519 San Juan basin 295, 317, 635 San Juan Formation 448 San Juan graben 477, 492 San Juan-Azua basin 249, 291 San Juancito 155 San Luis Potosi 123, 137, 147 San Marcos Formation 237, 238, 242
SUBJECT I N D E X San Marcos unit 239 San Pedro basin 635, 637 San Pedro del Gallo 128, 130, 131,133, 136, 137, 138, 146 San Pedro del Gallo area 127 San Pedro del Gallo Fault 145 San Pedro del Gallo remnant 133, 137, 138, 139, 145, 147 San Pedro del Gallo terrane 109, 111, 123, 124, 129, 130, 133, 141, 144-147 San Pedro units 136 San Pedro Zacapa 159 San Salvador 168 San Vicente Member 87, 101,102, 113, 143, 144, 147 Sanarate limestone 213 Sandino forearc basin 9 Santa Ana field 450 Santa Barbara County 128 Santa B~irbara quadrangle 159 Santa Elena fault-Hess escarpment 152 Santa Marta Massif 621 Santa Marta-Bucaramanga fault 15 Santa Rita 220 Santa Rosa group 80 Santa Teresa 241 Santa Teresa Formation 105, 106, 114 Santa-Marta Massif 662 Santana 321,325 Santana-E1 Granado road 323 Santander Massif 621 Santaren Channel 188, 189 Santiago 239, 257, 267 Santiago Formation 140, 143 Santiago-Altamira highway 253 Santiago-Puerto Plata highway 258, 259, 267, 271,282 Santo Domingo 296, 312 Santonian time 5 Sarasota 87 Sarasota arch 88 Satellite altimetry 33, 34, 55 Scotland 313 Sea of Japan basin 391,392, 397 Seabeam map 627, 651 Seabeam survey 628 Seacarib 1 651 Seacarib 1 cruise 627, 629 Seacarib cruise 637 Seacarib profiles 637 Seafloor bathymetry 8 Sea-level fall 463 Seasat altimetry data 35, 36 Seaward-dipping reflectors 577, 583, 585, 600, 624 Sebastopol Complex 477 Secondary basal foredeep unconformity 463 Second-order sequence boundaries 419 Second-order T/R packages 472 Second-order transgressive-regressive cycle wedge 461 Second-order transgressive-regressive cycles 419 Sedimentary accretionary wedges 8 Sedimentary basins 3 Seed points 41
Seismic velocities 4 Septentrional block 635 Septentrional fault zone 248, 251,257, 267, 269, 272 Sepur clastics 115 Sepur foreland basin 12, 21 Sepur Formation 116 Sequence boundaries 443 Sequence boundary SB- 1 441 Sequence boundary SB-2 441,443 Serranfa del Interior 419, 420, 423, 425, 437, 439, 442, 447, 448, 450, 455, 465, 466, 472, 473, 477, 487, 494, 496, 501,503, 547, 548, 555 Serranfa del Interior belt 505, 546, 551, 552 Shale diapirs 14 Shallow subduction 9 Shallow water-deep basin 338 Shelf-break 463 Shelf-margin wedges 333 Shell 629 Shikoku basin 391,394, 397 Sicily 330, 333, 334, 336, 337 Sico River 220, 227, 231,232, 234 Sideswipe 565 Sierra Bahoruco 615 Sierra Bermeja 245 Sierra Cadnelaria 130, 146 Sierra Chiquita tectonic unit 116 Sierra de Bahoruco 295, 296, 300, 302, 309, 635, 637, 659 Sierra de Catorce 123, 130, 135, 138, 139, 145, 146 Sierra de Catorce remnant 138, 139, 147 Sierra de Escambray 93 Sierra de la Caja 123, 130, 135, 138, 146 Sierra de los Organos 83, 87, 88, 95, 97, 103, 105, 107, 109, 113, 123, 141, 143, 144, 147 Sierra de los Organos area 98 Sierra de los Organos belt 93, 95, 97-102, 105-108, 111-117 Sierra de los Organos remnants 111 Sierra de los Organos succession 112 Sierra de Martin Garcia 635 Sierra de Neiba 295, 305, 317, 635, 637 Sierra de Omoa 220 Sierra de Parras 135 Sierra de Perija 621 Sierra de Zuloaga 123, 138 Sierra del Abra 146 Sierra del Rosario 87, 88, 95, 100, 103, 109, 110, 123, 141,143, 144, 147 Sierra del Rosario belt 111 Sierra del Rosario meridional 83 Sierra del Rosario remnants 111 Sierra del Rosario septentrional 83 Sierra del Rosario succession 111 Sierra Jimulco 135, 146 Sierra la Gloria 133 Sierra Madre Oriental 123, 133, 145, 146, 147 Sierra Madre Oriental terrane 123, 130, 145 Sierra Martfn Garcfa 295, 304, 305, 317 Sierra Martfn Garcfa anticline 300 Sierra Nevada 141
SUBJECT INDEX Sierra Nombre de Dios 220 Sierra Ramirez 130, 146 Sierra Santa Rosa 135, 137, 138, 147 Sierra Sombreretillo 133 Sierra Sombretillo 130, 146 Sierra Zuloaga 130, 146 Siete Cabezas basalt, 237, 243 Siete Cabezas Formation 240 Siliciclastic wedge 495 Sinu belt 662 Sinu subduction zone 627, 663 Sinu trench 662, 663 Ski-jumps 566, 583 Slab pull 53 Slow-spreading ridges 565 Smackover carbonates 63 Smackover Formation 87 Small- and medium-offset fracture zones 37 Smith River subterrane 129 Smooth BI~basement 600 Smoothness 41 Snowshoe Formation 128, 141 Soldado High 488 Solim6es basin 618, 621,622, 624 Sombrerito Formation 295, 298, 301, 304 Sonic velocity 69 Sonobuoys 564 Soroa 96 South America 3, 13, 18, 35, 93, 108, 111, 112, 423, 546, 586, 591,593, 623 South America plate 3, 35, 391,403, 423, 425,435,450, 463, 471,477, 489, 495-498, 615, 617, 627, 663, 666 South America plate boundary 624 South America-Caribbean 43 South America-Caribbean plate boundary zone 499 South American Craton 617, 622 South American deformed belt 630, 650, 651,659 South American passive margin 539, 555 South American platform 398 South American Precambrian craton 472 South Atlantic 36, 289 South Australia 337 South Caribbean margin fault 16 South Caribbean marginal fault 15, 16 South Caribbean Plate boundary 477 South Coast fault 478 South Fiji basin 394, 415 South Florida platform 88 South Fork Member 129 South Martfn Garcfa fault zone 300 South Pacific 56 South Sandwich 391 South Sandwich basin 394 Southeast Indian Ocean 583 Southeastern Gulf of Mexico 21, 63, 73 Southeastern Trinidad offshore 425 Southern Basin of Trinidad 448, 489, 496, 506, 538, 539, 540, 542, 544, 551,555 Southern Basin sidewall 492
697 Southern Boreal paleolatitudes 137 Southern Boreal Province 109, 123, 126, 127, 137, 139, 144 Southern Boreal/Northern Tethyan faunas 103, 109, 147 Southern Canada Bonita section 259, 260, 263, 266, 267, 282 Southern Central America 24 Southern Compressional Zone 480, 482, 487, 488, 489, 493 Southern Cuba 281,283 Southern Great Bahama Bank 171, 182, 190 Southern Haiti 576 Southern Hispaniola 5, 52 Southern Mexico 19, 22, 498 Southern Middle America trench 9 Southern Peru 617 Southern platform 83 Southern province 145 Southern Range of Trinidad 477, 478, 488, 489, 491,493, 496, 506, 536, 548, 555 Southern Rosario belt 95, 96, 98, 99, 102, 103, 105-108, 113, 114, 117 Southern Senegal 441 Southern Taino ridge 643 Southern Trinidad 425,437 Southern United States 153 Southern Yucatan margin 115 Southwest Pacific 415 Southwestern Oregon 128 Southwestern Panama 15 Southwestern Puerto Rico 128 Spears of iron 336 Spreading ridge 13 Springvale Formation 482, 487, 488, 492, 511,513, 523 St. Croix 13,343, 345, 347, 353, 354, 357, 358, 360, 362, 363, 364 St. Croix basement rocks 357 St. Croix basin 357, 362 St. Croix Ridge 357, 360 St. John 354, 358 St. John' s/Judith Fancy area 349 St. Thomas 357, 358 Stage poles 42 Stage vectors 44 Stanley Mountain 127, 128 Stanley Mountain cherts 129 Statistical parameters 40 Step 7 Stoneyford 128 Straits of Andros 188 Straits of Florida 65, 74, 168, 188, 190 Stratigraphic and metamorphic terranes 94 Stress inversion technique 208 Strike-slip basins 6, 7 Stuart City Formation 162 Submarine canyons 462, 463 Subsatellite basement topography profiles 37 Subsidence mechanisms 6 Sula graben 201 Sula Islands 111,126 Sula Valley 220, 223 Sulu basin 397, 414
Sumidero Member 88, 102, 103, 143 Superior Oil Company 296, 297, 304 Superterrane 111 Sutawala Valley 153, 162 Swan Islands 197, 201,211,219, 222, 223 Swan Islands fault 197, 201 Swan Islands fault zone 197, 199, 200, 208, 214, 215, 216 Sylvie ridge 660 Symmetric plate accretion 41 Symon 123, 130, 146 Symon remnant 145 Syn-rift phase 419 Syn-rift sequence 63,435 Syn-rift stage 65, 93, 117 Tabasco 145 Tacutu Graben 441,473 Tacutu rift system 441 Taino ridge 627, 629, 641,642, 643, 649, 653, 662, 665 Tairona ridge 629, 653,659 Talanga 153 Talparo Formation 488, 511,513, 524, 526 Taman Formation 140, 141, 146 Taman, S.L.E 140 Tamana Formation 510 Tamana limestone 540 Taman-Tamazunchale 137 Tamaulipas 145 Tamaulipas Formation 146 Tamayo 331 Tamayo-Vuelta Grande road 327, 328, 330 Tampa embayment 83, 87, 88 Tampico 146 Tampico Embayment 146 Tampico-Ciudad Valles line 145 Taninul quarry 161 Taraises Formation 136 Taulab6 area 162 Tavera Group 284 Tectonic evolution 16 Tectono-paleogeographic maps 82 Tectono-stratigraphic megasequence 441 Tectonostratigraphic terrane 144 Tegucigalpa 153, 155 Tela 13, 219, 220 Tela Basin 197, 201, 213-216, 219, 234 Telemaque member 519 Temblador Group (well 'H') 447 Tepehuafio 144 Tepehuafio terrane 145 Tepexic Limestone 140 Terranes 94 Test hole M 10 354 Tethyan ammonites 137, 139 Tethyan Realm 123, 126 Tethys 161 Texaco, Inc 168 Texas 159, 162, 241 Texas megashear 124 Thermal subsidence 89 Third-order cycles 419 Third-order sequences 464 Third-order unconformity 435
698 Three-plate system 18 Tiburon Ridge 48, 49, 55 Tiburon Rise 49, 55 Tigre Formation 447 Time migration 68 Tirisne Cliffs 153, 155, 157, 159, 162 Tobago 14 Tobago trough 14 Todos Santos Formation 155 Toe thrusts 437 Tonga Trench 415 Tongue of the Ocean 167, 171, 174, 185, 191 Tonosf basin 15 Top basement unconformity 435, 438 Top Berriasian 69 Top evaporites/Angostura reflector 301 Top Sombrerito reflector 301 Trans-Atlantic Geotraverse 36 Transfer faults 8 Transform-ridge intersections 39 Transgressive-regressive cycles 443, 472 Transitional crust 74 Transpression 41, 50, 551 Transpressional folded belt 472 Transpressional foredeep 420 Transtension 41 Transtensional basins 477 Trench suction 53 Trinchera Formation 295, 296, 300, 304, 305, 317, 334, 336 Trinidad 6, 14, 419, 421,423,425, 435, 439, 447, 465, 472, 473, 487, 488, 492, 495,497, 498, 503 Trinidad and Tobago 533 Trinidad area 24 Trinidad belt 505 Trinidad-Venezuela boundary 507 Trujillo 219, 220, 223, 227, 234 Tucutunemo 15ormation 112 Tulip-type cross-sectional fault structures 185 Tumbadero Member 88, 101,102, 143 Tumbitas Member 88, 101,102 U.S. 161 U.S. Gulf Coast 155, 163 U.S. Virgin Islands 343 Uncertainties 35 Uncertainty ellipses 56 Unconformity surfaces 63 Unit II 443 Unit III 443 Unit IV 443 United Nations 219 University of Texas at Austin 629 University of Texas at Austin Institute for Geophysics 392 University of Texas seismic data base 65 Upper Atima Formation 162 Upper Cruse Formation 538, 555 Upper La Pica Formation 487 Upper Merecure Formation 466 Upper Nicaragua block 666 Upper Nicaragua rise 666 Upper Talparo Formation 526 Upper Valle de Angeles Group 151, 155
SUBJECT I N D E X Urica 499 Urica fault 493, 503 Urica fault system 420 Urica fault zone 501 Utila 197, 201,223, 224 Utila Island 201, 215 Valanginian Tumbitas Member 143 Valle de Angeles 159 Valle de Angeles Group 155, 156, 157, 159, 161,162, 163, 223 Valle de Angeles redbeds 156 Valle de Pons tectonic unit 97, 106 Valle Formation 126 Vector diagram 18 Vectorial closure condition 18 Vema 629 Vena del Gesso basin 323 Venezuela 4, 6, 50, 241,419, 472, 477, 498, 621,622 Venezuelan abyssal plain 622 Venezuelan Andes 621 Venezuelan Atlantic offshore 473 Venezuelan Basin 5, 13, 16, 28, 33, 50, 51, 56, 241,398, 561,564, 565, 573, 576, 577, 580, 581,582, 584, 586, 587, 591,592, 593, 599, 602, 611, 613, 614, 615, 617, 618, 622, 623, 624, 649, 650, 651,657, 659 Venezuelan Basin rough crust 565 Venezuelan Caribbean Mountains 618, 622 Venezuelan Coastal Ranges 424 Venezuelan Cordillera 477 Venezuelan crust 624 Venezuelan deformed belt 627, 663 Venezuelan fold-and-thrust belt 482 Venezuelan foredeep 471, 618 Venezuelan microplate 51, 52, 627, 662, 663, 666 Venezuelan offshore 465, 448 Venezuelan oil industry 6 Venezuelan Plate 33, 56, 57, 592, 615, 637 Venezuelan shelf 15 Venezuelan-Colombian microplate 665 Veracruz 123, 145, 145, 147 Vicente Noble block 300 Victoria segment 145 Vidofio Formation 448 Vieja Member 108 Vieques 357, 358, 360 Villa de Cura nappe 505 Villa La Reine 349, 350, 354 Villa La Reine type section 351 Villa Trina Formation 247, 251,257, 272, 279, 282-285 Villa Vasquez series 253 Vifiales Limestone 143 Virgin Islands 6, 14 Virgin Islands basin 357, 360, 362, 363, 364 Virgin Islands basin/Anegada Passage area 360 Virgin Islands platform 354, 360, 362, 363, 364 Virgin Islands Trough 343 Vizcaino Peninsula 126
Volcanic plateau 628 VCring Plateau 583 Vcring volcanic margin 573 Vuelta Grande 331 Walkers Cay 190 Walper Megashear 109, 111, 123, 124, 130, 131, 133, 144-147 Walton-Plantain Garden fault 663 Walton-Plantain Garden-Enriquillo fault 663 Wampti-Patuca region 161 Wampusirpi quadrangle 159, 162 Warao rise 653 Warm Springs fault 477, 480, 482, 487, 488, 489, 492, 496, 552 Warm Springs fault system 488, 492 Warm Springs fault zone 496, 498, 501, 506, 507, 521,524, 539, 541,542, 555 Warm Springs Member of the Snowshoe Formation 129 Warm Springs-Central Range 499 Warm Springs-Central Range fault zone 495, 501,503, 533, 554 Warm Springs-Central Range-Caigual fault zone 496, 542, 540, 555 Washita Group 159 Well M 1 350, 356 Well M2 350 Well M4 352, 356 Well M5 356 Well M10 349, 350 West Fiji basin 397 West Florida 446 West Indies 49 West Texas 289, 333 West YucaUin basin 13 West-central Honduras 151 West-central Mexico 130 Western Atlantic 623 Western Brazil basin 241 Western Canada 289 Western Chortfs block 8 Western coast of Hispaniola 637 Western Cordillera 51, 56 Western Cuba 6, 12, 22, 52, 83, 87, 88, 93, 94, 130, 138, 141,144, 147, 552 Western Mediterranean 289 Western Pacific 575 Western Pangea 63, 83 Western Straits of Florida 64 Western Venezuela 22 Wichita megashear 124 Wide-angle reflection data 564 Wilbur Springs 128 Wild Cane complex 244 Wood River Formation 88 Wrangellia terrane 124 Wrench models 477 Wyoming 313 Yojoa Group 151,153, 155, 157 Yoro 220 Yucat~in 63, 87, 89, 93, 108, 145 YucaUin basin 7, 13, 22, 28, 52, 95, 392, 499
699
SUBJECT I N D E X Yucatan block 18, 19, 63, 65, 80, 83, 109, 111,114 Yucatan block edge 96 Yucatan borderland 95, 109, 111 Yucatan margin 111, 114-118 Yucatan Peninsula 12, 13, 22, 52, 144, 199, 546
Yucatan platform 89, 93, 109, 109, 144, 145, 147 Yucatan terrace 69, 79, 80, 81, 89 Yucat~in-Florida Straits 114 Zacar~as Member 100 Zacatecas area 145
Zarza Member 143, 143 Zaza terrane 11 l, 115 Zaza volcanic arc 115 Zero edge 465 Zuloaga Limestone 133, 137, 138
