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.
VI
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.
SEDIMENTARY BASINS OF THE WORLD - - AN INTRODUCTION TO THE SERIES
VII
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.
VIII
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
W.A. van den Bold
X
DEDICATION
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
DEDICATION
XI
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
XII
DEDICATION
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-
DEDICATION
XIII
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-
XIV
DEDICATION
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.
XVI
SEDIMENTARY BASINS OF THE CARIBBEAN - - PREFACE
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.
XVIII
SEDIMENTARY BASINS OF THE CARIBBEAN m PREFACE
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
XX
SEDIMENTARY BASINS OF THE CARIBBEAN m PREFACE
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
XXIV
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.
4
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
5
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
7
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
9
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.
10
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
11
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.
12
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
13
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.
1
4
E
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).
15
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).
1
6
R
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
20
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.
22
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
23
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.
REFERENCES
Av6 Lallemant, H.G., 1997. Transpression, displacement partitioning, and exhumation in the eastern Caribbean/South American plate boundary zone. Tectonics, 16: 272-289. Ball, M.M., Martin, R.G., Bock, W.D., Sylvester, R.E., Bowles, R.M., Taylor, D., Coward, E.L., Dodd, J.E. and Gilbert, L., 1985. Seismic structure and stratigraphy of the northern edge of the Bahamian-Cuban collision zone. Am. Assoc. Pet. Geol. Bull., 69: 1275-1294. Bird, D.E., Hall, S.A., Casey, J.E and Millegan, ES., 1993. Interpretation of magnetic anomalies over the Grenada basin. Tectonics, 12: 1267-1279. Bouysse, E, 1988. Opening of the Grenada back-arc basin and evolution of the Caribbean plate during the Mesozoic and early Paleogene. Tectonophysics, 149: 121-143. Bowland, C.L. and Rosencrantz, E., 1988. Upper crustal structure of the western Colombian basin, Caribbean Sea. Geol. Soc. Am. Bull., 100: 534-546. Bralower, T.J. and Iturralde-Vinent, M.A., 1997. Micropaleontological dating of the collision between the North American plate and the Greater Antilles arc in western Cuba. Palaios, 12: 133-150. Bueno Salazar, R., 1989. Hydrocarbon exploration and potential of the Pacific coastal basin of Colombia. In: G.E. Ericksen, M.T. Pinochet and J.A. Reinemund (Editors), Geology of the Andes and its Relation to Hydrocarbon and Mineral Resources. Earth Science Series, 11, Circum-Pacific Council for Energy and Mineral Resources, Houston, TX, pp. 335-343.
29
Burkart, B. and Self, S., 1985. Extension and rotation of crustal blocks in northern Central America and effect on the volcanic arc. Geology, 13: 22-26. Burke, K., 1988. Tectonic evolution of the Caribbean. Annu. Rev. Earth Planet. Sci., 16: 201-230. Burke, K., Cooper, C., Dewey, J.E, Mann, E and Pindell, J.L., 1984. Caribbean tectonics and relative plate motions. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162:31-63. Case, J.E., MacDonald, W.D. and Fox, EJ., 1990. Caribbean crustal provinces: seismic and gravity data. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Vol. H, Geological Society of America, Boulder, CO, pp. 15-36. Cobbold, ER., Davy, E, Gapais, D., Rossello, E.A., Sadybakasov, E., Thomas, J.C., Tondji Biyo, J.J. and de Urreiztieta, M., 1993. Sedimentary basins and crustal thickening. Sediment. Geol., 86: 77-89. Coffin, M.E, Gahagan, L.M., Lawver, L.A., Lee, T.-Y. and Rosencrantz, E., 1992. Atlas of Mesozoic/Cenozoic reconstructions (200 Ma to Present Day), PLATES Progress Report No. 1-0192, University of Texas, Institute for Geophysics, Technical Rep. 122, 49 pp. Collins, L.S., Coates, A.G., Jackson, J.B.C. and Obando, J.A., 1995. Timing and rates of emergence of the Lfmon and Bocas del Toro basins: Caribbean effects of Cocos Ridge subduction? In: E Mann (Editor), Geologic and Tectonic Development of the Caribbean Plate Boundary in Southern Central America. Geol. Soc. Am. Spec. Pap., 295: 263-289. Corrigan, J.D., Mann, E, Ingle, J.C., Jr., 1990. Forearc response to subduction of the Cocos Ridge, Panama-Costa Rica. Geol. Soc. Am. Bull., 102: 628-652. DeMets, C., Gordon, R., Argus, D. and Stein, S., 1994. Effect of recent revisions to the geomagnetic reversal time scale on estimates of current plate motions. Geophys. Res. Lett., 21: 2191-2194. Denny, W.M., III, Austin, J.A., Jr. and Buffler, R.T., 1994. Seismic stratigraphy and geologic history of middle Cretaceous through Cenozoic rocks, southern Straits of Florida. Am. Assoc. Pet. Geol. Bull., 78: 461-487. Dewey, J.W. and Su(trez, G., 1991. Seismotectonics of Middle America. In: D.B. Slemmons, M.D. Zoback and D.D. Blackwell (Editors), Neotectonics of North America. Decade Map Volume I, Geological Society of America, Boulder, CO, pp. 309-321. Dillon, W., Austin, J.A., Jr., Scanlon, K.M., Edgar, N.T. and Parson, L.M., 1992. Structure and development of the insular margin north of western Hispaniola: A tectonic accretionary wedge on the Caribbean plate boundary. Marine and Petroleum Geology, 9: 70-88. Dixon, T.H., Farina, E, DeMets, C., Jansma, E, Mann, E and Calais, E., 1998. Relative motion between the Caribbean and North American plates and related plate boundary zone deformation from a decade of GPS observations. J. Geophys. Res. 15,157-15,182. Dolan, J.F., Mann, E, de Zoeten, R., Heubeck, C. and Shiroma, J., 1991. Sedimentologic, stratigraphic, and tectonic synthesis of Eocene-Miocene sedimentary basins, Hispaniola and Puerto Rico. In: E Mann, G. Draper and J.F. Lewis (Editors), Geologic and Tectonic Development of the North AmericaCaribbean Plate Boundary in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 217-263. Dolan, J.F., Mullins, H.T. and Wald, D.J., 1998. Active tectonics of the north-central Caribbean: oblique collision, strain partitioning, and opposing subducted slabs. In: J.E Dolan and E Mann (Editors), Active Strike-Slip and Collisional Tectonics
30 of the Northern Caribbean Plate. Geol. Soc. Am. Spec. Pap., 326, 1-61. Donnelly, T.W., Beets, D., Carr, M.J., Jackson, T., Klaver, G., Lewis, J., Maury, R., Schellekens, H., Smith, A.L., Wadge, G. and Westercamp, D., 1990. History and tectonic setting of Caribbean magmatism. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Vol. H, Geological Society of America, Boulder, CO, pp. 339-374. Draper, G., Guti6rrez, G. and Lewis, J.E, 1996. Thrust emplacement of the Hispaniola peridotite belt: orogenic expression of the mid-Cretaceous Caribbean arc polarity reversal? Geology, 24: 1143-1146. Duncan, R.A. and Hargraves, R.B., 1984. Plate tectonic evolution of the Caribbean region in the mantle reference frame. In: W.E. Bonini, R.B. Hargraves and R. Shagam, R. (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162:81-93. Eva, A.N., Burke, K., Mann, E and Wadge, G., 1989. Four-phase tectonostratigraphic development of the southern Caribbean. Mar. Pet. Geol., 6:9-21. Frisch, W., Meschede, W. and Sick, M., 1992. Origin of the Central American ophiolites: evidence from paleomagnetic results. Geol. Soc. Am. Bull., 104:1301-1314. Gealey, W.K., 1980. Ophiolite obduction mechanism. In: Ophiolites, Proc. Int. Ophiolite Symp., Cyprus Geological Survey Department, Nicosia, pp. 228-243. Gonzalez de Juana, C., Arozena, J. and Picard-Cadillat, X., 1980. Geologfa de Venezuela y de sus cuencas petrolfferas. Foninves Ediciones, Caracas (2 volumes), 1051 pp. Gordon, M.B., 1991. The Chortfs block is a continental, pre-Mesozoic terrane. In: D.K. Larue and G. Draper, G. (Editors), Transactions of the 12th Caribbean Geological Conference, St. Croix, U.S. Virgin Islands. Miami Geological Society, Miami, FL, pp. 505-512. Gordon, M.B. and Muehlberger, W.R., 1994. Rotation of the Chortfs block causes dextral slip along the Guayape fault. Tectonics, 13: 858-872. Gordon, M.B., Mann, E, Cficeres, D. and Flores, R., 1997. Cenozoic tectonic history of the North American-Caribbean plate boundary zone in Cuba. J. Geophys. Res., 102: 10,05510,082. Gose, W.A., 1985. Paleomagnetic results from Honduras and their bearing on Caribbean tectonics. Tectonics, 4: 565-585. Grindlay, N.R., Mann, E and Dolan, J.E, 1997. Researchers investigate submarine faults north of Puerto Rico. Eos, Trans. Am. Geophys. Union, 78: 404. Hempton, M.R. and Barros, J.A., 1993. Mesozoic stratigraphy of Cuba: depositional architecture of a southeast-facing continental margin. In: J.L. Pindell and R.E Perkins, R.E (Editors), Mesozoic and Early Cenozoic Development of the Gulf of Mexico and Caribbean Region: A Context for Hydrocarbon Exploration. Gulf Coast Section, Society of Economic Paleontologists and Mineralogists, Houston, TX, pp. 193-209. Heubeck, C., Mann, E, Dolan, J. and Monechi, S., 1991. Diachronous uplift and recycling of sedimentary basins during Cenozoic tectonic transpression, northeastern Caribbean plate. Sediment. Geol., 70: 1-32. Holcombe, T.L., Ladd, J.W., Westbrook, G.K., Edgar, N.T. and Bowland, C.L., 1990. Caribbean marine geology: ridges and basins of the plate interior. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Vol. H, Geological Society of America, Boulder, CO, pp. 231260. Hutson, E, Mann, E and Renne, E, 1998. 4~ dating of single muscovite grains in Jurassic siliciclastic rocks (San Cayetano Formation): constraints on the paleoposition of western Cuba. Geology, 26: 83-86. Jany, I., Scanlon, K.M and Mauffret, A., 1990. Geological in-
E MANN terpretation of combined Seabeam, Gloria, and seismic data for Anegada Passage (Virgin Islands, North Caribbean). Mar. Geophys. Res., 12: 173-196. Kellogg, J.N. and Vega, V., 1995. Tectonic development of Panama, Costa Rica, and the Colombian Andes: constraints from Global Positioning System geodetic studies and gravity. In: E Mann (Editor), Geologic and Tectonic Development of the Caribbean Plate Boundary in Southern Central America. Geol. Soc. Am. Spec. Pap., 295: 75-90. Kerr, A.C., Tarney, J., Marriner, G.E, Nivia, A. and Saunders, A.D., 1997. The Caribbean-Colombian Cretaceous igneous plateau: the internal anatomy of an oceanic plateau. In: J.J. Mahoney and M.E Coffin (Editors), Large Igneous Provinces: Continental, Oceanic, and Planetary Flood Volcanism. Am. Geophys. Union, Geophys. Monogr., 100: 123-144. Kolarsky, R.A., Mann, E and Montero, W., 1995a. Island arc response to shallow subduction of the Cocos Ridge, Costa Rica. In: E Mann (Editor), Geologic and Tectonic Development of the Caribbean Plate Boundary in Southern Central America. Geol. Soc. Am. Spec. Pap., 295: 235-262. Kolarsky, R.A., Mann, E and Monechi, S., 1995b. Stratigraphic development of southwestern Panama as determined from integration of marine seismic data and onshore geology. In: E Mann (Editor), Geologic and Tectonic Development of the Caribbean Plate Boundary in Southern Central America. Geol. Soc. Am. Spec. Pap., 295: 159-200. Ladd, J.W. and Buffler, R.T. (Editors), 1985. Middle America trench off western Central America. Atlas 7, Ocean Margin Drilling Program, Regional Atlas Series, Marine Science International, Woods Hole, MA. Ladd, J.W., Holcombe, T.L., Westbrook, G.K., Edgar, N.T., 1990. Caribbean marine geology: active margins of the plate boundary. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Vol. H, Geological Society of America, Boulder, CO, pp. 261-290. Lara, M.E., 1993. Divergent wrench faulting in the Belize Southern Lagoon: Implications for Tertiary Caribbean plate movements and Quaternary reef distribution. Am. Assoc. Pet. Geol. Bull., 77: 1041-1063. L6bron, M.C. and Perfit, M.R., 1994. Petrochemistry and tectonic significance of Cretaceous island-arc rocks, Cordillera Oriental, Dominican Republic. Tectonophysics, 229: 69-100. Leroy, S., Mercier de Ldpinay, B., Mauffret, A. and Pubellier, M., 1996. Structural and tectonic evolution of the eastern Cayman trough (Caribbean Sea) from seismic reflection data. Am. Assoc. Pet. Geol. Bull., 80: 222-247. Lugo, J. and Mann, E, 1995. Jurassic-Eocene tectonic evolution of Maracaibo basin, Venezuela. In: A.J. Tankard, R. Sufirez Soruco and H.J. Welsink (Editors), Petroleum Basins of South America. Am. Assoc. Pet. Geol. Mem. 62: 699-725. MacDonald, W.D., 1980. Net tectonic rotation, apparent tectonic rotation, and the structural tilt correction in paleomagnetic studies. J. Geophys. Res., 85: 3659-3669. Malfait, B.T. and Dinkelman, M.G., 1972. Circum-Caribbean tectonic and igneous activity and the evolution of the Caribbean plate. Geol. Soc. Am. Bull., 83: 251-272. Mann, E, 1996. Model for the formation of large, transtensional basins in zones of tectonic escape. Geology, 25: 211-214. Mann, E and Burke, K., 1990. Transverse intra-arc rifting: Palaeogene Wagwater belt, Jamaica. Mar. Pet. Geol., 7: 410427. Mann, E and Kolarsky, R.A., 1995. East Panama deformed belt: structure, age, and neotectonic significance. In: E Mann (Editor), Geologic and Tectonic Development of the Caribbean Plate Boundary in Southern Central America. Geol. Soc. Am. Spec. Pap., 295:111-130. Mann, E, Schubert, C. and Burke, K., 1990. Review of Caribbean neotectonics. In: G. Dengo and J.E. Case (Editors),
C A R I B B E A N SEDIMENTARY BASINS: CLASSIFICATION AND TECTONIC SETTING The Caribbean Region. The Geology of North America, Vol. H, Geological Society of America, Boulder, CO, pp. 291-306. Mann, E, Draper, G. and Lewis, J.E, 1991. An overview of the geologic and tectonic development of Hispaniola. In: E Mann, G. Draper and J.E Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 1-28. Mann, E, Taylor, E W., Edwards, R.L. and Ku, T.L., 1995. Actively evolving microplate formation by oblique collision and sideways motion along strike-slip faults: an example from the northeastern Caribbean plate margin. Tectonophysics, 246: 1-69.
Manton, W.I., 1987. Tectonic interpretation of the morphology of Honduras. Tectonics, 6:633-651. Masson, D.G. and Scanlon, K.M., 1991. The neotectonic setting of Puerto Rico. Geol. Soc. Am. Bull., 103: 144-154. McCann, W.R. and Pennington, W.D., 1990. Seismicity, large earthquakes, and the margin of the Caribbean plate. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Vol. H, Geological Society of America, Boulder, CO, pp. 291-306. Molnar, E and Sykes, L., 1969. Tectonics of the Caribbean and Middle America regions from focal mechanisms and seismicity. Geol. Soc. Am. Bull., 80:1639-1684. Ori, G.G. and Friend, EE, 1984. Sedimentary basins formed and carried piggyback on active thrust sheets. Geology, 12: 475-478. Pindell, J.L., 1995. Circum-Caribbean sedimentary basin development and timing of hydrocarbon maturation as a function of Caribbean plate tectonic evolution. In: R.L. Miller, G. Escalante, J.A. Reinemund and M.J. Bergin (Editors), Energy and Mineral Potential of the Central American-Caribbean Region. Earth Science Series, 16, Circum-Pacific Council for Energy and Mineral Resources. Springer, Berlin, pp. 47-56. 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, The Geological Society of America, Boulder, CO, pp. 405-432. PlaNer, G., 1976. Tectonic aspects of the Guatemala earthquake of 4 February 1976. Science, 193: 1201-1208. Protti, M., Gtiendel, E and McNally, K., 1995. Correlation between the age of the subducting Cocos plate and the geometry of the Wadati-Benioff zone under Nicaragua and Costa Rica. In: E Mann (Editor), Geologic and Tectonic Development of the Caribbean Plate Boundary in Southern Central America. Geol. Soc. Am. Spec. Pap., 295: 309-326. Reed, D.L. and Silver, E.A., 1995. Sediment dispersal and accretionary growth of the North Panama deformed belt. In: E Mann (Editor), Geologic and Tectonic Development of the Caribbean Plate Boundary in Southern Central America. Geol. Soc. Am. Spec. Pap., 295: 213-223. Renne, E, Mattinson, J.M., Hatten, C.W., Somin, M., Onstott, T.C., Millan, G. and Linares, E., 1989. 4~ and U-Pb evidence for Late Proterozoic (Grenville-age) continental crust in north-central Cuba and regional tectonic implications. Precambrian Res., 42: 325-341. Rosencrantz, E., 1990. Structure and tectonics of the Yucatan basin, Caribbean Sea, as determined from seismic reflection studies. Tectonics, 9:1037-1059.
31
Rosencrantz, E., Ross, M.I. and Sclater, J.G., 1988. Age and spreading history of the Cayman Trough as determined from depth, heat flow, and magnetic anomalies. J. Geophys. Res., 93: 2141-2157. 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. 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. Schubert, C., 1984. Basin formation along the Bocon6-Mor6nE1 Pilar fault system, Venezuela. J. Geophys. Res., 89: 5711-5718. Sen, G., Hickey-Vargas, R., Waggoner, D.G. and Maurrasse, F., 1988. Geochemistry of basalts from the Dumisseau Formation, southern Haiti: implications for the origin of the Caribbean Sea crust. Earth Planet. Sci. Lett., 87: 423-437. Sinton, C.W., Duncan, R.A. and Denyer, E, 1997. Nicoya Peninsula, Costa Rica: a single suite of Caribbean oceanic plateau magmas. J. Geophys. Res., 102: 15,507-15,520. Speed, R.C. and Westbrook, G.K. (Editors), 1984. Lesser Antilles arc and adjacent terranes. Atlas 10, Ocean Margin Drilling Program, Regional Atlas Series, Marine Science International, Woods Hole, MA. Speed, R.C., Torrini, R., Jr. and Smith, EL., 1989. Tectonic evolution of the Tobago trough forearc basin. J. Geophys. Res., 94:2913-2936. St6phan, J.-F., Mercier de L6pinay, B., Calais, E., Beck, C., Tardy, M., Carfantan, J.-C., Thery, J.-M., Olivet, J.-L., Bouysse, E, Tournon, J., Vila, J.-M., Mauffret, A., Bourgois, J., Blanchet, R. and Dercourt, J., 1990. Rdconstitutions gdodynamiques du domaine cara'fbe du Lias h l'Actuel. Bull. Geol. Soc. Fr., 8: 915-919. Van der Hilst, R.D. and Mann, E, 1994. Tectonic implications of tomographic images of subducted lithosphere beneath northwestern South America. Geology, 22:451-454. Von Huene, R. and 16 others, 1995. Morphotectonics of the Pacific convergent margin of Costa Rica. In: E Mann (Editor), Geologic and Tectonic Development of the Caribbean Plate Boundary in Southern Central America. Geol. Soc. Am. Spec. Pap., 295: 291-307. Westbrook, G.K., Hardy, N.C. and Heath, R.E, 1995. Structure and tectonics of the Panama-Nazca plate boundary. In: E Mann (Editor), Geologic and Tectonic Development of the Caribbean Plate Boundary in Southern Central America. Geol. Soc. Am. Spec. Pap., 295: 91-109. Williams, G.D., Powell, C.M. and Cooper, M.A., 1989. Geometry and kinematics of inversion tectonics. In: M.A. Cooper and G.D. Williams (Editors), Inversion Tectonics. Geol. Soc. London Spec. Publ., 44: 3-15. Wilson, J.T., 1966. Are the structures of the Caribbean and Scotia arcs analogous to ice rafting? Earth Planet. Sci. Lett., 1: 335-338. Wortel, R. and Cloetingh, S., 1981. On the origin of the Cocos-Nazca spreading center. Geology, 9: 425-430.
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) or overestimated (~ >> 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.
REFERENCES
Adamek, S., Tajima, E and Wiens, D.A., 1987. Seismic rupture associated with subduction of the Cocos Ridge. Tectonics, 6: 757-774. Algar, S.T. and Pindell, J.L., 1993. Structure and deformation history of the Northern Range of Trinidad and adjacent areas. Tectonics, 12: 814-829. Biju-Duval, B., Mascle, A., Rosales, H. and Young, G., 1984. Episutural Oligo-Miocene basins along the North Venezuelan margin. In: J.S. Watkins and C.L. (Editors), Drake Studies in Continental Margin Geology. Am. Assoc. Pet. Geol. Mem., 34: 347-358. Bralower, T.J., Hutson, E, Mann, E, Iturralde-Vinent, M., Renne, R and Sliter, W.V., 1993. Tectonics of oblique arc-continent collision in western Cuba, 1. Stratigraphic constraints. Eos, Trans. Am. Geophys. Union, 74: 546. Burke, K. and Seng6r, C., 1986. Tectonic escape in the evolution of the continental crust. In: Barazangi, M. and Brown, L. (Eds.), Reflection Seismology: The Continental Crust. Geodyn. Ser. 14, Am. Geophys. Un., Washington, D.C., pp. 4153. Burke, K., Fox, R and Seng6r, A.M.C., 1978. Buoyant ocean floor and the evolution of the Caribbean. J. Geophys. Res., 83: 3949-3954. Burke, K., Grippi, J. and Seng6r, A.M.C., 1980. Neogene structures in Jamaica and the tectonic style of the northern Caribbean plate boundary zone. J. Geol., 88: 375-386. Byrne, D.B., Suarez, G. and McCann, W.R., 1985. Muertos Trench subduction microplatetectonics in the northern Caribbean? Nature, 317: 420-421. Calais, E., St6phan, J.-F., Beck, Ch., Carfantan, J.-Ch., Tardy, M.,
58 Thery, J.-M., Olivet, J.-L., Bouysse, Ph., Mercier de L6pinay, B., Tournon, J., Vila, J.-M., Mauffret, A., Blanchet, R., Bourgois, J. and Dercourt, J., 1989. l~volution pal6og6ographique et structurale du domaine cara'/be du Lias ~t l'Actuel: 14 6tapes pour 3 grandes p6riodes. C. R. Acad. Sci. Paris, 309: 14371444. Calais, E., B6thoux, N. and Mercier de L6pinay, B., 1992. From transcurrent faulting to frontal subduction: a seismotectonic study of the Northern Caribbean plate boundary from Cuba to Puerto Rico. Tectonics, 11: 114-123. Cande, S.C. and Kent, D.V., 1995. Revised calibration of the geomagnetic time scale for the late Cretaceous and Cenozoic. J. Geophys. Res., 100: 6093-6098. Cande, S.C., LaBrecque, J.L. and Haxby, W.E, 1988. Plate kinematics of the South Atlantic: chron 34 to present. J. Geophys. Res., 93: 13,479-13,492. Cande, S.C., Raymond, C.A., Haxby, W.E and Haxby, W.E, 1995. Geophysics of the Pitman Fracture Zone and PacificAntarctic plate motions during the Cenozoic. Science, 270: 947-953. Case, J.E., Holcombe, T.L. and Martin, R.G., 1984. Map of geological provinces in the Caribbean region. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The CaribbeanSouth America Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162: 1-30. Chang, T., 1987. On the statistical properties of estimated rotations. J. Geophys. Res., 92:6319-6329. Chang, T., 1988. On reconstructing tectonic plate motion from ship track crossing. J. Am. Stat. Assoc., 83:1178-1183. Chang, T., Stock, J. and Molnar, P., 1990. The rotation group in plate tectonics and the representation of uncertainties of plate reconstructions. Geophys. J. Int., 101: 649-661. Collette, B.J., Slootweg, A.P., Verhoef, J. and Roest, W.R., 1984. Geophysical investigations of the floor of the Atlantic Ocean between 10~ and 38~ (Kroonvlag-project). Proc. Kon. Ned. Akad. Wet., Ser. B, 87: 1-76. DeMets, C., Gordon, R.G., Argus, D.E and Stein, S., 1990. Current plate motions. Geophys. J. Int., 101: 425-478. Dixon, T.H. and Mao, A., 1997. A GPS estimate of relative motion between North and South America. Geophys. Res. Lett., 24: 535-538. Dolan, J.E, Beck, Ch. and Ogawa, Y., 1989. Upslope deposition of extremely distal turbidites: an example from the Tiburon Rise, west-central Atlantic. Geology, 17: 990-994. Dolan, J.E, Beck, Ch., Ogawa, Y. and Klaus, A., 1990. EoceneOligocene sedimentation in the Tiburon Rise/ODP Leg 110 area: an example of significant upslope flow of distal turbidity currents. Proc. ODE Sci. Results, 110: 47-62. Ego, E, Sebrier, M. and Yepes, H., 1995. Is the Cauca-Patia and Romeral fault system left or right lateral? Geophys. Res. Lett., 22: 33-36. Elsasser, W.M., 1971. Sea-floor spreading as thermal convection. J. Geophys. Res., 76:1101-1112. Engebretson, D., 1982. Model presentation for displacements between western North America, eastern Eurasia, and adjacent oceanic plates for the past 180 million years. In: CircumPacific energy and mineral resources conference, AAPG Bull. 66: 966, Engebretson, D.C., Cox, A. and Gordon, R.G., 1985. Relative motions between oceanic and continental plates in the Pacific Basin. Geol. Soc. Am. Spec. Pap., 206, 59 pp. England, E and Wortel, R., 1980. Some consequences of the subduction of young slab. Earth Planet. Sci. Lett., 47: 403415. Eva, A.N., Burke, K., Mann, E and Wadge, G., 1989. Four-phase tectonostratigraphic development of the southern Caribbean. Mar. Pet. Geol., 6:9-21. Gripp, A.E. and Gordon, R.G., 1990. Current plate velocities
R.D. MLILLER et al. relative to the hotspots incorporating the NUVEL-1 global plate motion model. Geophys. Res. Lett., 17:1109-1112. Gradstein, EM., Agterberg, EP., Ogg, J.G., Hardenbol, J., van Veen, P., Thierry, J. and Huang, Z., 1994. A Mesozoic time scale. J. Geophys. Res., 99:24,051- 24,074. Harrison, C.G.A. and Lindh, T., 1982. A polar wandering curve for North America during the Mesozoic and Cenozoic. J. Geophys. Res., 87:1903-1920. Hellinger, S.J., 1981. The uncertainties in finite rotations in plate tectonics. J. Geophys. Res., 86: 9312-9318. Heubeck, C. and Mann, P., 1991. Geologic evaluation of plate kinematic models for the North American-Caribbean plate boundary zone. Tectonophysics, 191: 1-26. Heubeck, C., Mann, P., Dolan, J. and Monechi, S., 1991. Diachronous uplift and recycling of sedimentary basins during Cenozoic tectonic transpression, northeastern Caribbean plate margin. Sediment. Geol., 70: 1-32. Kent, D.V. and Gradstein, EM., 1986. A Jurassic to Recent chronology. In: P.R. Vogt and B.E. Tucholke (Editors), The Western North Atlantic Region. The Geology of North America, Vol. M, Geological Society of America, Boulder, CO, pp. 45-5O. Klitgord, K.D. and Schouten, H., 1986. Plate kinematics of the central North Atlantic. In: P.R. Vogt and B.E. Tucholke (Editors), The Western North Atlantic Region. The Geology of North America, Vol. M, Geological Society of America, Boulder, CO, pp. 351-404. Ladd, J.W., 1976. Relative motion of South America with respect to North America and Caribbean tectonics. Geol. Soc. Am. Bull., 87: 969-976. Ladd, J.W., Truchan, M., Talwani, M., Stoffa, P.L., Buhl, P., Houtz, R., Mauffret, A. and Westbrook, G., 1984. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The Caribbean-South America Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162: 153-160. Ladd, J.W., Holcombe, T.L., Westbrook, G.K. and Edgar, N.T., 1990. Caribbean marine geology; active margins of the plate boundary. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Geological Society of America, Vol. H, Boulder, CO, pp. 261-290. Mann, P., 1997. Model for the formation of large, transtensional basins in zones of tectonic escape. Geology, 25:211-214. Mann, P. and Burke, K., 1984. Neotectonics of the Caribbean. Rev. Geophys. Space Phys., 22: 309-362. Mann, P. and Corrigan, J., 1990. Model for late Neogene deformation in Panama. Geology, 18: 558-562. Mann, P., Schubert, C. and Burke, K., 1990. Review of Caribbean neotectonics. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Vol. H, Geological Society of America, Boulder, CO, pp. 307-338. Mann, P., Taylor, E, Edwards, L. and Ku, T., 1995. Actively evolving microplate formation by oblique collision and sideways motion along strike-slip faults: an example from the northeastern Caribbean plate margin. Tectonophysics, 246: 169. Maschenkov, S. and Pogrebitsky, Y., 1992. Preliminary results of Canary-Bahamas Project. Eos, Trans. Am. Geophys. Union, 73: 393-397. Mascle, J., Blarez, E. and Marinho, M., 1988. The shallow structures of the Guinea and Ivory Coast-Ghana transform margins: their bearing on the Equatorial Atlantic Mesozoic evolution. Tectonophysics, 155: 193-210. McKenzie, D., 1972. Active tectonics of the Mediterranean region. Geophys. J. R. Astron. Soc., 30: 109-185. Meijer, P.T., 1992. The dynamics of motion of the South America plate. J. Geophys. Res., 97: 11,915-11,931. Molnar, P. and Stock, J., 1987. Relative motions of hotspots in
N E W C O N S T R A I N T S ON T H E PLATE T E C T O N I C E V O L U T I O N OF T H E C A R I B B E A N the Pacific, Atlantic and Indian Oceans since late Cretaceous time. Nature, 327: 587-591. Mtiller, R.D. and Roest, W.R., 1992. Fracture zones in the North Atlantic from combined Geosat and Seasat data. J. Geophys. Res., 97: 3337-3350. Mfiller, R.D. and Smith, W.H.F., 1993. Deformation of the oceanic crust between the North America and South America plates. J. Geophys. Res., 93: 8275-8292. Mtiller, R.D., Sandwell, D.T., Tucholke, B.E., Sclater, J.G. and Shaw, ER., 1991. Depth to basement and geoid expression of the Kane Fracture Zone: a comparison. Mar. Geophys. Res., 13: 105-129. Mfiller, R.D., Royer, J.-Y. and Lawver, L.A., 1993. Revised plate motions relative to the hotspots from combined Atlantic and Indian Ocean hotspot tracks. Geology, 21: 275-278. Mfiller, R.D., Roest, W.R., Royer, J.-Y., Gahagan, L.M. and Sclater, J.G., 1997. Digital isochrons of the world's ocean floor. J. Geophys. Res., 102:3211-3214. Perfit, M.R. and Heezen, B.C., 1978. The geology and evolution of the Cayman Trench. Geol. Soc. Am. Bull., 89:1155-1174. Pindell, J.L. and Barrett, S.F., 1990. Geological evolution of the Caribbean region; a plate tectonic perspective. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Geological Society of America, Vol. H, Boulder, CO, pp. 405-432. Pindell, J.L. and Dewey, J.F., 1982. Permo-Triassic reconstruction of western Pangea and the evolution of the Gulf of Mexico/Caribbean region. Tectonics, 1: 179-212. Pindell, J.L., Cande, S.C., Pitman, W.C., III, Rowley, D.B., Dewey, J.F., LaBrecque, J. and Haxby, W., 1988. A plate-kinematic framework for models of Caribbean evolution. Tectonophysics, 155: 121-138. PlaNer, G. and Ward, S.N., 1992. Thrust faulting and tectonic uplift along the April 1922, 1991 Costa Rica earthquake. Tectonics, 11: 709-718. Roest, W.R. and Collette, B.J., 1986. The Fifteen Twenty Fracture Zone and the North American-South America plate boundary. J. Geol. Soc. London, 143: 833-843. Rona, EA., 1980. The central North Atlantic Ocean Basin and continental margins: geology, geophysics, geochemistry, and resources, including the trans-Atlantic geotraverse (TAG). National Oceanic and Atmospheric Administration/NOAA Atlas 3. Rosencrantz, E., 1990. Structure and tectonics of the Yucatan basin, Caribbean Sea, as determined from seismic reflection studies. Tectonics, 9:1037-1059. Rosencrantz, E. and Mann, E, 1991. SeaMARC II mapping of transform faults in the Cayman trough, Caribbean Sea. Geology, 19:2141-2157. Rosencrantz, E., Ross, M.I. and Sclater, J.G., 1988. Age and spreading history of the Cayman trough as determined from depth, heat flow, and magnetic anomalies. J. Geophys. Res., 93: 2141-2157.
59
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
66
G.L. MARTON and R.T. BUFFLER
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
b"
C3
4~
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.
Z
>.
0
X
0
E"
Z
>. -]
0
0 Z 0
< O t" -]
86
G.L. MARTON and R.T. BUFFLER
...-:,
~ ,...~
o
.,..~
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
A. PSZCZ()LKOWSKI
N
Igl
~FLORIPA
SGM ~
jKIj~_ I i ~
U ~ / ~ , ~_
,../
.A.Np.BAHA.M.AS!: ~"ql'qlNNNNNNN'qI'qlN'
~) "~11-1~ .
_
NINIR"q"
HAVANA
Ill,, K'~N e \.d,
f
Camagoey Province
200 km
-~.~.,
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
I
95
0A
TERRAN
83c,00 ' I
o APS
0 S LM
22030 '
/Y
o P i n a r del R~o
0 Pinar del R{o 0L
,
25 km
jv / / Mantua
(Y
LP La Co[oma <
N
g
~
/7
-
0 I
84o00'
I
20km l
83030 .
Fig. 2. Location maps (A, C) and tectonic map (B) of the Guaniguanico terrane in the Pinar del Rfo and Havana provinces of western Cuba. (A) Location of the area shown in (B). (B) Tectonic map of the Guaniguanico terrane (partly simplified, based mainly on data taken from: Pszczdtkowski et al., 1975" Pszcz6tkowski, 1977, 1978, 1994b; Piotrowska, 1978; Martfnez and V~izquez, 1987); tectonic units of the Sierra de los Organos belt: VP = Valle de Pons, I = Infierno, G = Sierra de Guane and Paso Real, V = Vifiales, PG = Pico Grande, SG = Sierra de la Gtiira, A = Anc6n, APS = Alturas de Pizarras del Sur; CB = metamorphosed tectonic units of the Cangre belt; tectonic units of the Southern Rosario belt: Z = La Zarza, T = Taco Taco, C = Caimito, CP = Cinco Pesos, LT = Los Tumbos, NP = Niceto P6rez, M = Mameyal, LB = Los Bermejales, PU = Loma del Puerto, LP = La Paloma, LM = Loma del Muerto, J = gabbro and serpentinite of the Jagua massif in the southwestern part of the Alturas de Pizarras del Norte; tectonic units of the Northern Rosario belt: B V = Bel6n Vigoa, NO = Naranjo, D = Dolores, LS = La Serafina, CE = Cangre, CH = Sierra Chiquita, QS = Quifiones; GA = Guajaib6n-Sierra Azul tectonic unit; N - Q = Neogene and Quaternary deposits south of the Cordillera de Guaniguanico; barbed lines denote thrusts. (C) Location map of the Guaniguanico terrane belts: SO = Sierra de los Organos belt, CB = Cangre belt, SR = Southern Rosario belt (in the Sierra del Rosario between Soroa and La Palma, and in the Alturas de Pizarras del Norte between La Palma, Mantua and Guane), NR = Northern Rosario belt, LE = La Esperanza belt, GA = Guajaib6n-Sierra Azul belt; arrows indicate sense of the movement along the Pinar fault.
boundary of the Guaniguanico terrane is located in the southeastern Gulf of Mexico, north of the Pinar del Rio Province (Fig. 1B). To the southwest, the Guaniguanico terrane is covered by NeogeneQuaternary deposits. The wells drilled in the Guanahacabibes Peninsula revealed the metamorphosed rocks of the Guaniguanico terrane (Cangre belt) beneath the Oligocene and Miocene rocks, about 55 km to the southwest of Guane (Fig. 1A). According to Rosencrantz (1996) the Guaniguanico terrane continues to the southwest as a fault-bounded, wedgeshaped block occurring between the Yucatfin basin and the Yucatfin borderland (see also Fig. 1B). If this interpretation is correct, the Guaniguanico terrane may be about 400 km long. The thrust nappes of the Guaniguanico terrane consist of Jurassic to Paleogene rocks (Hatten, 1957; Pszcz6lkowski, 1971) and were formed during the Early Eocene (Pszcz6tkowski, 1977, 1994b). The Eocene tectonic deformation affected north-central Cuba as well. Rigassi-Studer (1963) distinguished the Sierra de los Organos and Sierra del Rosario as two distinct stratigraphic successions. The Meso-
zoic successions of the northern Sierra del Rosario tectonic units and the La Esperanza belt (Fig. 2) are equivalents (Pszcz6tkowski, 1982, 1994a; Rodriguez, 1987). Consequently, the Esperanza belt is considered here as a continuation of the Northern Rosario belt. In the present paper, the Guaniguanico terrane is subdivided into the following tectonostratigraphic belts (from south to north): Cangre (CB), Sierra de los Organos (SO), Southern Rosario (SR), Northern Rosario belt (NR) and the Guajaib6n-Sierra Azul (GA) (Fig. 2C). The stratigraphic successions of the Cangre and Sierra de los Organos belts are similar (Pszcz61kowski, 1985); however, the Cangre belt consists of metamorphic (mainly metasedimentary) rocks. The metamorphic Cangre belt consists of three tectonic units: Mestanza, Cerro de Cabras and Pino Solo (Piotrowska, 1978). These units occur in the southeastern part of the Guaniguanico terrane, as a narrow tectonic belt along the Pinar fault. The Sierra de los Organos belt comprises the tectonic units of (1) the Mogote zone and (2) Alturas de Pizarras del Sur (APS in Fig. 2B). In general, the SO
96 and CB represent the Jurassic platform that subsided in the Tithonian and remained submerged in deepwater pelagic conditions during the Cretaceous and Paleocene. The Rosario belts occur in the eastern and northern parts of the Guaniguanico terrane (Fig. 2). The Southern Rosario belt (SR) extends from Soroa to Mantua and Guane. The 1:250,000 scale geological map of Pszcz6lkowski et al. (1975), published as a part of the geological map of Cuba (Puscharovsky, 1988), and the 1:50,000 scale map (Martfnez and V~izquez, 1987) revealed that the southern Sierra del Rosario tectonic units continue to the west, between La Palma and Mantua (Fig. 2B and C). In fact, a distinction between the SR and SO Jurassic lithology is difficult in this area. Near Mantua, the Jurassic (pre-upper Oxfordian) formations display features characteristic for the Sierra de los Organos belt; to the northeast these formations gradually change their facies features reaching SR characteristics in the Minas de Matahambre-La Palma area. The Northern Rosario belt (NR) occurs north of the SR (in the Sierra del Rosario), and has its counterpart in the Esperanza belt to the west and southwest. The Jurassic to Paleogene rocks exposed in the Martin Mesa tectonic window in the Havana Province (located around the Martin Mesa 1 well, but too small to be shown in Fig. 1) are similar to those known in the Guaniguanico terrane. The limestones, shales and sandstones of Early Cretaceous age drilled there in the Martin Mesa 1 well (Fig. 3) are equivalents of rocks occurring in the Northern Rosario belt (Fig. 4). The Jurassic-Cretaceous and Paleocene rocks of this belt (NR) differ from those occurring in the SO and SR to the south. In general, the Rosario belts are interpreted as the continental margin slope and (partly) adjacent basin floor. The Guajaib6n-Sierra Azul belt 1 (GA) occurs as a single (and narrow) tectonic unit exposed between the Northern Rosario belt and the Bahia Honda terrane (Fig. 2). The GA consists mainly of the Cretaceous shallow-water carbonates (Pardo, 1975), named the Guajaib6n Formation (Herrera, 1961; Pszcz6lkowski, 1978, 1982). Contrary to earlier interpretations (Pardo, 1975; Pszcz6lkowski, 1982) the pre-tectonic location of the Guajaib6n-Sierra Azul belt was probably to the south, but not necessarily atop the deep-water Lower Cretaceous deposits of the Northern Rosario belt as suggested by Iturralde-Vinent (1994, 1996). Rather, this belt could be situated somewhere at the Yucat~in block edge.
1The original name of this belt (Cacarajfcara Belt Pardo, 1975) cannot be maintained because of the existence of the Cacarajfcara Formation established earlier (Hatten, 1957) in western Cuba. This formation occurs in the Rosario and La Esperanza belts only.
A. PSZCZOLKOWSKI
Fig. 3. Lithologic column of the Martin Mesa 1 well, located northeast of Cayajabos, in the Havana Province (data from Segura Soto et al., 1985, simplified): 1 -- pelagic limestones, with intercalations of sandstone and shale; 2 -- detrital limestones; 3 = sandstones, shales and marls; 4 --- tectonic contacts. The Lower Cretaceous rocks and Campanian-Maastrichtian deposits (the Cacarajfcara Formation?) are equivalents of the formations exposed in the Northern Rosario belt of the Guaniguanico terrane (see Fig. 2B).
If so, the GA unit may represent a separate terrane not related to the Guaniguanico tectonostratigraphic belts. The Paleogene thrusting changed the original relative positions of the belts and tectonic units of the Guaniguanico terrane. Contrasting opinions have been expressed on the problem of the pre-tectonic restoration of the Guaniguanico belts and units (Hatten, 1957; Rigassi-Studer, 1963; Pszcz6tkowski, 1978; Mossakovskiy and de Albear, 1979; Piotrowska, 1993; and others). In the present paper, the author accepts the idea, that during the Paleogene thrusting the relative positions of the belts and tectonic units were completely reversed (IturraldeVinent, 1994, 1996). According to this interpretation, the ophiolites and the Cretaceous volcanic arc rocks of the Bahia Honda composite terrane are the structurally highest belts of the Pinar del Rio Province (Hatten, 1957; Pszcz6tkowski and de Albear, 1982; Pszcz6lkowski, 1990; Iturralde-Vinent, 1994).
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
97
Fig. 4. Generalized lithostratigraphic scheme of the Guaniguanico terrane in western Cuba (for location of belts see Fig. 2C); lithology: 1 -- sandstones and shales with intercalations of limestone, 2 -- mafic rocks, 3 = fossiliferous limestones and shales, 4 ---=thick-bedded to massive carbonates (Jurassic in age), 5 = thin- and medium-bedded limestones, 6 = limestones, shales and sandstones (Polier Formation), 7 = radiolarian cherts and shales, 8 = massive, shallow-water limestones (Cretaceous in age), 9 = shales and sandstones (Late Cretaceous and Paleogene in age), 10 - detrital limestones and breccia (Cacarajfcara Formation), 11 = limestone and chert breccias (Anc6n Formation), 12 - Paleogene olistostrome; lithostratigraphic units (circled letters): SC = San Cayetano Formation, A C -Arroyo Cangre Formation (equivalent to the San Cayetano Formation), ES = E1 Sfibalo Formation, J Jagua Formation, F = Francisco Formation, S V = San Vicente Member of the Guasasa Formation, G = Guasasa Formation, AR -- Artemisa Formation, PL -- Polier Formation, L = Lucas Formation, ST = Santa Teresa Formation, GB -- Guajaib6n Formation, P N = Pons Formation, CT = Carmita Formation, PA = Pinalilla Formation, MR = Moreno Formation, PS = Pefias Formation, CA = Cacarajfcara Formation, A N -- Anc6n Formation, M N -- Manacas Formation. =
The presence of the Upper Jurassic shallow-water limestones in the Pinar 1 well (Fig. 5) is a strong argument in favor of the above-mentioned tectonic restoration. This well was located 4 km south of Pons in the Sierra de los Organos (L6pez Rivera et al., 1987; Pszcz6lkowski, 1994), in a tectonic window occurring in the central, most uplifted zone of the Guaniguanico terrane. The Valle de Pons unit (VP in Fig. 2B) is the lowermost tectonic element exposed in the Sierra de los Organos (Piotrowska, 1978) and probably in the whole Guaniguanico terrane. At least three tectonic units were drilled in the Pinar 1 deep well (Fig. 5). The bottom unit, reaching below 3300 m, contains very thick Upper Jurassic shallow-water limestones (1500 m). L6pez Rivera et al. (1987) suggested that these limestones belong to the autochthonous unit. However, there is no conclusive evidence that the Pinar 1 well really penetrated the whole Guaniguanico nappe pile and entered into the autochthonous sedimentary succession. In any case, the Upper Jurassic shallow-water carbonates attain their m a x i m u m thickness (1500 m) in the subsurface. These rocks are thinner in the
higher tectonic units of the Sierra de los Organos belt ( 3 0 0 - 6 0 0 m), and eventually wedge out in the overlying tectonic units of the Southern Rosario belt (Pszcz6tkowski, 1978).
OUTLINE OF STRATIGRAPHY R e m a r k s on lithostratigraphic s c h e m e The lithostratigraphic scheme for the Guaniguanico terrane, as used in this paper (Figs. 4 and 6), was developed during the last 18 years (Pszcz6tkowski et al., 1975; Pszcz6tkowski, 1978, 1982, 1994a). Other authors (Iturralde-Vinent, 1994, 1996; Cobiella-Reguera, 1996) accepted this scheme, although with some minor changes. The results of a recent micropaleontological study on the Paleogene deposits (Bralower et al., 1993; Bralower and Iturralde-Vinent, 1997) also have been considered. The work of Hatten (1957, 1967) was fundamental for developing the m o d e m stratigraphic scheme for the Sierra de los Organos belt (see also Khudoley and Meyerhoff, 1971). Herrera (1961)
98
A. PSZCZOLKOWSKI PINAR
Sierra de los Organos belt
1 EARLY EOCENE
m 0
Southern Rosario belt
MANACAS FM. ;'i... . . . . . . . i
Northem Rosario belt
1
Vieja Member ...................... Pica Pica Member I
PALEOCENE
1000
CACARAJiCARA FM.
IVlAASTRICHTIAb
Lower Cretaceous (Aptian - Albian)
CAMPANIAN
Lower Cretaceous (Valanginian)
S~-E~-%5]q]-R~--
Lower Cretaceous (Aptian - Albian)
CONIACIAN TURONIAN CENOMANIAI~ APTIAN-ALBIAI~ BARREMIAN
"-~,~'~.~'t Upper Jurassic
CARMITA FM. PONS FORMATION
VALANGINIAN
~<,.,,<
BERRIASIAN
Upper Jurassic 2000
'< ''< ''< ,~-.~-)
! Tumbitas Mb. ~u-r~i~aer~' ' Member
KIMMERIDGIAh
u) ~ : , Member < p ............... ~ i D i San Vicente
o o,
Member i
Lower Eocene -
I u_m.i,:r~Mb' :1 ~ ARTEMISA
< ~- FE~i~,?rie-a-ca-n~ .......
TITHONIAN
,'-,>,3 ,,-<
SANTA TERESA FM.
5
La Zarza Member
q I ', FORMATION
~;:,
',
....
i
FRANCISCO FM.
Lower Cretaceous (Aptian - Albian)
-L_ FM.
300C
Lower Cretaceous (Berriasian - ?Barremian)
--
Upper Jurassic Lower Cretaceous (Berriasian - Valanginian) Lower Cretaceous (Aptian - Albian)
" I
SAN CAYETANO FORMATION
Lower Cretaceous (Berriasian - ?Barremian) N M'-N'-M ~ - ~ - ~
mm--
~ - ~ - ~
5
Fig. 6. Lithostratigraphic scheme of the Sierra de los Organos and Rosario belts in the Guaniguanico terrane: L. = lower (Oxfordian), U. = upper (Oxfordian), S.V. = San Vicente Member of the Artemisa Formation.
Upper Jurassic ~ - ~
5000
--'
~ ~ ~
Fig. 5. Lithologic column of the Pinar 1 well, located 4 km south of Pons in the Sierra de los Organos belt (after L6pez Rivera et al., 1987, simplified): 1 = massive, shallow-water limestones, 2 - Berriasian-?Barremian pelagic limestones, 3 -- AptianAlbian pelagic limestones, 4 -- Lower Eocene olistostrome, 5 -tectonic contacts.
introduced many lithostratigraphic names for the Sierra de los Organos area, but only few were valid and have been incorporated into the modern scheme (Fig. 6). Imlay (1942), de la Torre (1960, 1988), Furrazola-Bermddez (1965), Judoley and Furrazola-Bermddez (1968), Wierzbowski (1976), Myczyfiski (1976, 1977, 1989, 1994a,b), Myczyfiski and Pszcz6tkowski (1976, 1990, 1994) and Kutek et al. (1976) studied fossils and/or biostratigraphy of the Jurassic and Lower Cretaceous formations in the Sierra de los Organos belt and Rosario belts.
San Cayetano Formation (?Lower Jurassic to middle Oxfordian) Lithology and facies model. This formation consists of shales, siltstones and sandstones with some intercalations of conglomerates and limestones. Limestones occur mostly in the upper part
of the San Cayetano Formation. The rocks are dark-gray to black; they are rhythmically bedded (Fig. 7). The formation is deeply weathered so that the exposures favorable for sedimentological observations occur mainly in some streams and rivers. The sedimentary structures vary among distinct tectonic units (Meyerhoff and Hatten, 1974). Any reliable lithostratigraphic subdivision of the monotonous and often severely tectonized formation, 1000 to 3000(?) m thick, has not been well established so far. Haczewski (1976) proposed a general descriptive model of sedimentation of the San Cayetano Formation. This author distinguished nine facies (A-I) within the studied sections. In his opinion, the facies A - F occur in the Sierra de los Organos belt. However, the type sections of the facies A - C and E are now considered to belong to the Southern Rosario belt (Pszcz61kowski, 1994b). The facies A D consists of sandstones with some subordinate pebbly sandstones or conglomerates, fine-grained sandstones with trough cross-lamination, of bedded siltstones, shales, and rarely very fine-grained sandstones, siltstones, fine-grained sandstones and thin-bedded shales. The rocks belonging to these facies have been interpreted as deposited in a fluvial environment, and partly in a shallow marine or beach environment. The facies E and F consist of black shales, sometimes with septarian nodules and pyrite spheres in facies E deposits. Both facies
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
99
Fig. 7. Shales, siltstones and fine-grained sandstones of the San Cayetano Formation at Cinco Pesos (Southern Rosario belt). These deposits are similar to facies G and H of Haczewski (1976). most likely have originated in extensive lagoons with restricted circulation. The facies G-I (rhythmic sandstones and shales with graded and ripple bedding, alternating graded sandstones and shales, and thick-bedded sandstones) occur only in the Southern Rosario belt. The facies G consists of fine-grained sandstones, siltstones and shales. Deposits similar to facies G and H are exposed in the Cinco Pesos area (Fig. 7). The thick-bedded, coarse-grained sandstones, sometimes with pebbles up to 7 cm long, are also known in the Cinco Pesos area; the sandstones belong to the facies I of Haczewski (1976). Fossiliferous pebbles, containing late Paleozoic foraminifers and bryozoans, have been found in these sandstones (Pszcz6tkowski, 1989b). The foraminifers belong to Fusulinacea (Schwagerina sp. and Parafusulina sp.) of Permian age. One specimen was identified as Tetrataxis sp. Fossils and age. The age of the San Cayetano Formation (?Early Jurassic-middle Oxfordian) is well established in the uppermost part of this unit only. In the Sierra del Rosario belt, Myczyfiski and Pszcz6tkowski (1976) have found some ammonites in the uppermost part of the San Cayetano Formation, southeast of La Palma. These ammonites belong to the following taxa: Perisphinctes (?Dichotomosphinctes) cayetaensis Myczyfiski, 1976, P. (?Dichotomosphinctes) cf. anconensis Sfinchez Roig, and P. (Discosphinctes) cf. pichardoi Chudoley et Furrazola-Berm6dez. The ammonites indicate that the uppermost part of the San Cayetano Formation is of middle Oxfordian age. In the Sierra de los Organos belt, the siliciclastic deposits of the San Cayetano contain very scarce macrofossils, mainly bivalves. Pugaczewska (1978) identified the following taxa: Eocallista (Hemicorbula) spp., Vaugonia (Vaugonia)
spp., Gervillaria sp., and Neocrassina spp. In the Sierra de los Organos belt, bivalves belonging to the genus Gryphaea Lamarck probably also appear in the coquinid limestones of the upper part of the San Cayetano Formation. Facies interpretation. According to Haczewski (1976), the San Cayetano rocks were deposited on a coastal alluvial plain (facies A-C) by a river transporting the material a few hundred kilometers from the south (or southwest). The sediments debouched to the sea formed an arcuate delta and some of it was redistributed by a longshore drift and turbidity currents. Deposits distinguished as facies G probably accumulated on the slope of the continental margin. Facies H and I were characterized as deposits of a submarine fan accumulating at the base of the slope. The development of facies H is intermediate between normal and proximal turbidites, while facies I is of a proximal character (Haczewski, 1976). The relationship between the diagrams of paleocurrent measurements and the geographical distribution of the localities studied by Haczewski (1976) changes after the restoration of tectonic units. The resulting paleocurrent pattern indicates that the directions from south or south-southwest predominate in the northwestern part of the San Cayetano sedimentary basin (APS in the Sierra de los Organos belt m Fig. 2B), and the directions from northeast prevail in its southeastern sector (Southern Rosario belt).
El S~ibalo Formation (Oxfordian) Lithology and boundaries. The E1 Sfibalo Formation occurs in the Northern Rosario belt (Pszcz6tkowski, 1994a). This unit, up to 400 m thick, consists of basalts and diabases with interbed-
100 ded limestones, marls and sometimes shales. These horizons of pyroclastic rocks also have been observed. The basalts are massive or pillowed flows (Pszcz6tkowski and de Albear, 1983). The lower boundary of the E1 S~ibalo Formation is tectonic; this unit contacts with various Cretaceous and Paleogene formations. Locally, the E1 S~ibalo Formation is overlain by thin San Cayetano Formation siliciclastics, but often contacts with the ?late OxfordianKimmeridgian limestones of the Artemisa Formation. The E1S~ibalo/Artemisa boundary is tectonically disturbed in some sections, but in places a thin bed of sedimentary breccia separates the two formations. Age. The Jurassic age of the E1 S~ibalo Formation (Oxfordian-?early Kimmeridgian) was defined on the basis of infrequent microfossils occurring in the limestones intercalating with the basalts and diabases (Pszcz6tkowski, 1989a, 1994a). According to Cobiella-Reguera (1996), these rocks overlie the San Cayetano Formation clastics and may be as old as Callovian in the lowermost part of the E1 S~ibalo Formation. The alleged position of the E1 S~ibalo Formation rocks above the San Cayetano Formation is not supported by the geological relations visible in outcrops described so far from the Sierra del Rosario. In fact, the San Cayetano-type clastic deposits (up to 15 m thick) were observed in few places above the rocks of the E1 S~ibalo Formation in the Northern Rosario belt (Pszcz6tkowski, 1994b). Therefore, it seems that the E1 S~ibalo Formation is (locally?) older than the Francisco Formation (latemiddle to late Oxfordian). Consequently, the age of the E1 S~ibalo Formation should be pre-late Oxfordian, in some sections at least (Fig. 6). The supposed Callovian age of the lower part of the E1 S~ibalo Formation is, however, still not confirmed by any paleontological or radiometric data. Facies interpretation, The limestones of the E1 S~ibalo Formation have been deposited in anaerobic conditions, below the wave-base, in the outer neritic zone. This origin of this formation was related to the more advanced rifting stage during ?Middle to early Late Jurassic times. According to Iturralde-Vinent (1995a), the magmatic activity of continental passive margin represented in Cuba was time coincident with continental rifting until the Oxfordian, and with oceanic spreading in the Caribbean area since the Late Jurassic.
Jagua Formation (middle and upper Oxfordian) Subdivision and lithology. The Jagua Formation occurs in the Sierra de los Organos belt. This formation consists of limestones and shales subdivided into four subunits, namely the Pan de Azfcar, Zacarfas, Jagua Vieja and Pimienta members. The Pan de Azficar Member (Fig. 6) consists of black,
A. PSZCZOLKOWSKI well-bedded coquinas and bioclastic limestones. The beds and lenses of silicified limestones are common. The coquinas beds are 0.2 to 1.5 m thick. The bivalve-echinoderm and bivalve-oolitic microfacies are very common in the Pan de Azfcar Member. The coquina beds consist mainly of Gryphaea mexicana Felix (Pugaczewska, 1978). The bivalve shells are commonly encrusted by agglutinated foraminifers. The limestones of this member were deposited at a depth ranging from a few up to some tens of meters, in proximity of some oolitic shoals. The Zacarfas Member, not shown in Fig. 6, consists of ammonite-beating shales with thin intercalations of siltstones and bivalve coquinas with Liostrea, Ostrea, Exogyra and Plicatula (Wierzbowski, 1976). This member attains 40 m in thickness and is of middle Oxfordian age (op. cit.). The Jagua Vieja Member comprises black shales and marly limestones, up to 60 m thick. The calcareous concretions contain numerous ammonites, fish remains, marine reptile bones (Iturralde-Vinent and Norell, 1996) and bivalves. The Pimienta Member occurs in the upper part of the Jagua Formation (Fig. 6). This unit, up to 60 m thick, consists of well-bedded, dark-gray to black micritic limestones with some minor shale intercalations in the lower part of the member. Age. The ammonites indicate that the Zacarras Member represents the middle Oxfordian (Wierzbowski, 1976). As judged from the relation to the Zacarfas and Jagua Vieja members, the Pan de Azt~car Member is probably middle Oxfordian in age. The Jagua Vieja Member has been assigned to the middle Oxfordian on the basis of well preserved ammonites (Wierzbowski, 1976). Myczyfiski (1976) described from the Pimienta Member several ammonite species assigned to Mirosphinctes and Cubaspidoceras. One specimen of Taramelliceras (Metahaploceras) sp. also has been found. This ammonite assemblage was assigned to the upper Oxfordian (Myczyfiski, 1994a). Facies interpretation, The bioclastic limestones and coquinas of the Pan de Azfcar Member were deposited at a paleodepth ranging from a few to some tens of meters, in the proximity of some oolitic shoals. The shales of the Zacarfas Member were accumulated in a deeper part of the shelf environment adjacent to a delta. The shales and limestones of the Jagua Vieja Member and the limestones of the Pimienta Member were deposited in an outer shelf environment.
Francisco Formation (middle to upper Oxfordian) Lithology. In the Sierra del Rosario, the Francisco Formation is an equivalent of the Jagua Formation. The Francisco Formation consists of shales, mi-
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
Fig. 8. Spicules of Didemnidae (Didemnum carpaticum Migik et Borza and Didemnum sp.) in a microsparitic limestone of the Francisco Formation (middle and upper Oxfordian), Southern Rosario belt; • 220. critic limestones, and thin sandstone intercalations. Calcareous concretions occur in shales, sometimes with bivalves and/or ammonites. At Cinco Pesos, a volcanic rock (basalt) half a meter thick occurs within limestones and sandstones. The formation attains 25 m in thickness. Age. The Francisco Formation contains some ammonites, rare bivalves, fish and plant remains. Some limestone beds contain Globochaete alpina Lombard and spicules of Didemnidae (Fig. 8), while radiolarians were not observed in thin sections. The ammonites indicate the upper part of the middle Oxfordian and also the lower part of the upper Oxfordian (Kutek et al., 1976; Myczyfiski, 1976, 1994a). Facies interpretation. The deposits of the Francisco Formation were accumulated in an outer shelf environment. This environment was probably slightly deeper than in the case of the Jagua Formation deposits.
Guasasa Formation (?upper OxfordianValanginian) Subdivision and lithology. The Guasasa Formation has been distinguished in the Sierra de los Organos belt. This formation is subdivided into four subunits, namely the San Vicente, E1 Americano, Tumbadero and Tumbitas members (Fig. 6). The San Vicente Member, 300-650 m thick, occurs in the lower part of the Guasasa Formation. This member consists of massive or thick-bedded, gray to black limestones, often dolomitized, sometimes with chert nodules and lenses. Micritic limestones with Favreina predominate in the lower part of the member, while oncolitic and algal calcarenites occur in the upper part. At the top, there are also well-bedded limestones up to a dozen meters thick. The San Vicente Member includes a sedimentary
101
limestone breccia (sharpstone) separating the massive limestones of the Guasasa Formation from the Jagua Formation (Hatten, 1957). A description of the microfacies of San Vicente Member was given by Pszcz6lkowski (1978). The E1 Americano Member comprises well-bedded, dark-gray to black limestones, up to 45 m thick. The Saccocoma-Didemnidae microfacies is characteristic for the limestones of the lower part of this member, while biomicrites with calpionellids and radiolarians occur in its upper part (Myczyfiski and Pszcz6lkowski, 1990). The E1 Americano Member terminates the Jurassic part of the Guasasa Formation (Fig. 6). The Tumbadero Member consists of well-bedded, often laminated, radiolarian biomicrites and calcilutites with black chert intercalations. The thickness of this unit ranges from 20 to 50 m. The Tumbitas Member consists of thick-bedded, light-gray calpionellid, calpionellid-radiolarian and nannoconid biomicrites with infrequent thin intercalations of dark-gray or reddish limestones. The thickness of these deposits is about 40 m, but in some sections attains 80 m. Age. The ?late Oxfordian-Kimmeridgian age of the San Vicente Member is defined mainly on the basis of the lithostratigraphic position of this unit (Fig. 6). The late Oxfordian age of the basal San Vicente Member limestones is suggested by the ammonitebeating Jagua Formation occurring below the massive limestones of this unit (Wierzbowski, 1976; Myczyfiski, 1976). However, the limestone breccia locally separating the San Vicente Member from the Pimienta Member of the Jagua Formation indicates that upper Oxfordian limestones had been partly eroded before the deposition of the shallow-water San Vicente carbonates. A stratigraphic hiatus of unknown duration could be related with this (latest Oxfordian?) erosive event. The Kimmeridgian age of the San Vicente Member carbonates was confirmed by some microfossils (Fermindez Carmona, 1989). The E1 Americano Member of the Guasasa Formation is Tithonian in age. Ammonites collected at the base of the E1 Americano Member are early Tithonian in age (Houga, 1974; Myczyfiski, 1989). Recently, few specimens belonging to the genus Hybonoticeras have been identified in the Sierra de los Organos belt (Myczyfiski, 1996a). These ammonites indicate the Hybonoticeras-Mazapilites Zone in the Sierra de los Organos belt (Fig. 9). Their presence suggests that the age of the upper boundary of the San Vicente Member is close to the Kimmeridgian-Tithonian boundary. The upper Tithonian ammonite assemblage (Myczyfiski, 1989, 1994a, 1996a,b; Myczyfiski and Pszcz6~kowski, 1990) contains some cosmopolitan taxa (Durangites, Corongoceras, Kossmatia), those known from
102
A. PSzCZOLKOWSKI
LU
!-O9
z <
Substage
Upper
z 0 -1!-
~_
Ammonite zones
Durangites - Himalayites Hildoglochiceras (Salinites) Paralytohoplites carribeanus Pseudolissoceras spp.
Lower
Hybonoticeras - Mazapilites Fig. 9. Ammonite zones established in the Tithonian limestones of the E1 Americano Member (the Guasasa Formation), eastern part of the Sierra de los Organos belt (Myczyfiski, 1996a,b and pers. commun., 1996).
Mexico (Hildoglochiceras (Salinites), Proniceras), as well as endemic taxa (Vinalesites rosariensis (Imlay), Protancyloceras hondense Imlay, Butticeras butti (Imlay) and Butticeras antilleanum (Imlay)). Also bivalves Buchia aft. B. okensis (Pavlov) and Buchia aft. B. piochii (Gabb) have been reported from the upper Tithonian limestones of the Sierra de los Organos belt (Myczyfiski, 1989). The ammonites are rare and poorly preserved in the deposits of the Tumbadero and Tumbitas members of the Guasasa Formation. Therefore, stratigraphy of these members is based on calpionellids. The Tumbadero Member is Berriasian in age, while the Tumbitas Member is of late Berriasian-early Valanginian age (Pszcz61kowski, 1978). Facies interpretation. The ?late OxfordianKimmeridgian rocks of the San Vicente Member
were deposited on the shallow-water carbonate bank (Pszcz6ikowski, 1978, 1981) or platform. The Lower Tithonian S a c c o c o m a - D i d e m n i d a e biocalcisiltites of the E1 Americano Member accumulated in a deeper (outer neritic) environment. The upper Tithonian biomicrites of this member were deposited in an outer neritic to bathyal environment. The radiolarian and/or calpionellid biomicrites of the Tumbadero and Tumbitas members are bathyal deposits and were laid down below the aragonite compensation depth.
Artemisa Formation (upper OxfordianValanginian) Subdivision and lithology. The Artemisa Formation is subdivided into three members: San Vicente, La Zarza and Sumidero (Fig. 6). The San Vicente Member was recognized but in a few sections of the Southern Rosario belt. In the Northern Rosario belt, the thick-bedded dolomitized limestones occurring in the lower part of the Artemisa Formation are lithologic equivalent to the San Vicente Member. The La Zarza Member consists of bedded (0.1 to 0.8 m), gray to black micritic limestones (Figs. 10 and 11) with some intercalations of shale, siltstone and fine-grained sandstone in the lower part of this unit. The limestone beds contain rare aptychi and fish remains. In the upper part of the member, there are calcilutites interbedded with dark-gray to black bioclastic limestones and coquinas composed of ammonite shells and aptychi. The thickness of the La Zarza Member attains 200 m. The limestones of this member are overlaid by light-brown, rose and gray to black biomicrites with intercalations of radiolarian chert, assigned to the Sumidero Member (Pszcz6tkowski, 1978). The limestones contain
Fig. 10. Kimmeridgian limestones of the La Zarza Member (Artemisa Formation) exposed in a quarry west of La Palma, Southern Rosario belt.
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
Fig. 11. Black Tithonian limestones of the La Zarza Member, Artemisa Formation, exposed in a road-cut west of Cinco Pesos, Southern Rosario belt.
abundant radiolarians. Calpionellids are common in the lower part of the member, while numerous aptychi occur in its upper part. The Sumidero Member attains 200 m in thickness in some sections. Fossils and age. The Cubaspidoceras-Mirosphinctes assemblage has been identified in the basal limestone beds of the Artemisa Formation (Kutek et al., 1976). These ammonites indicate the late Oxfordian age of these deposits. Consequently, the lower part of the La Zarza Member is late Oxfordian-Kimmeridgian in age. The Tithonian ammonite assemblage occurs in the upper part of the La Zarza Member. These ammonites were studied by Imlay (1942), Judoley and Furrazola-Bermfidez (1968), Houga (1974), and Myczyfiski (1989, 1994b). The taxa Paralytohoplites, Butticeras, Vinalesites, Protancyloceras and Hildoglochiceras (Salinites) are characteristic for the Tithonian of the Rosario belts (Myczyfiski and Pszczdtkowski, 1994; Myczyfiski, 1996a,b). The bivalves occurring in the Tithonian limestones of the Rosario belts belong mainly to the genus Inoceramus (Parkinson, 1819), although representatives of Buchiidae are also present (Myczyfiski and Pszczdtkowski, 1994; Myczyfiski, 1994b). The Saccocoma-Didemnidae microfacies predominates in the lower Tithonian limestones of the Southern Rosario belt, being replaced gradually by the radiolarian microfacies in the Chitinoidella spp. Zone and lower part of the Crassicollaria Zone (Myczyfiski and Pszcz6tkowski, 1994).
103
The calpionellids document a late Berriasianearly Valanginian age of the lower and middle parts of the Sumidero Member. Ammonites attributed to Thurmanniceras cf. novihispanicus (Imlay) indicate a Valanginian age (Myczyfiski, 1977) and nannofossils (Nannoconus spp.) suggest a late Valanginian age for the upper part of the Sumidero Member. The mentioning of the presence of Southern Boreal/Northern Tethyan faunas (30~ in the Valanginian deposits of the Sierra de los Organos and Sierra del Rosario (Pessagno et al. in Chapter 5) needs a comment. Two specimens of Buchia sp. figured by Myczyfiski (1977, pl. 8: 6-7) were collected from the Artemisa Formation at La Catalina. The Tithonian limestones are exposed at this locality (Myczyfiski, 1989; Myczyfiski and Pszcz6tkowski, 1994), but no Cretaceous rocks with macrofauna are known there. Therefore, these specimens of Buchia sp. are late Tithonian (or earliest Berriasian?) in age. The specimens of Vinalesites rosariensis (Imlay) figured by Myczyfiski (1977, pl. 3: 1), were also found in the Artemisa Formation at La Catalina and are late Tithonian in age. In fact, specimens of Buchia from the (stratigraphically well documented) Lower Cretaceous deposits of the Sierra del Rosario and Sierra de los Organos, were not figured in any paper. In this situation, the presence of a Southern Boreal/Northern Tethyan macrofauna in the Valanginian deposits of the Guaniguanico terrane needs confirmation or should be rejected. Facies interpretation. The deposits of the lower part of the La Zarza Member accumulated in relatively quiet conditions of a partly restricted (lagunal?) environment. The limestones of the upper part of the La Zarza Member were deposited in the outer neritic environment, as indicated by its faunal content. Pelagic deposits of the Sumidero Member accumulated in bathyal zones below the aragonite compensation depth, as evidenced by the relative abundance of aptychi and scarcity of ammonites in the radiolarian limestones.
Polier Formation (upper Berriasian-?Albian) Lithology. This formation consists of gray pelagic limestones with intercalations of turbiditic sandstones and shales (Fig. 12). Radiolarian or radiolarian-spicule microfacies are typical of the limestones. Sometimes the sandstones and shales may be distinguished as a distinct subunit in the topmost part of the formation (the Roble Member). The Polier Formation attains its maximum thickness (about 300 m) in the Cangre tectonic unit (CE in Fig. 13). The Polier Formation occurs in the Northern Rosario belt. Similar rocks of Early Cretaceous age are also known in the Esperanza belt (Fig. 2). Age. Calpionellids indicate the late Berriasian age
104
A. P S Z C Z O L K O W S K I
Fig. 12. Limestones, sandstones and shales of the Polier Formation (Lower Cretaceous), exposed west of Cayajabos, Northern Rosario belt; hammer 28 cm long.
Fig. 13. Palinspastic restoration of relative position and width of the major tectonic units occurring in the Guaniguanico terrane (Fig. 2B). The highest units of the Northern Rosario belt were originally located far to the southeast (Iturralde-Vinent, 1994; Pszcz6tkowski, 1994b), while the structurally lowermost units exposed in the Sierra de los Organos belt (VP, /) are shown at the opposite, northwest extreme of this pre-tectonic reconstruction. Tectonic units: VP = Valle de Pons, I = Infierno, V = Vifiales, A = Anc6n, PG & SG = Pico Grande and Sierra de la Gtiira, A P S = Alturas de Pizarras del Sur, CB -- Cangre belt (Cerro de Cabras, Mestanza and Pino Solo metamorphosed units), C + T = Caimito and Taco Taco, Z = La Zarza, A P N = Alturas de Pizarras del Norte (Loma del Muerto and La Paloma units), CP -- Cinco Pesos, MA = Mameyal, B V = Bel6n Vigoa, NO -- Naranjo, CE = Cangre, CH = Sierra Chiquita, QS = Quifiones. Lithology: 1 = sandstones and shales, 2 -- mafic rocks, 3 = fossiliferous limestones and shales, 4 -- massive shallow-water limestones, 5 = thin- to medium-bedded limestones, 6 = limestones, shales and sandstones (Polier Formation), 7 = radiolarian cherts and shales, 8 = sandstones and shales (Late Cretaceous and Paleogene in age), 9 -- detrital limestones and breccia (Cacarajfcara Formation), 10 = limestone and chert breccias (Anc6n Formation), 11 = Paleogene olistostrome; lithostratigraphic units (encircled letters) as in Fig. 4.
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA of the basal beds of the Polier Formation. Ammonite imprints are common in the middle part of this formation. Valanginian to ?Albian taxa have been identified by Myczyfiski (1977). However, the bulk of the deposits is Hauterivian-Barremian in age. Rhyncholites (Hou~a, 1969) and aptychi are common in the Polier Formation. The lower boundary of the Polier Formation is diachronous. Facies interpretation. The turbidites of siliciclastic and mixed (siliciclastic and calcarenitic) composition are the most characteristic constituent of the Polier Formation. The sandstone beds are 0.02-1.0 m thick and most of them are graded, with horizontal and cross-lamination in their upper part (the Bouma sequence). Flute, groove and prod casts are very frequent on the soles of sandstones. Except some thick sandstone beds of the topmost part of the formation, these deposits display features of distal turbidites. Paleocurrent measurements indicate paleoflow toward the south and southeast (Pszcz6tkowski, 1982). The pelagic limestones and turbidites accumulated in the deep-water part of the basin, above the calcite compensation depth.
Lucas Formation (upper HauterivianBarremian) Lithology. The Lucas Formation consists of thin-bedded black limestones with occasional intercalations of fine-grained sandstones and shales. The limestones are radiolarian biomicrites. The thickness of this formation attains 300 m. Age. The Lucas Formation is the stratigraphic equivalent of the Polier Formation (Fig. 6) in the Quifiones tectonic unit (Fig. 13). The aptychi and ammonite imprints are the only macrofossils collected from the Lucas Formation. The ammonites indicate the late Hauterivian to early Barremian age of this formation (Myczyfiski, 1977). The top of the Polier and Lucas formations marks the upper boundary of the Upper Jurassic-Lower Cretaceous limestone sequence in the Northern Rosario belt. Both formations occur directly below the radiolarian cherts and shales of the Santa Teresa Formation. In the NR this important facies change occurred significantly later than in the SR (Figs. 4 and 6). Facies interpretation. The ammonite-bearing radiolarian limestones of the Lucas Formation are interpreted as pelagic deposit accumulated in bathyal environment, less influenced by influx of terrigenous sediment. Santa Teresa Formation (?HauterivianCenomanian) Lithology and distribution. The Santa Teresa Formation is composed of green, red or reddish-
105
brown and sometimes black radiolarian cherts and radiolarites with thin interbeds of silicified shales. Nannoconus-radiolarian biomicrites and thin turbidites occasionally occur in some sections. The thickness of the formation attains 40 m. This formation is exposed in the La Esperanza belt, the Northern and Southern Rosario belts, in the Matanzas Province, in central Cuba and the Camagtiey Province. The author has not seen the Santa Teresa Formation deposits in the Sierra de los Organos belt. Nevertheless, Iturralde-Vinent (1996) suggests the presence of a radiolarian chert unit within the pelagic limestones and cherts of the Pons Formation. Age. Infrequent planktonic foraminifers of Albian-Cenomanian age occur in the Santa Teresa Formation. The lithostratigraphic position suggests that deposition of this formation began during the Hauterivian in the Southern Rosario belt and during the Aptian in the Northern Rosario belt. Radiolarians identified by Aiello and Chiari (1995) indicate a Valanginian to middle Aptian age of one sample taken in the transitional beds from the Polier Formation to Santa Teresa Formation. Facies interpretation. The siliceous rocks of this formation originated in the deep-water environment, near (and sometimes below) the calcite compensation depth. Minor terrigenous input was still active during deposition of the lower part of the Santa Teresa Formation in the Northern Rosario belt.
Pons Formation (?upper Valanginian-Turonian) Lithology. The Pons Formation consists of gray to black micritic limestones interbedded with cherts. The amount of chert intercalations is variable vertically and laterally. The nannofossil-radiolarian limestones and radiolarian cherts are common in the Pons Formation, although other microfacies types are also present in this pelagic sequence. The Pons Formation is developed in the lower tectonic units of the Sierra de los Organos belt only (Fig. 13). The total thickness of Pons Formation is 120 to 150 m. Age. The ?late Valanginian age of the lowermost part of Pons Formation may be assumed on the basis of (1) the lithostratigraphic position of this unit, and (2) the presence of calpionellids (Tintinnopsella carpathica Murgeanu et Filipescu) in the limestones exposed at the base of the type section south of Pons (Pszcz6tkowski, 1978; de la Torre, 1988). Hatten (1957) considered the upper boundary of the Pons Formation to be of Turonian age. Facies interpretation. The pelagic limestones and radiolarian cherts were deposited in a deep bathyal environment, between the aragonite compensation depth and calcite compensation depth.
106 Terrigenous influx was negligible during deposition of the Pons Formation limestones.
Carmita Formation (Cenomanian-Turonian) Lithology. The Carmita Formation is composed of green, red and gray limestones (biomicrites and calcarenites), radiolarian cherts and shales. The calcarenites (calciturbidites) are up to 10 m thick. These graded limestones contain minor amounts of sharp-edged quartz, wacky sandstone fragments, and plagioclase detrital grains. In thin sections, ooids, shallow-water and pelagic limestone fragments are also seen. Maximum thickness of this formation is 70 m. This unit is better developed in the Northern Rosario belt, mainly because of Maastrichtian submarine erosion in the Southern Rosario belt (Figs. 3 and 6). Age. The planktonic foraminifers identified in thin sections by de la Torre (1988) indicate a Cenomanian-Turonian age for the Carmita Formation. The younger age of the marls and calcareous shales of the topmost part of this formation, although possible, has not been confirmed so far. Facies interpretation. The deep-water deposits of the Carmita Formation were accumulated in bathyal environment, below the aragonite compensation depth. The calciturbidites contain a shallowwater debris transported from a carbonate platform. Pinalilla Formation (Cenomanian-Turonian) Lithology. This formation is built of thick-bedded to massive, gray-green limestones, up to 170 m thick. These carbonate rocks are mainly biomicrites containing planktonic foraminifers. The Pinalilla Formation is developed in the Quifiones tectonic unit of the Northern Rosario belt only (Fig. 13). Age. The planktonic foraminifers identified in thin sections demonstrate a Cenomanian-Turonian age. Limestones of the upper part of the formation are middle to late Turonian in age. Facies interpretation. The thick-bedded pelagic limestones of the Pinalilla Formation do not contain any chert intercalations. These limestones were deposited in a bathyal environment, probably less deep than in the case of the Santa Teresa and Carmita formations. Pefias Formation (Campanian-Maastrichtian) Lithology. This formation consists of dark-gray to black, thin-bedded limestones with abundant black chert intercalations (Fig. 14). The limestones are biomicrites containing profuse calcified radiolarians and less numerous planktonic foraminifers. The thickness of the limestones and cherts attains 80 m
A. PSZCZOLKOWSKI in the type section situated south of Pons, in the Sierra de los Organos belt. This section occurs in the Valle de Pons tectonic unit (Fig. 13), while the rocks of the Pefias Formation have been eroded in the higher tectonic units of the Sierra de los Organos belt. Age. The planktonic foraminifers, studied in thin sections, indicate a Campanian-Maastrichtian age for this formation (de la Torre, 1988). According to Hatten (1957) and Meyerhoff (in Khudoley and Meyerhoff, 1971) the Pefias Formation is TuronianCampanian in age. Recently, Iturralde-Vinent (1994, 1996) reported the presence in the Sierra de los Organos belt of a hiatus that spans the Coniacian, Santonian and Campanian. Unfortunately, the results of the paleontological study mentioned by this author are still not published. In the present paper, a hiatus that comprises the late Turonian-Santonian is accepted (Fig. 6). Facies interpretation, Pelagic deposits of the Pefias Formation were accumulated in a bathyal environment, below the aragonite compensation depth. This part of the basin was effectively protected from the terrigenous influx.
Moreno Formation (Campanian) Lithology. The Moreno Formation is composed of marly limestones, detrital limestones, shales, siltstones and sandstones. The marly limestones prevail in the lower part of this formation. The detrital limestones are calciturbidites, up to 5 m thick. These deposits contain common shallow-water bioclasts. Volcanic lithoclasts and quartz are minor constituents of calciturbidites. Shales and sandstones (siliciturbidites) constitute the upper part of the formation. The sandstones contain abundant angular volcanic lithoclasts and plagioclase. This formation in known mainly in the Northern Rosario belt. The maximum thickness of the Moreno Formation rocks is 240 m, but in places this unit was eroded partly or completely at the base of the late Maastrichtian Cacarajfcara Formation (Figs. 4 and 6). Age. The Campanian age of the Moreno Formation is based on identification of planktonic foraminifers in thin sections (Pszcz6tkowski, 1994a). There is a hiatus between the Pinalilla Formation and the overlying Moreno Formation, which spans the Coniacian and Santonian (Fig. 6). A similar hiatus exists between the Carmita and Moreno formations, although the age of the uppermost deposits of the former unit is still not too precise. Facies interpretation. The deep-water, hemipelagic deposits of the lower part of the Moreno Formation originated in a bathyal environment, in conditions of increasing influx of shallow-water and volcaniclastic debris from a volcanic arc terrane.
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA
107
Fig. 14. Pelagic limestones and cherts of the Pefias Formation (Campanian-Maastrichtian); the Las Piedras river, south of Pons, Sierra de los Organosbelt. Hammerlength 28 cm. The shales and sandstones of the upper part of the formation are interpreted as an evidence for convergence between the Upper Cretaceous passive margin and a volcanic arc terrane approaching from the southwest. These terrigenous deposits were accumulated in front of this arc terrane. They are preserved mainly in the southernmost tectonic units of the Guaniguanico terrane (Fig. 13).
Cacarajicara Formation (upper Maastrichtian) Lithology. This formation is developed as detrital limestones composed of breccia, calcarenite and calcisiltite or calcilutite (Hatten, 1957; Pszcz6tkowski, 1978, 1986c, 1994a). The breccia consists of limestone and chert clasts 0.1 to 5 m across, subordinately of shale fragments; upward, it passes gradually into fine calcirudite and coarse calcarenite. Locally, this breccia is up 180 m thick (the Los Cayos Member). The fine calcarenite and calcisiltite constitute the upper part of the formation. Abundant redeposited shallow-water bioclasts of Cretaceous age and relatively infrequent grains of volcanic rocks occur throughout the Cacarajfcara Formation. The planktonic foraminifers are frequent in the upper, fine-grained part of the formation. The thickness of the Cacarajfcara Formation rises from 5-30 m in the SR up to 450 m in the NR. Age. Planktonic foraminifers indicate the upper Maastrichtian Abathomphalus mayaroensis Zone. The list of the identified taxa was given by Pszcz6tkowski (1994a). Facies interpretation. The Cacarajfcara Formation was interpreted as a clastic unit that originated as a result of a single, unusual depositional event close to the Cretaceous/Tertiary boundary (Pszcz6tkowski, 1986b). The Pefialver and Amaro
formations, in western and central Cuba, respectively, are stratigraphic and facies equivalents of the Cacarajfcara Formation. The origin of these massive megabeds was probably related to an extraordinary earthquake and tsunami wave at the end of the Maastrichtian. That unusual event could be associated with the Chicxulub structure of YucaUin (Hildebrand et al., 1991), although an accurate stratigraphic correlation of the Cuban megabeds with the suggested impact site is still not known.
Anc6n Formation (Paleocene) Subdivision and lithology. In the Sierra de los Organos belt, the Anc6n Formation is subdivided into the La Gtiira and La Legua Members (Fig. 6), and the informal member of micritic and marly limestones. In the Southern Rosario belt, the formation is undivided, although the terrigenous rocks occurring locally in the Cinco Pesos tectonic unit may be considered as an informal unit (member). The Anc6n Formation consists of biomicrites, marly limestones, breccias and locally also of terrigenous deposits. Gray, green, and reddish biomicrites contain planktonic foraminifers, calcareous nannoplankton and, sometimes, radiolarians. In the Sierra de los Organos belt, there are some chert nodules and/or lenses in the green biomicrites of the lower part of the formation. The breccias composed of limestone and chert clasts are a characteristic lithology of the formation in the Sierra de los Organos belt. The terrigenous rocks occurring in the Southern Rosario belt contain polymictic sandstones with abundant volcaniclastic fragments. The total thickness of the Anc6n Formation attains 50 m in the Sierra de los Organos and 120(?) m in the Southern Rosario belt. Only few outcrops of this
108
A. PSZCZOLKOWSKI
formation are known in the Northern Rosario belt and the yellowish shales of the Manacas Formation are the main component of the Paleocene deposits there. Age. Thin section analyses of planktonic foraminifers indicate that in the Sierra de los Organos belt the Anc6n Formation is Paleocene in age (Pszcz6~kowski, 1978; de la Torre, 1988). In the Southern Rosario belt, this formation is of late Early Paleocene to Late Paleocene age (Pszcz6~kowski, 1994a). The earliest Paleocene deposits are missing in the Southern Rosario belt. Facies interpretation. The breccia units originated in proximity of synsedimentary faults affecting the Jurassic-Cretaceous limestones (Pszcz61kowski, 1978). Some uplifted limestone blocks supplied a carbonate debris during the Paleocene. Pelagic biomicrites and marly limestones accumulated in deeper areas of the sea bottom (submarine depressions). The volcaniclastic debris locally appeared in the Southern Rosario belt beginning an early stage of the foreland basin development during the (late) Early Paleocene.
belt (Pszcz6~kowski, 1994b). The Vieja Member is Early Eocene in age (Pszcz6~kowski, 1994a,b), but in some sections its lower part may be as old as Late Paleocene (Bralower et al., 1993; Bralower and Iturralde-Vinent, 1997). Facies interpretation. Large amounts of volcanic rock fragments, serpentinite and other types of exotic material in the Manacas Formation define this unit as a foreland basin sedimentary fill (Iturralde-Vinent, 1995b). The deep-water stage with four phases characterize evolution of this basin: (1) deposition of fine-grained sediments (mostly shales), (2) turbiditic sedimentation, (3) development of multi-component olistostrome and (4) serpentinite slide. Phase (4) of the Guaniguanico foreland basin evolution was proposed by Iturralde-Vinent (1996). The deposits of a shallow-water stage, typical of many foreland basins (Covey, 1986), have not been reported from the Manacas Formation.
Manacas Formation (Paleocene-Lower Eocene)
Plate tectonic reconstructions of the Caribbean region
Subdivision and lithology. This formation is subdivided into the Pica Pica Member (lower) and Vieja Member (upper) (Fig. 6). The Pica Pica Member is composed of layered terrigenous deposits: shales, sandstones and conglomerates. Usually, yellow- or red-weathering shales are developed in the lower part of this member. In these shales, there are some breccia interbeds in the southwestern part of the Sierra de los Organos belt. The lithoclastic sandstones and conglomerates, sometimes with infrequent intercalations of marly and detrital limestones, predominate in the upper part of the Pica Pica Member. The Vieja Member consists of cobbles, boulders and larger olistoliths of various rocks enclosed in an argillaceous and/or silty matrix. These olistoliths were derived from the igneous, metamorphic and sedimentary rocks. Fragments of rocks derived from the Guaniguanico stratigraphic successions occur throughout the whole Vieja Member. Olistoliths of the serpentinite and other rocks of the ophiolitic complex are very frequent in this unit. Also volcanic rocks, derived from the Cretaceous volcanic arc, are common, especially in the Rosario belts. The total thickness of the Manacas Formation attains about 500 m in the most complete sections. Age. Planktonic foraminifers occurring in the Manacas Formation are Paleocene-Early Eocene in age (Pszcz6ikowski, 1978, 1994a, 1988). The lower boundary of the formation is diachronous; the oldest deposits (late Early Paleocene) of the Manacas Formation occur in the Northern Rosario
JURASSIC TO EARLY PALEOCENE EVOLUTIONOF THE PASSIVE MARGIN IN WESTERN CUBA
The Jurassic to Early Paleocene evolution of the North American margins exposed in Cuba are considered within the tectonic history of the Yucat~in and Bahamas platform margins, as well as the protoCaribbean basin development. The plate tectonic reconstructions of the Caribbean region (or the Gulf of Mexico-Caribbean system) have been outlined in several papers during the last 15 years (e.g. Pindell and Dewey, 1982; Anderson and Schmidt, 1983; K1itgord et al., 1984; Ross and Scotese, 1988; Pindell et al., 1988; Pindell and Barrett, 1990; Marton and Buffler, 1994). In the present paper, the plate tectonic reconstruction by Marton and Buffler (1994) is accepted as a general paleotectonic basis for the Jurassic. These authors presented the six-stage (Early Jurassic(?) to Berriasian) rift and drift evolution of the Gulf of Mexico and adjacent regions using the plate reconstruction software Plates 2.0. A rotation pole for the Yucat~in block in the southeastern Gulf of Mexico is proposed for the Callovian to Berriasian drifting stage (Marton and Buffler, 1994). The plate tectonic models proposed by Ross and Scotese (1988) and Pindell and Barrett (1990) are adopted for the Cretaceous. These models assume (1) the formation of a proto-Caribbean Sea due to separation of North and South America, (2) the development of the Greater Antilles Arc along the (south)western margin of this sea, and (3) the insertion of the Farallon plate between North and
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA ~ooovv
Callovian
l
/t~
,oovv
/-~",,
I
v 30ON
--
====
'7
~ 1
,'~
.
--~
I
/
I
b*+*+*+1 % % % " 1.+_,.+,+.,_.112 [. =.=,...eJ 3 I"'
--
B
,,
p," '
J4 ~ 5
Fig. 15. Reconstruction of the continental blocks and rift basins between North America an South America during the Callovian (after Marton and Buffler, 1994, simplified and partly modified by the author): BFZ = Bahamas fracture zone" BP = Blake-Bahamas plateau; C = Camajuanf belt of north-central Cuba (Upper Jurassic-Paleogene); E = Escambray terrane; G = Guaniguanico terrane; MSM = Mojave-Sonora megashear; NA = North America; NWB = northwestern Bahamas; P = Pinos terrane; PS -- Placetas belt of north-central Cuba (Proterozoic marbles, Early to Middle Jurassic granitoids and Late Jurassic to Paleogene deposits)" R = Remedios belt of north-central and northeastern Cuba (Jurassic?-Late Cretaceous)" SA = South America; T M V - - Trans-Mexican Volcanic belt; WMT = Western Main Transform; 1 -- continental blocks, 2 -- Proterozoic marbles and Early to Middle Jurassic granitoids of the Placetas belt, 3 = San Cayetano Formation and its metamorphic equivalents now exposed in western and south-central Cuba, 4 - rift basins in the Gulf of Mexico area, 5 = Central Atlantic.
South America resulting in the subduction of protoCaribbean crust. The position of the Guaniguanico terrane in respect to the Yucatfin block during the Jurassic is presented herein (Figs. 15-17) mainly after Iturralde-Vinent (1994, 1996). This author located the Guaniguanico, Pinos and Escambray terranes on the eastern margin of the Yucatfin platform. In fact, the exact original position of these Cuban terranes during the Jurassic and Cretaceous is still uncertain. The location of the Guaniguanico terrane very close to the present northeastern coast of the Yucatan Peninsula, as suggested by Iturralde-Vinent (1994, 1996), may be difficult to sustain, mainly because the remnants of the Mesozoic successions occurring in western Cuba were not reported from the eastern margin of the Yucat~.n block (Viniegra, 1971; Lopez Ramos, 1975). Probably, the Guaniguanico Mesozoic successions were deposited about 100 km to the east and northeast of the present northeast Yucatan coast. Such a possibility is not contradicted by the existing data on the Yucatan borderland topography and structure (Rosencrantz, 1990, 1996). This province extends from western Cuba to Honduras and has an average width of 100 km. In general, the Yucatfin
109
borderland represents the eastern extension of the Yucatfin platform. However, the northern part of this borderland and western Cuba may be structurally continuous (Fig. 1B). The Guaniguanico terrane rocks and the northern part of the Yucatfin borderland together form a large wedge-shaped block now isolated within the transform domain (Rosencrantz, 1996, figs. 2 and 4). Deformed metaterrigenous rocks dredged from the northern Yucatan borderland (Pyle et al., 1973) may represent equivalents of the Cangre belt located in the Guaniguanico terrane and in the subsurface of the Guanahacabibes Peninsula. According to Pessagno et al. (Chapter 5) the Jurassic and Early Cretaceous successions in western Cuba (Sierra del Rosario and Sierra de los Organos) show lithostratigraphic, paleobathymetric, and paleolatudinal signatures which are nearly identical to those of San Pedro terrane remnants in central Mexico. Pessagno et al. conclude: "These Cuban remnants are allochthonous when compared to surrounding Central Tethyan successions in the Blake Bahama basin and elsewhere in Cuba. They contain high latitude bivalves such as species of B u c h i a that can only be derived (exclusive of Greenland) from a Pacific source. The presence of Southern Boreal/Northern Tethyan faunas (30~ in the Sierra de los Organos and Sierra del Rosario remnants as late as the Early Cretaceous (Valanginian) suggests much later tectonic transport by northwest to southeast movement along the Walper Megashear and by subsequent southwest to northeast movement as the Caribbean plate plowed its way through the gap between the North American and South American plates." Also Pszcz6tkowski and Myczyfiski 2 suggested that the occurrence of the bivalves B u c h i a and the radiolarians P a r v i c i n g u l a sp. and Pantanellidae in the Tithonian of western Cuba may indicate (according to the criteria of Pessagno et al., 1993) the Southern Boreal Province or the Northern Tethyan Province, and that the Jurassic sequences of western Cuba could have been located in the Pacific during Tithonian time. They considered, however, that this interpretation requires further investigations. A paleomagnetic study of hand-samples collected from the Jurassic rocks of the Sierra de los Organos and Sierra del Rosario revealed postfolding magnetisation in mafic rocks of the E1 Sfibalo Formation, but no meaningful results were derived from samples of the San Cayetano and Artemisa formations due to the weak magnetization and/or large scatter of paleomagnetic data (Bazhenov et al., 1996). P6rez Lazo et al. (1995) studied their samples from the San Cayetano Formation of the Sierra del Rosario, and from the Collantes Formation marbles of the Escam2 'Information and 1994 Annual Report', Institute of Geological Sciences, Polish Academy of Sciences, Warsaw, 1995, p. 15.
110
A. PSZCZOLKOWSKI
Fig. 16. (A) Reconstruction of the continental blocks around the proto-Caribbean Sea during the middle Oxfordian (partly adapted from Marton and Buffler, 1994 and Iturralde-Vinent, 1994): NR = Northern Rosario belt, SR = Southern Rosario belt, SO = Sierra de los Organos belt, E = Escambray terrane, P -- Pinos terrane, BFZ = Bahamas fracture zone. (B) Enlarged rectangle shown in Fig. 1A; 1 = oceanic crust, 2 -- E1 Sfibalo Formation (mainly mafic rocks), 3 = clastic deposits of the San Cayetano Formation and equivalent metamorphosed units (Arroyo Cangre Formation, and others), 4 -- ammonite-bearing clastic deposits of the uppermost San Cayetano Formation in the Southern Rosario and La Esperanza belts, 5 = bivalve-bearing deposits of the uppermost part of the San Cayetano Formation in the Sierra de los Organos and southwestern part of the Southern Rosario belt (mainly the Loma del Muerto tectonic unit), 6 = land areas.
Fig. 17. (A) Reconstruction of the proto-Caribbean basin for the Tithonian, partly adapted after Marton and Buffler, 1994. (B) Enlarged area indicated by rectangle in (A). Yucatfin passive margin: SO = Sierra de los Organos belt, SR = Northern and Southern Rosario belts; Florida-Bahamas passive margin: P = Placetas belt, C -- Camajuanf belt, R = Remedios belt; PC = proto-Caribbean Sea, BFZ --- Bahamas fracture zone; 1 = oceanic crust, 2 -- pelagic ammonite-bearing biomicrites and calcarenites, 3 = terrigenous deposits, 4 = shallow-water carbonates, 5 - land areas. b r a y terrane. T h e s e s a m p l e s y i e l d e d p a l e o m a g n e t i c i n c l i n a t i o n indicative of a p a l e o l a t i t u d e of about 12 ~ Pdrez L a z o et al. (1995) c o n c l u d e d that their paleo m a g n e t i c result c o r r o b o r a t e s the earlier g e o l o g i c a l i n t e r p r e t a t i o n s c o n c e r n i n g the original l o c a t i o n of the C u b a n Jurassic rocks " n o t too far f r o m H o n d u r a s and G u a t e m a l a . . . " (op. cit.), that is, s o u t h w e s t of its p r e s e n t - d a y position. M o r e o v e r , Pdrez L a z o et al.
(1995) inferred f r o m their p a l e o m a g n e t i c data conc e r n i n g the L o w e r C r e t a c e o u s and A p t i a n - T u r o n i a n rocks, that the d e v e l o p m e n t of the volcanic arc in C u b a t o o k p l a c e at p a l e o l a t i t u d e values o f a b o u t 16-17~ C o n s e q u e n t l y , the e n t r a n c e to the protoC a r i b b e a n b a s i n c o u l d be b l o c k e d by the C u b a n or the G r e a t e r A n t i l l e s volcanic arc (see also R e n n e et al., 1991 for the A p t i a n - C e n o m a n i a n p a l e o l a t i t u d e
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA of their Zaza terrane). This result constrains (or even contradicts?) a possibility of tectonic transport of the 'Sierra de los Organos and Sierra del Rosario remnants' by southwest (from Pacific) to northeast movement after the Barremian?-Aptian. The location of the San Cayetano basin along the Yucat~in margin during Middle and Late Jurassic (Oxfordian) times cannot be defined exactly by the available geological data. The Mesozoic successions of western and south-central Cuba suffered severe tectonic deformations, including large-scale thrusting and metamorphism. During the Paleogene, these successions were thrust to the north, and probably sheared off from the Yucatan block, as the inactive volcanic arc passed along the eastern margin of this platform (Ross and Scotese, 1988; Pindell and Barrett, 1990; Hutson et al., 1998). The presence of a transform boundary parallel to the Yucatan borderland shows that the arc moved northward from a location south of the Yucat~in Peninsula (Rosencrantz, 1996). Nevertheless, various indirect geological evidences constrain the inferred position of the Guaniguanico terrane during the Jurassic. Only a part of these arguments may be expressed herein, because of limited space. The San Cayetano Formation facies configuration, reconstructed from the present-day internal tectonostratigraphic pattern of the Guaniguanico terrane, results as roughly parallel to the Yucat~in eastern margin. The directions of sediment transport from south or south-southwest measured in the San Cayetano Formation sandstones of the Sierra de los Organos belt (Haczewski, 1976), are compatible with the Caribbean paleotectonic reconstructions proposed by Ross and Scotese (1988) and Pindell and Barrett (1990) for the Middle JurassicOxfordian and Late Triassic-Early Jurassic, respectively. In general, the composition of the San Cayetano sandstones (Pszcz6tkowski, 1986a; Hutson et al., 1998) can be explained on the basis of the geological structure and Jurassic history of Yucat~in (Lopez Ramos, 1975). The shallow-water to neritic San Cayetano, Jagua and Francisco formations gradually disappear from northwest to southeast. These units are laterally replaced by the deep neritic E1 S~ibalo Formation. The carbonate bank of the Sierra de los Organos belt (Pszcz6lkowski, 1978, 1981) is a clear evidence of shallowing phase close to the Oxfordian/Kimmeridgian boundary; this phase is characteristic for the Late Jurassic successions of western Cuba. A deepening phase occurred at the Kimmeridgian/Tithonian boundary or during the earliest Tithonian. Both events seems to be unknown in Mexico, especially in the sections described by Pessagno et al. (Chapter 5) as belonging to the San Pedro del Gallo terrane.
111
The Oxfordian faunal assemblages of the San Cayetano and Jagua formations do not contain any taxa characteristic for high latitudes (see also Wierzbowski, 1976 and Myczyfiski, 1976, 1994a). These taxa were also not reported from the Kimmeridgian carbonate rocks of the Sierra de los Organos and Sierra del Rosario belts. In the Tithonian limestones of the Sierra de los Organos belt there are frequent specimens of bivalves belonging to the genera Anopaea and Buchia (Myczyfiski, 1999). Some specimens of Anopaea sp. resemble taxa described earlier from Tithonian deposits of Antarctica, Himalayas, New Zealand, and Sula Islands (Myczyfiski, 1999). In the Sierra del Rosario, the Tithonian bivalves belong mainly to the genus Inoceramus (Parkinson, 1819), although representatives of Buchiidae are also present (Myczyfiski, 1994b). The occurrence of Tithonian bivalves characteristic for northern and southern high latitudes may be explained by upwelling of deeper waters in the northwestern part of the proto-Caribbean basin (Coleman et al., 1995). The Tithonian ammonites of western and central Cuba (Imlay, 1942, 1980; Millfin and Myczyfiski, 1978; Myczyfiski, 1989, 1994a; Myczyfiski and Pszcz6tkowski, 1990, 1994) are so similar, that it would be unrealistic to place them in two different, widely separated, oceanic basins. Therefore, any attempt to locate the Guaniguanico terrane at a higher paleolatitudinal position (30~ during the Tithonian, far from the proto-Caribbean basin, should be accompanied by a similar shift of the Escambray terrane, as well as of the Placetas and Camajuanf belts of central Cuba. In this case, the Guaniguanico, Escambray (and Pinos) terranes and the Placetas and Camajuanf belts must be considered as parts (fragments) of a composite terrane, or superterrane, 8001000 km long and 150-200 km wide, transported (1) from NW to SE, and later (2) from SW to NE (or NNE). However, this tectonic shift would leave the Bahamas platform and the Yucatan block without deep-water successions of their proto-Caribbean continental slope and adjacent basin floor. Considering the aforementioned problems, it seems that the hypothesis of tectonic transport of the Sierra de los Organos and Sierra del Rosario successions along the Walper Megashear still needs a lot of evidences and probably additional studies.
Syn-rift stage (Lower Jurassic-?Callovian/early Oxfordian) South America was close to the Yucatan block during the late Middle Jurassic (Fig. 15). Some leftlateral transform motion probably occurred between these two continental blocks (Marton and Buffler, 1994). Such a situation could exist also during the
112 early Oxfordian with a rift zone (Iturralde-Vinent, 1994, 1996) or rift/spreading center (Marton and Buffler, 1994) between the two continental blocks. The basement of the Mesozoic sedimentary successions, that originally belonged to the Yucat~in eastern margin, is unknown in the Guaniguanico terrane. The thrust units of this terrane do not contain any basement rocks at their base, because the initial tectonic detachment occurred above the basement/sedimentary cover boundary. During the sedimentation of the San Cayetano deltaic deposits, their source areas probably consisted of metasedimentary and terrigenous rocks, and also granitoids (Pszcz6tkowski, 1978). A prolonged transport, recycling of the pre-Middle Jurassic terrigenous rocks, and weathering resulted in the high content of quartz in the San Cayetano sandstones (Pszcz6tkowski, 1986a). In the Southern Rosario belt, rare pebbles of the heavily silicified limestones occur in the thick-bedded pebbly sandstones. Some of those pebbles contain the late Paleozoic fossils (Pszcz6tkowski, 1989b). Evidently, these pebbles were derived from the limestone succession situated some hundreds kilometers from the San Cayetano basin during the early to middle Oxfordian. According to Donnelly et al. (1990) a correlation of fossiliferous pebbles found in the San Cayetano Formation with Permian beds of eastern Guatemala is not apparent. However, any conclusive explanation of the provenance of these pebbles is difficult, mainly because of small amount and size of fossiliferous clasts. At present it is not possible to demonstrate any direct connection between the San Cayetano basin and the northwestern edge of South America as a source of the clastic material for the Jurassic deposits in western Cuba. The hypothesis for a South American provenance for the San Cayetano clastics expressed by some authors (Anderson and Schmidt, 1983; Klitgord et al., 1984; Ryabukhin et al., 1984) is still to be proved. Apparently, the fossils and microfacies of the silicified limestone pebbles from the San Cayetano Formation are not similar to the Permian limestones occurring in the Palmarito and Tucutunemo formations known in northern South America (Benjamini et al., 1987). Instead, some data are consistent with a possible connection of the San Cayetano depocenter with Central America (Meyerhof, in Khudoley and Meyerhoff, 1971; Meyerhoff and Hatten, 1974; IturraldeVinent, 1975; Haczewski, 1976; Pszcz6ikowski, 1987). Prior to the Oxfordian, the San Cayetano clastics were located in the originally narrow, but steadily widening, rift zone formed between Yucat~in and South America (Fig. 15). According to Hutson et al. (1998), the presence of grains with Taconic and Grenvillian ages supports a Yucat~in
A. PSZCZOLKOWSKI source for the San Cayetano Formation. The extent of the San Cayetano depocenter to the southwest, along the Yucat~in margin, is still uncertain. The model of propagating westward rift/spreading center between Yucat~in and South America, with simultaneous counterclockwise rotation of the former block (Marton and Buffler, 1994), requires facies shift from the northeast to the southwest in the San Cayetano Formation. Indeed, some existing data agree with this paleotectonic model. The most obvious observation pertains to the facies changes within the San Cayetano deposits in the Southern Rosario belt between La Palma and Mantua (Pszcz6ikowski, 1994b). From NE to SW, the facies G-I disappear and facies A-C and E (Haczewski (1976) are dominant southwest of Minas de Matahambre, mainly in the Loma del Muerto tectonic unit (Fig. 2). These facies changes are parallel to the regional strike of the present-day tectonic structures (NE-SW). The San Cayetano Formation deposits, about 1400 m thick, also occur in the subsurface of the southwestern part of the La Esperanza belt (Los Arroyos 1 w e l l - Fern~indez et al., 1987; for location see Fig. 1). Deposition of the San Cayetano clastics was accompanied by syn-sedimentary magmatic activity, mainly of a mafic character (Piotrowski, 1977; Pszcz6tkowski, 1978; Iturralde-Vinent, 1995a).
Drift stage (?Callovian/middle OxfordianSantonian) During the middle Oxfordian, the gap between Yucat~in and northwestern South America widened, and the narrow proto-Caribbean seaway was formed. A paleogeographic location of the Guaniguanico, Escambray and Pinos terranes during middle Oxfordian time is shown in Fig. 16A. The presence of the mafic rocks in the Northern Rosario belt (the E1 S~ibalo Formation) suggests that the oceanic crust formation in this gap commenced during the middle Oxfordian or earlier, south of the Northern Rosario and La Esperanza belts. Recently, the Callovian sea-floor spreading was interpreted to have started simultaneously in the Gulf of Mexico and in the Caribbean (Marton and Buffler, 1994). Marine conditions developed before the middle Oxfordian in all belts of the Guaniguanico terrane. In the Sierra de los Organos succession, the first limestone intercalations with marine fauna occur about 400 m below the top of the San Cayetano Formation. Nevertheless, exact timing of the onset of marine deposition in all belts of the Guaniguanico terrane is still to be established. Fig. 16B shows a reconstruction of facies during the middle Oxfordian. Infrequent ammonites have been found only in the uppermost San Cayetano clastics in the Southern Rosario
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA belt. These ammonite-beating clastic deposits accumulated between the E1 S~balo limestones and basalts to the southeast and the shallow-water bivalve-bearing sandstones and shales to the northwest. The advance of the middle Oxfordian transgression resulted in a major facies change, when the San Cayetano deltaic sediments were replaced by limestones with bivalves and shales with ammonites (Sierra de los Organos and Cangre belts) and ammonite-bearing limestones and shales (Southern Rosario belt). Facies differentiation in the W - E (or NW-SE) direction existed during deposition of the Jagua and Francisco formations. Surprisingly, marine macrofossils were not found in the sedimentary rocks of the E1 S~balo Formation. The limestones occurring between the mafic rocks of this formation contain an impoverished microfossil assemblage dominated by Globochaete alpina Lombard indicating a deeper (and partly restricted?) depositional environment. The contact between the E1 S~balo and Artemisa formations is erosional and, in some sections, tectonic. Locally, thin breccia with volcanic clasts occurs at the E1 S~balo/Artemisa formations boundary. This boundary may be interpreted as the unconformity below the basal Artemisa Formation of ?late Oxfordian-early Kimmeridgian age roughly correlatable with the limestone breccia at the base of the Guasasa Formation in the Sierra de los Organos belt. The onset of the carbonate shallow-water sedimentation in the Sierra de los Organos and Cangre belts occurred in the late Oxfordian or earliest Kimmeridgian. Shallow-water carbonates are known also from the Southern Rosario and Northern Rosario belts. Development of the shallow-water bank above the E1 S~balo Formation in the Northern Rosario belt (in the Bel6n Vigoa and Naranjo tectonic units Fig. 13), with manifestations of erosion of basalts and diabases, indicates that this volcano-sedimentary sequence has been locally uplifted during late Oxfordian or earliest Kimmeridgian time. These local uplifts (rotated fault blocks?) could be a barrier inhibiting free communication of the Sierra de los Organos and Southern Rosario belts with the open, but still narrow, proto-Caribbean Sea. The fine-grained limestones (Fig. 10) and clastics (mainly shales and siltstones) accumulated in the inner, semirestricted, part of the Southern Rosario belt. During the Kimmeridgian, subsidence kept pace with the relatively high rate of sedimentation. About 400 to 650 m of shallow-water limestones and dolomitic limestones formed in the Sierra de los Organos belt (San Vicente Member in Figs. 6 and 13). The transition from shallow-water deposition to pelagic conditions of sedimentation occurred close to the Kimmeridgian/Tithonian boundary. In the Sierra de los Organos, this change was rather grad-
113
ual, with appearance of some pelagic microfossils (Saccocoma sp., Colomisphaera spp. ) and deposition of a few thin-bedded limestone units within the thick-bedded to massive calcarenites of the upper part of the San Vicente Member. In the Northern Rosario belt, the lower Tithonian limestones overlay the Kimmeridgian shallow-water dolomitic limestones. Drowning of the shallow-water carbonates resulted in a considerable uniformity of facies in all belts of the Guaniguanico terrane (Fig. 17). At Tithonian time, the sedimentation rate (8-10 m/m.y.) of the black pelagic limestones (Fig. 11) was some ten times lower than that of the Kimmeridgian shallow-water carbonates (80-100 m/m.y.). The Saccocoma-Didemnidae microfacies that predominated in the lower Tithonian ammonite-bearing limestones, was replaced gradually by a radiolarian microfacies (Myczyfiski and Pszcz6~kowski, 1994). The latter microfacies is typical for the upper Tithonian and Lower Cretaceous limestones in all Guaniguanico belts. In the Rosario belts, favorable conditions for radiolarians existed since the early Tithonian. Moderate fertility of the northern protoCaribbean surface waters is also suggested by an elevated content of phosphatic grains, mainly abundant fish debris (bones, scales and teeth) in the upper Tithonian limestones of the Sierra de los Organos. Radiolarian limestones are also common in the Camajuanf succession of central Cuba (Fig. 17) and in the Camagfiey Province. In some sections, there are frequent specimens of inoceramids belonging to the genera Anopaea and Buchia (identified by Dr. R. Myczyfiski). The occurrence of Tithonian bivalves characteristic for high latitudes is consistent with upwelling of cold, oxygenated, and nutrient-rich deeper waters in the northern (or northwestern) part of the proto-Caribbean basin. According to Baumgartner (1987), off-shore winds created upwelling and high fertility of the surface waters in the Late Jurassic 'Caribbean Tethys'. A Jurassic/Cretaceous boundary event, marked by positive shift in the both carbon and oxygen curves seems to be specific for the proto-Caribbean basin, or even for the Caribbean-Gulf of Mexico region (Coleman et al., 1995). The possible explanation for this proto-Caribbean event is the invasion of geochemically different water masses across the Jurassic-Cretaceous boundary (op cit.). During the Berriasian and Valanginian, pelagic limestones with chert interbeds accumulated in the northwestern part of the proto-Caribbean basin. By late Berriasian-early Valanginian, gray nannoconid limestones with abundant calpionellids were deposited in the Sierra de los Organos belt. Basinward, these thick-bedded, pure pelagic limestones passed gradually into black, thin-bedded radiolar-
114
A. PSZCZOLKOWSKI
ian limestones occurring in the Northern Rosario belt. The ammonites are uncommon in the pelagic limestones of Berriasian and Valanginian age (Myczyfiski, 1977). At the Tithonian/Berriasian boundary, the basin floor probably descended below the aragonite compensation depth. In the Northern Rosario belt (farthest south sections: BV, NO, CE, CH and QS, in palinspastic reconstruction Fig. 13), the first turbidite sandstones appeared in the late Berriasian deposits. Nevertheless, the main influx of siliciclastic material in this belt occurred during the ValanginianBarremian. A similar petrographic composition of Hauterivian-Barremian siliciturbidites occurs in the Northern Rosario belt, La Esperanza belt and Placetas belt in the Matanzas Province and north-central Cuba indicating a common source for the clastic material (Fig. 18). The Northern Rosario and La Esperanza belts were situated nearer to the source area. This conclusion results from established differences in abundance and thickness of siliciturbidites between the Northern Rosario and La Esperanza belts and Placetas belt (Pszcz6{kowski, 1982, 1987). The source area probably was located at the northeastern end of the Yucatfin block (Fig. 18). However, terrigenous deposits do not appear in the Sierra de los Organos and Southern Rosario belts, although they were situated relatively close to this hypothetical source area. Accepting the idea that the Paleogene thrusting completely reversed the relative positions of the belts and tectonic units (Fig. 13), one should account for the lack of terrigenous ma-
terial in the above-mentioned belts. Probably, the Northern Rosario and Placetas belts belonged to the deep-water sector of the basin which extended between the Yucatfin and Bahamas passive margins and the speculative spreading zone (ridge?) responsible for the generation of the proto-Caribbean oceanic crust. This deep-water part of the basin was much narrower at the Yucatfin-Florida Straits, creating a fan-like mode of the sediment transport and dispersal. Nevertheless, turbidity currents could not reach more marginally (and upslope?) located sedimentary successions now occurring in the Sierra de los Organos and Southern Rosario belts. In some sections, thin turbidites occur also among the radiolarian cherts in the lower part of the Santa Teresa Formation. The influx of the terrigenous material ceased during the Aptian-Albian(?). At the end of the Early Cretaceous, the siliceous deposition extended across the entire deeper part of the northwestern proto-Caribbean basin. The Santa Teresa Formation appears in all belts related to the Yucat~in passive margin, except the Sierra de los Organos and the Guajaib6n-Sierra Azul belts. This formation occurs also in the Placetas belt, originally located south of the Bahamas platform, represented in central Cuba mainly by the Remedios belt (Fig. 18). These radiolarian cherts of Early Cretaceous to early Cenomanian age are not known in the southeastern Gulf of Mexico. Probably, their accumulation in the northwestern part of the proto-Caribbean basin was a net result of several different reasons (basin margins subsidence since the Tithonian, eustatic sea-
Hauterivian - Barremian
~Gulf of Mexico):~ ~
"~o#/^'~,/'"-,. I
FB
,
"~-~
'
~
/
'
""".
"
I- V-.-2..:.:.'.
:...." ~
,
500 km !
Fig. 18. (A) Lower Cretaceous (Hauterivian-Barremian) paleogeography of the proto-Caribbean and adjacent areas. (B) Enlarged area indicated by rectangle in (A), with suggested provenance and distribution of terrigenous sediments (siliciturbidites). NA - North America, SA = South America, FB -- Florida-Bahamas block, SO = Sierra de los Organos belt, SR = Southern Rosario belt, NR = Northern Rosario belt, LE = La Esperanza belt, P -- Placetas belt, C = Camajuanf belt, R - Remedios belt; 1 - oceanic crust, 2 = radiolarian cherts and shales, 3 -- pelagic limestones, 4 - shallow-water carbonates, 5 = suggested distribution of siliciturbidites, 6 = land area. Black arrow indicates probable provenance of the terrigenous sediments.
THE EXPOSED PASSIVE MARGIN OF NORTH AMERICA IN WESTERN CUBA level rise, paleogeographic and paleoceanographic conditions). During the Cenomanian, the pelagic carbonate sedimentation was restored in the Rosario and Placetas belts. The Turonian deposits occur in the Pinalilla Formation and in the upper part of the Carmita Formation, and were also reported from the Pons Formation of the Sierra de los Organos belt (Hatten, 1957). The Coniacian-Santonian, or even late Turonian-Santonian, deposits are very scarce, or even entirely missing (?) in western Cuba, due to non-deposition and, sometimes, Late Cretaceous and/or Paleocene erosion (Figs. 4 and 6). The Carmita Formation (Cenomanian-Turonian) pelagic limestones are more clayey in their uppermost part and are overlain by the marls and limestones of the Campanian Moreno Formation. The thick-bedded pelagic limestones of the Pinalilla Formation occur immediately below the thin-bedded limestones and shales of the Moreno Formation, and no traces of erosion are discernible along the Pinalilla/Moreno formations boundary. The origin of a regionally extensive late Turonian-Santonian hiatus in the deep-water, pelagic sequence is probably related to paleoceanographic conditions existing during Late Cretaceous times in the northwestern part of the proto-Caribbean basin. For example, this hiatus is coincident with the maximum flooding of the South American continent in northwestern Venezuela during the Late Cretaceous highstand of sea level (Lugo and Mann, 1995). In addition, during the Turonian-Santonian, the Nicaraguan Rise-Greater Antilles Arc partially closed the connection of the proto-Caribbean basin with the Pacific (Pindell and Dewey, 1982; Pindell, 1991). However, the influence of this paleogeographic change on the sedimentation in the proto-Caribbean basin is still to be evaluated.
The beginning of the active margin stage (Campanian-Paleocene) According to Pindell and Dewey (1982), Ross and Scotese (1988) and Pindell (1991), the protoCaribbean ocean basin has been progressively subducted beneath the Caribbean plate during the Late Cretaceous-Eocene. An arc terrane collided with the southern Yucatan margin during the Campanian to early Maastrichtian (Donnelly, 1989; Pindell and Barrett, 1990). The Campanian Sepur clastics filled the foredeep in front of the arc terrane (Sp in Fig. 19). The northeastern passive (proto-Caribbean) margin of Yucatan was also affected by the approaching arc terrane in the Campanian. In the Northern Rosario belt of western Cuba, the Moreno Formation (MR in Fig. 19) contains abundant volcaniclastic
115
Fig. 19. Generalized paleogeographic reconstruction of the protoCaribbean basin for the late Campanian (tectonic framework adapted from Ross and Scotese, 1988 and Pindell et al., 1988). SP = Sepur foredeep, MR = Moreno depocenter, VB = Vfa Blanca deep-water basin (western Cuba), SJM San Juan y Martfnez shallow-water basin (western Cuba); 1 = Caribbean plate oceanic crust, 2 = proto-Caribbean and Atlantic oceanic crust, 3 = pelagic limestones and cherts (Pefias Formation in the Sierra de los Organos belt), 4 = shales, volcaniclastic sandstones (turbidites) and marly limestones of the Moreno Formation, 5 = lack of the Campanian deposits, 6 = shallow-waterlimestones of the Remedios belt in north-central Cuba, 7 -- deep-water clastics of the Vfa Blanca Formation, 8 = shallow-water limestones and conglomerates of the San Juan y Martfnez Formation. Heavy line with triangles denotes subduction of the proto-Caribbean crust beneath the Greater Antilles Arc. =
material, mainly in its upper part (Pszcz6~kowski, 1994a,b). The calciturbidites with volcanic lithoclasts occur in the lower part of this formation. The petrographic composition of the detrital limestones and sandstones clearly indicates the volcanic arc source for the Moreno Formation turbidites. During the Campanian, 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), as proposed by Pindell and Dewey (1982) and Pindell et al. (1988). The position of the GAA shown in Fig. 19 is, in general, in accordance with the tectonic reconstructions by Ross and Scotese (1988) and Pindell and Barrett (1990). The results of the paleomagnetic investigations published so far (Renne et al., 1991; Chauvin et al., 1994; P6rez Lazo et al., 1995; Bazhenov et al., 1996) indicate between 550 and 1600-+-600 km northward displacement of the Zaza volcanic arc (or Zaza terrane) during the Late Cretaceous, Paleocene and Early Eocene. Although the Moreno depocenter could be connected with the Sepur foredeep, their Campanian deposits were accumulated in different tectonic and paleobathymetric settings (Fig. 19). The Sepur For-
116
A. PSzCZOLKOWSKI
Fig. 20. Paleogeographic reconstruction of the proto-Caribbean basin for the latest Maastrichtian; modified tectonic framework adapted from Ross and Scotese (1988, fig. 9): Ca -- Cacarajfcara Formation (western Cuba), Am -- Amaro Formation (north-central Cuba), Pr -- Pefialver Formation (western Cuba), SJM San Juan y Martfnez shallow-water basin (western Cuba), Cf = Cienfuegos basin (south-central Cuba); 1 = Caribbean plate oceanic crust, 2 -- proto-Caribbean and Atlantic oceanic crust, 3 = pelagic limestones and cherts of the Pefias Formation, 4 -- detrital limestones of the Cacarajfcara, Amaro and Pefialver megabeds (maximum thickness of 200-450 m), 5 -- pelagic limestones and cherts of the Camajuanf belt in north-central Cuba, 6 = shallow-water limestones of the Remedios belt in north-central Cuba, 7 = shallow-water limestones of the San Juan y Martfnez Formation in western Cuba and marls of the lowermost part of the Vaquerfa Formation in the Cienfuegos basin. Heavy line with black triangles denotes subduction of the proto-Caribbean crust, while heavy line with open triangle marks underthrusting. =
mation was deposited on a carbonate shelf in southern Yucatan. This shelf was depressed and buried by the Sepur serpentinite-bearing flysch (Pindell and Barrett, 1990). The Moreno marly limestones and siliciturbidites were laid down in a deep-water basin on the thinned continental crust and partly on oceanic crust. During the Campanian, the Sierra de los Organos belt was not affected by the influx of volcaniclastic debris, as this material was not reported from the pelagic limestones and cherts of the Pefias Formation. Campanian sediments were not preserved (or deposited?) across the vast area situated between the Moreno depocenter, Sierra de los Organos belt and the Bahamas platform and slope (Fig. 19). The Maastrichtian deposits of the passive margin of Yucatan are represented by the Cacarajfcara Formation in the Rosario belts, and by the pelagic limestones and cherts of the Pefias Formation in the Sierra de los Organos belt (Figs. 4, 6 and 14). The location of these deposits is shown in paleogeographic reconstruction for the latest Maastrichtian (Fig. 20). South of the Bahamas plat-
form, the limestones and cherts of the Lutgarda Formation are known in the Camajuanf belt, while the Amaro Formation is a characteristic unit occurring in the deep-water Placetas belt (Am in Fig. 20). The Amaro Formation is an equivalent of the Cacarajfcara and Pefialver formations in western Cuba. The peculiar character and origin of the late Maastrichtian Cacarajfcara, Amaro and Pefialver megabeds was studied by Pszczdlkowski (1986b) and was also discussed by Iturralde-Vinent (1992). The latest Maastrichtian paleogeography, shown in Fig. 20, is partly based on conclusions presented in these papers. However, the position of GAA in respect to the Bahamas platform is more southerly in Fig. 20 than that assumed earlier by Pszcz6tkowski (1986b) and Iturralde-Vinent (1992). Such a position of the GAA at the end of the Cretaceous is inferred from the paleotectonic reconstruction and lithology of the late Maastrichtian deposits, as the Cacarajfcara Formation does not contain a significant amount of a coarse-grained volcaniclastic material (> 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.
REFERENCES
Aiello, I.W. and Chiari, M., 1995. Cretaceous radiolarian cherts of western Cuba. Ofioliti, 20:123-128. Anderson, T.H. and Schmidt, V.A., 1983. The evolution of Middle America and the Gulf of Mexico-Caribbean Sea region during Mesozoic time. Geol. Soc. Am. Bull., 94 (8): 941-966. Baumgartner, EO., 1987. Age and genesis of Tethyan Jurassic radiolarites. Eclogae Geol. Helv., 80 (3): 831-879. Bazhenov, M.L., Pszczolkowski, A. and Shipunov, S.V., 1996. Reconnaissance paleomagnetic results from western Cuba. Tectonophysics, 253: 65-81. Benjamini, C., Shagam, R. and Menendez V., A., 1987. (Late?) Paleozoic age for the 'Cretaceous' Tucutunemo Formation, northern Venezuela: stratigraphic and tectonic implications. Geology, 15: 922-926. Bradley, D.C. and Kidd, W.S.E, 1991. Flexural extension of the upper continental crust in collisional foredeeps. Geol. Soc. Am. Bull., 103: 1416-1438. Bralower, T.J., Hutson, E, Mann, E, Iturralde-Vinent, M., Renne, E and Sliter, W.V., 1993. Tectonic of oblique arc-continent collision in western Cuba, 1. Stratigraphic constraints. Eos, 74: 546. Bralower, T.J. and Iturralde-Vinent, M.A., 1997. Micropaleontological dating of the collision between the North American plate and Greater Antilles arc in western Cuba. Palaios, 12: 133-150. Case, J.E., Holcombe, T.L. and Martin, R.G., 1984. Map of geologic provinces in the Caribbean region. In: W.E. Bonini,
A. PSZCZOLKOWSKI R.B. Hargraves, and R. Shagam (Editors), The CaribbeanSouth American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162: 1-30. Case, J.E., MacDonald, W.D., Fox, EJ., 1990. Caribbean crustal provinces; Seismic and gravity evidence. In: J.E. Case and G. Dengo (Editors), Decade of North American Geology, Vol. H. The Caribbean Region. Geological Society of America, Boulder, CO, p. 15-36. Chauvin, A., Bazhenov, M.L. and Beaudouin, T., 1994. A reconnaissance paleomagnetic study of Cretaceous rocks from central Cuba. Geophys. Res. Lett., 21: 1691-1694. Cobiella-Reguera, J.L., 1996. E1 magmatismo junisico (caloviano?-oxfordiano) de Cuba occidental: ambiente de formaci6n e implicaciones regionales. Rev. Asoc. Geol. Argent., 51 (1): 15-28. Coleman, M.L., Gruszczyfiski, M., Pszcz61kowski, A. and Myczyliski, R., 1995. Geochemical isotopic and elemental record from Cuban Late Jurassic-earliest Cretaceous: a protoCaribbean sea-water evolution. 1st SEPM Congr. Sedimentary Geology, August 13-16, 1995, St. Pete Beach, FL, Progr. Abstr. Vol. 1, p. 41. Covey, M., 1986. The evolution of foreland basins to steady state: evidence from the western Taiwan foreland basin. In: EA. Allen and E Homewood (Editors), Foreland Basins. Int. Assoc. Sedimentol., Spec. Publ., 8: 77-90. de la Torre, A., 1960. Fauna de la formaci6n Cayetano del Junisico Medio de Pinar del Rio. Mem. Soc. Cubana Hist. Nat., 25 (1): 65-72. de la Torre, A., 1988. Resultados micropaleontol6gicos m~s importantes del levantamiento geol6gico de la provincia de Pinar del Rfo. In: Contribuci6n a la geologfa de la provincia de Pinar del Rfo. Edit. Cient. T6cn., La Habana, pp. 170-184. Donnelly, T.W., 1989. Geologic history of the Caribbean and Central America. In: J.E. Case and G. Dengo (Editors), Decade of North American Geology, Vol. A. The Geology of North America - - An Overview. Geological Society of America, Boulder, CO, pp. 299-321. Donnelly, T.W., Horne, G.S., Finch, R.C. and L6pez-Ramos, E., 1990. Northern Central America; the Maya and Chortis blocks. In: J.E. Case and G. Dengo (Editors), Decade of North American Geology, Vol. H. The Caribbean Region. Geological Society of America, Boulder, CO, pp. 37-76. Fernandez Carmona, J., 1989. Primer reporte de microf6siles fndices del Kimmeridgense en Cuba. In: Geologfa/Geology '89 (Primer Congreso Cubano de Geologfa, La Habana), Resumenes y programa, p. 64. Fermindez, J., Areces, A. and Dfaz, M.L., 1987. Estratigraffa del ~irea Los Arroyos, provincia de Pinar del Rfo basada en datos de perforaci6n profunda. In: Memorias del III encuentro cientffico-t6cnico de geologfa en Pinar del Rfo. Soc. Cubana de Geologfa, pp. 51-56. Furrazola-Bermfidez, G., 1965. Tres nuevas especies de Tintfnidos del Jur~isico Superior de Cuba. Inst. Cubano Rec. Min., Publ. Esp., 2: 1-29. Gradstein, EM., Agterberg, EE, Ogg, J.G., Hardenbol, J., van Veen, E, Thierry, J. and Huang, Z., 1994. A Mesozoic time scale. J. Geophys. Res., 99:24051-24074. Gordon, M.B., Mann, E, C~iceres, D. and Flores, R., 1997. Cenozoic tectonic history of the North America-Caribbean plate boundary in western Cuba. J. Geophys. Res., 102: 1005510082. Haczewski, G., 1976. Sedimentological reconnaissance of the San Cayetano Formation: an accumulative continental margin in the Jurassic of Western Cuba. Acta Geol. Pol., 26:331-353. Hatten, C.W., 1957. Geologic report on Sierra de los Organos (unpubl.). Fondo Geol. (MINBAS), La Habana. Hatten, C.W., 1967. Principal features of Cuban geology: discussion. Am. Assoc. Pet. Geol. Bull., 51: 780-789.
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 Herrera, N.H., 1961. Contribuci6n a la Estratigraffa de la Provincia de Pinar del Rfo. Rev. Soc. Cub. Ing., 61: 22-44. Hildebrand, A.R., Penfield, G.T., Kring, D.A., Pilkington, M., Camargo, Z.A., Jacobsen, S.B. and Boynton, W.V., 1991. Chicxulub crater: a possible Cretaceous/Tertiary boundary impact crater on the Yucatan Peninsula, Mexico. Geology, 19: 867-871. Hou~a, V., 1969. Neocomian rhyncholites from Cuba. J. Paleontol., 43:119-124. Houga, V., 1974. Informe final sobre los trabajos de campo realizados para el estudio bioestratigrafico y la recolecci6n de los ammonites del Tithoniano y Cretacico Inferior en algunas localidades en la Provincia de Pinar del Rfo, Cuba. Unpubl. Rep., Inst. Geol. Paleontol. MINBAS, La Habana. Howell, D., Jones, D. and Schermer, E., 1985. Tectonostratigraphic terranes of the circum-Pacific region. In: D. Howell (Editor), Tectonostratigraphic Terranes of the Circum-Pacific. Circum-Pacific Council for Energy and Mineral Resources, Houston, Texas, Earth Science Series 1, pp. 3-30. Hutson, F., Mann, E and Renne, E, 1998. 4~ dating of single muscovite grains in Jurassic siliciclastic rocks (San Cayetano Formation): constrains on the paleoposition of western Cuba. Geology, 26: 83-86. Imlay, R.W., 1942. Late Jurassic fossils from Cuba and their economic significance. Geol. Soc. Am. Bull., 53: 1417-1478. Imlay, R.W., 1980. Jurassic paleobiogeography of the conterminous United States in its continental setting. Geol. Surv. Prof. Pap., 1062: 1-134. Ingersoll, R.V., 1988. Tectonics of sedimentary basins. Geol. Soc. Am. Bull., 100: 1704-1719. Iturralde-Vinent, M.A., 1975. Problems in application of modern tectonic hypotheses to Cuba and Caribbean region. Am. Assoc. Pet. Geol. Bull., 59: 838-855. Iturralde-Vinent, M., 1992. A short note on the Cuban late Maastrichtian megaturbidite (an impact-derived deposit?). Earth Planet. Sci. Lett., 109: 225-228. Iturralde-Vinent, M.A., 1994. Cuban geology: a new plate-tectonic synthesis. J. Pet. Geol., 17: 39-70. Iturralde-Vinent, M.A., 1995a. Implicaciones tect6nicas de magmatismo de margen continental pasivo en Cuba. In: A.C. Riccardi and M.E Iglesia Llanos (Editors), Jurasico de Cuba y Am6rica del Sur. Rel. Conf. de Campo, PICG 322 en Cuba, Noviembre 22-29, 1994, Buenos Aires, pp. 14-30. Iturralde-Vinent, M.A., 1995b. Cuencas sedimentarias del Paleoceno-Eoceno de Cuba. Bol. Soc. Venez. Geol., 20: 7580. Iturralde-Vinent, M.A., 1996. Introduction to Cuban geology and tectonics. In: M.A. Iturralde-Vinent (Editor), Ofiolitas y arcos volcanicos de Cuba (Cuban ophiolites and volcanics arcs). IUGS/UNESCO Project 364, Contrib. 1, pp. 3-35. Iturralde-Vinent, M. and Norell, M.A., 1996. Synopsis of Late Jurassic marine reptiles from Cuba. Am. Mus. Novitates, 3164, 17 pp. Judoley, K.M. (= Khudoley, K.M.) and Furrazola-Berm6dez, G., 1968. Estratigraffa y fauna del Jurasico de Cuba. La Habana, 126 pp. Khudoley, K.M. and Meyerhoff, A.A., 1971. Paleogeography and geological history of Greater Antilles. Geol. Soc. Am. Mem., 129: 1-199. Klitgord, K.D., Popenoe, E and Schouten, H., 1984. Florida: a Jurassic transform plate boundary. J. Geophys. Res., 89: 7753-7772. Kutek, J., Pszcz6[kowski, A. and Wierzbowski, A., 1976. The Francisco Formation and an Oxfordian ammonite faunule from the Artemisa Formation. Acta Geol. Polon., 26:299-319. Lewis, J.F. and Draper, G., 1990. Geology and tectonic evolution of the northern Caribbean margin. In: J.E. Case and G. Dengo (Editors), Decade of North American Geology, Vol. H. The
119
Caribbean Region. Geological Society of America, Boulder, CO, pp. 77-140. Lopez Ramos, E., 1975. Geological summary of the Yucatan Peninsula. In: A.E.M. Nairn and EG. Stehli (Editors), The Ocean Basins and Margins, 3. The Gulf of Mexico and the Caribbean, Plenum Press, New York-London, pp. 257-282.. L6pez Rivera, J.G., L6pez Quintero, J.O., Fernandez Carmona J. and Fernandez Rodriguez, G., 1987. Analisis geol6gico del corte del pozo param6trico Pinar 1. In: Memorias del III Encuentro Cientffico-t6cnico de Geologfa en Pinar del Rfo. Soc. Cubana de Geologfa, pp. 40-45. Lugo, J. and Mann, E, 1995. Jurassic-Eocene tectonic evolution of Maracaibo basin, Venezuela. In: A. Tankard, G. Suarez and H. Welsink (Editors), Petroleum Basins of South America. Am. Assoc. Pet. Geol., Mem., 62: 699-725. Mann, E, Draper, G. and Lewis, J.E, 1991. An overview of the geologic and tectonic development of Hispaniola. In: E Mann, G. Draper and J.F. Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 128. Martfnez, D. and Vazquez, M., 1987. Alturas Pizarrosas del Norte, extremo occidental de la subzona 'Sierra del Rosario'. In: Memorias del III Encuentro Cientffico-t6cnico de Geologfa en Pinar del Rfo. Soc. Cubana de Geologfa, pp. 57-71. Marton, G. and Buffler, R.T., 1994. Jurassic reconstruction of the Gulf of Mexico Basin. Int. Geol. Rev., 36: 545-586. Meyerhoff, A.A. and Hatten, C.W., 1974. Bahamas salient of North America: Tectonic framework, stratigraphy, and petroleum potential. Am. Assoc. Pet. Geol. Bull., 58: 12011239. Millan, G. and Myczyfiski, R., 1978. Fauna jurasica y consideraciones sobre la edad de las secuencias metam6rficas del Escambray. Inf. Cient. T6cn. Acad. Cien. Cuba, 80: 1-14. Mossakovskiy, A. and de Albear, de, J.E, 1979. Estructura de cabalgamiento de Cuba occidental y norte e historia de su formaci6n a la luz del estudio de los olistostromas y las molasas. Cien. Tierra Espacio, 1:3-31. Myczyfiski, R., 1976. A new ammonite fauna from the Oxfordian of the Pinar del Rfo Province, western Cuba. Acta Geol. Pol., 26:261-299. Myczyfiski, R., 1977. Lower Cretaceous ammonites from Sierra del Rosario (western Cuba). Acta Paleontol. Pol., 22: 139173. Myczyfiski, R., 1989. Ammonite biostratigraphy of the Tithonian of western Cuba. Ann. Soc. Geol. Pol., 59: 43-125. Myczyfiski, R., 1994a. Caribbean ammonite assemblages from Upper Jurassic-Lower Cretaceous sequences of Cuba. Stud. Geol. Pol., 195:91-108. Myczyfiski, R., 1994b. Lower Tithonian ammonites from the Sierra del Rosario, Western Cuba. Palaeopelagos Spec. Publ. 1, Proc. 3rd Pergola Int. Symp., pp. 287-298. Myczyfiski, R., 1996a. Discusi6n sobre la presencia de los g6neros Parodontoceras Spath, 1923 y Lytohoplites Spath, 1925 en el Tithoniano de Cuba. In: E O16riz and EJ. Rodrfguez-Tovar (Editors), Cephalopods - - Present and Past. IV Int. Symp. Granada, July 15-17, 1996, Abstracts Volume, pp. 125-126. Myczyfiski, R., 1996b. Interpretaci6n estratigrafica y revisi6n sistematica de algunos ammonites del Tithoniano de Cuba. Terra Nostra, 8/96, p. 103. Myczyfiski, R., 1999. Inoceramids and buchids in Tithonian deposits of western Cuba and their paleogeographic significance for the Caribbean region (in prep.). Myczyfiski, R. and Pszcz6~kowski, A., 1976. The ammonites and age of the San Cayetano Formation from the Sierra del Rosario, western Cuba. Acta Geol. Pol., 26:321-330. Myczyfiski, R. and Pszcz6~kowski, A., 1990. Tithonian stratigra-
120 phy in the Sierra de los Organos, western Cuba: correlation of the ammonite and microfossil zones. Atti II Conv. Int. EE.A. Pergola, 1987, pp. 405-415. Myczyriski, R. and Pszcz6lkowski, A., 1994. Tithonian stratigraphy and microfacies in the Sierra del Rosario, Western Cuba. Stud. Geol. Pol., 105: 7-38. Pardo, G., 1975. Geology of Cuba. In: A.E.M. Nairn and EG. Stehli (Editors), The Ocean Basins and Margins, 3. The Gulf of Mexico and the Caribbean. Plenum Press, New York, pp. 553-615. P6rez Lazo, J., Fundora Granda, M., Garcfa, A., Krop~i~ek, V. and Honi6ek, J., 1995. Paleomagnetic investigations in Cuba from late Jurassic to middle Eocene times and tectonic implications. Acta Univ. Carolinae Geol. 38 (1994): 3-19. Pessagno, E.A., Jr., Hull, D.M., Longoria, J.E and Kelldorf, M.E., 1993. Tectonostratigraphic significance of the San Pedro del Gallo area, Durango, western Mexico. In: G. Dunn and K. McDougall (Editors), Mesozoic Paleogeography of the Western United States, II. Pac. Sect. SEPM, Book 71, pp. 141-156. Pessagno, E.A., Jr., Cantd-Chapa, A., Hull, D.M., Kelldorf, M., Longoria, J.E, Martin, C., Meng, X., Montgomery, H., Fucugauchi, J.U. and Ogg, J.G., 1999. Stratigraphic evidence for northwest to southeast tectonic transport of Jurassic terranes in central Mexico and the Caribbean (western Cuba). In: E Mann (Editor), Caribbean Basins - - Sedimentary Basins of the World, vol. 4. Elsevier, Amsterdam, pp. 123-149. Pindell, J.L., 1991. Geologic rationale for hydrocarbon exploration in the Caribbean and adjacent regions. J. Pet. Geol., 14 (3): 237-257. Pindell, J.L. and Barrett, S.E, 1990. Geological evolution of the Caribbean region; A plate-tectonic perspective. In: J.E. Case and G. Dengo (Editors), Decade of North American Geology, Vol. H. The Caribbean Region. Geological Society of America, Boulder, CO, pp. 405-432. Pindell, J.L. and Dewey, J.E, 1982. Permo-Triassic reconstruction of western Pangea and the evolution of the Gulf of Mexico/Caribbean region. Tectonics, 1:179-211. Pindell, J.L., Cande, S.C., Pitman III, W.C., Rowley, D.B., Dewey, J.E, LaBreque, J.L. and Haxby, W., 1988. A platekinematic framework for models of Caribbean evolution. Tectonophysics, 155: 121-138. Piotrowska, K., 1978. Nappe structures of Sierra de los Organos, western Cuba. Acta Geol. Pol., 28: 97-170. Piotrowska, K., 1993. Interrelationship of the terranes in western and central Cuba. Tectonophysics, 220: 273-282. Piotrowski, J., 1977. First manifestations of volcanism in the Cuban geosyncline. Bull. Acad. Pol. Sci., S6r. Sci. Terre, 24: 227-234. Pszczolkowski, A. (-- Pszcz6lkowski, A.), 1994. Interrelationship of the terranes in western and central Cuba - - Comment. Tectonophysics, 234: 339-344. Pszcz61kowski, A., 1971. Jurassic, Cretaceous and Paleogene deposits of Sierra del Rosario (Cuba). Bull. Acad. Pol. Sci., S6r. Sci. Terre, 19 (3/4): 249-259. Pszcz6lkowski, A., 1977. Nappe structure of Sierra del Rosario (Cuba). Bull. Acad. Polon. Sci., S6r. Sci. Terre, 24 (3-4): 205-215. Pszcz6lkowski, A., 1978. Geosynclinal sequences od the Cordillera de Guaniguanico in western Cuba; their lithostratigraphy, facies development, and paleogeography. Acta Geol. Pol., 28 (1): 1-96. Pszcz6tkowski, A., 1981. E1 banco carbonatado junisico de la Sierra de los Organos, Provincia de Pinar del Rio: su desarrollo y situaci6n paleotect6nica. Cien. Tierra Espacio, 3: 3750. Pszcz6{kowski, A., 1982. Cretaceous sediments and paleogeog-
A. P S z C Z O L K O W S K I raphy in the western part of the Cuban miogeosyncline. Acta. Geol. Pol., 32: 135-161. Pszcz6tkowski, A., 1983. Tect6nica del miogeosinclinal cubano en el ~ea limftrofe de las provincias de Matanzas y Villa Clara. Cien. Tierra Espacio, 6: 53-61. Pszcz6~kowski, A., 1985. Sobre la edad del metamorfismo y la estructura tect6nica de la faja Cangre, provincia de Pinar del Rio, Cuba. Cien. Tierra Espacio, 10:31-36. Pszcz6tkowski, A., 1986a. Composici6n del material chistico de las arenitas de la Formaci6n San Cayetano, en la Sierra de los Organos (Provincia de Pinar del Rio, Cuba). Cienc. Tierra Espacio, 11: 67-79. Pszcz6tkowski, A., 1986b. Secuencia estratignifica de Placetas en el ~ea limitrofe de las provincias de Matanzas y Villa Clara (Cuba). Bull. Pol. Acad. Sci., Earth Sci., 34: 67-79. Pszcz61kowski, A., 1986c. Megacapas del Maestrichtiano de Cuba occidental y central. Bull. Pol. Acad. Sci., Earth Sci., 34: 81-87. Pszcz6tkowski, A., 1987. Paleogeography and paleotectonic evolution of Cuba and adjoining areas during the Jurassic-Early Cretaceous. Ann. Soc. Geol. Pol., 57: 127-142. Pszcz6{kowski, A., 1989a. La edad y posici6n de la secuencia vulcan6geno-sedimentaria (Formaci6n E1 S~ibalo), en la estructura geol6gica de la Sierra del Rosario (Cuba occidental). Geologfa/Geology '89 (Primer Congreso Cubano de Geologfa, La Habana), Resumenes y Programa, p. 66. Pszcz6tkowski, A., 1989b. Late Paleozoic fossils from pebbles in the San Cayetano Formation, Sierra del Rosario, Cuba. Ann. Soc. Geol. Pol., 59: 27-40. Pszcz6ikowski, A., 1990. Late Cretaceous volcanic arc in the Bahia Honda terrane, western Cuba. In: Tectonostratigraphic Correlation of Late Cretaceous-Tertiary Island Arc Rocks in the Caribbean Region. Annual Meeting of the Geological Society of America, Dallas, Abstracts with Programs, T 23, p. A338. Pszcz6{kowski, A., 1994a. Lithostratigraphy of Mesozoic and Palaeogene rocks of Sierra del Rosario, western Cuba. Stud. Geol. Pol., 105: 39-66. Pszcz6{kowski, A., 1994b. Geological cross-sections through the Sierra del Rosario thrust belt, western Cuba. Stud. Geol. Pol., 105: 67-99. Pszczd~kowski, A. and de Albear, J.E, 1982. Subzona estructuro-facial de Bahfa Honda, Pinar del Rio; su tect6nica y datos sobre la sedimentaci6n y paleogeograffa del Cret~icico Superior y del Pale6geno. Cien. Tierra Espacio, 5: 3-24. Pszcz6{kowski, A. and de Albear, J.E, 1983. La secuencia vulcanogeno-sedimentaria de la Sierra del Rosario, provincia de Pinar del Rio, Cuba. Cien. Tierra Espacio, 6: 41-52. Pszcz6lkowski, A., Piotrowska, K., Myczyfiski, R., Piotrowski, J., Grodzicki, J., Skupinski, A., Haczewski, G., Danilewski, D., 1975. Texto explicativo para el mapa geol6gico a escala 1:250000 de la provincia de Pinar del Rio. Unpubl. Rep., Inst. Geol. Paleontol. Fondo Geol. (MINBAS), La Habana. Pugaczewska, H., 1978. Jurassic pelecypods from Cuba. Acta Palaeontol. Pol., 23:165-186. Puscharovsky, Yu. (Editor), 1988. Mapa geol6gico de Cuba de la Reptiblica de Cuba escala 1" 250000 (42 sheets). Acad. Sci. Cuba. Pyle, T.E., Meyerhoff, A.A., Fahlquist, D.A., Antoine, J.W., McCrevey, J.A. and Jones, EC., 1973. Metamorphic rocks from north-western Caribbean Sea. Earth Planet. Sci. Lett., 18: 339-344. Renne, ER., Scott, G.R., Doppelhammer, S.H., Linares Cala, E. and Hargraves, R.B., 1991. Discordant Mid-Cretaceous paleomagnetic pole from the Zaza terrane of central Cuba. Geophys. Res. Lett., 18: 455-458. Rigassi-Studer, D., 1963. Sur la g6ologie de la Sierra de los
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
121
model of the Gulf of Mexico and Caribbean region. Tectonophysics, 155: 139-148. Ryabukhin, A.G., Tchekhovich, V.D., Zonenshein, L.P. and Khain, V.E., 1984. Development of the Caribbean basin and the western part of the Tethys. 27 Int. Geol. Congr., Coll. 03, Palaeoceanography, Rept. 3, Moscow, pp. 104-113. Segura Soto, R., Millfin, E. and Fern~indez, J., 1985. Complejos litol6gicos del extremo noroccidental de Cuba y sus implicaciones estratigrfificas de acuerdo con los datos de las perforaciones profundas. Rev. Tecnol., XV, Ser. Geol., 1: 32-36. Viniegra, O.F., 1971. Age and evolution of salt basins of southeastern Mexico. Am. Assoc. Pet. Geol. Bull., 55: 478-494. Wierzbowski, A., 1976. Oxfordian ammonites of the Pinar del Rfo Province (western Cuba); their revision and stratigraphical significance. Acta Geol. Pol., 26:137-260.
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 (<22~ during the latest Tithonian (Late Jurassic) and the Berriasian (Early Cretaceous). Jurassic and Early Cretaceous successions in western Cuba (Sierra del Rosario and Sierra de los Organos, Pifiar del Rfo Province) show lithostratigraphic, paleobathymetric, and paleolatitudinal signatures which are nearly identical to those of San Pedro del Gallo terrane remnants in central Mexico. They clearly represent portions of the North American Plate and are treated as remnants of the San Pedro del Gallo terrane herein. The Cuban remnants of the San Pedro del Gallo terrane were carried to eastern Yucatan by the Walper Megashear. By the Middle Cretaceous terrane amalgamation had occurred between the San Pedro del Gallo and Coahuiltecana terranes and all movement along the Walper Megashear had ceased. Subsequent southwest to northeast movement of the Caribbean Plate during the Late Cretaceous and Early Tertiary bulldozed the Cuban remnants of the San Pedro del Gallo terrane into their present position. Once the Cuban San Pedro del Gallo remnants were carried northward by the advancing Caribbean Plate, it is likely that they became part of an Atlantic-type margin.
Caribbean Basins. Sedimentary Basins of the World, 4 edited by P. Mann (Series Editor: K.J. Hs~i), pp. 123-150. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
124 INTRODUCTION Mexico is the key component in plate tectonic reconstructions involving the break-up of Pangea and the subsequent formation of the Gulf of Mexico and the Caribbean (Figs. 1-3 and 7). The well-known overlap position of Mexico and South America in Atlantic reconstructions necessitates a mechanism for moving much of the Mesozoic succession of Mexico away from its present-day position (see Fig. 7, Inset A). Finding such mechanism has been a major task in geology since the first fit of the continents was suggested more than three decades ago by Carey (1958) and Bullard et al. (1965). Walper and Rowett (1972) suggested that the offset of the Mexico-Marathon-Ouachita-Appalachian structural belt resulted from transform movement along the Texas and Wichita megashears. Almost all plate tectonic reconstructions of the Atlantic, Gulf of Mexico, and Caribbean regions invoke megashears or transcurrent faults as mechanisms to transport crustal blocks in Mexico and to explain the overlap position of Mexico in the reconstruction of Pangea (Fig. 7, Inset A). More than twenty tectonic models have been proposed during the last two decades to explain the paleogeographic reconstruction of Mexico, the Gulf of Mexico, and the Caribbean (e.g., Bullard et al., 1965; Dietz and Holden, 1970; Walper and Rowett, 1972; Van der Voo et al., 1976; Buffter et al., 1981; Dickinson and Coney, 1981; Walper, 1981; Anderson and Schmidt, 1983; Pindell, 1985; Longoria, 1985a,b, 1987, 1994; and so forth). The purpose of this report is to present stratigraphic data that not only demonstrate the presence of the Walper Megashear, but also demonstrate the northwest to southeast translation of remnants of the
E.A. PESSAGNO et al. San Pedro del Gallo terrane along the southwest side of this megashear from higher Boreal 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 ~
\ ~. J
~._
\
~ \
r
~.'~ \,------., "
- 7
.....\ .
"'~43
30 ~ '~'~.
"31
$,),.,,,,
if,, g '/~\
9
,.~.._
.,',,,,,>
~,,
~.
e
Fig. 3. Megashear map of Longoria (1994).
\
\!
"-h
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
<
BOREAL REALM
...............
o~ ~
Weyla
?
~
I
9
"
~
NORTHERN BOREAL PROVINCE
~ > !
< ............................ ............................... BOREAL AMMONITES
.d
~
SOUTHERN BOREAL PROVINCE I
~
Acanthocircus dciranocanthos=~,l
z
I ~
o
I
"' . . . . . .
......
~,~,~~Buchia
........
spp. ,........
,
PANTANELLIIDS~
~ ............
..............................~I .... i
..............
I
...... Parvicingula/Praeparvicingula ....................
Tripocyclia ...... , ~
i i
U3
CALPIoNELLIDS, .... ~>, i
z~. .~. .
GINKGoALES t g Dakota) 'e.-., '~-'---'
i I
-
~
. . . . .
~
.....
~
C R I T I C I S M O F T H E M O D E L O F P E S S A G N O AND
BLOME (1986) BY H A G S T R U M AND M U R C H E Y (1996)
Xiphostylus .: . . . . ~s,
u uJ oO
tudinal position of a given terrane from the study of the radiolarian assemblage alone. Conversely, in successions where megafossils are present and the radiolarians remain unstudied, it is possible to predict the character of the radiolarian assemblage from that of the paleogeographic character of the megafossil assemblage. This was, in fact, the case in our initial studies of the San Pedro del Gallo area. Because there was a mixture of Boreal ammonites such as Amoeboceras sp., cf. alterans (von Busch) (Burckhardt, 1930; Imlay, 1980) and Buchia associated with Tethyan ammonites, the senior author was able to predict that the radiolarian assemblage, if present in these strata, would be assignable to the Southern Boreal Province.
.....
~
.....
Z
. . . . .
:!i:~ ......... ........N .... " i ....... ,:iT:'~;1 ............
Fig. 5. Multiple criteria for use in paleobiogeographic reconstructions in the Northern Hemisphere.
Hagstrum and Murchey (1996) in a report entitled 'Paleomagnetism of Jurassic radiolarian chert above the Coast Range Ophiolite at Stanley Mountain, California and implications for its paleogeographic origins' challenged the validity of the methodology of Pessagno and Blome (1986) and Pessagno et al. (1987, 1993a). In spite of such criticism, we are appreciative of the paleomagnetic data that these authors obtained from the upper part of the volcanopelagic succession at Stanley Mountain (i.e., 32 + 8~ These data support the previous conclusions of Pessagno et al. (1984) and Hopson et al. (1996) that this Coast Range Ophiolite remnant and its overlying sedimentary cover was at mid to high latitudes (Southern Boreal Province) by Tithonian times. As noted in the rebuttal by Hull et al. (1997),
128 Hagstrum and Murchey's interpretation of the paleogeographic model presented by Pessagno and Blome (1986) and Pessagno et al. (1986) is inaccurate. The following points merit discussion herein. (1) The distribution of Praeparvicingula and Parvicingula (= 'homed parvicingulids' of Hagstrum and Murchey) cannot be linked to the environment of deposition based on rock type as suggested by these authors. (2) The distribution of pantanelliids as a criterion is still useful although it must be used with caution. (3) Pessagno and Blome (1986) and Pessagno et al. (1986, 1987, 1989, 1993a,b 1993b)stressed the use of multiple criteria (see P A L E O L A T I T U D I NAL RECONSTRUCTION USING MULTIPLE CRITERIA) rather than just the presence of pantanelliids and Praeparvicingula and Parvicingula as indicated by Hagstrum and Murchey. (4) The Pessagno and Blome model is not dependent on paleomagnetic data from Stanley Mountain as indicated by these authors. These points will be discussed below. For a more in depth discussion the reader should refer to Hull et al. (1997). (1) The distribution of Praeparvicingula and Parvicingula (= 'horned parvicingulids' of Hagstrum and Murchey) cannot be linked to the environment of deposition based on rock type as suggested by these authors. Hagstrum and Murchey (1996, p. 649) follow Baumgartner (1987) in stating that "In general, homed parvicingulids occur in hemipelagic rocks such as tuffaceous mudstone or gray, mudstone, and siltstone". This statement is erroneous. It infers that 'homed parvicingulids' only occur in coastal environments. Parvicingula and Praeparvicingula not only occur in hemipelagic strata, but they also occur in pelagic strata such as red manganiferous ribbon cherts at a number of localities throughout the world, some of which occur in Hagstrum and Murchey's own backyard in the California Coast Ranges. Examples of such occurrences are as follows: (a) thousands of meters of Upper Jurassic oceanic plateau-type basalt interbedded with red manganiferous ribbon chert in the Franciscan Complex at Wilbur Springs and Stoneyford (locs. 17, 18 of Hopson et al., 1981); (b) red ribbon cherts throughout the Greater Antilles in the Caribbean (e.g., volcanic member of ophiolite complex in La D6sirade, ophiolite remnants in Bermeja Complex of southwestern Puerto Rico, and ophiolite remnants in the Duarte Complex of the Dominican Republic (Montgomery et al., 1992, 1994a,b); (c) subduction complex in the Philippines (Yeh and Cheng, 1996). Red ribbon cherts from all of these localities contain Parvicingula and Praeparvicingula and clearly lack any sort of terrestrial input. It is apparent that these taxa flourished in a wide range of sedimen-
E.A. PESSAGNO et al. tary environments ranging from open ocean beyond the reach of terrigenous or volcanic input (red ribbon cherts and also some pelagic limestone such as that occurring in the pillow lava at La D6sirade) to open ocean down-wind from an island arc (e.g., tuffaceous chert above Coast Range Ophiolite at Point Sal, Santa Barbara County, California) to backarc, interarc, and forearc environments (e.g., 'chert member' of La Caja Formation at San Pedro del Gallo (= backarc), Rogue Formation, Klamath Mountains, Southwestern Oregon (= interarc), and Snowshoe Formation, Izee terrane, east-central Oregon (= backarc: see Pessagno and Blome, 1986)). (2) The distribution of pantanelliids as a criterion is still useful although it must be used with caution. As noted by Pessagno et al. (1986, p. 8), the abundance and diversity of pantanelliids in radiolarian chert, siliceous mudstone is controlled by diagenesis. However, it is also influenced by the method of extracting the microfossils from rock samples using the hydrofluoric acid (Blome and Reed, 1993). The fragile nature of many pantanelliid taxa prevents them from being preserved in sedimentary strata which were metamorphosed or underwent lithification subsequent to deposition. Crushing and stretching of specimens extracted from radiolarian chert and shale is common and usually results in the total destruction of all fragile radiolarians. Only radiolarians with thick-walled, sturdy tests are preserved (e.g., Parvicingula, Praeparvicingula, Mirifusus, Archaeocenosphaera) although these forms may often be quite abundant. The best recovery of well-preserved pantanelliids, and indeed all radiolarians occurring in Mesozoic strata, comes from limestone. Radiolarian diversity is at least three times greater in limestone strata than it is in adjacent chert or mudstone layers. This is in part due to two factors: (1) the use of the hydrofluoric acid technique (Pessagno and Newport, 1972) to extract the radiolarians from siliceous strata and the use of HC1 to extract radiolarians from limestone, and (2) the early lithification of limestone strata at the time of deposition as opposed to the post-depositional lithification of the chert. Blome and Reed (1993) demonstrated that the use of the HF technique invariably results in the destruction of radiolarians with more fragile tests (e.g., most pantanelliids). The most dramatic example exemplifying the early lithification of limestone close to the time of deposition comes from a study of Pessagno et al. (1993a) of the volcanopelagic succession overlying the Josephine Ophiolite (Smith River Subterrane, Klamath Mountains, Northwestern California). At this locality (Middle Fork of Smith River) volcanopelagic strata consisting of dark-gray to greenish-gray tuffaceous chert and light-gray pelagic limestone overlie the Josephine Ophiolite and underlie the flysch of the Galice Formation.
NW TO SE TECTONIC TRANSPORT OF JURASSIC TERRANES IN MEXICO AND THE CARIBBEAN Identical radiolarian chert occurs within the volcanic member of the Josephine Ophiolite. In the interval including the volcanic member of the Josephine Ophiolite and the overlying volcanopelagic strata, the limestone strata produced about three times more radiolarian taxa than did the chert strata. Moreover, about four times more pantanelliid taxa occur in limestone strata than in chert strata. Other examples, of this sort are cited in studies by Blome and Reed (1993) and Hull (1995). Although there are undoubtedly cases where pantanelliids can not be used in paleogeographic reconstructions because of the factors noted above, it is important to point out that this group of radiolarians has been successfully utilized to establish paleolatitudes by Pessagno and Blome (1986) for the Izee terrane (east-central Oregon) during the Late Triassic, Early Jurassic and Middle Jurassic, by Pessagno et al. (1984, 1987) for the Huayacocotla remnant of the San Pedro del Gallo terrane during the Late Jurassic and Early Cretaceous, and by Pessagno et al. (1993a) for the Smith River Subterrane during the Middle and Late Jurassic. In all of these examples, radiolarians were extracted from either bedded limestone or from limestone nodules using hydrochloric acid, were exposed to no higher than prehnitepumpellyite to greenschist metamorphism, and were well-preserved. In the Izee terrane, for example, well-preserved, abundant and diversified pantanelliids associated with Tethyan megafossils characterize the Late Triassic (Karnian-Norian) Rail Cabin Formation, the Early Jurassic Nicely Formation (late Pliensbachian), the Early Jurassic Hyde Formation (early to middle Toarcian), and the Early Jurassic part of the Warm Springs Member of the Snowshoe Formation (middle to late Toarcian, part). Moreover, such an assemblage also characterizes the Aalenian, early Bajocian, and late Bajocian (Middle Jurassic) parts of the Snowshoe Formation (Warm Springs Member to lower part of the South Fork Member. However, as noted by Pessagno and Blome (1986) pantanelliid diversity and abundance and diversity drops dramatically in the late Bathonian part of the South Fork Member of the Snowshoe Formation and in early Callovian Lonesome Formation. The great drop in diversity and abundance of pantanelliids can be directly related to the first occurrence of Boreal ammonites such as Kepplerites and Pseudocardoceras in the upper part of the South Fork Member (Pessagno and Blome, 1986; Pessagno et al., 1986). Hull et al. (1997) stated "Thus facing both successes and questions about pantanelliids, we agree that much remains to be solved concerning the paleoceanographic and/or paleolatitudinal preferences of this group of radiolarians. It must be remembered that the recovery of pantanelliids and, indeed,
129
all radiolarians, is greatly influenced by the care taken in sample processing. The recovery of abundant pantanelliids in the Oxfordian and Kimmeridgian parts of other CRO remnants could also be influenced by other factors, such as preferred highfertility upwelling areas. Although Hagstrum and Murchey selectively point to pantanelliids as favoring high-fertility upwelling areas, it could be equally expected that parvicingulids flourished in such areas, subject, however, to different water temperature (paleolatitude) controls. The published literature notably lacks any reference to the comparative preference of pantanelliids over parvicingulids for high fertility upwelling areas. Regarding this link between high fertility and radiolarian assemblages, we also question Hagstrum and Murchey's use (their fig. 8) of Lisitizin's 1972 map, showing spatial distribution of the annual production of silica by marine organisms in the world ocean, for the purpose of advancing their thesis that 'the tuffaceous Stanley Mountain cherts (CRO) were likely deposited at ~30~ within a high productivity zone near the western margin of North America'. This map for the modem ocean illustrates silica production predominantly for diatoms near continental coast lines. Diatoms are chief silica producers in modem oceans and are responsible for more than seventy percent of the total marine silica (Kennett 1982); silica production from radiolarians ranks a distant second among siliceous plankton, and presumably, silicoflagellates and siliceous sponge spicules contribute silica as well. Moreover, it is well known that radiolarians generally are not as abundant in nearshore waters of modern oceans (Kennett 1982) as diatoms in such settings, particularly in eastern boundary regions. We believe that the high production of biogenic silica in the coastal regions bordering North America during the Recent (Hagsrum and Murchey, fig. 8) reflects diatom production and therefore, once again, provides no clue to the geographic distribution of Stanley Mountain radiolarians in the Late Jurassic ocean." (3) Pessagno and Blome (1986) and Pessagno
et al. (1986, 1987, 1989, 1993a)stressed the use of multiple criteria rather than just the presence of pantanelliids and Praeparvicingula and Parvicingula as indicated by Hagstrum and Murchey. As noted by Hull et al. (1997), in their discussion of the paleogeographic foundation for the Pessagno and Blome model Hagstrum and Murchey neglect to mention that the model's realms and provinces, are by definition (Pessagno and Blome, 1986), always constructed on multiple criteria. This thesis was repeatedly stressed subsequently in a series of reports by Pessagno et al. (1986, 1987, 1993a,b) and Blome (1987) (see PALEoLATITUDINAL RECONSTRUCTION USING MULTIPLE CRITERIA).
130 (4) The Pessagno and Blome model is not dependent on paleomagnetic data from Stanley Mountain as indicated by these authors. Hagstrum and Murchey (p. 650) indicate that model presented by Pessagno and Blome (1986) incorporates little quantitative paleolatitudinal control. Moreover, they suggest that the placement of the Central Tethyan-Northem Tethyan boundary at 22~ was based on the Stanley Mountain paleomagnetic data (i.e., 14 -4- 7~ presented by McWilliams and Howell (1982). This statement is totally erroneous. The province boundary was tentatively placed at 22~ because of the presence of Central Tethyan Berriasian faunas occurring at 20~ at DSDP Site 534 (Blake Bahama Basin: Ogg, 1983; Baumgartner, 1984; Pessagno et al., 1987, p. 7). In addition, it is now supported by new data from the Philippines (Yeh and Cheng, 1996). Quantitative data was also cited by Hopson et al. (1996) for the Llanada and Point Sal remnants of the CRO and by Pessagno (1995) for west-central Mexico. There is little question, however, that more paleomagnetic data are needed by workers in future studies. It should be pointed out, however, that 'quantitative' paleomagnetic data for the Phanerozoic is completely dependent on biostratigraphically derived chronostratigraphic data. The Phanerozoic chronostratigraphic scale is based on fossil biozones. Moreover, the geochronologic scale ('geologic time scale') is derived from the integration of biostratigraphic and chronostratigraphic data with geochronometric data (e.g., U/Pb dates).
REMNANTS OF THE SAN PEDRO DEL GALLO TERRANE
The San Pedro del Gallo terrane (SPG) was first defined by Pessagno et al. (1993b) for the area first studied in 1910 by the famous Swiss geologist, Carlos Burckhardt in the state of Durango (Fig. 2: Loc. 12; Fig. 8). Burckhardt's contributions to the geology of San Pedro del Gallo were five-fold: (1) he established the succession/superposition of Jurassic and Cretaceous strata by utilizing ammonite biostratigraphy and chronostratigraphy; (2) he monographed the rich megafossil assemblage (largely ammonites; Burckhardt, 1912); (3) he made the first geologic map as well as the first topographic map of the area; (4) he interpreted the structure and produced numerous structural profiles; and (5) he assessed the mineral deposits of the area. Burckhardt's study was so thorough that few workers have been able to improve on it. Fig. 6 shows a comparison of the geology of the San Pedro del Gallo terrane (SPG) to that of adjacent Parral and Coahuiltecana terranes. The SPG terrane
E.A. PESSAGNO et al. is flanked to the north by the flysch of the Parral terrane (Coney and Campa, 1984) and to the east by the Sierra Madre Oriental terrane (= Coahuiltecano terrane of Sedlock et al., 1993; emend, herein; see Fig. 7). According to Coney and Campa (1984, p. D-3) the 'Parral Terrane' includes "highly deformed, partly calcareous Upper Jurassic and Lower Cretaceous turbiditic sandstone". "The Sierra Madre Oriental Terrane" includes "deformed upper Mesozoic sedimentary rocks of the Gulf of Mexico transgressive sequence and their diverse basement rocks which include, at different places, Pre-Cambrian crystalline rocks and structurally juxtaposed Paleozoic sedimentary rocks, Lower Jurassic sedimentary rocks, structurally associated with Pre-Cambrian and Paleozoic rocks, and pre-Late Jurassic red beds and volcanic rocks." Our data indicate that the San Pedro del Gallo area includes strata that were deposited in the Boreal Realm during the Late Jurassic. The San Pedro del Gallo succession is juxtaposed (along the Walper Megashear) against strata to the east ('Sierra Madre Oriental Terrane') which are both miogeoclinal and Tethyan in character. West-southwest of the Walper Megashear the geology is considerably more complex and is indicative of a back-arc setting. The presence of common basaltic andesite clasts in coeval Upper Jurassic to Lower Cretaceous olistostromal deposits coupled with beds of green and red silty tuff and associated graywacke at the village of Cinco de Mayo to the north of San Pedro del Gallo (Fig. 8) reflects a more proximal back-arc origin. Pessagno et al. (1993b) suggested that the Jurassic successions at Mazapil and in the Huayacocotla segment of the Sierra Madre Oriental (Longoria, 1984) are genetically related to that at San Pedro del Gallo and were possibly remnants of the SPG terrane. Subsequent investigations have proven this to be the case. In this report we regard the Upper Jurassic and Lower Cretaceous successions at San Pedro del Gallo, Symon and Sierra Ramirez, Mazapil (Sierra Santa Rosa), Sierra de la Caja, Sierra Zuloaga and Sierra Sombretillo, Sierra Cadnelaria, Sierra de Catorce, and in the Huayacocotla Anticlinorium to represent remnants of the SPG terrane (see Fig. 2: Locs. 12, 9, 5, 8, 7, 6, and 3; Burckhardt, 1930, 1931; Imlay, 1980). Moreover, it is clear that the Middle Jurassic to Lower Cretaceous succession in western Cuba (Fig. 2: Loc. 24) also includes remnants of the SPG terrane. In addition to displaying the similar lithostratigraphic and paleobathymetric signatures, all of the remnants of the SPG terrane show evidence of tectonic transport from higher latitudes to lower latitudes. The mechanism for this displacement in Central Mexico is the Walper Megashear (Figs. 2, 3 and 7). Tectonostratigraphic data to support this thesis is presented below.
131
N W T O SE T E C T O N I C T R A N S P O R T O F J U R A S S I C T E R R A N E S IN M E X I C O A N D T H E C A R I B B E A N
Coahuilatecana terrane
San Pedro del Gallo terrane
Parral terrane
Characteristics
Massively bedded quartzite and Thin-bedded to massively interbedded nerineid bedded gray sandstone, limestone, organic-rich, black pink silty siliceous shale and shale, and mudstone with mudstone with interbedded numerous limestone nodules, pink to buff, silty micritic black radiolarian chert, limestone, and chalky graywacke, tuff, and thin mudstone, etc. bedded micritic limestone.
Lithologic Characteristics
Flysch: Rhythmically bedded graywacke and shale. "TurbiditJc sandstone" of Coney and Campa (1984). Thickness unknown. Rocks partially coeval with those of San Pedro del Gallo terrane.
Domain
Eugeoclinal: Back Arc
Eugeoclinal: Back Arc
Miogeoclinal
Bathyal to Abyssal
Bathyal to upper abyssal. Above depth of compensation of aragonite
Neritic. Mostly shallow ned'dc
Paleolatitudinal Signature during Jurassic and Cretaceous
Unknown
Southern Boreal to Tethyan
Entirely Tethyan
Geomorphic Expression
Hummocky, irregularly trending ranges
Hummocky, irregularly trending ranges
Parallel, linearly arranged ranges
Water Depth
Structural Style
Folds relatively small scale. Folds relatively small scale. Folding tight, intricate. Folding tight, intricate. Rocks more faulted than those" Rocks more faulted than those of SMO terrane of SMO terrane
Folds open, parallel, more regularly trending, very large scale..
Fig. 6. Comparison of San Pedro del Gallo terrane to Coahuiltecano terrane.
[----~ .... r ...... ~...... . ....... ~ ...... v,..
"~... /.. :,
j,,'
Inset A
[ .... \,,.,..',.. ~ ' , , ~ %
Li 7-.! ,7i~
'.,)~;-,
Inset B
' i '!: ;i,.,,:.i-;,,-7:~,. ~{~.,' ~---~-...../.~-.. .... "
A .Z : : "~. ' "-~'<.,, .~--.~.',.'-;
"hi i>
i
'~..
.t,t"
i~...........7;,.... I
E
,'---.-.~ ....
/ ..........i.;'A../
g
]
~.1
' ~,
~'-
~" ~ ~i
,,~
k
"-~,
"';
~-,
-.,':...It ". '.....!-:::-<"/-. ! /""
.:," "; " . ~;<,i:../ ",
"~I
>
/ "
j
i
!-"/
"-~~~?-<;.">--~---~-.'
<:
"_.,,...... /.."
_ 30 ~ .,.;"
GULF
(
OF Cuba MEXICO
........................................ Tropic of Cancer ............................ Cuban r e m n a n t s
SPG T e r r a n e
~---J'r-~i~~'-"~[
N PACIFIC T
OCEAN
CARIBBEAN +""~//
-,o '":'~
S C A L E
o
leo
t
j
I
I
t:-,,:'~
~......~Y--"
t -
..............'%\
,~
#
(
9
"
""-':'.re, 9
+2
SEA
Guatemala
Fig. 7. New map showing terranes, first-order megashears, Sedlock et al.'s terrane map, and Pangea at end of Pe.
ANALYSES OF SAN PEDRO DEL G A L L O T E R R A N E REMNANTS
Figs. 2 and 7 shows the position of San Pedro del Gallo (SPG) terrane remnants to the west of
the Walper Megashear. Moreover, Fig. 7 shows the terranes utilized by Sedlock et al. (1993), some of which are emended or abandoned herein (see Inset B of Fig. 7). Fig. 9 is a correlation chart showing the correlation of lithostratigraphic units in each
132
E.A. PESSAGNO et al. 104
~
/z
26 ~
:~ceanls~cGoeaOl~~176 l
'~I
"~%
_
Gom~
,~i
U S a n Luis del Cordero
~
I 104
~
0
l
Kilometers
100
Fig. 8. Index map for San Pedro del Gallo.
,,,,i ._.1 Be
ins
Mazapil Sierra Santa San Pedro del Gallo, Rosa, Zacatecas Durango Chapulhuacan Chapulhuacan Limestone Limestone
Taman--Tamazunchale, Western Cuba Megafossil San Luis Potosi Si ,~rra Sierra de del Fosario los Organos Biohorizons Chapulhuacan ~E~ ~ Limestone ~ ~ ~
Sierra Catorce, San Luis Potosi unnamed Ims.
Radiolarian Biozones
Subzone 4or
PimientaFormation .s o z ~-, (./.) ,-, '< = c/3 oz rv" -1223
La Caja Formation (= "Capas de San Pedro", part, Burckhardt, 1910)
La Caja Formation Units E-B
__
crw 13... 13_
,..
.._
, ...
.~_ '~ E o " ~'
,._
Taman Formation "upper thin-bedded member"
o
sensu
--
Pessagno et al. 1987a
Pastor Mum. :!:i:!:!:i:i:i:i:!:!:!:!:i:i:l ,
...
...
N
~
LaCajaFmUnitE+ , upperUnilF .
:
.
-
.
.
~
E
i
T
"
_.
.
1'
~. ii5 "5) i15 ~. 5.1......15 ii/iI
C ) , ~
.//-
//"
.//-
-
.///
//-
.
Sz. 2ot2
i!~i~iI~Siiiiliiiiiiiiiii
u_ ~ ZuloagaLimestone ZuloagaLms. tongue I oX ~,.-- tonguebetweenupper in imddle of unnamed I o & lowerquartziteunits quartzite unit J z g. ........ -.... " " -.... -........ ;:::~ -........ -.-I z , , ~ ; ,r-., = " //
Sz. 2cz 1
!!~!i!i!~ii!ii~ "lowermassiv_.___~e mum."
z ~ ~:~i~~i::!::::ii::iHiatLi.S::ii::iii:: ......."
Zone 3 .
~
. . a. .m. . a. .n. . . .Fm T
~rancisco Ja0ua
- -
~F:
N
TamanFm.
unnamedred shale "::1: - -~ unnamedredshale,siltstone, unnamed_red shale, [::ii!i!i!i::iii!::!iiiiiiiii!ii ~ &limestone(UnitsF&GI siltstone,&limestone ::::::::::::::::::::::::::::::::::::::: or'v"
~
..~,~....~. ~
La Ca'a Fm PastorMum
-
.
~
C
:::::::i:::!:!
~
Subzone 4p
o
.~
,.,
iiiii;ii~!~i~ii!i!ii:iiiiiiiiiSi~tij~,iiiiiil
_o ~
o
C
,-...
!::!::i::iiiiii!iiiii!ii::!!!ii! :::::::::::::::::::::::::::: /i:!:!:i:!~!;;:!~i:i:!:i:i:i:
o~
El Verde Member
""~"'''""
Fn.
""
Zuloaga Limestone
?
Fm
C
t~
Sz. 213 - Sz. 2y
Sz. 26
Santiano Formation
%--
ii!i!iiiiiiiii!i!iiiii!iiii
ii!ii!ilHi~itL~Sil;iiiil
" - -- Zone 11
=o o
.
Zone 1 H
- - ~ i ~
--, = ........ .......... ......... / . ...... E~ c,-- O < S" c_)! ~ "/ z..-" ..... ~s~n,~!e.~,.,..se 9 i~ 9 Z z~, c)l~ -7- i Nt--. r--c-~
az
........ /
...... ........ .
! i !iiii!i!iii!!iii!iiiii!i
:!:?i?i?i?i?i?i?:?:?ii:?:!:
~L__
03
Zone 1G
Zone 1 F
Fig. 9. Correlation chart showing radiolarian biozones and ammonite and Buchia markers, and correlation of litho units in all SPG terrane remnants. SPG remnant with chronostratigraphic and biostratigraphic units. Fig. l0 shows the depth distribution of important Mesozoic fossil groups used in paleobathy-
metric interpretations. Fig. 11 shows the correlation of lithostratigraphic units with paleobathymetry and paleogeographic position; moreover, it attempts to
NW TO SE TECTONIC TRANSPORT OF JURASSIC TERRANES IN MEXICO AND THE CARIBBEAN
133
Fig. 10. Depth distribution of various Mesozoic fossil groups. demonstrate the significance of lithostratigraphic features (e.g., unconformities) that record regional North American tectonostratigraphic events. Figs. 12-17 are detailed stratigraphic summaries of each terrane remnant that supplement Figs. 9 and 11. San Pedro del Gallo remnant
The analysis of the San Pedro del Gallo remnant is based on ongoing investigations since 1990 (Pessagno et al., 1993b; Martin, 1996; Meng, 1997) as well as previous observations by Burckhardt (1910, 1912, 1930) and Imlay (1939). Fig. 2 shows the position of the San Pedro del Gallo terrane to the west of the Walper Megashear (see Fig. 2: Loc. 12). The understanding of the stratigraphy of the San Pedro del Gallo succession was hampered for many years by miscorrelation of several lithostratigraphic units. Web Fig. 5.11 is a space shuttle image showing the western front of the Sierra Madre Oriental (right). The boundary between the Coahuiltecana terrane (emended herein) and displaced terranes to the west occurs immediately west of the mountain front along the Walper Megashear. Examination of the type Zuloaga Limestone in the Sierra Sombreretillo (Zacatecas) has established that Burckhardt's 'nerineid limestone' at San Pedro del Gallo is its lithostratigraphic equivalent. At San Pedro del Gallo the Zuloaga Limestone (Web Figs. 5.2 and 5.3) occurs as a tongue in Burckhardt's (1910) quartzite unit which Imlay incorrectly correlated with the La Gloria Formation (type locality -- Sierra la Gloria, ~50 km east-southeast of Parras (Fig. 2: Loc. 17), Imlay, 1937). The base of the succession at San Pedro del Gallo is not exposed. In that neither the Zuloaga Limestone nor the lower and upper quartzite units possess age diagnostic fossils, their age can only be established 1 Available at: http://www.elsevier.nl/locate/caribas/
as middle Oxfordian or older via the superposition of overlying strata containing middle Oxfordian ammonites (Figs. 9, 11 and 12). Martin's (1996, p. 105) microfacies analysis of the Zuloaga Limestone indicates that it consists of bioturbated micrite ('lime mudstone') and peloidal lime grainstone. The faunal assemblage of the Zuloaga Limestone is characteristic of inner neritic depths. It contains nerineids, other gastropods, rare corals, echinoid fragments, and rare benthonic Foraminiferida (Miliolina and Textulariina) (Martin, 1996). The presence of miliolids in the Zuloaga Limestone indicates that it was deposited at depths no greater than 100 m (Fig. 10). Martin noted that most of the allochems are well-sorted peloids. He noted that many of the peloids have poorly defined rims that may indicate that they were formed by the rolling action of near-shore waves. Moreover, he suggested that the bioturbated micrite was deposited upon a gently sloping carbonate bank in an area that was isolated from the winnowing action of wave energy. The 'upper quartzite unit' consists of thick-bedded, well-sorted quartz arenite with symmetrical ripple marks, trough cross-beds (Martin, 1996, p. 105), and occasional large gastropods. Martin indicated that the uppermost beds of this unit are peloid lime grainstone similar to those occurring within the Zuloaga Limestone tongue. Moreover, according to Martin the well-sorted nature of the quartz sand and micrite peloids suggests that the 'upper quartzite unit' may have been deposited at even shallower inner neritic depths near a shore face where sediments would be subjected to higher wave energy. The 'upper quartzite unit' (Figs. 9, 11 and 12) is overlain by a unit consisting of pink siltstone, mudstone, and silty limestone with middle and upper Oxfordian ammonites, brachiopod shell fragments, Buchia concentrica, common nodosarids (benthonic foraminifers), and miliolids (benthonic foraminifers) (senior author's observation and those of Martin, 1996). Martin noted that the calcareous siltstone is
134
E.A. P E S S A G N O et al.
Fig. 11. Lithostratigraphy, chronostratigraphy, paleobathymetry, and paleobiogeography. (1) "Lower quartzite unit" of Burckhardt (1910). Unfossiliferous massively bedded, white to pink sandstone. (2) "Upper quartzite unit" of Burckhardt (1910). Unfossiliferous massively bedded sandstone. Overlies Burckhardt's "nerineid limestone" -- Zuloaga Limestone of Imlay (1938). (3) Unnamed pink silty limestone, mudstone, and siltstone. Contains Buchia, common ammonites, and Radiolaria (upper part only). Chronostratigraphically significant megafossils include the ammonites Dichotomosphinctes and Discosphinctes and the Buchia concentrica (middle to upper Oxfordian). (4) Lower Kimmeridgian Ataxicoceras Zone and probably part of Idoceras Zone (ammonites) missing. Upper part of upper Oxfordian probably missing. Regional unconformity in much of western North America. Corresponds approximately to onset of deposition of flysch during middle Oxfordian times in Klamath Mountains of northwestern California and southwestern Oregon (Galice Formation) and in Sierra Nevada (Mariposa Formation, Monte del Oro Formation). See Pessagno et al. (1993a). Possibly reflecting pre-Nevadian orogenic pulse in backarc domain. (5) "Lower shale member" of La Caja Formation = lower part of "Capas de San Pedro" of Burckhardt (1910). Dark gray calcareous to siliceous mudstone with micrite nodules containing abundant Radiolaria, common ammonites and Buchia. Basal strata assignable to Idoceras Zone (upper half of lower Kimmeridgian) and radiolarian Subzone 2 alpha-1 (Meng, 1997; Meng and Pessagno, in prep.). (6) Regional unconformity in western North America recognizable in Nevadian back arc terranes (e.g., all San Pedro del Gallo remnants in Mexico and in Cuba) and in Nevadian forearc terranes (e.g., volcanopelagic (VP) strata overlying Stanley Mountain remnant of Coast Range Ophiolite, San Luis Obispo Co., California and Point Sal remnant of Coast Range ophiolite, Santa Barbara Co., California). See Hull (1991, 1995), Hull et al. (1993); Hopson et al., (1996). (7) Includes upper part of "Capas de San Pedro" of Burckhardt (1910) and "chert", "upper shale", and "Cerro Panteon quarry unit 2" members of La Caja Formation herein (see Figs. 8 and 10 for more detailed litho description, biostratigraphic data and chronostratigraphic data). Note that the La Caja Formation at San Pedro del Gallo was miscorrelated lithostratigraphically by Imlay (1939) with his La Casita Formation. All members of La Caja Formation at San Pedro del Gallo with abundant Radiolaria. Sudden influx of silty wacke at Cerro Panteon and at La Pefia (10 km north of San Pedro del Gallo) reflects onset of Nevadian orogeny. Contact of Great Valley Supergroup (flysch) and underlying VP sequence at Stanley Mountain above Stanley Mountain remnant of Coast Range Ophiolite occurs in lower part of Subzone 4 alpha (uppermost upper Tithonian). At San Pedro del Gallo equivalent strata contain the ammonite Durangites and Buchia piochii. (8) Unnamed thin-bedded, tan to pink mudstone and micrite with common ammonites and calcified Radiolaria. These strata overly the massive to medium bedded tan micrites of the Chapulhuacan Formation (type area in Taman-Tamazunchale area to southeast. Imlay (1937) miscorrelated these strata with shallow neritic Taraises Formation (type area = Sierra de Parras). (9) Paleomagnetic data ("upper quartzite unit") from Ogg indicates ~40~ of Jurassic paleoequator (Pessagno, 1995). Presence of the Boreal megafossils Buchia concentrica and Amoeboceras sp. in overlying Oxfordian strata associated with Tethyan ammonites such as Dichotomospinctes indicate Southern Boreal Realm. Overlying La Caja Formation with Buchia concentrica, B. rugosa, B. mosquensis, and B. piochii associated with Southern Boreal radiolarian assemblage characterized by high diversity and abundance of Parvicingula and Praeparvicingula. Upper part of La Caja Formation with Buchia associated with Parvicingula/Praeparvicingula and abundant calpionellids (Tethyan: see Pessagno et al., 1996). Late Tithonian portion of La Caja Formation formed "~ at boundary between Boreal Realm and Tethyan Realm. (10) Abundant calpionellids together with lack of Buchia and presence of only Tethyan ammonites suggests Northern Tethyan Province (Meng, 1997). (11) Units G and F of Fig. 13. (12) Regional unconformity noted in (4) above. (13) Unit E (pt. cf. Fig. 11). The discovery of the lower Kimmeridgian ammonite Idoceras in a limestone nodule 1.5 m below contact between units E and F, indicate that the uppermost part of unit F is lower Kimmeridgian (identification by Dr. Cant4-Chapa, Instituto Politecnico Nacional, Mexico); the middle Oxfordian ammonite, Dichotomosphinctes was recovered 7 m below this horizon (identification by Dr. Cantf-Chapa, Instituto Politecnico Nacional, Mexico). (14) Units E (pt)-B of Fig. 13. Note that silty wacke occurs in upper part of Unit B in lower Subzone 4 alpha. This horizon occurs below final occurrence of Durangites and Substeueroceras (Pessagno et al., in prep.; identification by Dr. Canti-Chapa, Instituto Politecnico
NW TO SE TECTONIC TRANSPORT OF JURASSIC TERRANES IN MEXICO AND THE CARIBBEAN slightly laminated and contains 10 to 35% angular to subangular, well-sorted quartz grains. The upper part of this u n n a m e d Oxfordian unit contains c o m m o n radiolarians. The sudden occurrence of the radiolarians in this part of the succession reflects a rapid shift in paleobathymetry from inner neritic to outer neritic depths during the late Oxfordian (Figs. 10 and 11). Imlay (1939) correlated the informal unit which Burckhardt called the 'Capas de San Pedro' with the La Casita Formation (see Figs. 9, 11 and 12). Although these lithic units are approximately equivalent chronostratigraphically, it is clear from our examination of the La Casita in its type area (Sierra de Parras, Fig. 2: Loc. 17) as well as in the Sierra Jimulco (Fig. 2: Loc. 18) that the 'Capas de San Pedro' are not correlative lithostratigraphically with the La Casita Formation (see Martin, 1996). The La Casita Formation consists of gypsiferous gray to pinkish-gray silty, calcareous to siliceous mudstone, silty micritic limestone, and siltstone deposited at inner neritic depths containing ammonites, brachiopods, bivalves, and a sparse, poorly diversified foraminiferal assemblage (five species, largely Textulariina: senior author's observations). In contrast, Burckhardt's 'Capas de San Pedro' consists of upper abyssal dark-gray calcareous to siliceous mudstone with c o m m o n black, thin-bedded radiolarian chert in its upper part and c o m m o n radiolarian-rich micrite
135
nodules in its lower part. Although the 'Capas de San Pedro' was informally n a m e d by Burckhardt, it is clearly genetically related to the La Caja Formation. We have observed black radiolarian chert in the La Caja Formation thus far at Cation San Matias (Sierra Santa Rosa) near Mazapil, at its type locality in the Sierra de la Caja, in the Sierra de Catorce, and at other localities where the La Caja Formation has been reported (see Imlay, 1980). The La Caja Formation ( = 'Capas de San Pedro' part of Burckhardt, 1910) is divided into this report into four informal m e m b e r s (in ascending order): (1) the 'lower shale m e m b e r ' , (2) the 'chert m e m b e r ' , (3) the 'upper shale m e m b e r ' , and (4) the 'Cerro Panteon quarry unit 2' member. (1) 'Lower shale member'. The 'lower shale m e m b e r ' consists of 52 m (min.) of dark-gray siliceous to calcareous mudstone with c o m m o n dark-gray micrite nodules. Lenticular masses of thin-bedded, dark-gray micrite are present locally. The bedded micrite and micrite nodules contain abundant radiolarians, rare to c o m m o n ammonites, and Buchia (see Figs. 11 and 12). This unit is assignable to late early K i m m e r i d g i a n to the early late K i m m e r i d g i a n (Figs. 9, 11 and 12). It rests unconformably on the middle to early late Oxfordian strata of the u n n a m e d red, siltstone, limestone and shale unit (Figs. 9, 11 and 12).
Nacional, Mexico). (15) Unnamed limonitic mudstone and limestone of Burckhardt (1930 = Burckhardt's unit B). (16) Change from inner neritic to outer neritic occurs in upper part of unit F (Fig. 13). Martin (1996) noted the first occurrence of common Radiolaria in upper unit F. This horizon may also correspond to regional unconformity noted in 12 above. (17) Called La Joya Formation by Imlay (1980). Lower Jurassic strata below this unit contain ammonites and probably are equivalent to the Huayacocotla Group (see Imlay, 1980). (18) E1 Pastor Member of La Caja Formation (Verma and Westermann, 1973). Massively, bedded medium gray micrite with thin-beds of black radiolarian chert and wacke. Wacke often with displaced shallow neritic megafossils. Overlying E1 Verde member consisting of thin-bedded dark gray micrite and black radiolarian chert together with wacke. Graded-bedding and displaced shallow water fossils noted in wacke (19). Incorrectly correlated with shallow neritic La Taires Formation by Verma and Westermann (ibid.). (20) Overlies Huayacocotla Group in Huayacocotla remnant. (21) Canti-Chapa, (1969) recovered the Boreal ammonite Kepplerites in the subsurface of the Huayacocotla remnant from the shallow neritic Palo Blanco Formation. This ammonite is common in terranes in the Sierra Nevada, in the Izee Terrane of east-central Oregon, and in western terranes north to Alaska. It is indicative of the uppermost Bathonian or lower Callovian (Imlay, 1980) In the surface the Cahuasas Formation is overlain by the Tepexic calcarenite (see Fig. 13 for more detailed description). (22) All but uppermost part of Santiago Formation contains an inner neritic molluscan assemblage. Common Radiolaria (including Praeparvicingula) first occur near top of unit (Pessagno et al., 1987a). (23) Taman Formation characterized by abundant Radiolaria, common pectenacids, and rare ammonites. The rarity of ammonites suggests deposition in the upper abyssal depth zone just below the depth of composition of aragonite. See Pessagno et al. (1987a). Co-occurring throughout the Taman Formation in the area south of Taman (e.g., near Huauchinango Puebla) are discontinuous masses of inner neritic calcarenite (San Andres Member, Canti-Chapa, 1969, 1971; Imlay, 1980) that either represent shallow neritic carbonate sedimentation on sea mounts or turbidites. Definition of Taman Formation following that of Pessagno et al. (1987a). (24) The Pimienta Formation differs from the Taman Formation by consisting of light to medium gray, thin-bedded micrite interbedded with black radiolarian chert and occasional layers of green vitric tuff. (25) See annotation 21 above. Taman Formation with rich Northern Tethyan radiolarian assemblage including Parvicingula and Praeparvicingula and abundant diversified pantanelliids associated with Tethyan ammonites and calpionellids (Tithonian). (See Pessagno et al. 1987a). (26) Pimienta Formation and overlying Chapulhuacan Limestone with abundant calpionellids, tethyan ammonites, and lacking Parvicingula/Praeparvicingula. Central Tethyan Province. (27) Data from Imlay (1980), Haczewski (1976), Lewis and Draper (1990), and Myczyfiski and Pszcz61kowski (1994). Imlay indicates that marine bivalves of probable Middle Jurassic age occur in upper part of this unit. (28) San Vincente Member (Myczyfiski, 1994) is anomalous and appears to be an analog of San Andres Member of Taman Formation in Huayacocotla remnant. See Annotation 23 above. (29) Radiolarian-rich strata comprising all of the Artemisa Formation and most of Guasasa Formation (except for San Vincente, see above) were deposited at upper abyssal depths either above or slightly above the CCD of aragonite. (30) The presence of Parvicingula/Praeparvicingula in the radiolarian assemblage associated with Tethyan ammonites and Buchia (Myczyfiski, 1994) indicates that the Cuban SPG remnants were at Southern Boreal paleolatitudes. (31) Northern Tethyan paleolatitudes indicated by same association as in (30), but with common to abundant calpionellids (Myczyfiski and Pszcz61kowski, 1994).
136
E.A. PESSAGNO et al. Lith0 Unit
Chapulhuacan Limestone
Age
Description Medium to massive bedded light gray to tan, very aphanitic micrite with abundant calpionellids, common Radiolaria, and rare ammonites. Some horizons with phosphate nodules.
Berriasian
I
r
oo~
.,.., c'- ..,,.~_ c" =~_ 0
~.-
=
= o .,,_,
"uppershale member"
Red, pink, and pinkish gray limestone, calcareous siltstone (wacke), and calceous mudstone with common ammonites, common belemnites, Buchia, common Radiolaria, and abundant calpionellids. Thickness = 6 to 77 m. Dark gray siliceous mudstone minor thin-bedded, dark gray micrite and dark gray micrite nodules. Rare ammonites. Buchia rugosa & B. mosquensis. Common calpionellids. Thickness = 36 m.
J late Tithonian
DiagnosticFaunal Elements
Paleobathymetry
N. Tethyan Province
Abundant calpionellids (1) Durangites, Substeueroceras belemnites, Buchia pioch# + abundant calpionellids, abundant Radiolaria
Faunal Realm/Province
ct3
(Subzone 40 0 .
late Tithonian
Ourangites + Buchiarugosa, B.
mosquenfsis Radiolaria
Q
+ Subzone 4or
C
E k.. o LL.
Buchia rugosa, B.
mosquensis Dark gray siliceous mudstone interbedded with black, thin-bedded radiolarian chert and minor dark gray micrite. Common ammonites & Buchia + abundant Radiolaria. Common calpionellids Thickness = - 316 m.
late early Tithonian to
late Tithonian
Subzone 413 Radiolaria. See Meng (1997).
J
gray siliceous to calcareous mudstone with common dark gray micrite nodules throughout. Lenticular masses of thin-bedded, dark gray micrite present locally. Micrite nodules and bedded limestone with abundant Radiolaria, rare to common ammonites and Buchia. Rests unconformably on Oxfordian strata below. Thickness = 52 m (minimum).
Dark ~r = , _
o
=
E
unnamed red siltstone, limestone, and shale
Interbedded red sility limestone, silty mudstone, and siltstone containing Buchia and common ammonites. Radiolaria first occurring in upper part. Thickness = 21.2 m.
"upper quartzite unit"
Massivelybedded white to red sandstone. Cross-bedsand symmetricalripple marks. See Martin (1996). Paleoma.clneticdata = 40~ =3.3 m. Massively bedded micritic limestone with nodules of black chert.
"lower quartzite unit"
Massively bedded red and white fine grained sandstone. Base not exposed.
c-.}
middle Oxfordian to late Oxfordian
Discospinctes Buchia concentrica Dichotomosphinctes No
or
older
fossils
Nerinea, bivalves, corals, and sponge spicules See Burckhardt (1910, 1 9 3 0 No
cO
O r'n
Buchia concentrica, B. early Kimmeridgian rogusa, & B. mosquensis to in same bed with Glochicera,, early late gp. fialar sensu Burckhardt Kimmeridigan Idoceras spp. Subzone 2otl Radiolaria
middle Oxfordian
Zuloaga Limestone
O L_
4Kossmatia, Durangites +
e-L._ e---
Outer Neritic
O
Inner Neritic
fossils
Fig. 12. Stratigraphic summary for SPG remnant.
(2) 'Chert member'. This unit includes ~ 3 1 6 m of early to late Tithonian dark-gray siliceous mudstone interbedded with thin-bedded, black radiolarian chert, and minor dark-gray micrite which rest unconformably on the underlying 'shale member' (Web Fig. 5.4). Occasional thin-layers of quartzrich silty wacke often display graded bedding and may possibly represent turbidites. This unit contains common ammonites, abundant radiolarians, abundant siliceous sponge spicules, common calpionellids, and common Buchia (Fig. 1). (3) 'Upper shale member' (= top of Burckhardt's 'Capas de San Pedro'). The 'upper shale member' consists of 36 m of dark-gray siliceous mudstone and minor amounts of thin-bedded dark-gray micrite. The mudstone contains common micrite limestone nodules. Abundant radiolarians and rare ammonites occur in the siliceous mudstone and in the micrite. Rare ammonites, Buchia, and abundant calpionellids occur in the micrite beds and nodules. The late Tithonian strata of the 'upper shale member' rest conformably above the 'chert member' and below 'Cerro Panteon quarry unit 2' (Figs. 9, 11 and 12). (4) 'Cerro Panteon quarry unit 2' member. This member of the La Caja Formation at San Pedro del Gallo consists of 6 to 77 m of red, pink, and pinkish-gray micritic limestone, calcareous siltstone, and calcareous mudstone with common ammonites, bel-
emnites, calpionellids, and radiolarians (reddish color probably result of hydrothermal alteration by Tertiary intrusives) (Web Fig. 5.5). At La Pefia, 10 km to the north of San Pedro del Gallo, the senior author observed 92 m of interbedded black siliceous shale, thin-bedded siltstone (wacke), and thin-bedded dark gray micrite containing belemnites, abundant radiolarians, Buchia, and common ammonites. Abundant calpionellids were reported from this unit by Adatte et al. (1995). Contreras-Montero et al. (1988)recorded abundant ammonites, Buchia piochii as well as abundant radiolarians and belemnites from this locality. Imlay (1939) correlated strata assignable to 'Cerro Panteon quarry unit 2' and the Chapulhuacdn Limestone (Figs. 9, 11 and 12) with the inner neritic Taraises Formation (type area -- Sierra de Parras, Fig. 2: Loc. 17). Where the latter unit has been observed during the course of this study, it consists of rhythmically bedded chalky mudstone and interbedded medium-gray, medium-bedded micritic limestone. Whereas the Taraises Formation contains a poorly diversified foraminiferal assemblage, brachiopods, echinoids, and ammonites, the San Pedro units contain common ammonites and foraminifers as well as abundant radiolarians and calpionellids (see Fig. 10). The Chapulhuacdn Limestone (type area = Chapulhuacdn, Hildago near Taman-Tamazunchale, San
NW TO SE TECTONIC TRANSPORT OF JURASSIC TERRANES IN MEXICO AND THE CARIBBEAN Luis Potosi; Fig. 2: Loc. 3) consists of ~20 m of medium- to massively-bedded light-gray to tan very aphanitic micrite with abundant calpionellids, common radiolarians, and rare ammonites assignable to the Berriasian. Some horizons contain large phosphate nodules (12 cm). The La Caja Formation as well as the Chapulhuac~in Limestone contain abundant radiolarians and siliceous sponge spicules, rare benthonic foraminifers, and common ammonites (see Figs. 9, 11 and 12). Deposition took place at upper abyssal depths somewhat above the ACD (compensation level of aragonite) during early Kimmeridgian to Berriasian times and continued at these depths through the Late Cretaceous (see Burckhardt, 1930). The radiolarian cherts usually contain nearly 50% radiolarian tests. As a result, it is likely that they formed as radiolarian ooze. The Oxfordian to upper Tithonian part of the succession is characterized by containing a mixture of Tethyan and Boreal ammonites (e.g., Amoeboceras, Idoceras, Durangites), common Buchia (e.g., Buchia concentrica, B. mosquensis, B. rugosa) as well as an abundance of Parvicingula and Praeparvicingula and rare pantanelliids among the radiolarians. The megafossil and radiolarian assemblage coupled with preliminary paleomagnetic data indicate that this terrane remnant originated at Southern Boreal paleolatitudes (~40~ Ogg, in Pessagno et al., 1995) during the Oxfordian (see Figs. 4 and 5). The appearance of abundant calpionellids coupled with the presence of Buchia and the presence of common Parvicingula and Praeparvicingula in the 'Cerro Panteon quarry unit 2' member of the La Caja Formation demonstrate that the San Pedro del Gallo remnant had been transported to close to the boundary (~30~ between the Northern Tethyan Province and the Southern Boreal Province by the latest Tithonian (Late Jurassic) (Figs. 4 and 5). The lack of Boreal elements such as Buchia in overlying Early Cretaceous strata may suggest transport of the San Pedro del Gallo remnant to the Northern Tethyan Province (>22 ~ to <30~ by the Berriasian. Based on Ogg's paleomagnetic data and the faunal data cited above, Meng (1997) estimated the rate of movement of the San Pedro del Gallo remnant along the Walper Megashear to be 4.9 cm/yr.
The Mazapil remnant The Mazapil remnant of the SPG terrane was examined at Canyon San Matias in the Sierra Santa Rosa (Fig. 2: Loc. 5). As in the case of the San Pedro del Gallo remnant, the base of the succession at Canyon San Matias is not exposed (Figs. 9, 10 and 14). The oldest unit exposed at this locality is the Zuloaga Limestone (Unit H in Fig. 13).
137
The Zuloaga consists of massive to medium-bedded, medium-gray micritic limestone strata with nodules of black chert (Web Fig. 5.4). Microfacies analysis of the Zuloaga at this locality indicates that the micrite contains encrusting coralline algae, nerineid gastropods, bivalves, foraminifera, and siliceous sponge spicules (Martin, 1996, p. 67). The faunal and floral data suggests that the Zuloaga Limestone at Canyon San Matias was deposited at inner neritic depths (<50 m). on a carbonate bank free of wave energy. The age of the Zuloaga, like the oldest beds at San Pedro del Gallo, can only be established as middle Oxfordian or older via the superposition of overlying strata (Units G and F) containing middle Oxfordian ammonites (Fig. 13). Unnamed units G and F consists of pink, silty mudstone, micritic limestone, and siltstone rich in bivalves and with common ammonites. Common radiolarians are present in the upper part of Unit E All of the Zuloaga Limestone, Unit G, and all but the upper part of Unit F were deposited at inner neritic depths. The sudden appearance of common radiolarians in the upper part of Unit F reflects a rapid change in paleobathymetry from inner neritic depths to outer neritic depths (~200 m) in the late Oxfordian (Web Figs. 5.6 and 5.7). Units E, D, C, and B (lower Kimmeridgian to Berriasian) are included in the La Caja Formation of Imlay (1938, 1939); see Figs. 9, 11 and 13. All La Caja units at Canyon San Matias are characterized by the presence of common to abundant beds of thin-bedded black, radiolarian chert identical to that in the 'chert member' of Burckhardt's (1910) 'Capas de San Pedro' (Web Figs. 5.8 and 5.9). The chert is interbedded with thin- to medium-bedded, dark-gray micritic limestone and dark-gray siliceous to calcareous mudstone commonly containing dark-gray micritic limestone nodules. Unit D, as noted by Burckhardt (1930), is unique in that it is characterized by the presence of beds of phosphate and phosphatic limestone. All La Caja strata are characterized by containing a microfauna with abundant radiolarians, abundant siliceous sponge spicules, and rare benthonic foraminifera and a megafossil assemblage with common to abundant ammonites. Deposition of La Caja strata at this locality during the Late Jurassic and Early Cretaceous (Berriasian) occurred at upper abyssal depths, or perhaps lower bathyal depths, above the ACD (compensation level of aragonite) and continued at these depths until the end of the Cretaceous (see Burckhardt, 1930). Radiolarian chert formed consists of about 50% by volume of radiolarian tests and test fragments; hence, it is likely that it formed as a radiolarian ooze. The phosphate horizon occurring in Unit D is puzzling. Frequently, phosphate-rich sediments occur today along coast lines with narrow continental shelves
E.A. PESSAGNO et al.
138
Litho Unit
Description
Unnamed limestone Unit (1)
Medium bedded buff calcareous mudstone and micritic limestone. Micrite with common limonite nodules. Thickness (fide Burckhardt, 1930+ 50-70 m.
Chapulhuacan Limestone "UNIT A"
Massively bedded, very fine grained micritic limestone weathering to cream or buff color. Calpionellids, Radiolaria. Sparse ammonites. Thickness (fide Burckhardt 1930) = 15 m.
"UNIT B"
Age Valanginian
"UNIT C"
E o ii
"UNIT D"
...._,
Medium-bedded to thick-bedded (0.9--1.2 m) micrite, thin to medium-bedded black chert, and minor siltstone Abundant Radiolaria and ammonites. Thickness = 21.1 m.
Berriasian ?
"UNIT F"
"UNIT G"
"UNIT H" Zuloaga Limestone
13.
o
c--
, --~
late Tithonian
<s
cc~ c--
5 c---
!
Hybonoticeras spp . . . Kossmatia spp. +Subzone 4~ Radiolaria. i
Thin-bedded black chert, dark gray, silicerous mudstone often with interbedded limestone nodules (up to - 4 m in maximum dimension). Red calcareous mudstone in upper part. Some mudstone beds up to 0.6 m in lower part. Abundant Radiolaria in all lithofacies. Abundant ammonites. Thickness = 27.9 m.
early late Kimmeridgian + hiatus early to late Tithonian
Idoceras spp, Glochiceras grp. fialar, Buchia concentrica, Hyboniticeras, and Zone 2o~1, Zone 3, and Zone 4, Subzone4J3 Radiolaria. ! , i
Red silty calcareous mudstone with 1.5 m dark gray micrite nodules in upper part. Common ammonites, bivalves. Common Radiolaria at top. Thickness = 7.8 m
late early Kimmeridgian at top
Idoceras spp. at top of 'UNIT F"
Massively bedded micritic limestone with nodules of black chert. Base not exposed.
O
Kossmatia spp. Boundary between Subzone 413 and 4~ in upper part of Unit B. i
lateTithonian
Red medium-bedded silty limestone and mudstone. Thickness = 6.6 m
,
:>~
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
(<30~ to >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 (<22~ during the Berriasian (Early Cretaceous).
Remnants of the San Pedro del Gallo terrane in western Cuba (Figs. 9, 11, 16) Remnants of the SPG terrane in western Cuba crop out in the Sierra del Rosario and the Sierra de los Organos (Figs. 9, 11, 16 and 17). Our data are derived from field observations by Longoria and by examination of Jurassic rocks from western Cuba by Pessagno in the collections of the U.S. Geological Survey. Moreover, they are derived from data presented by Br6nnimann (1954), Arkell (1956), Meyerhoff (1964), Khudoley and Meyerhoff (1971), Kutek et al. (1976), Imlay (1980), Myczyfiski (1989, 1994), Lewis and Draper (1990), Myczyfiski and Pszcz6~kowski (1976, 1994).
The Sierra del Rosario remnant The succession in the Sierra del Rosario begins with the Middle Jurassic (Bajocian?) to Upper Jurassic (middle Oxfordian) San Cayetano Formation. The San Cayetano Formation includes 1500 to 3000 m of reddish weathering carbonaceous shale, white to grayish quartzose siltstones and sandstones, micarich gray shales, and friable arkoses (see Lewis and Draper, 1990 and Haczewski, 1976). Imlay (1980, p. 39) indicates that its upper part contains marine bivalves of probable Middle Jurassic age. The San Cayetano is the lithic equivalent of the Cahuasas Formation in the Huayacocotla remnant of the SPG terrane in east-central Mexico. Most of the San Cayetano except for its upper 609 m appears to be non-marine. The upper 609 m of the San Cayetano contains bivalves like Ostrea and Vaugonia which are interpreted herein as being inner neritic. The overlying Francisco Formation includes 13 to 25 m of shale, limestone, and some sandstone with middle to early late Oxfordian ammonites such as Discosphinctes and Dichotomosphinctes. These strata are interpreted herein as being neritic. Whether these late Oxfordian strata contain radiolarians as do coeval strata in the Mexican remnants of the San Pedro del Gallo terrane cannot be established at present. The Francisco Formation is overlain conformably by the late Oxfordian to
142
E.A. PESSAGNO et al.
Litho Unit
Description
~o Thin-bedded black micrite and black radiolarian chert. co
Age Berriasian to Valanginian
D i a g n o s t i c Faunal
Paleobathymetry
Elements
Faunal Realm/Province
Leptoceras Protancyloceras Olcostephanus Buchia
I:D r t-'o 1_ r r-
..-.-..
c~ .c: v
(D b'-
c o
03 03
(33 ..Q
E E o u_
Thin-bndded dark gray to black micritic limestone with thin interbeds of shales, siltstones, and sandstones. Abdundant Radiolaria and ammonites. Thickness -- 40 m.
early Tithonian to Berriasian
c~ N
Peeddolissoceraszitteli Virgatosphinctes Corongocerassp. Salinites Durangites Microacanthoceras Paradontoceras Buchia
r'-
r'....., L.. o
Calpionellids (1) Radiolaria
co
Francisco Formation
San Cayetano Formation (part)
Shale, limestone and sandstone with ammonites. Thickness = 13 -- 25 m.
Reddish weathering carbonaceous shale, white to grayish quartzone siltstone and sandstone, mica-rich shale, and friable quartzite. Upper 609 m with marine bivalves (e.g., Ostrea). See Lewis and Draper (1990). Thickness =1500 -- 3000 m.
early to late 0xfordian miildle 0xfordian to Bajocian
Neritic. Proably mostly inner neritic.
Dichotomosphinctes 4-
Discosphinctes
All but upper 609 m consisting of continental red beds. Upper 609 m inner neritic.
None
Fig. 16. Stratigraphic summary for Sierra de Rosario, Cuba.
Litho Unit
..--.. 4-,
D i a g n o s t i c Faunal
Faunal
Description
Age
Tumbitas Member
Light gray, thin-bedded limestone with intercalations of dark gray chert.
Valanginian
Ammonites poorly preserved Abudant calpionellids i +
Dark gray, laminated, medium-bedded with thin intercalations of shale, chert, and numerous dark gray chert nodules. Thickness = - 15 m.
0
Tumbadero Member
Berriasian
Radiolaria
13-
Paleobathymetry
Elements
Realrn~rovince d3
C O3 03
v t--
F.+.-, I--
o 4..., c13
E
Hard, compact dark gary to black micritic, thin to thick-bedded micrititc limestone. Abundant Radiolaria throughout. Abundant calpionellids in upper lower Tithonian and upper Tithonian. Abundant ammonites. This unit includes the Vifiales Limestone of older literature. Thickness = 300-400 m.
o LI_
early to late Tithonian
Mazapilites Hyboniticeras Salinites grossicostatum Durangites Kossmatia Buchia
0 Z
C3..
c.)
S Vicente Member
Jagua Formation
San Cayetano Formation (part)
Dark gray to black mostly massively bedded micrite with some lenses and concretions of chert. Poorly fossiliferous. Divided into 3 members (ascending order):1] Azucar Membv (48-76 m): Gray i to black, thinibedded micritic limestone Ioally sandy & oolitic (Hatten, 1967, p. 782); 2] Jagua Vieja Member (50-60 m): dark gray silty to sandy shale, mudstone, and limestone with many limestone nodules; 3] Pimienta Member (40-60 m): Gray dense, platey limestone. Overlies San Cayetanof Formation unconformably (See Figure 9). Reddish weathering carbonaceous shale, white to grayish quartzone siltstone and sandstone, mica-rich shale, and friable quartzite. Upper 609 m with marine bivalves (e.g., Ostrea). See Lewis and Draper (1990). Thickness =1500 -- 3000 m.
Fig. 17. Stratigraphic summary for S. de los Organos, Cuba.
Poorly dated. Possibly Kimmeridigian.
o ~._ m
None
Inner Neritic (2) o 1:13 c--
early to late Oxfordian
miildle Oxfordian to Bajocian
Dichotomosphinctes 4-
Discosphinctes (1)
None
Mostly inner neritic.
c.4--, 0 CO
All but upper 609 m i consisting of i continental red beds. Upper 609 m inner neritic.
NW TO SE TECTONIC TRANSPORT OF JURASSIC TERRANES IN MEXICO AND THE CARIBBEAN Valanginian Artemisa Formation. The Artemisa Formation as a whole consists of dark-gray, mostly thin to medium-bedded, dense cherty limestone, and tuffaceous shale (Imlay, 1980, p. 39). Pszcz6~kowski (1978) divided the formation into three members; they are in ascending order: (1) the San Vicente Member (upper Oxfordian-Kimmeridgian; (2) the La Zarza Member (Tithonian); and (3) the Sumidero Member (Berriasian to Valanginian). The San Vincente Member consists of massive inner neritic limestone probably formed as a bank deposit (Myczyfiski, 1994). The massive limestone strata are apparently localized in their distribution and may be analogous to those of the San Andres Limestone of the Huayacocotla remnant (SPG) in east-central Mexico (see Canttl-Chapa, 1969). The Zarza Member in the southern part of the Sierra del Rosario includes about 40 m of thin-bedded black to dark-gray micritic limestone with thin interbeds of shale, siltstone, and sandstone (see Myczyfiski and Pszcz6~kowski, 1994). These strata include black ammonite-beating limestones, aptychi, and ammonite shell coquina in the upper part of the member. The Zarza Member grades up into the overlying Sumidero Member. The Sumidero Member includes 'ammonite-free' black, thin-bedded micrite with interbedded radiolarian chert. Examination of limestones from the Tithonian and Berriasian parts of the Artemisa Formation in Mesozoic USGS collections from Cuba indicate that many of these rocks contain abundant radiolarians. This observation was also confirmed by Myczyfiski and Pszcz6~kowski (1994, p. 11, fig. 3). The presence of radiolarians as well as calpionellids in the La Zarza Member indicates that deposition occurred at bathyal to upper abyssal depths above the ACD (compensation level of aragonite). The absence of the ammonites in the Sumidero Member coupled with the presence of abundant radiolarian chert suggests that deposition during the early Berriasian (Early Cretaceous) was at upper abyssal depths below the ACD (compensation level of aragonite). For the most part the Upper Jurassic faunal assemblage of the Artemisa Formation includes Tethyan ammonites and calpionellids. However, species of the Boreal bivalve Buchia are present throughout the Tithonian (Upper Jurassic) to Valanginian (Lower Cretaceous) interval. Praeparvicingula/Parvicingula was observed by Pessagno in micritic limestones from the Mesozoic collections of the USGS formerly housed at the US National Museum. The composite faunal data suggest that the Sierra del Rosario remnant of the SPG terrane remained near the boundary between the Northern Tethyan Province and Southern Boreal Province during the Late Jurassic and Sierra de los Organos remnant (Figs. 9, 11 and 17)
143
As in the case of the Sierra del Rosario remnant the succession begins in the Sierra de los Organos with the deposition of the Middle Jurassic (Bajocian?) to Upper Jurassic San Cayetano Formation (see description above). The San Cayetano Formation in the Sierra de los Organos is overlain by the Jagua Formation (Hatten, 1967; Wierzbowski, 1976). The Jagua Formation is divided locally into three members: (1) a lower Azucar Member; (2) a middle Jagua Vieja Member; and an upper Pimienta Member. The Azucar Member (48 to 76 m)consists of gray to black, thin-bedded micritic limestone that is in some places oolitic and sandy (Hatten, 1967, p. 782). The Jagua Vieja Member (50 to 60 m) consists of dark-gray silty to sandy shale, marl, and limestone and contains many limestone nodules in the shale beds. The upper Pimienta Member (40-60 m) consists of gray, dense platy thin-bedded limestone. Imlay (1980), p. 39) indicates that ammonites from the Azucar Member are of middle or late Oxfordian age. The Jagua Formation appears to be lithic equivalent of the Santiago Formation in the Huayacocotla remnant (SPG). The remainder of the Jurassic succession (Kimmeridgian to upper Tithonian) above the Jacaguas Formation in the Sierra de los Organos is included in the Guasasa Formation (Herrera, 1961; Kutek et al., 1976; Wierzbowski, 1976; Imlay, 1980). The Guasasa Formation in the more recent literature includes the Vifiales Limestone of the older literature. The Jurassic part of the Guasasa includes two members: (1) a lower San Vicente Member (1000 m) consisting of dark-gray to black, mostly massively bedded micrite with some lenses and concretions of chert, and (2) an upper E1 Americano Member (300-400 m) consisting of hard, compact gray to black micritic, highly fossiliferous thinto thick-bedded limestone. The presence of common to abundant radiolarians in the E1 Americano Member in association with ammonites and calpionellids suggests that these strata were deposited at upper abyssal to bathyal depths above the ACD (compensation level of aragonite). The Lower Cretaceous part of the Guasasa Formation includes three members: the Tumbadero Member (Berriasian), the Valanginian Tumbitas Member, and the AlbianAptian Infierno Member (cf. Myczyfiski, 1989). The Tumbadero Member consists of 15 m of dark and dark-gray, laminated medium-bedded limestone with thin intercalations of shale, chert, and numerous dark-gray chert nodules. The Tumbitas Member contains about 50 m of light gray, thin-bedded limestone with intercalations of dark-gray chert. The uppermost member, the Infierno Member, consists of about 50 m of dark-gray micritic limestone interlayered with lighter-colored micritic limestone and dark-gray chert. The occurrence of radiolarians as
144 well as calpionellids in the Tumbadero and Tumbitas members coupled with the absence of ammonites suggests that these strata were deposited below the ACD (compensation level of aragonite) at upper abyssal depths. The presence of Tethyan ammonites together with calpionellids, Parvicingula/Praeparvicingula, and Buchia (Myczyfiski, 1994) in the Upper Jurassic Tithonian part of the Guasasa Formation indicate a paleolatitudinal position close to the boundary between the Northern Tethyan Province and Southern Boreal Province (~30~ It should be noted that Buchia is unknown from the North Atlantic Province except for occurrences in Greenland. Elsewhere it is known from Jurassic and Lower Cretaceous strata from Baja California Sur to Alaska. Hence, its occurrence in the Jurassic of western Cuba is entirely anomalous.
Origin of San Pedro del Gallo terrane (= Guaniguanico terrane) in western Cuba In Chapter 4, Pszcz6tkowski includes what we refer to as remnants of the 'San Pedro del Gallo terrane' in the 'Guaniguanico terrane' of IturraldeVinent (1994, 1996). We treat the Guaniguanico terrane as a junior synonym of the San Pedro del Gallo terrane (Pessagno et al., 1993b) herein. Pszcz6tkowski followed Iturralde-Vinent (1994, 1996) in suggesting that the Guaniguanico terrane (= San Pedro del Gallo terrane) was situated along the eastern margin of the Yucat~in platform. He admits, however, that the original position of the Guaniguanico terrane is difficult to establish with any degree of certainty during the Jurassic and Cretaceous. In this report we advocate a similar origin for the Guaniguanico terrane as that advocated by Iturralde-Vinent (1994, 1996). As can be seen from the examination of Figs. 2, 3 and 7, the Walper Megashear cuts the Yucat~in Peninsula. Hence, there is a strong case for an eastern Yucat~in origin for the Guaniguanico terrane as advocated by IturraldeVinent (1994, 1996). As noted, previously both the San Pedro del Gallo terrane and the Coahuiltecana terrane show similar paleobathymetric records and stratigraphic records by the Middle Cretaceous. We suggest that terrane amalgamation had occurred by the Middle Cretaceous and that all movement along the Walper Megashear had ceased. Subsequent southwest to northeast movement of the Caribbean Plate during the Late Cretaceous and Early Tertiary bulldozed the Cuban remnants of the San Pedro del Gallo terrane into their present position (see Montgomery et al., 1992, 1994a,b). Jurassic and Early Cretaceous successions in western Cuba (Sierra del Rosario and Sierra de los Organos, Pifiar del Rio Province) show lithos-
E.A. PESSAGNO et al. tratigraphic, paleobathymetric, and paleolatitudinal signatures which are nearly identical to those of San Pedro del Gallo terrane remnants in central Mexico (Figs. 9, 11, 16 and 17). Even the presence of the inner neritic 'San Vicente Member' (Guasasa Formation) has an analogue in the San Andres Limestone of the Huayacocotla remnant (Fig. 2: Loc. 3). This record, at least during the Jurassic and Early Cretaceous, is distinctly North American. Unconformities and hiatuses are regional in distribution and can be traced as far north as the California Coast Ranges and Klamath Mountains (see Fig. 11). Moreover, these unconformities reflect tectonic events that affect the Nevadian forearc, interarc, and backarc in western North America. Although a backarc origin is advocated for the pre-latest Tithonian (radiolarian Subzone 4 alpha: Fig. 9), such an origin is related to the Nevadian island arc and not the Antillean island arc. Once the Cuban San Pedro del Gallo remnants were carried northward by the advancing Caribbean Plate, it is likely that they became part of an Atlantic-type margin as suggested by Gordon et al. (1998).
ANALYSIS OF PREVIOUSLY DESCRIBED TERRANES IN MEXICO (SEE FIG. 7: INSET B): VALIDITY OF MAYA, GUACHICHIL, TEPEHUAlqO, AND COAHUILTECANO TERRANES OF SEDLOCK ET AL. (1993)
According to Howell et al. (1985, p. 4; see also Howell, 1995): "A tectonostratigraphic terrane is a fault-bounded package of rocks of regional extent characterized by a geologic history which differs from that of neighboring terranes. Terranes may be characterized internally by a distinctive stratigraphy, but in some cases a metamorphic or tectonic overprint is the most distinctive characteristic. In cases where juxtaposed terranes possess coeval strata, one must demonstrate different and unrelated geologic histories as well as the absence of intermediate lithofacies that might link the two terranes. In general, the basic characteristic of terranes is that the present spatial relations are not compatible with the inferred geologic histories." The Mexican SPG terrane remnants occur within the Tepehuafio, Guachichil, and Maya terranes of Sedlock et al. (1993). It is clear that these terranes, as presently defined, are incompatible with the SPG terrane. The focus of Sedlock et al. (1993) in their definition of these terranes seems to have been mostly on the Paleozoic and Pre-Cambrian basement rocks rather than on important differences in the Mesozoic stratigraphic record. This Mesozoic stratigraphic record is critical in plate-tectonic reconstructions beating on the break-up of Pangea, the opening of the Gulf of Mexico, and the origin of the
NW TO SE TECTONIC TRANSPORT OF JURASSIC TERRANES IN MEXICO AND THE CARIBBEAN Caribbean Plate. The break-up of Pangea heralds a new cycle of terrane formation that crosscuts terrane and plate boundaries that characterized the Pre-Cambrian and Paleozoic. The Tepehuafio and Guachichil terranes are too generalized and all encompassing to be useful in plate reconstructions. Sedlock et al. (1993) failed to relate the geologic history of coeval rock packages (see Longoria, 1994). The Maya terrane and the Coahuiltecano terrane in an emended form may still be usable. The Coahuiltecano terrane is emended herein (see below).
Objections to the Maya terrane Sedlock et al. (1993, p. 28) divide their Maya terrane into three 'geographic' provinces: (1) A northern province. This province includes southern Tamaulipas, Veracruz as far southeast as the Isthmus of Tehuantepec, and thin transitional crust along the western margin of the Gulf of Mexico. (2) The Yucat(m platform. The Mexican states of Tabasco, Campeche, Quintana Roo, and Yucatan as well as Belize, northern Guatemala, and thinned transitional crust in the adjacent Gulf of Mexico. (3) A southern province. This province is said to include central Guatemala, Chiapas, and northeastern Oaxaca. Although all three provinces of the Maya terrane may show remnants of Pre-Cambrian and Paleozoic Gondwana crust that remained behind after the break-up of Pangea and the opening of the Gulf of Mexico, the northern province displays two completely different stratigraphic records during the Jurassic and Early Cretaceous. Most of the Mayan terrane in the state of Veracruz should be assigned to the SPG terrane (Huayacocotla remnant) as described herein. That portion of Mayan terrane said to be in the state of Tamaulipas should be reassigned to the Coahuiltecano terrane (emend. herein); it is clearly a portion of the 'Victoria segment' of the Sierra Madre Oriental as described by Longoria (1985a,b, 1986, 1987, 1994). In east-central Mexico the Huayacocotla remnant of the SPG terrane has been juxtaposed against the Coahuiltecano terrane along the Walper Megashear at approximately the latitude of the Tampico-Ciudad Valles line.
Objections to the Guachichil terrane and Coahuiltecano terrane Sedlock et al. (1993) divided the Guachichil terrane into provisional northern and southern subterranes based on the outcrop of Paleozoic sedimentary and metamorphic rocks. The southern terrane corresponds to the Huayacocotla remnant of the SPG terrane herein and to the Huayacocotla segment of the Sierra Madre Oriental of Longoria (1985a,b) and the
145
Huayacocotla terrane of Longoria (1994). The northern subterrane corresponds to the Victoria segment of the Sierra Madre Oriental of Longoria (1984). We see no significant differences in the Mesozoic (Pangea and post-Pangea) paleobathymetric, paleolatitudinal, and lithostratigraphic signatures of the northern subterrane of Sedlock et al.'s Guachichil terrane and their Coahuiltecano terrane. In this report the northern subterrane is included within the Coahuiltecano terrane and the term Guachichil is abandoned. The Coahuiltecano terrane is emended herein to include the northern subterrane and exclude the southern subterrane of the Guachichil terrane; moreover, it is emended to include all parts of the Tepehuafio terrane east of the Walper Megashear. As so defined, the southern boundary between the Coahuiltecano terrane and the Huayacocotla remnant of the SPG terrane would correspond to the Walper Megashear (Fig. 7).
Objections to Tepehuafio terrane The Tepehuafio terrane as defined by Sedlock et al. (1993) encompasses the San Pedro del Gallo remnant, the Symon remnant, and the Sierra de Catorce remnant of the SPG terrane. Moreover, it includes the Parral and Sierra Madre Oriental Terranes of Campa (1983), Campa and Coney (1983) and Coney and Campa (1984) (= in part Coahuiltecano terrane emended herein). As presently defined it embraces totally unrelated scraps of real estate on both sides of the Walper Megashear. It is obvious that the chief problem with the definition of the Tepehuafio terrane is the placement of the Mojave Sonora Megashear in that the alleged position of this structure defines the Tepehuafio terrane's eastern boundary. As noted under the discussion of the Coahuiltecano terrane above, all of the Tepehuafio terrane east of the Walper Megashear is included in the Coahuiltecano terrane. At present the geology of the area west of the Walper Megashear is still too poorly understood and too complex to warrant establishing terranes as large as the Tepehuafio terrane. Stratigraphic packages are present in different structural blocks each representing different tectonostratigraphic settings: Parral terrane (Upper Jurassic and Lower Cretaceous flysch: backarc? or forearc?), SPG terrane (Upper Jurassic to Upper Cretaceous distal backarc), Zacatecas area (Triassic to Jurassic Interarc). The northern boundary of the San Pedro del Gallo terrane occurs at the village of Cinco de Mayo (Fig. 8) and is tentatively taken to be Longoria's (1994) San Pedro del Gallo Fault (see Fig. 3). At Cinco de Mayo one can observe a complex jumble of structural blocks that include Upper Jurassic-Lower Cretaceous sandstones (graywacke), shales, limestone, green and red tuffaceous siltstone, and olistostromal units with basaltic andesite clasts.
146 COMPARISON OF PALEOBATHYMETRY IN THE SAN PEDRO DEL GALLO TERRANE (SPG) AND THE ADJACENT COAHUILTECANO (COAH) TERRANE (EMEND. HEREIN)
Fig. 18 shows a comparison of the composite paleobathymetry of the SPG terrane and the COAH terrane along opposing sides of the Walper Megashear. The paleobathymetry of the SPG remnants has been discussed above. As can be seen in Fig. 18, the paleobathymetric signature of the COAH is totally different. The COAH paleobathymetric signature can be substantiated by examining the succession exposed at Peregrina Canyon and elsewhere along the eastern front of the Sierra Madre Oriental. Moreover, it can be documented by examining published well records from the Tampico Embayment area presented by Burckhardt (1930), Muir (1936), Imlay (1980), L6pez-Ramos (1985), and numerous other workers. Burckhardt (1930, p. 95) reported inner neritic megafossils (including Ostrea and hydrocorals) in upper Tithonian oolitic limestone from a depth of 986-1029 m in well Chocoy No. 2 about 50 km northwest of Tampico. Inner neritic strata continue upward into the Berriasian and Valanginian (lower part of Tamaulipas Formation). To the southwest of Tampico near Panuco Well Panuco No. 82 includes Upper Jurassic (lower Tithonian) black carbonaceous limestones and shales with Aptychus, pectenacid Aulacomyella, and ammonites
Fig. 18. Comparison of paleobathymetry between Cohuiltecana terrane and SPG terrane.
E.A. PESSAGNO et al. like Mazaplites zitteli Burckhardt. These strata lithologically appear to be similar to those of the Taman Formation and the La Caja Formation. Ostrea, bryozoans, and the remains of conifers occur at two horizons. These may represent inner neritic forms that have been displaced by turbidity currents to bathyal or abyssal depths. By Late Cretaceous times the paleobathymetric record of the Mexican SPG terrane remnants and COAH terrane remnants become similar. Both terranes show a similar lithostratigraphic record during the remainder of the Late Cretaceous. At this point in time (~Albian/Cenomanian) it would appear that terrane amalgamation had occurred and movement along the Walper Megashear had ceased. Anomalies to the scenario described above are inner neritic platform deposits (e.g., E1 Abra Limestone: rudistid reef complex in Sierra del Abra west of Tampico) that formed during the Albian to Turonian interval (Murray, 1961). These strata (rudistid and miliolid limestones) were probably deposited on seamounts at inner neritic depths and relate to a remnant horst and graben topography resulting from previous rifting.
CONCLUSIONS
(1) Stratigraphic data from displaced terranes situated to the west of the Walper Megashear (Mexico) demonstrate similar records of lithostratigraphy, paleobathymetry, and tectonic transport from higher latitudes to lower latitudes. (2) In general, the stratigraphic successions in each of these areas show the same paleobathymetric fingerprint: (a) marine deposition at inner neritic depths during the Callovian to early Oxfordian (Middle to Late Jurassic); (b) marine deposition at outer neritic depths during the late Oxfordian (Late Jurassic); (c) sudden deepening to bathyal or upper abyssal depths (above the ACD of aragonite) from the early Kimmeridgian (Late Jurassic) until the end of the Cretaceous. (3) This paleobathymetric fingerprint differs markedly from that occurring to the east of the Walper Megashear in the Coahuiltecano terrane (emend.) (e.g., Sierra Jimulco, Coahuila; Peregrina Canyon, Tamaulipas). (4) In this report we regard the Upper Jurassic and Lower Cretaceous successions at San Pedro del Gallo, Symon and Sierra Ramirez, Mazapil (Sierra Santa Rosa), Sierra de la Caja, Sierra Zuloaga and Sierra Sombretillo, Sierra Cadnelaria, Sierra de Catorce, and in the Huayacocotla Anticlinorium to represent remnants of a single terrane, the San Pedro del Gallo terrane, that has undergone dismemberment and tectonic transport to varying degrees (NW
NW TO SE TECTONIC TRANSPORT OF JURASSIC TERRANES IN MEXICO AND THE CARIBBEAN to SE) along the west (southwest) side of the Walper Megashear (see Fig. 2: Locs. 3, 5, 6, 7, 8, 9, and 12; Burckhardt, 1930; Imlay, 1980). The San Pedro del Gallo remnant of the San Pedro del Gallo terrane originated at Southern Boreal paleolatitudes (~40~ according to preliminary paleomagnetic data) during the Oxfordian and was tectonically transported to Northern Tethyan paleolatitudes (22 ~ to 29~ by latest Tithonian or earliest Berriasian times. Faunal data (radiolarians and megafossils) from the Mazapil succession (Sierra Santa Rosa) indicates 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 south in the state of San Luis Potosi, the Sierra de Catorce remnant was situated in the Southern Boreal Province during the Oxfordian to Kimmeridgian interval. During the Tithonian the Sierra de Catorce remnant was probably situated in the Northern Tethyan Province. Farther to the east (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 (Middle Jurassic) to Northern Tethyan paleolatitudes (22~ 29~ during the Kimmeridgian and Tithonian (Late Jurassic) to central Tethyan paleolatitudes (<22~ during the Berriasian (Early Cretaceous). The kinetics of terrane remnants along the Walper Megashear can be likened to blocks of ice in an ice flow with most blocks being episodically rotated during transport, some blocks moving along at steady rate (e.g., Huayacocotla remnant and San Pedro del Gallo remnant), and still others moving very little while rotating in place (e.g., Mazapil remnant). (5) In western Cuba the Sierra del Rosario and Sierra de los Organos successions are likewise regarded to be remnants of the SPG terrane and show stratigraphic, paleobathymetric, and paleolatitudinal signatures which are nearly identical to those of San Pedro del Gallo terrane remnants in Mexico (Figs. 9, 11, 16 and 17). Even the presence of the inner neritic 'San Vicente Member' (Guasasa Formation) has an analogue in the San Andres Limestone of the Huayacocotla remnant (Fig. 2: Loc. 3). This record at least during the Jurassic and Early Cretaceous is distinctly North American. Unconformities and hiatuses are regional in distribution and can be traced as far north as the California Coast Ranges and Klamath Mountains (see explanation for Fig. 11). Moreover, these unconformities reflect tectonic events that affect the Nevadian forearc, interarc, and backarc in
147
western North America. Although a backarc origin is advocated for the pre-latest Tithonian (radiolarian Subzone 4 alpha: Fig. 9), such an origin is related to the Nevadian island arc and not the Antillean island arc. In addition, these Cuban remnants are allochthonous when compared to surrounding Central Tethyan successions in the nearby Blake Bahama Basin and elsewhere in Cuba. They contain high latitude bivalves such as species of Buchia that can only be derived (exclusive of Greenland) from a Pacific source. The presence of Southern Boreal/Northern Tethyan faunas (~30~ in the Sierra de los Organos and Sierra del Rosario remnants as late as the Early Cretaceous (Valanginian) suggests much later tectonic transport by northwest to southeast movement along the Walper Megashear and by subsequent southwest to northeast movement as the Caribbean Plate plowed its way through the gap between the North American and South American plates. As suggested by Pszcz6lkowski (see Chapter 4) and by Iturralde-Vinent (1994, 1996), the Guaniguanico terrane (= San Pedro del Gallo terrane) was situated along the eastern margin of the Yucatfin platform. This hypothesis is supported by the fact that the Walper Megashear cuts the Yucatfin Peninsula. We suggest that terrane amalgamation had occurred by the Middle Cretaceous and that all movement along the Walper Megashear had ceased. Subsequent southwest to northeast movement of the Caribbean Plate during the Late Cretaceous and Early Tertiary bulldozed the Cuban remnants of the San Pedro del Gallo terrane into their present position (see Montgomery et al., 1992, 1994a,b). Once the Cuban San Pedro del Gallo remnants were carried northward by the advancing Caribbean Plate, it is likely that they became part of an Atlantic-type margin as suggested by Gordon et al. (1998).
ACKNOWLEDGEMENTS
This investigation was in part supported by grants from the National Science Foundation: EAR-9418194 to Pessagno and Montgomery and EAR-9304459 to Pessagno, Hull, and Ogg. It has also been supported by a grant from CONOCYT to Urrutia Fucugauchi. We also wish to thank Dr. Jurgen Remane (Institut de G6ologie, Universit6 de Neuch~tel, Switzerland) for identifying Late Jurassic and Early Cretaceous calpionellids. Finally, we would like to thank Dr. Norman Silberling (U.S.G.S., Denver, Colorado), Dr. Charles D. Blome (U.S.G.S., Denver, Colorado), and Dr. Gyorgy L. Marton, Amoco Exploration and Production Company, Houston, Texas for their review of the manuscript. Contribution No. 826, Programs in Geosciences, The University of Texas at Dallas.
148 REFERENCES
Adatte, T., Stinnesbeck, W. and Remane, J., 1995. The JurassicCretaceous boundary in Northeastern Mexico. Confrontation and correlations by microfacies, clay minerals, mineralogy, calpionellids, and ammonites. Geobios, 17: 37-56. Aita, Y. and Grant-Mackie, J.A., 1992. Late Jurassic Radiolaria from Kowhai Point Siltstone, Murihku terrane, North Island, New Zealand. In: K. Ishizaki and T. Saito, (Editors), Centenary of Japanese Micropaleontology. Terra Scientific Publishing Co., Tokyo, pp. 375-382. Anderson, T.H. and Schmidt, V.A., 1983. The evolution of Middle America and the Gulf of Mexico-Caribbean sea region during Mesozoic time. Geol. Soc. Am. Bull., 94: 941-966. Arkell, W.J., 1956. Jurassic Geology of the World. Oliver and Boyd, London, 806 pp. Baumgartner, P.O., 1984. A Middle Jurassic-Early Cretaceous radiolarian zonation based on Unitary Associations and age of Tethyan radiolarites. Eclogae Geol. Helv. 77: 729-837. Baumgartner, P.O., 1987. Age and genesis of tethyan Jurassic radiolarites. Eclogae Geol. Helv., 80:831-879. Blome, C.D., 1987. Paleogeographic significance of lower Mesozoic radiolarians from the Brooks Range, Alaska. In: L. Taileur, and P. Weimer (Editors). Alaska North Slope Geology. Society of Economic Paleontologists and Mineralogists, Pacific Section. 1: pp. 371-380. Blome, C.D. and Reed, K.M., 1993. Acid processing in pre-Tertiary radiolarian cherts and its impact on faunal content and biozonal correlation. Geology, 21: 177-180. BrOnnimann, P., 1954. On the occurrence of calpionellids in Cuba. Ecologae Geol. Helv., 46 (2): 263-268. Buffler, R.T., Watkins, J.S., Shaub, EJ. and Worzel, J.L., 1981. Structure and early geologic history of the deep central Gulf of Mexico Basin. In: R.H. Pilger, Jr. (Editor), The Origin of the Gulf of Mexico and Early Opening of the Central North Atlantic Ocean. Proc. Symp. Louisiana State University, Baton Rouge, March 1980, pp. 41-51. Bullard, E.C., Everett, J.E. and Smith, A.G., 1965. The fit of the continents around the Atlantic: a symposium on continental drift. Philos. Trans. R. Soc. London, A, 258: 3-16. Burckhardt, C., 1910. Estudio geologico de la regi6n de San Pedro del Gallo (Durango). Instituto Geol. de M6xico, Parergones, III, 6, pp. 307-357. Burckhardt, C., 1912. Faunes jurassiques et cr6taciques de San Pedro del Gallo. Inst. Geol. M6x., 29:264 pp. Burckhardt, C., 1927. Cefal6podos del Jur~isico medio de Oaxaca y Guerrero. Inst. Geol. M6x. Bol., 47:108 pp. Burckhardt, C., 1930. l~tude synth6tique sur le M6sozoique M6xicain. Soc. Pal6ontol. Suisse M6m., 49: 1-123. Burckhardt, C., 1931. l~tude synth6tique sur le M6sozoique M6xicain. Soc. Pal6ontol. Suisse M6m., 50: 123-280. Campa, M.E, 1983. The Mexican thrust belt. In: D.G. Howell (Editor), Tectonostratigraphic Terranes of the Circum Pacific Region. Circum-Pacific Council for Energy and Mineral Resources, Earth Science Series, 1, pp. 299-313. Campa, M.F. and Coney, P.J., 1983. Tectonostratigraphic terranes and mineral resource distributions in Mexico. Can. J. Earth Sci., 20:1040-1051. Cantti-Chapa, A., 1969. Estratigraffa del Junisico medio-superior del subsuelo de Poza Rica, Ver. Inst. Mex. Pet. Rev., 1 (1): 3-9. Cantti-Chapa, A., 1971. La serie Huasteca (Jur~isico medio-superior) del centro este de Mexico. Inst. Mex. Pet. Rev., 3 (2): 17-40. Cantti-Chapa, A., 1989. Precisiones sobre el limite JurasicoCretacico en el subsuelo del este de M6xico. Rev. Mex. Paleontol., 2 (1): 26-69. Carey, S.W., 1958. Symposium on the Present Status of the
E.A. P E S S A G N O et al. Continental Drift Hypothesis. Symposia of the Geology Department, University of Tasmania, 2, 375 pp. Carrillo-Bravo, J., 1961. Geologia del Anticlinorio HuizachalPeregrina al N.W. de Ciudad Victoria, Tamaualipas. Asoc. Mex. G6ol. Pet. Bol., 13" 1-98. Coney, P.J. and Campa, M.E, 1984. Lithotectonic terrane map of Mexico. In: N.J. Silberling and D.L. Jones (Editors), Lithotectonic Terrane Maps of the North American Cordillera. U.S.Geol. Surv., OF84-0506, pp. D 1-D 14. Contreras-Montero, B., Martinez-Cortez, A. and Gomez-Luna, M., 1988. Bioestratigrafia y sedimentologia del Jurasico Superior en San Pedro del Gallo, Durango, Mexico. Rev. Inst. Mex. Pet., 20 (3): 5-49. Davila-Alcocer, V., 1986. Biostratigraphic Studies of Jurassic, Cretaceous Radiolaria of the Eugenia and Asuncion Formations, Vizcaino Peninsula, Baja California Sur. M.S. Thesis, University of Texas at Dallas, 157 pp. Dickinson, W.R. and Coney, P.J., 1981. Plate tectonic constraints on the origin of the Gulf of Mexico. In: R.H. Pilger, Jr. (Editor), The Origin of the Gulf of Mexico and Early Opening of the Central North Atlantic Ocean. Proc. Symp. Louisiana State University, Baton Rouge, March 1980, pp. 27-36. Dietz, R.S. and Holden, R.C., 1970. Reconstruction of Pangea: Breakup and dispersion of continents, Permian to present. J. Geophys. Res., 75: 4937-4956. Douglas, R.G., 1969. Upper Cretaceous planktonic foraminifera in Northern California. Micropaleontology, 15 (2): 151-209. Erben, H.K., 1956. E1 Junisico Inferior de M6xico y sus amonitas, M6xico D.F. Int. XX Congr. Geol. Internacional, 393 pp. Gordon, M.B., Mann, P., C~iceres, D., Flores, R., 1998. Cenozoic tectonic history of the North American-Caribbean plate boundary zone in western Cuba. J. Geophys. Res. (in press). Haczewski, G., 1976. Sedimentological reconnaissance of the San Cayetano Formation; an accumulative continental margin in the Jurassic of Western Cuba. Acta Geol. Pol., 26:331-353. Hagstrum, J.T. and Murchey, B.L., 1996. Paleomagnetism of Jurassic radiolarian chert above the Coast Range ophiolite at Stanley Mountan, California, and implications for its paleogeographic origins. Geol. Soc. Am. Bull. 108: 643-652. Hatten, C.W., 1967. Principal features of Cuban geology. Am. Assoc. Pet. Geol. Bull., 51: 780-789. Herrera, N.M., 1961. Contribuci6n a la estratigraffa de la provincia de Pinar del Rfo. Soci. Cubana Ing., 61 (1-2): 1-24. Hopson, C.A., Mattinson, J.M. and Pessagno, E.A., Jr., 1981. Coast Range ophiolite, western California. In: W.G. Ernst (Editor), The Geotectonic Development of California. Prentice-Hall, Englewood Cliffs, NJ. Rubey 1: pp. 419-510. Hopson, C.A., Pessagno, E.A., Jr., Mattinson, J.M., Luyendyk, B.P., Beebe, W., Hull, D.M., Mufioz, I.M. and Blome, C.D., 1996. Coast Range ophiolite as paleoequatorial midocean lithosphere. In: W.R. Dickinson, C.A. Hopson and J.B. Saleeby (Editors), Alternate Origins of the Coast Range Ophiolite (California): Introduction and Implications. GSA Today, 6 (2): 1-10. Howell, D.G., 1995. Principles of Terrane Analysis: New Applications for Global Tectonics. Topics in the Earth Science 8, Chapman and Hall, London, 2nd ed., 245 pp. Howell, D.G., Jones, D.L. and Schermer, E.R., 1985. Tectonostratigraphic terranes of the Circum-Pacific region. In: D.G. Howell (Editor), Tectonostratigraphic Terranes of the Circum-Pacific Region. Circum-Pacific Council for Energy and Mineral Resources, Earth Science Series, 1, pp. 1-581. Hull, D.M., 1991. Upper Jurassic Radiolarian Biostratigraphy of the Lower Member of the Taman Formation, East-Central Mexico and of Volcanopelagic Strata Overlying the Coast Range Ophiolite, Stanley Mountain, Southern California Coast Ranges. Ph.D. dissertation, The University of Texas at Dallas, 696 pp.
N W TO SE T E C T O N I C T R A N S P O R T OF JURASSIC T E R R A N E S IN M E X I C O AND THE C A R I B B E A N Hull, D.M., 1995. Morphologic diversity and paleogeographic significance of the Family Parvicingulidae (Radiolaria). Micropaleontology, 41 (1): 1-48. Hull, D.M., Pessagno, E.A., Jr., Blome, C.D., Hopson, C.A. and Mufioz, I.M., 1993. Chronostratigraphic assignment of volcanopelagic strata above the Coast Range ophiolite. In: G. Dunn and K. McDougall (Editors), Mesozoic Paleogeography of the Western United States, II. Pacific Section SEPM, Book 71, pp. 157-170. Hull, D.M., Blome, C.D. and Pessagno, E.A., Jr., 1997, Paleomagnetism of Jurassic radiolarian chert above the Coast Range ophiolite at Stanley Mountain, California, and implications for its paleogeographic origins: Discussion and reply. Discussion. Geol. Soc. Am. Bull., 109 (12): 1633-1639. Imlay, R.W., 1937. Geology of the middle part of the Sierra de Parras, Coahuila, Mexico. Geol. Soc. Am. Bull., 48: 587-630. Imlay, R.W., 1938. Studies in the Mexican Geosyncline. Geol. Soc. Am. Bull., 49:1651-1694. Imlay, R.W., 1939. Upper Jurassic ammonites from Mexico. Geol. Soc. Am. Bull., 50: 1-78. Imlay, R.W., 1980. Jurassic paleobiogeography of the conterminous United States in its continental setting. U.S. Geol. Surv. Prof. Pap., 1062: 1-134. Iturralde-Vinent, M.A., 1994. Cuban Geology, A new plate-tectonic synthesis. J. Pet. Geol., 17: 39-70. Iturralde-Vinent, M.A., 1996. Ofiolitas y arcos volcanicos de Cuba: Project 364, Caribbean ophiolites and volcanic arcs. Special Contribution No. 1, Miami, FL: pp. 1-254. Jones, D.L., Silberling, N.J. and Hillhouse, J., 1977. Wrangellia - - a displaced terrane in northwestern North America. Can. J. Earth Sci., 14 (11): 2565-2577. Kiessling, W., 1995. Pal0kologische Verwertbarkeit Ober-Jurassisch Unterkretazischer Radiolarienfaunen mit Beisp. Ph.D. dissertation, Erlangen University, Erlangen. Kiessling, W. and Scasso, R., 1996. Ecological perspectives of Late Jurassic radiolarian faunas from the Antarctic Peninsula. In: A.C. Riccardi (Editor), Advances in Jurassic Research. GeoRes. Forum, 1-2:317-326. Khudoley, K.M. and Meyerhoff, A.A., 1971. Paleogeography and geological history of the Greater Antilles. Geol. Soc. Am. Mem., 129:200 pp. Kutek, J., Pszcz6tkowski, A. and Wierzbowski, A., 1976. The Francisco Formation and an Oxfordian ammonite faunule from the Artemisa Formation, Sierra del Rosario, Western Cuba. Acta Geol. Pol., 26 (2): 299-319. Lewis, J.E and Draper, G., 1990. Geology and tectonic evolution of the northern Caribbean margin. In: G. Dengo and J.E. Case (Editor), The Geology of North America, H. The Caribbean Region. Geological Society of America, Boulder, Colo., pp. 77-140. Longoria, J.F., 1984. Mesozoic tectostratigraphic domains in east-central Mexico. In: G.E.G. Westermann (Editor), Jurassic-Cretaceous Biochronology and Paleogeography of North America. Geol. Assoc. Can. Spec. Pap., 27: 65-76. Longoria, J.E, 1985a. Tectonic transpression in the Sierra Madre Oriental, northeastern Mexico: an alternative model. Geology, 13 (7): 453-456. Longoria, J.E, 1985b. Tectonic transpression in northeast Mexico: its relation to sea floor spreading in the Gulf of Mexico. Gulf Coast Assoc. Geol. Soc., Trans., 35: 199. Longoria, J.E, 1986, Tectonic transpression in the Sierra Madre Oriental, northeast Mexico: an alternative model. Reply. Geology 14 (9): 809-810. Longoria, J.E, 1987. Oblique subduction and kinematics of the American Plate: evidence from the stratigraphic record of Mexico. In: T.W.C. Hilde and R.L. Carlson (Conveners), Geodynamics Symposium. Texas A and M University, College Station, 2 pp.
149
Longoria, J.E, 1994. Recognition and characteristics of a strikeslip fault system in Mexico and its Mesozoic transpressional regime: implications in plate tectonics and paleogeographic reconstruction. Bol. Dept. Geol. Uni-Son., 11 (1): 77-104. L6pez-Ramos, E., 1985. Geologfa de M6xico, Tomo II. Impresiones Resendiz, Mexico, D.E, 453 pp. Martin, C.B., 1996. Tectonostratigraphic analysis of the Upper Jurassic through Lower Cretaceous stratigraphy in the Mazapil, Sierra de Jimulco, and San Pedro del Gallo regions of North-central Mexico. Thesis (M.S.), Geosciences Department, The University of Texas at Dallas: pp. 1-148. McWilliams, M.O. and Howell, D.G., 1982. Exotic terranes of western California. Nature, 297:215-217. Meng, X., 1997, Radiolarian biostratigraphy of the Upper Jurassic of San Pedro del Gallo terrane, north-central Mexico and the Lower Cretaceous of Nooksack Group, Nooksack Terrane, northwestern Washington. Ph.D. Dissertation, The University of Texas at Dallas, 351 pp. Meyerhoff, A.A., 1964. Review of the geological formations of Cuba (Bermudez, 1961). Int. Geol. Rev., 6 (1): 149-156. Montgomery, H.A., Pessagno, E.A., Jr. and Mufioz, I., 1992. Jurassic (Tithonian) Radiolaria from La D6sirade (Lesser Antilles): preliminary paleontological and tectonic implications, Tectonics, 11 (6): 1426-1432. Montgomery, H.A., Pessagno, E.A., Jr. and Pindell, J.L., 1994a. A 195 Ma terrane in a 165 Ma sea: Pacific origin of Caribbean Plate. GSA Today, 4 (1): 3-6. Montgomery, H., Pessagno, E.A., Jr., Lewis, J.A. and Schellekens, J.H., 1994b. Paleogeography of the Jurassic fragments in the Caribbean, Tectonics, 13: 725-732. Muir, J.M., 1936. Geology of the Tampico Embayment area. Am. Assoc. Pet. Geol., 280 pp. Murray, G.E., 1961. Geology of the Atlantic and Gulf Coastal Province of North America. Harper, New York, pp. 1-692. Myczyfiski, R., 1989. Ammonite biostratigraphy of the Tithonian of western Cuba. Ann. Soc. Geol. Pol., 59: 43-125. Myczyfiski, R., 1994. Caribbean ammonite assemblages from Upper Jurassic-Lower Cretaceous sequences of Cuba. Stud. Geol. Pol., 105:91-108. Myczyfiski, R. and Pszcz6~kowski, A., 1976. The ammonites and age of the San Cayetano Formation from the Sierra del Rosario, western Cuba. Acta Geol. Pol., 26 (2): 321-329. Myczyfiski, R. and Pszcz6lkowski, A., 1994. Tithonian stratigraphy and microfacies in the Sierra del Rosario, western Cuba. Stud. Geol. Pol., 105: 7-38. Nestell, M.K., 1979. Lower Permian fusulinids from the vicinity of Tianguistengo, Hidalgo, Mexico. Texas Academy of Sciences Meeting, Arlington, Texas, Prog. Abstr., p. 40. Ogg, J.G., 1983. Magnetostratigraphy of the Upper Jurassic and lowest Cretaceous sediments, Deep Sea Drilling Site 534, Western North Atlantic. In: S. Orlofsky, Initial reports of the Deep Sea Drilling Project, Washington, D.C., U.S. Government Printing Office. 76: pp. 685-697. Pessagno, E.A., Jr., 1995. Stratigraphic evidence for Northwest to Southeast movement along the west side of the Walper Lineament. Geol. Soc. Am., Abstr. Prog., 27 (6): A75. Pessagno, E.A., Jr. and Blome, C.D., 1986. Faunal affinities and tectonogenesis of Mesozoic rocks in the Blue Mountains Province of eastern Oregon and western Idaho. In: T.L. Vallier and H.C. Brooks (Editors), Geology of the Blue Mountains Region of Oregon, Idaho, and Washington: Geologic Implications of Paleozoic and Mesozoic Paleontology and Biostratigraphy. U.S. Geol. Surv. Prof. Pap., 1435: 65-78. Pessagno, E.A., Jr. and Hull, D.M., 1999. Upper Jurassic (middle Oxfordian Radiolaria) from the Sula Islands (East Indies): their taxonomic, biostratigraphic, and paleogeographic significance (in prep.). Pessagno, E.A., Jr. and Newport, R.L., 1972. A technique for ex-
150 tracting Radiolaria from radiolarian cherts, Micropaleontology, 18 (2): 231-234. Pessagno, E.A., Jr., Blome, C.D. and Longoria, J.E, 1984. A revised radiolarian zonation for the Upper Jurassic of western North America. Bull. Am. Paleontol., 87 (320): 1-51. Pessagno, E.A., Jr., Whalen, EA. and Yeh, K.Y., 1986. Jurassic Nassellariina (Radiolaria) from North American Geologic Terranes. Bull. Am. Paleontol., 91 (326): 1-75. Pessagno, E.A., Jr., Longoria, J.E, MacLeod, N. and Six, W.M., 1987. Upper Jurassic (Kimmeridgian-upper Tithonian) Pantanelliidae from the Taman Formation, east-central Mexico: tectonostratigraphic, chronostratigraphic, and phylogenetic implications. In: S.J. Culver (Editor), Studies of North American Jurassic Radiolaria, Part I. Cushman Found. Foraminiferal Res. Spec. Publ., 23:1-51. Pessagno, E.A., Jr., Blome, C.D., Hull, D. and Six, W.M., Jr., 1993a. Middle and Upper Jurassic Radiolaria from the Western Klamath terrane, Smith River subterrane, northwestern California: their biostratigraphic, chronostratigraphic, geochronologic, and paleolatitudinal significance. Micropaleontology, 39 (2): 93-166. Pessagno, E.A., Jr., Hull, D.M., Longoria, J.E and Kelldorf, M.E., 1993b. Tectonostratigraphic significance of the San Pedro del Gallo area, Durango, Western Mexico. In: G. Dunn and K. McDougall (Editors), Mesozoic Paleogeography of the Western United States, II. Pacific Section SEPM, Book 71, pp. 141-156. Pessagno, E.A., Jr., Six, W.M. and Yang, Q., 1989. The Xiphostylidae Haeckel and Parvivaccidae, n. fam., (Radiolaria) from the North American Jurassic. Micropaleontology, 35: 193-255. Pindell, J.L., 1985. Alleghenian reconstruction and subsequent evolution of the Gulf of Mexico, Bahamas, and ProtoCaribbean. Tectonics, 4: 1-39. Pszcz6tkowski, A., 1978, Geosynclinal sequences of the Cordillera de Guaniguanico in western Cuba; their lithostratigraphy, facies development and paleogeography. Acta Geol. Pol., 28: 1-96. Pujana, I., 1989. Stratigraphical distribution of the multicyrtid Nassellariina (Radiolaria) at the Jurassic-Cretaceous boundary in the Neuqu6n Basin, Argentina. Zbl. Geol. Pal~iontol. I (56): 1043-1052. Pujana, I., 1991, Pantanelliidae (Radiolaria) from the Tithonian of the Vaca Muerta Formation, Neuqu6n, Argentina. Neues Jahrb. Geol. Paleontol. Abh., 180 (3): 391-408. Pujana, I., 1993. Middle Jurassic (Bathonian-Callovian) Radiolaria from Chacay Melehue, Cordillera del Viento, Province of Neuqu6n, Argentina. Master of Science thesis, Univ. of Texas at Dallas, pp. 1-87. Pujana, I., 1996. Occurrence of Vallupinae (Radiolaria) in the Neuqu6n Basin: Biostratigraphic Implications. In: A.C. Riccardi (Editor), Advances in Jurassic Research. GeoRes. Forum, 1-2: 459-456. Salvador, A., Westermann, G.E.G., O16ritz, E, Gordon, M.B. and Gursky, H.J., 1992. Meso-America. In: G.E.G. Westermann (Editor), The Jurassic of the Circum-Pacific. Cambridge University Press, Cambridge, pp. 1-676.
E.A. PESSAGNO et al. Sandoval, J. and Westermann, G.E.G., 1988. The Bajocian (Jurassic) ammonite fauna of Oaxaca, Mexico. J. Paleontol., 60: 1220-1271. Sandoval, J., Westermann, G.E.G. and Marshall, M.C., 1990. Ammonite fauna, stratigraphy, and ecology of the BathonianCallovian (Jurassic) Tecocoynca Group, south Mexico. Paleontographica, A210: 93-149. Sedlock, R.L., Ortega-Guti6rrez, E and Speed, R.C., 1993. Tectonostratigraphic terranes and tectonic evolution of Mexico. Geol. Soc. Am. Spec. Pap., 278: 1-153. Smith, EL., 1980. Correlation of the members of the Jurassic Snowshoe Formation in the Izee basin of east-central Oregon. Can. J. Sci., 17 (12): 1603-1608. Taylor, D.G., Callomon, J.H., Hall, R., Smith, EL., Tipper, H.W. and Westermann, G.E.G., 1984. Jurassic ammonite biogeography of western North America: the tectonic implications. In: G.E.G. Westermann (Editor), Jurassic-Cretaceous Biochronology and Paleogeography of North America. Geol. Assoc. Can. Spec. Pap., 27: 121-141. Tipper, H.W., 1981. Offset of an upper Pliensbachian geographic zonation in the North American Cordillera by transcurrent fault movement. Can. J. Earth Sci., 18: 1788-1792. Van der Voo, R., Mauk, EJ. and French, R.B., 1976. PermianTriassic continental configurations and the origin of the Gulf of Mexico. Geology, 4: 177-188. Verma, H.M. and Westermann, G.E.G., 1973. The Tithonian (Jurassic) ammonite fauna and stratigraphy of Sierra de Catorce, San Luis Potosi, Mexico. Bull. Am. Paleontol., 63: 107-278. von Hillebrandt, A., Smith, E, Westermann, G.E.G. and Callomon, J.H., 1992. Ammonite zones of the circum-Pacific region. In: G.E.G. Westermann (Editor), The Jurassic of the Circum-Pacific. Cambridge University Press, Cambridge, 676 PP. Walper, J., 1981. Tectonic evolution of the Gulf of Mexico. In: R.H. Pilger, Jr. (Editor), The Origin of the Gulf of Mexico and Early Opening of the Central North Atlantic Ocean. Proc. Symp. Louisiana State University, Baton Rouge, March 1980, pp. 27-98. Walper, J. and Rowett, C.L., 1972. Plate tectonics and the origin of the Caribbean Sea and the Gulf of M6xico. Gulf Coast Assoc. Geol. Soc. Trans., 22: 105-116. Whalen, EA., 1985. Lower Jurassic Radiolarian Biostratigraphy of the Kunga Formation, Queen Charlotte Islands, British Columbia and the San Hipolito Formation, Baja California Sur. Ph.D. dissertation, University of Texas at Dallas, 440 pp. Whalen, EA. and Pessagno, E.A., Jr., 1984. Lower Jurassic Radiolaria, San Hipolito Formation, Vizcaino Peninsula, Baja California Sur. In: V.A. Frizzell, Jr. (Editor), Geology of the Baja California Peninsula. Pacific Section SEPM, pp. 53-65. Wierzbowski, A., 1976. Oxfordian ammonites of the Pifiar del Rfo Province (western Cuba); their revision and Stratigraphical significance. Acta Geol. Pol., 26 (2): 138-260. Yeh, K. and Cheng, Y., 1996. Jurassic Radiolarians from the northwest coast of Busuanga Island, North Palawan Block, Philippines. Micropaleontology, 42 (2): 93-124.
Chapter 6
Cretaceous Carbonate B iostratigraphy and Environments in Honduras
ROBERT W. SCOTT and R.C. FINCH
A carbonate platform developed on the Chortfs block beginning in the Berriasian-Aptian and ending in the Albian. This platform is represented in Honduras by carbonate strata of the Yojoa Group, mainly the Atima Formation. We report new stratigraphic and paleontological data from exposures of this carbonate package and superjacent Upper Cretaceous strata in easternmost Honduras. In addition, new paleontological data from west-central Honduras are presented confirming previous interpretations there. In west-central Honduras, carbonate deposition of the Atima Formation was interrupted in the Aptian by an influx of non-marine, terrigenous sediments of the 'Mochito shale', and terminated at the beginning of the Cenomanian by coarse redbed strata of the lower Valle de Angeles Group. The exact nature of the Atima/Valle de Angeles contact in west-central Honduras is not known, but may be gradational. In the Montafias de Col6n area in eastern Honduras, no equivalent of the 'Mochito shale' has yet been identified, and late Albian deposition of the Atima was ended by an influx of fine-grained, shallow marine, terrigenous sediments of the Cenomanian 'Krausirpi beds'. Although the Krausirpi beds are only known locally, we include them in the Yojoa Group as an informal member. After deposition of the Krausirpi beds, local subaerial exposure and erosion occurred prior to deposition of the redbeds of the lower Valle de Angeles Group. In west-central Honduras, carbonate platform deposition was resumed during the Cenomanian represented by the Jaitique Formation and by the slightly younger and lithologically distinct Esqufas Formation just east of the outcrop area of the Jaitique. These limestone units conveniently divide the Valle de Angeles Group into lower and upper redbed sections. No equivalent limestone unit has been found in the Montafias de Col6n area. Instead, deposition of continental redbeds began here in the Cenomanian and continued into latest Cretaceous, as indicated by radiometrically dated volcanic flow units within the redbeds. The Jaitique Formation consists of a lower thick-bedded limestone member overlain by a thin shaly limestone member in the area south and west of Lake Yojoa. In central Honduras the Esqufas Formation consists of interbedded limestone, marl and shale. The remainder of Cretaceous deposition was fluvial clastic sediments and volcanic rocks which belong to the upper Valle de Angeles Group. New paleontologic data from limestone of the Atima Formation, shale of the Krausirpi beds, limestone conglomerate in the Valle de Angeles Group, and from the Jaitique Formation confirm the ages and depositional environments. The Atima ranges from Barremian-Aptian to late Albian based on ten foraminiferal taxa, one caprinid rudist and three calcareous algae. The Krausirpi beds are Albian-Cenomanian based on nine palynomorph taxa and two planktic foraminifers. The Jaitique is middle to late Cenomanian with the benthic foraminifer Biconcava, which is reported for the first time in the Caribbean Province. A Cenomanian dasyclad, Dissocladella undulata, was recovered in a limestone clast in the Valle de Angles Group suggesting the possibility that Cenomanian as well as older limestone units were eroded during the Late Cretaceous. The Atima records at least two shoaling upward depositional cycles, and the Jaitique was deposited in an open shelf environment that became restricted near the end of deposition.
INTRODUCTION
Cretaceous marine and non-marine strata cover much of the Chortfs block of northern Central America, and represent Early to mid-Cretaceous carbonate platform sedimentation buried by Late Cretaceous terrigenous redbeds. Honduras comprises most of
the Chortfs block, and published detailed stratigraphic and paleontologic data were, until recently, restricted mainly to the western and central portions of the republic. The geology of easternmost Honduras, however, was known from reconnaissance work by Mills et al. (1967), Mills and Hugh (1974), and Weiland et al. (1992). The first detailed strati-
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 151-165. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
152
R.W. SCOTT and R.C. FINCH
graphic data from this part of the Chortfs block has recently been provided by Rogers (1994, 1995) and Mills and Barton (1996). Mills and Barton (1996) used these Cretaceous carbonate strata in southeastern Honduras to define the Col6n platform. The geology and tectonic history of the Chortfs block has been summarized by Azema et al. (1985), Donnelly et al. (1990), Finch and Dengo (1990), and Gordon and Muehlberger (1994). The Chortfs block, which today forms the northwest comer of the Caribbean plate, is bounded on the southwest by the Middle America Trench subduction zone and on the north by the Cayman transform and its continental extension in the Motagua-Chixoy-Polochfc fault zones. To the south the Chortfs block is thought to be separated from the Chorotega block of the Caribbean plate by the Santa Elena fault-Hess escarpment. The eastern limit is undefined, but the Nicaraguan rise is generally included as a submarine extension of the Chortfs block (e.g., Pindell and Barrett, 1990). Prior to its Late Cretaceous suturing against the Maya block (Donnelly et al., 1990), the Chortfs was independent of the Maya, as indicated by differences in basement rocks (Gordon, 1989a,b; Gordon and Gose, 1989; Donnelly et al., 1990) and paleomagnetic data (Gose, 1985a,b; Gordon and Gose, 1989). Gose's paleomagnetic data also strongly imply that prior to its emplacement adjacent to the Maya block, Chortfs was not part of the Caribbean plate of today (Gose and Finch, 1992). Since mid-Cretaceous time the Chortfs block has rotated counterclockwise into
~,
+
I
\
MEXICO
''
+
/
r
,
" /"
L--
84"
+
#" San Pedro Sula
/~'
~,-oc.,,o i ~ ,,,,,,,,,, ,,. o.o HONDURAS . . = , o .... ,,
.r_
Guateemlla G
+',,,.GUATEMALA
+~'J
"---4 "
~
/"
W -P 1 6 ~ N
/
r ,t
k.
CARIBBEAN SEA
i BELl
"-,
_t
,"--
its present position. Consensus is growing, based on paleomagnetic data, lithologic similarities between Chortfs and the Guerrero block of Mexico, and plate reconstructions, that Chortfs formerly lay along the present southwest coast of Mexico (e.g, Anderson and Schmidt, 1983; Dengo, 1985; Gose, 1985a; Fourcade and Michaud, 1987; Pindell and Barrett, 1990; Burkart, 1994). As a consequence of its location west of Mexico during the Early Cretaceous, the new paleontological data reported here have biogeographic significance. The biota of the Chortfs block is very similar to that of the Caribbean Province. The primary purpose of this paper is to report new paleontological results from samples collected in easternmost Honduras by R.C. Finch and T.J. Weiland in a traverse across the Montafias de Col6n fold belt, and along the Rio Patuca and Rio Wampfi, and to place these results in the context of stratigraphic interpretations made by Rogers (1994, 1995). This report provides new data from a region lying some 200 km east of most of the previously studied areas of the Honduran Cretaceous. A secondary purpose is to describe other collections yielding new paleontological data from the region south and west of Lake Yojoa in west-central Honduras (Fig. 1). In 1984 Amoco Production Co. sponsored a field party to eastern Honduras to investigate and sample outcrops in the Montafias de Col6n-Rio Wampfi area near the border with Nicaragua (Fig. 1). This area consists of three distinct physiographic elements: (1)
.9.Metaphn
~v,~.
~ ,4,
"*~
L ~1,.,,.._jp-~.- 1~.
EL SALVADOR
~
/. " ~ t
*Ta~'n." ,
~.'~~" ::.
13.,t
",,,~
MQSQUITI&
Co,t
. . . . .o ./.-
Tegulligalpa
2
.-.,~.
.."
""
Jr-
r
+
?
PACIFIC OCEAN 3 %
0 ~- _
50 _
~
100
1,. ~ . . . )
NICARAGUA
200 I
km
- 4 - 12 ~ N
Fig. 1. Location map of selected quadrangles, geographic reference points; stippled pattern shows the approximate distribution of Mesozoic strata in Olancho east of the Guayape fault (from Donnelly et al., 1990, figs. 5 and 6). The Montafias de Col6n area is northeast of its name. Numbered quadrangles: ! = Confluencia Rfos Patuca y Wampfi quadrangle; 2 = Krausirpi quadrangle; 3 = Wampusirpi quadrangle; 4 = Agalteca quadrangle.
CRETACEOUS CARBONATE BIOSTRATIGRAPHY AND ENVIRONMENTS IN HONDURAS [
I
I
I
I
I
I
$ i ~0 ~/
:
ic.I h Cerro /Asangbusna
~/"
Rio Wampu
I
I
Boca ~A~_._...VVdl i I pU
"~ _ _ . , ..
-
~
,~-~,
. . . . . . . . . .. ... . . . . . . . -
/
...............
!
58
"~"
M...., "..~A-3 ~
9
6
./.. ~"~., " "1
, I
. . u
I
.
~,r~
9 9
~@"
. 1
z Km
20 I
I
I
de
~
.
""~ "/"
9
X2,~.'o
,-,Y ~"/
'~*' ~
,,.
9
--
~,
A 10
Colon 25 I
~
-,..=
r"
J_ -I1-
~"
J / 55
,,~
~
Tirisne
Montanas
I
9 e
ml,uJ.,,L'l.,,uummuun1,...unuukv
-
.,r l,-
Are~
9
.o % "~
9
9,,-
_
9
/,:L-I9 " ' - .9. , "9 .o,
~f~g,r.. - . - ~
5~
oo,OO
,4"
// Ie Wankibila .. /.; .: 9
I
.
,~ X,
153
"~'~ HA-1 "'.l ~ I 219
I
I
I
50 ..-J
Fig. 2. Location map of key samples in the Rfo Sutawala area; base map from the Confluencia Rfos Patuca y Wamp6 quadrangle scale 1:50.000. Southwest map corner begins at UTM grid coordinates 16PGM150407.
a northeast-trending belt of folded and faulted limestone mountains (the Montafias de Col6n), flanked on the northwest by (2) a low terrain of clastic strata flooring the broad northeast-trending valley of the Rio Patuca, flanked, in turn, to the northwest by (3) a mountainous mixed metamorphic/sedimentary terrain drained by the Rio Wamp6, a southeast-flowing tributary to the Patuca (Fig. 2). Paleontological results were obtained from a measured section representing a thin interval of the uppermost part of the Atima Formation exposed at Tirisne Cliffs and from other locations along the Sutawala Valley cutting across the Montafias de Col6n (Fig. 2). Additional collections were made along the Rios Patuca and Wamp6. Paleontological results from the present study demonstrate that Chortfs mid-Cretaceous carbonate strata were part of the same biogeographic province as coeval carbonate rocks of the Caribbean Province in Jamaica, northern Mexico and the Gulf of Mexico. Previous paleontological data from west-central Honduras (Mills et al., 1967; Lozej, 1976) suggested such a relationship. Likewise, outcrops of Atima Formation limestone reported by Emmet et al. (1992) in central Honduras (some 50 km north-northeast of Tegucigalpa in the Agalteca quadrangle, 4 in Fig. 1) contain late Aptian to middle Albian planktic foraminifers, colomiellids, and calcispheres. These forereef basinal fossils also demonstrate biogeographic similarities and exchange with coeval carbonate environments in Guatemala, Jamaica, Mexico and southern United States.
STRATIGRAPHY
Lithostratigraphy The Cretaceous stratigraphic terminology of Honduras (Fig. 3) was revised by Finch (1981) and summarized by Donnelly et al. (1990). The mid-Cretaceous platform carbonates are widely exposed across Honduras and have been designated the Yojoa Group (Mills et al., 1967), consisting of the locally developed Cantarranas Formation overlain by the widespread Atima Formation (see Gose and Finch, 1992, fig. 4, and Donnelly et al., 1990, fig. 7, for stratigraphic columns in west-central Honduras). The Cantarranas Formation (Carpenter, 1954) is a shaly limestone of variable lithology ranging in thickness from 30 to 190 m. The unit is dated as Valanginian-Hauterivian by ammonites (Gordon and Young, 1993) and represents a shelf lagoon. Outside of its type area near Talanga northeast of Tegucigalpa, the Cantarranas has been found to be of limited value as a map unit at the 1:50,000 scale used for geologic quadrangle mapping in Honduras, but has been mapped by Gordon (1990). We did not recognize the Cantarranas in the Rio Wampfi-Montafias de Col6n area, nor did Rogers (1994, 1995), although Mills et al. (1967) reported Cantarranas exposed on a tributary to the Wampfi. The Atima Formation, named by Mills (1959), is a widely recognized map unit consisting primarily of thick-bedded micritic limestone ranging in thickness from 90 to 1400 m (Finch, 1981). In the area of the
154
R.W. SCOTT and R.C. FINCH t...= rE o t.-
pper redbeds, conglomerate volcanics
8 Cenomanian
UpperAlbianLower Cenomanian Albian
=tique Formation limestone FTC-3, P-l, wer redbeds
2, 4, A-11
Krausirpi beds shale A-9, HA-1 ~.tima Formation T-1 to-29, A-3,-7.-4,-5 FSB-48, -47, -3, -4; limestone FZ-2S.-32 FM-21, -33; FSB~100
Pre-Albian
"Mochito Shale"
Honduras Group sandstone, shale & metasedimentary rocks above basement rocks ,m O N O
Agua Fria Formation
(9
=E
Basement Rocks igneous-metamorphic rocks
Fig. 3. Generalized, composited stratigraphic section of Mesozoic strata in eastern and western Honduras showing distribution of key samples listed in Tables 1 and 2. Krausirpi beds found only in the eastern region and the Jaitique Formationonly in the western region. Mochito Mine west of Lake Yojoa (Fig. 1) the Atima is divided into informal lower and upper intervals separated by about 100 m of green and gray mudstone with limestone interbeds and some red shale. In the Mochito area Lozej (1976) dated the lower limestone of the Atima as upper Barremian through Aptian by "possible 'early' Orbitolina" sp. and by Tintinnopsella cf. carpathica (Murgeanu and Filipescu), which in the Gulf Coast ranges from Berriasian through Barremian. A rudistid facies is present about 210 m below the top of the lower limestone. The upper Atima was dated as latest Aptian through upper Albian (Lozej, 1976) by the presence of orbitolinids, Cuneolina, Coskinolina/Dictyoconus, and Dicyclina. An interval between the orbitolinid facies contains the offshore shelf protozoan Colomiella recta Bonet, which ranges from Barremian through lower Albian in the Gulf Coast (Bonet, 1956). Rudistid facies are developed about 150 m below the top of the upper limestone. The shale and mudstone interval within the Atima Formation has been informally designated the "Mochito shale", and comprises a mappable unit at the
Mochito Mine near E1 Mochito (Fig. 1) and in the area north and west of the mine (Finch, 1973, 1985). This shale was described by Lozej (1976) as "of 'red bed' affinity (somewhat similar lithologically to the overlying and underlying formations)...". He reported conformable boundaries with the lower and upper Atima limestones "marked by green and/or red mottled mudrock interbeds". The transitional beds at the top of the lower Atima contain ostracodes and the base of the upper limestone contains charophytes, gastropods and ostracodes (Lozej, 1976). He concluded that the Mochito was late Aptian by its position between the Barremian-lower Aptian lower Atima limestone and the Albian upper Atima limestone. At three widely separated exposures in central Honduras, Gallo and Van Wagoner (1978) described sections 130 to 150 m thick of predominantly clastic strata within the Atima, which they correlated with the Mochito shale. In the Agalteca quadrangle (4 in Fig. 1), Emmet (1983a,b) also subdivided the Atima into upper and lower intervals separated by a shaly unit. It is not known if Emmet's un-named shale correlates with the Mochito shale, but it is
CRETACEOUS CARBONATE BIOSTRATIGRAPHY AND ENVIRONMENTS IN HONDURAS possible to infer that at least one major shale break in carbonate deposition is relatively widespread in the Atima Formation of central Honduras. This shale interval represents a regressive depositional phase, which contrasts with a widespread transgressive marine shale break between mid-Cretaceous carbonates in northern Mexico and the U.S. Gulf Coast (Scott et al., 1988). In the Montafias de Col6n area no Mochito shale-equivalent has been identified (Rogers, 1995). For the present study, the upper Atima Formation was sampled in the Montafias de Col6n in a traverse along the Rfo Sutawala that drains a prominent valley transecting the folded and faulted limestone mountains approximately perpendicular to stratigraphic and structural strike (Fig. 2). Along this traverse the strata dip southeastward 20 to 65 degrees and strike generally northeast-southwest. Brownish-gray limestone is found near the confluence of the Rfo Sutawala with the Rfo Patuca, in the Sutawala channel. Several kilometers upstream, and about 100 m higher topographically, a section of 54.8 m of the upper part of the Atima was measured and sampled bed by bed at the Tirisne Cliffs exposure, which is 4.5 km east of the junction of Sutwala with Rfo Patuca directly south of the Sutawala Valley trail in the Confluencia Rfos Patuca y Wamp6 quadrangle (Fig. 1, quadr. 1; 16PGM253530; Fig. 2, Table 1). Farther up the Sutawala valley from this measured section, shaly strata of the 'Krausirpi beds' are exposed, apparently in sedimentary contact overlying the upper Atima (Rogers, 1994, 1995). Additional collections were made northeast of Rfo Wamp6 in the Krausirpi and Wampusirpi quadrangles (Fig. 1, quadr. 2, 3). The base of the Yojoa Group is not exposed in the Montafias de Col6n. However, wherever the basal contact can be seen, the limestone overlies siliciclastic strata now generally assigned to the Honduras Group (Ritchie and Finch, 1985; Finch and Ritchie, 1990; Donnelly et al., 1990). The Honduras Group includes: (1) strata mapped as E1 Plan Formation near Tegucigalpa (Carpenter, 1954), (2) unnamed siliciclastic beds (Simonson, 1977, 1981), 3) strata formerly mapped on the Chortfs block as Todos Santos Formation (Finch, 1985; Ritchie and Finch, 1985; Emmet et al., 1992; Rogers, 1995), and (4) a thick succession of dark marine shale, marine and non-marine sandstone and conglomerate designated the Agua Frfa Formation in eastern and southeastern Honduras. The Agua Frfa Formation was named by Ritchie and Finch (1985) after a long section exposed near Danlf, which was first described by Roberts and Irving (1957). Agua Frfa strata have been widely mapped in eastern Honduras (Kozuch, 1989, 1991; Gordon, 1992; Finch and Ritchie, 1990; Weiland
155
et al., 1992; Rogers, 1994, 1995; Mills and Barton, 1996). Ammonites from marine portions of the Agua Frfa indicate a Middle Jurassic age (Ritchie and Finch, 1985; Gordon and Young, 1993). The Honduras Group underlies Lower to middle Cretaceous carbonates of the Yojoa Group with no detected unconformity or in fault contact. If there is no hiatus in deposition between the two groups, then deposition of the upper Honduras strata continued into the Early Cretaceous. However, in several areas, including the Rfo Wamp6 area, strata of the Agua Frfa have undergone a weak dynamic metamorphism (Kozuch, 1989; Gordon, 1990; Rogers, 1994, 1995) that did not affect the Yojoa Group limestone, leaving open the possibility that a significant hiatus occurred between deposition of the Honduras Group and the Yojoa Group. At the present time we do not know if Honduras Group deposition continued into the Early Cretaceous. In west-central Honduras carbonate strata of the Atima Formation are overlain, apparently gradationally, by red terrigenous clastic strata of the Valle de Angeles Group, named by Carpenter (1954) for exposures between Tegucigalpa and San Juancito (Fig. 1). In the area west and south of Lake Yojoa (Fig. 1) the Valle de Angeles redbeds are divided into lower and upper intervals by the presence of a prominent limestone unit, the Jaitique Formation (Finch, 1981). Somewhat to the east of Lake Yojoa and north of Tegucigalpa, the Esqufas Formation (Home et al., 1974) divides the Valle de Angeles. The Jaitique Formation includes two units, an unnamed cliff-forming limestone member, overlain by a thin section of distinctive petroliferous limestone designated the Guare Member (Finch, 1981). The lower member is about 100-150 m thick, comprised of thick-bedded shelfal limestone, well-dated paleontologically as Cenomanian. No unconformity was seen in the field at the base of the Jaitique, and paleomagnetic data suggest that the contact between the basal Jaitique and the underlying redbeds of the lower Valle de Angeles Group is conformable (W.A. Gose, in Finch, 1981). The Guare is about 10 m thick, consisting of thin- to medium-bedded, flaggy weathering, oil-stained, laminated, dark-gray to black limestone with black shale interbeds. The Guare is in turn overlain by laminated gypsum that grades upwards into fine-grained red clastic strata of the upper Valle de Angeles Group (Finch, 1981). Farther east, in central Honduras near the village Esqufas (Fig. 1) marly and argillaceous limestone of the Esqufas Formation (Weaver, 1942; Atwood, 1972; Home et al., 1974; Finch, 1981) divides the redbeds. Like the Jaitique Formation, the Esqufas apparently is in conformable contact with the enclosing redbeds (Home et al., 1974). The age of the unit is not well defined, paleontologically,
156
R.W. S C O T T and R.C. F I N C H
Table 1 Age-diagnostic taxa in samples from the Tirisne Cliffs measured section, Honduras, along cliffs directly south of the Sutawala valley trail and about 5 km east-southeast of the Rfo Patuca Field No.: Cum. thickn. (m):
Barkerina barkerensis Cuneolina walteri Globigerina delrioensis Globigerina planispira? Globochaeta alpina Lenticulina sp. Heterohelix globulosa Nezzazata simplex Nodosaria sp. Praeg lob igerina delrioensis Pseudocyclammina hebergi Radiolaria Sponge spicules
Micritosphaera ovalis Pithonella ovalis Pithonella sphaerica Saccomma sp.
T1 0-1.5
T2 3.6
T3 3.9
T4 4.5
T5 5.4
T6 6.3
T7 7.9
T8 8.9
T9 9.5
T10 10.0
R
R?
Tll 10.6
Barkerina barkerensis Cuneolina walteri Globigerina delrioensis Globigerina planispira? Globochaeta alpina Lenticulina sp. Heterohelix globulosa Nezzazata simplex Nodosaria sp. Praeglobigerina delrioensis Pseudocyclammina hebergi Radiolaria Sponge spicules
Micritosphaera ovalis Pithonella ovalis Pithonella sphaerica Polystrata alba Saccomma sp.
T13 T14 1 6 . 1 17.8
R? R? R?
R
R
R
R?
R?
R R
R
F
R
R R?
R
Oc
F
R F
F
R?
R? R?
F
F
R
R
T20 29.1
R
R
R R F
R
R
R
R
R
R
F
R R
R R R
F
R
R
R A
R
Oc
Oc
R R
R
R
R R
T26 44.8
Indet. calc. algae Indet. caprinids Field No." Cum. thickn. (m)"
T12 13.4
T15 20.6
R
T16 22.0
T17 22.4
T18 24.3
T19 27.8
R
R?
R? R?
T22 33.4
R?
R R
T23 35.7
T24 39.9
T25 41.8
R?
R?
R?
R
R
R
T21 31.8
R
R
R
R
T28 52.1
T29 54.8
R
R
F
R
R
R
R
R R?
R
T27 50.0
R
F
R
R F F
F
Oc
R?
R?
R
R
R
R
R R
R
R
R
R
R
R? R
R
Indet. calc. algae Indet. caprinids Indet. radiolitids?
R R
R?
R
R Oc
9
Long. 84~ lat. 14o 56 ! 35 I! N. Collected by T. Weiland, measured from lowest bed exposed at west edge of cliff face. Relative abundance scheme: R = rare, F = few, Oc = occasional, A -- abundant.
but a Cenomanian-Turonian age is suggested, and its paleopole position indicates that the Esqufas is somewhat younger than the well-dated Cenomanian Jaitique Formation (Gose and Finch, 1992) Although the Jaitique and Esqufas formations differ in lithology and somewhat in age, where present, each conveniently divides the thick Valle de Angeles Group into upper and lower redbed intervals, and provide some constraint on the age of the lower redbeds. The lower Valle de Angeles redbeds lie above the paleontologically well-dated upper Atima
Formation of mainly Albian age, and below the Jaitique of Cenomanian age or the Esqufas of upper Cenomanian or Turonian age. The age of the lower Valle de Angeles redbeds are, therefore constrained to upper Albian to Lower Cenomanian. Thus, in western and central Honduras, Yojoa carbonate deposition appears to have ceased toward the close of the Albian and to have been superseded by redbed clastic deposition by early Cenomanian time. In west-central Honduras the upper contact of the Valle de Angeles Group is an angular unconformity,
CRETACEOUS CARBONATE BIOSTRATIGRAPHY AND ENVIRONMENTS IN HONDURAS with the redbeds overlain by middle Tertiary volcanic rocks of the Matagalpa Formation or by the Miocene-Pliocene Padre Miguel Group (Williams and McBirney, 1969; Everett, 1970; Finch, 1972, 1981; Emmet, 1983b). It is not known if redbed sedimentation continued into the early Tertiary. Neither the Jaitique nor the Esqufas Formation has been found within the Valle de Angeles Group in the Rfo Wampfi-R/o Patuca area (Weiland et al., 1992; Rogers, 1994, 1995). Instead, in the Montafias de Col6n region of easternmost Honduras, the upper Albian Atima Formation is overlain apparently conformably by non-red, olive tan shale and silty, shallow marine strata. These strata were designated the 'Krausirpi beds' by Rogers (1994, 1995), who reported the unit to include gray and tan shale, arkosic sandstone, graywacke, and minor conglomerate. Rare wood fragments were observed in clastic beds. However, the shale and sandstone beds are calcareous, and thin limestone beds occur within the unit. In thin section one sample is a silty planktic foram lime mudstone. A sample collected from the Krausirpi beds by Finch and Weiland contained a latest Albian to Early Cenomanian biota. At the village of Krausirpi (Fig. 2), Rogers found an erosional surface developed on top of the Krausirpi beds with the Valle de Angeles redbeds overlying them in a slightly angular unconformity. Where the Krausirpi beds are missing, Valle de Angeles redbeds directly and unconformably overlie the Atima (Rogers, 1995). South of Cerro Wampfi (Fig. 2) karstic features are developed at the top of the Atima (Rogers, 1995). Rogers suspects that the Atima was subaerially exposed prior to deposition of the Valle de Angeles in this region (R.D. Rogers, pers. commun., 1996). Thus, in the Montafias de Col6n region, marine deposition continued into the Cenomanian, and deposition of the Valle de Angeles did not commence until latest Cretaceous. Although Rogers (1995) noted that it is "possible that the Krausirpi beds are a local unit of the Yojoa Group", he concluded that they "should not be correlated with other mapped units in Honduras". However, we feel that, in spite of their local nature, it is appropriate to include the Krausirpi beds in the Yojoa Group as an informal member because of their marine character, the absence of redbeds, their conformable contact with the Atima, and their unconformable contact with the Valle de Angeles. Within the Valle de Angeles Group throughout Honduras, conglomerates of various compositions form a significant portion of the unit. In west-central Honduras the lower Valle de Angeles section is more conglomeratic than the upper, with quartz pebble conglomerate beds being prominent ledge-formers in outcrop. However, the most distinctive conglomerate units consist of limestone pebble- to boulder-con-
157
glomerate, with the clasts embedded within a redbed matrix. These limestone conglomerates, commonly designated the Ilama Formation (Mills et al., 1967; Southernwood, 1986; Mills and Barton, 1996) are found throughout Honduras. Available paleontological evidence indicates that the limestone clasts were derived from the Atima Formation and the conglomerates certainly indicate widespread erosion of Yojoa Group carbonates during Valle de Angeles time. However, there is no good evidence that limestone conglomerate production was coeval across Honduras, or that the various conglomerate exposures should all be assigned to one formation. Indeed, Southernwood's stratigraphic correlation chart (Southernwood, 1986, plate 4) indicates a wide variety of stratigraphic positions for the conglomerates of the 'Ilama Formation', as would be expected for such coarse, fluvial deposits, which possibly were deposited as alluvial fans shed from fault blocks. 'Ilama-type' limestone conglomerate deposits occur in the valley of the Rio Patuca, along the northwest flank of the Montafias de Col6n, within redbeds assigned to the Valle de Angeles Group (Rogers, 1994, 1995; Mills and Barton, 1996). Limestone clasts in these conglomerates sampled by Finch and Weiland yielded Cretaceous microfossils and megafossils that suggest that most clasts were derived from the Atima Formation. A single sample contains a dasyclad alga known only in Cenomanian rocks elsewhere in the Tethys suggesting that Cenomanian limestone was a local source in the Montafias de Col6n, although no in-situ outcrops are mapped. The Krausirpi beds contain local limestone beds that may be of Cenomanian age and could have been a source. Rogers (1994, 1995) has demonstrated that the volcanic units dated by Weiland et al. (1992; K Ar dates from unaltered plagioclase separations) at 80-70 Ma are interbedded with the Valle de Angeles redbeds. Thus, the Valle de Angeles redbed sedimentation continued at least into the latest Cretaceous in easternmost Honduras, but at present there is, as in west-central Honduras, no evidence to show that redbed deposition continued as late as Tertiary time.
Biostratigraphy The Atima Formation in the Montafias de Col6n region is as young as upper Albian. The age of the 55-m-thick section of the uppermost Atima Formation at Tirisne Cliffs section is upper Albian based on the co-occurrence of Cuneolina walteri Cushman and Applin (Fig. 4A-C), Pseudocyclammina hedbergi Maync (Fig. 4D,E), and Praeglobotruncana delrioensis (Plummer). The first two species occur in middle and upper Albian carbonates in northwestern Mexico (Scott and Gonzales-Leon, 1991) and elsewhere in the Gulf Coast (Conkin and Conkin,
158
R.W. SCOTT and R.C. FINCH
CRETACEOUS CARBONATE BIOSTRATIGRAPHY AND ENVIRONMENTS IN HONDURAS 1958). The third species, a planktic foraminifer, ranges from late Albian to late Cenomanian (Bolli et al., 1985). Samples from the limestone section in the Rfo Sutawala stratigraphically below the Tirisne Cliffs section yield Pseudonummoloculina [Nummoloculina] heimi (Bonet) (Fig. 4F), which ranges through Albian and lower Cenomanian rocks in the Gulf Coast region (Scott and Gonzales-Leon, 1991). The genetic reassignment of this species was made by Hottinger et al. (1989) because the apertural structure differs from the Neogene genus Nummoloculina. From west-central Honduras we can demonstrate the presence of the key benthic foraminifer Orbitolina (Mesorbitolina) subconcava Leymerie (Fig. 5D,E) in the Atima Formation in several places, which probably was the more advanced type reported by Lozej (1976). This species ranges from the late Aptian to the late Albian and is reported in lower Albian strata of Texas and Mexico (Schroeder and Neumann, 1985). This species occurs in the uppermost Atima a few meters below its contact with the Valle de Angeles Formation southwest of Lake Yojoa (in the Montafia de Santa B~irbara, Table 2). This suggests that the upper Atima is not younger than upper Albian. The species also occurs in a limestone clast within a conglomerate in the Valle de Angeles along the Rfo Patuca (Fig. 2). Other benthic foraminifers support the lower to upper Albian age of the upper Atima limestone. Coskinolinoides sp. cf. C. texanus Keijzer is found southwest and west of E1 Mochito near Lake Yojoa (in the San Pedro Zacapa and the Santa Bfirbara quadrangles, Table 2). This species occurs in the middle to lower-upper Albian Fredericksburg Group in central Texas and unnamed species of this genus range up into the upper Albian part of the Washita Group (Coogan, 1977). Praechrysalidina infracretacea Luperto Sinni, which we found in the Atima with Orbitolina texana (Roemer) and Cuneolina walteri southwest of Lake Yojoa (in the Santa Bfirbara quadrangle), is a Hauterivian-Albian species in the Middle East (Banner et al., 1991). Another Albian taxon in the upper Atima in this area is Coskinolinella sp., which is well illustrated by Schroeder and Neumann (1985). Outcrop grab samples from the Olancho district (Fig. 1, in the area indicated by the name Olancho) collected by M.B. Gordon (pers. commun.,
159
1990) contain key taxa of rudists and foraminifers that confirm the age of the Atima to range from Barremian-lower Aptian to upper Albian. Poorly preserved rudist specimens from a caprinid packstone in the upper 50 m of the Atima Formation are close to either Kimbleia sp. (late Albian), or Caprinula sp. (Albian-Turonian) (Table 2). Other samples of the Atima contain the Barremian-early Aptian foraminifer Choffatella decipiens Schlumberger with Pseudocyclammina hedbergi, the bivalve Chondrodonta sp., and the alga Lithocodium aggregatum Elliott. All of these taxa are common in theGulf of Mexico part of the Caribbean Province. The post-Atima Krausirpi beds exposed in the bottom of Rfo Sutawala are represented by two samples. One yielded late Albian-early Cenomanian palynomorphs identified by R.W. Aurisano (written commun., 1985): Reticulisporitesjardinus, Xenascus ceratioides, Dinopterygium sp., Diconodinium sp., Oligosphaeridium sp., Classopollis sp., Coronifera sp., Scriniodinium sp., Trichodinium sp., and tricolpate pollen. A thin section of a second sample contains Globigerinoides cushmani and Heterohelix globulosa, which together indicate a latest Albian to early Cenomanian age for the Krausirpi beds, showing that shallow marine sedimentation continued into the Cenomanian in the eastern part of the Chortfs block. Limestone conglomerate in the Valle de Angeles Group exposed along the Rfo Patuca in the Krausirpi and Wampusirpi quadrangles also contains late Albian fossils (Table 2). Important taxa are Orbitolina subconcava, Cuneolina walteri, Pseudonummoloculina [Nummoloculina] heimi, Pseudocyclammina hedbergi, Favusella [Globigerina] washitensis (Carsey), and Globigerinoides cushmani, the caprinid Mexicaprina sp., and calcareous algae, Polystrata alba (Pfender) and Parachaetetes texana Johnson (Fig. 5A). Evidently these cobbles were eroded from the Atima Formation At Boca Wampfi a conglomerate with a reddish brown sandy matrix and cobbles of red siltstone and well rounded, light-gray limestone was identified in the Valle de Angeles Group, although the limestone cobbles are lighter colored than most of the Atima Formation and they contain miliolids, textulariids, and bivalves. A limestone clast from this conglomerate yields the calcareous dasyclad alga, Dissocladella undulata (Raineri) (Fig. 4G), which is
Fig. 4. (A-G) Photomicrographs of key fossils found in the Atima Formation in eastern Honduras and the Jaitique Limestone in central Honduras. (A-C) Cuneolina walteri Cushman and Applin (1947); (A) proloculus, sample A-4, 75 x; (B) transverse section, sample A-4, 75• (C) longitudinal section, sample T-29, 47x. (D-E) Pseudocyclammina hedbergi Maync (1953); (D) axial view, sample T-7, 30x; (E) longitudinal view, sample T-7, 30x. (F) Pseudonummoloculina [Nummoloculina] heimi (Bonet, 1956), emend. Conkin and Conkin (1958), generic reassignment by Hottinger et al. (1989), sample A-4, 75x. This specimen has flatter chambers than is typical of P. heimi. (G) Dissocladella undulata (Raineri, 1922; Pia, 1936), sample A-11, Valle de Angeles Group, 55x. (H) Biconcava n. sp., Jaitique Limestone, sample FTC-3, 60•
160
R.W. SCOTT and R.C. FINCH
CRETACEOUS CARBONATE BIOSTRATIGRAPHY AND ENVIRONMENTS IN HONDURAS
161
Table 2 Location and description of key samples collected in Honduras by R.C. Finch during several field seasons (locations given using the Universal Transverse Mercator Grid system)
Krausirpi quadrangle P-1 Valle de Angeles Gp. at 16PGM257617, Krautara on Rfo Patuca; cobble conglomerate of light-gray limestone and lesser amounts of volcaniclastics, quartz, and quartzite of silicified siltstone in a matrix of reddish sandstone. Fossils in limestone clasts: Orbitolina subconcava, Nummoloculina heimi, Cuneolina walteri, Pseudocyclammina hedbergi, Parachatetes texana, Lithocodium aggregatum. P-2 Valle de Angeles Gp. at 16PGM346681, Pimienta on Rfo Patuca; 'Ilama' limestone conglomerate as above. Fossils in limestone clasts: Mexicaprina sp., Lithocodium aggregatum. Wampusirpi quadrangle P-4 Valle de Angeles Gp. at 16PGM523723, Walpatanta on Rfo Patuca; 'Ilama' limestone conglomerate as above. Fossils in limestone clasts: Globigerina delrioensis, Favusella washitensis, Globigerellinoides cushmani, Parachaetetes texana, Polystrata alba, Lithocodium aggregatum. Confluencia Rfos Patuca y Wampft quadrangle A series Atima Fm. collected by core; all are limestone except A-9 which is shale of the Krausirpi beds and A-11 which is conglomerate in Valle de Angeles Gp.; A-l: 16PGM221544; A-2: 16PGM221543; A-3: 16PGM223540; A-4: 16PGM234539; A-6: 16PGM227540; A-7: 16PGM231542; A-8: 16PGM247533; A-9: 16PGM262529; A-10: 16PGM279520; A-11: 16PGMI74547. T series Atima Fm. collected in measured section at Tirisne Cliffs by T. Weiland; T-1 (base) through T-29 (top) at 16PGM253530, 84054, 19"W, 14~ HA-locality in Rfo Sutawala estimated at 16PGM289512. Santa B6rbara quadr. (west side of Lake Yojoa in Fig. 1) FSB-3 Uppermost Atima Fm. no more than a few tens of meters below contact with Valle de Angeles Fm. at 16PCM711409. Orbitolina subconcava. FSB-4 Uppermost Atima Fm., no more than a few tens of meters below contact with Valle de Angeles Fm. in a quarry near 16PCM844433. Coskinolinella sp. FSB-46 Upper Atima Fm., very probably the upper part on top of Montafia Santa B~irbara near 16PCM788454. Coskinolinella texanus, Praechrysalidina infracretacea. FSB-47 Upper Atima Fm., probably the upper part, from near 16PCM828419 in the Mochito graben. Coskinolinella texanus. FSB-100 Atima Fm., low on the southwest flank of Montafia Santa B~irbara at about 16PCM717535 apparently above the 'Mochito shale'. Orbitoliina texana, Praechrysalidina infracretacea. FM-21 Upper Atima Fm., possibly lower part from inside the Mochito Mine at level 1350, cross-cut 9444E. Praechrysalidina infracretacea (16PCM9394230). FM-33 Upper Atima Fm., lower part near contact with 'Mochito shale' inside Mochito Mine at level 850, cross-cut 992. Colomiella coahuilensis (16PCM939423). Taulab# quadr. (southeast of Lake Yojoa in Fig. 1) FTC-3 Uppermost part of lower, unnamed member of Jaitique Fm. directly below basal contact of Guare Mbr. at 16PCM998218 near Carrizal-La Mision. See fig. 4 in Finch (1981) for stratigraphic position. Specimens of Biconcava n. sp. in thin section. San Pedro Zacapa quadr. (southwest of Lake Yojoa in Fig. 1) FZ-28, 32 Uppermost Atima Fm. near 16PCM870343; FZ-32 is approximately 50 m above FZ-28. Coskinolinella texanus.
characteristic of the C e n o m a n i a n in the Tethys (Bassoullet et al., 1978). This is the first report of this species in the C a r i b b e a n P r o v i n c e . This l i m e s t o n e clast m a y have b e e n d e r i v e d f r o m one of the C e n o m a n i a n l i m e s t o n e units within the Valle de A n g e l e s , or f r o m a l i m e s t o n e b e d within the K r a u s i r pi beds. If this is the case, it is the first t e n u o u s e v i d e n c e that C e n o m a n i a n l i m e s t o n e d e p o s i t i o n m a y h a v e e x t e n d e d this far east. To date, no l i m e s t o n e unit has b e e n m a p p e d within the Valle de A n g e l e s r e d b e d s in the W a m p d - P a t u c a r e g i o n ( W e i l a n d et al., 1992; R o g e r s , 1995).
T h e rudist M e x i c a p r i n a sp. is r e p o r t e d f r o m H o n duras for the first time. This g e n u s is charact eri s tic of u p p e r A l b i a n rocks in M e x i c o and the U.S. T h e s p e c i m e n was f o u n d in a c o n g l o m e r a t e within the Valle de A n g e l e s G r o u p and is c o m p l e t e l y recrystallized so that internal structures are not visible (Fig. 5F). This s p e c i m e n has the distinctive ridges at the corners similar to M e x i c a p r i n a sp. illustrated by C o o g a n (1977, pl. 17, figs. 8), w h i c h is f r o m the E1 A b r a F o r m a t i o n at Taninul quarry in M e x i c o . T h e H o n d u r a n s p e c i m e n , h o w e v e r , is a b o u t 10 m m across and the M e x i c a n s p e c i m e n is a b o u t 23 m m
Fig. 5. Photomicrographs of fossils from limestone clasts in conglomerates of the Valle des Angeles Group (A, D, E, F) in eastern Honduras and the Jaitique Limestone (B, C) in west-central Honduras. (A) Parachaetetes texana Johnson, sample P-4, 55• (B, C) Biconcava n. sp, sample FTC-3, sagittal and axial sections, 151 x. (D, E) Orbitolina (Mesorbitolina) subconcava Leymerie, sample P-1, 47 x. (F) Mexicaprina sp., sample P-2, transverse section, 7.4 x.
162 in longest diameter. A similar species, Mexicaprina minuta Coogan, is 8 to 11 mm across, but it does not have the large ridges. The Honduran specimen may be a new species of this very specialized genus. Coogan (1973) described the genus from the E1Abra Formation and from the Stuart City Formation in the subsurface of Texas (Coogan, 1977, pl. 17, fig. 5). The age of the Stuart City is middle to lowerupper Albian by its foraminifers and rudists (Scott, 1990). The age of the E1 Abra is less clear; Coogan (1973, 1977) placed it in the Cenomanian based on the radiolitid rudists, Pecten roemeri and Parkeria sphaerica, and the absence of characteristic Albian taxa. The age of the oldest radiolitids is indefinite; Pecten roemeri is found in lower Cenomanian rocks of north Texas, and Parkeria sphaerica is originally known from the uppermost Albian beds in England and its Cenomanian attribution comes from its occurrence in the E1 Abra (Dieni and Turn~ek, 1979). Recently, Mexicaprina minuta and Mexicaprina cornuta have been found with the late Albian rudists Kimbleia and Caprinuloidea in northern Mexico (Alba and L6pez-Casillas, 1993). Therefore, the age significance of Mexicaprina spp. is late Albian. The Jaitique Formation has yielded specimens of the foraminifera Biconcava n. sp. (Fig. 4H, Fig. 5B,C), which is reported from the Caribbean Province for the first time. The sample was collected southeast of Lake Yojoa in the Taulab6 area from the uppermost part of the lower, unnamed member directly below basal contact of Guare Member (fig. 4 in Finch, 1981). These specimens possess the major features of the genus as defined by Hamoui and Saint-Marc (1970): the planispiral test is slightly involute and expands gradually; the proloculus is simple, globular; chamber cross-section is lunate but the periphery of these specimens is more convex than in the genotype, Biconcava bentori Hamoui and Saint-Marc (1970). The wall of both species is microgranular calcite. The aperture appears to be a simple pore in the septal face (Fig. 5B, thirdlast septa in final whorl). B. bentori ranges through the middle and upper Cenomanian in its type area of Lebanon and Israel, but similar specimens have been reported in upper Albian and Senonian strata (Schroeder and Neumann, 1985), so the range of the genus is yet to be defined precisely.
PALEOENVIRoNMENTS
The upper Atima Formation exposed near the mouth of the Rfo Sutawala records a transgressive succession. The beds dip southeast and the samples were collected from outcrops in the valley floor. Top: rudist-peloid packstone; rudist packstone-wackestone; quartz-peloid packstone; echi-
R.W. SCOTT and R.C. FINCH noid wackestone with pycnodont oysters; red shale with plant debris (probably not Atima); silty peloidostracode packstone. Farther upstream in the Sutawala Valley, at Tirisne Cliffs, 55 m of upper Atima Formation were measured and closely sampled. This section records a shoaling upwards facies succession (Fig. 6). Two major microfacies were collected: a lower peloidforam-spicule wackestone/packstone and an upper peloid-foram-bivalve wackestone/grainstone. These facies indicate an environmental change from the middle shelf to the inner shelf. Overall, taxonomic diversity decreases slightly near the top of the section where carbonate sands were more common. The upwards shoaling is indicated by the decrease in number of planktic taxa, the decrease in sponge spicules, and the replacement of Lenticulina sp. and Nodosaria sp. by shallow-water benthic foraminifers. Bottom waters were well oxygenated and water energy increased up-section where grainstone is more common. The post-Atima Krausirpi beds are poorly exposed. This unit represents an influx of terrigenous sediments into a shallow marine environment terminating the Atima carbonate platform. Bottom waters may have been poorly oxygenated. The depositional environment may have been prodeltaic (Rogers, 1994, 1995). The limestone conglomerates within the Valle de Angeles Group along the Rfo Patuca in the Krausirpi and Wampusirpi quadrangles were deposited in non-marine to nearshore environments. The limestone clasts were derived mainly from the Atima Formation, which was deposited in a shallowwater, moderate- to high-energy environment with rudists. The common facies are rudist packstone, intraclast-peloid grainstone and rudist-peloid grainstone. Some clasts represent quiet water lagoonal or shelf environments with foram wackestone. One clast contains radiolarian-planktic foram wackestone in contact with peloid grainstone turbidite. So the cobbles represent a suite of environments ranging from the shelf margin to the slope. A major shift in depositional regime from carbonate shelf to terrigenous shelf, possibly prodeltaic, in eastern Honduras is recorded by the termination of Atima deposition by an influx of terrigenous sediments of the lower Cenomanian Krausirpi beds. Water depths were up to approximately 50 m to accommodate the planktic foraminifers and dinoflagellates. Deposition of the Krausirpi beds was followed by a period of exposure, during which an erosional unconformity developed that cut out part of the Krausirpi beds and, locally, exposed the top of the Atima Formation. This exposure event on the eastern part of the Chortfs block may correspond with one of several middle Cenomanian sequence boundaries
CRETACEOUS CARBONATE BIOSTRATIGRAPHY AND ENVIRONMENTS IN HONDURAS M 55
"=''I
I
I
I
I
I
I
I
I
I
I
I
I
I
163
I
W UJZ
",2
I=6.
<
50-
~,,,,
U.I
O Z u.O OW
40"
,,-I,
_
g
S
~3
Pseudoc~ .,.~edber!
-
~..
7 N
"n
oer" ILl
20' w !11 z
e~
!
g
,,
,,, UJ
10
="1 w q~ ~ v ~
IJ.l
-Praeglob - delrio -j ,
0
BASEOF EXPOSURE
5
,PBarkerina ~
,barkerensis ~
, i lO 15 NO. TAXA
~
"
' A ' o' ' ' ' '5'0' o ' ~ ' ' A' 0 .... % PLANKTIC SPICULES BENTHIC DEPOSITIONAL TAXA FORAMS ENVIRONMENT
Fig. 6. Paleoenvironmental interpretation of the Atima Formation at the Tirisne Cliffs section, Sutwala valley, Rio Wampd, Honduras. Relative abundance ranges from absent (0), to rare, few, occasional, common, and abundant (A). (Haq et al., 1988); the unconformity at the lowermiddle Cenomanian boundary is widespread in the U.S. Gulf Coast and Arabian shelf and platform sections (Scott et al., 1988). Deposition of redbeds of the Valle de Angeles Group buried this exposure surface, which, locally, overlies karstic Atima (Rogers, 1995). In west-central Honduras a similar, but gradual environmental shift deposited conglomeratic redbeds with Atima limestone clasts prior to deposition of the Jaitique and Esqufas formations. However, in the area south and west of Lake Yojoa some occurrences of limestone conglomerate containing abundant light-gray limestone clasts that do not look like typical Atima limestone were mapped by Finch (1972, 1985) as within the upper Valle de Angeles. So a second, younger local erosional event may be recorded in this area.
SUMMARY New collections from Cretaceous carbonate strata in eastern Honduras yield taxa new to the area that confirm the upper Albian age of the upper part of the Atima Formation. Some of these taxa have been reported to the generic level, but not illustrated nor identified to species: Cuneolina walteri (Fig. 4A-C), Pseudocyclammina hedbergi (Fig. 4D,E), and Pseudonummoloculina [Nummoloculina] heimi. Some have never before been re-
ported in Honduras: Mexicaprina sp., Kimbleia sp. (late Albian), or Caprinula sp. (Albian-Turonian), and Praeglobotruncana delrioensis. The presence of Choffatella decipiens with Pseudocyclammina hedbergi on the western margin of the Montafias de Col6n indicates that the lowermost part of the Atima crops out there. In west-central Honduras the upper Aptian to upper Albian Orbitolina subconcava is documented for the first time, as is Coskinolinoides sp. cf.
C. texanus, Orbitolina texana, Cuneolina walteri, Coskinolinella, and Coskinolinella sp. The Atima Formation in eastern Honduras is overlain conformably by a marine shale, the Krausirpi beds (Rogers, 1994, 1995), that contain late Albian to early Cenomanian dinoflagellates and planktic foraminifers. The limestone conglomerates in the Valle de Angeles Group were derived mainly from the Atima; however, some may have been derived from a Cenomanian limestone unit, as suggested by the presence of the dasyclad alga Dissocladella undulata, which is reported here from the Caribbean Province for the first time. Accepting the placement of Chortfs near the Guerrero block on the Pacific side of Mexico during the mid-Jurassic (Anderson and Schmidt, 1983; Dengo, 1985; Pindell and Barrett, 1990), the biota reported here confirm that by Aptian and Albian time, the Chortfs block had moved into biogeographic connection with the fauna of the Caribbean
164
R.W. SCOTT and R.C. FINCH
P r o v i n c e . A n e a s t w a r d c o m p o n e n t o f t r a n s l a t i o n is i m p l i e d , likely s o u t h e a s t w a r d a l o n g a structure like the A c a p u l c o - G u a t e m a l a
megashear of Anderson
and S c h m i d t (1983).
REFERENCES
Alba, J.A. and L6pez-Casillas, 1993. Mid-Cretaceous rudists (Bivalvia-Hippuritacea) from Durango, north-central Mexico. Proc. 3rd Int. Conf. Rudists, UNAM, Inst. de Geologfa, pp. 2-3. Anderson, T.H. and Schmidt, V.A., 1983. The evolution of Middle America and the Gulf of Mexico-Caribbean Sea region during Mesozoic time. Geol. Soc. Am. Bull. 94: 941-966. Atwood, M.G., 1972. Geology of the Minas de Oro Quadrangle, Honduras, Central America. Unpubl. M.A. thesis, Wesleyan University, Middletown, CT, 88 p. Azema, J., Biju-Duval, B., Bizon, J.J, Carfantan, J.C., Masclr, A.J. and Tardy, M., 1985. Le Honduras (Am6rique centrale nucl6aire) et le bloc d'Oaxaxa (Sud du Mexique): deux ensembles comparables du continent Nord-Am6ricain s6par6s par le jeu d6crochant s6estre des failles du syst~me Polochic-Motagua. Symp. Geodynamique des Caribes, Paris, 5-8 Fevrier 1985. Editions Technipp, Paris, pp. 427-438. Banner, ET., Simmons, M.D. and J.E. Whittaker, 1991. The Mesozoic Chrysalidinidae (Foraminifera, Textulariacea) of the Middle East: the Redmond (Aramco) taxa and their relatives. Bull. Br. Mus. Nat. Hist. (Geol.), 47: 101-152. Bassoullet, J.P., Bernier, P., Conrad, M.A., Deloffre, R. and Jaffrezo, M., 1978. Les algues dasycladales du Jurassique et du Cr6tac6. Geobios, M6m. Sp6c., 2, 330 pp., 40 pl. Bolli, H.M., Saunders, J.B. and Perch-Nielsen, K., 1985. Plankton Stratigraphy. Cambridge Univ. Press, London, 1032 pp. Bonet, E, 1956. Zonifacaci6n microfaunistica de las calizas cret~icicas del este de M6xico. Asoc. Mex. Ge61. Petrol. Bol., 8: 389-488. Burkart, Burke, 1994. Northern Central America. In: S.K. Donvan and T.A. Jackson (Editors), Caribbean Geology: An Introduction. U.W.I. Publisher's Association, Kingston, Jamaica, pp. 265-284. Carpenter, R.H., 1954. Geology and ore deposits of the Rosario mining district and the San Juancito Mountains, Honduras, Central America. Geol. Soc. Am. Bull., 65: 23-38. Conkin, J.E. and Conkin, B.M., 1958. Revision of the genus Nummoloculina and emendation of Nummoloculina heimi Bonet. Micropaleontology, 4: 149-150. Coogan, A.H., 1973. Nuevos rudistas del Albian y Cenomaniano de Mexic6 y del sur de Texas. Rev. Inst. Mex. Pet. 5:51-82. Coogan, A.H., 1977. Early and middle Cretaceous Hippuritacea (rudists) of the Gulf Coast. In: D.G. Bebout and R.G. Loucks (Editors), Cretaceous Carbonates of Texas and Mexico. Bur. Econ. Geol., Univ. Texas Austin, Rept. Invest., 89: 32-70. Cushman, J.A. and Applin, E.R., 1947. Two new species of Lower Cretaceous foraminifera from Florida. Cushman Lab. Foraminiferal Res. Contrib., 23: 29-30. Dengo, G., 1985. Mid America: tectonic setting for the Pacific margin from southern Mexico to northwestern Colombia. In: A.E.M. Nairn and EG. Stehli (Editors), The Ocean Basins and Margins. Plenum Press, New York, 7, pp. 123-180. Dieni, I. and Turngek, D., 1979. Parkeria sphaerica Carter, 1877 (Hydrozoan) in the Vraconian (Lower Cretaceous) of Orosei (Sardinia). Boll. Soc. Paleontol. Ital., 18: 200-206. Donnelly, T.W., Horne, G.S., Finch, R.C. and L6pez-Ramos, E., 1990. Northern Central America: The Maya and Chortfs blocks. In: G. Dengo and J.E. Case (Editors), The Caribbean
Region. The Geology of North America, Vol. H, Geological Society of America, pp. 37-76. Emmet, P.A., 1983a. Mapa geol6gico de Honduras, cuadningulo de Agalteca, escala 1:50.000. Tegucigalpa, D.C., Instituto Geognifico Nacional, 1 sheet. Emmet, P.A., 1983b. Geology of the Agalteca Quadrangle, Honduras, Central America. Unpubl. MS thesis, University of Texas at Austin, 210 pp. Emmet, P.A., White, R.J. and Curry, R.P., 1992. Measured stratigraphic sections, biostratigraphy and sequence stratigraphy of Mesozoic strata in the Agalteca Quadrangle, Honduras. Geological Society of America, South-Central Section, Abstracts with Programs, p. 10. Everett, J.R., 1970, Geology of the Comayagua quadrangle, Honduras, Central America. Unpubl. PhD. dissertation, University of Texas at Austin, 152 pp. Finch, R.C., 1972. Geology of the San Pedro Zacapa Quadrangle, Honduras, Central America. Unpubl. PhD. dissertation, University of Texas at Austin, 238 pp. Finch, R.C., 1973. Mochito number 1 exploration 1973. Unpubl. report, Rosario Resources Corp., E1 Mochito, Honduras, 28 PP. Finch, R.C., 1981. Mesozoic stratigraphy of central Honduras. Am. Assoc. Pet. Geol., Bull., 65: 1320-1333. Finch, R.C., 1985. Mapa Geol6gico de Honduras, Cuadningulo de Santa B~irbara, escala 1" 50.000. Instituto Geognifico Nacional, Tegucigalpa, D.C., Honduras, 1 sheet. Finch, R.C. and Dengo, G., 1990, NOAM-CARIB plate boundary in Guatemala: a Cretaceous suture zone reactivated as a Neogene transform fault. Geological Society of America Annual Meeting, Field Trip No. 17 Guidebook, 46 pp. Finch, R.C. and Ritchie, A.W., 1990. Mapa Geol6gico de Honduras, Cuadr~ingulo de Danlf, escala 1:50.000. Instituto Geognifico Nacional, Tegucigalpa, D.C., Honduras, 1 sheet. Fourcade, E. and Michaud, E, 1987. L'ouverture de l'Atlantique et son influence sur les peuplements des grands foraminif~res de plates-formes p6ri-oc6aniques au M6sozoique. Geodin. Acta (Paris), 1 (4/5): 247-262. Gallo, J. and Van Wagoner, J.C., 1978. Stratigraphy and facies analysis of Honduras. Exxon Production Research Co., unpubl, special report. Gordon, M.B., 1989a. The Chortfs block is a continental, pre-Mesozoic terrane. In: D.K. Lame and G. Draper (Editors), Transactions of the 12th Caribbean Geological Conference, St. Croix, U.S.V.I. Miami Geological Society, pp. 505-512. Gordon, M.B., 1989b. Mesozoic igneous rocks on the Chortfs block: implications for Caribbean reconstructions (abstr.). Eos, 70: 1342. Gordon, M.B., 1990. Strike-slip faulting and basin formation at the Guayape fault-Valle de Catacamas intersection, Honduras, Central America. Ph.D. dissertation, University of Texas at Austin, 260 pp. Gordon, M.B., 1992. Northern Central America: The Chortfs block. In: G.E.G. Westermann (Editor), Jurassic of the Circum-Pacific Region. World and Regional Geology 3, Cambridge University Press, New York, pp. 93-121. Gordon, M.B. and Gose, W.A., 1989. The Chortfs block: a raft of Mesozoic sediments and Cenozoic volcanics on a solid foundation. Geol. Soc. Am. Abstr. Prog., 21: 12. Gordon, M.B. and Muehlberger, W.R., 1994. Rotation of the Chortfs block causes dextral slip on the Guayape fault. Tectonics, 13: 858-872. Gordon, M.B. and Young, K., 1993. Bathonian and Valanginian fossils from Honduras. In: S. Elmi, C. Mangold and Y. Alm6ras (Editors), 3~me Symposum International: C6phalopodes Actuels et Fossiles, Symposium F. Roman. G6obios, M6m. Sp6c., 15: 175-179.
CRETACEOUS C A R B O N A T E B I O S T R A T I G R A P H Y AND E N V I R O N M E N T S IN H O N D U R A S Gose, W.A., 1985a. Paleomagnetic results from Honduras and their bearing on Caribbean tectonics: Tectonics, 4: 565-585. Gose, W.A., 1985b. Caribbean tectonics from a paleomagnetic perspective. In: EG. Stehli and D. Webb (Editors), The Great American Biotic Interchange. Plenum Press, New York, pp. 285-301. Gose, W.A. and Finch, R.C., 1987. Magnetostratigraphic studies of Cretaceous rocks in Central America. Actas Fac. Ciencias Tierra U.A.N.L. Linares, E1 Cret~icico de M6xico y Am6rica Central, Resumenes, 2, pp. 233-241. Gose, W.A. and Finch, R.C., 1992. Stratigraphic implications of palaeomagnetic data from Honoduras. Geophys. J. Int., 108: 855-864. Hamoui, M. and Saint-Marc, E, 1970. Microfaunes et microfacies du C6nomanien du Proche-Orient. Bull. Cen. Rech. Pau-SNPA, Pau, 4 (2): 257-352. Haq, B.U., Hardenbol, J. and Vail, ER., 1988. Mesozoic and Cenozoic chronostratigrapy and cycles of sea-level change. Soc. Econ. Paleontol. Mineral. Spec. Publ., 42: 71-108. Horne, G.S., Atwood, M.G. and King, A.E, 1974. Stratigraphy, sedimentology and paleoenvironment of Esqufas Formation in Honduras. Am. Assoc. Pet. Geol., Bull., 58: 176-188. Hottinger, L., Drobne, K. and Caus, E., 1989. Late Cretaceous, larger, complex miliolids (Foraminifera) endemic in the Pyrenean faunal province. Facies, 21 : 99-134. Kozuch, M.J., 1989. Mapa Geol6gico de Honduras, Cuadr~ingulo de San Francisco de Becerra, escala 1:50.000. Instituto Geogr~fico Nacional, Tegucigalpa, D.C., 1 sheet. Kozuch, M.J., 1991. Mapa Geol6gico de Honduras, Segunda Edici6n, escala 1:50.000. Instituto Geogr~ifico Nacional, Tegucigalpa, D.C., 3 sheets. Lozej, G.E, 1976. Stratigraphy and petrography of the Mochito limestone, central Honduras (preliminary report). Unpubl. report, Rosario Resources Corp., E1 Mochito, 22 pp. Maync, W., 1953. Pseudocyclammina hedbergi n. sp. from the Urgo-Aptian and Albian of Venezuela. Cushman Found. Foraminiferal Res. Contrib., 4:101. Mills, R.A., 1959. Habr~ petr61eo en Honduras?. Pet. Interam., 17 (5): 39-44. Mills, R.A. and Barton, R., 1996. Geology of the Ahuas area in the Mosquitia Basin of Honduras: preliminary report. Am. Assoc. Pet. Geol., Bull., 80:1627-1640. Mills, R.A. and Hugh, K.E., 1974. Reconnaissance geologic map of Mosquitia region, Honduras and Nicaragua Caribbean coast. Am. Assoc. Pet. Geol., Bull, 58: 189-207. Mills, R.A., Hugh, K.E., Feray, D.E. and Swolfs, H.C., 1967. Mesozoic stratigraphy of Honduras. Am. Assoc. Pet. Geol., Bull., 51:1711-1786. Pia, J., 1936. Calcareous green algae from the Upper Cretaceous
165
of Tripoli (North Africa). J. Paleontol., 10: 3-13. Pindell, J.J. and Barrett, S.E, 1990. Geological evolution of the Caribbean region; a plate tectonic reconstruction. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Vol. H, Geological Society of America, pp. 405-432. Raineri, R., 1922. Alghi sifonee fossili della Libia. Soc. Ital. Sci. Nat., Milano, Atti, 61: 72. Ritchie, A.W. and Finch, R.C., 1985. Widespread Jurassic strata on the Chortfs block of the Caribbean plate. Geol. Soc. Am., Abstr. Prog., 17:700-701. Roberts, R.J. and Irving, E.M., 1957. Mineral deposits of Central America. U.S. Geol. Surv. Bull., 1034, 205 pp. Rogers, R.D., 1994. Preliminary stratigraphy and structure along the Rfo Patuca and Rfo Wampd, La Mosquitia, Honduras. Geol. Soc. Am., Abstr. Prog., 26 (7): A-247. Rogers, R.D., 1995. Geology along the Rfo Patuca and Rfo Wamp6, La Mosquitia, Honduras. Open-file report, Instituto Geogr~ifico Nacional, Tegucigalpa, D.C., 21 pp. Schroeder, R. and Neumann, M., 1985. Les grands foraminif~res du Cr6tac6 moyen de la r6gion M6diterran6nne. Geobios, M6m. Sp6c., 7, 161 pp., 68 pl. Scott, R.W., 1990. Models and stratigraphy of Mid-Cretaceous reef communities, Gulf of Mexico. Soc. Econ. Paleontol. Mineral., Concepts Sedimentol. Paleontol., Vol. 2, 102 pp. Scott, R.W. and Gonzales-Leon, C., 1991. Paleontology and biostratigraphy of Cretaceous rocks, Lampazos area, Sonora, Mexico. Geol. Soc. Am., Spec. Pap., 254:51-67. Scott, R.W., Frost, S.H. and Shaffer, B.L., 1988. Early Cretaceous sea level curves, Gulf Coast and southeastern Arabia. Soc. Econ. Paleontol. Mineral., Special Publ., 42: 275-284. Simonson, B.M., 1977. Geology of the E1 Porvenir quadrangle, Honduras, Central America. Open-file report, Instituto Geogr~fico Nacional, Tegucigalpa, D.C., 84 pp. Simonson, B.M., 1981. Mapa geol6gico de Honduras, cuadrangle de E1 Porvenir, escala 1:50.000. Instituto Geogr~fico Nacional, Tegucigalpa, D.C., 1 sheet. Southernwood, R., 1986. Late Cretaceous limestone clast conglomerates of Honduras. Unpubl. MS thesis, University of Texas at Dallas, 299 pp. Weaver, C.E., 1942. A general summary of the Mesozoic of South America and Central America. Proc. 8th Am. Sci. Congr., Geol. Sci., 4:179-180. Weiland, T.J., Suayah, I.B. and Finch, R.C., 1992. Petrologic, stratigraphic and tectonic significance of Mesozoic volcanic rocks in the Rfo Wampd area, Eastern Honduras. J. S. Am. Earth Sci., 6: 309-325. Williams, H. and McBirney, A.R., 1969. Volcanic history of Honduras. Univ. Calif. Publ. Geol. Sci., 85, 101 pp.
Chapter 7
Jurassic-Cenozoic Structural Evolution of the Southern Great B ahama Bank
JOSE L. MASAFERRO and GREGOR E EBERLI
Multichannel seismic reflection lines from the southern Great Bahama Bank (GBB), with a complete record down to 5 s (two way travel time), reveal the nature of basement and the evolution of the bank, which was strongly influenced by tectonic activity. The reflection seismic profiles display a fragmented internal anatomy of the bank that is tectonically controlled by deep basement faults of different ages. Three distinct episodes of deformation have shaped the architecture of the southern GBB. During the first episode high-amplitude reflections overlaying the acoustic basement are displaced by faults creating a fault-bounded topographic relief. The seismic facies of the basement with faint continuous, horizontal internal reflections overlain by a continuous high-amplitude reflection horizon is reminiscent of continental or transitional crust with a sedimentary cover. The episode of extensional tectonism that affected these two seismic facies probably corresponds to the rifting phase in the Jurassic. The GBB established on this faulted crust but subsequent growth of the bank immediately leveled the topographic relief. During the growth of the bank in the Cretaceous about 5 km of shallow-water platform carbonates were accommodated by passive-margin subsidence. The second deformational episode occurred probably in mid-Cretaceous and was characterized by the reactivation of some pre-existing structures which first segmented GBB. In southern GBB it created a WNW-ESE-trending margin towards the Tongue of the Ocean. We speculate that this tectonic phase is a consequence of the reorientation of the stress regime that caused the plate reorganization in the Caribbean realm. The third deformational episode occurred during the Late Cretaceous-Middle Eocene when Cuba collided with the southern edge of the Bahamas platform. As a result of the collision, the bank was dissected by WNW-striking oblique-slip faults forming long, narrow, symmetric depressions within the bank. The deepest depressions were produced by a long WNW-ESE straight master fault zone consisting of oblique-slip faults with both normal and left-lateral strike-slip components of displacement. The faults offset the entire bank and diverged upwards in convergent fault slices, forming negative flower structures. Activity of most faults ceased with the end of the collision in the Late Eocene. The subsequent infilling by the highly productive carbonate environment healed the depressions masking the scars produced by tectonism. Since the late Cenozoic the southern GBB is mostly a flat-topped aggrading carbonate bank.
INTRODUCTION
The northwestern portion of the Bahamas archipelago has generally been considered as a carbonate platform on a passive continental margin that was not significantly affected by tectonism. Some authors have, however, reported the existence of both compressional structures close to the Cuban orogen (Austin, 1983; Ball et al., 1985) and faulting along the Great Abaco fracture zone, in the Providence Channel and in the Tongue of the Ocean (Mullins and Lynts, 1977; Mullins and Van Buren, 1981; Sheridan et al., 1981; Ladd and Sheridan, 1987). Although most scientists have recognized that a rift topography underlies the Bahamas archipelago, a controversy has emerged around the question of how the Early Jurassic rift structures influence the present-day bank-
trough configuration of the Bahama banks. Some scientists have argued that the modern bank-channel configuration is the result of an inherited Early Jurassic rift pattern in which the intraplatform seaways are remnants of the grabens, while the platforms grow on ancient horsts (Mullins and Lynts, 1977). Others have speculated that the Jurassic structures were healed by Early Cretaceous platform growth and are buried below a Cretaceous megabank. They view the modern pattern of platforms as the result of a partial drowning of the megabank in the mid-Cretaceous caused by a worldwide crisis in carbonate and reef growth where only isolated platforms survived (Bryant et al., 1969; Paulus, 1972; Meyerhoff and Hatten, 1974; Sheridan, 1974; Schlager and Ginsburg, 1981). With the first multichannel seismic reflection data from the top of the Great Bahama Bank (GBB) it be-
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 167-193. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
168 came obvious that the modern platform configuration is not the result of one or the other event alone, but is controlled by episodic tectonic segmentation and coalescence of the banks during their growth (Eberli and Ginsburg, 1987, 1989). Although Eberli and Ginsburg (1987, 1989) recognized the role of tectonic deformation in the segmentation of the GBB, the limited depth control of their seismic data did not provide information about possible deeper structures. For this study on the southern GBB seismic reflection data that extend to a depth of 5 s twoway travel time (twtt) are available, allowing us to evaluate structural deformation of the bank since its inception in the Jurassic (Fig. 1A). We develop a structural interpretation based on recognition of structural styles from seismic reflection profiles in the southern GBB (Fig. 1B). In particular, we investigate the structural response of the southern GBB to major tectonic changes. The purpose of this paper is to document how the carbonate platform has preserved a record of the major tectonic events that have affected the southern portion of the Bahamas archipelago since the breakup of Pangea.
Database The data for this study consist of two seismic reflection datasets and a set of gravity data from the same area (Fig. 1). A grid of 1300 km of unmigrated multichannel seismic profiles was acquired with air guns with variable volumes (1600-2600 cub. in.) at intervals of 25 m. 400 km of migrated seismic profiles were acquired with an average charge size of 520 cubic inch at intervals of 100 m. These data were provided by Texaco, Inc. In addition, 180 km of migrated multichannel seismic profiles were provided by Pecten International. The spacing in the grid varies between 9 km in the northern portion and 5 km in the south. The seismic data are not connected to a well. Thus, age assignments to the main reflections in the upper part of the seismic reflection profiles were extrapolated from the Doubloon Saxon 1 well which is located 80 km to the east of our seismic grid (Walles, 1993). Correlation of the deeper part of the profiles was based on similarities of reflection character with seismic reflection data from the available literature (Ladd and Sheridan, 1987; Schlager et al., 1988; Sheridan et al., 1988). Time to depth conversion of seismic profiles was performed using ProMAX TM processing software. Gravity data from the study area, also released by Texaco Inc., provided additional evidence for some of our interpretations.
Regional setting and paleogeographic evolution The Bahamas carbonate platform was probably established in the Late Jurassic on a thinned crust
J.L. MASAFERRO and G.E EBERLI during the break-up of Pangea. During the break-up, the Bahama-Cuban area was a transform zone linking the young Gulf of Mexico and the Atlantic Ocean (Klitgord et al., 1984; Dillon et al., 1987; Ross and Scotese, 1988; Sheridan et al., 1988; Buffler and Thomas, 1994). Sedimentation in this passive-margin setting was characterized by the deposition of volcaniclastic and terrigenous material (Meyerhoff and Hatten, 1974; Mullins and Lynts, 1977). In the Late Jurassic, the southern edge of the Bahamas archipelago and northern Cuba were part of a large evaporite basin (Lewis, 1990; Walles, 1993). Carbonate sedimentation began with shallow-water carbonates that alternated with evaporites, indicating a restricted marine environment, and on localized basement highs. Shallow-water conditions persisted throughout the Early Cretaceous in some areas of the Bahamas archipelago, although deepwater sediments have been described in the Cayo Coco area in northern Cuba (Roque-Marrero and Iturralde-Vinent, 1987; Hempton and Barros, 1993), and in the northeast Providence Channel (Austin et al., 1988), and south of San Salvador (Schlager et al., 1984), suggesting the existence of deep re-entrants in the Bahamas carbonate province at that time. Seismic reflection data in the Straits of Florida also suggest the existence of such a re-entrant between GBB and Cay Sal Bank (Masaferro and Eberli, 1995). The Mid-Cretaceous was a time of change in plate motions and global plate re-arrangements (Ross and Scotese, 1988). In the Pacific, intraplate volcanism caused an increase in volume of oceanic plateaus resulting in an increase of global spreading rates and a relative rise of sea level (Larson, 1991; Mullins et al., 1992; Vaughan, 1995). In the Caribbean region, a change in the polarity of the subduction zone caused the insertion of the Farallon plate into the proto-Caribbean realm carrying the Greater Antilles towards the northwest (Burke et al., 1984; Ross and Scotese, 1988; Pindell and Barrett, 1990). Movement of the Greater Antilles arc and consumption of the proto-Caribbean crust continued up to the Late Cretaceous-Paleocene when the arc collided with the southern edge of the Bahamas platform (Burke, 1988; Ross and Scotese, 1988), forming the Cuban orogenic belt. The Cuban arc, which consists of obducted ophiolites, imbricated platform carbonates, melanges and arc-derived plutonic and volcanic rocks, was thrust onto the passive margin of the Bahamas archipelago in a northward-verging fold and thrust system (Gealey, 1980). The timing and the polarity of subduction is still controversial and several interpretative models have been proposed (Goodell and Garman, 1969; Malfait and Dinkelman, 1972; Mattson, 1973, 1974, 1979; Gealey, 1980; Pindell and Dewey, 1982; Pindell, 1985; Dillon et al., 1987;
CENOZOIC STRUCTURAL EVOLUTION OF THE SOUTHERN GREAT BAHAMA BANK
169
Fig. 1. (A) Location map of the seismic profiles, industry boreholes and core borings across the Great Bahama Bank. A - A t is a trace of cross-section in Fig. 20. (B) Multichannel seismic data used in this study. Bold line segments indicate parts of the seismic profiles shown in this paper. Burke, 1988; Draper and Barros, 1994; IturraldeVinent, 1994). With the collision and loading of Cuba onto the margin the Bahamas area became a foreland basin (Ball et al., 1985; Walles, 1993; Denny et al., 1994). The collision ceased in the middle Tertiary, as indicated by the stratigraphic relationships in Cuba, where Middle Eocene and younger sediments lie
unconformably on the highly deformed collisional deposits (Iturralde-Vinent, 1972, 1975, 1988; Pardo, 1975; Pindell and Barrett, 1990; Lewis, 1990). From this time onward the NW Bahamas-Cuban area has been tectonically stable and the foreland basin became again part of the passive continental margin of the North American plate (Fig. 2). The platforms of the Bahamas archipelago were influenced by these
170
J.L. MASAFERRO and G.R EBERLI
Fig. 2. Present-day plate boundary faults of the northern Caribbean plate modified from Mann et al. (1995). Shaded area indicates study area.
tectonic changes but succeeded for the most part to exist. Subsequent growth of the platforms and the maintenance of their flat-topped geometry hide the structural features, which are revealed by the multichannel seismic reflection data we present below.
SEISMIC REFLECTION DATA
Seismic facies description The seismic data contain five distinct seismic facies defined by the acoustic character of the seismic reflections. The first seismic facies (S.0) is the basal seismic facies occurring below 3.3 s (twtt); it consists of horizontal discontinuous, medium-amplitude to chaotic reflections. It is overlain by seismic facies (S.1), which is a group of strong regional highamplitude seismic reflections that extend regionally and appear in the lower half of the seismic profiles from 2.8 to 3.3 s (twtt) (Fig. 3). Displacement of the S.1 horizons displays the geometry of the deep structure. The third seismic facies (S.2) consists of an interval of approximately 2 s (twtt) in thickness with incoherent to chaotic seismic reflections (Fig. 3). Using interval velocities, we estimate the average thickness of S.2 to be about 5000 m (Table 1).
The S.2 seismic facies is bounded at the top by a high-amplitude reflection (R.1, Fig. 3). Occasionally, this seismic package contains internally strong, more continuous reflections that can be used for correlation within the otherwise chaotic package (correlation horizons in Fig. 3). Overlying the high-amplitude reflection R.1 at approximately 0.8 s (twtt), is the fourth seismic facies that contains fairly continuous, moderate- to high-amplitude reflections, which are locally interrupted by chaotic and incoherent reflections (S.3, Fig. 3). A lateral change of the acoustic character within this unit indicates lateral facies changes. The most common change is from thin low-amplitude and transparent reflections to more continuous reflections of a higher amplitude. S.3 is overlain seismic facies S.4 at the top of the seismic reflection profiles. S.4 is transparent at the base and contains horizontally continuous high-amplitude reflections at the top.
Seismic facies interpretation The nature of the crust underneath the modem banks is unknown as no drill hole ever reached the basement. Refraction seismic data suggest a 1214-km-thick transitional crust (Sheridan, 1974). The seismic facies S.0 on the reflection seismic data gives
CENOZOIC STRUCTURAL EVOLUTION OF THE SOUTHERN GREAT B A H A M A BANK
171
Table 1 Seismic facies units identified in the southern GBB and their internal descriptions Seismic facies units
Seismic facies
Interval velocities (km/s)
Estimated ages (top of seismic facies)
Interpretation of seismic facies
S.4 S.3 S.2 S.1 S.0
Transparent low-amplitude High-amplitude continuous Chaotic High-amplitude continuous Discontinuous chaotic
1.9-2.5 3.2-3.6 4.0-5.1 5.4 5.5
Recent Middle Paleocene Late Cretaceous Middle Jurassic? Early Jurassic?
Platform carbonates Periplatform basin fill Platform carbonates Evaporites, clastics and carbonates Continental to transitional crust
evidence that the southern GBB might be located on such a transitional or extended continental crust but not oceanic crust. Seismic images of oceanic crust are characterized by chaotic internal reflections and a hummocky top (Sheridan et al., 1981; Shipley et al., 1989; Mountain and Tucholke, 1989). Such a chaotic facies is not present in the southern GBB but horizontally continuous reflections are observed below the continuous, high-amplitude reflection horizons S.1. These seismic facies suggest a continental crust overlain by a sedimentary unit, consisting of lithologies with large impedance variations. These lithologies could be alternations of volcaniclastics, evaporites and limestones. The chaotic seismic reflection signature (S.2) has been interpreted by many authors as shallow-water platform carbonates (Ball et al., 1985; Eberli and Ginsburg, 1987, 1989; Denny et al., 1994). This interpretation is corroborated by seismic modeling which indicates that the lack of contrast of acoustic impedance within the shallow-water platform carbonates is the main factor that causes the incoherent character of the reflections in the seismic sections (Anselmetti et al., 1997). Thus, the seismic facies at the top 2.7 s (twtt) of the seismic profiles probably consists exclusively of carbonates (Walles, 1993). Across reflection horizon R.1, however, occurs a drastic change in seismic facies from low-amplitude and chaotic reflections to more continuous, high-amplitude reflections (S.3). This facies is best developed in areas were depressions formed within the bank. The depressions were filled by onlapping and prograding packages (Fig. 4), while more shallow-water carbonates were deposited on the adjacent topographic highs. Within the depressions, the change in seismic facies probably reflects the transition from platform deposits to more basinal deposits and might be a mixture of periplatform ooze and various amounts of neritic debris and redeposited material (Eberli and Ginsburg, 1989). The continuous reflections on the adjacent highs are probably a result of alternating shallow-water and deeper-water, chalky lithologies. The prograding packages of S.3 indicate margin progradation into the depressions. No predominant direction of progradation is observed, and although westward progradation is seen,
progradation from the south and southwest is at least as common (Fig. 4). This progradation pattern differs markedly from the prevailing east-to-west progradation in the northeastern Great Bahama Bank (Hine and Neumann, 1977; Eberli and Ginsburg, 1987). The top portions of all the seismic profiles are characterized by horizontal reflections (S.4), indicating homogeneous platform aggradation in the interior of southern GBB. Shallow core borings from other parts of the GBB penetrated this seismic facies and consisted of shallow-water carbonates punctuated by exposure horizons (Beach and Ginsburg, 1980).
Age assignments to seismic horizons In order to give relative ages to the deformational events and to better describe the structural features, we have divided the section into four stratigraphic units based on both the character of the seismic facies and the geometry of the reflections that characterize each unit: (1) a high-amplitude seismic unit (S.1), (2) a thick, chaotic, intermediate seismic unit we interpret as platform carbonate facies (S.2), (3) an upper, more continuous seismic unit that corresponds to the filling of the depressions (S.3), and (4) an overlying upper relatively undeformed seismic unit (S.4). Determination of geologic ages and lithologies of the seismic facies are difficult because of limited well control in the area. No well is situated in the area of our seismic reflection data. As a result, our age assignments are based on the correlation of the depth of the seismic reflections to depths of the ages in the nearby Doubloon Saxon 1 well and similarities in seismic character with reflections described in the literature (Sheridan et al., 1981, 1988; Ladd and Sheridan, 1987; Walles, 1993). The S.1 seismic facies in the southern GBB is similar in seismic character and position with respect to the overlying seismic facies to a group of reflections observed under the Tongue of the Ocean, Exuma Sound and northeast Providence Channel (Ladd and Sheridan, 1987; Sheridan et al., 1988). The tops of these high-amplitude reflections were correlated with Middle Jurassic limestones, dolomites and evaporites (Sheridan et al., 1981, 1988). We have used line 337 of Ladd and Sheridan (1987),
172
J.L. M A S A F E R R O and G.R EBERLI II
II
~r.,5 ~g ~A O
O
. ,..,,~ ~99
O
O
U=
.,..~
.,,a
. ,...~ ,,.o cA =
9
9.= .~ .,..~ . ,...,
II
,.~
=
O
..,
(D ....~
9
,..,
U::
=
O
O
" ~ ~
.,.~ ,~
~ 0
~ ~
~
~.~ e'~
9,'-~
~.~
~',~
~
~
~
=
N
~ i
9
"~
9
9, -
~
~
9
9
CENOZOIC STRUCTURAL EVOLUTION OF THE SOUTHERN GREAT BAHAMA BANK
173
Fig. 4. Uninterpreted and interpreted unmigrated seismic profile showing the first segmentation of the southern GBB. Direction of progradation is towards the NE. See Figs. 1B and 4 for location.
174
J.L. MASAFERRO and G.E EB ERLI
located in the southern part of the Tongue of the Ocean, to extrapolate these reflections into our seismic grid, and assign a Middle Jurassic age to the top of seismic facies S. 1. As the Tongue of the Ocean is a basin and our grid is within the shallow-water platform of the GBB, it is not possible to trace younger reflections from southern Tongue of the Ocean into our seismic reflection grid. Therefore, we compare the depth of three reflections (R.1, R.2 and R.3) of a depthconverted seismic section to ages in the Doubloon Saxon 1 well (Fig. 5). The Doubloon Saxon 1 is a deep (6631 m TD) exploration well located 80 km west of our seismic reflection grid (Fig. 1A) (Walles, 1993). From the depth section, we correlated the age of reflection R.1 with the top of the Upper Cretaceous section in the Doubloon Saxon 1 (Fig. 5). Reflection R.2 can be correlated with the top of the middle Paleocene section and reflection R.3 is just above the top of the Middle Eocene and could be as young as Early Oligocene.
Internal anatomy of the southern GBB The seismic reflection profiles reveal the complicated internal structure of the Great Bahama Bank south of Tongue of the Ocean. An isochron map SSW
Doubloon Saxon 1
(Fig. 6) of reflection R. 1, which marks the base of intra-platform depressions, shows five platforms separated by depressions of different size and orientation. A large, asymmetric depression, termed here Cochinos Sound (Fig. 6), was a wide N W - S E depression south of the modem Tongue of the Ocean from which it was separated by a narrow platform ridge. This narrow WNW-ESE-trending margin formed a barrier between the Tongue of the Ocean and the Cochinos Sound (Figs. 4 and 5). To the south, the Cochinos Sound borders the northern platform, and turns to the south into a more N-S trend to form a narrow N-S depression on the eastern side of the eastern platform (Fig. 6). We speculate that this depression was probably connected with the Old Bahama Channel in the south. Using average interval velocities the depth of the Cochinos Sound is estimated to be about 300-500 m and the depth of the eastern depression approximately 700 m In the area southwest of the Cochinos Sound, where the coverage by seismic lines is extensive, the platform is dissected by narrow, more symmetric depressions (Figs. 3 and 6). The northwestern depression extends N W - S E for about 15 km. Toward the east, this long depression turns into a more E - W direction (Fig. 6). The depth of the northwestern depression is estimated at about 500 m in its axial valNNE
Depth-converted seismic profile from Figure 3 (see location map)
Fig. 5. Age correlation of the Doubloon Saxon 1 well (Walles, 1993) with a depth-converted seismic profile. See text for discussion.
CENOZOIC STRUCTURAL EVOLUTION OF THE SOUTHERN GREAT BAHAMA BANK
175
Fig. 6. Isochron map on top of reflection R.1 (Late Cretaceous) showing the orientation of both depressions and adjacent platforms. Contour interval is 50 ms. ley. Farther to the south, at the southern limit of our seismic reflection grid, a third depression (southern depression, Fig. 6) dissects the southern portion of the platform. It is possible that this depression was an ancient northern boundary of the Old Bahama Channel.
GRAVITY DATA The gravity data (Web Fig. 7.1, available at http://www.elsevier.nl/locate/caribas/), provided by Texaco Inc., consist of an uncorrected Bouguer anomaly map. To remove the regional effect that obscures the local anomalies, we corrected the data. The regional gradient is a component of the gravity anomaly that has longer wavelength and is thought to be caused by density contrasts of deep-seated masses located at any depth. Removal of this regional effect from the Bouguer gravity map yields the residual anomalies which represent the anomalies created by a density contrast at or near the surface. Different methods can be used to remove the regional gradient from the Bouguer gravity map. We applied a two-dimensional surface fitting technique (Nettleton, 1971) to fifteen N-S cross-sections of the gravity field, using a polynomial of second order to smooth
the observed gravity and to estimate the regional gradient. A regional gradient was estimated from each of the profiles and residual gravity profiles were obtained by subtracting the estimated regional from the observed gravity at all points along the profiles. The residual gravity map (Fig. 7) displays the enhanced local anomalies caused by the density contrast between high and less-dense sediments. The areas of negative values coincide approximately with the depressions while positive values follow the adjacent shallow-water areas. Thus, the residual gravity anomalies support the seismic facies interpretation where the chaotic, incoherent seismic reflections are thought to represent shallow-water platform carbonates (higher density), and the more continuous, highamplitude seismic reflections form the less-dense sediments in the depressions.
STRUCTURAL ANALYSIS The structural style of the southern GBB is dominated mainly by faults of different ages accompanied by localized compressional structures. Each of the four stratigraphic units (S.1-S.4) contains a characteristic set of structural features that formed during
176
J.L. MASAFERRO and G.E EBERLI
Fig. 7. Residual gravity map obtained subtracting the estimated regional from the observed gravity (Web Fig. 7.1). The minimumvalues coincide with depressions and the maximumvalues with adjacent platforms (shaded areas, compare with Fig. 6). Contour interval is 0.7 mGal. different tectonic stages in the evolution of the carbonate platform.
Fault geometries Seismic reflection profiles through the southern GBB image three generations of faults that disrupt the carbonate platform at different levels. The oldest set of faults has a planar to more curved profile and displace the top of the high-amplitude reflection horizon of S.1. Fig. 8 shows a migrated seismic profile that crosses two of these faults, which displace the S.1 reflections from 2.8 s to 3 s (twtt) and from 3 s to 3.2 s (twtt) which translates into approximately 600 m of vertical displacement. These faults separate extensional fault blocks. Fig. 9 is a structure map that shows the distribution and orientation of the faults that affected the S.1 seismic facies. The old
faults have a predominantly NNE strike and most of them dip steeply to the west. The lack of continuous reflections within the overlying chaotic unit S.2 does not show whether these faults propagated upward into this unit. However, some migrated seismic lines with more continuous reflections in the chaotic facies show propagation of some faults into this unit where they seem to die out. The second generation of faults observed in the area dissects S.2 seismic facies. An example is shown in Fig. 10 from the northern end of our seismic reflection grid, where two steep normal faults form a WNW-ESE-striking narrow asymmetric trough. The southern fault disrupts the S.2 unit at approximately 1.7 s (twtt) based on the southward inclined reflection terminations of unit S.3 (shaded area, Fig. 10). The dipping reflections thicken to the southwest into the fault plane. The northern fault
C E N O Z O I C S T R U C T U R A L E V O L U T I O N OF THE S O U T H E R N GREAT B A H A M A B A N K
177
Fig. 8. Uninterpreted and interpreted migrated seismic profile showing old faults affecting S.1 seismic facies (Middle Jurassic?). See text for discussion. See Fig. 1B for location.
Fig. 9. Isochron map contoured on top of the S.1 seismic facies. Inferred pre-Cretaceous faults have NNE-SSW orientation. Contour interval is 50 ms.
17 8
J.L. MASAFERRO and G.E EBERLI
Fig. 10. Part of migrated seismic profile and interpreted line drawing showing a narrow, fault-bounded trough. The SW-dippingreflections within the shaded area of S.3 thicken towards the fault plane indicating that the deposition of the strata was coeval with fault activity. Subsequent filling of the trough, however, is by NE-prograding sequences. See Fig. 1B for location. displays contraction in the upper part whereas in the deeper part the same fault shows normal separation. A third generation of faults is widespread and displaces strata of Paleocene and younger age. High-angle faults of this generation have mostly normal separation and propagate towards the surface, disrupting almost the entire section of S.3 (Figs. 11-13).
In map view, this main fault system consists of three fault zones (fault zones A, B and C, Fig. 14) characterized by their linearity, by the presence of localized contractional structures such as folds and reverse faults coexisting with the normal dip-slip faulting, and by an uplifted zone that blocks the main fault corridor. The major W N W - E S E fault zone A is about 10 km in
CENOZOIC STRUCTURAL EVOLUTION OF THE SOUTHERN GREAT BAHAMA BANK
179
Fig. 11. Uninterpreted and interpreted part of unmigrated seismic profile. Faults dip steeply and have both normal and reverse separation. Steep faults seem to offset S. 1, while reverse faults continue to younger strata. Box is enlarged in Fig. 12. See Fig. 1B for location.
width and 65 km in length and displays a sharp bend in its northwest portion where contractional structures such as reverse faults and folds are concentrated (Fig. 14). To the east, fault zone A merges with the E W-trending fault zone B to form a braided geometry. This combined fault zone widens southeastward and terminates at an uplifted and folded structure. A third fault zone, C, is located in the southernmost extension of our seismic grid and is not connected to the other two fault zones (Fig. 14).
In cross-section, the entire fault system is characterized by steep faults that show both normal and reverse separation where the S.1 unit is involved in the deformation. The narrow area where the S.1 unit is affected by splay-upward structures is generally devoid of significant reflections, producing a sharp change in the continuity of the reflections. Four seismic reflection profiles (Figs. 11-13 and 15) cross fault zone A, and cut the fault system at a fight angle illustrating the marked lateral changes in fault geom-
180
J.L. MASAFERRO and G.E EBERLI
Fig. 12. Uninterpreted and interpreted migrated upper part of seismic profile shown in Fig. 11, showing normal and reverse fault separation, some contractional features and the creation of a depression within the fault system. These are characteristics of a negative flower structure. Note how the infilling sequences heal the structural deformation and how the seismic reflections are horizontal above R.2. etry within a distance of 25 km (Fig. 14). Figs. 12 and 13 show faults that have an almost symmetrical upward-branching form showing reverse faults and folding on the northern side. These fault splays
converge at depth. North of this branching-upward structure a fault with normal separation (called here D), displaces both seismic unit S.1 (with throws up to 1200 m) and probably the lower part of the
CENOZOIC STRUCTURAL EVOLUTION OF THE SOUTHERN GREAT BAHAMA BANK
181
Fig. 13. Uninterpreted and interpreted part of seismic profile shown in Fig. 11. Some faults displace S.1 seismic facies. Fault D is an old normal dip-slip fault with approximately 1000 m of displacement. See Fig. 1B for location.
chaotic unit S.2. Fault D has a slightly listric profile at depth with a gentle rollover in the hanging wall (Fig. 13). Passing from the eastern side to the central area of the fault zone A, the geometry changes from an upward-branching form to a single strand fault where the strata thicken toward its plane (Fig. 15). The second fault zone B strikes east-west and probably continues to the southwest out of our seis-
mic grid (Figs. 8 and 16). The geometry of this fault zone is characterized by a series of normal faults that converge at depth. In contrast to fault zone A, no contraction or reverse faults are observed. Fig. 16 illustrates the structural style of faults in zone B and the style of the infilling sequences. The development of aggrading and prograding sequences within this fault-controlled depression seems to be
182
J.L. MASAFERRO and G.E EBERLI
Fig. 14. Tectonic map of structures mapped on the R.1 reflection and position of seismic profiles. Three fault zones can be identified. Fault zones A and B merge towards the east and terminate at an area of tectonic uplift. Fault zone C parallels the two other zones further in the south. Shaded areas indicate extension. related to the temporal evolution of the fault system. Initial and in some cases syntectonic deposition is aggradational (Fig. 16). As soon as the fault system becomes inactive prograding sequences start to fill the depression from south to north (Fig. 16). Progradation fills the deepest portion of the depression and shifts the basin axis to the north. Before the infilling is completed another tectonic pulse deepens again the basin. Onlapping sequences fill this young embayment. Finally, shallow-water conditions reestablished over the entire bank, masking the former depression and infilling sequences. The third fault zone, C (Figs. 14 and 17), has the following geometric characteristics: a narrow fault zone, convergent steeply dipping faults showing normal separation at depth, and coeval normal and reverse faults with folding in the upper unit. On the migrated seismic profile (Fig. 17) that crosses this fault zone, the base of unit S.3 is down-faulted and tilted towards the southernmost fault creating a steep southwestward-dipping ramp. The sediments start to fill in the space created by extensional faults progressively onlapping the ramp towards the northeast. After deposition of the onlapping sequences the strata in the center of the depression are locally folded producing a symmetrical anticline. The southern flank of the an-
ticline is down faulted (fault B) almost coeval with the local convergent phase. A local angular unconformity separates the folded sequences from the flat-lying sequences of unit S.4 above (Fig. 17). Folds Two types of folds are recognized based on their cross-sectional geometry: (1) folds related to reverse faults that lie within the main fault system, and are mostly parallel and subparallel to the displacement zone (Figs. 12 and 13), and (2) localized anticlines and monoclinal flexures associated with vertical to subvertical normal faults (Fig. 18). These flexures are located outside and at the southeastern end of the main fault zone (Fig. 14). Folds that are associated with the upward-branching fault system lie within or immediately adjacent to the fault zone (Fig. 14). The crestal traces of the folds trend subparallel and oblique to the main fault zone direction.
STRUCTURAL INTERPRETATION The seismic reflection data show two main structural styles in the southern GBB: a predominantly
CENOZOIC STRUCTURAL EVOLUTION OF THE SOUTHERN GREAT BAHAMA BANK
183
Fig. 15. Uninterpreted and interpreted part of seismic profile illustrating a change in fault geometry across fault zone A. See Figs. 1B and 14 for location. extensional style with NNE-SSW-striking normal faults, and a late sinistral divergent strike-slip style within a main WNW-ESE-striking deformation zone. The deformation is dominated by faults that show mostly normal separation on the seismic profiles.
Extensional phase Due to the limited amount of dip seismic reflection profiles that cross the extensional fault system it is difficult to infer the regional extension direction.
Most of the extensional faults are confined to the S. 1 unit that we interpret to be pre-Early Cretaceous in age. Therefore their origin is probably related to the late stages of rifting caused by the separation between North America and Africa (Mullins and Lynts, 1977; Ross and Scotese, 1988; Buffer and Thomas, 1994). The timing of faulting is not constrained by well control, but rifting is reported to begin during the Early-Middle Jurassic (Sheridan and Osburn, 1975; Sheridan et al., 1981, 1988; Klitgord et al., 1984; Ladd and Sheridan, 1987). The extensional faults are buried by a thick pile of platform car-
Fig. 16. Uninterpreted and interpreted upper part of migrated seismic profile showing the geometry of fault zone B. Faults with normal separation form a depression that is subsequently filled by prograding sequences. See text for discussion.
t-
t~
C~
t-
CENOZOIC STRUCTURAL EVOLUTION OF THE SOUTHERN GREAT BAHAMA BANK bonates deposited during the Late Jurassic-Early Cretaceous. Some of the old faults propagate into the chaotic unit (S.2), e.g. creating a WNW-striking margin at the southern end of Tongue of the Ocean (Fig. 6), and narrow depressions within the Upper Jurassic-Lower Cretaceous shallow-water platform (Fig. 10). We interpret these propagations as related to a younger tectonic event in the mid-Cretaceous that reactivated some of the Jurassic faults.
Transtensional phase The extensive coverage of the main deformation zone by the seismic reflection grid defines both the orientation and the geometry of the structural features of a left-lateral strike-slip system with an important component of extension. The difficulty in the interpretation of areas where strike-slip is a main component of deformation from vertical sections has led different authors to establish numerous subsurface criteria to identify such areas (Harding, 1983, 1985, 1990; Gibbs, 1986, 1987). The following characteristics of structural features suggest a left-lateral strike-slip component in the deformation of the southern GBB. (1) The main deformation zone is restricted to an elongated, straight, narrow area about 8 km wide. Fig. 3 shows the width of this zone in a seismic reflection line that crosses the structure at almost a right angle. The horizontal reflections north and south of the deformation zone are tectonically undisturbed. (2) High-angle faults that converge at depth disrupt the entire section down and beyond the S.1 seismic facies. On several seismic reflection profiles (e.g., Figs. 12 and 16) the continuation of these faults is seen in the upper S.3 unit. (3) The bending geometry of the main displacement fault zone defines an area of local convergence that can be interpreted as a restraining bend (Crowell, 1974b) (Fig. 14). Folds and faults with reverse separation are concentrated in this area. The existence of these localized contractional structures allows us to infer a general sinistral sense of displacement (Fig. 14). The fault trace also bends into a more E - W orientation to the southeast, and creates a zone of extension or an incipient releasing bend (Crowell, 1974b). (4) The upward-branching geometry observed in some seismic reflection profiles (Figs. 12, 13, 16 and 17) across the deformation zone that includes both normal and contractional elements in the upper S.3 unit defines negative flower structures (Harding, 1983, 1985; Harding et al., 1985). The almost coeval development of folds and extensional faults within the negative flower structure is characteristic of strike-slip systems.
185
(5) The abrupt change in the geometry alongstrike of the main deformation zone, from an upward-branching structure to a single-strand fault (Figs. 13 and 15). (6) External folds oriented oblique and outside the main deformation zone (Fig. 18) that look like flexures or monoclinal knees in seismic sections (Harding et al., 1985) indicate local convergence associated with the displacement of the fault zone. (7) The change in both seismic reflection facies and thickness across the fault zone from continuous, high-amplitude reflections in the southwestern side to a chaotic, thicker, discontinuous reflections in the northeastern side (Fig. 12). (8) The presence of local unconformities with a limited lateral extent within the subsided areas indicates pulses of syntectonic sedimentation associated with the temporal development of the strike-slip system (Fig. 17). This structural assemblage of the southern GBB cannot be interpreted in terms of either a pure extensional nor a pure strike-slip model. Most of the deformation is dominated by faults that look extensional on the seismic data, but the coexistence of folds, reverse and normal faults is reminiscent of a strike-slip system with a component of extension. This mixture of structural styles of both extensional and associated contractional elements is characteristic for divergent wrench zones (Harding et al., 1985).
Structural model for the southern GBB Experimental models on fault reactivation by Richard and Krantz (1991) and oblique slip by Mandl (1988) support our interpretation. Richard and Krantz's models suggest that a previously faulted zone can control the deformation of a strikeslip zone, and the pre-existing basement structure controls also the location of the deformation imposed during the subsequent strike-slip episode. The set of structures that Richard and Krantz (1991) obtained from the experimental models are analogous to the structures we described in the main fault system (Fig. 19A). Mandl (1988) interpreted the formation of 'tulip'-type cross-sectional fault structures or negative flower structures (Harding, 1985) as a result of oblique slip induced by a basement fault. His models suggest that the fault pattern in the overburden depends on the predominance of the strike-slip or dip-slip component in the basement. If the strikeslip component dominates, the result is a 'tulip'-type fault structure, whereas if the strike-slip component is small the structure resembles a single graben. In either model the pattern of the overburden faults caused by a strike-slip component in the deformation of the basement is comparable with the fault pattern that characterizes our seismic sections. By analogy,
Fig. 17. Uninterpreted and interpreted migrated part of a seismic profile showing the geometry of fault zone C. Seismic profile illustrates the coeval coexistence of both extensional and contractional structures within the fault system. See text for discussion.
t-
t~
~z
9
t-"
CENOZOIC STRUCTURAL EVOLUTION OF THE SOUTHERN GREAT BAHAMA BANK
187
Fig. 18. Uninterpreted and interpreted migrated part of a seismic profile showing folding outside the main deformation zone.
the strike-slip structures seen in the southern GBB can be interpreted to be induced by a basement fault system located at depth. In particular, fault D (Fig. 19B) may be such a pre-existing dip-slip fault subsequently reactivated in a strike-slip mode creating extensional and contractional structures in the overburden similar to the structures observed in the models.
DISCUSSION
The internal configuration of the southern GBB documents the change of a passive-margin carbonate
platform into a foreland tectonic setting. Multichannel seismic reflection data across the southern GBB suggest that the bank underwent three major tectonic events during its geologic history. These three tectonic episodes can also be seen in other parts of the GBB. In the following we discuss the timing of these events and a possible regional correlation. The first deformation event was extensional associated with a transform system that operated probably through the Early-Middle Jurassic in a broad zone across the Florida-Bahamas area (Klitgord et al., 1984) (Fig. 20A). Klitgord et al. (1984) proposed a transform plate boundary located between the
188
Fig. 19. Tectonic model for the southern GBB. Basement-induced strike-slip structures. (A) Reactivation of a normal dip-slip basement fault in strike-slip mode (modified from Richard and Krantz, 1991). (B) Part of seismic profile (Fig. 13) showing similarities with model. See text for discussion.
Bahamas and Cuba fracture zones that connected the Central Atlantic and Gulf of Mexico spreading systems. The southern GBB established on a stretched transitional/continental crust controlled by a set of NW-striking strike-slip faults and NNE extensional faults created by the Jurassic transform zone (Fig. 20B). Rift-related sediments consisting of clastics, evaporites and hemipelagic carbonates were deposited soon after breakup (Sheridan et al., 1981; Ladd and Sheridan, 1987). In the northwestern GBB, this extensional event is indicated in seismic reflection profiles by two NNE-striking normal faults that probably first segmented the bank and formed the Straits of Andros (Fig. 21A). Seismic reflection profiles from the Straits of Florida and Santaren Channel suggest that part of this area was a fault-controlled deep-water re-entrant (Masaferro and Eberli, 1995). Schlager et al. (1984) also reported the existence of Early Cretaceous deep-water sediments in the northeast Providence Channel (Fig. 21A) giving evidence for an old re-entrant at this place. The second deformational episode probably occurred during the middle Cretaceous after the deposition of Lower Cretaceous shallow-water carbonate platform sediments in a post-rift subsidence phase (Fig. 20C). The middle Cretaceous was a time of major plate rearrangement and a change in the orientation of tectonic stresses affecting the Caribbean
J.L. MASAFERRO and G.E EBERLI realm (Ross and Scotese, 1988; Scotese et al., 1989). In the southern GBB we see evidence that this reorganization of the stress regime caused the reactivation of older Jurassic faults and their propagation into the Upper Jurassic-Lower Cretaceous carbonate platform (Fig. 20D). Some of these faults die out within the Upper Jurassic-Lower Cretaceous section without producing significant changes in the architecture of the platform. Others propagated through the entire section creating deep, narrow, WNW-striking troughs within the Upper Jurassic-Lower Cretaceous shallow-water carbonate platform (Fig. 20D). Further north, both Andros Bank and Bimini Bank were segmented and backstepped probably during this time as a result of extensional faulting (Fig. 21B) (Eberli and Ginsburg, 1987, 1989). In the Santaren Channel, seismic reflection profiles show that the Lower Cretaceous megabank was also segmented by deep, NW-striking extensional faults (Fig. 21B) with throws up to 1200 m (Masaferro and Eberli, 1995) creating a northeast-facing platform margin. The third deformational episode occurred during the Late Cretaceous/middle Tertiary, when the Bahamas archipelago became involved in the Caribbean-American collision (Fig. 20E). The southern GBB started to be affected by subduction beneath the Caribbean plate, which some authors postulated to be oblique under the Caribbean plate producing a left-lateral wrench system (Mullins and Sheridan, 1983; Mullins, 1984; Pindell and Barrett, 1990; Mann and Burke, 1990; Mann et al., 1995). In the southern GBB, this collision resulted in renewed movements on older Jurassic faults and the formation of new WNW-striking strike-slip zones. The pre-existing Jurassic transform faults (Klitgord et al., 1984; Ladd and Sheridan, 1987; Sheridan et al., 1988) probably played an important role in defining the newly formed strike-slip system reactivated by the Cuban convergent orogen. The structural expression of this deep wrenching in the carbonate platform was a composite of extension and strikeslip elements. This phase of left-lateral strike-slip faulting, which was primarily accommodated on three narrow fault zones, caused the fragmentation of the Cretaceous carbonate platform in the area of the southern GBB. Individual faults branched out from the deep basement showing extensional and compressional elements that defined negative flower structures. Transtension associated with these faults controlled localized subsidence and sedimentation. Sedimentary sequences developed within the subsided depressions. They show in some cases internal angular unconformities as a result of the episodic tectonic pulses associated with the activity of the strike-slip faults. During this time, further north the Straits of Andros started to be infilled by aggrading sequences followed by successive pulses of progra-
C E N O Z O I C S T R U C T U R A L E V O L U T I O N OF THE S O U T H E R N GREAT B A H A M A B A N K
189
Fig. 20. Interpretation of the tectonic evolution of the southern GBB along cross-section A-A' (Fig. 1A). A rift topography is buried by Late Jurassic-Early Cretaceous sedimentation. Mid-Late Cretaceous tectonism segments the platform during the collision of Cuba with the North American plate. In the tectonic quiet post-Eocene to Recent time the platform heals these tectonic scars again. See text for discussion.
dation (Fig. 21C) (Eberli and Ginsburg, 1989). In the Santaren Channel area, a major extensional border fault east of Cay Sal that had controlled the location of the platform margin, was inverted due to compressional forces of the Cuban-Bahamas collision (Masaferro and Eberli, 1995). Reactivation led to the development of faults with reverse separation and broad anticlines along this margin (Masaferro and Eberli, 1995). During the collision and subsequent orogeny in Cuba, the southern extension of the Bahamas plat-
form became part of a foreland basin (Fig. 20F). Soon after the collision the productive carbonate system responded hiding the depressions created by tectonism. During the Middle/Late Eocene, the convergence ceased when Cuba was detached from the Caribbean plate and was added to the North American plate (Ross and Scotese, 1988; Mann et al., 1995). In the seismic sections, the cessation of tectonic activity is well represented by a sharp, horizontal reflection (R.2) at approximately 0.4 s (twtt). Horizontal seismic reflections above this reflection
190
J.L. M A S A F E R R O and G.E E B E R L I
Fig. 21. Compilation of structural features and paleogeographic evolution of the Great Bahama Bank area. Arrows indicate main direction of progradation. Data compiled from Schlager et al. (1984), Eberli and Ginsburg (1987, 1989), Sheridan et al. (1988) and Masaferro and Eberli (1995).
record vertical growth of the carbonate platform up to the present. Some faults, however, seem to affect very young seismic reflections that might indicate recent tectonic activity. Such Cenozoic tectonic events are also documented in various places in the northwest Bahamas, e.g. in Walkers Cay (Mullins and Van Buren, 1981), Bimini Bank (Eberli and Ginsburg, 1987) and at the western side of the Straits of Florida close to the eastern edge of the former larger Cay Sal Bank (Eberli et al., 1995). The origin of these post-collisional tectonic events is uncertain, but it might imply that deformation was still occurring between the Caribbean and North American plates during the Neogene.
CONCLUSIONS
Seismic reflection profiles across the southern GBB show that the internal architecture of the bank is primarily a result of the interaction between tectonic destructive processes and platform recovery through sedimentation. Seismic data suggest that the southern GBB experienced three segmentation events. The first segmentation event was probably related to the Jurassic rifting and created a series of fault blocks. Subsequent sedimentation buried large portions of these rift structures to form the southern GBB. Probably during mid-Cretaceous time, a second tectonic event that was also extensional re-
C E N O Z O I C S T R U C T U R A L E V O L U T I O N OF THE S O U T H E R N GREAT B A H A M A B A N K
activated WNW-trending old structures as normal faults. It created in the southern GBB a NE-facing margin towards the Tongue of the Ocean and several deep water re-entrants further north. The third segmentation event, during the Late Cretaceous to middle Tertiary, was transtensional and synsedimentary. Oblique-slip faults with reverse and sinistral strike-slip components disrupted the bank, forming symmetric intraplatform depressions. Subsequently, sedimentation filled the depressions with aggrading/prograding packages burying the inactive structures. The internal structure of the GBB reveals a competition between relief-forming tectonism and masking sedimentation. Beneath the flat top of the southern GBB lies a complex history of changing geological regimes from a passive continental margin into a collisional plate boundary, the effects of which became masked by sediments of the highly productive carbonate platform.
ACKNOWLEDGEMENTS
We thank Texaco Inc. for providing us with the seismic data and the gravity data, and Pecten International for additional migrated seismic profiles. Numerous discussions with Chris Avenius, Steve Hook, Tim Dixon and John Hurst were of great benefit to the ideas presented in the paper. Critical and thorough reviews by Henry Mullins, Robert Sheridan, Jim Dolan, and Paul Mann improved the manuscript. Kevin Cunningham read the final version and improved the English. Financial support for this project was provided by NSF Grant OCE-9314586 (to G.P. Eberli), and several smaller grants from GSA, 1APG, Sigma Xi and GCAGS (to J.L. Masaferro). The Fulbright Commission is acknowledged for supporting J.L. Masaferro in the early stages of the project.
REFERENCES
Anselmetti, ES., Eberli, G.R and Bernoulli, D., 1997. Seismic modeling of a carbonate platform margin (Montagna della Maiella, Italy): variations in seismic facies and implications for sequence stratigraphy. In: I. Palaz and K.J. Marfurt (Editors), Carbonate Seismology. S.E.G. Geophys. Dev., 6: 373406. Austin, J.A., Jr., 1983. OBC 5-A. Overthrusting in a deep-water carbonate terrane. In: A.W. Bally (Editor), Seismic Expression of Structural Styles. Am. Assoc. Pet. Geol., Stud. Geol., 3: 3.4.2-167-172. Austin, J.A., Jr. et al., 1988. Proceedings of the Ocean Drilling Program, Scientific Results, Vol. 101, 501 pp. Ball, M.M., Martin, R.G., Bock, W.D., Sylwester, R.E., Bowles, R.M., Taylor, D.E. Coward, L., Dodd, J.E. and Gilbert, L., 1985. Seismic structure and stratigraphy of northern edge of
191
Bahaman-Cuban collision zone. Am. Assoc. Pet. Geol. Bull., 69: 1275-1294. Beach, D.K. and Ginsburg, R.N., 1980. Facies succession of Pliocene-Pleistocene carbonates, northwestern Great Bahama Bank. Am. Assoc. Pet. Geol. Bull., 64: 1634-1642. Bryant, W.R., Meyerhoff, A.A., Brown, N.K., Furrer, M.A., Pyle, T.E. and Antoine, J.W., 1969. Escarpments, reef trends, and diapiric structures, eastern Gulf of Mexico. Am. Assoc. Pet. Geol. Bull., 53: 2506-2542. Buffler, R.T. and Thomas, W.A., 1994. Crustal structure and evolution of the southeastern margin of North America and the Gulf of Mexico basin. In: R.C. Speed (Editor), Phanerozoic Evolution of North American Continent-Ocean Transitions. Geol. Soc. Am., Decade of North American Geology Project, Continent-Ocean Transect Volume. Geological Society of America, pp. 219-264. Burke, K., Cooper, C., Dewey, J.E and Pindell, J.L., 1984. Caribbean tectonics and plate motions. In: W.E Bonini, R.B. Hargraves and R. Shagam (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162:31-63. Burke, K., 1988. Tectonic evolution of the Caribbean. Annu. Rev. Earth Planet. Sci. Lett., 16:201-230. Crowell, J.C., 1974b. Sedimentation along the San Andreas fault. In: R.H. Dott Jr. (Editor), Modern and Ancient Geosynclinal Sedimentation. Soc. Econ. Paleontol. Mineral. Spec. Publ., 19: 292-303. Denny, W., Austin, J.A. and Buffler, R.T., 1994. Seismic stratigraphy and geologic history of Mid-Cretaceous through Cenozoic rocks, Southern Straits of Florida. Am. Assoc. Pet. Geol. Bull., 78: 461-487. Dillon, RS., Austin, J.A., Scanlon, N. and Klitgord, K.D., 1987. Geology of the Caribbean. Oceanus, 30: 42-52. Draper, G. and Barros, J.A., 1994. Cuba. In: S. Donovan and T.A. Jackson (Editors), Caribbean Geology An Introduction. University of West Indies Publishers' Association, pp. 65-83. Eberli, G. and Ginsburg, R.N., 1987. The segmentation and coalescence of Cenozoic carbonate platforms, northwestern Great Bahama Bank. Geology, 15: 75-79. Eberli, G. and Ginsburg, R.N., 1989. Cenozoic progradation of northwestern Great Bahama Bank, a record of lateral platform growth and sea-level fluctuations. In: RD. Crevello, J.A. Wilson, J.F. Read and F.J. Sarg (Editors), Controls on Carbonate Platform and Basin Systems. Soc. Econ. Paleontol. Mineral. Spec. Publ., 44: 339-351. Eberli, G., Anselmetti, E and Swart, R, 1995. The Bahamas transect: Neogene/Quaternary sea-level fluctuations and fluid flow in a carbonate platform. Safety Package for Leg 166, ODP data file (unpublished). Gealey, W.K., 1980. Ophiolite obduction mechanism. Proc. Int. Ophiolite Symp., Nicosia, pp. 228-243. Gibbs, A.D., 1986. Strike-slip basins and inversion: a possible model for the southern North Sea gas areas. In: J. Brooks, J.C. Goff and B. Van Hoorn (Editors), Habitat of Paleozoic Gas in N.W. Europe. Geol. Soc. London Spec. Publ., 23: 23-35. Gibbs, A.D., 1987. Development of extensional and mixed mode sedimentary basins. In: M.R Coward, J.E Dewey and EL. Hancock (Editors), Continental Extensional Tectonics. Geol. Soc. London Spec. Publ., 28: 19-33. Goodell, H.G. and Garman, K., 1969. Carbonate geochemistry of Superior deep test well, Andros Island, Bahamas. Am. Assoc. Pet. Geol. Bull., 53: 513-536. Harding, T.R, 1983. Divergent wrench fault and negative flower structure, Andaman Sea. In: A.W. Bally (Editor), Seismic Expression of Structural Styles. Am. Assoc. Pet. Geol. Bull., Studies in Geology, Series 15, 4.2-1-4.2-8. Harding, T.E, 1985. Seismic characteristics and identification
192 of negative flower structures, positive flower structures and positive structural inversion. Am. Assoc. Pet. Geol. Bull., 69: 582-600. Harding, T.R, 1990. Identification of wrench faults using subsurface structural data; criteria and pitfalls. Am. Assoc. Pet. Geol. Bull., 74: 1590-1609. Harding, T.R, Vierbuchen, R.C. and Christie-Blick, N., 1985. Structural styles, plate-tectonic setting, and hydrocarbon traps of divergent (transtensional) wrench faults. In: K.T. Biddle and N. Christie-Blick (Editors), Strike-Slip Deformation, Basin Formation and Sedimentation. Soc. Econ. Paleontol. Mineral. Spec. Publ., 37: 51-77. Hempton, M.R. and Barros, J.A., 1993. Mesozoic stratigraphy of Cuba: depositional architecture of a southeast facing continental margin. GCSSEPM Foundation 13th Annu. Res. Conf. Proc., pp. 193-209. Hine, A.C. and Neumann, A.C., 1977. Shallow carbonate bank margin growth and structure, Little Bahama Bank. Am. Assoc. Pet. Geol. Bull., 61: 374-406. Iturralde-Vinent, M., 1972. Principal characteristics of Oligocene and lower Miocene stratigraphy of Cuba. Am. Assoc. Pet. Geol. Bull., 56: 2369-2379. Iturralde-Vinent, M., 1975. Problems of the application of two modern tectonic hypotheses to Cuba and the Caribbean region. Am. Assoc. Pet. Geol. Bull., 59: 838-855. Iturralde-Vinent, M., 1988. Naturaleza geologica de Cuba. La Editorial Cientifico-Tecnica, Habana, 246 pp. Iturralde-Vinent, M., 1994. Cuban geology: A new plate-tectonic synthesis. J. Pet. Geol., 17: 39-70. Klitgord, K.D., Popenoe, R and Shouten, H., 1984. Florida: a Jurassic transform plate boundary. J. Geophys. Res., 89: 7753-7772. Ladd, J.W. and Sheridan, R.E., 1987. Seismic stratigraphy of the Bahamas. Am. Assoc. Pet. Geol. Bull., 7: 719-736. Larson, R.L., 1991. Latest pulse of Earth: evidence for a mid-Cretaceous superplume. Geology, 19: 547-550. Lewis, J.E, 1990. Cuba. In: J.E Lewis and G. Draper (Editors), Geological and Tectonic Evolution of the Northern Caribbean Margin, the Caribbean. Geol. Soc. Am., Vol. H, pp. 77-140. Malfait, B.T. and Dinkelman, M.G., 1972. Circum-Caribbean tectonic and igneous activity and the evolution of the Caribbean plate. Geol. Soc. Am. Bull., 83: 251-271. Mandl, G., 1988. Mechanics of tectonic faulting, models and basic concepts. In: H.J. Zwart (Editor), Developments in Structural Geology, 1. Elsevier, Amsterdam, pp. 1-407. Mann, R and Burke, K., 1990. Transverse intra-arc rifting: Paleogene Wagwater belt, Jamaica. Mar. Pet. Geol., 7: 410-427. Mann, R, Taylor, E W., Edwards, R.L. and Ku, T., 1995. Actively evolving microplate formation by oblique collision and sideways motion along strike-slip faults: an example from the northeastern Caribbean plate margin. Tectonophysics, 246: 169. Masaferro, J. and Eberli, G., 1995. Tectonic destruction and segmentation of the Bahamas carbonate platform. Geol. Soc. Am., Abstr. Prog., Session 81, p. A228. Mattson, RH., 1973. Middle Cretaceous nappe structures in Puerto Rico ophiolites and their relation to the tectonic history of the Greater Antilles. Geol. Soc. Am. Bull., 84:21-37. Mattson, RH., 1974. Cuba. In: A.M. Spencer (Editor), Mesozoic-Cenozoic Orogenic Belts. Geol. Soc. London Spec. Publ., 4: 625-638. Mattson, RH., 1979. Subduction, buoyant breaking, flipping and strike-slip faulting in the northern Caribbean. J. Geol., 87: 21-37. Meyerhoff, A.A. and Hatten, C.W., 1974. Bahamas salient of North America. In: C.A. Burk and C.L. Drake (Editors), The Geology of Continental Margins. Springer-Verlag, New York, pp. 429-446.
J.L. M A S A F E R R O and G.R E B E R L I Mountain, G.S. and Tucholke, B.E., 1989. Abyssal sediment waves. In: A.W. Bally (Editor), Atlas of Seismic Stratigraphy, Vol. 3. American Association of Petroleum Geologists, Tulsa, OK, pp. 233-236. Mullins, H.T., 1984. Structural controls of contemporary carbonate continental margins. In: Platform Margin and Deep Water Carbonates. Soc. Econ. Paleontol. Mineral., Short Course 12, pp. 57. Mullins, H.T. and Lynts, G.W., 1977. Origin of the northwestern Bahama platform: review and reinterpretation. Geol. Soc. Am. Bull., 88: 1447-1461. Mullins, H.T. and Sheridan, R.E., 1983. Wrench tectonic origin for the northern Bahama Platform. Geol. Soc. Am. Abstr. Prog., 15: 648-649. Mullins, H.T. and Van Buren, H.M., 1981. Walkers Cay Fault, Bahamas: evidence for Cenozoic faulting. Geo-Mar. Lett., 1: 225-231. Mullins, H.T., Breen, N., Dolan, J., Wellner, R., Petruccione, L., Gaylord, M., Andersen, B., Melillo, A., Jurgens, A. and Orange, D., 1992. Carbonate platforms along the southeast Bahamas-Hispaniola collision zone. Mar. Geol., 105: 169209. Nettleton, L.L., 1971. Elementary gravity and magnetics for geologists and seismologists. Monogr. Ser., Soc. Explor. Geophys., 1,121 pp. Pardo, G., 1975. Geology of Cuba. In: A.I.M. Nairn and E Stehli (Editors), The Ocean Basins and Margins, The Gulf and Mexico and the Caribbean. Plenum Press, New York, pp. 553-615. Paulus, EJ., 1972. The geology of Site 98 and the Bahamas platform. Init. Rep. DSDP, 11: 877-897. Pindell, J.L., 1985. Alleghenian reconstruction and subsequent evolution of the Gulf of Mexico, Bahamas, and protoCaribbean. Tectonics, 4: 1-35. Pindell, J.L. and Barrett, S.E, 1990. Caribbean plate tectonic history. In: G. Dengo and J.E. Case (Editors), The Caribbean region. The Geology of North America, Vol. H, Geological Society of America, Plate 12. Pindell, J.L. and Dewey, J.E, 1982. Permo-Triassic reconstruction of western Pangea and the evolution of the Gulf of Mexico/Caribbean. Tectonics, 1:179-211. Richard, P. and Krantz, R.W., 1991. Experiments on fault reactivation in strike-slip mode. Tectonophysics, 188:117-131. Roque-Marrero, E and Iturralde-Vinent, M., 1987. Redefinicion de la zona de Cayo Coco en la provincia de Camaguey. Rev. Tecnol., XVII (1): 18-22. Ross, M. and Scotese, C., 1988. A hierarchical tectonic model of the Gulf of Mexico and Caribbean region. Tectonophysics, 155: 139-168. Schlager, W. and Ginsburg, R.N., 1981. Bahamas carbonate platforms m the deep and the past. Mar. Geol., 44: 1-24. Schlager, W., Austin, J.A., Jr., Corso, W., McNulty, C.L., FRiegel, C.V., Renz, O. and Steinmetz, J.C., 1984. Early Cretaceous platform re-entrant and escarpment erosion in the Bahamas. Geology, 12: 147-150. Schlager, W., Bourgeois, E, Mackenzie, G. and Smith, J., 1988. Boreholes Great-Isaac-l, ODP 626 and the history of the Florida Straits. Proc. ODP, Sci. Results, 101: 425-438. Scotese, C.R., Gahagan, L.M. and Larson, R.L, 1989. Plate tectonic reconstructions of the Cretaceous and Cenozoic ocean basins. Tectonophysics, 155: 27-48. Sheridan, R.E., 1974. Atlantic continental margins of North America. In: C.A Burk and C.L. Drake (Editors), The Geology of Continental Margins. Springer-Verlag, New York, pp. 391407. Sheridan, R.E. and Osburn, W.L., 1975. Marine geological and geophysical studies of the Florida-Blake Plateau-Bahamas area. In: C. Yorath, E.R. Parker and D. Glass (Editors),
C E N O Z O I C S T R U C T U R A L E V O L U T I O N OF THE S O U T H E R N GREAT B A H A M A B A N K Canada's Continental Margins and Offshore Petroleum Exploration. Can. Soc. Pet. Geol., Mem., 4: 9-33. Sheridan, R.E., Crosby, J.T., Bryan, G.M. and Stoffa, EL., 1981. Stratigraphy and structure of southern Blake plateau, Northern Florida Straits, and Northern Bahama Platform from multichannel seismic reflection data. Am. Assoc. Pet. Geol. Bull., 65: 2571-2593. Sheridan, R., Mullins, H., Austin, J.A., Jr., Ball, M. and Ladd, J., 1988. Geology and geophysics of the Bahamas, the Atlantic Continental Margin, Vols. 1-2. U.S. Geological Society of America, pp. 329-364.
193
Shipley, T.H., Mayer, L.A. and Winterer, E.L., 1989. Seismic reflections and sedimentation in the pelagic equatorial Pacific. In: A.W. Bally (Editor), Atlas of Seismic Stratigraphy, Vol. 3. American Association of Petroleum Geologists, Tulsa, OK, pp. 224-232. Vaughan, A.E, 1995. Circum-Pacific mid-Cretaceous deformation and uplift: a superplume-related event? Geology, 23: 491494. Walles, EE., 1993. Tectonic and diagenetically induced seal failure within the south-western Great Bahamas Bank. Mar. Pet. Geol., 10: 14-28.
Chapter 8
Deformation History of Roat in Island: Implications for the Origin of the Tela Basin (Honduras)
HANS G. AVI~ L A L L E M A N T and M A R K B. G O R D O N
The rocks of Roat~n Island, Honduras, have experienced several phases of ductile (D1 to D3) and brittle (F1 to Fs) deformations related to the NE to E passage of the Caribbean plate along the Motagua and Swan Islands faults which lie in the North American-Caribbean plate boundary zone. The older, ductile structures formed at metamorphic conditions during Late Cretaceous-early Tertiary left-oblique collision of the Chortfs and Maya blocks. Most brittle structures are younger and formed after uplift and exhumation of the metamorphic rocks. New isotopic data indicate that the exhumation occurred in Late Eocene or Early Oligocene time. The most prominent faults on Roat~in Island are NNE- to NNW- and WNW- to ENE-trending normal faults. The younger NNE- to NNW-trending normal faults may have formed when the island was lying in a releasing bend of the Swan Islands fault zone, whereas the older normal faults formed by 'transtension' consistent with displacement partitioning of the relative divergence rate vector along the Swan Islands fault zone. The deformation history of Roat~in Island suggests that the Tela Basin lying between the island and the mainland of Honduras subsided since Eocene time first by N-S extension and subsequently by E-W extension.
INTRODUCTION The Bay Islands of Honduras, consisting from west to east of Utila, Roat~in, Barbareta, Guanaja, and many smaller islands, together with the Swan Islands to the east are part of the ENE-trending Bonacca Ridge which lies south of the similarly ENE-trending Swan Islands fault zone (Figs. 1 and 2). This fault zone is a left-lateral transform and is the southern boundary of the Cayman Trough, a deep-marine basin formed by sea-floor spreading (Holcombe et al., 1973). South of the Bonacca Ridge lies the Tela Basin on the northern continental margin of Honduras and the northern margin of the Nicaraguan Rise. The Cayman Trough, Swan Islands fault zone, Bonacca Ridge, Tela Basin, and the northern portion of mainland Honduras are all part of the North A m e r i c a n - C a r i b b e a n plate boundary zone (Figs. 1-3). The present study of the rocks of Roat~in Island was undertaken to investigate their deformation history and relate the structures to the relative motions between the North American and Caribbean plates as the present motions (e.g., DeMets et al., 1994) and Cretaceous-Tertiary ones (e.g., Pindell et al., 1988; Pindell, 1993) are quite well known. A
secondary goal of the study was to investigate if and how the deformation history of Roat~n Island constrains the origin and evolution of the Tela Basin.
REGIONAL GEOLOGY AND TECTONICS Prominent features of the North A m e r i c a n Caribbean plate boundary (Figs. 1 and 2) in the region are the western Cayman Trough, the Swan Islands fault zone, and the Motagua fault zone. The Motagua fault is currently the most active of the left-lateral strike-slip faults cutting across northern Central America. South of the Motagua fault zone lies the Chortfs block. North of it lies the Maya block (e.g., Dengo, 1969). Chortis b l o c k
The Chortfs block (Fig. 1) occupies all of Honduras and E1 Salvador, the southern part of Guatemala, and the northern part of Nicaragua and it extends eastward to the Nicaraguan Rise (e.g., Dengo, 1969; Arden, 1975; Case et al., 1984; Donnelly et al., 1990). It is thought to be an al-
Caribbean Basins. Sedimentary Basins of the World, 4 edited by P. Mann (Series Editor: K.J. Hsti), pp. 197-218. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
198
H.G. AVlff L A L L E M A N T and M.B. G O R D O N
Fig. 1. Simplified tectonic map of the Caribbean from Case and Holcombe (1980). Plate motion vectors are given with respect to a fixed Caribbean plate (DeMets et al., 1994). Based on the spreading rate of the Cayman Trough (15 mm/yr; Rosencrantz et al., 1988) plus the slip on strike-slip faults south of the plate boundary, Rosencrantz and Mann (1991) determined that the North America-Caribbean plate motion is 20 mm/year which is notably higher than the Nuvel-lA rate shown here. Box outlines map of Fig. 2. Abbreviations: OF = Oriente fault; P R T = Puerto Rico Trench; SIF -- Swan Islands fault.
Fig. 2. Tectonic map of the Gulf of Honduras and adjacent areas (location on Fig. 1) after Pinet (1976), Case and Holcombe (1980), Manton (1987), and Mills and Barton (1996). Bathymetric contours in km. Thick lines with arrows are strike-slip faults (from Case and Holcombe, 1980; Manton, 1987; Gordon and Muehlberger, 1994; Mills and Barton, 1996); thick lines with tick marks are normal faults (from Pinet, 1976; Manton, 1987; Kozuch, 1991). Abbreviations: A F = Agufin fault; B = Barbareta Island; CF -- Chamelec6n fault; EF = La Esperanza fault; G -- Guanaja Island; GF -- Guayape fault; GoH = Gulf of Honduras; L C F -- La Ceiba fault; M F = Motagua fault; OF -- Oriente fault zone; R L G -- Rio Lenin graben; R V F -- Rfo Viejo fault; R -- Roat~in Island; SG -- Sula graben; SI = Swan Islands; SIF -- Swan Islands fault zone; U -- Utila Island. Focal-mechanism diagrams are from Harvard CMT Database (see Table 1). Box outlines area of Fig. 3.
lochthonous terrane of continental character (Case
Mesozoic
et al., 1984; G o r d o n , 1991) that h a d a c o m p l i c a t e d
1991). I s o t o p i c a g e s o f the m e t a m o r p h i c r o c k s (of-
m i g r a t i o n h i s t o r y ( G o s e , 1985). T h e C h o r t f s b l o c k consists of a generally poorly dated metamorphic
ten o f q u e s t i o n a b l e q u a l i t y ) v a r y f r o m P r e c a m b r i a n ( H o m e et al., 1976a) to C e n o z o i c ( S o u t h e r n w o o d ,
basement
1986). T h e o v e r l y i n g s e d i m e n t a r y units c o n s i s t o f
overlain
unconformably
by
well-dated
a n d C e n o z o i c f o r m a t i o n s (e.g., G o r d o n ,
DEFORMATION HISTORY OF ROATAN ISLAND
199
Fig. 3. Simplified geologic map of Gulf of Honduras region (location on Fig. 2) after McBirney and Bass (1969a), Manton (1987), and Kozuch (1991). Faults in Tela Basin from Pinet (1976), Tyburski (1992), and Paul Mann (pers. commun., 1996). Bathymetric contours (in m) from Pinet (1976) and von der Hoya (1986) (200 m contour along Honduran coast and 1200 and 1400 m contours in Tela Basin are dashed; Tela Basin deeper than 1000 m is shaded. Symbols: i = Late Cretaceous or early Tertiary intrusives; K -- Cretaceous sedimentary rocks; m --- Precambrian, late Paleozoic, or Late Cretaceous-Tertiary metamorphic rocks; Q --- Quaternary sedimentary rocks; v -Tertiary or Quaternary volcanic rocks. Middle Jurassic and Lower Cretaceous clastic rocks, Lower Cretaceous (Aptian and Albian) carbonates, and middle to Upper Cretaceous red beds with interbedded Cenomanian limestones (Mills et al., 1967; Home et al., 1974; Finch, 1981; Gordon, 1993). These rocks were intruded by Cretaceous and Tertiary granitic plutons (Home et al., 1976b; Southernwood, 1986). Cenozoic deposits consist mostly of volcanic rocks with some intercalated (volcani)clastic rocks and red beds (Olson and McGrew, 1941; Williams and McBirney, 1969; Heiken et al., 1991; Bargar, 1991). The Nicaraguan Rise is often thought to be the easterly extension of the Chortfs block. From west to east the crust thins and becomes more island arc in character (Arden, 1975).
Maya block The Maya block (Fig. 1) lies north of the Motagua fault zone and occupies the northern part of Guatemala, all of Belize and the Yucatfin Peninsula of eastern Mexico (Dengo, 1969). It consists of an igneous-metamorphic basement, overlain unconformably by upper Paleozoic, Cretaceous, and Cenozoic sedimentary rocks (e.g., Donnelly et al., 1990; Burkart, 1994). The metamorphism of the basement
rocks occurred in the Paleozoic, but Precambrian inheritance has been shown by Gomberg et al. (1968). The basement is intruded by early Paleozoic granitic plutons (Steiner and Walker, 1996). The sedimentary sequence consists of upper Paleozoic siliciclastics and carbonates, Upper Jurassic to Lower Cretaceous evaporites and red beds, Lower Cretaceous platform carbonates, Upper Cretaceous flysch, and Tertiary clastic rocks (Burkart, 1994). Late Cretaceous to Paleogene granitic intrusions have been reported (Donnelly et al., 1990; Burkart, 1994).
Cayman Trough and the Swan Islands and Motagua fault zones The Cayman Trough (Figs. 1 and 2) is an ENEtrending, approximately 100-km-wide basin locally more than 6000 m deep (e.g., Case and Holcombe, 1980). In the north, the Cayman Trough is bounded by the Oriente transform fault zone which continues eastward all the way to the Puerto Rico Trench (Fig. 1) (e.g., Burke et al., 1984; Calais and Mercier de Ldpinay, 1991; Leroy et al., 1996). The trough is bounded to the south by the Swan Islands transform fault zone (Figs. 1 and 2) which continues westward through the Gulf of Honduras into Honduras and Guatemala. The basin formed by sea-floor
200
H.G. AVE LALLEMANT and M.B. GORDON
spreading (e.g., Holcombe et al., 1973; Perfit, 1977). N-S-trending magnetic anomalies (e.g., MacDonald and Holcombe, 1978) indicate E - W spreading. Rifting may have started at 55 Ma (Land, 1979). Results of earthquake focal-mechanism studies are consistent with the location and orientation of the spreading center (e.g., Molnar and Sykes, 1969; Deng and Sykes, 1995). Spreading rates are difficult to constrain, but may have been 15 m m / y e a r since 30 Ma and 30 m m / y e a r from 44 to 30 Ma (Rosencrantz et al., 1988). Total strike-slip displacements may be of the order of 1100 km (e.g., Rosencrantz and Sclater, 1986; Rosencrantz et al., 1988) whereas normal fault displacements may amount to several kilometers. The generally ENE-trending Swan Islands fault zone consists of two en-echelon, right-stepping fault segments between which a restraining bend developed with NW-trending folds (Rosencrantz and Mann, 1991). On the basis of side-scan sonar and seismic reflection studies, Mann et al. (1991) have shown that the fault segment through the Swan Islands is transpressive. The islands are not only underlain by a major ENE-trending anticline, but also by several ENE-trending, N-vergent thrust faults consistent with the thrust mechanism deduced from the earthquake of December 25, 1995 (No. 13 on Fig. 2 and Table 1). The Swan Islands fault zone trends west as it continues into mainland Honduras and Guatemala, where it is comprised of several major faults, with the Motagua fault currently the most active (e.g., Schwartz et al., 1979; Burkart, 1994). Earthquake focal-mechanism studies (e.g., Molnar and Sykes,
1969; Kanamori and Stewart, 1978; White, 1991; Deng and Sykes, 1995) and surface faulting (PlaNer, 1976) studies have both indicated that the Swan Islands and Motagua fault zones are left-lateral transform faults (Fig. 2).
Amalgamation of the Chortis and Maya blocks According to Donnelly et al. (1990), the Chortfs block, part of the 'Great Arc of the Caribbean' of Burke (1988), collided in Late Cretaceous time with the Maya block approximately along the Motagua fault zone causing partial to total overprint of the Paleozoic metamorphic structures in a divergent orogenic belt with N-vergent thrust faults in the north and S-vergent ones in the south. Because of the post-Eocene displacements along the Motagua fault zone, it is likely that the terrane that collided with the Maya block is the Nicaraguan Rise rather than the Chortfs block proper. Slabs of Late Cretaceous ophiolites (El Tambor Group; McBirney and Bass, 1969b; Donnelly et al., 1990), possibly representing oceanic crust and mantle that once occurred between the Maya and Chortfs blocks, were emplaced both north and southward onto the Maya and Chortfs blocks (e.g., Burkart, 1994). Many plate tectonic models (e.g., Pindell, 1993) suggest that the Chortfs block and Nicaraguan Rise represent the plate that overrode the Maya block and the intervening ocean basin. Based on the age of sedimentary rocks in strike-slip basins, Burkart (1994) suggested that the Motagua fault zone and the other faults parallel to the Motagua became active in Eocene time.
Table 1 Mechanisms of selected earthquakes which occurred since 1976 in the region from the Harvard CMT Database (Dziewonski et al., 1981) and a single event from 1969 (Dean and Drake, 1978) No.
Yr
Mon
Dy
Hr
Mn
Sec
Lat
Lon
D
Mo
1 2 3 4 5 6 7 8 9 10 11 12 13
69 77 77 80 80 80 81 82 90 91 92 94 95
2 8 8 3 8 9 6 4 7 1 7 6 12
25 20 20 20 9 2 11 10 27 29 21 8 25
7 2 3 16 5 10 18 16 0 5 15 20 18
39 46 51 54 45 28 34 25 54 29 31 53 18
2.0 11.8 54.7 18.3 9.5 8.8 20.6 34.1 57 3.6 58.7 25.9 17
15.30 16.61 16.70 16.84 15.89 15.91 16.72 17.53 16.06 16.88 17.52 16.76 17.51
-87.40 -86.85 -86.61 -85.71 --88.52 --88.29 --86.11 --83.37 --86.25 --85.53 --83.66 --85.85 --82.87
24 14 36 27 22 25 20 10 15 33 10 10 10
1.58 • 3.81 x 1.48 • 6.66 x 9.62 x 8.89 x 3.71 x 1.62 • 1.16 x 6.29 x 2.01 x 1.43 •
1025 1025 10 24
1025 1023 1024 1025 10 24
1024 1023 10 24 10 24
Mw
strl
dipl
rkl
str2 dip2
rk2
5.4a 6.1 6.4 5.4 6.5 5.3 5.9 6.3 5.4 5.3 5.2 5.5 5.4
339 157 156 74 352 176 72 79 357 254 176 67 257
42 68 61 80 67 90 72 70 17 81 90 70 37
-168 -168 9 --162 180 1 --16 --136 3 --180 --17 41
86 62 60 342 255 266 341 175 224 164 266 163 133
-23 -29 170 --25 0 162 --159 --78 171 0 --159 120
75 79 80 81 74 90 89 75 78 87 90 74 67
Most mechanisms have been published in Physics of Earth and Planetary Interiors from 1983 to 1997, except for those in press and the 1969 one (No. 1) from Dean and Drake (1978). Abbreviations: No., Yr, Mon, Dy, Hr, Mn, Sec, Lat, Lon, D, Mo, Mw, strl, dipl, rkl, str2, dip2, and rk2 stand for number, year, month, day, hour, minute, second, latitude, longitude, depth (in km), seismic moment (in dyne-cm), moment magnitude (calculated from the seismic moment), strike of first fault plane, dip of first plane, slip direction (given as rake) of first plane, strike of second plane, dip of second plane, and slip direction (rake) of second plane, respectively. Time is given as Universal Time. a mb.
DEFORMATION HISTORY OF ROAT/kN ISLAND
201 Quaternary deposits. Most of Utila Island is covered by Quaternary deposits, but the most easterly tip of the island is underlain by Holocene alkali-basalt (Wadge and Wooden, 1982).
The northern portion of the Chortfs block is disrupted by many approximately ENE-trending strikeslip faults and N- to NNE-trending normal faults. The strike-slip faults are all thought to be left-lateral. The Rio Viejo fault has left-lateral kinematic indicators (M.B. Gordon, unpubl, data) and an earthquake focal-mechanism plot shows that the Aguan fault is left-lateral as well (Gordon and Muehlberger, 1994; see also Fig. 2). Some of the normal faults are seismically still active, e.g., the Sula graben (Osiecki, 1981; Mann and Burke, 1984; White, 1991; Gordon and Muehlberger, 1994) and the Rio Lean graben (Manton, 1987). The fact that the Rio Lean graben abuts against the La Ceiba fault suggests that the latter is a left-lateral strike-slip fault as well.
Lithology of Roatfin Island Roatan Island is underlain by metamorphic and igneous rocks. McBirney and Bass (1969a) divided the metamorphic rocks into two packages: a highgrade package overlain by a low-grade one, the latter containing large masses of serpentinite (Fig. 5a). They assumed that the contact between the two was a south-dipping low-angle thrust fault. The high-grade, amphibolite facies assemblage consists of biotite gneiss (often with large plagioclase augen), biotite and muscovite schists with bands and lenses of metagabbro, pyroxene hornblendite, amphibolite, and marble. They are invaded by many quartz and rare pegmatite veins. The low-grade, greenschist facies assemblage consists of phyllite, (stretched) pebble conglomerate, chlorite (with or without muscovite) schist, serpentinite, (meta)sandstone, chert, and marble and limestone. The smaller island Barbareta, just east of Roatan, is underlain by chlorite schist, serpentinite, and granite porphyry.
STRUCTURAL GEOLOGY OF ROATAN ISLAND
The Bonacca Ridge stretching from Utila Island in the west to the Swan Islands in the east (Fig. 2) is an ENE-trending ridge separated from the Cayman Trough by the sinistral Swan Islands transform fault as well as by a similarly ENE-trending normal fault (Figs. 3 and 4). The ridge, consisting of several en-echelon, right-stepping segments (Pinet, 1976) is mostly submarine, except for the Bay Islands and the Swan Islands. The deepest sites of the ridge are only about 650 m below sea level (Case and Holcombe, 1980). The ridge is separated from the Tela Basin by a normal fault. Thus, the Bonacca Ridge is a horst, but one that has been interpreted to be in part antiformal in structure (Pinet, 1976). The major Bay Islands, consisting from west to east of Utila, Roatan, Barbareta, and Guanaja islands have been mapped by McBirney and Bass (1969a). Roatan, Barbareta, and Guanaja islands are underlain by metamorphic rocks, a few plutonic rocks, and
Age constraints Marble We collected a sample of low-grade marble from the eastern part of Roatan (sample H-94-14; locality No. 33; for location see Fig. 5b, Table 2). J. Tom Dutro, Jr. (pers. commun., 1995) identified gastropod fragments that constrain the age of the marble to between Triassic and Recent. Because Cretaceous limestones are common in the Chortfs block, we as-
LI2 ~2
L~
g
Ud
o
lz I3
I
SIA
~174
-I-
'0 ,o
k
~J
~
3
~ ,~
0
Q:
3
- -
W/v
+
+ ' ~~ L' , + +
o
................i :
q+
-.+
+ 2),+
', ~
r
Tela
~~176Coymon ~ Trough
~
.......
.
+ ' ,
N +
+',,+ +""
Basin
+
+
.
.
.
.
.
A/N
.
~__z
+ +
.
I0
km.
....
7
.
Fig. 4. Speculative N-S cross-section through Cayman Trough, Roat~n Island, Tela Basin, and northern portion of Honduras along line A-A' on Fig. 3. Thickness of continental crust (crosses) and of oceanic crust in Cayman Trough (v-pattern) from Case et al. (1984); thickness of sediments in Tela Basin from yon der Hoya (1986). Faults from Manton (1987), Kozuch (1991), and Tyburski (1992); fault dips are speculative. Faults merging into basal detachment along the Moho-discontinuity (M) is speculative as well9 Earthquake mechanisms (lower-hemisphere, equal-area) are projected onto a vertical N-trending plane (numbers 3 and 9 refer to same earthquake plots in Fig. 3).
202
H.G. AVI!~LALLEMANT and M.B. GORDON
.Q o ~
"i L
~
,,_
.Iti
9
~ (D
0
Ol~
0
~| |
1,1
~
9,-- Y:,,| ,.. L
0,
"0~
9
" "
--~-~ ~
~,.,
?
9
oR
. . _
br
E
114111111
9 .,..o 9
\ ',..(,
.,,=~
o
~o
'gloge~
' glo 90 '
~
~..,.~
r
9
I' ...-:.
,,0s '~o9
"~163 ,,z~>,,ge..
<~ $
=I-" I"
~g
eat
0
7~ .,..a
.,.,~
i-i .Io
==i
~=-~-.><-....,b . i
I=
)
~~v,.~
"-
i= v, ~'..~'~-.-.,d~ V',
~
*.W~"..,#
~
~
#
b
i i
DEFORMATION HISTORY OF ROATAN ISLAND
203
Table 2 Geographic location of localities mentioned in the text and figures Site No.
Site name
UTM coordinates
Latitude oN
Longitude oW
Rock type
Walk Point Wild Mines Wild Mines Jonesville Ranch Hong Kong Hill Further Wetl Hill West End Point
16QEP741163 16QEP514046 16QEP511043 16QEP678141 16QEP651127 16QEP646126 16QEN424984
16~ ' 16~ 16o19.2' 16o24.5' 16023.7' 16o23.7' 16~ '
86018.4' 86~ 86031.3 ' 86021.9' 86023.5 ' 86023.7 ' 86036.2 '
metaconglomerate muscovite schist muscovite schist metaconglomerate amphibolite amphibolite amphibolite and marble
11 12 12A 13 14 15 20B 22 24 24C 25 25C 27 32 33
Bodden Bight Blue Rock Smooth Bay Coxen Hole Coxen Hole Coxen Hole Big Bight Big Plan Mid Isle Mid Isle Key Point West End Mid Isle Dixon Hill Diamond Rock
16QEP674116 16QEP892172 16QEP906182 16QEP503039 16QEP494043 16QEP495048 16QEP602103 16QEP610104 16QEP617109 16QEP622108 16QEN435998 16QEN428988 16QEP630110 16QEP503052 16QEP709148
16024.0' 16~ ' 16o26.6' 16o19.0' 16o19.2' 16o19.4' 16o22.4' 16o22.5' 16022.7' 16022.7' 16~ 16o16.2' 16022.8' 16o19.7' 16024.9'
86~ 86009.9 ' 86~ ' 86031.8 ' 86032.3 ' 86032.2' 86026.2 ' 86025.7 ' 86025.3 ' 86~ ' 86035.6 ' 86036.0 ' 86024.6' 86031.7 ' 86020.2 '
34
Alligator Nose
16QEP732164
16025.7'
86o18.9 '
1 4 4B 5B 7 7B 8
sume that the Roat~in limestone is Cretaceous as well in agreement with this fossil identification. We collected several other samples of low-grade limestone for conodont analyses, but all samples were barren.
Amphibolite A sample of amphibolite was collected from the w e s t e n d o f Roat~in I s l a n d ( s a m p l e n u m b e r H - 9 3 - 1 2 ; locality No. 8; for location, see Fig. 5b, Table 2). H o r n b l e n d e o f this r o c k w a s d a t e d b y the 4 ~
Result
fault mversmn fault inversion fault inversion fault inversion fault mverslon fault inversion 4~ sample, conodont analysis, Fig. 10c metalimestone conodont analysis serpentinite fault inversion granite porphyry zircon fission-track sample chlorite schist fault inversion serpentinite fault inversion serpentinite fault inversion biotite gneiss Fig. 10a,b quartz-feldspar schist fault inversion, Fig. 14 amphibolite and pegmatite fault inversion amphibolite and pegmatite fault inversion metalimestone conodont analysis metalimestone conodont analysis biotite gneiss fault inversion, Figs. 15, 16 chlorite schist fault inversion metalimestone fossil analysis (Jurassic-Recent) metalimestone and chlorite fault inversion and conodont schist analysis
Ductile deformation structures Most
rocks
o f Roatfin I s l a n d h a v e u n d e r g o n e
several ductile f o l d i n g (D) and brittle faulting (F) p h a s e s o f d e f o r m a t i o n (Table 5). T h e first ductile d e f o r m a t i o n (D1) is s y n m e t a m o r p h i c w h i l e the seco n d (D2) and third (D3) p o s t - d a t e the p e a k m e t a morphism
and
refold
the
first g e n e r a t i o n
struc-
tures. Five g e n e r a t i o n s o f faults (F1 to Fs) w e r e
'ot
m e t h o d and y i e l d e d a L a t e E o c e n e p l a t e a u age o f 36.0 :i: 1.2 M a (Fig. 6, Table 3).
I
I
A sample of granite porphyry was collected from
I
I
I
I
I
H-93-12 hornblende
7O
Granite porphyry
I
60 m
B a r b a r e t a I s l a n d ( s a m p l e H - 9 3 - 1 3 ; l o c a l i t y No. 13; for location, see Fig. 5b, Table 2), but it c o n t a i n e d insufficient z i r c o n for a U - P b age d e t e r m i n a t i o n . T h e z i r c o n c o n t e n t w a s sufficient t h o u g h for fissiontrack a n a l y s i s (Table 4), w h i c h p r o v i d e d an age o f 39.4 :k 2.8 M a ( L a t e E o c e n e ) .
v
50-ol
,,r
m
40--
r
I--
Q. Q. <:
30-20--
36.0 + 1.2 M a
10--
Conglomerates McBirney and Bass (1969a) assigned a Tertiary age to deposits of apparently unmetamorphosed conglomerates and sedimentary breccia from a small island off Guanaja Island. The cobbles and fragments contain fossils of Cretaceous age.
O--
I 0
I 20
I
I 40
I
I 60
I
I 80
I
100
Cumulative % 39 Ar Released
Fig. 6. 4~ release spectrum of hornblende from amphibolite of Roat~in Island (locality No. 8; location on Fig. 5b and Table 2; more data in Table 3).
204
H.G. AVE LALLEMANT and M.B. GORDON -d m. qq 9 0 0
< "0
,< o'b
"0
O~
o
=o 0 9
0
,<
70 7 07 7~7 7 7 707 07 7 7~7 7~ 70
o'b
X
X
X
X
X
X
X
0
X
X
~ 0
X
X
0 0
X
X
0
X
X
7 0 0
0 9
o
0 0 o O~
,< r
r
0
u'b o'b ,--~ c q
O~ 0
~'b t'~ ~
O~
0 0
,<
<
0
<
0
0 "0 0 o
< 0
,<
r
0 0
0 0 ~
0 ~
0 ~ 0
0 0
~
00
"0
.<
0
<
9
r~
"0
o
0
[-
o 0 0
0
X
9
0
X OO 0
0
~
~
~
It3
0
~rb ,~"
O~ OO t'~. ~
r
r
r
r
r
r
r
"0
, - d d o o o o d o d o o d o o
0 0 "0
0 0 O
~ o. o. o o
O
t'..
b
O
t"~ OO O~
U
'
O
0 O
0
0 O
0
0
,"~
~
0
0 ~
0
0
0
0
C',l C",l O'b c~'b ~'~
0 0
9
~dd~ "0
,< m
"~ "~' ~" ~" <<<
~~~
~ ~'
0
0
0
0 r
0 0 0 r
r
0
sssssssssssss 0
0
t'xl t'~l C',I r
0
r
0
r
O 0
r
r
~
r
cq
<<< C',I ~',1 ~',1 ~',1 ~,1 t"q t",l ~q
~-1 ~1
~',1 ~,1 ~,1 t',l
6 2:
~',1
0 9
DEFORMATION HISTORY OF ROATAN ISLAND
205
Table 4 Zircon fission-track data from granite porphyry on Barbareta Island Sample Mineral No. of grains Standard track density Fossil track density Inducedtrack density Chi square prob. Fission-track age (xl06 cm-2) (xl06 cm-2) (xl06 cm-2) (%) (Ma) H-93-13 zircon
11
0.21 (1359)
556.30 (413)
502.42 (373)
89
38.9 4- 2.8
Locality No. 12b, for location see Fig. 5b and Table 2. Brackets show number of tracks counted. Standard and induced track densities measured on mica external detectors (g - 0.5), and fossil track densities on internal mineral surfaces. Ages determined by Ann E. Blythe using r = 335 + 20 (zircon) for dosimeter glass SRM 962a (e.g. Hurford and Green, 1983). Ages are pooled. Table 5 Summary of deformation structures of Roatfin Island and their kinematic interpretation Phase
Structure
Orientation
Kinematic interpretation
D1
S1 cleavage B1 fold axis L1 lineation
NE strike, moderate SE dip shallow SW plunge shallow SW plunge
Shear on Sl with top to SW sense
D2
$2 axial planes
NNW-SSE contraction
B2 fold axis L2 lineation
(1) E-W strike, moderate S dip (2) N-S, strike, moderate W dip shallow SW plunge shallow SW plunge
D3
$3 kink bands B3 fold axis
NNW strike, steep shallow NNW and SSE plunge
ENE-WSW contraction
Fl
Thrusts faults
(1) N-S strike, shallow W dip (2) N-S strike, shallow E dip
E-W contraction
F2
Strike-slip faults
(1) (LL) NNW strike, steep (2) (RL) ENE strike, steep
NNE-SSW extension
F3
Normal faults
(1) E-W strike, moderate N dip (2) E-W strike, moderate S dip
N-S extension
F4
Strike-slip faults
(1) (RL) NW strike, steep (2) (LL) NE to ENE strike, steep
E-W extension
F5
Normal faults
(1) N-S strike, moderate W dip (2) N-S strike, moderate E dip
E-W extension
LL and RL stand for left-lateral and right-lateral, respectively.
observed to cut across the ductile deformation structures. The first recognized phase of folding on Roatfin (D1) is synmetamorphic and is related to mylonitization. It is expressed by small isoclinal folds (B1) with a strong axial-planar cleavage (S1) and mineral, stretching, and intersection lineations (L1) all parallel to the fold axes (Fig. 7a, Figs. 8 and 9). Generally the cleavage trends ENE, parallel to the long axis of Roatfin Island and it dips moderately to the south-southeast. The lineations and fold axes tend to plunge gently to the west-southwest. The deformation is clearly non-coaxial as plagioclase augen in biotite gneiss and quartz lenses in marble have asymmetric tails indicating shear with a top-to-the-SW sense (Fig. 10). The S1 cleavage is folded by the D2 deformation, expressed by a conjugate set of chevron-style folds with fold axes (B2) generally coaxial with the first fold axes (B1) and axial planes ($2) moderately
dipping to the south and west (Fig. 7b, Fig. 9). In fact, the entire island is a major D2 antiform which plunges to the west-southwest (Figs. 8 and 9). Rare kink bands (D3) refold all the previous structures and their axial planes ($3) are steep and strike NNW. Kink axes (B3) are sub-horizontal and plunge to the north-northwest or south-southeast (Fig. 7c, Figs. 9 and 10b). Inspection of the geological map of Guanaja Island of McBirney and Bass (1969a) suggests that the rocks of Guanaja have undergone at least the D1 and D2 deformations. The orientations of these structures are very similar to the D1 and D2 structures on Roat~n, implying that no appreciable vertical-axis rotation has occurred between the two islands. Metamorphic rocks south of the Motagua fault zone in Honduras have all three (D1, 02, and D3) folds approximately of similar orientation as those on Roatfin Island, although the $1 cleavage is steeper on the mainland (unpubl. data, M.B. Gordon and
206
H.G. AV]~LALLEMANT and M.B. GORDON
N
b
?
N II
| |
9
| illl
e
9 I
|
9
a# e
|| e
~e
i
|
0
II I
.
d "
~
9 9
9
I
9
la"
el 9
.
e
| |
D,
D2
D~
Fig. 7. Equal-area, lower-hemisphere projections of mesoscopic ductile fabric elements from Roat~in Island. Solid symbols are poles to planes (dot = cleavage; square - axial plane of fold); open symbols are lineations (dotted circle = fold axis; open circle --- stretching lineation [arrows show sense of displacement; arrows pointing away from center of diagram show 'normal' motion and arrows pointing toward center show 'thrust' motion]; triangle -- mineral lineation; square -- intersection lineation). Diagrams (a), (b), and (c) show the data of the first (D1), second (D2), and third ductile deformation (D3), respectively.
N
b
N
\
Fig. 8. Equal-area, lower-hemisphere projections of poles to D1 cleavages (S]) in southeastern (a) and northwestern domain (b); domain boundary follows the divide of Roat~in. Contours in (a) at 3, 6, 9, 12, 15, and 18% per 1% area and in (b) at 5.5, 11.1, 16.5, 22.2, and 27.8% per 1% area. Dashed great circle connects the two maxima of (a) and (b); dotted circle is pole to dashed great circle and probably megascopic B2 fold axis.
or tilting of the Bay Islands probably has occurred though (see below).
Brittle deformation structures
Fig. 9. Interpretative block diagram showing schematically the D1 to D3 structures as shown in Fig. 7. S and L stand for axial planes or cleavage and lineation, respectively, with subscripts referring to the phases D1, D2, and D3. L1 (lozenges) is stretching lineation; L2 and L3 are intersection lineations ($2 with $1 and $3 with $1, respectively).
H.G. Av6 Lallemant, 1994). Thus, vertical-axis rotations between the mainland and the Bonacca Ridge have not occurred either. Horizontal-axis rotation
Roat~in Island is covered by tropical vegetation and plantations. Therefore mesoscopic analysis of brittle structures relies mostly on road and coastal outcrop. Megascopic faults that might be present are not exposed because of erosion and vegetation. We did not observe the low-angle thrust fault inferred by McBirney and Bass (1969a) between the lowand high-grade metamorphic rocks. However, rare mesoscopic thrust faults do occur near the contact.
Aerial photographs To assess the presence of megascopic faults, aerial photographs were analyzed to find preferred orientations of lineaments. Lineaments on aerial pho-
DEFORMATION HISTORY OF ROAT,/~N ISLAND
207
Fig. 10. Photographs of biotite gneiss [(a) and (b); locality No. 20B, see Fig. 5b, Table 2] and marble [(c); locality No. 8; see Fig. 5b, Table 2]. Asymmetric plagioclase augen in gneiss and asymmetric quartz lenses in marble show top-to-the-WSW shear (top to the left). Kink fold in (b) is D3 structure. tographs may be the result of lithological contacts, but generally they result from differential weathering along fault planes. Two factors have to be consid-
ered in interpreting aerial photographs taken at low angles of incidence of sun rays: (1) shadows may cause apparent lineations; and (2) linear valleys may
208
H.G. AVI~ LALLEMANT and M.B. GORDON
N
I0
5 %
0 /
5
I0
I0"
Fig. 11. Rose diagram of trends of lineaments determined from air photos (Fig. 5b) expressed as percentage of lineament length per 10~ Total length of lineaments is 241 km. Note the large maximum of 12.2% at N25W and two lesser maxima of 9.6% and 6.4% at N85W and N65E, respectively. be invisible if they are parallel to the direction of incidence, but valleys perpendicular to that direction are strongly accentuated. Fortunately, both factors are not important in the present study, because the photos were taken on April 1, 1989 between 11:00 AM and 1 : 0 0 PM.; thus, the incidence angle of the sun rays was generally greater than 75 ~. The lineaments observed in the aerial photographs were plotted on the geological map of Roat~in Island on Fig. 5b. The orientation and length of the lineaments were measured and plotted in a rose diagram (Fig. 11) showing that the most common lineaments trend N25W, almost perpendicular to the Swan Islands fault zone. Two lesser-preferred orientations are N85W and N65E, the latter direction parallel to the Swan Islands fault zone.
Mesoscopic faults As many fault planes and slickenside striations as possible were measured in each outcrop or locality, with the sense of slip of each fault determined by techniques described by Petit (1987) and Twiss and Gefell (1990). The relative age of a fault with respect to other faults was determined by studying crosscutting relationships of the fault planes or slickenside striations. Using the stress inversion technique of Angelier (1990, 1991 a), the fault data for each locality or a set of localities were separated into several inter-
nally consistent sets. In this study, the results of the inversions are not given as principal stresses, but as principal strains, because stresses are calculated and thus derivatives of the displacements and strains. The fault data are presented in Fig. 12 and Table 6, while simplified block diagrams are shown in Fig. 13. Our fault data from Roat~in and Barbareta Islands can be separated into five sets, but these sets do not occur in all outcrops. We found some relative age relationships between the fault sets, but not sufficient to uniquely constrain the relative timing of the deformation history. No relative age relationship between the F1 fault set and the other sets could be established. The only relationships that are certain between the fault sets F2 to F5 are: (1) F2 > (is older than) Fs; (2) F3 > F4; and (3) F4 > Fs. There are thus only three possibilities for sets F2 to Fs: (1) F2 > F3 > F4 > Fs; (2) F3 > F2 > F4 > Fs; and (3) F3 > F4 > F2 > F5 (Angelier, 1991b). How then was the relative timing scheme F1 > F2 > F3 > F4 > F5 chosen? Timing relationships of set F1 are completely absent, but the kinematics of this set of faults are compatible with those of the third phase of folding (D3). None of the other fault sets is compatible with the earlier folding phases (D] and D2). Thus we speculate that the F] faults are the oldest brittle structures. In all areas F5 is evidently the youngest. Because the strains of F4 and F5 are similar (both are related to E - W extension) and because both sets F2 and F3 are characterized by N - S extension, we suggest that the first scheme (F1 > F2 > F3 > F4 > F s ) a p p l i e s to Roat~in Island. Only four F1 thrust faults (Fig. 12a, Fig. 13a) were found and only in a small area (localities Nos. 13, 14, and 15; see Fig. 5b, Table 2). They indicate E - W contraction. They occur near the contact of the high- and low-grade metamorphic rocks that McBirney and Bass (1969a) had interpreted as a thrust fault. Although our F1 data are consistent with that interpretation, the superposition of low-grade on high-grade metamorphic rocks suggests that the contact is a low-angle normal fault. The mesoscopic F2 faults (Fig. 12b, Fig. 13b) occur in two orientations. The steep NNW-striking ones are left-lateral strike-slip faults, whereas steep ENE-striking ones are right-lateral strike-slip faults. These structures indicate N N E - S S W extension.
Fig. 12. Equal-area, lower-hemisphereprojections of mesoscopic faults (great circles) and slip directions; small arrows pointing outwards indicate normal faults, pointing inwards indicate thrust or reverse faults; half arrows indicate strike-slip faults (sense of displacement follows usual practice). Fault analysis was performed using the method of Angelier (1990, 1991a); resulting principal strain axes are shown as 3-, 4-, and 5-pointed stars indicating, respectively, the maximum principal extension (X), the intermediate principal strain axis (Y), and the maximum principal shortening axis (Z). Large solid arrows pointing outward show the trend of the X-axes and those pointing inward are the trend of the Z-axes. Localities where fault analysis was performed are shown here and in Table 2. Diagrams (a) to (e) show F1 to F5 faults, respectively. Symbols: triangles = poles to joints; open squares = poles to tension gashes; solid squares = fiber lineations.
DEFORMATION HISTORY OF ROAT/~N ISLAND
209
210
H.G. AVt~ LALLEMANT and M.B. GORDON t",l
r tt~
~h 9 9 .,-~
b
- ....
~
~
/__
,,1
0
Zd .,..~
9
~5 .,..~
,.o
~ ,.o
Fig. 13. Interpretative block diagrams showing graphically the five sets (F1 to Fs) of mesoscopic faults depicted in Fig. 12. (a) F1 thrust faults resulting in E-W shortening. (b) Set of conjugate strike-slip faults (F2) resulting in N-S extension. (c) Set of normal faults (F3) resulting in NNE extension. (d) Set of conjugate strike-slip faults (F4) causing E-W extension. (e) Set of conjugate normal faults (Fs) causing E-W extension.
The F3 faults are WNW- to ENE-striking normal faults dipping moderately to the north and south (Fig. 12c, Fig. 13c). They are consistent with NNE to NNW extension suggesting that F3 is related to F2. These faults may be correlated with the N85W and N65E lineaments observed on the aerial photographs (Fig. 5b, Fig. 11). Well developed vertical joints striking parallel to the faults may be related to the F3 event. The F4 faults consist of steep NE- to E-striking left-lateral strike-slip faults and steep NW- to N-trending left-lateral ones (Fig. 12d, Fig. 13d). These faults indicate E - W extension. The NE- to ENE-striking left-lateral faults are parallel to the Swan Islands fault and they may be correlated with the N65E lineaments on the aerial photographs (Figs. 5 and 11). The F5 structures are N- to NNW-striking normal faults (Figs. 14-16) and may be related to F4 as displacements along both caused E - W extension. This generation of faults is by far the most common structure on Roatfin Island and can be seen in many roadside outcrops (e.g., Fig. 14). These faults are not only the most numerous ones as the aerial photographs show (Fig. 5b, Fig. 11), but they appear to have caused very large strains. In one outcrop (Fig. 15, Fig. 16c) E - W extensile strains are at least 80%. North-trending joints and tension gashes with W-trending mineral fibers (generally quartz) may be related to this deformation.
9
9
,.o
2=
c~
t",l
9
~.~
~.~ ~'~
N~ r~
0 9
DEFORMATION HISTORY OF ROAT.A,NISLAND
211
Table 6 Orientation data and statistics of fault analysis of Roatdn and Barbareta Islands using inversion technique of Angelier (1990, 1991 a) Phase
Program
o1
02
5 5 3 4 5 5 4 2 3 1 4 5 5 2 3 5 4 5 3 5 3 4 3 5 3 5
INVD INVD INVD INVD INVD INVD INVD INVD INVD INVD INVD R4DT R4DT INVD INVD INVD INVD INVD INVD INVD INVD R4DT INVD INVD INVD INVD
277.0 86.3 11.9 62.1 351.1 79.0 3.9 32.7 113.6 73.1 230.1 71.3 22.6 18.0 286.4 4.3 17.5 80.4 270.1 0.4 187.7 27.1 272.2 73.2 156.0 87.2 294.4 32.3 296.1 83.2 357.8 76.6 189.8 21.1 39.3 71.2 295.2 84.1 52.1 70.7 120.3 87.4 221.4 1.2 302.3 63.7 44.4 70.3 305.5 63.7 44.7 74.2
170.0 180.9 97.2 154.5 342.6 14.5 157.9 29.5 283.0 180.1 41.2 157.4 357.2 88.7 111.2 157.9 302.3 184.4 111.5 185.5 332.6 352.5 63.9 206.0 70.7 206.2
03 1.1 27.4 3.1 53.6 11.3 15.4 65.4 71.7 0.8 8.7 58.5 7.2 2.6 55.0 6.8 12.6 44.8 15.6 5.9 13.5 2.2 88.1 14.5 18.8 15.9 15.0
79.9 273.3 187.8 264.5 250.1 107.4 287.2 195.1 192.9 2.5 285.5 65.4 267.2 196.6 201.3 248.9 82.4 277.3 201.5 278.8 242.5 131.3 159.7 298.0 166.8 297.5
3.5 4.5 10.6 14.2 12.4 10.4 16.2 17.8 9.6 81.3 14.9 15.1 1.0 12.2 0.6 4.4 37.7 10.2 0.4 13.5 1.4 1.4 21.4 5.8 20.4 4.8
fi0
Ang
Rup
No. faults
Site name
Site number
0.558 0.244 0.795 0.511 0.395 0.390 0.771 0.372 0.402 0.409 0.428 0.537 0.159 0.697 0.540 0.205 0.387 0.456 0.533 0.421 0.638 0.425 0.219 0.625 0.160 0.563
2 5 8 3 6 9 12 12 9 25 9 2 13 5 7 15 8 4 7 5 10 9 4 13 5 13
12 24 18 30 19 29 44 27 31 54 37 37 29 22 22 30 31 15 21 27 32 50 30 41 30 37
5 6 5 4 7 10 5 7 12 4 7 5 4 13 8 4 4 6 9 12 6 8 5 7 6 9
Alligator Nose Blue Rock Blue Rock Blue Rock CoxenHole Coxen Hole CoxenHole CoxenHole Coxen Hole Coxen Hole Coxen Hole Hong Kong Hill Hong Kong Hill Mid Island Mid Island Mid Island Mid Island Mid Island Mid Island Mid Island Big Plan Walk Point Wild Mines Wild Mines Wild Mines and Dixon Hill Wild Mines and Dixon Hill
34 12 12 12 14 15 15 13, 14,15 13, 14, 15 14, 15 13, 14, 15 5B, 7 7 24C 24C 24C 24C 27 24C, 27 24, 24C, 27 22 1 4 4 4, 32 4, 32
Note that in the text, principal strain axes X, Y, and Z are being used instead of 03, 0"2, and 0.1, respectively.
,-
I EN E
i
~
~
-
----
i
-- ~ - = - - ~ ~ ~
~~-,~~'-%=--
~-
-
-
:
W SW
. ----...5.
Fig. 15. ENE-trending road section showing stretching (F5 deformation) of competent quartz-feldspar schist (boudins A to D) in incompetent phyllite. Extensional strain parallel to outcrop measured from center of boudin B to center of boudin D is about 80%. Note quartz in pressure shadows of boudins and the late F5 faults post-dating the boudinage; in locality No. 27 (location on Fig. 5b and Table 2).
Tectonic evolution of Roat~n Island
It has been proposed that left-lateral displacement along the Swan Islands and Motagua fault zones since the 55 Ma (Early Eocene) initial opening of the Cayman Trough (e.g., Rosencrantz et al., 1988) is about 1100 kin. Thus, Roat~in Island could have originated far to the northwest off the west coast of Mexico. However, the island may not always have been part of the Caribbean plate, but could at times have been attached to the North American plate, if major left-lateral strike-slip displacements occurred along the Agu~in, La Ceiba, La Esperanza, or Rio Viejo faults (Figs. 2 and 3; e.g., Gordon
and Av6 Lallemant, 1995). Thus, Roat~in Island may have been displaced eastward for much less than 1100 km, but for how much cannot be ascertained because the metamorphic rocks on Roatdn Island are lithologically different from those in the Chortfs and Maya blocks. The rocks most similar to the rocks exposed on Roatdn are the ophiolitic rocks of the E1 Tambor Group (first defined as the E1 Tambor Formation diabase dikes, basalts and chert have not been encountered in the Bay Islands. However, they are relatively rare in the Motagua Valley and may be metamorphosed beyond recognition in the Bay Islands. The chlorite schist found in the Bay Islands is not generally considered part of the E1 Tambor
212
H.G. AVI~ L A L L E M A N T and M.B. G O R D O N
Fig. 16. Photographs of F5 structures. (a) Boudins of quartz-feldspar gneiss in mica schist showing ENE extension; locality No. 8, Fig. 5b). (b) Northward view of conjugate set of N- to NW-trending normal faults in quartz-feldspar gneiss with quartz veins at the fault intersection (locality No. 22, Fig. 5b). (c) Southward view of boudins of quartz-feldspar gneiss in mica schist with quartz in the pressure shadows (locality No. 27, Fig. 5b and 15).
DEFORMATION HISTORY OF ROATAN ISLAND Group. However, Donnelly et al. (1990) state that the San Diego phyllites of the Motagua Valley are commonly difficult to distinguish from rocks of the E1 Tambor Group; thus this group is similar to the rock sequence found in the Bay Islands. Wilson (1974) showed that the 'Sanarate limestone' is in normal contact with the graywacke-chlorite schist of the E1 Tambor Formation, and Lawrence (1975) included it in the E1 Tambor Group calling it the Cerro de la Virgen Limestone. Wilson (1974) found rudist fragments of certain Cretaceous age and probable Coniacian-Campanian age. Thus a major metalimestone unit of the E1 Tambor Group may also be similar to the metalimestone on Roatfin.
Displacement partitioning It is at present well known that deformation structures in boundary zones of two obliquely converging plates are not directly related to the relative convergence rate vector, but to the two components parallel and sub-perpendicular to the boundary (e.g., Fitch, 1972; Walcott, 1978; Beck, 1983; Av6 Lallemant and Guth, 1990; McCaffrey, 1991; Cashman et al., 1992). Many deformation structures on Roat~in Island can be interpreted to be the result of partitioning of the convergence rate vector between the Caribbean and North American plates. Other structures seem to indicate that during oblique divergence the relative plate motion vector was partitioned as well.
D~ folds The original metamorphic ages of the rocks along the Motagua fault zone are Precambrian to Paleozoic (Gomberg et al., 1968), but they were overprinted during Late Cretaceous time as a result of the collision of the Chortfs block with the Maya block (Donnelly et al., 1990; Burkart, 1994). Although N- and S-vergent thrust faults, north and south of the Motagua fault zone, respectively, are the main results of the collision, mylonites resulting from left-lateral strike-slip displacement are ubiquitous in the fault zone (Meschede et al., 1993). We propose that the metamorphic structures along the Motagua fault zone can be correlated with the D~ structures on Roatfin Island. We believe that the convergence between the two blocks was left-oblique and that the thrust faults in the Chortfs and Maya blocks are related to the plate boundary-normal component of convergence, while the mylonites in the Motagua fault zone and the Bay Islands are related to the plate boundary-parallel vector component (Fig. 17a).
D2 folds The second generation of folds on Roatfin refolded the D~ folds and the mylonites. They are chevron-type folds with axes plunging moderately to
213 the west-southwest. They are related to NNW-SSE contraction and thus may have formed during the Late Cretaceous collisional event as well. The S1 cleavages on Roatfin dip moderately to the southsoutheast and the sense of shear is ENE over WSW. To bring these structures into coincidence with the D1 structures in the Motagua fault zone, the island must have been tilted to the south which is most easily done by back thrusting (Fig. 17b). Loading by this back thrust may have caused subsidence to the south and the birth of the Tela Basin.
D3 folds and F1 faults During the third folding phase kink folds developed that plunge moderately to the north-northwest and south-southeast. The F1 faults are N-trending thrust faults. Both folds and faults caused E - W contraction that may have taken place in a restraining bend along the Motagua fault zone (Fig. 17c). The F~ faults are the only mesoscopic thrust faults that we encountered, and although they may be related to the megascopic thrust fault that McBirney and Bass (1969a) proposed for the contact between the highand overlying low-grade metamorphic rocks, such contact is better explained as a low-angle normal fault (see below).
Unroofing The D2 and D3 deformations could have been responsible for uplift of the metamorphic rocks of Roatfin. Erosion, tectonic denudation, or both caused the exhumation. The uplift, exhumation, and cooling through the closure temperatures of hornblende (500 ~ to 550~ McDougall and Harrison, 1988) and zircon (fission track; ~280~ M. Brix, pers. commun., 1996) occurred in Late Eocene to Early Oligocene times (36 Ma 4~ hornblende and 39 Ma zircon fission track). The contact between the high-grade and the overlying low-grade metamorphic rocks has been interpreted by McBirney and Bass (1969a) as a thrust fault. We observed a few mesoscopic thrust faults near the contact consistent with their model. Generally, however, such contacts are interpreted as low-angle normal faults. We found many mesoscopic F3 normal faults with the correct orientation for a megascopic normal fault near the contact (Fig. 12c), and although most F3 faults are of Miocene and younger age (see below), tectonic denudation and exhumation of the metamorphic rocks cannot be ruled out.
F2 and F3 faults The F2 faults are a mostly conjugate set of strikeslip faults and the F3 faults a mostly conjugate set of normal faults (Fig. 12b,c, Fig. 13b,c). They are discussed together, because they both caused NNS-SSW to NNW-SSE extension (transtension).
214
The formation of F3 faults may be related to displacement partitioning. Heubeck and Mann (1991) have shown that most Caribbean plate motion models (e.g., Stein et al., 1988) applied to the Swan Islands fault zone imply extension along the fault zone (at least since 5 Ma). It is possible that the relative plate motion vector (possibly oriented E - W ) was partitioned into a component of left-lateral slip along the Swan Islands fault zone and a component perpendicular to the fault causing the formation of normal faults parallel to the Swan Islands fault zone on and just south of Roat~in Island. The WNW trend of many normal faults may be related to the WNW trend of the Honduran coast line (Fig. 17d). The F2 faults could be interpreted as tear faults. As the F3 faults are related to the major phase of subsidence of the Tela Basin, the age of the faulting is Miocene (vonder Hoya, 1986).
F4 and F5 faults The F4 and F5 faults both caused approximately E - W extension. The F4 phase consists of a mostly conjugate set of strike-slip faults, one set of which (the left-lateral one) parallels the Swan Islands fault zone. The F5 phase consists of a set of approx-
H.G. AVI~ L A L L E M A N T and M.B. G O R D O N
imately conjugate normal faults that strike NNW to NNE. It is the most penetrative fault system on Roatfin Island: e.g., Fig. 14 shows a representative E - W road section with numerous, approximately NNW-trending normal faults and tension gashes (see also Fig. 12e, locality No. 22) and Fig. 15 shows a road cut with E-W-trending extensile strain of at least 80% (see also Fig. 16c). Normal faults of this orientation can easily form near major strike-slip faults (e.g., Harding, 1974), but they tend to disappear at short distances from the strike-slip fault. However, similar normal faults occur across the Tela Basin in northern Honduras (Figs. 2 and 3). Furthermore, ENE-trending strike-slip faults extend toward the northern Honduran coastline, but are not shown to continue into the Tela Basin (Kozuch, 1991). We speculate that they do continue into the basin and that the E - W extension is related to left-stepping of the Swan Islands fault zone, and thus, that the Tela Basin during F4 and F5 faulting is a pull-apart basin (Fig. 17e). As the Swan Islands fault as well as the La Ceiba and Agu~in faults are seismically active, it seems that E - W as well as N-S extension is occurring at present.
Fig. 17. Tectonic model of deformation history of Roat~in Island and Tela Basin. (a) Late Cretaceous to early Tertiary NE convergence of Chortfs and Maya blocks was left-oblique resulting in displacement partitioning: Chortfs block is thrust northward along fault No. 1 upon Maya block and it moved eastward along E-W-trending proto-Motagua left-lateral strike-slip fault (fault No. 2) (D1 deformation). (b). Major southward tilting and folding (D2) of Roat~in Island segment of proto-Motagua fault zone caused by back thrusting along fault No. 3 which may have caused loading and subsidence and the origin of the Tela Basin; strike-slip displacements occur now along fault No. 4. (c). E-W contraction, possibly in restraining bend, resulted in N S-trending (D3) folds and (F1) thrusts. During the Late Eocene deformation D2 and/or D3 may have caused uplift followed by erosional exhumation of metamorphic and igneous rocks of Roat~in Island (MF = Motagua fault; PF = Polochic fault). (d). In Oligocene time, the relative displacement direction between the North American and Caribbean plates changed to approximately E-W causing left-oblique divergence along the Motagua (MF) and Swan Islands fault (SIF) zones. Displacement partitioning resulted in strike-slip displacements along these faults and normal faulting along parallel ENE-trending faults; WNWtrending normal faults formed along the Honduran coast line. Displacement along these normal faults caused the second phase subsidence of the Tela Basin. (e). The youngest and most pervasive brittle deformation on Roat~in Island is related to approximately E-W extension by a conjugate set of strike-slip faults (F4) and N-S-trending normal faults (Fs). The Rio Lenin graben (RLG) on the Honduran mainland may be related to F5 and as it terminates against the La Ceiba fault (LCF), the latter must be a left-lateral strike-slip fault. It is speculated that during the Miocene-Recent, the Tela Basin was lying in a releasing bend of the Swan Islands fault zone. Seismic activity along many of these faults indicate, however, that both E-W and N-S extension is still occurring today.
DEFORMATION HISTORY OF ROATAN ISLAND Tela Basin Previous work The Tela Basin lies between the Bonacca Ridge and mainland Honduras (Figs. 2-4). It is wedgeshaped narrowing toward the west. Few data are available in the public domain. In the 1960s to 1980s the petroleum industry carried out several geophysical surveys (gravity, magnetics, seismic reflection) and drilled test wells in the Tela Basin and on the mainland of Honduras (e.g., Caceres Avila et al., 1984). Although economic hydrocarbon deposits were not found, few data have been released. Pinet (1975) published several interpretations of N-S-trending seismic reflection lines. Von der Hoya (1986) showed several seismic reflection lines, however, without location maps. Paul Mann provided us with an unpublished E-W-trending seismic reflection line through the central part of the basin (with location map) as well as SeaMARC II side-scan sonar data. Magnetic studies were done by Pinet (1971) and vonder Hoya (1986). Bathymetric studies were carried out by Pinet (1976) and vonder Hoya (1986). The deepest part of the Tela Basin is a narrow E-W-trending trough which at present is the main depocenter. It is segmented by N-S-trending ridges (Figs. 3 and 4; Pinet, 1976). SeaMARC II side-scan sonar data of the central deep, south of Roat~n Island, indicate that the sea floor is quite smooth with only a few ENE-trending lineaments (Paul Mann, pers. commun., 1996). Some of these lineaments may be seismically active faults (see focal-mechanism plot of earthquake in the Tela Basin on Fig. 2 and Table 1). Pinet (1971) presented a magnetic anomaly map of the Tela Basin and suggested that three positive anomalies may be related to serpentinite bodies thrust onto non-magnetic sedimentary rocks. Based on the seismic reflection characteristics of an area where such magnetic anomaly occurs, von der Hoya (1986) suggested that they are better explained as being caused by basaltic intrusives, perhaps related to the basalts exposed on Utila Island. Seismic reflection lines were interpreted by Pinet (1975) and von der Hoya (1986) as indicating that the Tela Basin is underlain by a basement of metamorphic and plutonic rocks overlain unconformably by up to 5000 m of Tertiary sedimentary rocks. Von der Hoya (1986) proposed that in the early Tertiary the Tela Basin was formed by down-warping and that the basin was not controlled by faulting. He proposed that during the Oligocene (35 to 30 Ma) basement uplifts were formed in the axial portions of the Tela Basin along ENE trends. The older sedimentary rocks were faulted and tilted and eroded. The basin started subsiding in Middle Miocene time
215 and the tilted rocks were unconformably overlain by fine-grained clastic material (von der Hoya, 1986). At about 10 Ma deepening of the basin accelerated and turbidites were deposited in the deepest part of the basin (Pinet, 1976; von der Hoya, 1986). Structure and origin of the Tela Basin The only structures in the Tela Basin, discussed in the literature, are ENE- to WNW-trending normal faults (Fig. 3; Pinet, 1975; Paul Mann, pers. commun., 1996). The earthquake of July 27, 1990 (No. 9 in Figs. 3 and 4 and Table 1) in the Tela Basin might be the result of slip along such normal fault. The early history of the Tela Basin may have been 'transpressional' because it formed during the collision and overthrusting of the Chortfs block onto the Maya block. The back thrusting event that we proposed above (D2) may have caused loading and down-warping. The uplift of axial regions of the basin (von der Hoya, 1986) may have been related to late D2 folding as well. This period may have started in the Late Eocene (our new isotopic ages) and continued until the Middle Miocene ( v o n d e r Hoya, 1986). On the basis of the existence of ENE- to WNWtrending normal faults on Roatfin Island (F3) and in the Tela Basin as deduced from seismic reflection lines (Pinet, 1976; v o n d e r Hoya, 1986; Paul Mann, pers. commun., 1996), we suggest that in the Miocene the Tela Basin started to subside again due to N-S stretching related to transtension across the Swan Islands fault zone. The ENE trend of the faults is related to the ENE trend of the Swan Islands fault zone and the WNW trend to the WNW trend of the Honduran coast line. The most penetrative brittle deformation on Roatfin Island (Fs) caused major E - W extension along N-S-trending faults. These faults have to extend to the south until they abut against an E W-trending strike-slip fault. All approximately E W-trending faults in the Tela Basin have been identified as normal faults. The Rio Lefin graben on the mainland of Honduras (Fig. 3) trends NNE toward the Tela Basin in which, however, it has not been identified. Toward the south the Rio Le~n graben abuts against the ENE-trending La Ceiba fault which consequently must be a strike-slip fault. The La Ceiba fault has been mapped eastward until it reaches the coastline. We speculate that the La Ceiba fault continues eastward along the Honduran shelf into the Tela Basin (Fig. 17e). Furthermore, we speculate that the Rio Lefin graben continues northward (Fig. 17e) until it reaches the Swan Islands fault zone west of Utila Island. We also suggest that the ridge separating the two subbasins is a N S-trending horst. Thus, we propose that during the
216 F5 phase the Tela Basin was lying in a releasing bend of the Swan Islands fault zone. The areas where we propose the existence of normal and strike-slip faults in the Tela Basin have not been mapped by side-scan sonar and we do not have access to seismic reflection lines through these areas to test our speculative model. However, a similar deformation scheme that we propose for the Tela Basin has previously been proposed for another part of the plate boundary: Aldrich et al. (1991) described E-W-trending grabens cut across by N S-trending ones in northern Honduras, just south of the Motagua fault zone and the time of deformation is about the same as in the Tela Basin.
SUMMARY AND CONCLUSIONS
Three generations of folds (D1 to D3) and five generations of faults (F1 to Fs) were identified in the metamorphic rocks of Roat~in Island. The folding occurred in Late Cretaceous to early Tertiary and the faulting from early Tertiary to the Present. Displacement partitioning is an important factor when interpreting the structural evolution of Roat~in Island, because the deformation structures are generally not directly related to the relative convergence or divergence rate vectors between the North American and Caribbean plates, but to their components parallel and perpendicular to the plate boundary: the Swan Islands and Motagua fault zones. The first folding (D1) is synmetamorphic and related to mylonitization along the left-lateral Motagua and Swan Islands fault zones. The D1 structures are similar in style and orientation as those on Guanaja Island and along the Motagua fault zone in Honduras and Guatemala. N- and S-vergent thrusting in Honduras and Guatemala may be coeval with the formation of the mylonites. The second generation of folds (D2) is related to S-vergent thrusting and refolding. Loading by these thrust sheets may have caused subsidence and the origin of the Tela Basin. The D3 structures indicate E - W contraction possibly in a restraining bend along the Motagua fault zone. These two contractional phases (D2 and D3) resulted in uplift of the metamorphic rocks which were subsequently exhumed by erosion and possibly by tectonic denudation. New isotopic ages indicate that this event occurred in the Late Eocene. The first generation faults (F1) are N-S-trending thrust faults that resulted in E - W contraction. This phase may be coeval with the D3 folding. The second generation faults (F2: a mostly conjugate set of strike-slip faults) and the third one (F3: a mostly conjugate set of normal faults) are related to N-S extension. These structures indicate that the
H.G. AVt~ LALLEMANT and M.B. GORDON E-W-trending relative plate motion vector was partitioned into a N-S component (normal faulting) and an ENE-trending component (strike-slip on the Swan Islands fault zone). These normal faults caused renewed subsidence of the Tela Basin in Miocene time. An earlier generation of F3 faults may have been partly responsible for tectonic denudation and exhumation of the metamorphic rocks of Roat~in. The fourth generation of faults (F4: a conjugate set of strike-slip faults) and the fifth (F5: conjugate set of normal faults) resulted in E - W extension. The normal faults can be correlated with the NNE-trending Rfo Lenin graben on the mainland of Honduras. Toward the south this graben abuts against the La Ceiba fault which is a left-lateral strike-slip fault suggesting that Roat~in Island and the Tela Basin were lying at the time in a releasing bend of the Swan Islands fault zone. As the Swan Islands fault between Roatfin and the Motagua fault is seismically active, it seems that both E - W and N-S extension are still occurring at present.
ACKNOWLEDGEMENTS
This study was made possible by a grant from the National Science Foundation (EAR-9219384). We thank A.R. McBirney for air photos, maps, and comments; E. Phelgar, Defense Mapping Agency, Instituto Geogrfifico Nacional, and Direcci6n General de Minas e Hidrocarburos (Honduras) for logistical support; R.J. Phillips for supplying us with the thesis of v o n d e r Hoya; R Mann, for sharing with us unpublished material (side-scan sonar map and seismic reflection line) of the Tela Basin; J.T. Dutro Jr. for fossil identification; G.K. Merrill for helping us in our (unsuccessful) search for conodonts; Peter Copeland for the 4~ dating; Ann Blythe for the fission-track dating; Jinny Sisson for the amphibole analysis; and D. Chu for help with the focal-mechanisms and plate motions. We are especially grateful to Burke Burkart, Gren Draper, Pete Emmet, and Paul Mann for their extensive reviews of the original manuscript which improved it considerably. However, errors, mistakes, and misinterpretations are clearly our responsibility.
REFERENCES
Aldrich, M.J., Jr., Adams, A.I. and Escobar, C., 1991. Structural geology and stress history of the Platanares geothermal site, Honduras: implications on the tectonics of the northwestern Caribbean plate boundary. J. Volcanol. Geotherm. Res., 45: 59-69. Angelier, J., 1990. Inversion of field data in fault tectonics to obtain the regional stress, III. A new rapid direct inversion by analytical means. Geophys.J. Int., 103: 363-376.
D E F O R M A T I O N HISTORY OF ROAT,/~N I S L A N D Angelier, J., 1991a. Inversion directe et recherche 4-D: comparison physique et math6matique de deux modes de d6termination des tenseurs des pal6oconstraintes en tectonique de failles. C. R. Acad. Sci. Paris, Ser. II, 312: 1213-1218. Angelier, J., 199 lb. Analyse chronologique matricielle et succession r6gionale des 6v6nements tectoniques. C. R. Acad. Sci. Paris, Ser. II, 312 1633-1638. Arden, D.D., 1975. Geology of Jamaica and the Nicaragua Rise. In: A.E.M. Nairn and EG. Stehli (Editors), The Ocean Basins 3: The Gulf of Mexico and the Caribbean. Plenum, New York, pp. 617-662. Av6 Lallemant, H.G. and Guth, L.R., 1990. Role of extensional tectonics in exhumation of eclogites and blueschists in an oblique subduction setting. Geology, 18: 950-953. Bargar, K.E., 1991. Fluid inclusions and preliminary studies of hydrothermal alteration in core hole PLTG-1, Platanares geothermal area, Honduras. J. Volcanol. Geotherm. Res., 45: 147-160. Beck, M.E., Jr., 1983. On the mechanism of tectonic transport in zones of oblique subduction. Tectonophysics, 93: 1-11. Burkart, B., 1994. Northern Central America. In: S.K. Donovan and T.A. Jackson (Editors), Caribbean Geology: An Introduction. Univ. West Indies Publ. Assoc., Kingston, Jamaica, pp. 265-284. Burke, K., 1988. Tectonic evolution of the Caribbean. Annu. Rev. Earth Planet. Sci., 16:201-230. Burke, K., Cooper, C., Dewey, J.F., Mann, E and Pindell, J.L., 1984. Caribbean tectonics and relative plate motions. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162, 31-64 pp. Caceres Avila, E, Tappmeyer, D.M., Aves, H.S., Gillett, M. and Klenk, C.D., 1984. Recent studies of basins are encouraging for future exploration in Honduras. Oil Gas J., 82: 139-140. Calais, E. and Mercier de L6pinay, B., 1991. From transtension to transpression along the northern Caribbean plate boundary off Cuba: implications for the Recent motion of the Caribbean plate. Tectonophysics, 186: 329-350. Case, J.E. and Holcombe, T.L., 1980. Geologic-tectonic map of the Caribbean region. U.S. Geol. Surv. Misc. Invest. Map I- 1100, Scale 1 : 2,500,000. Case, J.E., MacDonald, W.D. and Fox, P.J., 1984. Caribbean crustal provinces; seismic and gravity evidence. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Vol. H, Geological Society of America, Golden, CO, pp. 15-36. Cashman, S.M., Kelsey, H.M., Erdman, C.E, Cutten, H.N.C. and Berryman, K.R., 1992. Strain partitioning between structural domains in the forearc of the Hikurangi subduction zone, New Zealand. Tectonics, 11: 242-257. Dean, B.W. and Drake, C.L., 1978. Focal mechanism solutions and tectonics of the Middle America arc. J. Geol., 8 6 : 1 1 1 128.
DeMets, C., Gordon, R.G., Argus, D.E and Stein, S., 1994. Effect of recent revisions to the geomagnetic reversal time scale on estimates of current plate motions. Geophys. Res. Lett., 21: 2191-2194. Deng, J. and Sykes, L.R., 1995. Determination of Euler pole for contemporary relative motion of Caribbean and North American plates using slip vectors of interplate earthquakes. Tectonics, 14: 39-53. Dengo, G., 1969. Problems of tectonic relations between Central America and the Caribbean. Trans. Gulf Coast Assoc. Geol. Soc., 19:311-320. Donnelly, T.W., Horne, G.S., Finch, R.C. and L6pez-Ramos, E., 1990. Northern Central America; the Maya and Chortfs blocks. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Vol. H, Geological
217 Society of America, Golden, CO, pp. 37-76. Dziewonski, A.M., Chou, T.-A. and Woodhouse, J.H., 1981. Determination of earthquake source parameters from waveform data for studies of global and regional seismicity. J. Geophys. Res., 86: 2825-2852. Finch, R.C., 1981. Mesozoic stratigraphy of central Honduras. Am. Assoc. Pet. Geol. Bull., 65: 1320-1333. Fitch, T.J., 1972. Plate convergence, transcurrent faults, and internal deformation adjacent to southeast Asia and the western Pacific. J. Geophys. Res., 77: 4432-4469. Gomberg, D.M., Banks, EO. and McBirney, A.R., 1968. Guatemala: preliminary zircon ages from Central Cordillera. Science, 162: 121-122. Gordon, M.B., 1991. The Chortfs block is a continental, pre-Mesozoic terrane. In: D.K. Larue and G. Draper (Editors), Transactions of the 12th Caribbean Geological Conference, St. Croix, U.S. Virgin Islands. Miami Geological Society, Miami, FL, pp. 505-512. Gordon, M.B., 1993. Revised Jurassic and Early Cretaceous (Pre-Yojoa Group) stratigraphy of the Chortfs block: Paleogeographic and tectonic implications. In: J.L. Pindell and B.F. Perkins (Editors), Mesozoic and Early Cenozoic Development of the Gulf of Mexico and Caribbean Region. Trans. 13th Annu. Res. Conf. Gulf Coast Sect., Society of Economic Paleontologists and Mineralogists Foundation, Austin, TX, pp. 143-154. Gordon, M.B. and Av6 Lallemant, H.G., 1995. Cryptic strike-slip faults of the Chortfs block. Geol. Soc. Am. Abstr. Progr., 27: 226-227. Gordon, M.B. and Muehlberger, W.R., 1994. Rotation of the Chortfs block causes dextral slip on the Guayape fault. Tectonics, 13: 858-872. Gose, W.A., 1985. Paleomagnetic results from Honduras and their bearing on Caribbean tectonics. Tectonics, 4: 565-585. Harding, T.E, 1974. Petroleum traps associated with wrench faults. Am. Assoc. Pet. Geol. Bull., 58: 1290-1304. Heiken, G., Ramos, N., Duffield, W., Musgrave, J., Wohletz, K., Priest, S., Aldrich, J., Flores, W., Ritchie, A., Goff, E, Eppler, D. and Escobar, C., 1991. Geology of the Platanares geothermal area, Departamento de Cop~n, Honduras. J. Volcanol. Geotherm. Res., 45: 41-58. Heubeck, C. and Mann, E, 1991. Geologic evaluation of plate kinematic models for the North American-Caribbean plate boundary zone. Tectonophysics, 191: 1-26. Holcombe, T.L., Vogt, ER., Matthews, J.E. and Murchison, R.R., 1973. Evidence for sea-floor spreading in the Cayman Trough. Earth Planet. Sci. Lett., 20: 357-371. Horne, G.S., Atwood, M.G. and King, A.E, 1974. Stratigraphy, sedimentology, and paleoenvironment of Esquias Formation of Honduras. Am. Assoc. Pet. Geol. Bull., 58: 176-188. Horne, G.S., Clark, G.S. and Pushkar, E, 1976a. Pre-Cretaceous rocks of northwestern Honduras: basement terrane in Sierra de Omoa. Am. Assoc. Pet. Geol. Bull., 60: 566-583. Home, G.S., Pushcar, E and Shafiqullah, M., 1976b. Laramide plutons on the landward continuation of the Bonacca ridge, northern Honduras. Publ. Geol. ICAITI, 5: 84-93. Kanamori, H. and Stewart, G.S., 1978. Seismological aspects of the Guatemala earthquake of February 4, 1976. J. Geophys. Res., 83: 3427-3434. Kozuch, M.J., 1991. Mapa Geol6gico de Honduras. Tegucigalpa, Honduras. Instituto Geogr~fico Nacional, Tegucigalpa scale 1 : 500,000.
Land, L., 1979. The fate of reef-derived sediment on the north Jamaican Island slope. Mar. Geol., 29: 55-71. Lawrence, D.E, 1975. Petrology and Structural Geology of the Sanarate-E1 Progreso Area, Guatemala. Ph.D. Thesis, Univ. New York, Binghamton, NY, 255 pp. Leroy, S., Mercier de L6pinay, B., Mauffret, A. and Pubellier,
218 M., 1996. Structural and tectonic evolution of the eastern Cayman Trough (Caribbean Sea) from seismic reflection data. Am. Assoc. Pet. Geol. Bull., 80: 222-247. MacDonald, K.C. and Holcombe, T.L., 1978. Inversion of magnetic anomalies and sea-floor spreading in the Cayman Trough. Earth Planet. Sci. Lett., 40:407-414. Mann, E and Burke, K., 1984. Cenozoic rift formation in the northern Caribbean. Geology, 12: 732-736. Mann, E, Tyburski, S.A. and Rosencrantz, E., 1991. Neogene development of the Swan Islands restraining-bend complex, Caribbean Sea. Geology, 19: 823-826. Manton, W.I., 1987. Tectonic interpretation of the morphology of Honduras. Tectonics, 6: 633-651. McBirney, A.R. and Bass, M.N., 1969a. Geology of Bay Islands, Gulf of Honduras. Am. Assoc. Pet. Geol. Mem., 11: 229-243. McBirney, A.R. and Bass, M.N., 1969b. Structural relations of Pre-Mesozoic rocks of northern Central America. Am. Assoc. Pet. Geol. Mem., 11: 269-280. McCaffrey, R., 1991. Slip vectors and stretching of the Sumatran forearc. Geology, 19: 881-884. McDougall, I. and Harrison, T.M., 1988, Geochronology and Thermochronology by the 4~ Method. Oxford Univ. Press, New York, 212 pp. Meschede, M., Ratschbacher, L. and Frisch, W., 1993. Kinematic information from fault-slip data in southern Mexico and along the Motagua-Polochic fault system in Guatemala. In: E Ortega-Guti6rrez, EJ. Coney, E. Centeno-Garcfa and A. G6mez-Caballero (Editors), Transactions of the 1st CircumPacific and Circum-Atlantic Terrane Conference, Univ. Nac. Aut6n. M6xico, Inst. Geol., pp. 81-85. Mills, R.A. and Barton, R., 1996. Geology of the Ahuas area in the Mosquitia Basin of Honduras. Am. Assoc. Pet. Geol. Bull., 80:1627-1640. Mills, R.A., Hugh, K.E., Feray, D.E. and Swolfs, H.C., 1967. Mesozoic stratigraphy of Honduras. Am. Assoc. Pet. Geol. Bull., 51: 1711-1786. Molnar, E and Sykes, L.R., 1969. Tectonics of the Caribbean and Middle America regions from focal mechanisms and seismicity. Geol. Soc. Am. Bull., 80: 1639-1684. Olson, E.C. and McGrew, EO., 1941. Mammalian fauna from the Pliocene of Honduras. Geol. Soc. Am. Bull., 52: 1219-1244. Osiecki, ES., 1981. Estimated intensities and probable tectonic sources of historic (pre-1898) Honduran earthquakes. Bull. Seism. Soc. Am., 71: 865-881. Perfit, M.R., 1977. Petrology and geochemistry of mafic rocks from the Cayman Trench: evidence for spreading. Geology, 5: 105-110. Petit, J.E, 1987. Criteria for the sense of movement on fault surfaces in brittle rocks. J. Struct. Geol., 9: 597-608. Pindell, J.L., 1993. Regional synopsis of Gulf of Mexico and Caribbean evolution. In: J.L. Pindell and B.E Perkins (Editors), Mesozoic and Early Cenozoic Development of the Gulf of Mexico and Caribbean Region. Trans. 13th Annu. Res. Conf. Gulf Coast Sect., Society of Economic Paleontologists and Mineralogists Foundation, Austin, TX, pp. 251-274. Pindell, J.L., Cande, S.C., Pitman, W.C., Rowley, D.B., Dewey, J.E, LaBrecque, J. and Haxby, W., 1988. A plate-kinematic framework for models of Caribbean evolution. Tectonophysics,
H.G. AVt~ L A L L E M A N T and M.B. G O R D O N 155: 121-138. Pinet, ER., 1971. Structural configuration of the northwestern Caribbean. Geol. Soc. Am. Bull., 82: 2027-2032. Pinet, ER., 1975. Structural evolution of the Honduras continental margin and the sea floor south of the western Cayman trough. Geol. Soc. Am. Bull., 86: 830-836. Pinet, ER., 1976. Morphology of northern Honduras, northwestern Caribbean Sea. Deep-Sea Res., 23: 839-847. Plafker, G., 1976. Tectonic aspects of the Guatemala earthquake of 4 February 1976. Science, 193: 1201-1208. Rosencrantz, E. and Mann, E, 1991. SeaMARC II mapping of transform faults in the Cayman Trough, Caribbean Sea. Geology, 19: 690-693. Rosencrantz, E. and Sclater, J.C., 1986. Depth and age in the Cayman Trough. Earth Planet. Sci. Lett., 79: 133-144. Rosencrantz, E., Ross, M.I. and Sclater, J.G., 1988. Age and spreading history of the Cayman Trough as determined from depth, heat flow, and magnetic anomalies. J. Geophys. Res., 93: 2141-2157. Schwartz, D.E, Cluff, L.S. and Donnelly, T.W., 1979. Quaternary faulting along the Caribbean-North American plate boundary in Central America. Tectonophysics, 52:431-445. Southernwood, S., 1986. Late Cretaceous Limestone Clast Conglomerates of Honduras. MS Thesis, Univ. of Texas, Dallas, 299 pp. Stein, S., DeMets, C., Gordon, R.G., Brodholt, J., Argus, D., Engelen, J.E, Lundgren, E, Stein, C., Wiens, D. and Woods, D.E, 1988. A test of alternative Caribbean plate relative motion models. J. Geophys. Res., 93: 3041-3050. Steiner, M.B. and Walker, J.D.,1996. Late Silurian plutons in Yucatan. J. Geophys. Res., 101: 17,727-17,735. Twiss, R.J. and Gefell, M.J., 1990. Curved slickenfibers: a new brittle shear sense indicator with application to a sheared serpentinite. J. Struct. Geol., 12: 471-482. Tyburski, S.A., 1992. Deformational Mechanisms along Active Strike-Slip Faults: SeaMARC II and Seismic Data from North America-Caribbean Plate Boundary. MA Thesis, Univ. of Texas, Austin, 195 pp. vonder Hoya, H.A., II, 1986. A Reflection Seismic and Magnetic Investigation of the Tela Basin: Northern Offshore Honduras. MS Thesis, Southern Methodist Univ., Dallas, Texas, 112 pp. Wadge, G. and Wooden, J.L., 1982. Late Cenozoic alkaline volcanism in the northwestern Caribbean: tectonic setting and Sr isotopic characteristics. Earth Planet. Sci. Lett., 57: 35-46. Walcott, R.I., 1978. Geodetic strains and large earthquakes in the axial tectonic belt of North Island, New Zealand. J. Geophys. Res., 83: 4419-4429. White, R.A., 1991. Tectonic implications of upper-crustal seismicity in Central America. In: D.B. Slemmons, E.R. Engdahl, M.D. Zoback and D.D. Blackwell (Editors), Neotectonics of North America. The Geology of North America, Decade Map Vol., Geological Society of America, Golden, CO, pp. 323338. Williams, H. and McBirney, A.R., 1969. Volcanic history of Honduras. Univ. California Publ. Geol. Sci., 85: 1-101. Wilson, H.H., 1974. Cretaceous sedimentation and orogeny in Nuclear Central America. Am. Assoc. Pet. Geol. Bull., 58: 1348-1396.
Chapter 9
The Southern Flank of the Tela Basin, Republic of Honduras
W.I. MANTON and R.S. MANTON
A wide strike-slip zone is formed at the northern margin of the Caribbean plate. Its northern portion, the Tela Basin, is submerged, but the north coast of Honduras, from Tela to Trujillo, is characterized by high, fault-bounded mountains. Shear zones exposed in drainages attest to a long history of faulting. A feature of some interest is the Agu~n fault, which lies south of the mountains, and today forms the southern boundary of the zone. It curves through 35-45 ~ and is in the west transpressional. In the east, however, it is transtensional and gives rise to the wide valley of the Agu~n River, which is probably a 'lazy' S-type of pull-apart basin. Poorly bedded conglomerates of unknown age outcropping at the eastern end of the valley may represent part of its infilling with fluvial sediments. Also at its eastern end are near-vertical mudstones and shales with thin interbeds of limestone, that were deposited in a hemipelagic to pelagic environment. Poorly preserved globeriginids suggest that the rocks are Middle Miocene or younger. About 50 km south of this occurrence, on the road between Trujillo and San Esteban, are exposures of rocks with a similar lithology that have been thrust from the northwest. They contain poorly preserved fossils, some of which may be Tertiary, and from both their environment of deposition and their direction of transport are correlated with the rocks of the Agu~n Valley. Both sequences are cut by mafic dikes, probably intruded before deformation. The presence of such rocks suggests that in the Miocene the north coast of Honduras was submerged with the formation of deep, turbidite-filled basins. When the region was uplifted some of these rocks were preserved by transtensional faults; others were deformed into flower structures by transpressive faults and were thrust to the south.
INTRODUCTION
High mountains whose northern slopes are covered with dense tropical vegetation, wide valleys filled with alluvium, few roads, and a geology consisting of a monotonous succession of poorly exposed phyllites cut by plutons or overlain by deeply weathered lavas have discouraged geological mapping of the north coast of Honduras, even though the region lies on the margin of a major strike-slip (Fig. 1) boundary that has accommodated more than 1000 km of movement during the Tertiary (Pindell and Barrett, 1990). Satellite images, however, show with clarity the great arcuate faults that shape the topography (Fig. 2), and several papers interpreting them have been written (Muehlberger, 1976; Letouzey, 1985; Manton, 1987; Gordon and Muehlberger, 1994). Apart from the mapping of the westernmost part of the coast by the United Nations (part of their map is reproduced in Home et al., 1976a), published geologic observations are restricted to those made by Sapper (1905) along
colonial roads that for the most part no longer exist, by Powers (1918) and Foye (1918), and by Roberts and Irving (1957), Williams and McBirney (1969) and Manton (1987) along the present-day roads. In contrast, the offshore region is much better known. The Bay Islands have been mapped (McBimey and Bass, 1969), as have the Swan Islands (Ivey et al., 1980), and several offshore wells have been drilled to the Mesozoic basement (Caceres Avila et al., 1984). Marine geophysics has detailed the topography, the sediment-filled basins, and the principal faults of the submerged part of the plate boundary (Pinet, 1975, 1976; Rosencrantz and Mann, 1991; Mann et al., 1991). The area between the Bay Islands and the mainland is known as the Tela Basin, and the geology of its northern margin is discussed elsewhere in this volume (Av6 Lallemant and Gordon, Chapter 8). Here we discuss the geology of its southern margin and elaborate on the occurrences of both continental and marine sediments on mainland Honduras that were mentioned but not described by Manton (1987).
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 219-236. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
220
.
.
.
.
W.I. M A N T O N and R.S. M A N T O N
.
.
'
~dw
'
86'w
'
Figs:3&4~
e4'w
'
17N -
/
15N ' I
15N -
HONDURAS
-13N ~ l ~ ,
....., 8B.w
....~
_//',_..j,,~-'"'-
~ ,
\
'%
~t .... e6w _,,_,
,
o
4,.
-
13N-
-
Fig. 1. Map of portion of Central America showing Honduras and the locations of the Landsat image (Fig. 2) and of the area discussed in this paper (Figs. 3 and 4).
THE NORTH
COAST:
PHYSIOGRAPHY AND A C C E S S
The coast of Honduras between Omoa and Trujillo consists of a narrow plain backed by mountain ranges, the Sierra de Omoa behind Puerto Cortez and the Sierra Nombre de Dios between Tela and Trujillo (Fig. 3). The ranges are breached between Omoa and Tela by the Sula Valley and south of La Ceiba by the gorge of the Cangrejal River. To the east of La Ceiba the Cordillera Nombre de Dios narrow, and a low saddle is eroded in them. South of the ranges lies the valley of the Agu~in, a broad, flat-floored depression, that curves gently northward to meet the Caribbean coast east of Trujillo. The town of San Pedro Sula, which lies in the Sula Valley, is the second city in Honduras and has an international airport with car rental agencies. The north coast is reached by driving to E1 Progreso on the east side of the valley and taking the tarred road that runs northeastwards to Tela (Fig. 3). This road follows the coast for 90 km before turning inland and crossing over the saddle into the Agu~in Valley, the southern margin of which it follows for 50 km before terminating at Trujillo. At the east end of the valley an unpaved road leads southwards through the town of Bonito Oriental to the Sico River and thence to E1 Carbon and San
Esteban. From E1 Progreso (Fig. 3) the tarred road continues south to Santa Rita where it turns east to Yoro. Yoro and the western end of the Agmin Valley are linked by an unpaved road that passes through Olanchito to meet the Jutiapa-Trujillo road at Saba. Distances are: San Pedro Sula-Tela, 99 km; Tela-La Ceiba, 103 km; La Ceiba-Trujillo, 171 km.
PRINCIPAL FEATURES OF THE STRIKE-SLIP MARGIN
The kinematics of the North AmericanCaribbean (NOAM-CARIB) plate boundary have proved particularly difficult to determine, both because the data are few and because the calculated rotation parameters are strongly biased by the initial choice of data. While the models of MacDonald (1976), Minster and Jordan (1978) and Stein et al. (1988) predict fairly well the azimuths of bounding Swan Islands and Oriente faults, those of Sykes et al. (1982) and of the NUVEL-1 model do not, although the latter can be brought to agreement if the Lesser Antilles data are omitted from the inversion (DeMets et al., 1990). Heubeck and Mann (1991), however, have shown that even the models yielding good fits fail to predict the pattern of transpression
Fig. 2. (A) Landsat image of the north coast of Honduras in the region of La Ceiba and Trujillo. Honduras P18 R49, 19 December, 1973. Color composite 1, 2, 3. (B) Some features of the Landsat image. C.V. = Cangrejal Valley; A.B. = Agu~in beds, Agu~in Valley locality. Linear features in the image parallel to the strike of the beds suggest that they extend for several kilometers southwest of where they are exposed by the road. C.T. = Cerro de Tarros (see Fig. 6). Triangle = Pico Bonito, 2643 m. D = possible dolerite dikes occupying deep linear valleys (within the Cangrejal Valley the relationship between dikes and valleys has been observed in the field); V = linear valleys of unknown origin; L = linear features of unknown origin; R.S.L. = Agu~in beds, Rio Sico locality. For reference, the approximate location of the Swan Islands fault is shown as is the NUVEL-1 NOAM-CARIB plate boundary vector (DeMets et al., 1990).
THE SOUTHERN FLANK OF THE TELA BASIN, HONDURAS
221
~u
o~
~
Roatan
I (~Utila LA
CEIBA
/
///"~2 ~ ._........_
L
222
W.I. MANTON and R.S. M A N T O N 1
;
I 88W
I
~
871W
'
i
86]W
/ o
20
40
60
. . . . . . . .
:
8,0
loo
I
KILOMETRES
ISLANDS
~GUANAJA
~UTI
LA ,~ o
- 16 N ~
~
C
O
R
T
E
~.~,o,~~./
s
Z
~
.
~'J
~.~ _.sAN , " ~ ~
~~ ~)
-i~~~o~uon ' ~ " ~0
/
,.,...
,o
/
J.OSta. 4 \ !
H 0 N D U R,4 S
88W 1
..... DE LA PAZ
L.Yojoo .J T Tegucigalpa i
16N -
~
j
/ ,,,.,.,,,...,,,'.. r( ,
~
To Jutiopo
...
I.
87W I
86W I
I
I
Fig. 3. Location map showing principal towns, roads, mountains and rivers.
and transtension observed along the plate boundary, and this observation has led them to divide the Caribbean plate into three segments each having slightly different poles of rotation. An interpretation of the geology of the north coast has been given by Manton (1987) whose map is reproduced with some revisions in Fig. 4. Principal changes are: (1) the addition of the Precambrian (Manton, 1996); (2) the change of the age of the Agu~n beds from Late Cretaceous to Tertiary (this paper, below); (3) the change of the age of the volcanics between La Ceiba and Olanchito from Cretaceous to Tertiary to make them consistent with the geological map of Honduras (Kozuch, 1991); and (4) the addition of faults to each side of the Agu~n Valley to make the feature a lazy S pull-apart basin (Mann et al., 1983). Offshore The present-day boundary between the North American and Caribbean plates is an active fault that appears on side-scan sonar images as a well defined linear feature extending from the Motagua Valley of Guatemala to the Swan Islands and closely following the base of the southern wall of the Cayman Trough (Rosencrantz and Mann, 1991). At the Swan Islands the fault fight steps to a parallel fault that continues to the Cayman spreading centre. Stresses associated
with a gentle restraining bend that previously existed between the two faults produced the broad anticlinal uplift upon which the islands lie (Mann et al., 1991). Flanking the southern margin of the western portion of the trough is a discontinuous series of en-echelon ridges (the Bonacca Ridge of Banks and Richards, 1969) which break sea level to form the Bay Islands of Roatan and Guanaja and elsewhere rise to within 1000 m of the surface. Between these ridges and the Honduran shelf break lies a triangular area composed of ridges and deep, fault-bounded, sedimentfilled basins elongated in the direction of the plate boundary (Pinet, 1976). On seismic reflection profiles Pinet (1975) identified an eroded pre-Eocene or Cretaceous surface unconformably overlain by two units, which from their widespread distribution and uniform thickness he ascribed to deposition during a tectonically quiet interval during the middle Tertiary. He assigned a Pliocene age to the faulting that produced the present bottom topography of ridges and basins and argued that the sediments filling the basins are turbidites, at the same time noting that some are faulted. Pinet (1971) recorded three magnetic anomalies, one between Roatan and the mainland, and the others off the mouth of the Agu~n and off the coast at longitude 84~ 30". These he attributed to en-echelon offsets of an elongate mass of ultramafic rock. In addition, Pinet (1972) drew attention to piercement structures lying off the coast,
THE SOUTHERN FLANK OF THE TELA BASIN, HONDURAS / / Basalt flows and cones
Alluvium
Aguan beds, ~-~ marine shale ~ limestone
Fluvial c o n g l o m e r a t e s unnamed beds
\871W
50o
223 /I
861W
~~176
~.~OU %
y 4000 ~
Andesitic volcanics -- ~
~o~
Cretaceous sediments undifferentiated
Schists, phyllites 8~ marbles l.OGa metamorphosed quartzo- felspathic rocks
-r
Granitic plutons of various
ages
\
CASTILLA + No I CASTANA
/
Nol
--16N
Geology
Unknown
Geology On now,,o
~'~b.
: < ~ ,;~; v ~
. . ~ - / _ \ Z _ g.,"4 Tv ..17 ~.'-~ t,t. >~-.r n
-15N
I
88 W I
I
87W I
I
86W I
Fig. 4. Geological sketch map of the region shown in Fig. 2, modified from Manton (1987). The earthquake solution from the western portion of the Agufin fault is from Gordon and Muehlberger (1994), and the arrows are the directions of the NUVEL-1 NOAM-CARIB plate boundary vector (DeMets et al., 1990). The vector of the 'alternative model' would be within 2~of the azimuth of the Swan Islands fault.
which coincided with the western and eastern magnetic anomalies but were not observed off the Agu4n mouth. He postulated that they were salt domes. Von der Hoya (1986) interpreted propriety seismic and magnetic data obtained between Roatan and the coast and concluded that the magnetic anomalies and the piercement structures were manifestations of one or more laccolith-shaped igneous bodies that reached to within 1.5 km of the top of the sediments. He argued that they were intruded in the Quaternary. The island of Utila is partly formed by a Quaternary volcano, and as such may be a manifestation or the same activity that produced the intrusions discovered by Von der Hoya (1986), but Roatan and Guanaja are composed of metamorphic rocks similar to those found in Guatemala (McBirney and Bass, 1969). The Swan Islands (Fig. 1), on the other hand, are composed of Oligocene to Lower Miocene turbidites containing a large amount of volcanic glass (Ivey et al., 1980). Two wells, Castilla No. 1 and Castana No. 1, have been drilled on the continental shelf northeast of Trujillo (see Caceres Avila et al., 1984, and Fig. 4). Both penetrated approximately 3000 m of Tertiary sediments before bottoming in red to brown shales, siltstones, and sandstones correlated with the Upper Cretaceous Valle de Angeles group (Aves, 1983). The lowermost units consisted of 670-820 m
of Middle Miocene shales and white quartzose sands interpreted as being deposited in a neritic environment. These may be the widespread Middle Miocene units of Pinet (1975). They are unconformably overlain by fine-grained clastic sediments deposited in ever increasing depths.
Onshore The strike-slip nature of the north coast is spectacularly displayed by the Cordillera Nombre de Dios which abruptly rise from the narrow coastal plain and attain their highest point behind La Ceiba, where the bare, pointed peak of Pico Bonito rises to 2435 m (Fig. 5). On the Landsat images (Fig. 2) the northern flank of the cordillera is demarcated by a sharp line that extends from the Sula Valley to east of La Ceiba and is clearly a recent fault, called by Muehlberger (1976) the La Ceiba fault. Southeast of La Ceiba is an equally strong linear feature that marks the Rfo Viejo fault, and between these two faults the cordillera have been recently uplifted (Manton, 1987). (For the radar image of this region, see Gordon and Muehlberger, 1994.) East of La Ceiba, the La Ceiba fault makes a 45 ~ clockwise turn and almost intersects the Rio Viejo fault, but thereafter the trace of both faults becomes unclear, with the La Ceiba fault probably turning offshore
224
W.I. MANTON and R.S. MANTON
Fig. 5. The Cordillera Nombre de Dios photographed from La Ceiba airport. The pointed peak in the backgroundis Pico Bonito, 2643 m. and the Rfo Viejo extending into a deep linear valley in the eastern part of the cordillera. The present-day faults closely follow the traces of older faults. For example, in Quebrada Juana Leandra, 11 km east of the bridge over the Cangrejal at La Ceiba, a waterfall exposes a ductilely deformed shear zone containing large boudins of the granitic country rock. It is about 100 m wide and strikes at 82 ~ with a 60N dip. Deformed leucocratic shear veins indicate that it is a high-angle thrust fault with upthrow to the north. It is cut by a highly fractured and altered dolerite dike of unknown age. The Cangrejal River, which cuts through the cordillera south of La Ceiba, reveals more of the long history of tectonism experienced by the region. About 15 km south of the town, a gneissose tonalite possessing a well developed C-S fabric indicative of dextral slip (Manton and Manton, 1989) is intruded by an undeformed granodiorite. Both rocks have been dated by Home et al. (1976b). Conventional K-Ar ages on hornblende gave 72.2 Ma for the deformed tonalite (their Piedras Negras tonalite) and 57.3 Ma for muscovite from the undeformed granodiorite (their Las Mangas tonalite). The age obtained from biotite from the Piedras Negras tonalite was 57.3 Ma. To the south the Piedras Negras tonalite passes into a 900-m-wide zone of mylonite. (For a general view of this rock, see frontispiece, Tectonics, 9 (2), 1990.) In contrast to that
of the tonalite, its sense of slip is sinistral (W.I. Manton, unpubl, information). The southern portion of the mylonite is developed in a granite of unknown age, which contains pseudo-tachylite veins up to 5 cm wide. Whether these were produced during the uplift of the cordillera along the Rfo Viejo fault or whether they belong to some earlier event is unknown. At some time the region experienced northsouth extension accompanied by the intrusion of a swarm of dolerite dikes along the N60E strike of the gneissosity of the Piedras Negras tonalite. These dikes weather more easily than the surrounding rock and give rise to the deep, narrow valleys that are conspicuous on the Landsat image (Fig. 2). Similar closely spaced, deep valleys occur in the cordillera further west in the vicinity of Pico Bonito. Although they have a more easterly orientation it seems likely that they were produced by weathering of a related swarm of dikes. All that is known of their age is that they are younger than the mylonite. They appear to be absent from the eastern end of the cordillera because the Landsat image of the region (Fig. 2) shows none of the deep linear valleys that would reveal their presence. It may be significant that they are restricted to the part of the coast opposite the volcanic island of Utila and the intrusions discovered by Von der Hoya (1986). The Agu~in Valley has by virtue of its length, breadth, and curvature been long regarded as the
225
T H E S O U T H E R N F L A N K OF T H E T E L A BASIN, H O N D U R A S
2'o
~.~
21 LAGUNA
6
3
|
2'2
23
2;
'1
2'6
28
~
~2 FAULT BR E C C IA
61
c,,,o
63Lo Piedro
-,',~,_~.__~. . . . .
61 -
~ AGUAN BEDS ~ c e ~
GRANITE 1 QUARR'Y
60-
J~;~~ ~ _~_/~
59-
58Monte Cristo
57-
CONGLOMERATES AND SANDSTONES 5655 CARIBBEAN
:/
SEA
,/
54
54-
Wta
TRUJ I
5:5
-'-----~ Fig6~L__L,L.
53-
_
52
52-
20
21
22
I
I
I
2~ I
24 I
25 I
26 I
27 I
28 I
Fig. 6. Location of outcrops at the eastern end of the Agu~n Valley.
site of a major strike-slip fault (Muehlberger, 1976), but only on its southeastern flank is it bounded by a youthful looking scarp. Elsewhere its walls are dissected and eroded back as if the faults that caused them have not been active for some time (Fig. 2), although several hot springs do occur along its length (Manton, 1987). The direction of its continuation to the west is unknown as there is no means of access to the region. Manton (1987) drew it following the course of the Guiamas River. If this extrapolation is correct the fault changes direction by 35-45 ~ along its length, so that its western portion, by virtue of its orientation with respect to the NOAM-CARIB rotation vector (Fig. 4), is in transpression while its eastern portion is in transtension. This interpretation is supported by an earthquake with a strong thrust component plotted by Gordon and Muehlberger (1994) on its presumed western extension (Fig. 4). On the other hand, the wide floor of the eastern part of the valley and the generally low relief of the surrounding region are typical
of a transtensional regime, and the valley itself is probably a 'lazy' S-type of pull-apart basin (Mann et al., 1983). See also Heubeck and Mann (1991). There are no published geophysical data to confirm this interpretation, but at its eastern end are low hills (Fig. 6) made up of southward-dipping coarse sandstones and quartz-pebble conglomerates (Fig. 7) which may represent part of a sedimentary filling. Some of these rocks are exposed in a road cut through Cerro La Penita and, because they are the only known examples of their kind in the Agu~in Valley, they are described below. The cut exposes 250 m of conglomerate with minor amounts of sandstone and shale that strike 4060 ~ and dip south at 36 ~. The first rock encountered at the northern end of the cut is a coarse-grained quartz sandstone which grades upwards into an indurated, weakly bedded or massive quartz-pebble conglomerate. The rock is clast-supported and consists of white vein quartz pebbles up to 4.5 cm in greatest dimension set in a coarse-grained sand or
226
W.I. MANTON and R.S. MANTON
Fig. 7. Northward-dipping, poorly bedded conglomeratic section exposed at Cerro La Penita, eastern end of the Agu~in Valley. The apparent decrease in dip results from distortion brought about by a wide-angle lens.
pebble matrix. With increasing height in the section, the clast composition begins to include gray limestone and a small quantity of brown sandstone; then the conglomerates become matrix-supported and sandstone clasts up to 15 cm appear. The conglomerates at the top of the section are massive and vary from matrix- to clast-supported. Clasts include several varieties of sandstone and weakly indurated claystone which could not have been transported a great distance. Shales in the lower part of the section are for the most part gray, becoming maroon in the upper portion, and may be associated with sandstones. Their most characteristic feature, however, is the way they have been diapirically deformed by the heavier, rapidly loaded conglomerates. Soft-sediment structures produced in the deformed shales include large-scale protrusions into the base of the overlying conglomerates, and a spectacular balloonshaped diapir (Fig. 8) forcibly intruded into the overlying beds. The coarseness of the sediments, the lack of bedding, and the rapidity of loading are all characteristic of alluvial deposition at the margin of a rapidly deepening fault-generated valley. Today the Agu~in Valley is the southern limit of the strike-slip margin, but in the vicinity of E1 Carbon (Fig. 4) the road crosses a major fault zone developed in a granitic pluton (Southemwood,
1986). The undeformed rock is coarse-grained and composed of large phenocrysts of pink orthoclase set against a matrix of white plagioclase and quartz. Biotite, however, is pervasively altered to green chlorite. As the shear zone is traversed, the granite becomes progressively deformed, first cataclastically, then ductilely until the quartz has flowed into stringers with aspect ratios between 5:1 and 10: 1. The orthoclase is also extended by about 5:1 but has done so brittlely by slipping along closely spaced microfaults. The zone, which is about 3 km wide, strikes at 55 ~ and the minerals are stretched in the vertical direction, implying the motion was dominantly dip-slip. Zircon separated from the undeformed granite is strongly discordant with inter+15 and 30 +6 cepts at 127_13 _ Ma (Appendix). The lower intercept, which falls in the Oligocene, may reasonably be interpreted as resulting from Pb loss during hydrothermal activity associated with the sheafing, raising the possibility that in the Miocene the strikeslip margin extended as far south as E1 Carbon.
THE AGUAN BEDS: POSSIBLE TERTIARY MARINE SEDIMENTS
Manton (1987) has already reported the presence of marine sediments in northern Honduras, naming
THE SOUTHERN FLANK OF THE TELA BASIN, HONDURAS
227
Fig. 8. Large balloon-shaped diapir of shale and thinly bedded sandstones disrupting conglomerates at Cerro La Penita. For scale, see man standing in the middle of the diapir. them the Agufin beds. They are well exposed in two localities, the one in road cuts through low hills (Fig. 6) in the Agufin Valley, southeast of Trujillo, and the other in a deep cut on the road between Trujillo and E1 Carbon, near the Sico River (Fig. 4). At each locality sections were measured, thin sections were made for petrographic study, the mineralogy of fine-grained samples was determined by X-ray diffraction, and interbedded limestones were analyzed for Sr isotope ratios with a view to determining their age by comparison with published seawater curves. Samples containing microfossils were submitted to micropaleontologists for identification.
Agu~n Valley locality Cuts made in a succession of low hills that straddle the road 4.5 km north of Cerro La Penita (Fig. 6) expose steeply dipping, thinly bedded mudstones and shales and thicker limestone units, intruded by deeply weathered basic dikes (Fig. 9). The total length exposed in outcrop is about 520 m. The exposures end on the north against a thickly bedded, green sandstone containing abundant spheroidal concretions. The contact is not exposed, nor does the rock resemble any in the measured section. There is no evidence to suggest duplication of the sequence by folding. Faults may exist, given gaps in the sec-
tion between the hills. Primary sedimentary geopetal structures were searched for in the field without success, but piercement structures associated with a dike suggest that the section youngs to the south. The rocks lie structurally beneath the shales and sandstones of Cerro La Penita, but the contact is obscured. Indurated, green-gray to dark tan, silty siliceous mudstones and thin bands of fissile gray shales dominate the lower part of the section (Fig. 9). The mudstones may be massive or may have planar lamination. Beds generally have thicknesses between 2.5 and 5 cm, with some units ranging up to 20 cm. Many mudstone beds have a pinch and swell appearance, but it is unclear whether this is a primary feature. Thin- to medium-bedded, black (but gray weathering), argillaceous, micritic limestones and interbedded mudstones dominate the upper half of the section. A few beds of coarser-grained calcarenite are also present. Planar lamination is the most common sedimentary structure in the limestones. In places, however, laminae appear to be disturbed, giving the rocks a mottled appearance. En-echelon, calcite-filled microfractures are found in some beds and thicker veins filled by sparry calcite infill fractures in the more thickly bedded limestones such as S- 138 (see Fig. 10).
228
W.I. M A N T O N and R.S. M A N T O N
Fig. 9. View of the mudstones and shales of the Agu~in beds exposed at the Agu~in Valley locality. South is to the right of the photograph.
The mudstones are too fine-grained for their composition to be resolved microscopically. In thin section they vary from silty packstones (S-114) to recrystallized pelagic foraminiferal wackestones (S-144). Bioclastic and organic content varies as does quartz content. Sedimentary structures were not generally recognized in thin section, though some samples contained darker layers of organic material. Compaction and broken bioclastic grains were frequently seen, and miniature flame structures were found in S-126. X-ray diffraction showed the mudstones to be composed predominantly of quartz with minor amounts of muscovite and calcite. The limestones vary from micrites to packstones depending on fossil content and several samples contain abundant globigerinids (Figs. 10 and 11) some of which may be Orbulina (Tony Eva, written commun.). In samples such as S-139, the fossils have been compacted and flattened. The composition of the limestones as determined by X-ray diffraction is calcite and quartz. Pink and yellow weathering, very fine-grained material weathering to clay, or powdery, loosely consolidated material was encountered at places in the section and may be layers of volcanic ash. Unfortunately, none of the friable material was collected for X-ray analysis.
The strike of the rocks is about 60 ~. They extend northeastwards as far as Cerro Las Lomas, 2 km from the road. They have not been traced to the west, but linear features on the Landsat image (Fig. 2) suggest that they continue in that direction for several kilometers. Their total thickness is unknown, but must exceed the 520 m of outcrop exposed on the road. The contact between the Agu~in beds and the overlying strata exposed at Cerro La Penita is not exposed, but the departure from a deep-water marine environment to a continental piedmont environment is so profound that a major unconformity must exist between them.
Rio Sico locality This locality, which is 50 km south-southeast of the Agu~in Valley locality described above, lies on the road joining the towns of Bonito Oriental and E1 Carbon. A cutting 4 km south of the Sico River exposes about 25 m of limestone and mudstones. The rocks, which strike east-west and dip gently to the south, are cut by minor faults and deeply weathered dikes. At the base of the section a fold is developed, and the contact with the underlying schists is a thrust (Fig. 12). The measured section is shown in Fig. 13, and it is seen that, as in the Agu~in locality, two rock
THE SOUTHERN
FLANK OF THE TELA BASIN, HONDURAS
229
i|-~-'I il, .J
~
Z
o
0
-I l l- ~
o
:=i
---
z
"'
r
o
II
t!~!i:ll;
iil~.
E
,~ 1 I!I "
r
<
l ~t ~ ,o ,
,
]t'!;
i
iji
:ll! [i;I~. !!" ~
I,LI
~
- "~ 9,...~
0
nD
o
~
'-~
.J
="
<
~
>
0<
z
='='!
o
o "E
o iT
o
7
< = <
.........................................
o
~i.:~ ;-ilt
,..1
i~
.
""
I~
~
.~i
....
L~. , - . ~
~
i~
t..a o "~i
,.~ .Q .,.. l.i.
-
i
%=
-
~
. . . . .
~
.~
,-e=i
i
~>t "I
!ijt
:--:.=
i]ii,=~
ltl'~ '~.~-=+. =
~,~111~
~,,__~,j,,~i,:,<~=-,~ " ,~,=,-
-.'~?it
.o-..
i!|
,..~-,.
:.,~
:
."; ! . l i =
~i :l.0--l. ~
~I
.ll 8 1..11.. =
=is.,<~
-" ~
lll=
t,..-~,~
~ = o,..~
I! ~ .- ~ J..~ l =
.,,~
= :,~~~
1
o >: o
_
.
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
9i-~ li
!i
~ ~'~
0
0
~
m
rT~"
~
:<
~..
l,,_
-"
,~
9
--
o
~
.--
'~
. '> ~:'~ :
.-,
l.,.~_l~"x."fllliil!llilllll;:illlp: 9: ;4tlliiiiiiiii!li!!!!iiiiiilbi~i!l{li!iIii!l':))':..)':i "
~
N
ik~!!',!!!l!i!!i!i ~
t:l!iJl ~l!!!!!i!)l
z ~"
,.-
,.
t
t
[
~
i~!~;iil .~ '""" ~ ~ :
N!NIIii!i!IltiiliI[ 9 ;]ttf!ii i !il i i !i i i!iiti-:;;::;i;T4 !iki i iTi i ji,i ~ iN~li i iJi!7i lfll!tlff!!ili!ilri{l!~ ETI
,~=~,"~
EN!!!i!tllll!{ :',::,,: 9... ;t.!Ii!i!.!i!I . i!!i![!j!j(ill!{!P!!!i!li!i!l:!t::~;~7;(;~;::~i ...... :~ ~ ik~i!i',!!Jf!!iii!~ 1!f!:!i !!!l!!!itt =-~ iEiItl ':':,,,,,"" ':: o_o=. < -~~~-~~: ~
fl
It
R
t
t
~l
fi1
t
-~
0
~I
It
~~
t
t
~i
fil
IiI
-~ ~
It
t
~~
t
t
~
fil
lil
~~ ~
~I
t
R
t
t
~
I
t
~
..~
~~" :~
230
W.I. M A N T O N and R.S. M A N T O N
Fig. 11. Photomicrograph of sample S-144 showing abundant globigerinids some of which may be Orbulina.
Fig. 12. Mudstones and shales at the Rio Sico locality folded and thrust over schists of the basement. The thrust plane is at the base of the Jacob staff. The fold does not persist upwards, so that in the area above the photograph the rocks are completely undisturbed. View from the north, i.e. west is on the right of the photograph.
THE SOUTHERN FLANK OF THE TELA BASIN, HONDURAS
231
AGUAN BEDS
78'
RIO SICO L O C A L I T Y _~____.-~- . 75'-- ...... ._-_-_-_-
u
. . .
-
- ':"~ ~ , S 62-3
_.
.
70'_
LEGEND
more rn~gNe;
--__2
m
Thickest bed I 8/10', lam mudstone"
65'_ 60'_
Section ~ green oolour
:2 ~J :2
:5
f~~--]
Mudstone
_4
:-_ . . . . . .
~-:--~--
'~
----
>
___~ - - -
,~
~ ~ - - ~ 3 Sh ale
v
[ !
!1
ll':.'. : . f : .'11
55'
Tough gray-green siliceous indurated mud=done, non-calcareous; calcite rifled shear surfaces interbedded with 1-2" shady beds; small faults in outcrop; gray colour but often weathered red Fe stained: shears 132/56E & 60/50SW; v fine grained ss, tough, could be silicified tutf(~
;,.-.
-~,S 59
50'_
~ , S 58 -~.S 57
i:.1: :'l(:J ~= 1i
I
Limestone
Impure
Limestone
Stylolite Calcareous
~.$60
45'_
"~'-'
Break between "benches" but section may not be missing
40'--
! !
1
!
! I
1
1
1
1
1
I
I ]
I
i
!
1
1
!
l
I
'i
I
I
~
1
I
I
I
, ,,r
]
',, 1
I i I i I
I i I
I "
I 1 I ,
I 1
1 I I
:.][
I
i i ~
I 1 I I ~
~ ~" >. " Massive gray micritic fossilif Is, calcite veined, fractured, and stylolitized. Shell fragments, foraminifera?, contact .~,S 70of limestone with =li calcareous green lain mudstone or rnicritic shale (?)
'.--F/
I 1
-,=.S 71 ~
--~..'~-
'L
Igneous Dike Faulted Contact Mudstone used her~ as a textural term
i i
F~
Fossiliferous
-q,S 72-3
30'
~.$64
Amorphous non-calcaleous mudstone (containing some fossils in thin section)
-.*S 65
25' -_ ..... 20'm
"5~C:::
___:_:_ _
Mudstone (not $ilicified); sandstone chertified (?), ss is medium grained
.....................
_: _
_
N -
-
-: (S 67) non calcareous, v fine gr fiat lamin mudstone; laminae disappearing upwards
15'--
Fining up sequence - beginning of bench 2 I
i
10'--
i
[ t
[
t l
[
1 i
I
! i
I
i I s
I
I
t i I=l~. I
I
"i;,
t
l
I
J i
a
t
i l
i
t I
t
t I I i i I i J i I ! I
!
I
i
1
i
i
I
! t
i
i
,
i t
I
i i
i
i
I
l
I l
i
! 1 ! i i 94 I I I" i l
1
I
I
~
!
lq'-I
i
J
i
i
i
i
!
,
i
I
i
l !
i I I
II
l'~
Section down-faulted about 10', Tough, gray, calcite veined micriSc limestone
l]
It
i
I
!/ rE- ./ ) tv !
id~'i ! i i I 1":'.! I I I
l
J) .'q.S 76
S
75 $74
Interbedded medium bedded micritic limestone and marly shale (field
d~-~ion)
Limestone & mudstone in westward dipping thrust contact with basement
F ?-9
Highly contorted schists
Fig. 13. Agufin beds as measured at the Rio Sico locality. (The map reference is approximately 40350E,13600N, E1 Carbon sheet, HOJA 3062II.) The beds are near-horizontal; so true thickness is shown.
232 types are present. The one is a tough, gray-green, indurated, calcite-veined mudstone with individual beds up to 50 cm thick. Sedimentary structures within the mudstones are limited to planar parallel lamination with rare ripples and cross-lamination. Millimeter-scale bands of coarser-grained silt-sized material were sometimes present. The other rock type is a massive, gray, stylolitized, micritic limestone which occurs in two units, each about 5 m thick. Within each unit individual beds may be up to 30 cm thick and are separated by thin beds of argillaceous shale. The rocks, however, are much more indurated and silicified than those of the Agu~in Valley locality. The mudstones are microcrystalline and range in texture from packstones to wackestone; none contain more than about 15% bioclastic material. Some of the more silty varieties contain plant fragments. In thin section the mudstones, e.g. samples S-59, S-63, and S-67, are siliceous or cherty, and contain recrystallized spherical microfossils filled with micritic or siliceous material. In a few samples, such as S-65, millimeter-scale planar or wavy laminae and pseudo-breccias are preserved. Thin sections of the limestones, such as S-70, show them to be micrites containing abundant plant fragments and recrystallized fossil fragments. X-ray diffraction showed that
W.I. MANTON and R.S. MANTON
both mudstones and limestones contain calcite and quartz; thus there are no pure calcareous or siliceous rocks present in the section. The state of preservation in the rocks is poor and fossil identification difficult, but sample S-59, a siliceous wackestone, contains (Fig. 14) spumellarian radiolaria and foraminifera, some of which resemble Orbulina (Tony Eva, written commun.).
About 7.5 km north-northeast of the Sico River locality a road cut exposes about 5 m of folded, thinly bedded sediments thrust over andesitic lavas of the Oligocene Matagalpa formation (Fig. 15). They have not been examined in any detail, but superficially resemble the rocks studied at the Rio Sico locality. The apparent strike of the thrust plane is east-west, whereas at the Rio Sico locality it is approximately north-south, so that it seems that the rocks have been transported from a source lying within the northwestern quadrant.
Paleoenvironment Composition, sedimentary structures, and fauna are consistent with the interpretation that the rocks of the two localities are hemipelagic to pelagic slope deposits formed near a carbonate platform. No
Fig. 14. Photomicrograph of sample S-59 showing spumellarian radiolaria and poorly preserved foraminifera. The globular specimens may be Orbulina.
THE SOUTHERN FLANK OF THE TELA BASIN, HONDURAS complete turbidite units were seen, but parallel lamination indicates modification by bottom currents. The minor beds of calcarenite were probably emplaced by some type of grain flow mechanism or are the result of reworking by currents.
Lithologic age The microfossils are poorly preserved. If Orbulina has been correctly identified the rocks are no older than Middle Miocene (zone N9 of Blow, 1969) but could be as young as his zone N23, i.e. Holocene. No finer control is possible due to the absence of age-diagnostic globorotalids. In an attempt to overcome the shortcomings of the micropaleontology, strontium isotope ratios were measured in limestones from the Agu~in Valley locality. The 87sr/g6sr ratios range from 0.70598 to 0.70693 (Table 1). These values are far lower than those encountered in seawater over the Tertiary and are found only in the Late Jurassic (McArthur, 1994). Most of the samples, however, were collected from a part of the section that had been extensively intruded by mafic dikes (Fig. 10), and it is likely that Sr of low isotope ratio derived from these equilibrated with the pool of Sr of high isotopic ratio originally incorporated in the sediments. From the difference in induration between the rocks of the two localities it is clear that they had
233
Table 1 87Sr/86Sr ratios of limestones from the Agu~n beds Sample
87Sr/86Sr
R~ Sico locali~ $66A $79
0.70636 0.70606
Agudn Riverlocali~ S-134 0.70657 S-136 0.70603 S-141 0.70598 S-145 0.70663 S-148 0.70693 S-159 0.70686
Analyzed on Finnigan MAT 261 multicollector mass spectrometer. Analysis of SRM987 yielded 87sr/g6sr + 0.71023 • 1.
different histories and may have been derived from different basins. We correlate them on the basis of being deposited in similar environments, from rather tenuous fossil evidence, and in the knowledge that the Rfo Sico rocks were transported from the northwest. If the assumption is made that these rocks were laid down in marine basins associated with the strike-slip margin, it may be argued from a tectonic standpoint that they can be no older than Miocene, because before this time the part of Honduras where they are found was bounded on the north by the
Fig. 15. Agu~in-typebeds folded and thrust over Oligocene Matagalpa lavas. South is to the left of the photograph.
234
W.I. M A N T O N and R.S. M A N T O N
continental mass of Guatemala and Belize. This fact may be deduced from the magnetic anomalies in the Cayman Trough (Rosencrantz et al., 1988), which indicate that the present location of Trujillo did not lie east of the Guatemala-Belize coastline before the Early Miocene, i.e. before 20 Ma. See Rosencrantz et al. (1988, fig. 11) and Mann et al. (1991, fig. 2C). A Miocene age is compatible with the presence of ash in the rocks of the Agufin locality, which could have been derived from the voluminous ignimbritic volcanism associated with the Middle America Trench that produced the Padre Miguel volcanics coveting much of central Honduras (Donnelly et al., 1990).
is problematic, but it could be due to their being restricted to the region between the valley walls, so that they would be hidden beneath alluvium. Their presence is suggested by the magnetic anomaly off the mouth of the Agufin (Pinet, 1971). To explain the thrust sequences near the Sico River it is proposed that the transtensional phase ended and new strikeslip faults were initiated. Where the basins were cut by segments of these faults that were in transpression, the sediments were deformed into flower structures, and uplift and thrusting of the deepest parts of the basin occurred, cf. Mann and Gordon (1996).
Timing of dike emplacement
ACKNOWLEDGEMENTS
A critical question is whether the deformation of the Agufin beds occurred before or after the emplacement of the dikes. Unfortunately, no direct evidence exists because in neither locality were the dikes seen to cut across folds. In the Agufin locality, however, the dike with which piercement structures were associated presently dips north at 30 ~ but the upturning of the beds is symmetric on either side of it as if the sediments had been horizontal when it was intruded. No petrography, geochemistry or dating was done on the dikes because they are extensively altered. It may be stated that the apparent homogenization of strontium isotopes is consistent with the dikes having been intruded before faulting, because the environment of a fault or thrust zone, where fluids readily circulate through fractured rock, presents ideal conditions for homogenization.
We thank Tony Eva and Donna Meyerhoff Hull for assistance with the paleontology and J.B. Barton of Rand Afrikaans University for use of an X-ray diffractometer. Rob Rogers kindly checked on some of our observations.
APPENDIX Zircon was separated from a sample of E1 Carbon granite and was sized into fractions. Ten to twenty milligrams of each were spiked with 235U and were decomposed in Teflon capsules at 200~ An aliquot was taken and spiked with 2~ to measure Pb concentration. Lead and uranium were separated by standard ion-exchange techniques. Isotope ratios were measured on a single collector NBS-type mass spectrometer. The Pb blank was 1 ng Pb, large by today's standards but nonetheless an acceptable fraction (1 to 2%) of the Pb processed. The results are shown in Table A1, and are plotted in a Tera-Wasserberg diagram in Fig. A1. The points are highly discordant, which is attributed to Pb loss from zircon having exceptionally high contents of U.
Tectonic implications The occurrence of deep-water marine strata on the north coast of Honduras is not easily explained, and it is difficult to make regional tectonic syntheses from sedimentologic observations at two small road cuts. The most conservative and least provocative hypothesis is one which envisages that the strike-slip margin extended as far south as E1 Carbon in the Tertiary and was totally submerged, so that the Tela Basin was twice the width it is today. In its southern portion a number of deep strike-slip basins formed that were filled with turbidites. The region was uplifted, but faults curving at a steep angle to the plate boundary cut across some of these basins, preserving the sediments by downfaulting. Such faults necessarily produced highly transtensional regions which were locally intruded by mafic dikes. In this scenario the Agufin locality would represent a block of sediments preserved by downfaulting. The lack of evidence for dikes at the eastern end of the valley
Q_ o e~
0.04780 8 ~ ~~i ~ ~ 0.04775
.5*-z3my
Q_ e,l
5O
004770
'
',oo
. . . .
,;o
. . . .
2oo'
27,8U / 206 Pb
Fig. A1. Tera-Wasserberg plot of zircon data from E1 Carbon granite sample, X-3.
235
THE S O U T H E R N F L A N K OF THE T E L A BASIN, H O N D U R A S Table A 1 Uranium and lead concentrations, lead isotope ratios and corrected daughter/parent ratios for zircons from E1 Carbon granite Mesh range (~tm)
U (ppm)
Pb (ppm)
2~176
2~176
2~176
2~
2~
100-200 170-200 200-230 230-270 270-325
1477 1747 1261 1447 2489
20.0 20.2 15.8 17.7 18.7
0.0031 0.0027 0.0014 0.0014 0.0023
0.09449 0.08778 0.06864 0.06889 0.08179
0.2214 0.2101 0.1589 0.1654 0.2010
0.07460 0.06472 0.07630 0.07427 0.04222
0.01122 0.00980 0.01152 0.01120 0.00647
Common Pb correction from K-feldspar: 2~176
-- 18.695; 2~176
REFERENCES
Aves, H.S., 1983. Tela Basin. Unpublished Report. Helchris Associates, Dallas, TX, 38 pp. Banks, N.G. and Richards, M.L., 1969. Structure and bathymetry of western end of Bartlett Trough, Caribbean Sea. In: A.R. McBirney (Editor), Tectonic Relations of Northern Central America and the Western Caribbean - - the Bonacca Expedition. Am. Assoc. Pet. Geol. Mem., 11: 229-243. Blow, W.H., 1969. Late Middle Eocene to Recent planktonic foraminiferal biostratigraphy. Proceedings First International Conference on Planktonic Microfossils Geneva, 1967, Vol. 1, pp. 199-422. Caceres Avila, E, Tappmeyer, D.M., Aves, H.S., Gillett, M. and Klenk, C.D., 1984. Recent studies of basins are encouraging for future exploration in Honduras. Oil Gas J., 42:139-149. DeMets, C., Gordon, R.G., Argus, D.E and Stein, S., 1990. Current plate motions. Geophys. J. Int., 101: 425-478. Donnelly, T.W., Horne, G.S., Finch, R.C. and Lopez-Ramos, E., 1990. Northern Central America; the Maya and Chortis blocks. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Vol. H, The Geological Society of America, Boulder, CO, pp. 37-76. Foye, W.G., 1918. Notes on a collection of rocks from Honduras, Central America. J. Geol., 26:524-531. Gordon, M.B. and Muehlberger, W.R., 1994. Rotation of the Chortis block causes dextral slip on the Guayape fault. Tectonics, 13: 858-872. Heubeck, C. and Mann, E, 1991. Geologic evaluation of plate kinematic models for the North American-Caribbean plate bounding zone. Tectonophysics, 191: 1-26. Horne, G.S., Clark, G.S. and Pushkar, E, 1976a. Pre-Cretaceous rocks of Northwestern Honduras: basement terrane in Sierra de Omoa. Am. Assoc. Pet. Geol. Bull., 60: 566-583. Horne, G.S., Pushkar, E and Shafiquallah, M., 1976b. Laramide plutons on the landward continuation of the Bonacca Ridge. Publ. Geol. ICAITI, 5: 84-90. Ivey, M.L., Breyer, J.A. and Britton, J.C., 1980. Sedimentary facies and depositional history of the Swan Islands, Honduras. Sediment. Geol., 27:195-212. Kozuch, M.J., 1991. Mapa geologico de Honduras. Segunda Edition. Instituto Geografico Nacional, Tegucigalpa, Honduras. Letouzey, E, 1995. Interpr6tation gdologique de la r6gion du Rio de Aguan (Honduras) et du bassin de la basse Magdal6na, Sierra de San Jacinto (Colombie) ~ l'aide des images Landsat multispectrales tra~tdes numdriquement h l'6chelle du 1/500,000. In: A. Mascle (Editor), Gdodynamique des Cara'/bes. Editions Technip, Paris, pp. 439-450. MacDonald, W.D., 1976. Cretaceous-Tertiary evolution of the Caribbean. Transactions of the 7th Caribbean Conference, Guadeloupe, pp. 69-81. Mann, E and Gordon, M.B., 1996. Tectonic uplift and exhumation of blueschist belts along transpressional strike-slip fault zones. In: G.E. Bebout, D.W. Scholl, S.H. Kirby and J.E Platt
= 15.646; 2~176
-- 38.566.
(Editors), Subduction: Top to Bottom. Am. Geophys. Union, Geophys. Monogr., 96:143-154. Mann, E, Hempton, M.R., Bradley, D.C. and Burke, K., 1983. Development of pull-apart basins. J. Geol., 91: 529-554. Mann, E, Tyburski, S.A. and Rosencrantz, E., 1991. Neogene development of the Swan Islands restraining-bend complex, Caribbean Sea. Geology, 19: 823-826. Manton, R.S. and Manton, W.I., 1989. A major Late Cretaceous dextral strike-slip zone on the north coast of Honduras. Geol. Soc. Am., Abstr. Progr., 21: 203-204. Manton, W.I., 1987. Tectonic interpretation of the morphology of Honduras. Tectonics, 6: 633-651. Manton, W.I., 1996. The Grenville of Honduras. Geol. Soc. Am., Abstr. Progr., 28: 187. McArthur, J.M., 1994. Recent trends in strontium isotope stratigraphy. Terra Nova, 6: 331-358. McBirney, A.R. and Bass, M.N., 1969. Geology of the Bay Islands, Gulf of Honduras. In: A.R. McBirney (Editor), Tectonic Relations of Northern Central America and the Western C a r i b b e a n - the Bonacca Expedition. Am. Assoc. Pet. Geol. Mem., 11: 229-243. Minster, J.B. and Jordan, T.H., 1978. Present-day plate motions. J. Geophys. Res., 83: 5331-5354. Muehlberger, W.R., 1976. The Honduras depression. Publ. Geol. ICAITI, 5: 43-54. 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, The Geological Society of America, Boulder, CO, pp. 405432. Pinet, ER., 1971. Structural configuration of the northwestern Caribbean plate boundary. Geol. Soc. Am. Bull., 82: 20272032. Pinet, ER., 1972. Diapirlike features offshore Honduras: implications regarding tectonic evolution of Cayman Trough and central America. Geol. Soc. Am. Bull., 83: 1911-1922. Pinet, ER., 1975. Structural evolution of the of the Honduras continental margin and the sea floor south of the western Cayman Trough. Geol. Soc. Am. Bull., 86: 830-838. Pinet, ER., 1976. Morphology off northern Honduras, northwestern Caribbean Sea. Deep-Sea Res., 23: 839-847. Powers, S., 1918. Notes on the geology of eastern Guatemala and northwestern Spanish Honduras. J. Geol., 26: 507-523. Roberts, R.J. and Irving, E.M., 1957. Mineral deposits of central America. U.S. Geol. Surv. Bull., 1034: 1-205. Rosencrantz, E. and Mann, E, 1991. SeaMarc II mapping of transform faults in the Cayman Trough, Caribbean Sea. Geology, 19: 690-693. Rosencrantz, E., Ross, M.I. and Sclater, J.G., 1988. Age and spreading history of the Cayman Trough as determined from depth, heat flow, and magnetic anomalies. J. Geophys. Res., 93: 2141-2157. Sapper, K., 1905. Uber Gebirgsbau und Boden des s~idlichen
236 Mittelamerikas. Petermanns Mitteilungen, Erganzungsheft 151, Justus Perthes, Gotha, 82 pp. Southernwood, R., 1986. Late Cretaceous Limestone Clast Conglomerates of Honduras. M.S. Thesis, Univ. of Texas at Dallas, 300 pp. Stein, S., DeMets, C., Gordon, R.G., Brodholt, J., Argus, D., Engeln, J.E, Lundgren, P., Stein, C., Wiens, D.A. and Woods, D.E, 1988. A test of alternative Caribbean Plate relative motion models. J. Geophys. Res., 93:3041-3050. Sykes, L.R., McCann, W. and Kafka, A.L., 1982. Motion of
W.I. M A N T O N and R.S. M A N T O N Caribbean Plate during the last 7 million years and implications for earlier Cenozoic movements. J. Geophys. Res., 87: 10,656-10,676. Von der Hoya, H.A., 1986. A Reflection and Seismic and Magnetic Investigation of the Tela Basin: Northern Offshore Honduras. M.S. Thesis, Southern Methodist Univ., Dallas, TX, 112 pp. Williams, H. and McBirney, A.R., 1969. Volcanic history of Honduras. Univ. Calif. Publ. Geol. Sci., 85: 1-101.
Chapter 10
Cretaceous Microfaunas of the Blue Mountains, Jamaica, and of the Northern and Central Basement Complexes of Hispaniola
HOMER MONTGOMERY and EMILE A. PESSAGNO, JR.
Radiolaria and foraminifera isolated from matrix and observed in thin section refine or extend previously known or suspected chronostratigraphic assignments for Mesozoic rocks from the Blue Mountains of Jamaica and from the Puerto Plata basement complex, the Dajabon area, the Rfo San Juan complex, and the Siete Cabezas basaltic rocks of the Dominican Republic. Microfossils are found only rarely and in isolated sites some of which have questionable field relations with the surrounding complex. New chronostratigraphic assignments indicate that at least some of the basalt of the Blue Mountains, Jamaica, is older than suggested by previous paleontological studies. In the Dominican Republic, Upper Cretaceous radiolaria in the San Marcos Formation and within volcanics of the Puerto Plata basement complex refine previous chronostratigraphic assignments. Upper Cretaceous radiolaria from chert associated with the Siete Cabezas basaltic rocks of Arroyo Bermejo apparently constrain the age of basalt in this area. Upper Cretaceous foraminifera from Dajabon and from the Rfo San Juan complex are the first reported fossils from these areas.
INTRODUCTION
The Late Cretaceous to Eocene Great Arc (Fig. 1) of the Caribbean existed as an east-facing volcanic island chain spanning the gap between North and South America. During the Eocene-Oligocene the arc collided with the Bahama platform in a diachronous event that produced an east-west-trending strike-slip plate boundary on the north- and the east-facing subduction zone present today along the eastern boundary. Recognition, dating, and correlation of components of the Great Arc remain an on-going puzzle. Several relatively well-studied terrane fragments of prominence in discussions of plate tectonic interactions remain so poorly age-constrained that labels such as Mesozoic may be applied only with caution. Examples of such units include the metamorphic complexes of Rio San Juan and Samami in the Dominican Republic. Other areas are broadly mapped with little lithological inventory and with no age control of some of the units. An example is the Dajabon area of the Dominican Republic. In this report we add to the body of paleontological information from problematic Cretaceous units in Jamaica and the Dominican Republic.
Isotopic ages, where determined, may date uncertain events which could be wide-scale magmatism and/or plate collision but are not necessarily ages of deposition. Paleontological control associated with isotopic ages remains almost nonexistent. Rare sedimentary rock bodies, many of which go unrecognized, have never been sampled or at least have never been rigorously investigated for fossils. A note relevant to GPS locations of outcrops in this report is in order. All locations in Jamaica and locations in the Dominican Republic of DR92 and DR95 samples were located using map datum WGS84. Use of this datum produces a map error of approximately 300 m in latitude. DR96 samples were located using NAD 1927-Caribbean. These localities should be accurately located.
MICROFAUNAS AND AGES
The Cretaceous faunas presented in this report are derived from rocks collected in Jamaica and in the Dominican Republic (Fig. 2). Most of the rocks were collected in situ. As noted below, two samples collected in Jamaica are from fiver gravels. The
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 237-246. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
238
H. MONTGOMERY and E.A. PESSAGNO
Fig. 1. Late Cretaceous snapshot of possible paleogeographic setting of the Great Arc of the Caribbean.
gravels were sampled because no fossiliferous rocks were discovered cropping out in this area of the Blue Mountains. Radiolaria and foraminifera were isolated from siliceous rocks via extraction in dilute hydrofluoric acid. Fossils vary from poorly to well-preserved. Foraminifera extracted from chert via the same process are fluoritized and are only rarely identifiable or are identifiable only to genus. Thin sections were cut of most of the siliceous samples that had produced no microfossils in residue. Radiolaria were present
Fig. 2. Location map of outcrops sampled for this report.
in one of those samples, a cobble from the Devil's River, Jamaica. Thin sections of limestone revealed identifiable larger and planktic foraminifera. The paleontological details presented modify or amplify what was previously known about several units (Blue Mountains, Jamaica; San Marcos Formation, and Arroyo Bermejo, Dominican Republic) or are the first assignments made (Puerto Plata basement complex of Maimon Bay, Dajabon cherts, and Rfo San Juan u all of the Dominican Republic).
CRETACEOUS MICROFAUNAS OF JAMAICA AND HISPANIOLA
239
MICROFAUNAS OF THE DOMINICAN REPUBLIC
Kg DR95.8a
Puerto Plata basement complex (DR92.8a) Limestone nodules in volcanic matrix are found along the margin of Maimon Bay (Fig. 3). Limestone is present in four distinct forms. Light gray nodules approximately 10 to 15 cm in diameter are too highly metamorphosed to yield fossils. A few dark gray nodules produced radiolaria and planktic foraminifera. Buff-colored limestone adjacent basalt pillows yielded no fossils. Select pieces of brecciated, greenish-gray limestone (sample DR92.8a) and crudely bedded but discontinuous dark gray, silicified limestone proved variably fossiliferous. Two pieces of greenish-gray limestone within highly weathered volcanic matrix and adjacent crumbly red umber yielded well-preserved radiolaria and planktic foraminifera. Pieces of all limestone samples were dissolved in HC1 and thin sectioned. The planktonic foraminifera of sample DR92.8a include Globotruncana arca, Globotruncana bulloides, Globotruncana hilli, Globotruncana lapparenti s.s., Globotruncana linneiana, Globotruncanita calcarata, Globotruncanita elevata, Globotruncanita stuartiformis, and Heterohelix sp. The biostratigraphic determination is upper Campanian. Radiolaria of sample DR92.8a are some of the best we have observed in Mesozoic Caribbean rocks. HC1 residues include Alievium gallowayi (base of Coniacian-top of Campanian), Amphipyndax pseudoconulus (base of Campanian-top of Maastrichtian), Archaeodictyomitra squinaboli (upper Albian-top of Campanian), Dictyomitraformosa (base of Coniacian-top of Campanian), Patellula verteroensis (Campanian), Pseudoaulophacus pargueraensis (base of Santonian-top of Campanian). Although ranges vary somewhat among authors, DR92.8a radiolaria are clearly Campanian.
San Marcos Formation, Puerto Plata basement complex (DR95.1a; Pindell and Draper (1991) locality No. 31), UTM: 2188790; 307775 Abundant and well-preserved radiolaria are present in a dark red jasper boulder located near Imbert along the east side of the highway that runs
Maimon Bay Puerto Nm
--x
Kv f
Plata
N
l
Nm Pi
0
i
I
2 km
I
"N to Santiago Fig. 3. Location map for sampled outcrops of the Puerto Plata area, Dominican Republic. Kv -- volcanics; Kg -- gabbro; Ks -serpentinite; Pi = Imbert Fm." Nm = San Marcos unit.
from Puerto Plata to Santiago (Fig. 3). The exposed portion of the boulder is the size of a small car and is associated with a mass of fractured serpentinite and limestone in muddy matrix. Radiolaria from sample DR95.1a are lower Santonian to lower Campanian (Table 1) and are the best-documented radiolarian assemblage of the San Marcos unit.
Dajabon subcomplex(?) of the Upper Duarte complex(?) (DR95.12), UTM: 2167651; 219445 Abundant, multicolored (red, orange, blue, green, black, brown, gray) chert rubble is scattered across several low hills northeast of the town of Dajabon (Fig. 4). No bedded rocks crop out in this area of low relief. Dissolution of a few of the chert samples yielded Albian to upper Cenomanian foraminifera (Table 2).
Table 1 Radiolaria from San Marcos Fm., Puerto Plata basement complex, Dominican Republic sample DR95.1a Radiolaria present
Chronostratigraphic assignment
Alievium gallowayi Amphipyndax stocki Dictyomitraformosa Praeconocaryoma universa Pseudoaulophacus lenticulatus
base Santonian-middle-upperCampanian Coniacian-lower Campanian base Coniacian-middle Campanian lower Coniacian-upper Campanian
240
H. MONTGOMERY and E.A. PESSAGNO Table 4 Foraminifera from the Rio San Juan complex, Dominican Republic sample DR96.1 Foraminifera present
Chronostratigraphicassignment
Orbitoides sp. Heterohelix sp.
upper Campanian-upper Maastrichtian upper Albian-upper Maastrichtian
Rio San Juan complex (DR96.1), UTM: 2165512; 380118
DR95.12
N
Dajabon t'4
km
Limestone at this site is located in the fiver at the end of the road south of Magante (Fig. 5). Much of the limestone is metamorphosed, especially where interbedded with schist. Farther away from the schist the limestone is sporadically but richly fossiliferous. Among other microfossils and broken megafossil fragments, thin sections revealed upper Campanian to Maastrichtian foraminifera (Table 4).
I
Arroyo Bermejo (DR96.9), UTM: 2056799; 385880 Fig. 4. Location map for chert sampled in the Dajabon area, Dominican Republic.
Rio San Juan complex (DR95.7a), UTM: 216820; 37675 Sandy red mudstone dug from the muddy shale in the roadcut at the top of the first hill south of the town of Magante (Fig. 5) yielded poorly preserved foraminifera (Table 3). The identifications, if correct, assign the sample to the uppermost Cenomanian on the basis of the concurrence of Dicarinella and Thalmanninella. Table 2 Foraminifera from the Dajabon subcomplex(?) of the Upper Duarte complex(?), Dominican Republic sample DR95.12 Forams p r e s e n t
Chronostratigraphic assignment
Heterohelix moremani Hedbergella sp. Spongodiscacea radiolaria
upper Albian-lower Turonian Barremian-Maastrichtian
Microfossiliferous, black, gray, and greenish bedded chert of the Siete Cabezas Formation is found in the cut bank of the stream approximately 100 m south of the bridge that leads east of the Duarte highway (Fig. 6). Radiolaria are probably middle Campanian to Maastrichtian (Table 5). The critical specimen for this chronostratigraphic assignment is Dictyomitra multicostata, and while slightly broken, is clearly this taxon. Other specimens suggest a Coniacian assignment. Problematically, we have not observed any specimens of Dictyomitra formosa which is usually found associated with Dictyomitra multicostata. We suggest the Coniacian specimens are reworked.
SIGNIFICANCE OF THE DOMINICAN REPUBLIC MICROFAUNAS
Puerto Plata basement complex The Puerto Plata basement complex is composed of fault-bounded sections of serpentinized peridotite, harzburgite, layered cumulate ultramafics and gab-
Table 3 Foraminifera from the Rfo San Juan complex, Dominican Republic sample DR95.7a Foraminifera present
Chronostratigraphic assignment
Hedbergella sp. cf. delrioensis Thalmaninella (poorlypreserved) Double keeled forms such as: Dicarinella sp., Marginotruncana sp.
uppermost Cenomanian uppermost Cenomanian
CRETACEOUS MICROFAUNAS OF JAMAICA AND HISPANIOLA
241
ATLANTIC OCEAN to
to Rio San Juan
ar Hernandez
N
DR95.7a
DR96.1
Fig. 5. Location map for sampled outcrops of the Rfo San Juan area, Dominican Republic. S = Gaspar Hernandez serpentinites; H = Hicoteca schists. bros, massive gabbros, and basic to intermediate volcanics (Pindell and Draper, 1991). The age of the Puerto Plata basement complex is an enduring problem. The Puerto Plata basement complex was interpreted as a pre-Paleocene ophiolite complex (Pindell and Draper, 1991) emplaced as part of a forearc/accretionary prism of the Great Caribbean Arc (Mann et al., 1991). But the paleogeographic relationship of the Puerto Plata basement complex both to other Hispaniola terranes and to the Caribbean
Ud\
Bonao \ < ) \
]
,-o\ , o \, N K \ U s i~ ,i~ .. Villa'~k--~ ~ I Altagracia-~ \X / ~'
-. us
0"Q" 0'
4 km I~"
9 to I, Santo Domingo \
Ws~X
Id
",~X._~ \ ~ "N~a
Fig. 6. Location map for chert sampled along Rfo Isabela at Arroyo Bermejo in the Siete Cabezas Formation. Us = Siete Cabezas Fm.; Ud = Duarte Fm.
plate remained conjectural without a chronostratigraphic framework. The discovery of Campanian faunas allows comparison with several other circumCaribbean terranes. Campanian radiolaria are well-known in the Caribbean and surrounding regions from Puerto Rico, Cuba, Texas, Panama, and DSDP Sites 146) (Venezuelan basin), 4 (adjacent to the Bahama platform), 95 (200 km northwest of Cuba), and 24 (western Brazil basin). Recently, Campanian radiolaria identical to the Puerto Plata basement complex fauna were isolated from red chert (samples provided by M. Iturralde-Vinent of the Museo Nacional de Historia Natural, Habana) in a chert and basalt sequence of the Santa Teresa or Quifiones Formation of the Bahfa Honda quadrangle (H. Montgomery, unpubl. data). Brecciated greenish gray foraminiferal and radiolarian limestone of Campanian age drilled at Site 146 in the Caribbean Sea between Hispaniola and Venezuela (Edgar et al., 1973) and examined by Morin (1982) are the age and lithological equivalents of the Puerto Plata basement complex limestones. A fauna examined by Foreman (1973) containing many of the same species as the Puerto Plata basement complex was recovered in the Gulf of Mexico at DSDP Site 95 from olive-gray, dolomitic limestone, some of which was interbedded with black chert. Based on the planktic microfaunal assemblage, the
242
H. MONTGOMERY and E.A. PESSAGNO
Table 5 Radiolaria from Arroyo Bermejo, Dominican Republic sample DR96.9 Radiolaria present
Chronostratigraphic assignment
Acanthocircus spp. Amphipyndax sp. cf. A. stocki Alievium praegallowayi Archaeosponoprunum triplum Archaeodictyomitra sp. Artostrobium sp. Dictyomitra multicostata Microsciadiocapsa cortinaenis Patellula sp. Pseudoaulophacus sp. Pseudodictyomitra sp. Theocampe sp. cf. tina
Triassic-Upper Cretaceous lower Coniacian-lower Santonian Coniacian Middle Jurassic-Upper Cretaceous middle Campanian-Maastrichtian Coniacian Cenomanian-Maastrichtian middle Turonian-upper Campanian Upper Jurassic-middle Turonian
Puerto Plata basement complex fauna was deposited at abyssal depths probably slightly above the CCD.
San Marcos Formation The San Marcos Formation was described as having been emplaced by mud diapirism and lateral flow sampling all of the pre-Quaternary stratigraphic units of the Puerto Plata area (Pindell and Draper, 1991). Within this emplacement scenario the jasper block sampled for this report would have been derived from the Puerto Plata basement complex. Such a scenario is consistent with chronostratigraphic assignments for Campanian limestone of the Puerto Plata basement complex at Maimon Bay. Ages reported for the San Marcos are Early Cretaceous (Bourgois et al., 1982) and Miocene (Nagle, 1966). Lower Santonian to lower Campanian jasper blocks such as our sample DR95.1a have not been previously recognized in any other part of the San Marcos.
Dajabon subcomplex(?) of the Upper Duarte complex(?) The age of the Dajabon rocks is important to decide whether they are a westerly component of the Duarte complex or are they of some other association. A Jurassic age for the cherts would lend credence to a Duarte association. Dajabon sample DR95.12 presented an older age for the Dajabon area than any previously reported, but the chert is clearly not as old as the Jurassic red ribbon chert of the Duarte complex in the Central Cordillera (Montgomery et al., 1994a). The Dajabon region remains poorly mapped and there is a problem with identification of the Duarte lithologies in this region. Other units in the Dajabon area are mapped as Tertiary, and DR95.12 is clearly not associated with these rocks. A less metamor-
phosed version of the Duarte may be confused with basaltic rocks of the Magua Formation (considered Eocene). There is even a chance that these rocks are part of the Maimon-Amina schists complex. Further mapping will be required.
Rio San Juan complex The Rio San Juan complex of the northeastern Dominican Republic is composed predominantly of igneous and metamorphic rocks. Conglomerate in sandstone/tuff sequences is similar to strata of the Imbert Formation of the Puerto Plata area (Draper and Nagle, 1991). No fossil-bearing rocks have been reported for the Rio San Juan complex. Superposition is not particularly age-defining as sedimentary rocks surrounding the Rio San Juan complex are of the Upper Eocene La Toca Formation. Microfossils of DR95.7a and DR96.1 are the first reported from the Rio San Juan complex. In spite of this seemingly important discovery, it remains unclear if the fossils date metamorphic rocks of the complex, or are themselves two separate overlying or incorporated units. The site is mapped as La Toca Formation of the Mamey group which is Upper Eocene, presenting the possibility that the foraminifera are somehow reworked. G. Draper (written communic., 1997) suggested the sample may be nonconformable on the coherent part of the Rio San Juan high-pressure rocks comparing such an occurrence to a similar relationship from the Purial area in Cuba. The significance of the upper Cenomanian chronostratigraphic assignment for DR95.7a remains tentative. The foraminifera are poorly preserved. Having been collected from a mud matrix not unlike the San Marcos Formation in appearance, we assume that this rock is not a genetic component of the Rio San Juan metamorphic complex. The Imbert Formation is found in the northern part of the Rio San Juan
CRETACEOUS MICROFAUNAS OF JAMAICA AND HISPANIOLA
Bath Fountain
Spring Bank
Kbd
Kcp
g.,-:
/
i
/
,
Kcp
Kbd (/)\
Kcp
~-...~.,.
\ !
243
j
DR95.5
.
! DR9S.10
Kbd ~--_~_~.~,-,,___: \. K b d ~ N Bath
to Port Morant
I
Fig. 7. Location m a p for s a m p l e d outcrops near Bath, Jamaica. Kcp -- Cross Pass Fm.; Kr : Rfo G r a n d e Fm.; Kbd : B a t h - D u n r o b i n Fm.
complex (Draper and Nagle, 1991), but this material is unlike the sandstone, serpentinite conglomerate, and tuff components of the Imbert. The Imbert is also younger, being Paleocene(?)-Eocene (Pindell and Draper, 1991). The limestone of sample DR96.1 is rich in calcium carbonate fossils and fossil fragments including lagoonal miliolid foraminifera, larger foraminifera, various algae, echinoderm fragments, and other calcareous debris normally encountered in thin sections of sediments deposited in shallow water. The seemingly inconsistent element is the presence of deeper-water planktic foraminifera. Such faunal mixing is not unknown in the Caribbean in various sections with shallow-water fossils having cascaded into adjacent deeper waters (Montgomery, 1997). The Campanian to Maastrichtian chronostratigraphic assignment for the DR96.1 limestone is consistent with the 4~ cooling age of 85.1 Ma for a hornblende-phengite-bearing block from the Jagua Clara M61ange (Draper and Nagle, 1991). The dated block is from the m61ange part of the Rfo San Juan, but no dates were forthcoming from the area that contains sample DR96.1. Draper (G. Draper, written commun., 1996) is uncertain whether the fossiliferous limestone is part of the metamorphic part of the Rfo San Juan complex. Over the hill from the fossiliferous limestone are outcrops of Imbert Formation. The question remains whether this fossiliferous limestone is a part of an unmetamorphosed Upper Cretaceous to Paleocene sequence. More mapping will be required to resolve this problem.
and Coniacian assignments. Chris Sinton (J. Lewis, written commun., 1996) obtained Ar-Ar ages of 69 + 0.7 Ma whole rock and 68.5 + 0.5 Ma plagioclase for the Siete Cabezas basalt located several kilometers north of Arroyo Bermejo. Assuming reworking of the Coniacian specimens, our assignment agrees relatively well with the Ar-Ar ages.
M I C R O F A U N A S OF J A M A I C A
All of the samples were collected in the Blue Mountains of southeastern Jamaica (Fig. 7).
Devil's River (JAM95.5a), UTM: 1976631; 344535 A dark red jasper cobble collected from the Devil's River stream bed where it crosses the road east of Bath yielded Santonian to lower Campanian radiolaria (Table 6) with possibly reworked Pseudoaulophacus sp. cf. praefloresensis. The origin of this sample is uncertain but the Devil's River drains the Blue Mountains block.
Devil's River (JAM95.5b), UTM: 1976631; 344535 A thin section of a red jasper cobble from the Devil's River stream bed at the above locality (JAM95.5a) yielded Coniacian to lowermost Santonian radiolaria (Table 7).
Arroyo Bermejo
Devil's River (No. 802), locality: third cascade of the Devil's River northwest of the village of Spring Bank
This chronostratigraphic assignment for the Arroyo Bermejo chert is not completely consistent with that of Boisseau (1987) who presented Cenomanian
This sample of reddish calcareous mudstone was collected by John Lewis and Edward Robinson (University of the West Indies museum sample No. 802).
244
H. MONTGOMERY and E.A. PESSAGNO
Table 6 Radiolaria from Devil's River, Jamaica sample JAM95.5a Radiolaria present
Chronostratigraphic assignment
Alievium gallowayi Dictyomitra formosa Dictyomitra torquata Praeconocaryoma universa Pseudoaulophacus sp. cf. praefloresensis Patellula sp.
base Santonian-middle-upperCampanian Coniacian-lower Campanian Coniacian-lower Campanian base Coniacian-middle Campanian middle-upper Turonian
Table 7 Radiolaria from Devil's River, Jamaica sample JAM95.5b Radiolaria present
Chronostratigraphic assignment
Alievium praegallowayi/ superbum Dictyomitra formosa Dictyomitra torquata
base Coniacian-lower Santonian Coniacian-lower Campanian Coniacian-lower Campanian
Table 8 Radiolaria from Devil's River, Jamaica sample No. 802 Radiolaria present
Chronostratigraphic assignment
Alievium superbum/praegallowayi Archaeodictyomitra sp. Crucella sp. Dictyomitra formosa Pseudodictymitra n. sp.
middle Turonian-upper Campanian Jurassic-Cretaceous Jurassic-Cretaceous Coniacian-lower Campanian Upper Jurassic-middle Turonian
Dissolution produced upper Turonian to lower Coniacian (Table 8) radiolaria. This rock probably contains the oldest Jamaican microfossils collected for this report.
Bath (JAM95.10), locality: adjacent to Churchill's Vineyard along the road from the town of Bath to Bath Spring The Wild Cane complex outcrop sampled is red, thinly bedded chert. Relatively well-preserved middle-upper Turonian to upper Coniacian radiolaria (Table 9) suggest that JAM 95.10 is one of the older samples collected for this report from the Blue Mountains block.
SIGNIFICANCE OF THE JAMAICAN MICROFAUNAS
Blue Mountains The Bath-Dunrobin Formation and Wild Cane complex of Jamaica is composed mostly of basalt, dolerite, gabbro, and tonalite with rare limestone and chert. Foraminifera from the Back Rfo Grande limestone of the Blue Mountains inlier were reported as upper Campanian (Krijnen and Chin, 1978). The three radiolarian assemblages presented in this report are all older than Campanian. Assuming that the chert, jasper, and red mudstone of this report are/were interbedded with the basalts, then some of the Bath-Dunrobin basalts are older than previously thought. The JAM 95.10 chert sampled is clearly
Table 9 Radiolaria from Bath, Jamaica sample JAM95.10 Radiolaria present
Chronostratigraphic assignment
Alievium superbum/gallowayi Artostrobium urna Dictyomitra napaensis Praeconocaryoma universa Vitorfus sp.
middle Turonian-upper Campanian middle-upper Turonian-upper Coniacian base Coniacian-middle Campanian Albian-?Maastrichtian
CRETACEOUS MICROFAUNAS OF JAMAICA AND HISPANIOLA incorporated within a volcanic sequence. Thus the middle-upper Turonian to upper Coniacian chronostratigraphic assignment for JAM 95.10 indicates the same assignment for the volcanic sequence. The age for JAM 95.10 suggests that the ophiolitic sequence of the southeast mountains is older than Campanian, a proposition long suspected by G. Draper (written commun., 1997). The Campanian limestone is at the top of, or unconformable on, the ophiolitic sequence.
CONCLUSIONS
Tectonic scenarios describing the complex origin and development of the Caribbean plate are difficult to construct partly because of poor age-control. One persistent problem involves the earliest paleogeographic history of the Great Arc of the Caribbean (for an overview see Mann et al., 1991). Montgomery, Pessagno, and co-workers (Montgomery et al., 1992, 1994a,b) described pre-Great Arc paleogeography with emphasis on the Pacific origin of sedimentary rocks that collected in the west-facing trench during the Jurassic. The timing of arc polarity reversal from west-facing to east-facing marking the beginning of the Great Arc remains speculative. Based on the age of the youngest chert in the Bermeja accretionary complex of Puerto Rico, Montgomery et al. (1994a) suggested that the flip occurred shortly after 90 Ma. Incorporating the results of this report with other current information, observations are forthcoming better characterizing the origin of the Jamaican and Dominican units as regards the earliest history of the Great Arc. (1) Most if not all of the sedimentary rocks described in this report formed in deep basins some probably below the CCD. (2) Lithologies are significant. The importance of multicolored chert, jasper, red mudstone, and limestone of the Upper Cretaceous metamorphic and sedimentary complexes of the leading (and colliding) edge of the Caribbean plate should be clear. Nowhere have we or has anyone else reported finding post-Jurassic red ribbon chert in the Antilles nor has red ribbon chert been dredged or drilled in the Atlantic basin. This is notable because red ribbon chert at E1 Aguacate, Dominican Republic (Montgomery et al., 1994a), Sierra Bermeja, Puerto Rico (Montgomery et al., 1994b), and La Ddsirade, Guadeloupe (Montgomery et al., 1992) is remnant, sub-CCD, oceanic floor originating in the Pacific basin. As far as concerns the Caribbean, red ribbon chert is exotic and is a defining characteristic of Pacific rather than Atlantic origin. Red ribbon chert accumulated only during the west-facing phase of subduction.
245
(3) The various Cretaceous rocks of Jamaica, the Dominican Republic, and other Cretaceous Great Arc subduction complexes and associated sedimentary sequences are a collection of Atlantic bank and basin material with no incorporated rocks of Pacific affinity. Multicolored Cretaceous chert and jasper are common. Red mudstone is present, a rock not observed in the Jurassic sequences. Most obvious is the abundance of limestone in Great Arc material. No carbonate rocks are known from the oldest Jurassic terrane fragments which are found in the Dominican Republic and Puerto Rico. While not abundant nor apparent, the oldest limestone is interbedded with red ribbon chert in the youngest Jurassic terrane fragment in La D6sirade. (4) Based on the chronostratigraphy presented in this report and assuming an approximate 90 Ma date for polarity reversal, various components of the Great Arc were deposited before reversal within the arc or on the Atlantic side. The JAM samples are pre-Great Arc. The Rfo San Juan mudstone and the Dajabon cherts are pre-Great Arc. The other samples are too close to call or they formed after polarity reversal. (5) Based on the chronostratigraphic assignment of limestone in the Rio San Juan complex, and assuming that this limestone is incorporated within the subduction complex, the Great Arc of the Caribbean apparently encountered and subducted part of a carbonate platform (too early to be the Bahama platform?) during the latest Cretaceous. The striking admixture of shallow- and deep-water constituents may suggest tectonic disruption.
ACKNOWLEDGEMENTS
This work was supported by National Science Foundation grants NSF-EAR 9418194 and NSFEAR-9117397 to E.A. Pessagno and H. Montgomery. Some initial collection and site evaluation by H. Montgomery was supported by funding from the University of Puerto Rico through the NSF EPSCoR program. J. Lewis and G. Draper are due much credit for their guidance during many days of field work and during many nights of highly constructive contemplation. Falconbridge Dominicana provided vehicles, lodging, and advice. Special thanks go to Salvador Brouwer for his help. Paul Mann offered a thorough and constructive review which greatly improved the manuscript.
REFERENCES
Boisseau, M., 1987. Le flanc nord-est de la Cordillbre Centrale dominicaine (Hispaniola, Grandes Antilles); un 6difice de
246 nappes Cr6tac6 polyphase. Doctoral Thesis 3rd Cycle, Universit6 Marie et Pierre Curie, Paris, 215 pp. Bourgois, J., Vila, J.-M., Llinas, R. and Tavares, I., 1982. Datos geologicos nuevos acera de la region de Puerto Plata (Republica Dominicana). In: Transactions, 9th Caribbean Geological Conference, Santo Domingo, pp. 39-50. Draper, G. and Nagle, E, 1991. Geology, structure, and tectonic development of the Rfo San Juan Complex, northern Dominican Republic. In: P. Mann, G. Draper and J.E Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 77-95. Edgar, N.T., Kaneps, A.G. and Herring, J.R., 1973. Initial Reports of the Deep Sea Drilling Project, Vol. 15. U.S. Government Printing Office, Washington, DC, 1137 pp. Foreman, H.P., 1973. Radiolaria of Leg 10 with systematics and ranges for the families Amphypyndacidae, Artostrbiidae, and Theoperidae. In: J.L. Worzel, W. Bryant et al. (Editors), Initial Reports of the Deep Sea Drilling Project, Vol. 10. U.S. Government Printing Office, Washington, DC, pp. 407-474. Krijnen, J. and Chin, A., 1978. The geology of the northern central and southeastern Blue Mountains, Jamaica with a provisional compilation of the entire inlier. Geol. Mijnbouw, 57: 243-250. Mann, P., Draper, G. and Lewis, J.E, 1991. An overview of the geologic and tectonic development of Hispaniola. In: P. Mann, G. Draper and J.E Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary
H. M O N T G O M E R Y and E.A. P E S S A G N O in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 1-28. Montgomery, H., 1997. Paleogene stratigraphy and sedimentology of the North Coast, Puerto Rico. In: E. Lidiak and D.K. Larue (Editors), Tectonics of the Northeastern Caribbean. Spec. Publ. Geol. Soc. Am. 322: 177-192. Montgomery, H., Pessagno, E.A., Jr. and Mufioz, I.M., 1992. Jurassic (Tithonian) Radiolaria from La D6sirade (Lesser Antilles): preliminary paleontological and tectonic implications. Tectonics, 11: 1426-1432. Montgomery, H., Pessagno, E.A., Jr. and Lewis, J.A., Schellekens, J.H., 1994a. Paleogeography of the Jurassic fragments in the Caribbean. Tectonics, 13: 725-732. Montgomery, H., Pessagno, E.A., Jr. and Pindell, J.L., 1994b. A 195 Ma terrane in a 165 Ma ocean: Pacific origin of the Caribbean Plate. GSA Today, 4: 1-6. Morin, K.M., 1982. Analysis and Definition of Campanian (Late Cretaceous) Radiolarian Populations Characteristic of Tropical to Subtropical Latitudes. Ph.D. Dissertation, Univ. of Texas at Dallas, 375 pp. Nagle, E, 1966. Geology of the Puerto Plata Area, Dominican Republic. Ph.D. Thesis, Princeton University, Princeton, NJ, 171 pp. Pindell, J.L. and Draper, G., 1991. Stratigraphy and geological history of the Puerto Plata area, northern Dominican Republic. In: P. Mann, G. Draper and J.E Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 97-114.
Chapter 11
Cenozoic E1 Mamey Group of Northern Hispaniola: A Sedimentary Record of Subduction, Collisional and Strike-Slip Events within the North America-Caribbean Plate Boundary Zone
RUURDJAN DE ZOETEN and PAUL MANN
Cretaceous and Cenozoic paleogeographic and plate tectonic reconstructions of the Greater Antilles (Hispaniola, Cuba, Jamaica, and Puerto Rico) are complicated by large-offset, Eocene? to Recent strike-slip movements between the North America and Caribbean plates. Moreover, Eocene?-Recent oblique subduction of the Bahama carbonate platform presently affects the Hispaniola region and further complicates the reconstruction of this wide and complex plate boundary zone. This paper describes a detailed sedimentological study of deformed and uplifted Eocene to Lower Pliocene sedimentary rocks (El Mamey Group) between the North America and Caribbean plates in northern Hispaniola (Dominican Republic). Paleocene to Lower Pliocene siliciclastic and carbonate rocks of the E1 Mamey Group, which crop out within a 500 km 2 area of the central Cordillera Septentrional and are well exposed in road and stream cuts, formed the object of this regional tectonic study. On the basis of its compositional, age and facies character, we divide the sedimentary succession of the E1 Mamey Group of the central Cordillera Septentrional into three lithologically distinct, stratigraphic sequences which we relate to three tectonic phases that affected this segment of the North America-Caribbean plate boundary from the Eocene to Recent. Phase 1 (Paleocene to Middle Eocene). Sedimentary and facies characteristics of an approximately 250-m-thick section of Upper Paleocene to Lower Eocene siliciclastic and carbonate rocks (Los Hidalgos ~ormation), suggest that these rocks were deposited in a deep-marine, hemipelagic environment adjacent to an active volcanic arc. Calc-alkaline volcanic flows and sills are interbedded with these deep-marine sedimentary rocks. Termination of deposition and volcanism in Early to Middle Eocene time coincides with a major folding and uplift event, which we believe was caused by the early attempted subduction of the Bahama carbonate platform beneath the arc-related basin. This event terminated arc activity in the Hispaniola volcanic arc and forearc. Phase 2 (Late Eocene to Early Miocene). Sedimentary and facies characteristics of a 4000-m-thick, Upper Eocene to Lower Miocene siliciclastic succession (Altamira and Las Lavas formations of the E1 Mamey Group) suggest that these rocks were deposited as submarine turbidites and other types of mass-flow deposits within a west-northwest-trending, elongate basin. Petrographic analysis of framework grains of sandstones within the section shows two distinct sandstone populations separated by a linear, 100-400-m-wide left-lateral strike-slip fault zone. Petrographic differences across the fault zone are especially prominent in coeval Oligocene sedimentary rocks and suggest that the two basins were juxtaposed by lateral fault movement sometime after Oligocene time. The end of deep-marine siliciclastic deposition in both basins coincides with a gentle Middle Miocene folding event believed to be related to transpressional strike-slip faulting. Phase 3 (Late Miocene to Recent). Sedimentary and facies characteristics of an approximately 250-m-thick section of Upper Miocene to Lower Pliocene carbonate rocks (Villa Trina Formation) suggest that these rocks were deposited as a shallow carbonate bank above slightly folded, Early Miocene siliciclastic rocks. Carbonate deposition was terminated in Early Pliocene time by a folding and uplift event believed to be related to transpression along a restraining bend in the Septentrional fault zone.
INTRODUCTION
Objectives The island of Hispaniola consists of a 250-kmwide tectonic collage of fault-bounded igneous,
metamorphic, and sedimentary basement blocks of Late Cretaceous to Middle Eocene age that formed in an intra-oceanic island-arc setting (cf. Bowin, 1975; Sykes et al., 1982; Mann et al., 1991, for regional reviews; Fig. 1). The basement blocks are overlain by a cover of Upper Eocene to Pliocene
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsfi), pp. 247-286. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
248
R. DE ZOETEN and E MANN
Fig. 1. Present-day plate structure of the Caribbean region. Direction and rates of plate motion relative to the Caribbean plate are from DeMets et al. (1990) and Dixon et al. (1998). The island of Hispaniola straddles the active left-lateral strike-slip zone separating the North America and Caribbean plates. The large amount of plate convergence and topographic uplift of Hispaniola is related to transpression between two thick crustal blocks: the Bahama carbonate platform to the north and the Cretaceous Caribbean oceanic plateau (hatched pattern) on the Caribbean plate to the south. Box shows map area shown in Fig. 2. siliciclastic and carbonate sedimentary rocks that post-date island-arc activity and mainly record the initiation of the present period of left-lateral strike-slip motion between the North America and Caribbean plates (Fig. 2). Many previous tectonically oriented studies in northern Hispaniola have focused on the composition and structure of arcrelated basement rocks in order to better understand the origin and tectonic history of the Greater Antilles island-arc system (for example, Nagle, 1979; Palmer, 1979; Pindell and Draper, 1991; Joyce, 1991; Fig. 2). Following the earlier sedimentation and tectonics study of Dolan et al. (1991), the objective of this paper is to reconstruct the depositional history of Paleocene to Lower Pliocene sedimentary rocks of the central Cordillera Septentrional that either post-date or accompany the final stages of island-arc activity in Hispaniola (Fig. 1). The method of this paper is to reconstruct the depositional history and water depth of sedimentary basins overlying arc rocks through the use of: (1) regionally correlated measured sections of continuous stratigraphic sequences; (2) integration of biostratigraphic data as a way to reconstruct sedimentary environments, determine water depths through time, and date sedimentary, tectonic, and eustatic events; (3) integration of paleocurrent and sandstone petrographic data to determine siliciclastic provenance.
Based on the same outcrops discussed in this paper, de Zoeten and Mann (1991) published a Cenozoic structural history of the area using data on major and minor faults and folds observed in the same set of outcrops as described in this paper. Dolan et al. (1991) presented a summary of some of the preliminary results of this study in their discussion of the pattern of regional basin formation during the Cenozoic in Hispaniola and Puerto Rico. Neither of the previous studies presented the primary sedimentological data collected by de Zoeten (1988) that are fundamental to many tectonic interpretations.
Significance of this study The Cordillera Septentrional is a critical area for studies of post-Middle Eocene island-arc tectonics in Hispaniola because the area straddles the Septentrional fault zone, which is presently the main strikeslip fault zone separating the North America and Caribbean plates (Mann et al., 1991, 1998) (Fig. 2). Marine studies have shown that the Septentrional fault zone continues westward as fault zones bounding the Middle Eocene to present Cayman Trough pull-apart basin (Rosencrantz et al., 1988; Calais and Mercier de L6pinay, 1995) and eastward as active fault zones along the southern edge of the Puerto Rico Trench (Grindlay et al., 1997; Dolan et al., 1998; Fig. 1). The Cordillera Septentrional provides the largest sub-
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA aerial exposure of this 3600-km-long, interplate fault system outside of Central America (Fig. 1).
249
GEOLOGIC SETTING OF THE CORDILLERA SEPTENTRIONAL Ages and distribution of rock units
TECTONIC SETTING OF THE CORDILLERA SEPTENTRIONAL Plate-scale tectonic setting
The Cordillera Septentrional ('Northern Range') forms an elongate, east-northeast-trending mountain range that rises to a maximum elevation of 1249 m (Fig. 2). The range is partially bounded by seismically active, strike-slip and reverse faults related to left-lateral displacement between the North America and Caribbean plates across Hispaniola (Mann et al., 1991, 1998; Fig. 2B). Transpression across northern Hispaniola is probably a response to highly oblique subduction of the Bahama carbonate platform (Mullins et al., 1992; Mann et al., 1995; Dolan et al., 1998) (Fig. 2A). The relation between the general shape of the unsubducted Bahama Platform and the thrust front north of Hispaniola suggests that as much as 100 km of Bahamanian crust may been subducted beneath Hispaniola (Dolan et al., 1998). Local convergence in Hispaniola may also be related to the location of the island between the Bahama carbonate platform to the north and thicker-thanaverage oceanic plateau seafloor of the Caribbean Sea to the south (Mann et al., 1995; Diebold and Driscoll, Chapter 19; Figs. 1 and 2). Island-scale tectonic setting
A regional, unbalanced cross-section modified from Mann et al. (1991) illustrates several important features about the Cenozoic structural history of Hispaniola (Fig. 2B). (1) A prominent folding and thrusting event in central Hispaniola is Late Miocene and younger in age and verges southward to southwestward. (2) Late Miocene and younger reverse and oblique-slip faulting is responsible for the present pattern of morphotectonic units in central Hispaniola, including the distribution of the three major ramp, or thrust-bound, basins the Cibao, San Juan-Azua, and Enriquillo (Mann et al., 1991; Mann et al., Chapter 12). (3) Cretaceous-Eocene island-arc terranes of the northern and central part of the island are topographically high-standing and deeply eroded; the Cretaceous oceanic plateau terrane of the southern part of the island is relatively low-standing and less deeply eroded. The lower elevation of the oceanic plateau in the south may also reflect its footwall position relative to the higher-standing hanging wall block represented by the island-arc terranes in the north (Fig. 2B).
Basement units Igneous, metamorphic, and sedimentary rocks ranging in age from Cretaceous to Early Pliocene are exposed in the Cordillera Septentrional whereas siliciclastic sedimentary rocks of Mio-Pliocene age are exposed in the adjacent, asymmetric Cibao basin (de Zoeten and Mann, 1991) (Figs. 2B and 3). The oldest, arc-related rocks of the Cordillera Septentrional occur in three basement complexes of the Cordillera Septentrional that include the Samana Peninsula (Joyce, 1991) and Rio San Juan complex to the east of the study area, and the Puerto Plata complex (Pindell and Draper, 1991) to the north of the study area (Fig. 3). Mann et al. (1991) and Mann and Gordon (1996) proposed that all three inliers may have been uplifted by late Neogene restraining bend tectonics along the Septentrional fault zone (Fig. 3). Basement rocks of the Cordillera Septentrional and the Samana Peninsula can be divided into two lithologic provinces: a blueschist-serpentinite-sedimentary complex interpreted as an outer forearctrench assemblage by previous workers (Nagle, 1979; Bowin and Nagle, 1982; Pindell and Draper, 1991; Draper and Nagle, 1991; Joyce, 1991); and a volcanic-plutonic-sedimentary complex interpreted to be a forearc assemblage (Bowin and Nagle, 1982; de Zoeten and Mann, 1991; Calais et al., 1992). Based on a compilation of data, Mann et al. (1991) have classified these two areas as tectonostratigraphic terranes originating as fragments of the forearc or accretionary prism of the arc (Fig. 2B). We follow the block terminology of the central Cordillera Septentrional that was defined by de Zoeten and Mann (1991). This scheme subdivides the basement of our study area into the Altamira block to the west and the La Toca block to the east separated by the left-lateral Rio Grande fault zone (Fig. 4C). Basement of the La Toca block in the eastern part of the study area consists of Upper Cretaceous to Eocene andesitic tuff and tonalite of the Pedro Garcfa Formation (Eberle et al., 1982; Peralta-Villar, 1985) (Fig. 4C). Basement of the Altamira block in the western part of study area consists of Upper Paleocene to Lower Eocene pelagic carbonate rocks of the Los Hidalgos Formation that are crosscut by dikes and sills of the Palma Picada intrusives (Muff and Hernandez, 1986) (Fig. 4C). The Los Hidalgos Formation correlates well with the Paleocene-Eocene E1 Cacheal tufts described by Calais et al. (1992) in the western Cordillera Septentrional.
250
R. DE ZOETEN and E MANN
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA
251
Fig. 3. Map of northern Hispaniola showing ages of exposed rocks and the four main physiographic provinces which include the Cordillera Septentrional, Samana Peninsula, Cibao Valley, and Cordillera Septentrional. The area mapped in detail for this study is boxed and lies in the central part of the Cordillera Septentrional (cf. Fig. 4 for detailed maps). Major high-angle faults include: HFZ -- Hispaniola fault zone; SFZ = Septentrional fault zone; CFZ = Camti fault zone. Labelled basement complexes of the Cordillera Septentrional are discussed in the text.
Relation of major structures to outcrop pattern of sedimentary units Younger, s e d i m e n t a r y units above these three b a s e m e n t c o m p l e x e s g e n e r a l l y dip away from the inliers or are b o u n d e d by high-angle faults (Figs. 3 and 4C are based on a 1 : 1 5 0 , 0 0 0 geologic c o m p i l a t i o n map of the Cordillera Septentrional by de Z o e t e n et al., 1991). The ages of E o c e n e through Pliocene sedimentary units in the western Cibao basin and central and eastern Cordillera Septentrional defines a large synclinal structure with the axis of the syncline a p p r o x i m a t e l y parallel to the long axis of the Cibao basin (Fig. 3). The orientation and age of this postPliocene fold is consistent with the regional pattern of n o r t h e a s t - s o u t h w e s t shortening across Hispaniola seen on the cross-section in Fig. 2B. In our study area, ages of rock units indicate two large h a l f - d o m e or anticlinal structures adjacent to the Septentrional and Rio Bajabonico fault zones (Figs. 4C and 5). Tilting related to the Paradero half-
d o m e along the Septentrional fault zone accounts for the northeast dips o b s e r v e d over m u c h of the A l t a m i r a b l o c k (Fig. 5B). F o l d i n g related to the Pedro Garcfa anticline accounts for the northeast dips observed over m u c h of the La Toca block. The A1tamira fault zone abruptly truncates fold axes developed in the central part of the study area (Fig. 5A). Well-dated, U p p e r E o c e n e to L o w e r M i o c e n e d e e p - m a r i n e siliciclastic s e d i m e n t a r y rocks of the A l t a m i r a and Las Lavas formations u n c o n f o r m a b l y overlie arc-related b a s e m e n t of the A l t a m i r a block (Fig. 4B). T h e s e two formations together consist of about 4000 m of thin- to m e d i u m - b e d d e d sandstone i n t e r b e d d e d with c o n g l o m e r a t e . About 1200 m of O l i g o c e n e to L o w e r M i o c e n e siliciclastic sedimentary rocks of the La Toca F o r m a t i o n u n c o n f o r m a b l y overlie, or are locally faulted against, igneous rocks of the Pedro Garcfa F o r m a t i o n (Fig. 5). M i d d l e M i o c e n e to L o w e r Pliocene shallowmarine l i m e s t o n e of the Villa Trina F o r m a t i o n forms
Fig. 2. (A) Map of the northeastern Caribbean plate margin modified from Dolan et al. (1998). Crystalline basement rocks represent the exhumed core of the Caribbean island-arc that became inactive in Eocene to Oligocene time in this area. Rocks of the Bahama carbonate platform on the North America plate north of Hispaniola are part of a passive margin sequence formed following the Mesozoic rifting of North and South America. Rocks of the Caribbean plate and oceanic plateau were mainly formed as part of a large oceanic plateau in Late Cretaceous time. Convergence in Hispaniola is related to the local impedance of the eastward migration of the Caribbean plate by the salient formed by the Bahama Platform. Box shows map of northern Dominican Republic in Fig. 3. Key to abbreviations: OFZ = Oriente fault zone; SDB = Santiago deformed belt; EPGFZ = Enriquillo-Plantain Garden fault zone; SFZ = Septentrional fault zone; NHDB = North Hispaniola deformed belt; PRT = Puerto Rico trench; LMOB = Los Muertos deformed belt. (B) Cross-section across Hispaniola modified from Mann et al. (1991). Convergence between the Bahama Platform and the Caribbean oceanic plateau across the island of Hispaniola has led to extreme topographic uplift and erosion of the extinct arc core and the formation of three, thrust-bound ramp basins of Neogene age: the Enriquillo, San Juan and Cibao. Abbreviations of fault zones from north to south: CFZ = Cam6 fault zone; R G F Z = Rfo Grande fault zone; SFZ = Septentrional fault zone; HFZ - Hispaniola fault zone; GFZ = Guacara fault zone; SJRFZ = San Josd-Restauraci6n fault zone; SJLPFZ San Juan-Los Pozos fault zone; EPGFZ Enriquillo-Plantain Garden fault zone; and BAGFZ = Bahoruco fault zone. See Mann et al. (1991) for detailed descriptions of faults and terranes. - -
=
252
R. DE ZOETEN and E MANN
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA the youngest sedimentary unit in the Cordillera Septentrional and is at least 250 m thick (de Zoeten and Mann, 1991; Calais et al., 1992) (Fig. 6). The limestone exhibits both conformable and unconformable contacts with the underlying siliciclastic rocks of the La Toca and Altamira formations, respectively. Dips in the Villa Trina Formation define a large, post-Early Pliocene anticline that coincides with the topographically highest part of the Cordillera Septentrional (Fig. 6A). The angular contact between the Villa Trina Formation and the underlying siliciclastic formations is generally found in the western and central Cordillera Septentrional whereas conformable or disconformable contacts are found in the eastern part of the range (Calais et al., 1992).
PREVIOUS WORK AND STRATIGRAPHIC FRAMEWORK OF THE CENTRAL CORDILLERA SEPTENTRIONAL Previous studies
The stratigraphic nomenclature developed in previous studies of the central Cordillera Septentrional is summarized in Fig. 7. Previous work by Eberle et al. (1982) and Redmond (1982) on the sedimentary rocks of the Cordillera Septentrional has been on a regional scale (1 : 100,000) with little emphasis on detailed mapping of smaller areas. Vaughan et al. (1921) published the first reconnaissance map of the Cordillera Septentrional. Their proposed stratigraphy was based mainly on work done along the southern margin of the Cibao Valley and correlated units exposed in the Cibao Valley. Oil exploration by Dohm (1943) and Beall (1943) assembled a more detailed geologic map at a scale of 1 : 100,000 and used the nomenclature of Vaughan et al. (1921). Bermudez (1949) analyzed 22 samples from the Cordillera Septentrional for planktonic foraminifera to revise formation ages established previously by Dohm (1943); (Fig. 7). Bermudez (1949) found that the oldest sedimentary rocks of the 'Abuillot Formation' (the Los Hidalgos Formation in this paper) consisted of hard, thin-bedded limestone containing Lower Eocene radiolaria and planktonic foraminifera.
253
Eberle et al. (1982) compiled a lithologic and structural map of the central and eastern Cordillera Septentrional, based on reconnaissance mapping and detailed biostratigraphic sampling of calcareous nannofossils. One formation, the E1 Mamey, was established to include all siliciclastic rocks of Eocene to Early Miocene age. The E1 Mamey Formation was divided into two, laterally equivalent 'facies': the Luperon facies north of the Cam6 fault zone, and the Altamira facies to the south (Fig. 7). Eberle et al. (1982) were also the first to recognize the intrusive rocks in the Palma Picada area south of E1 Mamey (Figs. 4C and 7). The study by Redmond (1982) focused on the sedimentology of Cenozoic sedimentary rocks of the central Cordillera Septentrional and their relation to the occurrence of amber deposits in the area. Redmond refers to these turbiditic rocks as the Altamira Formation after the town of Altamira along the SantiagoAltamira highway (Fig. 4A). A Late Eocene age was determined for the Altamira Formation based on the planktonic and benthonic foraminifera from ten samples (E. Robinson, in Redmond, 1982). Calais et al. (1992) carried out stratigraphic and structural mapping in the western Cordillera Septentrional that dated a Lower Paleocene-Eocene fine-grained section equivalent to the Los Hidalgos Formation (El Cacheal tufts) unconformably overlain by a Lower Miocene to Upper Pliocene siliciclastic section (Gran Mangle and Villa Vasquez series). Stratigraphic n o m e n c l a t u r e of this study
We propose a stratigraphic scheme shown in the far right column of Fig. 7 based on data presented in this paper. We measured over 9000 m of stratigraphic sections and integrated these with 199 biostratigraphic analyses. Of these 199 biostratigraphic analyses, 136 were done on samples collected specifically for this study (all biostratigraphic data shown on Fig. 4B are compiled in Appendix 1 of de Zoeten, 1988). Combining biostratigraphic data with detailed lithostratigraphy, we recognized that the Altamira and La Toca blocks were characterized by distinct stratigraphy and sandstone composition (de Zoeten and Mann, 1991) (Fig. 4C). These two blocks are separated by the Rfo Grande fault zone, a
Fig. 4. (A) Map of central Cordillera Septentrional showing major towns, road system and streams with outcrops that are referred to in the text of this paper. Paved roads are shown by heavy lines; unpaved, secondary roads in the mid-1980's are shown as dashed lines. (B) Map of the central Cordillera Septentrional showing the age of exposed rock units based on microfossils from 140 sample localities shown by circles. Patterns indicate ages which and are constrained by both biostratigraphy and stratigraphic relationships. Microfossils used in this compilation include calcareous nannofossils, planktonic foraminifera, and benthonic foraminifera. Appendix 1 in de Zoeten (1988) provides a list of microfossils identified from each sample site. (C) Map showing lithostratigraphy and formations of the central Cordillera Septentrional. The study area is limited to the area between the Septentrional fault zone (SFZ) and the Camfi fault zone (CFZ), and is separated into the Altamira block and the La Toca block by the Rfo Grande fault zone (RGFZ). Key to abbreviations: A F Z = Altamira fault zone; RBFZ = Rio Bajabonico fault zone; CBA = Canada Bonita anticline; LS = Llanos syncline.
254
R. D E Z O E T E N and R M A N N
Fig. 5. (A) Major structural features of the central Cordillera Septentrional modified from de Zoeten and Mann (1991). Note the en echelon arrangement of major anticline and half-domes, which are shaded in gray. Key to numbered folds: 1 = Paradero half-dome; 2 = Canada Bonita syncline; 3 = Llanos syncline; 4 = Ocampo anticline; 5 = Pedro Garcfa anticline; 6 = Sonador anticline. Lines A - A I, B - B I, and C - U indicate cross-sections shown in (B). (B) One-to-one cross-sections of the central Cordillera Septentrional. Key to rock units: KTpg = Pedro Garcfa Formation (shaded); Tlh = Los Hidalgos Formation (shaded); Tp = Palma Picada intrusive rocks (shaded); Ta = Altamira Formation; Tt = La Toca Formation; T1 = Las Lavas Formation; Tvt = Villa Trina Formation. Heavy dot pattern indicates Mio-Pliocene sedimentary rocks of the Cibao basin; wavy lines are unconformities. SFZ -- Septentrional fault zone; R G F Z = Rfo Grande fault zone; CFZ = Camti fault zone. Dip symbols represent the dip of beds measured in outcrop.
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA
255
Fig. 6. Map showing the distribution and structure of the Villa Trina Formation in the Cordillera Septentrional. The elevation of the base of the Villa Trina Formation above sea level was determined from 1:50,000 topographic maps.
Fig. 7. Comparison of stratigraphic columns for the central Cordillera Septentrional including the stratigraphy proposed on the basis of data presented in this paper. Most previous work was reconnaissance in nature. This study used systematic biostratigraphy and detailed measured sections to better establish the character and age of lithologic units and correlate them across the central part of the mountain range. 100-400-m-wide, oblique-slip shear zone (de Zoeten and Mann, 1991). The name E1 M a m e y Group was proposed by Dolan et al. (1991) and in this paper to include all the Upper Eocene to Lower Miocene sedimentary rocks in the Cordillera Septentrional. The Altamira block
comprises the Altamira and Las Lavas formations and the La Toca block comprises the La Toca and Luperon formations (Figs. 4C and 7). The Luperon Formation has been described previously by Nagle (1979), Eberle et al. (1982), and Pindell and Draper ( 1991) (Fig. 7).
256
R. DE ZOETEN and R MANN
~
r~
r~ c~
c~
o
m
.~
0
o
~ ~ . ~:~ .,~ . ~
~
~
~
.,.,
.,.,
,.~
~ ,...,
~
;>
_=~.~:~
_
~
~,
~
-~ ..~
9
~
m.o
~
9
~
:~
o ~
~ .,.-, z~
~
~
;>, r~ ,..C:~
o
o
=~ '~ =~
~
~=~
E
0
~
c~
,o
9
~;~
r,.) ~ , . - ,.~
o
=o-~
~o
:~~ ~
m
m
o
o
.~.-~..'~.~.= ~ "~ :~ ~
"~ ~
= .-
o
0
i
i
i
i
b~
b~
o
o
o
.~
o ,"~
"~ .~
.
.~=.~
=o~
~o
~.~
~
~o
.
~
~
9,-.
o
o~
.~'~
~
.,,_.
o=
.,.. ,.r
~
0"~
0
o;~
~
;~.~
~
*"
"c:~ , ' ~
.~
.~
.~-
~ . ~ .~ o~,~ ~'~ .~ ,~ ,~
~
M
..,
~,~E
~
~
.
..~ .~ .~
0
.-,
o
c:= Ei
;~,.
~
.o.~
_
o,.~
o
"~
9 m
.,..~., o
~
,..-~ r~
=,-~
~
o o
=o ~ ~" . - . ~ r~
~,
~
o,..
m
'" 0
o~
~.~
r~l
r~
~s
r~
;
.~ ~= ~"
~
~
=~.~
~
o ~,
~o'~
~a
~
o
0
o
9,-.
i=
o
.,--~
.o
~"=
r./3 ..c:~ i
i
i
;>,
9"
-
, ~ ""
o
="=
.-.=
.~
9
~E
o
= =~'~
o
~'~ o = . ~
~-
~.~ ~ " ~
=.-= o
.~-.-.-~ ~
~"
,~==
-.-
==~i==~
~ ~
O~ 0
=~
-
"~
o ~ =~ ~ ~
~
o
=~
..
" ..
_o~.~ ~..=
~
~ . ~
~,T..
,
0
0
r~
!
~
0
.~ ~ . ~ =
~o
'c~ '-'-
,
~
~~
~
=. == . . . .
0
0
,
.~
.
o ~
c~
.
~
~.~ ~r~
~ 0
.
r~ ~1
.
. ,.._
~
c~ ~
r~ ~ r ""
~
~.,
r~
~o
~
,-~
0
0
.~...,
== ~ . ~
iaa==
==
;=" " - r.~
.
r
,.
0
cn
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA Facies classification used in this study
In our stratigraphic and sedimentologic study of the Altamira, Las Lavas and La Toca formations, we used the facies classification for deep-marine siliciclastic rocks that was developed by Pickering et al. (1986). This classification, a modification of Mutti and Ricci Lucchi's (1978) lithofacies classification, facilitates facies descriptions and interpretations in the field. The Pickering et al. (1986) classification is a comprehensive and purely descriptive scheme used to subdivide lithologies into mappable facies. Four of the seven facies classes proposed by Pickering et al. (1986) were recognized in the central Cordillera Septentrional. These commonly seen facies are summarized on Fig. 8. The letter-number code shown in Fig. 8 modified from Pickering et al. (1986) is used for all measured sections and outcrop photographs presented in this paper. Biostratigraphic age determinations used in this study
Ages of stratigraphic units are based on their fossil content (Fig. 4B). Reworked older fauna, however, are an inherent problem associated with resedimented turbidite deposits characteristic of the E1 Mamey Group and so the youngest ages were picked to represent the time of deposition. Location grid used in this study
The Universal Transverse Mercator (UTM) grid location system is used throughout this paper to precisely locate outcrops discussed in the text. This grid is printed on U.S. Defense Mapping Agency 1:50,000 topographic maps which were used as basemaps for our field investigations (see Preface of Mann et al., 1991, for a key to these maps in the Dominican Republic; the grid zone is 19Q).
STRATIGRAPHY OF THE ALTAMIRABLOCK Definition of the Altamira block
The Altamira block is bounded on the north by the Cam~ fault zone (Pindell and Draper, 1991), on the east by the Rio Grande fault zone (de Zoeten and Mann, 1991), and on the south by the Septentrional fault zone (Mann et al., 1998; Fig. 5A). The basement rocks of the Altamira block comprise Upper Paleocene to Lower Eocene, sedimentary and igneous rocks of the Los Hidalgos Formation. The deep-marine siliciclastic rocks of the Altamira Formation sit unconformably over the Los Hidalgos Formation and its intrusive Palma Picada
257
rocks south of E1 Mamey (UTM 847816). The Las Lavas Formation, which conformably overlies the Altamira Formation, consists of calcareous and terrigenous, deep-marine turbidite and other mass-flow deposits. The topographically higher regions of the Altamira block are capped by Upper Miocene to Lower Pliocene, shallow-marine carbonate rocks of the Villa Trina Formation (Fig. 6). Basement of the Altamira block Los Hidalgos Formation The Los Hidalgos Formation is composed of thinly laminated to medium-bedded, dark gray, red and green recrystallized biomicrite interbedded with minor amounts of volcaniclastic calciturbidites and tuffaceous shale (argillite). The base of the Los Hidalgos Formation is not exposed, but the formation is at least 250 m thick in single outcrops. The formation is exposed in three localities in the study area: (1) in an irregular-shaped area 3-12 km south of E1 Mamey (UTM 847816); (2) on a ridge 0.4 km wide, and 5 km long just north of Altamira (UTM 096752); and (3) as a fault-bounded sliver within the Rio Grande fault zone about 3 km north of Santiago (UTM 299594; Fig. 4C). Based on its outcrop distribution at these three localities, the Los Hidalgos Formation appears to underlie most of the Altamira block. Thin parallel laminae in the recrystallized biomicrites are laterally continuous, but locally bioturbated by Skolithos and Planolites. Laminated limestones contain matrix-supported, silt-sized, angular plagioclase grains and ellipsoidal silt-sized grains, interpreted to be calcified radiolarians or recrystallized planktonic globigerinid foraminifera. Acarinina and Morozovella spp. have been identified in thin-sections (E. Robinson, pers. commun., 1988) suggesting that the rocks range in age from Late Paleocene through Early Eocene and that they were deposited in bathyal to abyssal water depths (1506000 m; cf. Appendix 1 of de Zoeten, 1988). These rocks contain stylolites which are crosscut by complex micro- to mesoscopic calcite veins. The structure of the Los Hidalgos Formation is described in more detail by de Zoeten and Mann (1991) and Calais et al. (1992). Thin- to medium-bedded calciturbidites (1-30 cm) with partial Bouma sequences (Tad, Tae) are interbedded with the laminated limestone unit. The coarser fraction consists of poorly graded, fine- to medium-grained plagioclase crystals, volcanic rock fragments and bioclasts. An abrupt grain-size change occurs between the basal sand unit and the overlying recrystallized limestone, which composes the pelagic division (Te) of the sand bed. The limestone unit is commonly bioturbated.
258 Thin parallel laminae and abundant planktonic fossils in the Los Hidalgos Formation indicate slow deposition from suspension in a deep-marine environment. Pelagic to hemipelagic accumulation in the basin was periodically interrupted by low-density turbidity flows or other mass-flow processes introducing the coarser clastic material of the Ta turbidite intervals.
Palma Picada intrusive rocks Porphyritic rocks of the Palma Picada Formation intrude the Los Hidalgos Formation sedimentary rocks near Palma Picada (UTM 922778) (Fig. 4C). These compositionally diverse, porphyritic rocks consist of a series of vertical dikes and horizontal sills, which are approximately 250 m thick (Eberle et al., 1982; Muff and Hernandez, 1986). Stratigraphy of the Altamira Formation Outcrop distribution and general stratigraphy The Altamira Formation extends over a 200 km 2 area from the Rfo Grande fault zone to a poorly defined western limit near E1 Mamey (Fig. 4C). The Altamira Formation consists of thin- to medium-bedded sandstone and siltstone couplets, with minor interbedded conglomerate and thick-bedded sandstone. The Altamira Formation is divided into two members: (1) a 50-m-thick basal conglomerate, the Ranchete Member, which makes up a minor part of the total thickness of the Altamira Formation, and lies unconformably above rocks of the Los Hidalgos Formation, and the Palma Picada intrusions (UTM 906815; Fig. 4C); and (2) the Canada Bonita Member, an approximately 2500-m-thick section of alternating sandstone and siltstone with interbedded conglomerate; this member composes most of the thickness of the Altamira Formation (Fig. 4C). The Ranchete Member of the Altamira Formation is named here after the village of Ranchete, which is located 3.5 km southwest of E1 Mamey (UTM 838824; Fig. 4A). In this area, the Ranchete Member is exposed as a thin (<130-m-wide), 3.5-km-long belt rimming the northern margin of the Los Hidalgos Formation (Fig. 4C). Sedimentary rocks of the Altamira Formation overlie the basal conglomerate of the Ranchete Member in an arcuate belt stretching 40 km across the center of the study area (Fig. 4C). This unit is here named the Canada Bonita Member of the A1tamira Formation, after the village of Canada Bonita, 7 km north of Navarette (UTM 064703; Fig. 4A). The Canada Bonita Member is composed of 80% very thin- to medium-bedded, sandstone and siltstone couplets (facies types: C2.2, C2.3, D2.2), 15% conglomerate (AI.1, A2.1), and 5% thick-bedded sandstone (B2.1, C2.1).
R. DE ZOETEN and E MANN Sandstone and siltstone of the Canada Bonita Member are blue-gray, calcite-cemented, feldspathic litharenite, which weathers to an orange-tan color. The Bouma facies Tde of the sandstone and siltstone couplet of the Altamira Formation typically lack mud-sized particles. These upper divisions consist predominantly of coarse- to fine-grained silt with only minor clay. Clasts in conglomerates of the Canada Bonita Member range in size from granules to boulders, but most commonly range in size from pebbles to cobbles. The clasts are equidimensional or oblate in shape, and subrounded to well rounded. In general, clasts are composed of: (1) recrystallized limestone (~60%), including biomicrite, dark-gray, green, and banded argillite, and carbonaceous silt derived from the underlying Los Hidalgos Formation; (2) plutonic porphyries (~20%); (3) bioclastic limestone (~10%); and (4) sandstone and volcanic fragments (,-10%). The conglomerate matrix is a gray to light brown, fossiliferous volcaniclastic sand.
Age and paleobathymetry Biostratigraphic analysis on 37 samples indicates that the Altamira Formation ranges in age from Middle or Late Eocene to Late Oligocene (Fig. 4B). Ten samples from the lower part of the Altamira Formation exposed near E1 Mamey have been well dated as Middle to Late Eocene using calcareous nannofossils and foraminifera (Appendix 1 in de Zoeten, 1988). These dates constrain the upper age limit of the underlying Ranchete Member as Late Eocene. The lower age limit of the Ranchete Member is constrained by the Early Eocene age of the underlying Los Hidalgos Formation. Several samples from the Canada Bonita Member collected along the Santiago-Puerto Plata highway suggest Late Eocene ages (Appendix 1 in de Zoeten, 1988) (Fig. 4C). However, Bourgois et al. (1982, 1983) and S. Monechi (pers. commun., 1988) determined that three samples from beds stratigraphically below the Upper Eocene sample localities contain distinct Oligocene faunas (Appendix 1 in de Zoeten, 1988). No northwest-striking faults separating the sample localities along the Santiago-Puerto Plata highway were recognized (de Zoeten and Mann, 1991) (Fig. 4C). This suggests that Upper Eocene calcareous nannofossils are reworked and that deposition of the Altamira Formation continued into Late Oligocene time as we show on Fig. 7. Paleoenvironmental studies on the benthonic foraminiferal assemblage indicate upper bathyal water depths (150500 m). Measured sections of the Altamira Formation Seven sections were measured from west to east to better characterize rapid lateral facies changes
259
CENOZOIC EL M A M E Y GROUP OF NORTHERN HISPANIOLA
Plata highway (Fig. 11, inset map). The rocks are gently folded into an anticline-syncline pair whose structure is described in detail by de Zoeten and Mann (1991). These measured sections all describe variations within a 1000-1500-m-thick section of lateral equivalent rocks of the upper Canada Bonita Member of the Altamira Formation (Fig. 7). The Calabaza measured section is separated from the Northern Canada Bonita and Southern Canada Bonita sections by the north-south striking Altamira fault zone (Eberle et al., 1982; Figs. 4C and 5). Despite the structural complexity related to folding and the presence of the Altamira fault zone, there are many similarities between all five sections near the Santiago-Puerto Plata highway that allow the various sections to be correlated. For example, all three section are overlain by the basal conglomerate
within the Altamira Formation. The measured sections are divided into three sets based on geographic location and similar facies patterns. From west to east, the three sets are: (1) E1 Mamey and Guananico, (2) Southern Canada Bonita, Northern Canada Bonita, and Calabaza, and (3) Rio Perez and Llanos syncline (Fig. 4A). The basal conglomerate of the Ranchete Member was only measured in the westernmost E1Mamey section (Fig. 9, inset map; Fig. 10). The Guananico section lies 10 km to the west of Altamira and lies along strike of the E1Mamey section (Fig. 9, inset map). The Guananico and E1 Mamey sections are dominated by a repetitive series of thin- to medium-bedded sandstones and siltstones (C2.2, C2.3, D2.2 facies). Five sections of the Altamira Formation were measured along or adjacent to the Santiago-Puerto A. El M a m e y section" Altamira Fm.
B. Guananico section: Altamira Fm.
90C C2.1
m
TOP
r,, !
2 - 2.3
400
400
\ )2.1 \
r \
3 \
80C
\ \
300
'
l
\
\
r~
\ \ \
\
A1.1 OffsettoB
)G
\
\
\
\ 600
- 2.3 / 1~2.2
(~ F - G \ 600 B2.1 / C2.1 \
TOP
t
\\
u. Eoc
\
\m
1u Eoc \\\
100 ~A~ ~
\
soo
..... ', ~ Breccia (Fig. 10) 1 LOS tt=calg(s IJ~ U P a l l Eoc ForlTl(tion 0 ~ l i " ~ F " "" (Eocene) III .... Grain Size I I Gravel =~o,IMed. Sand
groove cast ripplemark
Mio Miocene Olig Oligocene Eoc Eocene
\
C2..2 2387
\ \ I
)G
C 2 . 2 - 2 . 3 ! D2.2
500t
C2.3
~-"- F
BASE I Gravel Med. Sand
Silt G R
\
100
~ooo
\
A22
' - 2.3 / D2.2
\
200
\
,2..2-2.3 \
700
\
\ \
, ~
\
300
\ 700
200
\ \
~2 2-'23/D22 \
=E,
\
~t
\
800
\
Silt F C
flute cast a-axis clast orientation
u m I
upper middle lower
Very thin to med. bedded Med. to very thick bedded 13;El Covered interval
" - ~
Stratigraphyinferred Fault Facies type
A2.1 2387 Sandstone sample (~ Paleocurrent direction
I
Offset in El Mamey "measured section
IE'MameY]--//~ I
41 J l l " / / " f =
~
0 5 km I
''
I
Guananico]
I ~,Mamo,sec,~on iJ "~ I II Guananico[ ~
I~Y:;," ~
se='~ I
%~
I
- Map trace of measured section
Fig. 9. Measured sections from the E1 Mamey (A) and Guananico (B) localities (cf. inset map for location of sections). The sections summarize facies types, paleocurrent indicators, biostratigraphic samples with age determinations, and sandstone sample localities used for point-count analyses. See Fig. 8 for explanation of letter-number codes of facies types. Note along-strike offset in the E1 Mamey section at 120 m above the base.
260
R. DE ZOETEN and E MANN
Fig. 10. Normally graded, sedimentary breccias from the Middle(?) to Late Eocene Ranchete Member of the Altamira Formation (location of photograph shown on section in Fig. 9). Clasts are composed of grey, recrystallized biomicrites of the underlying Upper Paleocene to Lower Eocene Los Hidalgos Formation. Outcrop is exposed in roadcut just south of Los Hidalgos Pass (UTM 847816). Jacob's staff is divided into 0.5 m increments. Dip to northeast. Color version at http://www.elsevier.nl/locate/caribas/
of the Las Lavas Formation and all three sections consist predominantly of thin- to medium-bedded sandstone and siltstone (C2.2, C2.3, D2.2) and associated conglomerate (AI.1, A2.1). The Calabaza section consists of a higher percentage of conglomerate (~50%) than the Northern Canada Bonita and Southern Canada Bonita sections (12%). The Rfo Perez and Llanos syncline sections are south and southwest of Altamira (Fig. 11, inset map). These sections consist of thick-bedded sandstonesiltstone couplets (B2.1, C2.1 facies), conglomerates (AI.1, A2.1 facies) and thin- to medium-bedded sandstones and siltstones (C2.2, C2.3, D2.2 facies). El Mamey and Guananico measured sections of the Altamira Formation Ranchete Member of the Altamira Formation The E1Mamey section is a composite section based on two traverses (Fig. 9, inset map). The basal 54 m of the section forms the type section of the Ranchete Member, which is exposed along the road near Los Hidalgos Pass (UTM 846817) (Fig. 4A). The overlying section was measured in the Arroyo Berraco Blanco stratigraphic section (UTM 927844; Fig. 4A). The contact between the Ranchete Member and the underlying Los Hidalgos Formation is unconformable, although we found that this contact is locally modified by faulting at two localities (UTM 846817; 871825). The lower part of the Ranchete Member consists of a massive, tightly packed sed-
imentary breccia (AI.1), which is approximately 15 m thick. The breccia grades into thick-bedded, coarse-tail graded conglomerates (A2.3; Fig. 10) and is capped by medium-bedded, parallel-stratified conglomerates (A2.1). Clast rounding increases upward in the section from angular in the A2.3 facies to subrounded clasts in the A2.1 facies. Gravel to cobble size, angular clasts near the base of the Ranchete Member are composed entirely of gray and dark gray, recrystallized limestone, which reflect the lithology of the underlying basement rocks of the Los Hidalgos Formation (Fig. 10). The percentage of recrystallized limestone clasts decreases upward in the section at the expense of marly packstone clasts composed mainly of red algae and larger reefal foraminifera. Lepidocyclina foraminifera in the packstone clasts suggest that they formed in a backreef environment (E. Robinson, Appendix 1 in de Zoeten, 1988). Laterally equivalent conglomerate to the northwest (UTM 906815) contain up to 50% dark purple porphyry clasts of the Palma Picada intrusions. This conglomerate also exhibits an upsection increase in shallow-water carbonate clasts. Canada Bonita Member of the Altamira Formation The overlying Canada Bonita Member in the E1 Mamey and Guananico sections consists mostly of thin- to medium-bedded, alternating sandstones and siltstones, which are mostly facies D2.2, D 1.2, C2.3, with minor C2.2 (Fig. 12). These facies are interbed-
CENOZOIC EL M A M E Y GROUP OF NORTHERN HISPANIOLA
261 O
,.~
O
""'~
O
O
o.~
O
O
.,,~
,,,.o
or,.)
E~
O O
rs~
.,.-~
O,, ~
~Z O ~
~
9
o
-~ r,..) ~
~< 0
~Zo
=~ ~
~,~ .~,.~
NN~
o ~c~
~
~ o
-~
,.~
" ~
~
r~
~ g.~
262
R. DE ZOETEN and R M A N N
9
r,.)
9 .,,,.
9
O
~,,,,~ 9 9
.=. 9
.< 9
N
r,.) .~
~ .,,a
r,,) ~: O
(D
Z
O
<~
r,.)
~
O
,x::
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA ded with minor, thick-bedded sandstone and siltstone couplets (C2.1). Conglomeratic facies (AI.1, A2.3) are rare, and mostly found near the base of the Canada Member. Thin-bedded sandstone and siltstone couplets of the Canada Bonita Member have a sand/silt ratio greater than 1 : 1, are tabular, laterally continuous for at least 10 m, exhibit sharp basal contacts, and are overlain by medium- to coarse-grained, structureless or poorly graded sand (Tbd, Tabd). Thin parallel laminae are very common throughout the lower division of beds. These laminae are defined by 0.5-4 mm biogenic carbonate grains composed mostly of red algae and benthonic foraminifera. The silty upper division of C2.2 and C2.3 sandstone and siltstone couplets contain isolated stringers of medium to coarse sand, are locally parallel-laminated, and are commonly bioturbated by Skolithos, Planolites, and Glockeria. The thick-bedded sandstones of the Canada Bonita Member have sand/silt ratios greater than 1:1 and exhibit erosional basal contacts with local load and groove casts. The lower division of beds are massive or coarse-tail graded and capped by parallel laminae which are defined by concentrations of coarser-grained carbonate material (Tbd, Tabd). Siltstones are commonly thin (3-10 cm), mud-poor, parallel-laminated, and occasionally show bioturbation by Skolithos and Glockeria. Beds are tabular and exhibit good lateral continuity at the outcrop scale. No vertical cycles were recognized in the Canada Bonita Member, except for rare, thin (1-3 m) thickening-upward cycles within rhythmically bedded, thin, sand-silt couplets (Fig. 12). Two 2-7-m-thick, thinning-upward intervals of C2.1-C2.3 facies are found 450 and 470 m above the base of the Guananico section (Fig. 9). Both intervals are composed of lithic-rich calcarenites (Tabd). These inferrals include the only calciturbidites found in the Altamira Formation. Their lateral extent or geometry could not be determined. Twenty-eight groove casts measured from the bottom of medium- to thick-bedded sandstones indicate a northwest to southeast (110 ~ mean paleocurrent trend. In the Guananico section, eight unimodal flute casts indicated paleoflow to the southeast (Fig. 9). Northern and Southern Canada Bonita measured sections of the Altamira Formation Canada Bonita Member of the Altamira Formation The Northern and Southern Canada Bonita sections consist mostly of very thin- to medium-bedded sandstone and siltstone couplets (C2.2, C2.3, D2.2), interbedded with minor conglomeratic facies (AI.1, A2.1), and thick-bedded sandstones (C2.1) (Fig. 12A,B). At the base of the Southern Canada Bonita
263
section to the south of the anticlinal axis (Canada Bonita anticline) (Fig. 11, inset), a conglomeratic sequence 40 m thick displays two depositional units. The base of the upper conglomerate unit is scoured at least 1 m into the underlying conglomerate (Fig. 13A). Both conglomerate units grade upwards from predominantly pebble to boulder, clastsupported disorganized conglomerates (AI.1 facies) (Fig. 14A), to pebble to cobble, poorly stratified conglomerate (A2.1, A2.5 facies) (Fig. 14B). Although poorly stratified conglomerate appears massive, the preferred orientation of the a-axis of conglomerate clasts parallel to bedding produces a stratified appearance in the conglomerate (Fig. 14B). Smaller conglomerate bodies (A 1.1, A2.1) form 2-5-m-thick concave-down channel structures within alternating thin-bedded sandstones and siltstones (630 m from the base of the Northern Canada Bonita section, and 170 m from the base af the Southern Canada Bonita section) (Fig. 12A). The lower part of the Northern Canada Bonita section (Fig. 12B) is dominated by thin- to medium-bedded sandstone and siltstone couplets (C2.2, C2.3, D2.2). These couplets have a 1:3 sand/silt ratio. Basal sands are medium- to very coarse-grained and massive to thinly laminated. Basal sands occasionally exhibit normal, inverse, and coarse-tail grading (Tdb). A sharp contact divides the upper and lower divisions of couplets. The upper division of most couplets are silt-rich, but increasing clay-sized fraction causes the beds to be transitional into the D2.2 and D1.2 facies. The siltstone contains parallel laminations, commonly bioturbated by Skolithos and Planulites, and contains lignite fragments with a-axes ranging from 0.5 to 10 cm. The a-axis of lignite fragments are oriented parallel to laminations. In the Td intervals, coarse-grained sand stringers and lenses are oriented parallel to bedding. Locally, outsized carbonate blocks with longest axes ranging from 25 to 160 cm are found isolated in the thin- to medium-bedded couplets (410 and 585 m from the base of the Northern Canada Bonita section, 225 m from the base of the Southern Canada Bonita section). In general, conglomerate displays thinning- and fining-up cycles with increasing internal organization of clasts (Fig. 13A). Thinning- and fining-up cycles are poorly expressed in the thin- to medium-bedded sandstone and siltstone couplets (C2.2, C2.3, D2.2), which range in thickness from 5 to 300 m. Thirty-three paleocurrent indicators were measured using sole marks, tipple marks, and orientation of a-axes of conglomerate clasts. At least fifteen measurements of the a-axes of conglomerate clasts were averaged to a mean direction, which is shown on the section in Fig. 12. The orientations of bimodal paleocurrent measurements were scattered, but generally show either N-S or NNW-SSE bidirectional
264
R. DE ZOETEN and E MANN
Fig. 13. (A) Drawing from photographs of conglomerate exposed in a roadcut along the Santiago-Puerto Plata highway (UTM 062706). Outcrop consists of two vertically stacked, fining-upwards conglomerate bodies within the Upper Oligocene Canada Bonita Member. This outcrop occurs 60 m above the base of the Southern Canada Bonita measured section shown in Fig. 12A. Letter-number codes indicate facies types of Pickering et al. (1986) (see Fig. 8 for explanation). Location of photograph shown in Fig. 14A is indicated by box. (B) Drawing from photographs of a poorly stratified conglomerate (A2.1) which forms the basal unit of the E1 Limon Member of the Las Lavas Formation and crops out in a roadcut along the Santiago-Puerto Plata highway (UTM 072690). Note the sharp contact with thinly bedded sandstones and siltstones (C2.3). paleoflow. Three unimodal measurements indicate paleocurrent flow both to the south and north.
Calabaza measured section of the Altamira Formation Canada Bonita Member of the Altamira Formation The base of the section starts at a 350-mthick, faulted conglomerate ridge (AI.1, A2.1) and is capped by a 400-m-thick section of thinto medium-bedded sandstone and siltstone couplets (Fig. 12C). This conglomerate is predominantly massive, disorganized, clast-supported, and decreases in thickness upward in the section. The few basal contacts observed are erosive. Thin- to medium-bedded sandstone and siltstone (C2.1, C2.2, D2.2) is tabular and laterally continuous for 5 m or more. These beds consist of silt to coarse-grained sand, and exhibit a graded or structureless lower division and an upper division characterized by parallel laminae (Tabd, Tbd). Rare
30-90-cm-thick beds (C2.1) are interbedded with the thin- to medium-bedded sand and silt couplets. These thicker beds are graded, and commonly contain groove marks (Tabd). The Calabaza section shows a distinct 1400 m thinning- and fining-up cycle, in which conglomerate decreases in thickness and abundance upward (Fig. 12C). Only one 10-m-thick thinning- and fining-up cycle was recognized. Eleven bidirectional paleocurrent indicators suggest northeast-southwest (70 ~ paleoflow (Fig. 12C). Well-preserved flute casts measured within 10 m of the base of the Las Lavas Formation indicate a paleoflow towards the northwest (285~
Rio Perez measured section of the Altamira Formation Canada Bonita Member of the Altamira Formation The R/o Perez section (Fig. l lA) is laterally equivalent to the Northern Canada Bonita, Southern
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA
265
Fig. 14. (A) Disorganized, clast-supported conglomerate (type A I.1) observed in the Upper Oligocene Canada Bonita Member. Location of photo is shown in drawing of Fig. 13A. Bar scale on card is 15 cm long. (B) Poorly stratified, clast-supported conglomerate of the Upper Oligocene Canada Bonita Member. Note absence of stratal boundaries and well-developed, parallel orientation of oblate clasts. Bar scale on card is 15 cm long. Color version at http://www.elsevier.nl/locate/caribas/ Canada Bonita, and Calabaza sections (Fig. 12). It is composed of 60% alternating, thin- to medium-bedded sandstone and siltstone (C2.2, C2.2, D2.2), 20% thick-bedded sandstone and siltstone couplets (B2.1, C2.1), and 20% conglomeratic facies (AI.1, A2.1). Characteristics of the thin- to medium-bedded facies are identical to those described for the Northern Canada Bonita and Southern Canada Bonita sections. Thick-bedded sandstones (B2.1, C2.1) are tabular and laterally continuous over distances of 10 m
and have a sandstone to siltstone ratio much greater than 1:1 (average 4 : 1 ) (Fig. l l A ) . Bouma intervals include Tbd and Tabd. The lower Ta intervals are medium- to coarse-grained, massive, coarsetail graded, or amalgamated. Tb intervals are very common. Higher concentrations of red algae, larger reefal foraminifera, and plant fragments define parallel laminae and rare cross-bedding. The overlying siltstone of the Td interval also contains abundant plant and lignite debris. The siltstone is parallel-lam-
266 inated, bioturbated, and rarely exhibit ripple marks. Conglomerate of the Rio Perez section consist of massive, clast-supported, disorganized facies (AI.1), and thick-bedded, clast-supported, stratified and poorly graded facies (A2.1, A2.3) (Fig. l lA). Conglomerate horizons range in thickness from 1 to 50 m. Basal contacts are scoured. Lateral geometry is difficult to discern on outcrop scale, but some smaller-scale beds are discontinuous over distances of 5-20 m and exhibit convex-downward channel cross-sections (e.g., 80 m from the base of Rio Perez section; Fig. 11). A total of 32 groove marks and flute clasts were measured from the bottom of C2.1 sandstone beds in the Rio Perez section (Fig. l lA). Twenty-five bidirectional measurements indicate a northwestsoutheast (~120 ~ paleocurrent trend, and seven unimodal current indicators suggest flow to the southeast. There are five coarsening- and thickening-upward sequences (5-20 m thick) capped by conglomerates (A1.1, A2.1). This coarsening- and thickening-upwards trend appears to reverse to a fining- and thinning-upwards trend toward the top of the section (400 m from the base), where beds define three, 10-50-m-thick, thinning- and fining-up cycles. Llanos syncline measured section of the Altamira Formation Canada Bonita Member of the Altamira Formation The Llanos syncline section is composed of about 70% thick-bedded sandstone and siltstone couplets (C2.1, B2.1), 25% thin- and medium-bedded facies (C2.2, C2.3, D2.2) and 5% conglomeratic facies (AI.1, A2.1) (Fig. llB). The basal contact of the Canada Bonita Member with the Los Hidalgos Formation is obscured by vegetation. Thick-bedded sandstone is predominantly C2.1 facies type, tabular-shaped, and laterally continuous. However, a few sandstone beds form discontinuous, broad convex-down channels (450 m above the base of the section). Although facies assemblages are quite different between the Llanos syncline and the Rio Perez sections, individual facies characteristics are similar. In the Llanos syncline section alternating sandstone and siltstone beds (B2.1, C2.1, C2.2, C2.3) form packages of 3-20-m-thick, thickening- and coarsening-upwards cycles. These cycles are rarely capped by conglomerate-filled (AI.1), concave-down scours which range from 1 to 5 m in thickness (Fig. 11B). B imodal paleocurrent indicators, based on 27 groove casts, suggest an east-west to northwestsoutheast (90-120 ~ flow direction (Fig. l lB). Unimodal measurements from flute casts and tipple marks indicate that the paleocurrents flowed towards the southeast (135~
R. DE ZOETEN and R MANN Facies analysis of the Altamira Formation Facies of the Ranchete Member The basal Ranchete Member breccia suggests local erosion and limited transport of the underlying Los Hidalgos Formation (Fig. 9A). The lithologically homogeneous, angular clasts suggest that the breccia was deposited as a rock fall (Fig. 10). It is unclear if the rockfall was subaerial or submarine. The graded nature of the overlying A2.3 conglomerates indicates a transition from a disorganized deposition to a more organized deposition from submarine, high-concentration turbidity or other mass-flow currents (Fig. 9A). The upward increase in clast rounding may reflect increased grain-to-grain contacts during transport or increasing distance from source area, or a combination of both factors. Facies of the Canada Bonita Member Several sedimentary features displayed by rocks of the Canada Bonita Member rocks support Redmond's (1982) interpretation that these rocks are deep-marine turbidity deposits. These features include graded beds, poor to well developed Bouma sequences (Taba, Tba), bathyal depths based on benthic foraminifera (Appendix 1 in de Zoeten, 1988), and lateral and vertical facies relationships. The absence of organization and the very coarse nature of A1.1 and A2.1 conglomerate horizons indicates rapid deposition from high-concentration turbidity currents or debris flows; AI.1 facies suggests rapid sedimentation by frictional freezing processes, and A2.1 facies implies deposition from traction bedload processes (Picketing et al., 1986). Smaller, convex-down conglomerate beds (1-7 m thick; A I.1, A2.1) suggest that the coarsest-grained material moved as channelized flows. Internal stratal features of thick-bedded sandstone (B2.1, C2.1 facies) suggest deposition from high-concentration turbidity currents. The B2.1 sandstone facies indicate more rapid deposition than the C2.1 sandstones (Pickering et al., 1986). Deposition of the repetitive, thin- to medium-bedded sandstones and siltstones resulted from low-concentration turbidity currents. The predominance of parallel-laminated and the structureless lower division of these beds suggests rapid deposition from high-velocity flows. Three distinct facies assemblages are recognized in the Canada Bonita Member of the Altamira Formation. The first is characterized by conglomerates (AI.1, A2.1) interbedded with alternating thin- to medium-bedded sandstone and siltstone (C2.2, C2.3, D2.2) found in the Northern Canada Bonita, Southern Canada Bonita, and Calabaza sections (Fig. 12). Facies D2.2, C2.3, and C2.2 appear to be organized into a few poorly developed, thinning- and
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA fining-upward sequences, which are about 100 m thick. Individual conglomeratic units average about 5 m in thickness and exhibit greater internal organization and fining-upwards trends. We infer that the thin-bedded sandstone and siltstone turbidite couplets represent channel/overbank or interchannel deposits (Mutti and Normark, 1987). In this interpretation, the conglomerate-filled channels serve as conduits for sediment transport from the slope. The coarse nature of the overbank sediment and channel deposits requires either very competent flows or a nearby source. The Rio Perez and the Llanos syncline sections (Fig. 11) illustrate the second major facies assemblage in the Canada Bonita Member, which consists of coarsening- and thickening-up packages of tabular-shaped, thick-bedded sand (B2.1, C2.1). This facies assemblage is interpreted as representing prograding depositional lobes in a submarine fan system (Mutti and Normark, 1987). Channel-fill conglomerates capping the coarsening- and thickening-up packages suggest that lobes were prograding (Shanmugan and Moiola, 1988). Fan lobe progradation is further supported by the reversal in vertical cycles at the top of the Rio Perez section, which suggests a transition from lobe facies assemblage to more proximal channel/overbank deposits (Fig. 11). The final facies assemblage consists chiefly of thin- to medium-bedded, sandstone and siltstone couplets (C2.2, C2.3, D2.2) which characterize the E1 Mamey and Guananico sections. Lithologically and sedimentologically these deposits resemble the overbank deposits described in the first facies assemblage, although the absence of conglomerates or thick-bedded sandstone beds argues against a direct correlation. These rocks are more likely basinal deposits at the distal regions of the fan or distal interchannel deposits. Sedimentary rocks of the Las Lavas Formation Outcrop distribution and general stratigraphy of the Las Lavas Formation The Las Lavas Formation crops out over 800 km 2 in the Cordillera Septentrional in a belt extending from Monte Cristi in the west to northeast of Santiago (Figs. 3, 4C). This study focused on the best-exposed sections along the south-central flank of the Cordillera Septentrional. Approximately 450 m of section is exposed in isolated outcrops along the Santiago-Puerto Plata highway, and more than 1600 m is exposed in the type section along the Arroyo Las Lavas (UTM 087645) (Fig. 4A). The Las Lavas Formation consists of lithic and carbonate conglomerate, lithic-rich calcarenite and thin- to medium-bedded sandstone and siltstone couplets. The Las Lavas Formation is divided into two
267
members: the E1 Limon and the overlying La Pocilguita Member (Fig. 7). The E1 Limon Member forms a resistant hogback which strikes 110~ for approximately 6 km. The ridge-forming E1 Limon Member is offset by north-south striking faults (Fig. 4C). The type section is found in the Arroyo Las Lavas near the village of E1 Limon (UTM 094689), 7 km to the northeast of Navarette (Fig. 4A). The E1 Limon Member disconformably overlies the Altamira Formation near the village of E1 Limon (UTM 097688) (Fig. 4A). The E1 Limon Member is approximately 260 m thick and consists of a basal conglomerate (A1.1, A2.1 facies comprising 40% of the members), thin- to medium-bedded couplets (C2.2, C2.3, D facies comprising 30% of the members), capped by interbedded lithic-rich calcarenites (AI.1, A2.3, C2.1, C2.2 facies comprising 30% of the members). The La Pocilguita Member conformably overlies the E1 Limon Member and is named after the village La Pocilguita del Limon, 1 km south of E1 Limon (UTM 088678) (Fig. 4A). The La Pocilguita Member is 1300 m thick with more than half the section consisting of alternating thin- to medium-bedded siliciclastic sandstone and siltstone (C2.2, C2.3, D facies) and the remainder consisting of lithic-rich calcarenite (AI.1, A2.3, A2.7, C2.1, C2.2 facies) and minor lithic conglomerate (A1.1). West-northwest striking shear zones increase in number southward in the Las Lavas Formation toward the Septentrional fault zone (de Zoeten and Mann, 1991). The regional effect on these shear zones on the stratigraphy of the Las Lavas Formation appears minor, because biostratigraphic dating is consistent with southward younging in the southdipping section (Fig. 4B). Age and paleobathymetry of the Las Lavas Formation Biostratigraphic analysis of four samples indicate that deposition of the E1 Limon Member occurred during Late Oligocene time (Appendix 1, in de Zoeten, 1988). Fourteen samples from the La Pocilguita Member indicate a Late Oligocene to Early Miocene age (Appendix 1, in de Zoeten, 1988). Very few depth-distinctive benthic foraminifera were recognized. One sample suggests deposition at bathyal depths. Measured sections of the Las Lavas Formation Three sections were measured from the Las Lavas Formation. From west to east, these sections include the Navarette, Las Lavas, and Rio Jacagua (Fig. 15). The Navarette section was measured along the Santiago-Puerto Plata highway south of the Southern Canada Bonita section (Fig. 11). The Las Lavas measured section is the type section for the Las Lavas Formation (Fig. 15A). The Las Lavas sec-
268
R. DE ZOETEN and E MANN
A:Z ~9
o . O,...,
o~ r.gZ
g~
Z
o~
,..z.
z~0~ r
,.q
o
..~
~
o
9
,.tzZ
,x:Z
9 ,.,,~
O
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA
269
Fig. 16. Measured section of the Lower Oligocene to Lower Miocene La Toca Formation. Inset map shows location of the measured section near Pedro Garcfa and location of the three along-strike offsets in this composite section (A, B, C). Also shown is the location of the Rfo Jacagua measured section that appears in Fig. 15C.
tion overlies the Calabaza section measured of the Altamira Formation in the Arroyo Las Lavas (Fig. 11, inset map). The Rfo Jacagua section is bounded on the north by the Rio Grande fault zone and on the south by the Septentrional fault zone (Fig. 16, inset map).
Las Lavas measured section: type locality of the Las Lavas Formation El Limon Member of the Las Lavas Formation The basal conglomerate of the E1 Limon M e m b e r exposed in Arroyo Las Lavas thickens westwardly
270 from 50 to at least 150 m before being truncated by the north-south striking Altamira fault zone (Fig. 4C). The contact between the basal conglomerate and the underlying Altamira Formation is inferred to be an erosional disconformity based on lateral alongstrike variations in thickness and distinct lithologic differences across the conglomerate. The conglomerate is massive to thick-bedded (Fig. 15A). Conglomerate beds of the E1 Limon Member are disorganized, clast-supported (AI.1), with some A2.3 beds showing marked scouting and poor grading, and some A2.1 beds showing tangential cross-stratification. Conglomerate clasts in the E1 Limon Member range in size from pebbles to cobbles. Clasts are equidimensional or oblate, subrounded to well rounded, and have a poorly defined a-axis orientation. Near E1 Limon (UTM 094689), clasts are composed predominantly of recrystallized limestone (>60%) (which include gray, black, green, and banded argillite derived from the Los Hidalgos Formation) and minor amounts of volcanic (15%), plutonic (5%), sandstone (5%), and packstone (5%) clasts. Some conglomerate beds contain higher concentrations of igneous clasts (e.g., andesite 32%, tonalite 12%). The basal conglomerate of the E1 Limon Member is overlain by about 80 m of monotonous thinto medium-bedded sandstones and siltstones (C2.2, C2.3, D2.2) (Fig. 15A). The beds are tabular and laterally continuous. The basal division of beds are massive or poorly graded, and the upper parts are mud-rich siltstones exhibiting parallel laminae and more rare bioturbation by Planolites and Skolithos. The thin-bedded facies of the E1 Limon Member is overlain by about 60 m of medium-bedded to massive lithic-rich carbonates. Lithic-rich carbonate conglomerate is clast-supported and exhibits both disorganized (AI.1) and organized (A2.3) internal character (Fig. 15A). Clasts are predominantly lithic to carbonate clasts. The texture and lithology of the lithic component resembles that of the basal conglomerate. The lithic fragments decrease upward in the graded beds. Larger carbonate clasts (760 cm) are irregularly shaped and are composed of coral heads, branching corals, and fossiliferous wackestone and mdstone, while the smaller grains are mainly red algae, larger reefal foraminifera, and coral fragments. Calciturbidites and calcarenite interbedded with the carbonate conglomerates are medium- to very thick-bedded, and contain abundant (5-40%) sandto pebble-sized lithic fragments. Medium- to thick-bedded, clastic limestone is tabular, laterally continuous, and has a sharp basal contact. Clastic limestone is commonly massive, exhibits coarse-tail grading and lacks sole marks. Calciturbidites and calcarenite are predominantly composed of biogenic material, including red algae, larger
R. DE ZOETEN and R MANN reefal foraminifera, coral and shell fragments with lesser lithic fragments. A few foraminiferal rudstones, found as outsized boulders (1-5 m), are composed entirely of large shallow-marine Eulepadina foraminifera which are diagnostic of Oligocene reef, forereef, or shelf environments (Appendix 1 in de Zoeten, 1988). La Pocilguita Member of the Las Lavas Formation Very thin- to medium-bedded siliciclastic facies (C2.2, C2.3, D2.2) conformably overlie the E1 Limon Member and make up more than half of the La Pocilguita Member (Fig. 15). Bed geometry and internal organization of these facies are identical to those previously described in the Altamira Formation. Thin-bedded facies are interbedded with individual and stacked lithic-rich clastic limestone beds, which increase toward the south. Calcarenites range in grain size from silt to small cobbles, and are organized into conglomerate (A2.3, A2.7) and sandstone and siltstone couplets (C2.1, C2.2, B2.1). Thinner-bedded (0.5-2 m) conglomeratic facies (A2.3, A2.5, A2.7) define 1-4-m-thick finingupward cycles, which begin at 420 m above the base of the section. The basal contact of conglomerate beds are erosional, but on the scale of the outcrop appear to be tabular and laterally continuous. Carbonate-rich, lithic pebble conglomerate grades up into structureless coralline gravel beds, parallel-laminated calcarenite, silty calcilutite, and back into thin-bedded siliciclastic rocks. The thicker and coarser conglomerate is mostly organized into A I.1 and A2.1 facies and consist of very few carbonate clasts. Gravel- to silt-sized calcarenite in the La Pocilguita Member commonly shows well developed Bouma sequences (Tabd, Tabcd). Beds are tabular and laterally continuous. Calcarenite beds have sharp stratal boundaries and display few sole marks. Lithic fragments are typically concentrated in the lower portion of the bed (Ta) and decrease in abundance upward. Basal sands of the Ta interval are coarse-tail graded or structureless, and grade into parallel laminae, tipple marks, and laminated calcilutite. Compositionally, the calcarenite is identical to that found in the E1 Limon Member. Calcarenite is composed mainly of red algae, larger reefal foraminifera, coral fragments, and fossiliferous packstone. It is difficult to discern any large-scale vertical organization in the Las Lavas Formation (Fig. 15). However, on a smaller scale, calcarenite beds increase in frequency upsection and are organized into small (1-5 m thick) thinning- and fining-up cycles with conglomeratic bases. The only matrixsupported, disorganized conglomerate (A1.4) recognized in the Altamira or Las Lavas formations is
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA present near the top of the Las Lavas measured section (Fig. 15A). Conglomerate clasts are composed of cobble-sized coral heads with a minor amount of lithic pebbles supported in a calcareous shale matrix. Paleocurrent indicators were measured on a few medium- and thick-bedded calcarenite and siliciclastic strata. Measurements on 24 bimodal current features, including groove casts and a-axis clast alignment of conglomerate indicate bidirectional, northwest-southeast (~ 130 ~ trending flow. Twentythree unimodal paleocurrent structures indicate that currents flowed predominantly to the northwest. Navarette measured section of the Las Lavas Formation El Limon Member of the Las Lavas Formation Along the Santiago-Puerto Plata highway, the basal conglomerate of the E1 Limon Member is approximately 50 m thick (Fig. 15). The conglomerate is massive to poorly bedded, clast-supported, and disorganized to organized (A2.1, AI.1) (Fig. 13B). Poor stratification is defined by oblate clasts preferentially oriented parallel to bedding (Fig. 14B). The basal contact with the underlying Altamira Formation is faulted south of Canada Bonita (UTM 071692). The lateral geometry of the conglomerate horizon is difficult to discern in the field (e.g. Fig. 14B). The conglomerate body appears to pinch out over a lateral distance of 10 m, and is replaced by pebble stringers, which are found interbedded with very thin- to thin-bedded sandstone-siltstone couplets. Conglomerate clasts have similar characteristics to those described in the Las Lavas section. Approximately 40 m of facies D siltstone and shale and minor amounts of C2.3 sandstone and siltstone overlie the basal conglomerate of the E1 Limon Member (Fig. 15). These siltstones and shales are pervasively sheared, which results in a distinctive blocky weathering pattern. There does not appear to be any vertical organization in this thin-bedded sequence. Along the Santiago-Puerto Plata highway, the relationship between the calcarenite and siliciclastic units is less clear. Carbonate conglomerate, foraminiferal rudstone and lithic-rich calcarenite (20 m thick) are found in the Arroyo Guanabano (UTM 073687), east of the highway (Fig. 4A). West of the highway (UTM 068687), limestones of the Late Miocene-Pliocene Villa Trina Formation caps the surrounding hilltops and occur as talus blocks at the level of the road. La Pocilguita Member of the Las Lavas Formation In the Navarette section (Fig. 15B) the following four units comprise the La Pocilguita Member over-
271
lying calcarenite of the E1 Limon Member: (1) thinto medium-bedded sandstone and siltstone (C2.2, C2.3, D2.2 comprising 75% of the La Pocilguita Member); (2) interbedded medium to very thick calcareous sandstone beds (C2.1 comprising 10% of the member); (3) 5-15-m-thick calcarenite packages comprising 10% of the member; and (4) pebblecobble lithic conglomerate comprising 5% of the member. The thin-bedded rocks (C2.2, C2.3, D2.2) have sharp basal contacts and are overlain by parallel-laminated to massive, medium- to coarse-grained sand. The thick sand beds (C2.1) are laterally continuous, although some are lenticular-shaped. The thick sandstone is composed of poorly graded to massive, coarse-grained carbonate-rich sand (Tabd, Tdb). Thick (5-15 m) packages of lithic-rich calcarenite and pebble conglomerate are found 240 m and 430 m above the base of the section (Fig. 16). At 240 m above the base, a 15-m-thick clastic limestone unit is in fault contact with underlying thin-bedded siliciclastic rocks. Stratal boundaries are poorly developed in the calcarenite, which is predominantly massive and contains few lithic fragments. Carbonate material consists predominantly of red algae, larger reefal foraminifera, coral fragments and carbonate rock fragments. The 20-m-thick conglomerate exposed 310 m above the base of the section is poorly stratified (A2.1 facies). Clasts are composed of similar lithologies as those of the basal conglomerate of this section. Secondary gypsum fills some fractures in the clasts and in the surrounding thin-bedded siliciclastic rocks. Vertical organization is not apparent in this section, although exposures are limited. Poorly defined a-axis clast orientations in the basal E1 Limon conglomerate suggests varied north-south to northwestsoutheast paleocurrent trends. Rio Jacagua measured section of the Las Lavas Formation
Beds of the Las Lavas Formation (Fig. 15) could not be directly correlated between the Arroyo Las Lavas and the Rio Jacagua sections (location maps in insets of Figs. 11 and 16). The thick lithic conglomerate that characterized the base of the Las Lavas and Navarette sections was not observed in the Rfo Jacagua section. For this reason, the medium-bedded to massive, lithic-rich calcarenite and carbonate conglomerate was inferred to form the base of the Las Lavas Formation in the Rio Jacagua section. These clastic limestones are about 15 m thick and exhibit similar characteristics to those in the E1 Limon Member in the Lavas and Navarette sections. The Rio Jacagua section is dominated by small-scale (110 m) thinning- and fining-up cycles composed of
272 carbonate-rich conglomerates (AI.1, A2.3), calciturbidites (C2.1, C2.2, C2.3) and siliciclastic turbidites (C2.1, C2.2, C2.3, D2.2) (Fig. 15C). Over 70 bidirectional paleocurrent structures indicate a uniform northwest-southeast (130 ~ mean trend to paleoflow direction. The five unidirectional indicators suggest a paleoflow towards the southeast. Facies analysis of the Las Lavas Formation
Vertical facies assemblages and lateral facies relationships in the Las Lavas Formation are not clearly developed and appear to be more heterogeneous than those described above from the Altamira Formation. The stratigraphy and lithology in the E1 Limon Member indicate rapid deposition of coarse terrigenous and intrabasinal clastic sediment from highconcentration turbidity currents and debris flows. The facies of the La Pocilguita Member are similar to those found in the Altamira Formation. The two formations differ mainly in their lithologic composition, with the Las Lavas Formation containing abundant calciturbidites and the Altamira Formation containing siliciclastic turbidites. The influx of intrabasinal, carbonate biogenic material in the Las Lavas Formation suggests an actively prograding carbonate platform at the basin edge. Deposition by deep-marine turbidity currents in the Las Lavas Formation is supported by the occurrence of graded bedding, good lateral continuity of beds in larger outcrops, development of partial to complete Bouma sequences, vertical facies cycles, and diverse facies associations. Calciturbidites of the Las Lavas Formation were deposited by gravity-driven processes similar to those which deposited the siliciclastic turbidites because the two rock types share the same sedimentary structures.
R. DE ZOETEN and E MANN because tectonic uplift has been less on the La Toca block (Fig. 6A).
STRATIGRAPHY OF THE LA TOCA BLOCK Definition of the La Toca block
The basement rocks and sedimentary cover of the La Toca block is abruptly separated from the basement rocks and sedimentary cover of the Altamira block by the Rio Grande fault zone (Fig. 4C). The La Toca block is bounded on the north by the Camt~ fault zone, on the southwest by the Rio Grande fault zone, and on the south by the Septentrional fault zone. The basement rocks and sedimentary cover of the La Toca block extend to the Rio San Juan area of the eastern Cordillera Septentrional (Draper and Nagle, 1991). The basement of the La Toca block consists of two types of Upper Cretaceous to lower Tertiary rocks: (1) volcanic rocks of the Pedro Garcfa Formation in the central Cordillera Septentrional (Eberle et al., 1982; Peralta-Villar, 1985); and (2) a heterogeneous assemblage of igneous and metamorphic rocks of the Rio San Juan complex in the eastern Cordillera Septentrional (Draper and Nagle, 1991). The basal section of the La Toca Formation consists of interbedded sandstone, siltstone and conglomerate and is faulted against both basement assemblages. The contact between the basement rocks and sedimentary cover of the La Toca Formation is inferred to be a fault-modified unconformity, because the compositions of the Upper Eocene to middle Oligocene basal conglomerates directly reflect the lithologies of the underlying Upper Cretaceous to lower Tertiary basement rocks. Basement complex of the La Toca block
Carbonate rocks of the Villa Trina Formation
Siliciclastic rocks of the Altamira block are capped by a widespread, little studied, --~250-m-thick Upper Miocene to Lower Pliocene shallow-water carbonate unit, the Villa Trina Formation (Fig. 6). The Villa Trina Formation is composed of a lower unit of medium-bedded to massive, marly limestones with few isolated deposits of reefal material. The Villa Trina Formation is capped by well-indurated, medium- to thick-bedded wackestones and packstones, which are interbedded with massive reefal deposits, and reefal talus deposits. This upper unit exhibits a karst topography. Similar lithologies of the Villa Trina Formation cap the La Toca block to the east of the Rio Grande fault zone (Fig. 7). The Villa Trina Formation is more extensive on the La Toca block than the Altamira block probably
Pedro Garcia Formation In the central Cordillera Septentrional, basement rocks of the La Toca block consist of the Pedro Garcfa Formation, which is exposed in a 45 km 2 inlier west of the village of Pedro Garcfa (UTM 266670) (Fig. 4A). The fault-bounded igneous rocks are composed mostly of volcanic rocks (aphanitic andesite, amygdaloidal andesite, tuff, and pyroclastic) with a minor amount of intrusive rocks (tonalite, basaltic dikes, and porphyritic volcanic rocks) (Eberle et al., 1982; Peralta-Villar, 1985). A single K-Ar radiometric date indicates that this igneous complex is at least 72 4- 6 Ma (Bowin and Nagle, 1982). Rio San Juan complex The modal distribution of sandstone grains studied in thin-section and the composition of clasts
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA within conglomerates of the La Toca block exposed in the central Cordillera Septentrional suggests a metamorphic and igneous source area. Likely siliciclastic source areas include the Rfo San Juan complex to the east (Draper and Nagle, 1991) and/or the Puerto Plata basement complex to the north (Pindell and Draper, 1991). Both complexes consist of serpentinite, gabbro, and mafic and felsic schist. Based on similar lithologic assemblages and some similar rock types, Draper and Nagle (1991) believe that the Rfo San Juan and the Puerto Plata basement complexes formed in the Upper Cretaceous to lower Tertiary forearc-trench environment and were offset by the Camfi fault zone in Neogene time. Because of their lithologic similarities and the possibility of lateral fault offset, Mann et al. (1991) included both basement complexes as part of the Rfo San JuanPuerto Plata-Pedro Garcfa disrupted terrane.
Stratigraphy of the La Toca Formation Outcrop distribution and general stratigraphy of the La Toca Formation The La Toca Formation crops out over an area of 100 km 2 in the central Cordillera Septentrional. The La Toca Formation is bounded on the south and west by the Rfo Grande fault zone, and on the north by the Cam6 fault zone (Fig. 4C). The base of the La Toca Formation is faulted against the Upper Cretaceous igneous rocks of the Pedro Garcfa Formation (UTM 233663) (Fig. 4C). The La Toca Formation consists of an approximately 300-m-thick basal conglomerate (AI.1, A1.4, A2.3, A2.5 facies), which crops out as hogback ridges, up to 4 km in length (Fig. 16). Sedimentary rocks overlying the basal conglomerate (and possibly laterally equivalent to the conglomerates) are about 500-m-thick and composed mostly of alternating, very thin- to medium-bedded, sandstone and shale couplets (C2.2, C2.3, D 1.2, D2.2) (Fig. 16). These couplets are capped by a 300-m-thick package of thick- to very thick-bedded, sandstones (B2.1, C2.1). These thick-bedded sandstones are exposed in a faulted synclinal ridge (La Cumbre Ridge), striking northwest, which separates north- and south-flowing streams in the central Cordillera Septentrional (Fig. 4A). High-angle faulting north of the Rio Grande fault zone has complicated the stratigraphy of the La Toca Formation (Fig. 5B). To the north of the Rio Grande fault zone, strata generally strike north- to northwesterly and dip to the east. In this area, sedimentary rocks of the La Toca Formation are deformed by high-angle faults. Of the three siliciclastic formations of the E1 Mamey Group described in this chapter, the La Toca Formation received the least amount of study, and the following descriptions should be considered as a preliminary report.
273
Age and paleobathymetry of the La Toca Formation Thirty-six biostratigraphic analyses were performed on twenty samples collected from the La Toca Formation (Appendix 1 in de Zoeten, 1988) (Fig. 16). No microfossils from the basal conglomerate were identified. The overlying sediments range in age from Early Oligocene to early-Middle Miocene. Sedimentary rocks of the La Toca Formation could be as old as Late Eocene, if they correlate with compositionally similar rocks described by Draper and Nagle (1991) that overlie the Rfo San Juan complex. Studies of benthic foraminifera in the La Toca Formation indicate deposition in middle to upper bathyal water depths (150-1500 m; Appendix 1 in de Zoeten, 1988). Measured sections of the La Toca Formation One composite section was measured for the La Toca Formation. Because of structural complexities, this section consists of three parts (Fig. 16, inset map). The base of the section (part A) lies northeast of the village of Altamira, where the Rfo Grande cuts through a steeply dipping basal conglomerate ridge (Fig. 4A). The majority of the section (parts B and C) was measured to the east in outcrops along the Rfo Yaroa (Fig. 4A). The base of the La Toca section consists mostly of a 300-m-thick package of amalgamated conglomeratic facies (A 1.4, A1. l, A2.3, A2.7) with minor interbedded sandstones (B2.1; Fig. 16). Disorganized, matrix-supported facies (A1.4) and a minor clastsupported facies (A1.1) comprise the lower 200 m of the conglomerate. Most conglomeratic beds are tabular, medium-bedded to massive, with planar basal contacts which rarely drape over clasts protruding from the underlying bed. Internal organization of the conglomerate appears to increase upward in the section. Locally, upsection, inversely graded facies (A2.3, A2.7) are interbedded with clast-supported, parallel-stratified conglomerates. Conglomerate hogbacks, extending for 5 km along strike, are separated by laterally equivalent, alternating, thin- to medium-bedded sandstones and shales (C2.2, C2.3, D2.2, D 1.2). Clasts in the base of the La Toca Formation range from granule to boulder in size. They are equidimensional to oblate in shape, and are subangular to rounded. The composition of clasts directly reflects the heterogeneous lithologies in the underlying Pedro Garcfa Formation (Eberle et al., 1982). Clasts in outcrops near the village of Pedro Garcfa (Fig. 16, inset) are composed of approximately 70% volcanic rocks, including tuff, andesite and amygdaloidal lava, 20% tonalite, and 10% sandstone, argillite (recrystallized limestone), vein quartz, serpentinite, and coralline rudstone. This indicates that the Pedro Gar-
274 cfa Formation was the major sediment source of the La Toca Formation. Above the basal conglomerate, the measured section continues in the east up the R/o Yaroa (Fig. 16). 350 m above the base of the section is a sedimentary package, 500 m thick, dominated by very thin- to medium-bedded sandstones and shales (D1.2, D2.2, C2.3 facies). Beds are tabular, laterally continuous, and exhibit sharp basal contacts. Basal sands are relatively coarse-grained, graded or structureless, and have mud-rich, massive, and bioturbated upper divisions (Tabd, Tbd). A few thin- to medium-bedded, lens-shaped calcareous sandstone beds are interbedded with the sandstone and shale couplets. Higher in the section lithic-rich calciturbidites and lithic conglomerate are interspersed between thin-bedded siliciclastic couplets (Fig. 16). Calciturbidites are organized into C2.1 and C2.2 facies (Tabd, Tbcd), but are less abundant than in the Las Lavas Formation (e.g. Fig. 15). Together with the conglomerate (1-2-m-thick AI.1 facies), the calciturbidites form small (1-5 m thick) coarsening- and thickening-up cycles. Thick-bedded sandstones (B2.1, C2.1, C2.2) and lesser amounts of conglomerate beds (AI.1, A2.7) form the upper 300 m of the La Toca Formation (Fig. 16). Sandstone beds have high sand to silt ratios ( 2 : 1 - 1 0 : 1), are commonly tabular, laterally continuous, and thick- to very thick-bedded (ranging from 0.5 to 3.5 m thick). Basal sands range from gravel to medium sand in grain size. Lower division of beds are coarse-tail graded, or massive, and grade up into parallel-laminated sands, rarely containing tipple marks (Tabd, Tbd). Load casts are the dominant type of sole marks. The upper division in many beds contain concentrations of lignite and amber fragments that define parallel laminae in the siltstones. The finer fraction of beds is mud-poor and parallel-laminated or massive. Locally interbedded with the sandstone beds are 1-10-m-thick pebble to small cobble conglomerate beds (AI.1, A2.7 facies) (Fig. 16). Clast-supported, disorganized conglomerates are dominant. Clasts are composed of a wide assortment of lithologies, which include volcanic, plutonic, metamorphic, and sedimentary rock fragments. Vein quartz and serpentinite fragments are present. In the 800-m-thick section above the basal conglomerate, increasingly coarse sand and conglomerate beds define a thickening- and coarsening-up cycle (>400 m) that is capped by almost 300 m of thick-bedded sandstones. Thin (1-5 m) isolated sandstone packets also reflect thickening-upward trends. The thick-bedded sandstone at the top of the section shows no large-scale thickening- or thinningupward trend, but locally there are both thinningand thickening-up cycles (Fig. 16).
R. DE ZOETEN and E MANN Few paleocurrent indicators were found in the La Toca Formation. Nine measurements of groove marks indicate a northwest-southeast mean paleocurrent trend. Another twelve paleocurrent structures were measured between anastomosing high-angle faults in the Rfo Grande fault zone and therefore are possibly subject to tectonic rotation. They suggest both northwest-southeast and northeast-southwest mean paleocurrent trend (Fig. 16). Facies analysis of the La Toca F o r m a t i o n
The internal organization of the basal conglomerate of the La Toca Formation is distinct from conglomerate described in both the Altamira and Las Lavas formations. Matrix-supported, basal conglomerate of the La Toca Formation indicates rapid deposition from cohesive debris flows (Middleton and Hampton, 1976). Whether deposition occurred from broad, unconfined sheets or from channelized flows is not clear. Very thin- to medium-bedded sandstones and siltstones (C2.2, C2.3, D2.2, D2.1) in the lower part of the La Toca Formation suggest deposition from lowconcentration turbidity currents and from high-concentration silt-dominated turbidity currents (Fig. 16). Coarsening and thickening of sandstone upsection suggests progradation of a marine depositional system. Thick-bedded sandstone capping the La Toca Formation section was deposited by high-concentration turbidity currents. Because vertical organization is lacking in these thick, tabular sandstone beds, it is unclear if they were deposited as lobes, channel-lobe transitional facies, or as delta frontal sands.
PALEOCURRENTS FROM THE EL MAMEY GROUP
More than 280 bidirectional and unidirectional paleocurrent indicators were recorded from the Upper Eocene to Lower Miocene siliciclastic rocks in the E1 Mamey Group (Fig. 17). For completeness another 90 measurements from work by Dolan et al. (1991) were added to the measurements collected during this study. Measurements taken from structurally tilted beds have been rotated around horizontal axes. Sole marks are generally scarce in the study area, but occur mostly in facies C2.1 and C2.2. (Fig. 8). Unimodal measurements are based on flute casts and ripple marks. Groove casts, channel axes and a-axis clast orientation provided bimodal current indicators. Rose plots of bidirectional measurements shown in Fig. 17 indicate a uniform northwest-southeast trend in paleocurrents continued from the Late Eocene through the Early Miocene. Unimodal paleocurrent indicators, however, suggest that the paleoslope reversed during latest Oligocene time. Uni-
C E N O Z O I C EL M A M E Y G R O U P OF N O R T H E R N H I S P A N I O L A
275
Fig. 17. Paleocurrent data from siliciclastic marine rocks from the Upper Eocene to Lower Miocene Altamira, Las Lavas, and La Toca formations in the central and western Cordillera Septentrional (modified from Dolan et al., 1991). Rose diagrams differentiate between unidirectional (shown in white) and bidirectional (shown in black) paleocurrent indicators. Unidirectional paleocurrent indicators from the Lower Miocene Las Lavas Formation near Monte Cristi and the Lower Miocene La Toca Formation from near Moca show a mean northwesterly paleoflow.
modal indicators found in the Altamira Formation (Upper Eocene to Upper Oligocene) show paleoflow towards the southeast (Fig. 17). Paleocurrents measured in the Las Lavas and La Toca formations (middle Oligocene to Lower Miocene) suggest paleoflow to the northwest (Fig. 17). A significant number (8%) of paleocurrent indicators are oriented at high angle (N-S) to the predominant current orientation, and suggest the possibility of lateral (north or south) sediment source areas. Alternatively, anomalous north-south oriented paleoflow structures may simply have formed from interchannel deposits. Dolan et al. (1991) compiled paleocurrent data from four Late Cretaceous to Miocene basins in both Hispaniola and Puerto Rico and included some of these data from the E1 Mamey Group. Dolan et al. (1991) showed that all four basins are characterized by elongate shapes with most paleocurrents oriented in a basin-parallel orientation. The elongate basins in Puerto Rico appear to have formed as intra-arc basins which accompanied volcanic activity in the arc. The elongate basins of Hispaniola have largely post-dated magmatic activity of the Hispaniola arc.
SANDSTONE PETROGRAPHY OF THE EL MAMEY GROUP
Methodology Thirty-six medium- to coarse-grained sandstone samples were selected for point-counting to determine the petrographic mode (Dickinson, 1970; Graham et al., 1976). Samples are representative of the 5-km-thick siliciclastic section of the Altamira, Las Lavas, La Toca formations, and range from Upper Eocene to Lower Miocene (location of petrographic samples shown on all measured sections). Samples were collected over a 500 k m 2 geographic area. The Gazzi-Dickinson point-count method described by Ingersoll et al. (1984) was followed. Using this method, a point is counted as a rock fragment if the cross-hairs fall on the aphanitic part of a rock fragment. A point is counted as a mineral grain if the cross-hair falls on a phenocryst greater than 0.0625 mm. For each thin-section, 300-350 points were identified. Thin-sections were stained with sodium cobaltinitrite to facilitate recognition of orthoclase. Because of the inherent difficulty of recognizing chert from felsic volcanic rock fragments,
276
R. D E Z O E T E N and E M A N N
Table 1 Point-count data based on 300-350 point counts per thin section following the conventions of the Gazzi-Dickinson method (Ingersoll, 1984) No.
Sample no. Age
Quartz
Feldspar Lithicfragments
Qm F
Lt
Qp Lvm Lsm Lv
1 1 2 18 1 0 1 0 0
19 3 22 50 14 14 9 13 46
80 96 76 32 85 86 90 87 54
1 0 1 13 1 0 0 0 0
19 3 22 50 14 14 9 13 46
80 97 77 37 85 86 91 87 54
0 1 1 15 0 0 1 0 0
100 99 85 57 93 82 70 92 100
0 0 14 28 7 18 29 8 0
100 100 85 70 93 82 70 92 100
0 0 0 0 0 0 0 0 0
0 0 15 30 7 18 30 8 0
Las Lavas Formation Eocene? 10 10187 Oligocene 11 6687 Early Oligocene 12 7 Early Oligocene 13 6187 Early Oligocene 14 6087 Early Oligocene-Early Miocene 15 12387A Early Oligocene-Early Miocene 16 5787 Early Miocene 17 6887 Early Miocene 18 1B Early Miocene 19 10987 Early Miocene 20 10887 Early Miocene 21 8387 Early Miocene 22 7387 Oligocene? 23 1787 Oligocene? 24 1887 Oligocene? 25 -521 Oligocene? 26 14287
3 1 1 2 8 2 1 15 18 27 0 0 34 18 27 3 1
33 58 24 20 10 17 10 37 42 25 30 39 56 20 38 23 39
64 41 75 78 82 81 89 48 40 48 70 61 10 62 35 74 60
2 0 1 2 2 1 1 10 16 13 0 0 14 7 13 1 1
33 58 24 20 10 17 9 37 42 25 30 39 56 20 38 23 39
65 2 42 1 75 0 78 1 88 7 82 1 90 1 53 9 42 6 62 24 70 0 61 0 30 64 73 16 49 29 76 3 60 0
97 98 98 98 89 98 99 89 92 75 100 100 4 84 71 86 99
1 1 2 1 4 1 0 2 2 1 0 0 32 0 0 8 1
99 99 98 98 96 99 100 98 99 93 100 100 89 100 100 90 99
0 0 0 0 1 0 0 0 0 5 0 0 0 0 0 0 0
1 1 2 1 3 1 0 2 1 2 0 0 11 0 0 10 1
La Toca Formation 27 9487 Oligocene 28 7887 middle Oligocene 29 8687 middle Oligocene 30 13987 Late Oligocene 31 8887 Late Oligocene 32 13187 Late Oligocene 33 10587 Early Miocene 34 8287 Early Miocene 35 9587 Early Miocene
5 40 33 51 30 41 74 29 28
16 21 46 41 60 39 12 60 54
79 39 21 8 10 20 14 11 18
4 9 26 9 21 20 15 7 18
16 21 46 41 60 38 11 60 54
80 70 28 50 19 42 74 33 28
97 46 70 8 51 50 1 29 56
1 1 5 0 0 0 6 2 9
99 0 70 28 90 1 100 0 100 0 100 0 28 41 96 0 91 1
1 1 9 0 0 0 31 4 8
Altamira Formation Early 1 15387 Early 2 2387 Early 3 1687 Early 4 4908 Early 5 -746 Early 6 -745 Early 7 -744 Early 8 4687 Early 9 15687
Eocene Eocene Oligocene Oligocene Oligocene Oligocene Oligocene Oligocene Oligocene
2 53 25 92 49 50 93 69 35
Lm Ls
Percentages are recalculated to 100%. Data indicate that sands from the La Toca Formation consist of much more quartz than coeval sandstones from the Altamira Formation. The quartz grains from the La Toca Formation are predominantly polycrystalline. Key to abbreviations: Q -- quartz, F = feldspars, L = lithic fragments, Qm = monocrystalline quartz, Lt = total lithic fragments and polycrystalline quartz, Qp = polycrystalline quartz, Lvm -- total volcanic and metavolcanic lithic fragments, Lsm -- total sedimentary and metasedimentary lithic fragments, Lv = total volcanic lithic fragments, Lm = total metamorphic lithic fragments, Ls = total sedimentary lithic fragments.
the former was grouped with the felsic fragments. Recalculating framework modes by assigning the questionable polycrystalline grains to chert showed no significant (<5%) change in their ternary distribution. Recalculated point-count data are shown in Table 1.
Results Altamira Formation Framework modal data indicate that the siliciclastic rocks in the central Cordillera Septentrional range in composition from volcanic arenite to quartz-rich,
lithic arkose (Fig. 18). Sandstone of the Late Eocene to Late Oligocene age Altamira Formation are characterized by: (1) abundant volcanic lithic fragments; (2) feldspars; (3) biogenic fragments; and (4) the absence of quartz (Fig. 19A). Las Lavas Formation The composition of Las Lavas Formation sandstone changed from Late Oligocene to Early Miocene time (Fig. 18A). Quartz and plagioclase grains gradually increase in abundance upsection, whereas the concentration of volcanic lithic fragments remains constant. Similar to the Altamira For-
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA
277
Fig. 18. (A) Ternary plots of quartz, feldspar, and lithic fragments (QFL) which illustrate change in composition with time for the (1) Altamira Formation, (2) La Toca Formation, and (3) Las Lavas Formation. From Late Eocene to Early Miocene the underlying arc and forearc crust of the area is subjected to progressively deeper levels of erosion related mainly to tectonic events. (B) QFL ternary diagram of sandstone framework grains from all three formations. Tectonic provenance fields from Dickinson et al. (1983) are shown. mation, sandstone of the Las Lavas Formation (Upper Oligocene) is volcanic arenite. Lower Miocene sandstone from the Las Lavas Formation, on the other hand, is feldspathic litharenite (Fig. 19B). Detrital serpentine was found in three samples near the top of the Lavas section (two samples of Late Oligocene age; one sample of Early Miocene age). In these three samples, serpentinite fragments make up almost 30-+- 10% of the framework grains in each sample. Several Miocene sandstones from the Las Lavas Formation contain albite-rich plagioclase, which contain parallel-oriented needle-shaped microlites, along with zoisite and epidote inclusions. These types of plagioclase grains are also commonly found in plagioclase grains from the La Toca Formation sandstone.
La Toca Formation Sandstone from the La Toca Formation has a much lower volcanic lithic component (24-1-3%), and a significantly higher quartz concentration (37 :i: 3%), than sandstone from the Altamira and Las Lavas Formations (Figs. 18A and 19C). Polycrystalline quartz is the most abundant and is probably metamorphic in origin. Plutonic quartz makes up most of the inclusion-rich monocrystalline quartz.
Detrital mineral grains such as hornblende, serpentinite, epidote, zoisite, and several types of mica make up a significant proportion (5-30%) of the framework grains.
Comparison of all formations of the El Mamey Group Sandstone from the Altamira, Las Lavas, La Toca formations exhibit several similar characteristics. They contain abundant intrabasinal biogenic material, such as red algae, larger reefal foraminifera, and lesser amounts of coral and shell fragments. The lithic fragments are predominantly volcanic, although, significant quantities of sedimentary rock fragments are recognized in sandstones from the Altamira Formation. Terrestrial organic debris (lignite and amber) is a minor constituent in all the sandstones. The majority of the lignite and amber deposits are concentrated in the thick-bedded sandstone at the top of the La Toca Formation (Redmond, 1982; Iturralde-Vinent and MacPhee, 1996) (Fig. 16). Overall, the sandstone of the E1 Mamey Group are well compacted, contain very little clayor mud-sized fraction (<5%), and are moderately well cemented with sparry calcite and much lesser amounts of laumonite and analcite. All sands have
278
R. DE ZOETEN and R MANN
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA very low porosities (<3%). Chloritization, epidotization, and clay alteration of volcanic rock fragments are likely to be pre-depositional, because alteration either appears to be localized within grain boundaries, or slightly affecting the matrix (Fig. 19A,B). Some plagioclase grains show partial calcite replacement and/or seritization.
Provenance of the Altamira Formation The Altamira Formation sandstones are characterized by very low quartz contents, high plagioclase/orthoclase values, and a volcanic-rich lithic component. These grain parameter values all point to a volcanic provenance (Fig. 18B). The volcanic rock fragments are composed predominantly of lathwork and felsitic (65-97%) grains with a lesser amount of microlitic grains. These grains suggest a diverse volcanic source with a composition that ranges from silicic to basaltic lavas (Dickinson, 1970). Framework grains of sandstones plotted on a QFL diagram fall within the provenance field of a magmatic arc setting as described by Dickinson and Suczek (1979) and Dickinson et al. (1983) (Fig. 18B). Plotting the modal distribution of sandstone from the Las Lavas Formation on a QFL ternary diagram (Fig. 18B) indicates that Miocene sandstone is derived from a transitional magmatic arc setting. The composition of Altamira Formation and Las Lavas Formation sandstone of Late Oligocene age suggests an undissected arc source (Dickinson et al., 1983).
Provenance of the La Toca Formation La Toca Formation sandstone is more heterogeneous and their framework modes plot in the feldspar and quartz realms of the QFL diagram, which correlates to the dissected magmatic arc field (Dickinson et al., 1983) (Fig. 18B). The presence of volcanic lithic fragments and the high plagioclase/orthoclase ratio suggests that the sedimentary source for the La Toca Formation continued to be a volcanic arc. The increase in compositional complexity in the La Toca Formation implies breaching of the shallow levels of the arc with exposure of plutonic-metamorphic rocks (Fig. 18B).
279
DISCUSSION
Three-phase tectonic and sedimentary history of northern Hispaniola Based on sedimentary data presented in this paper, the tectonic and basin evolution of northern Hispaniola can be divided into three phases, each of which is marked by distinct depositional facies (Fig. 20): (1) Paleocene to Middle Eocene phase records the termination of arc activity and uplift of arc basement rocks probably as the result of early interaction between the Hispaniola arc and the Bahama carbonate platform; (2) Late Eocene to Early Miocene phase marks the first major deposition of deep-marine siliciclastic rocks; and (3) Late Miocene to Early Pliocene phase records tectonic uplift of northern Hispaniola to near sea level and subsequent deposition of shallow-marine limestones of the Villa Trina Formation (Fig. 6). Data from this study of the central Cordillera Septentrional has refined the timing and extent of these phases and, therefore, the tectonic history of plate interactions in this part of the North AmericaCaribbean plate boundary zone
Phase 1: Paleocene to Middle Eocene Tectonics Coeval igneous and metamorphic rocks in northern Hispaniola are interpreted to have formed in an intra-oceanic island-arc environment at the leading edge of the Caribbean plate (Bowin, 1975; Nagle, 1979; Draper and Nagle, 1991; Pindell and Draper, 1991; Dolan et al., 1991; Calais and Mercier de Ldpinay, 1995) (Fig. 20A). Isotopic ages for northern Hispaniola suggest that island-arc development was continuous from a mid-Cretaceous orogenic event (Draper et al., 1996) to the Late Eocene or Early Oligocene (Kesler et al., 1991). Near the end of the arc phase, initial opening of the Cayman Trough marks a change from northeast to eastward Caribbean plate motion and the transition from a convergent to the present east-west strike-slip margin (e.g., Sykes et al., 1982; Mann et al., 1995). Several authors believe that the collision of the Caribbean plate with the buoyant Bahama Platform
Fig. 19. (A) Thin-section photomicrograph with crossed-polars of an Upper Eocene to Upper Oligocene Altamira sandstone consisting of felsitic and microlitic volcanic rock fragments, plagioclase, and carbonate rock fragments. The width of the photo is approximately 1.65 mm (sample no. 7-46). (B) Thin-section photomicrograph with crossed-polars of a typical Lower Miocene sandstone of the Las Lavas Formation which consists of polycrystalline quartz, plagioclase, felsite and lathwork type volcanic rock fragments, and bioclasts. The width of the photo is approximately 3.3 mm (sample no. 10887). (C) Thin-section photomicrograph with crossed-polars of a Lower Oligocene to Lower Miocene sandstone from the La Toca Formation. This photomicrograph shows polycrystalline and monocrystalline quartz, plagioclase, volcanic rock fragments, and zoisite. The large, twined plagioclase grain near the center of the photomicrograph contains unidentified, needle-shaped microlites which exhibit a preferred parallel orientation. The width of the photo is about 1.65 mm (sample no. 10587). Color version at http://www.elsevier.nl/locate/caribas/
280
R. DE Z O E T E N and R M A N N
Fig. 20. Block diagrams summarizing three main tectonic phases in the evolution of the North America-Caribbean plate boundary zone in northern Hispaniola. (A) Phase 1 is marked by Paleogene to Early Eocene deposition of hemipelagic, fine-grained turbidites (Los Hidalgos Formation = number 5) which are interbedded with arc-related dikes and sills of intermediate composition. Similar deep-marine sedimentary rocks are found to the north (Imbert Formation = number 4) and to the south (Magua Formation = number 6) of the study area and suggest a regionally extensive basin at least 40 km wide. The substrate of the Imbert Formation is a heterogeneous basement consisting of serpentinite, gabbros, volcanic rocks and blueschists (Puerto Plata basement and Rio San Juan complexes = number 3), whereas the substrate of the Magua Formation is greenschist metamorphic rocks intruded by granodiorite plutons (Duarte complex = number 7). Tuffaceous horizons are common in the Imbert, Los Hidalgos, and Magua formations and suggest an active arc environment probably to the south along the Hispaniola segment of the volcanic arc. We interpret these geologic relationships in terms of a forearc basin developed above a south- to southwest-dipping slab of subducted Atlantic ocean floor (number 2). Large-scale, Middle Eocene folding and uplift will terminate Phase 1 deposition. This compressive event is related to attempted subduction of the Bahama Platform (number 1) beneath the forearc area. (B) Phase 2 is marked by Upper Eocene to Lower Miocene deposition of several kilometers of siliciclastic turbidites (El Mamey Group = number 8; Tabera Group -- number 9) in west-northwesterly striking, elongate basins. Arrows indicate paleoflow directions based on paleocurrent studies in turbiditic rocks (cf. Fig. 17). Source areas for the Tabera Group include Lower Cretaceous metasedimentary rocks (Amina schists = number 10) and volcanic arc rocks exposed to the east. Source area for the E1 Mamey Group include folded, hemipelagic rocks of the Los Hidalgos Formation to the south (number 5), the Puerto Plata basement and
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA terminated subduction and arc-related volcanism and initiated east-west movement (Pindell and Draper, 1991; Dolan et al., 1991; Mann et al., 1995). Calais and Mercier de L6pinay (1995) have correlated the Late Eocene deformation seen in northern Hispaniola with a convergent event of the same age in southern Cuba and northern Haiti.
281
ship was recognized by us in the central Cordillera Septentrional between the tightly folded Upper Paleocene to Lower Eocene Los Hidalgos Formation and less folded overlying Upper Eocene Altamira Formation. In the western Cordillera Septentrional, Calais et al. (1992) recognized this same unconformity separating the Paleocene-Eocene E1 Cacheal tufts from the overlying Lower Miocene series.
Paleocene-Eocene rock record The oldest sedimentary rocks in northern Hispaniola are mainly turbidites and hemipelagic sedimentary rocks of Paleocene to Middle Eocene age. These rocks were deposited in a deep-marine environment above igneous and metamorphic arc basement rocks (Fig. 20A). This deep-marine depositional phase seems to have affected all components of the arc system, including: (1) deposition of the Imbert Formation in the 'outer forearc-trench setting' (Nagle, 1979; Pindell and Draper, 1991); (2) deposition of the Los Hidalgos Formation in the 'inner arc setting'; and (3) deposition of the Magua Formation on the volcanic arc (Palmer, 1979). Unlike the Los Hidalgos Formation, the Imbert and Magua formations contain a significant amount (10-60%) of siliciclastic rocks. Igneous activity during deposition of the Imbert and the Los Hidalgos formations is shown by interbedded calcareous tuffaceous rocks in the Imbert Formation and Palma Picada porphyritic rocks intruded into the Los Hidalgos Formation (Fig. 20A). Although no radiometric ages have been determined for these igneous rocks, pelagic foraminifera from the tuffaceous rocks of the Imbert Formation indicate a Paleocene-Early Eocene age (Nagle, 1979). The Palma Picada igneous rocks intruded the calcareous sediments of the Los Hidalgos Formation either during or after their deposition and must be Late Paleocene to Middle Eocene in age. Pindell and Draper (1991) report that the Irabert Formation is overlain unconformably by Upper Eocene sedimentary rocks of the Luperon Formation. The stratigraphic relationship between shallowmarine, Lower to Middle Eocene limestone of the La Isla Formation and the Imbert Formation is unclear, but the La Isla Formation limestone is believed to post-date Imbert deposition and pre-date Upper Eocene rocks above the unconformity (Pindell and Draper, 1991). A similar, unconformable relation-
Interpretation of phase 1 tectonic and sedimentary events Unconformable contacts between the basement and overlying sedimentary rocks indicate that basement blocks were locally uplifted and exposed by Late Eocene time, and possibly as early as the Late Paleocene. Lithologies of the Imbert and Magua formations indicate that arc-derived sediments were deposited on the edges of the deep-marine forearc basin (Fig. 20A). The composition of the Los Hidalgos Formation indicates that very little terrigenous sediment reached its position in the deep basin. Three reasons may be responsible for the absence of terrigenous sediment in the Los Hidalgos Formation: (1) the basin was effectively isolated from a siliciclastic source, possibly because of an irregular bottom topography; (2) the basin was a great distance from the source of siliciclastic material; or (3) because only a small area of arc rocks were subaerially exposed during this time. The deep-marine origin and present proximity of the Imbert and Los Hidalgos formations suggest that these formations may be laterally equivalent. Lessfolded shallow-marine limestone deposits of the La Isla Formation overlie folded rocks of the Imbert Formation and suggest that the Imbert Formation was uplifted with an outer 'arc-trench' assemblage to near sea level during Late Paleocene to Middle Eocene time (Pindell and Draper, 1991). A similar uplift history for the Los Hidalgos Formation is suggested based on the Middle Eocene angular unconformity separating the folded Los Hidalgos Formation and the less-folded Altamira Formation. The mixture of shallow-marine carbonate clasts and clasts derived from the Los Hidalgos Formation at the top of the basal conglomerate (Ranchete Member) of the Altamira Formation further suggests that the Los Hidalgos Formation had been uplifted to near sea level by Late Eocene time. Thus, the fore-
Rfo San Juan complexes and Pedro Garcfa Formation to the north (number 3). Regional uplift in Middle Eocene time is attributed to the initial attempted subduction of the Bahama Platform (number 1) beneath the Hispaniola arc and Oligocene to Miocene left-lateral, strike-slip faulting along the Rio Grande fault zone (RGFZ) and the Septentrional fault zone (SFZ). (C) Phase 3 is marked by Upper Miocene to Lower Pliocene deposition of shallow-watercarbonate rocks (Villa Trina Formation -- number 11). This limestone appears to have covered most of northern Hispaniola as shown by the wide distribution of its remnants (cf. map in Fig. 6). Late Pliocene to Present uplift of the Cordillera Septentrional along the transpressional Septentrional fault zone has folded the Villa Trina Formation and uplifted it to an elevation of 1250 m. Uplift of the Cordillera Septentrional has accompanied subsidence of coeval rocks in the Cibao basin to depths greater than 3500 m below sea level (Yaque Group = number 12).
282 arc may have uplifted as a single block during this time period. This folding and uplift event coincides with the cessation of most subduction-related processes and probably resulted from the early oblique collision of the Caribbean plate with the Bahama Platform (Pindell and Draper, 1991). Phase 2: Late Eocene to Early Miocene Tectonics Phase 2 covers the depositional period from the Late Eocene to the Early Miocene (that is, the period of deposition of the Altamira, Las Lavas, and La Toca formations; Fig. 7). The tectonic regime during this period has been interpreted by Sykes et al. (1982) and Mann et al. (1995) as transitional: collision with the Bahama Platform was ending, and the Caribbean plate was moving along strike-slip faults in a more easterly direction. In the western Cordillera Septentrional, Calais et al. (1992) provided important structural confirmation of this transition period. They mapped an older set of collision-related folds affecting Paleocene and Eocene rocks equivalent to the Los Hidalgos Formation with folds having north- to east-northeast-trending axial traces and a younger, strike-slip set of folds affecting Miocene rocks that have northwest-trending axial traces and sub-vertical fold axes. Late Eocene to Early Miocene rock record Systematic lateral facies and facies assemblage changes which commonly characterize submarine fans (e.g., Walker, 1984) are not well expressed in the siliciclastic deposits exposed in the central Cordillera Septentrional. The best documentation of lateral facies relationships is recorded in the measured sections from the Altamira Formation (Figs. 9, 11 and 12). The channel (AI.1, A2.1) and overbank deposits (C2.2, C2.3, D2.2) in the Northern Canada Bonita, Southern Canada Bonita, and Calabaza sections resemble middle submarine fan deposits (Mutti and Ricci Lucchi, 1978; Nilsen and Abbot, 1981). The stacked tabular sandstones (B2.1, C2.1) seen in the Rio Perez and Llanos syncline sections are similar to deposits interpreted as outer fan lobes or as large crevasse-splay lobes in the middle fan environment (Mutti and Ricci Lucchi, 1978; Shanmugan and Moiola, 1988). Thin- to medium-bedded sandstone and siltstone facies (C2.2, C2.3, D2.2) seen in the E1 Mamey and Guananico sections may represent basin plain or distal overbank deposits. The Rio Perez, Llanos Syncline, Southern Canada Bonita, Northern Canada Bonita, and Calabaza sections are all laterally equivalent. Their areal distribution in the E1 Mamey Group suggests north to northeast fan progradation from the Northern Canada Bonita to the Rio Perez section. Such progradation
R. DE ZOETEN and E MANN would require paleocurrents flowing from the south and southwest. This prediction, however, conflicts with the mean paleocurrent trend (125 ~ measured from 170 structures, 40 of which indicate that the current flowed to the southeast (Fig. 17). High concentrations of calcareous biogenic material in the turbidite rocks of the E1 Mamey Group indicate shallow-water carbonate production near the shelf margin. Limestone turbidites and debris sheets recognized in slope and basinal settings have commonly been reported in both ancient rocks (Cook and Mullins, 1983) and modem environments (Schlager and Chermak, 1979). Deep-marine fossiliferous limestones, like those from the Las Lavas Formation, commonly emanate from a line source dissected by several small channels and are rarely associated with submarine fans (Cook and Mullins, 1983). Vertical facies relationships show that the Oligocene section exposed along the SantiagoPuerto Plata highway is more conglomeratic than the Upper Eocene section near E1 Mamey. A gradual, upsection transition from lobe to channel/overbank deposition suggests that southeasterly prograding submarine fan systems developed during Late Oligocene time. Vertical relationships in the La Toca Formation point to a prograding submarine fan or delta during Early Miocene time. A major, Late Oligocene depositional event is marked by sudden appearance of limestone conglomerate and calciturbidite in the E1 Limon Member at the base of the Las Lavas Formation (Fig. 15). Intrabasinal, clastic limestone beds increase upward in the section. The section is overlain by platform carbonate of the Villa Trina Formation (Fig. 6). Sandstone petrography from the Altamira, Las Lavas and La Toca formations on the QFL diagram indicates that these sediments came from an arc environment (cf. Dickinson and Suczek, 1979) (Fig. 18B). Modal distribution of framework grains suggest two compositional trends: (1) sandstone from the Altamira and Las Lavas formations have a different sand grain composition from coeval sandstone from the La Toca Formation; and (2) quartz content increases upsection from the Altamira through the Las Lavas Formation. Compositional differences in the sandstones from the La Toca and Altamira formations imply two distinct source areas, with apparently no mixing between the two. This interpretation is further supported by the absence of serpentinite grains in the Altamira Formation sandstones. In contrast, serpentinite clasts are found in the La Toca Formation as well as in the Lower to Middle Eocene sedimentary rocks to the north of the Camfi fault zone. Only the very youngest sandstone (Lower Miocene) in the Las Lavas Formation have similar modal distributions as the sandstone from the La Toca Formation. Superimposed on this trend
CENOZOIC EL MAMEY GROUP OF NORTHERN HISPANIOLA is a time-dependent unroofing sequence from Upper Eocene undissected magmatic arc provenance to a Lower Miocene dissected arc provenance (Fig. 18B). Calais and Mercier de L6pinay (1995) have correlated Upper Eocene to Lower Miocene sedimentary formations in northern Hispaniola (Altamira Formation) with similar rocks in northern Haiti and southern Cuba. They interpret this period as a time of tectonic quiescence when topographic relief generated by the Late Eocene uplift event was eroding to sea level. The Late Eocene uplift and folding event may explain why there is little or no Oligocene preserved in the western part of the Cordillera Septentrional (Calais et el., 1992).
Interpretation of phase 2 tectonic and sedimentary events Sandstone petrography and facies relationships suggest that the La Toca Formation was probably deposited in a separate basin from the basin in which the Altamira and Las Lavas formations were deposited (Fig. 20B). However, it is also possible that these formations may have been deposited in different localities within a larger basin with a wide range of source areas. The La Toca Formation is separated from the coeval Altamira and Las Lavas formations by the 100-400-m-wide, left-lateral Rfo Grande fault zone (de Zoeten and Mann, 1991). The composition of sandstone and conglomerate changes abruptly across the Rfo Grande fault zone. Quartz-rich, La Toca Formation sandstone lie north of the Rfo Grande fault zone, whereas lithic-rich Altamira Formation sandstone occur to the south of the fault. Distinct Upper Cretaceous to lower Tertiary basement complexes underlying sedimentary rocks are likewise separated by the Rfo Grande fault zone. Differences in basement rock lithologies, composition of overlying sandstone, and sedimentary facies profiles suggest distinct origins for the sedimentary rocks now juxtaposed along the Rfo Grande fault zone. Left-lateral oblique-slip motion on the Rfo Grande fault zone appears to have juxtaposed the Altamira and La Toca blocks which Mann et al. (1991) have interpreted as part of much larger terranes (Fig. 20B). Similar relationships, in which blocks with distinctly different geologic histories are in fault contact, are documented for several of the ten other terranes proposed by Mann et al. (1991) in Hispaniola. Movement of blocks along multiple high-angle faults is consistent with large-offsets documented further west in the Cayman Trough (Rosencrantz et al., 1988) (Fig. 1). The time of 'docking' or present juxtaposition of the La Toca and Altamira blocks is not well constrained. The La Toca and Altamira blocks were possibly juxtaposed by Miocene time for two reasons: (1) similar detrital compositions of Lower
283
Miocene sandstones from both the La Toca and Las Laves formations (Fig. 18A); and (2) the widespread cover of Upper Miocene to Lower Pliocene Villa Trine Formation limestones over the siliciclastic formations (Fig. 6). The consistent northwest-southeast trend of paleocurrents measured in Upper Eocene to the Lower Miocene indicates a long-lived, elongate basin as proposed by Dolan et el. (1991) for the E1 Mamey and three other coeval deformed basin complexes in Hispaniola and Puerto Rico. Deposition of the E1Mamey Group rocks within an elongate depositional basin is further supported by: (1) complex lateral facies patterns, which do not form classical deep-sea fan morphologies (Link, 1982); and (2) regional stratigraphy which suggests that the 'outer forearc-trench' assemblage was bathymetrically shallow to the north of the basin from Early Eocene time to the Present (Pindell and Draper, 1991), and that the area of the Los Hidalgos Formation to the south of the basin also formed a bathymetric high during at least Late Eocene time and possibly into the Oligocene (Fig. 20B). The mechanism that produced the complex basins of the E1 Mamey Group is unclear. Although the basin is situated in the forearc region of Hispaniola, siliciclastic deposition began during the waning stages of arc activity in Late Eocene time (Fig. 20B). The basin may have formed by down-warping between the uplifted 'outer forearc-trench' assemblage (Imbert Formation) and the 'inner forearc' (Los Hidalgos Formation)' during the final stages of the collisional event with the Bahama Platform. Folding and uplift of these arc environments would provide igneous and metamorphic source areas that seem to be absent prior to the Middle Eocene folding and uplift event. Calais et el. (1992) show that an older set of folds affecting the Eocene Los Hidalgos equivalent rocks in the area (El Cacheal tufts) of the western Cordillera Septentrional has north to east-northeast axial trends while a younger set of folds with northwest-trending axial traces affecting post-Eocene rocks. They interpreted the older set of folds as related to the Bahama collision and the younger set as related to post-collisional strike-slip movement. Alternatively, the highs flanking the basin may have been formed by strike-slip faults roughly parallel to the axis of the basin (Fig. 20B). Whether the basin formed in a convergent or strike-slip setting is not clear from the sedimentary results of this study, but it is likely that both processes may have affected early basin formation and sedimentation given the transition from arc to strike-slip tectonics at the time of deposition of most of the sediments. Based on regional plate reconstructions and the presence of the Eocene to Recent Cayman Trough west of the area, strike-slip faulting may have played an impor-
284
R. DE ZOETEN and R MANN
tant role in the development of Upper Eocene source areas and basins (Fig. 1).
stones are found throughout the Cordillera Septentrional and Samana Peninsula (Fig. 6).
Eustatic controls on deposition in the El Mamey Group Coarse-grained, channel/overbank deposits in the Altamira Formation correlate with a global sea level lowstand in the Late Oligocene (Haq et al., 1987) and may represent a 'lowstand fan'. The introduction of calciturbidites of the Las Lavas Formation into the basin may indicate that the drop in sea level forced carbonate progradation on to the slope environment. An unstable slope environment would provide a source area for carbonate-rich gravity flows. Another possibility is that rising sea level, following the lowstand, produced a quiescent period inhibiting siliciclastic transport into the basin and increasing carbonate production on the shelf. Carbonate produced on the shelf could be mobilized by localized tectonic activity and moved into the deeper basins. Although Oligocene sedimentation patterns appear to coincide with a Late Oligocene fall in sea level, tectonic influences on sedimentation may have also been important. Upper Oligocene olistostromes containing outsized platform carbonate blocks have been documented in the Tavera Group (Palmer, 1979; Groetsch, 1983) and in the Ocoa Group (Heubeck et al., 1990) south of the study area (Fig. 2).
Interpretation of phase 3 tectonic and sedimentary events De Zoeten and Mann (1991) divided the Neogene uplift history of Hispaniola into two major events: (1) Middle Miocene uplift event, and (2) a post-Early Pliocene uplift event. The Middle Miocene uplift is recorded by a ubiquitous change from deep-marine deposition (Las Lavas and La Toca formations) in Early Miocene time to shallow-marine deposition (Villa Trina Formation) in Late Miocene and Early Pliocene time. Post-Early Pliocene restraining bend tectonics continues to uplift the Cordillera Septentrional to its present position at 1250 m above sea level (Fig. 6).
Phase 3: Late Miocene to Recent Tectonics Miocene strike-slip faulting and related deformation is documented in several areas of northern Hispaniola (de Zoeten and Mann, 1991). Gentle folding of the siliciclastic units of the E1 Mamey Group, which is not recognized in the overlying shallowwater carbonate rocks, indicates Middle Miocene folding, probably associated with transpressional strike-slip faulting. Calais et al. (1992) report the same folding event affecting rocks of the western Cordillera Septentrional. During post-Early Pliocene time, northern Hispaniola underwent transpression associated with the strike-slip 'restraining bend' in the plate boundary zone near Hispaniola (Mann et al., 1991). Obliqueslip movement on the Septentrional fault zone uplifted Upper Miocene to Lower Pliocene limestones to 1200 m above sea level (de Zoeten and Mann, 1991). Late Miocene to Pliocene rock record Few Middle Miocene rocks are found in northern Hispaniola (Figs. 4B, 7). This time period marks a change from Lower Miocene deep-marine sedimentation to Upper Miocene shallow-marine carbonate deposition. Upper Miocene to Lower Pliocene lime-
CONCLUSIONS
Detailed studies of several Paleocene through Miocene sedimentary formations exposed in the central Cordillera Septentrional, Dominican Republic, indicate at least three, distinct tectonic phases in the Cenozoic evolution of the North America-Caribbean plate boundary zone. Each tectonic phase is characterized by deposition of characteristic sedimentary facies and is punctuated by a short-lived folding event. The three phases include: (1) Paleocene to Early Eocene deposition of at least a 250-m-thick section of hemipelagic, finegrained turbidites (Los Hidalgos Formation) interbedded with arc-related dikes and sills of intermediate composition (Palma Picada intrusive rocks). Sedimentation was terminated by a folding and uplift event, which is thought to be related to early attempted subduction of the Bahama Platform beneath the Hispaniola arc (Fig. 20A). (2) Late Eocene to Early Miocene deposition of approximately 4000 m of deep-marine, siliciclastic turbidites (Altamira, Las Lavas, and La Toca formations) into at least two elongate basins subsequently juxtaposed by strike-slip faulting. We interpret sandstone compositions suggesting two distinct source areas during Late Eocene and Oligocene time as evidence that the basins were isolated from one another and later juxtaposed by a 400-m-wide, linear, strike-slip fault (Rfo Grande fault zone). Similarities in sandstone composition indicate that the two basins were juxtaposed in Miocene time. Siliciclastic sedimentation was terminated by a folding and uplift event, which is thought to be associated with transpressional strike-slip faulting related to North America-Caribbean plate motion (Fig. 20B). (3) Late Miocene to Early Pliocene deposition of more than 250 m of shallow-marine limestones
C E N O Z O I C EL M A M E Y G R O U P OF N O R T H E R N H I S P A N I O L A
(Villa Trina Formation). Carbonate sedimentation was terminated by a folding and uplift event related to the current pattern of restraining bend tectonics and active collisional underthrusting of the Bahama Platform. Maximum uplift of the limestone is associated with a large fold in the convex, uplifted side of the restraining bend (Fig. 20C).
ACKNOWLEDGEMENTS
This work formed part of a master's thesis by R. de Zoeten that was supervised by E Mann, E. McBride and M. Cloos at the University of Texas at Austin (de Zoeten, 1988). M. Perez, L. Pena, J. Guzman and G. Draper provided assistance in the field and W. Eberle, E Cepek, S. Monechi, B. Redmond, and E. Robinson generously provided us with unpublished map and biostratigraphic data. We thank J. Dolan, G. Draper, E. McBride, C. Heubeck, J. Lewis, M. Cloos and J. Pindell for useful discussions and J. Dolan, G. Draper, and E. Calais for their careful reviews of this paper. This work was supported by Grant 17068-AC2 from the Donors of the Petroleum Research Fund of the American Chemical Society to E Mann. University of Texas Institute for Geophysics contribution 1423.
REFERENCES
Beall, R., 1943. Geologic map of the eastern portion of the Cibao Basin, Dominican Republic. Scale 1 : 100,000. Dominican Seaboard Oil Co., New York Office, Santo Domingo (unpubl.). Bermudez, EJ., 1949. Tertiary smaller foraminifera of the Dominican Republic. Cushman Lab. Foraminiferal Res. Spec. Publ. 25, 322 pp. Bourgois, J., Vila, J.-M. and Tavares, I., 1982. Datos geol6gicos nuevos acerca de la regi6n de Puerto Plata (Republica Dominicana). 9th Caribbean Geological Conference, Santo Domingo, Dominican Republic, Vol. 2, pp. 633-635. Bourgois, J., Blondeau, A., Feinberg, H., Glaqon, G. and Vila, J., 1983. The northern Caribbean plate boundary in Hispaniola: tectonics and stratigraphy of the Dominican Cordillera Septentrional (Greater Antilles). Soc. G6ol. Fr. Bull., 25 (1): 83-89. Bowin, C.O., 1975. The geology of Hispaniola. In: A.E.M. Nairn and EG. Stehli (Editors), The Gulf of Mexico and the Caribbean. The Ocean Basins and Margins, Vol. 3, Plenum, New York, pp. 501-552. Bowin, C.O. and Nagle, F., 1982. Igneous and metamorphic rocks of the northern Dominican Republic: an uplifted subduction zone complex. 9th Caribbean Geological Conference, Santo Domingo, Dominican Republic, Vol. 1, pp. 39-50. Calais, E. and Mercier de L6pinay, B., 1995. Strike-slip tectonic processes in the northern Caribbean between Cuba and Hispaniola (Windward Passage). Mar. Geophys. Res., 17: 63-95. Calais, E., Mercier de L6pinay, B., Saint-Marc, E, Butterlin, J. and Schaaf, A., 1992. La limite de plaques d6crochante nord cara'fbe en Hispaniola: 6volution pal6og6ographique et structural c6nozoYque. Bull. Geol. Soc. Fr., 163: 309-324.
285
Cook, H.E. and Mullins, H.T., 1983. Basin margin. In: EA. Scholle, D.G. Bebout and C.H. Moore (Editors), Carbonate Depositional Environments. Am. Assoc. Pet. Geol. Mem., 33: 539-618. DeMets, C., Gordon, R.G., Argus, D.F. and Stein, S., 1990. Current plate motions. Geophys. J. Int., 101: 425-478. De Zoeten, R., 1988. Structure and Stratigraphy of the Central Cordillera Septentrional, Dominican Republic. MA thesis, University of Texas at Austin, 168 pp. (unpubl.). De Zoeten, R. and Mann, E, 1991. Structural geology and Cenozoic tectonic history of the central Cordillera Septentrional, Dominican Republic. In: E Mann, G. Draper and J. Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary Zone in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 265-279. De Zoeten, R., Draper, G. and Mann, E, 1991. Geologic map of the northern Dominican Republic (scale 1:150,000). In: E Mann, G. Draper and J. Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary Zone in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: Plate 1. Dickinson, W.R., 1970. Interpreting detrital modes of greywacke and arkose. J. Sediment. Petrol., 40: 695-707. Dickinson, W.R. and Suczek, C.A., 1979. Plate tectonics and sandstone composition. Am. Assoc. Pet. Geol., 63: 2164-2182. Dickinson, W.R., Beard, L.S., Brakenridge, G.R., Erjavec, J.L., Ferguson, R.C., Inman, K.F., Knepp, R.A., Lindberg, EA. and Ryberg, ET., 1983. Provenance of North American Phanerozoic sandstones in relation to tectonic setting. Geol. Soc. Am. Bull., 94: 222-235. Dixon, T., Farina, E, DeMets, C., Jansma, E, Mann, E and Calais, E., 1998. Relative motion between the Caribbean and North American plates and related plate boundary zone deformation based on a decade of GPS measurements. J. Geophys. Res., 103: 15,157-15,182. Dohm, C.F., 1943. Geologic map of and report on the western portion of the Cibao Basin, Dominican Republic. Scale 1 : 100,000. Dominican Seaboard Oil Co., New York Office, Santo Domingo (unpubl.). Dolan, J.E, Mann, E, de Zoeten, R., Heubeck, C., Shiroma, J. and Monechi, S., 1991. Sedimentologic, stratigraphic, and tectonic synthesis of Eocene-Miocene sedimentary basins, Hispaniola and Puerto Rico. In: E Mann, G. Draper and J. Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary Zone in Hispaniola. Geol. Soc. Am. Spec. Pap., 262:217-263. Dolan, J., Mullins, H. and Wald, D., 1998. Active tectonics of the north-central Caribbean region: oblique collision, strain partitioning and opposing slabs. In: J. Dolan and E Mann (Editors), Active Strike-Slip and Collisional Tectonics of the Northern Caribbean Plate Boundary Zone. Geol. Soc. Am. Spec. Pap., 326: 1-61. Draper, G. and Nagle, E, 1991. Geology, structure and tectonic development of the Rfo San Juan complex, northern Dominican Republic. In: E Mann, G. Draper and J. Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary Zone in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 77-95. Draper, G., Guti6rrez, G. and Lewis, J.F., 1996. Thrust emplacement of the Hispaniola peridotite belt: orogenic expression of the mid-Cretaceous Caribbean arc polarity reversal? Geology, 24:1143-1146. Eberle, W., Hirdes, W., Muff, R. and Pelaez, M., 1982. The geology of the Cordillera Septentrional (Dominican Republic). 9th Caribbean Geological Conference, Santo Domingo, Dominican Republic, Vol. 2, pp. 619-632. Graham, S.A., Ingersoll, R.V. and Dickinson, W.R., 1976.
286 Common provenance for lithic grains in Carboniferous sandstones from Ouachita Mountains and Black Warrior Basin. J. Sediment. Petrol., 46: 620-632. Grindlay, N., Mann, E and Dolan, J., 1997. Researchers investigate submarine faults north of Puerto Rico. EOS, 78: 404. Groetsch, G.J., 1983. Resedimented Conglomerates and Turbidites of the Represa and Janico Formations, North-Central Dominican Republic. M.S. Thesis, George Washington University, Washington DC, 108 pp. Haq, B.U., Hardenbol, J. and Vail, ER., 1987. Chronology of fluctuating sea levels since the Triassic. Science, 235: 1156-1167. Heubeck, C.E., Mann, E, Dolan, J. and Monechi, S., 1990. Diachronous uplift and recycling of sedimentary basins during Cenozoic tectonic transpression, northeastern Caribbean plate margin. Sediment. Geol., 70: 1-32. Ingersoll, R.V., Bullard, T.E, Ford, R.L., Grimm, J.E and Sares, S.W., 1984. The effect of grain size on detrital modes: a test of the Gazzi-Dickinson point-counting method. J. Sediment. Petrol., 54:103-116. Iturralde-Vinent, M.A. and MacPhee, R.D.E., 1996. Age and paleogeographical origin of Dominican amber. Science, 273: 1850-1852. Joyce, J., 1991. Blueschist metamorphism and deformation on the Samana Peninsula: a record of subduction and collision in the Greater Antilles. In: E Mann, G. Draper and J. Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary Zone in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 47-76. Kesler, S., Sutter, J., Barton, J. and Speck, R., 1991. Age of intrusive rocks in Hispaniola. In: E Mann, G. Draper and J. Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary Zone in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 165-172. Link, M.H., 1982. Slope and turbidite facies of the Miocene Castaic Formation and the lower part of the Marple Canyon Sandstone Member, Ridge Route Formation, Ridge Basin, southern California. In: J.L. Crowell and M.H. Link (Editors), Geologic History of Ridge Basin, Southern California. Pacific Section of the Society of Economic Paleontologists and Mineralogists, Los Angeles, CA, pp. 79-88. Mann, E and Gordon, M., 1996. Tectonic uplift and exhumation of blueschist belts along transpressional strike-slip fault zones. In: G. Bebout, D. Scholl, S. Kirby and J. Platt (Editors), Dynamics of Subduction Zones. Am. Geophys. Union Monogr., 96: 143-154. Mann, E, Draper, G. and Lewis, J., 1991. An overview of the geologic and tectonic development of Hispaniola. In: E Mann, G. Draper and J. Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary Zone in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 1-28. Mann, E, Taylor, E, Edwards, L. and Ku, T., 1995. Actively evolving microplate formation by oblique collision and sideways motion along strike-slip faults: an example from the northern Caribbean plate margin. Tectonophysics, 246: 1-69. Mann, E, Prentice, C., Burr, G., Pena, L. and Taylor, E, 1998. Tectonic geomorphology and paleoseismology of the Septentrional fault system, Dominican Republic. In: J. Dolan and E Mann (Editors), Active Strike-Slip and Collisional Tectonics of the Northern Caribbean Plate Boundary Zone. Geol. Soc. Am. Spec. Pap., 326: 63-123. Middleton, G.V. and Hampton, M.A., 1976. Subaqueous sediment transport and deposition by sediment gravity flows. In: D.J. Stanley and D.J.E Swift (Editors), Marine Sediment Transport and Environmental Management. Wiley, New York, pp. 197-218. Muff, R. and Hernandez, M., 1986. The hydrothermal alteration and pyrite-galena-sphalerite mineralization of a porphyrite
R. DE Z O E T E N and E M A N N intrusion at Palma Picada in the Cordillera Septentrional, Dominican Republic. Natural Resources and Development (Ttibingen, West Germany), Vol. 26, pp. 83-94. Mullins, H. and nine others, 1992. Carbonate platforms along the southeast Bahamas-Hispaniola collision zone. Mar. Geol., 105: 169-209. Mutti, E. and Normark, W.R., 1987. Comparing examples of modern and ancient turbidite systems: problems and concepts. In: J.K. Leggett and G.G. Zuffa (Editors), Marine Clastic Sedimentology: Concepts and Case Studies. Graham and Trotman, London, pp. 1-37. Mutti, E. and Ricci Lucchi, E, 1978. Turbidites of the northern Apennines: introduction to facies analysis (translation by T.H. Nilsen). Int. Geol. Rev., 20: 125-166. Nagle, E, 1979. Geology of the Puerto Plata area, Dominican Republic. In: B. Lidz and E Nagle (Editors), Hispaniola: Tectonic Focal Point of the Northern Caribbean, Three Tectonic Studies in the Dominican Republic. Miami Geological Society, Miami, FL, pp. 1-28. Nilsen, T.H. and EL. Abbot, 1981. Paleogeography and sedimentology of Upper Cretaceous turbidites, San Diego, California. Am. Assoc. Pet. Geol., 65 (7): 1256-1284. Palmer, H.C., 1979. Geology of the Moncion-Jarabacoa area, Dominican Republic. In: B. Lidz and E Nagle (Editors), Hispaniola: Tectonic Focal Point of the Northern Caribbean, Three Tectonic Studies in the Dominican Republic. Miami Geological Society, Miami, FL, pp. 29-68. Peralta-Villar, J., 1985. Geologie und Erzffihrung in der Umgebung des Intrusionsbrekzienkorpers von Los Jobos/Pedro Garcia. Dominikanische Republik. M.S. Thesis, Institut der Universit~it Heidelberg, 111 pp. Pickering, K., Stow, D., Watson, M. and Hiscott, R., 1986. Deep-water facies, processes and models: a review and classification scheme for modern and ancient sediments. Earth-Sci. Rev., 23: 75-174. Pindell, J.L. and Draper, G., 1991. Stratigraphy and geological history of the Puerto Plata area, northern Dominican Republic. In: P. Mann, G. Draper and J. Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary Zone in Hispaniola. Geol. Soc. Am. Spec. Pap., 262:97-114. Redmond, B.T., 1982. Sedimentary Processes and Products: an Amber-Beating Turbidite Complex in the Northern Dominican Republic. Ph.D. Dissertation, Rensselaer Polytechnic Institute, New York, 495 pp. Rosencrantz, E., Ross, M.I. and Sclater, J.G., 1988. Age and spreading history of the Cayman Trough as determined from depth, heat flow, and magnetic anomalies. J. Geophys. Res., 93: 2141-2157. Schlager, W. and Chermak, A., 1979. Sediment facies of platform-basin transition, Tongue of the Ocean, Bahamas. In: L.L. Doyle and O.H. Pilkey (Editors), Geology of Continental Slopes. Soc. Econ. Paleontol. Mineral. Spec. Publ., 27: 193-208. Shanmugan, G. and Moiola, R.J., 1988. Submarine fans: characteristics, models, classification, and reservoir potential. Earth-Sci. Rev., 24: 383-428. 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. Vaughan, T.W., Cooke, W., Condit, D.D., Ross, C.E, Woodring, W.E and Calkins, E C., 1921. A Geological Reconnaissance of the Dominican Republic. Gibson Brothers, Inc., Washington, 268 pp. Walker, R.G., 1984. Turbidites and associated coarse clastic deposits. In: R.G. Walker (Editor), Facies Models (2nd Ed.). Geological Association of Canada, Toronto, ON, pp. 171-188.
Chapter 12
Tectonic and Eustatic Controls on Neogene Evaporitic and Siliciclastic Deposition in the Enriquillo Basin, Dominican Republic
PAUL MANN, PETER E MCLAUGHLIN, JR., W.A. VAN DEN BOLD, S.R. LAWRENCE, and MICHAEL E. LAMAR
There are two distinct facies of Early Pliocene evaporites which formed in the center and edges of the tectonically active Enriquillo basin of the southwestern Dominican Republic (island of Hispaniola). The basin-central deposit may be a late Neogene analog to poorly understood ancient saline giants known in many other parts of the world. Both the basin-center and basin-edge evaporite deposits are important for understanding the complex interplay between tectonic and eustatic effects on the stratigraphy of this part of the North America-Caribbean oblique-slip plate boundary zone. The basin-central evaporite deposit (Angostura Formation) is composed mainly of halite and is approximately 1500 m thick in the Charco Largo-1 well in the center of the basin. A basin-margin evaporite composed mainly of gypsum interbedded in a 1-km-thick section of shallow-marine siliciclastic and carbonate rocks (Arroyo Blanco Formation) is exposed on the margins of the basin and in the Charco Largo-1 well (Las Salinas Formation). Analysis of the seismic reflection data, tied to the Charco Largo-1 and other exploration wells, shows that the deposition of these Lower Pliocene evaporites occurred in a major sub-circular deep in the center of the valley. The location of this deep was controlled by west-northwest to north-striking reverse faults and was separated from the Caribbean Sea to the east by a shallow sill. Source areas for the interbedded siliciclastic rocks lay to the north and northeast in the topographically elevated central range of Hispaniola (Cordillera Central). The depositional and climatic setting of the Enriquillo basin during the Early Pliocene was probably similar to the modern Enriquillo Valley, which is 80 m below sea level (BSL) at its lowest point, separated by a shallow sill from the Caribbean Sea, and receives siliciclastic sediment derived from the Cordillera Central by the Rfo Yaque del Sur. Outcrop exposures of these evaporites were examined along the southern edge of the basin at Loma Sal y Yeso. The exposures occur along a faulted diapir in which halite of the Angostura Formation is extruded as a narrow, 100-m-wide strip along a reverse fault separating the Angostura and Las Salinas Formations. Primary textures in halite and gypsum exposed there are overprinted by alteration, faulting and folding. Interbedded sedimentary rocks are fine-grained mudstone lacking sedimentary structures. The monoclinal Las Salinas Formation overlies the more deformed rocks of the Loma Sal y Yeso and consists of a lower shallow-marine siliciclastic interval, 180 m in thickness, which is overlain by a brackish interval, 1670 m in thickness, containing both siliciclastic and carbonate rocks. The lower shallow-marine interval terminated deposition of the basin-central halite of the Angostura Formation. Correlations between distinctive lithologies cropping out in the Angostura and Las Salinas Formations can be made to the lithologies present in the Charco Largo-1 well. A notable difference is the lack of gypsum in the well but its presence in the Loma Sal y Yeso. The Early Pliocene northeast margin of the Enriquillo basin consists of an originally southwest-facing shelf-slope which was tilted to the north and northwest 20-40 ~ in post-Early Pliocene times by tectonic folding. The tilted beds now form a natural cross-section of the gently sloped, Early Pliocene margin. The sedimentary rocks of the shelf-slope can be divided into a gypsum and oolite-bearing shallow-marine facies and a deeper-marine facies. Nine distinct lithologic groups are recognized, including overlying non-marine rocks of the Late Pliocene-Pleistocene Arroyo Seco Formation. The primary textures of gypsum in the Arroyo Blanco Formation suggests that these were deposited in several meters or less of water in a tidal-flat setting. The nine lithologic groups can be interpreted as three shoaling-upwards cycles consisting of: (1) a basal deeper-marine facies association characterized by basinward-prograding clinoforms composed of grainstones and coral debris; (2) a shallow marine oolitic limestone facies association; and (3) an upper gypsum facies association. The shallow-marine gypsum and oolitic limestone facies associations are wedge-shaped in cross section and occupy a position at the base of the slope. Interbedded marine units onlap and truncate the underlying gypsum units.
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 287-342. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
288
E MANN et al.
The three shallow-marine facies associations of the basin-edge evaporite deposit are interpreted as having been deposited during three eustatic falls in sea level during the Early Pliocene. These eustatic sea-level drops correlate reasonably well to the tectonically active Cibao basin of the northern Dominican Republic. Correlation between the basin-margin and basin-central evaporites in the Enriquillo basin is difficult because of uncertainty in the age of the units. Despite this correlation problem, we propose that the Enriquillo evaporites formed in a shallow water-deep basin setting for three reasons: (1) the 'ramp basin' setting of the basin would promote a 'deep' sub-sea level depression in a coastal setting; (2) shallow water within this 'deep' basin would promote efficient evaporation of seawater spilled over the coastal sill; and (3) multiple spills during times of higher Pliocene sea level would replenish brines to sustain the formation of thick basin central evaporites and explain the three cycles of basin margin evaporites observed.
INTRODUCTION
Significance of evaporites for studies of sedimentary basins The study of evaporites and their associated deposits can provide important information on the paleogeography, eustatic sea-level position and tectonic setting of ancient sedimentary basins. Moreover, several authors have pointed out the close relationship b e t w e e n hydrocarbon occurrence and the presence of evaporites in ancient sedimentary basins (for example, Kirkland and Evans, 1981; Friedman, 1982; Warren, 1986). Evaporites can serve as seals and interbedded anoxic sedimentary rocks can serve as hydrocarbon source rocks. The potential for seals and sources within comm o n l y thick stratigraphic sections makes evapor-
itic facies a desirable target for p e t r o l e u m exploration. Studies of the relationship b e t w e e n evaporite occurrence and paleogeography, tectonic setting, eustatic sea level, and hydrocarbons are best carried out in geologically young basins which have not been extensively modified by repeated tectonic deformation or diagenesis related to either deformation or subaerial exposure. The Enriquillo basin of the southern D o m i n i c a n Republic on the island of Hispaniola (Fig. 1) is an ideal place to study the controls on evaporite formation for the following reasons. (1) The basin contains a thick (up to 4.5 km) succession of mainly shallow-marine siliciclastic and evaporitic rocks deposited since the Late M i o c e n e (McLaughlin et al., 1991). The rocks are slightly d e f o r m e d by gentle folds and high-angle faults along the edges of the basin that were produced in an in-
Fig. 1. Present-day plate tectonics of the Caribbean region. Direction and rates of plate motion relative to the Caribbean plate are from DeMets et al. (1990) and Dixon et al. (1980). The island of Hispaniola straddles the active left-lateral strike-slip zone separating the North America and Caribbean plates. The large amount of plate convergence and topographic uplift of Hispaniola is related to transpression between two thick crustal blocks: the Bahamas carbonate platform to the north and the Cretaceous Caribbean oceanic plateau to the south. Box shows map area shown in Fig. 3.
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN terplate, transpressional setting (Mann et al., 1991b, 1995) (Fig. 1). (2) The basin contains both a massive basincentral halite and gypsum deposit up to 1500 m thick and three intervals of much thinner-bedded, coeval basin-margin gypsum deposits which occur in a siliciclastic section about 1 km in thickness. The basin-central deposit is known from wells in the basin center and from outcrops in a diapir (Loma Sal y Yeso) along the southern margin of the basin (Llinas, 1972a,b). One deep exploration well, Charco Largo-1, penetrates the entire undeformed thickness of the basin-central evaporite deposit and provides an opportunity to compare the geologic history of the basin center with the history of the more deformed basin edges known from outcrop mapping. (3) Sedimentary textures in the basin-central evaporite deposit are recrystallized, but textures in the basin-margin evaporites are primary and closely resemble textures described in other, better-studied areas like the Neogene of the western Mediterranean (Schreiber, 1988). (4) The present-day tectonic, physiographic, and climatic setting of the Enriquillo basin (Fig. 1) is similar to classic conceptions of silled, coastal evaporite basins described in other modern settings like Lake MacLeod, Australia (Logan, 1987) and postulated in ancient settings including the Middle Devonian Elk Point basin of western Canada (Kendall, 1989), the Early Cretaceous South Atlantic (Burke and Seng6r, 1988), the Late Jurassic Louann salt of the Gulf of Mexico (Winker and Buffler, 1988), the Permian Delaware basin of west Texas (Lowenstein, 1988), and the Late Miocene Mediterranean Sea (Hsii et al., 1973; Schreiber, 1988). Although massive basin-central evaporites have formed in all of these settings, a modem or late Neogene analog for these ancient 'saline giants' has never been identified with certainty and remains a major paradox in the study of evaporite deposits. Kendall (1984) proposed three, idealized models for silled basins which he proposed as possible depositional settings of subaqueous evaporites found in 'saline giants' (Fig. 2). (5) Rates of plate movement between the Caribbean and North America plates in Hispaniola are relatively slow (--~2 cm/year, Fig. 1) and it is likely that the physiography of the Enriquillo basin has not changed significantly since the formation of the basin about 5 million years ago (Mann et al., 1995; Dixon et al., 1998). A recent cycle of marine flooding and desiccation, similar to that postulated for the origin of the Pliocene evaporite deposits, has occurred as recently as 4000 years BP (Taylor et al., 1985). (6) There are reports of oil and gas seeps from the margins of the basin (Guerra Pena, 1956; Llinas, 1972a,b). The source, reservoir and relation of these evaporties is unknown but deserves further study.
289
Fig. 2. Three models for the deposition of massive evaporite deposits like the one present in the Enriquillo basin (modified from Kendall, 1984). Because it is presently about 50 m below sea level, the Enriquillo basin is a likely shallow-water, deep-basin setting for evaporites.
OBJECTIVES AND METHODS OF THIS STUDY
This study follows on several previously published studies by the authors and their colleagues on the surface and subsurface geology of the Enriquillo basin (Norconsult, 1983; Mann et al., 1984; Taylor et al., 1985; Mann et al., 1991a,c; McLaughlin et al., 1991; Fig. 3A). This study builds on these previous surface studies by attempting the following. (1) Integrate subsurface seismic reflection data collected by industry in the center of the Enriquillo basin (Norconsult, 1983) with surface mapping and section measuring of the age-equivalent sections exposed by Neogene deformation at the basin margins. (2) Integrate the stratigraphy of the Charco Largo-1 well in the center of the basin with surrounding seismic reflection data and surface outcrops. Use the biostratigraphy of the well to resolve correlation and age estimates of evaporitic and siliciclastic sedimentation in the center of the basin.
290
E M A N N et al.
9
,.
~~o~
~~~ o~
,s::
r
o 9
~
~J
O
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN (3) Establish the thickness, paleobathymetry, and facies of siliciclastic sedimentary rocks that enclose the evaporite deposits in the center of the basin and at its edges. Can these data be used to distinguish between the three types of silled basins proposed by Kendall (1984) (Fig. 2) which may have been the depositional setting for the 'saline giant' found today in the center of the Enriquillo basin? (4) Establish the textural variations in evaporite deposits in various parts of the valley and relate these regional variations to depositional environment and water depth within the Enriquillo basin (for example, Schreiber and Friedman, 1976; Schreiber, 1988). These data, in conjunction with paleobathymetric estimates derived from benthic foraminifera preserved in siliciclastic rocks enclosing the evaporites, can also be used to evaluated the silled basin models shown in Fig. 2. (5) Identify stratigraphic sequences and sequence boundaries and to relate these to changes in Miocene and Pliocene sea level. Traditionally, evaporites are interpreted as forming during times of major regression. However, in a sub-sea-level depression setting like the modern Enriquillo Valley and as shown schematically in Fig. 2 (bottom panel), marine transgression would produce a large area of seawater subject to evaporation in a dry, basinal setting. Do evaporites in these settings mark sea-level highs and spills into such depressions or do they mark regressions? (6) Compare sedimentation in the Enriquillo basin to three classic models of silled evaporite basins proposed by Kendall (1984) (Fig. 2). Does this comparison suggest that the Enriquillo basin could be considered a late Neogene analog to massive basin-central evaporite deposits known from the Gulf of Mexico, Permian basin, and other well-studied localities?
TECTONIC AND GEOLOGIC SETTING OF THE
ENRIQUILLO BASIN Tectonic setting The Enriquillo basin of the southern Dominican Republic is a thick (~5 kin), Plio-Pleistocene basin containing both massive halite and gypsum deposits and thin-bedded unrecrystallized gypsum deposits (McLaughlin et al., 1991) (Fig. 3). The basin forms an elongate valley bounded by active strike-slip and reverse faults related to left-lateral displacement between the North America and Caribbean plates across Hispaniola (Fig. 3B) and trends approximately east-west across the southern part of the island (Mann et al., 199 l a). Transpression in the Enriquillo basin appears to be a response to oblique collision of the Bahamas carbonate platform at a
291
major, offshore thrust front subparallel to the northern coast of Hispaniola (Mann et al., 1995; Dolan et al., 1998) (Fig. 3A). The relation between the general shape of the unsubducted Bahamas Platform and the thrust front north of Hispaniola suggests that as much as half of the island of Hispaniola is underlain by the subducted Bahamas Platform. Local convergence in Hispaniola with present topographic elevations up to 3 km may also be related to the location of the island between the Bahamas carbonate platform to the north and thicker-thanaverage oceanic plateau seafloor of the Caribbean Sea to the south (Fig. 3). Mann et al. (1995) propose that the Enriquillo-Plantain Garden strike-slip fault zone formed in late Neogene time as a response to oblique subduction of the Bahamas Platform because the subduction of the thicker-than-average crust of the Bahamas Platform resisted the eastward motion of the central and northern parts of Hispaniola. A regional, unbalanced cross-section modified from Mann et al. (1991b) and shown in Fig. 3B illustrates several important features of the Cenozoic structural history of Hispaniola. (a) The most prominent folding and thrusting event in Hispaniola is Late Miocene and younger in age and verges southward to southwestward. (b) South to southwest vergence of the central mountain range of the island, the Cordillera Central, is reflected in its slightly asymmetric topographic profile with straighter and steeper slopes along its southwestern flank. (c) Late Miocene and younger reverse and oblique-slip faulting is responsible for the present pattern of morphotectonic units in central Hispaniola, including the distribution of the three major ramp, or thrust-bound, basins the Cibao, San Juan-Azua, and Enriquillo (Mann et al., 1991a). (d) Cretaceous-Eocene island-arc terranes of the northern and central part of the island are topographically high-standing and deeply eroded; the Cretaceous oceanic plateau terrane of the southern part of the island is relatively low-standing and less deeply eroded. The lower altitude of the oceanic plateau in the south may also reflect its foot-wall position relative to the higher-standing hanging-wall block represented by the island-arc terranes in the north (Fig. 3B).
Geologic setting Present-day basin physiography Most of the 2000 km 2 floor of the Enriquillo basin and its westward extension into Haiti (Cul-deSac basin) falls within rain shadows of bordering, fault-bounded mountain ranges up to 2 km high in contrast to the semiarid valley that receives only ~60 cm of rainfall per year, with most falling in the month of June (Garcia, 1976) (Fig. 4). Most of the moisture transported off the Atlantic Ocean
292
E MANN et al.
Fig. 4. Amount of annual rainfall based on rainfall gauges around the Dominican Republic (modified from Garcfa, 1976). Areas of high rainfall correspond to mountain ranges while areas of low rainfall correspond to rain shadows in semiarid valleys separating mountain ranges. Note the correlation between Pliocene evaporite occurrence and the belt of low rainfall in the Enriquillo Valley. by northeasterly trade winds falls as rain on top of the Cordillera Central north of the Enriquillo basin (Fig. 4). This close association of high topography, rain shadows, and semiarid intermontane valleys is the reason why the eastern Enriquillo Valley is the site of the only Neogene occurrence of evaporites in the northern Caribbean. The 262 k m 2 Lago Enriquillo in the Dominican Republic and the 120 km 2 Etang Saumatre in Haiti occupy closed drainage basins in the central part of the valley and are separated by a sill a few hundred meters wide at the international border (Fig. 5). The sill between the two lakes lies between 20 and 40 m above sea level (ASL). All of the Haitian Cul-de-Sac basin is above sea level with the surface of Etang Saumatre at an elevation of about 14 m ASL. About half of the Enriquillo Valley is below sea level (BSL) with the surface of Lago Enriquillo at about - 4 2 m BSL. The northern, deeper part of the Lago Enriquillo has a maximum depth of - 4 0 m and forms the deepest known part of the valley at about - 8 0 m BSL (Taylor et al., 1985). Lago Enriquillo is separated from the Caribbean Sea at Bahia de Neiba at the eastern end of the valley by a sill about 4 m ASL (Fig. 5). This relief results mainly from flood plain deposition near the mouth of the Rio Yaque del Sur, a major river draining into the Bahia de Neiba. In the LANDSAT image of Fig. 5A, these floodplain
deposits appear as a large light-colored delta-shaped area at the eastern end of the valley. As sea level rose about 10,000 years ago, the Enriquillo Valley was flooded as the Caribbean Sea at Bahia de Neiba breached the sill at the eastern end of the valley (Mann et al., 1984; Taylor et al., 1985; Mann et al., 1995). A pristine coral reef, now subaerially exposed in the area around Lago Enriquillo (Fig. 5), records marine conditions from about 10,000 to 5000 years BE From the period from about 5000 to 2800 years BE conditions became increasingly brackish as the eastern end of the valley became blocked by deltaic deposits of the Rio Yaque del Sur (Taylor et al., 1985) (Fig. 5A). Evaporation of the seawater in the semiarid valley produced Lago Enriquillo which presently has about twice the salinity as seawater. Ostracodes from short cores (= 1.15 m) from the lake bed reflect the present period of hypersalinity of the lake waters (Bold, 1990). There are no modem evaporites in or around Lago Enriquillo. Mann et al. (1984) speculated that Pliocene sealevel incursions, similar to the Holocene transgression into the Enriquillo Valley, may have resulted in the deposition of evaporites reported from field studies in the eastern Enriquillo Valley (Llinas, 1972a,b; Bold, 1975) and from deep wells in the center of the basin (Bowin, 1975; de Leon, 1983).
N E O G E N E E V A P O R I T I C A N D S I L I C I C L A S T I C D E P O S I T I O N IN T H E E N R I Q U I L L O B A S I N
293
Fig. 5. (A) LANDSAT image of southern Hispaniola (Haiti and Dominican Republic). (B) Interpretation of LANDSAT image modified from Mann et al. (1995). Key to numbered localities and tectonic features" 1 = Cibao Valley (Dom. Rep.); 2 -- Bonao fault zone; 3 = Cordillera Central (Dom. Rep.) vertical hatching indicates areas over 2 km in topographic elevation; 4 -- Los Pozos-San Juan fault zone; 5 --- Massif du Nord (Haiti); 6 = Plateau Central (Haiti)" 7 = San Juan Valley (Dom. Rep.); 8 = Azua Valley (Dom. Rep.)" 9 = Muertos trench; 10 -- Beata fault zone; 11 -- Sierra Martfn Garcfa (Dom. Rep.); 12 = late Holocene delta of the Rfo Yaque (Dom. Rep.); 13 -- Enriquillo Valley (Dora. Rep.); 14 = southern Sierra de Neiba (Dom. Rep.); 15 = northern Sierra de Neiba; 16 -- Montaignes Noires-Chaine de Marmelade (Haiti); 17 = late Holocene delta of Rfo Artibonite (Haiti)" 18 = northern Chaine des Matheux; 19 = southern Chaine des Matheux; 20 = Cul-de-Sac Valley (Haiti); 21 -- Gulf of Gonave and offshore Rfo Artibonite delta; 22 -- Gonave Island; 23 = Enriquillo-Plantain Garden fault zone; 24 = Sierra de Bahoruco (Dom. Rep.) and Massif de la Selle (Haiti); 25 = Mirogoane Lakes (late Quaternary pull-apart basin); 26 = eastern Massif de la Hotte. (C) Interpretation of proposed boundaries for the Gonave microplate based on major fault features seen in (A) and (B).
294
R M A N N et al. ~D
9
E ,s= O
,s= ~
~
~0 9 9 ~D ,s= ~0
.=.
9
,..Z,
,-4
,SD
,.~ - ~ ~-~ c~
0
,.~
~'~
~
~
.,..~
~.~
~ ~E
~.~
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN
Main lithologic units of the Enriquillo basin Previous workers including McLaughlin et al. (1991) have mapped several major lithologic units in south-central Hispaniola that range in age from Late Cretaceous to Quaternary (Fig. 6). These units are shown on the geologic compilation map in Fig. 6 that has been simplified from a much more detailed l:150,000-scale geologic map compiled by Mann et al. (1991c). The present physiographic Enriquillo basin or valley is continuous with the Azua basin to the east and the San Juan basin to the north (Fig. 5). Many of the same lithologic units are present in all three basins although thicknesses and ages vary between the central Enriquillo basin and the northeastern margin of the Enriquillo basin (McLaughlin et al., 1991) (Fig. 6).
CENTRAL ENRIQUILLO BASIN
PLEISTOCENE
LAS SALINAS FM. (2182 M) PLIOCENE
Igneous basement The lowest unit consists of Cretaceous to Eocene igneous and volcanisiliciclastic rocks that crop out to the south of the Enriquillo basin in the Sierra de Bahoruco and to the north in the Sierra de Neiba (Fig. 6). The rocks in the Sierra de Bahoruco have been correlated with the Caribbean oceanic plateau province that underlies much of the areas of the Caribbean Sea (Diebold and Driscoll, Chapter 19) (Fig. 1) while the rocks to the north have more arc-like affinities and appear related to the Caribbean arc of Hispaniola (Mann et al., 1991b). Deep wells in the Enriquillo basin have not penetrated crystalline basement so it is not known whether the basin is underlain by the oceanic plateau crust, arc crust, or both. In the regional cross-section of Fig. 3B, the lower-standing oceanic plateau basement is inferred to extend northward as part of an underthrusting foot-wall block beneath northeast-dipping reverse faults along the southern flanks of the Sierra de Neiba and Cordillera Central. A fold-thrust belt of Paleocene-Eocene deepmarine rocks called the Peralta belt is found northeast and north of the Enriquillo basin (Fig. 6). Dolan et al. (1991) and Heubeck et al. (1991) interpret these rocks as back-arc basin deposits that were deformed during the late Neogene closure of the basin (Fig. 3B).
JIMANI FM. (210 M)
295 NE MARGIN ENRIQUILLO BASIN
ARROYO SECO FM. (~ 500 M)
ARROYO BLANCO FM. (~ 830 M)
ANGOSTURA FM. QUITA CORAZA FM (2oo- 700 M) (1562 M)
TRINCHERA FM. (271 M)
TRINCHERA FM. (~ 2500 M)
MIOCENE GAJO LARGO MB. (~ 200 M)
SOMBRERITO FM. (322 M) SOMBRERITO FM. (~ 2500 M)
Fig. 7. Correlation, ages and stratigraphic thicknesses proposed by McLaughlin et al. (1991) and McLaughlin and Sen Gupta (1994) and used in this paper for Neogene formations of the Enriquillo basin.
the youngest, the Sombrerito Formation of Early Miocene to latest-Middle Miocene age (Fig. 7). This unit, consisting of about 500 m of mainly micrite, was deposited in 500 to 1500 m of water in a basinal setting (McLaughlin et al., 1991). Deep wells in the Enriquillo basin have also penetrated rocks of the Sombrerito Formation, so it appears to underlie most, if not all, of the Enriquillo basin.
Neogene rocks Paleogene rocks The most extensive of the lithologic units of the Enriquillo basin shown on the map in Fig. 6 consists of Paleocene-Middle Miocene carbonate rocks. These rocks crop out in the cores of the Sierra de Bahoruco, the Sierra de Neiba, and the intervening Sierra Martfn Garcfa (Fig. 6). The Paleogene carbonate units have not been studied in detail but consist of limestone and local dolomite indicative of shallower-water facies (van den Berghe, 1983). The best-studied carbonate unit in the area is also
Neogene siliciclastic rocks of the Enriquillo basin are divisible into three main units (Fig. 7). A thick (~2 km), latest-Middle Miocene to earliest Pliocene section of deep-marine siliciclastic turbidites (Trinchera Formation) is exposed on the northern flank of the Sierra Martfn Garcfa at the eastern end of the Enriquillo basin (Fig. 6). This unit also crops out on the southern and northern flanks of the Sierra de Neiba and has been encountered in most of the wells of the eastern Enriquillo Valley. Previous studies by McLaughlin (1989, 1991)
296 have shown that this unit formed a large, timetransgressive siliciclastic wedge produced during the Late Miocene-Early Pliocene uplift and erosion of the Cordillera Central north of the Enriquillo basin (Fig. 3). The present outcrop area of the unit on the northern flank of the Sierra Martfn Garcfa appears to have been the main depocenter of the siliciclastic wedge as thicknesses of the unit drop off dramatically to the west in the Enriquillo Valley and to the east in the Azua basin. Early Pliocene to Quaternary shallow-marine and fluvial siliciclastic and evaporitic rocks conformably overlie the deep-marine Trinchera Formation. These rocks crop out most extensively along the north flank of the Sierra Martin Garcfa (Arroyo Blanco Formation), the north flank of the Sierra de Bahoruco (Las Salinas Formation), and in the Azua and San Juan basins (Arroyo Blanco Formation) (Fig. 6; for a detailed review of all stratigraphic nomenclature, see Mann et al., 1991a). All wells drilled in the Enriquillo and Azua area have penetrated the equivalent Arroyo B lanco and Las Salinas Formations which range in thickness from about 1 to 2 km. The Early Pliocene to Quaternary unit includes both outcrop areas of Pliocene evaporites in the Enriquillo Valley: a southern area in the Loma Sal y Yeso ('Hill of Salt and Gypsum') along the southern edge of the basin, where it is known as the Las Salinas Formation, and a northern area near the eastern edge of the basin where it is known as the Arroyo Blanco Formation (Fig. 7). The Loma Sal y Yeso is an elongate anticline bounded by thrust faults separating the Sierra de Bahoruco from the Enriquillo basin (Fig. 3B). Llinas (1972a,b) interpreted the massive and bedded salt and gypsum along the axis of the fold as a diapir mobilized during folding and classified the structure as a 'diapiric anticline'. In more modem terminology, the anticline probably represents a hanging-wall anticline developed along the frontal thrust of the Sierra de Bahoruco (Vann et al., 1986). The northern area of basin-edge evaporites interbedded with siliciclastic rocks (Arroyo Blanco Formation) lies about 30 km to the northeast where they crop out along the eastern bank of the Rio Yaque del Sur. Evaporites consist of beds of gypsum 0.1 to 10 m thick and are interbedded in a siliciclastic sequence up to 1 km in thickness. The section strikes east to northeast, has an average dip of 20 ~, and is well exposed in about a dozen stream sections draining eastward into the Rio Yaque del Sur. The exceptionally continuous exposure of the resistant beds of gypsum, conglomerate, and sandstone visible on aerial photographs is attributed to planation by the Rio Yaque del Sur prior to the incision of the present-day modem fiver channel. Evaporites are reported from three deep wells in the center of the Enriquillo basin about equidistant
R MANN et al. between the northern and southern outcrop areas discussed in this paper (Fig. 6). Llinas (1972a,b), Bowin (1975) and de Leon (1983) present a summary of the logs of the Mella 1 and 2 wells, which were drilled in the early 1970s. These wells bottomed at depths of 2673 m and 3328 m, respectively, but did not penetrate the entire thickness of the massive salt and gypsum. De Leon (1983) presents a summary log of the Charco Largo-1 well, which was drilled by Superior Oil Company (now Mobil) in 1980 and penetrated a 1966-m-thick section of massive salt and gypsum, and bottomed at a total depth of 4830 m. Prior to drilling the well, Superior Oil Company acquired a dense grid of multichannel seismic profiles over the central part of the valley (Fig. 6, inset map).
METHODS Field-based mapping Evaporites of Loma Sal y Yeso were mapped at a scale of 1:25,000 using 2x enlargements of U.S. Defense Mapping Agency 1:50,000 topographic maps available through the Dominican Cartographic Institute in Santo Domingo. Aerial photographs at a scale of 1:40,000 are used in this paper and by Mann et al. (1991a) to interpret the along-strike continuity of strike ridges for sequence stratigraphic interpretation and to identify major faults. Detailed field mapping focused on the area of an open-pit salt mine on the north-central flank of the Loma Sal y Yeso and the active open-pit gypsum mines operated by the Cooperaci6n Dominicana de Empresas Estatales (CORDE) of the Dominican government on the eastern and southeastern flanks of the hill near the town of Las Salinas. We measured one long (1850 m) section of siliciclastic rocks north of the Loma Sal y Yeso (Arroyo del Pozo), which dip northwards into the Enriquillo basin and presumably intersect the wells in the basin center at depth. We carried out a similar style of mapping of the evaporites to the west of the Rio Yaque del Sur. Because of the exceptional natural exposures of laterally continuous strike ridges of gypsum, we made a special effort to identify and measure the thickness of sedimentary sequences using a Brunton compass mounted on a Jacob's staff. We measured five sections which ranged from 140 m to 930 m in thickness.
Subsurface mapping In order to attempt correlations between the basinedge exposures of evaporites at Loma Sal y Yeso and along the western side of Rio Yaque del Sur, we integrated the results of subsurface mapping done by S. Lawrence (Norconsult, 1983) using seismic lines
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN originally collected by Superior Oil Company in the late 1970s and now owned by Mobil. Mobil released the use of these lines to Mann for use in this study.
Analysis of Charco Largo-1 well cuttings Mobil also made available the well report on the Charco Largo-1 well and released cuttings from the well to E McLaughlin and W. Bold who made a complete analysis of the age, facies and paleobathymetry of the well using planktonic foraminifera and ostracodes. These results are discussed with previously published biostratigraphic results from outcrop samples (Bold, 1975; McLaughlin et al., 1991) and shallow core samples from Lago Enriquillo (Bold, 1990).
GEOLOGY OF THE BASIN CENTER INFERRED FROM INTERPRETATION OF MULTI-CHANNEL SEISMIC LINES TIED TO WELLS Key reflectors used in mapping and their correlation to outcrops Subsurface analysis of seismic data is based on the basin-wide recognition of several major seismic
297
reflectors that are correlated to certain horizons at the base of or within sedimentary formations described from wells and surface mapping (Fig. 8). T i m e structure maps were then made in two-way time to these horizons. We illustrate these key reflectors using some of the longer lines in the seismic data set (locations of these lines shown on Fig. 8). Key seismic reflectors in the Enriquillo basin seen on selected lines shown on the inset map on Fig. 6 include the following units.
Top of the Neiba-Plaisance Formation This unit corresponds to the P a l e o c e n e - M i d d l e Miocene carbonate section shown on the regional map in Fig. 6. This contact has not been mapped in detail in the field so the origin of the seismic reflector remains unknown. This reflector is not present on all lines in the study area and for that reason no time map was made to the base or top of this unit. The reflector is present on regional line 111 through the area of the Charco Largo-1 and Mella wells (Fig. 9) as well as on lines 1410 and 410 in the Lago Enriquillo area (Fig. 10). On these lines, the top Neiba-Plaisance reflector is conformable with the overlying top Sombrerito reflector.
Fig. 8. Time structural map (contour interval 50 ms) of the reflector interpreted as the top of the earliest-Late Miocene Sombrerito Formation. Complete track map of lines used to map this reflector is shown in Fig. 6. Major faults subdivide the subsurface of the basin into several major structural blocks. The two central blocks have undergone the greatest amount of subsidence in post-Miocene time because they are being dynamically depressed by thrust or strike-slip faults at their margins. Much of the present-day surface of the Lago Enriquillo block is about 50 m below sea level and is isolated from the Caribbean Sea by a low sill about 10 m above sea level at the Bahfa de Neiba and Mella blocks. The lined pattern at the top of the map indicates the zone of recent left-lateral shearing along the Enriquillo-Plantain Garden fault zone of Mann et al. (1995). The inset box shows the GPS-derived rate and direction of Caribbean plate motion relative to a fixed point on the North America plate (Dixon et al., 1998).
298
P. MANN et al.
Fig. 9. Regional migrated multi-channel seismic line 111 from the central part of the Enriquillo basin (location on Fig. 8). The line crosses the deepest three blocks of the central part of the basin (Angostura, Mella, and Vicente Noble). Prominent reflectors are indicated in the left margin. Mapping of the upper four of these reflectors was used to make the four time structural maps presented in this paper.
Fig. 10. Regional migrated multi-channel seismic line 1410 (to left) and 410 (to right) from the central part of the Enriquillo basin (location on Fig. 9). The line crosses the Lago Enriquillo block and the edge of the Mella block. Prominent reflectors are indicated in the center margin. Mapping of the upper four of these reflectors was used to make the four time structural maps presented in this paper. The lenticular seismic anomaly is interpreted as a reefal mound at the top of the Sombrerito Formation.
Top of the Sombrerito Formation This horizon forms a p r o m i n e n t reflection surface over m o s t of the basin and offshore in Bahfa de N e i b a and has b e e n penetrated by all the wells shown on Fig. 8. The p r o m i n e n c e of the horizon m a y be related to the m a j o r velocity change sepa-
rating the m a i n l y carbonate rocks of the Sombrerito F o r m a t i o n from the overlying siliciclastic rocks of the Trinchera F o r m a t i o n (Fig. 7). The horizon is present on regional line 111 through the area of the Charco Largo-1 and M e l l a wells (Fig. 9), on lines 1410 and 410 in the L a g o Enriquillo area (Fig. 10),
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN
299
Fig. l l. Regional migrated multi-channel seismic line 133 across the westernmost part of the Enriquillo basin (location on Fig. 8). The line crosses the Lago Enriquillo block and the left-lateral Enriquillo-Plantain Garden fault zone from north to south (see Mann et al., 1995, for three other seismic lines crossing the EPGFZ to the east of Isla Cabritos). The fault is inferred to be the prominent down-to-the-south fault adjacent to the Cabritos-1 well. Two prominent reflectors are indicated. and on line 133 across the Cabritos anticline in Lago Enriquillo (Fig. 11). On these lines, the top Sombrerito reflector is conformable with the overlying base evaporites reflector but is disconformable with the overlying top evaporites reflector, probably as a result of deformation of the evaporite unit. In the Lago Enriquillo area, the top of the Sombrerito Formation is marked by a lenticular seismic anomaly interpreted as a reefal mound (Fig. 10).
Base of evaporites (Angostura Formation) This reflector is present only in the area of the basin-central evaporite deposit between the eastern edge of Lago Enriquillo and Laguna Rinc6n. The reflector is present on regional line 111 through the area of the Charco Largo-1 and Mella wells (Fig. 9) and on line 410 in the area east of Lago Enriquillo (Fig. 10). The horizon was penetrated in the Charco Largo-1 and Mella wells. The prominence of the horizon may be related to the major velocity change separating the mainly carbonate rocks of the
Sombrerito Formation from the overlying evaporitic rocks of the Angostura Formation (Fig. 7).
Mid-Las Salinas Formation This horizon forms a prominent reflection surface over most of the basin and offshore in Bahia de Neiba and has been penetrated by all the wells shown on Fig. 8. Although this horizon has been penetrated in all of the wells, its lithology is not clear since there are several abrupt lithologic changes between isolated beds of limestone and evaporites in the middle part of the Las Salinas Formation both in the Charco Largo-1 well and in the Loma Sal y Yeso section (Table 1). We speculate that the Mid-Las Salinas reflector may correspond to a 3-m-thick limestone bed penetrated at a depth of 7075 ft (2156 m) in the Charco Largo-1 well (Table 1). This is the highest occurrence of limestone in the well and suggests that this might have been the last marine influence until the Holocene marine incursion about 10 ka (Taylor et al., 1985). The horizon is present
300
R MANN et al.
Table 1 Correlation between Angostura and Las Salinas Formations exposed in Loma Sal y Yeso and in the Charco Largo-1 well (see Fig. 8 for location) Loma Sal y Yeso
Charco Largo-1 (Fig. 15)
Angostura Formation (Fig. 19) Unit 1 - halite (200 m-minimum thickness) Unit 2 - gypsum (350 m-minimum thickness)
interval 8700-13,750 ft or 2652-4191 m (1539 m of halite and shale; shallow-marine fossils present at 8390 ft or 2557 m) not present in thick sections
Unit 3 - gypsum and massive gray mudstone (50 m)
interval 8625-8700 ft or 2628-2651 m (23 m of shale)
Las Salinas Formation in Arroyo el Pozo (Fig. 21) Unit 4 - shallow-marine limestone of 'Razorback Ridge' (3 m)
interval 8610-8625 ft or 2625-2629 m (4.6 m of dense white limestone)
Unit 5 - ripple-marked sandstone and interbedded siltstone
(150 m)
interval 8510-8610 ft or 2594-2624 m (30 m of shale capped by a 6-m-thick bed of anhydrite)
Unit 6 - trough-cross bedded sandstone and interbedded siltstone (390 m)
interval 7075-8510 ft or 2156-2484 m (437 m of 25-100 ft or 8-30-m-thick beds of sandstone interbedded in shale)
Unit 7 - coquina beds interbedded with siltstone (95 m)
10 ft or 3-m-thick limestone bed at 7075 ft or 2156 m (?)
Unit 8 - growth position coral reefs, reworked coral debris interbedded in siltstone (225 m) Unit 9 - recrystallized micrite and calcarenite interbedded in siltstone (955 m)
7075 ft or 2156 m (?) to (?); no coral or limestone distinguished above this horizon
on regional line 111 through the area of the Charco Largo-1 and Mella wells (Fig. 9) and on lines 1410 and 410 in the Lago Enriquillo area (Fig. 10).
Main subsurface structural blocks of the Enriquillo basin Mapping of all of these reflectors indicates that the subsurface of the Enriquillo basin is divisible into four fault-bounded structural blocks. We discuss each of these blocks and their bounding faults as defined by the depth in two-way travel time to the top of the Sombrerito surface (Fig. 8).
Bahia de Neiba block This block is isolated between the South Martfn Garcfa fault zone, a probable oblique-slip thrust zone along the southern edge of the 1200-m-high Sierra Martin Garcfa, fault A, and the Bahoruco fault zone, the frontal thrust of the Sierra de Bahoruco (Fig. 8). To the east, this block is probably truncated by the Beata fault zone, a northeast zone of probable oblique-slip normal faulting that defines the straight, northeast-trending coast and shelf edge of the southcentral Dominican Republic (Mann et al., 1991a; Mauffret and Leroy, Chapter 21). Fault A is a subsurface fault with a transverse northerly to northwesterly strike that is present in the sill area separating Bahia de Neiba from the interior of the valley. The Bahfa de Neiba block exhibits a large, closed structural high along the Bahoruco fault zone and a low along the South Martin Garcfa fault zone (Fig. 8). The Palo Alto-1 well is a dry exploration well drilled
in the area between the high and low in the approximate center of the block (Fig. 8). A zone of complex northeast to northwest-striking faults in the eastern part of the block is perhaps related to movements along the northeast-striking Beata fault zone.
Vicente Noble block This block on the northeastern edge of the basin is isolated between the South Martfn Garcfa fault zone to the southeast, fault A, and the left-lateral Enriquillo-Plantain Garden fault zone to the north (Fig. 8). The Sierra Martfn Garcfa forms the probable extension of this block to the southeast as the intervening area is occupied by unfaulted strata of the Trinchera and Arroyo Blanco Formations dipping to the northwest (Fig. 6). The steep plunge of the Sierra Martfn Garcfa anticline may be terminated in the subsurface by fault A which exhibits a reverse throw to the southwest. Reverse throw on fault A may be the mechanism for monoclinally tilting the exposed late Neogene stratigraphic section of the western Vicente Noble block to the northeast (Fig. 9). Fault A is a subsurface fault with a transverse northwesterly strike in the area between outcrops of the Arroyo Blanco and Trinchera Formations on the northwest flank of the Sierra Martfn Garcfa and the area of Quaternary deltaic sedimentation of the Rfo Yaque del Sur (Fig. 6). Fault A is shown on seismic line 111 in Fig. 9 as a broad zone of transparent to chaotic reflections that bounds a closed structural high on its northeastern edge. Most of the reflections from the Vicente Noble block appear to originate from the Arroyo Blanco Formation (Fig. 9). Deeper
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN reflectors are not apparent, possibly because of the steeper dip of the strata within the block. Mella block This basin-central block is isolated between the subsurface and transverse faults A to the east and fault B to the west (Fig. 8). The block is bounded to the north and south by the Bahoruco and EnriquilloPlantain Garden range-front faults, respectively. The block is strongly downthrown along fault A with about 4 seconds of structural relief between the closed structural high on the upthrown Vicente Noble block and the closed structural low on the downthrown Mella block (Fig. 9). The block rises to the north and south and therefore exhibits a broad saddle shape with its largest low in the approximate center of the basin. The western edge of the Mella block is formed by reverse throw on fault B to the approximate level of the Mid-Las Salinas reflector (Fig. 9). The Charco Largo-1 and Mella wells are drilled into a hanging-wall anticline on the upthrown side of fault B. Two closed structural highs are present north of the well area and one closed structural high is present south of the well area (Fig. 8).
Angostura block This basin-central block is isolated between the subsurface and transverse faults B to the east and C to the west (Fig. 8). The block is bounded to the north and south by the Bahoruco and Enriquillo-Plantain Garden range-front faults, respectively. Fault B exhibits much less reverse throw (~250 ms) than fault A to the east (~4 s) (Fig. 9). Large structural lows are present in the center of the block along with some short, discontinuous faults with apparent normal throws (Fig. 10). The structural contours define an upward slope towards the Enriquillo-Plantain Garden and Bahoruco blocks from the structural lows in the approximate center of the basin. Strong westward deflection of the structural contours near the Enriquillo-Plantain Garden fault zone may reflect shear effects related to left-lateral displacement on this fault (Fig. 8). These deflected contours underlie a zone of en-echelon anticlines and tectonically uplifted Holocene reef deposits in the area of Lago Enriquillo (Mann et al., 1995).
Lago Enriquillo block This block is isolated between the subsurface and transverse faults C to the east and fault zone D to the west (Fig. 8). The block is bounded to the south and north by the Bahoruco and Enriquillo-Plantain Garden range-front faults, respectively. The Lago Enriquillo block is distinguished from the basin blocks to the east by a much thinner late Neogene siliciclastic and evaporitic section. The thinner siliciclastic section results from its greater distance from
301
source areas in the Cordillera Central and the thinner evaporitic section probably results from its greater distance from the Caribbean Sea at Bahfa de Neiba (Fig. 8). Fault C exhibits minor normal throw up to the level of the base of the evaporites as seen on line 410 in Fig. 10. The structural contours are relatively flat across the Lago Enriquillo block as seen on line 1410 in Fig. 10 although a central sag defines a large structural low between Isla Cabritos and the Bahoruco fault zone. This structural low is manifested in a large gravity minimum seen plotted along line 133 crossing the Cabritos faulted anticline from north to south (Fig. 11). North-striking faults of fault zone D, en-echelon normal faults in the area of Isla Cabritos, and the anticline defining Isla Cabritos itself may be secondary structures developed by leftlateral shear along the Enriquillo-Plantain Garden fault zone (Fig. 11). Mann et al. (1995) document a 500-m-wide swath of tectonically uplifted Holocene coral reef deposits above these subsurface features.
Evolution of subsurface blocks of the Enriquillo basin through time Comparison of the three subsurface structural maps for the times of latest-Middle Miocene (top Sombrerito reflector), Early Pliocene (top evaporites/Angostura reflector) and mid-Pliocene (Mid-Las Salinas reflector) reveals information on the evolution of the subsurface fault systems active in the basin during these times. Because deformation is cumulative, each map including those of the oldest top Sombrerito reflector reflects the youngest deformation event in addition to the effects of all older events. For this reason, these maps do not provide a true picture of the structure at the time each surface formed. However, the maps do show how fault movement ceased in some areas during the times of the younger, Mid-Las Salinas reflector.
Top Sombrerito reflector It is unclear from the structural maps how active the subsurface faults were during late-Middle Miocene times of deposition, since reflectors below the top Sombrerito reflector are not well imaged (Figs. 9 and 12). Surface studies of the Sombrerito Formation indicate a relatively uniform pelagic carbonate environment ranging in water depth from 1000 to 1500 m over a large area of south-central Hispaniola (McLaughlin et al., 1991). The large slopes observed on the top Sombrerito structural map therefore were likely produced by deformation after the deposition of the upper Sombrerito Formation in late-Middle Miocene time (Fig. 12).
302
P. MANN et al.
Fig. 12. Time structural map (contour interval 50 ms) of the reflector interpreted as the top of the Middle-earliest-Late Miocene Sombrerito Formation. Complete track map of lines used to make this map is shown in Fig. 7. Boxes indicate areas of detailed mapping along basin margins.
Base evaporites (Angostura Formation) It is unclear from the structural maps how active the subsurface faults were during the Early Pliocene, since stratigraphic growth adjacent to these faults is not imaged (Figs. 9 and 13). The Angostura Formation exhibits a large variation in thickness some of which may be related to deformation and/or compaction by the overlying siliciclastic formations (Fig. 9). The localization of the evaporites in the central area of the basin adjacent to subsurface faults A and B suggest that reverse movement on these faults may have been active at this time and created the sub-circular topographic low in which the evaporites formed (Fig. 13). It is interesting to note that these north-northwest to northwest-striking faults would be optimally oriented to accommodate reverse fault movement if one assumes that the same approximately east-west direction of interplate slip operated in the mid-Pliocene as observed today with GPS geodesy (Dixon et al., 1998) (Fig. 8). The more northwesterly striking faults bounding the evaporites in the central part of the valley would therefore be prone to reactivation as reverse faults whereas the more east-west-trending faults bounding the western part of the valley would be prone to reactivation as oblique-slip faults.
Mid-Las Salinas reflector During the mid-Pliocene, fault B is covered by undeformed reflectors correlated with the middle
part of the Las Salinas Formation (Fig. 14). Fault A, a much wider zone of deformation, extends much higher into an area of unresolved reflectors above one second of two-way travel time (Fig. 9). The arcuate pattern of closed structural lows on the Angostura and Lago Enriquillo blocks may reflect increased overthrusting of the similarly shaped frontal thrust of the Sierra de Bahoruco. Flowage of salt and gypsum southwestward towards the subaerial exposures at Loma Sal y Yeso from the Angostura block may also contribute to this large, closed structural low in the center of the basin (Fig. 14). Large folds in the Mid-Las Salinas reflector beneath Lago Enriquillo may reflect increasing activity and movement along the left-lateral EnriquilloPlantain Garden fault zone (Fig. 14). Three folds are present with east-west fold axial traces sub-parallel to the Enriquillo-Plantain Garden fault zone. The most northerly syncline north of the interpreted trace of the Enriquillo-Plantain Garden fault zone is a large syncline that overlies the deepest part of the Lago Enriquillo and the lowest part of the valley at - 8 0 m BSL. The central fold is a large, doubly plunging anticline with geomorphic expression on Isla Cabritos in the center of Lago Enriquillo. This fold tectonically elevates Holocene reef and algal tufa deposits that fringe the island (Mann et al., 1995). A dry exploration well, Cabritos-1, was drilled on the crest of this structure at the eastern end of the island (Mann et al., 199 l a) (Fig. 11). The
N E O G E N E E V A P O R I T I C A N D S I L I C I C L A S T I C D E P O S I T I O N IN THE E N R I Q U I L L O BASIN
303
Fig. 13. Time structural map (contour interval 50 ms) of the reflector interpreted as the base of the Early Pliocene massive evaporites of the Angostura Formation. Complete track map of lines used to make this map is shown in Fig. 7. Boxes indicate areas of detailed mapping along basin margins.
Fig. 14. Time structural map (contour interval 50 ms) of the reflector interpreted as the middle part of the middle to Late Pliocene siliciclastic rocks of the Las Salinas Formation. Complete track map of lines used to make this map is shown in Fig. 7. Boxes indicate areas of detailed mapping along basin margins.
304
R MANN et al.
MOBIL CHARCO LARGO NO. 1
southernmost fold is a large syncline that underlies the part of the lake between the island and the frontal thrust zone of the Bahoruco fault zone (Fig. 14). This syncline corresponds to the large gravity minimum plotted on the seismic line in Fig. 11.
A
FORMATION
ENVIRONMENT
-, ._-, ._-, ._-,
JtMANI
AND
QUAT.
MARINE
?
',MESTONC Regional pattern of strain partitioning in the Enriquillo basin The large folds seen in the structural map of the Mid-Las Salinas reflector may form by a process of 'strain partitioning', whereby part of the regional transpression across this part of the island is accommodated by left-lateral slip along the EnriquilloPlantain Garden fault zone and part of the transpression is accommodated by shortening normal to the trend of the Enriquillo basin (cf. regional cross-section in Fig. 3B). Based on GPS observations and modeling, Dixon et al. (1998) predict a minimum rate of 8 -t- 4 mm/year of left-lateral slip along the Enriquillo-Plantain Garden fault zone. This fault therefore accommodates about 38% of the total rate of 21 -t- 1 mm/year between the Caribbean and North America plates at the longitude of Hispaniola. Well-developed folding at the Mid-Las Salinas horizon in the western part of the valley and waning thrust movement on fault B in the central part of the valley may indicate a relative increase in the amount of strike-slip faulting at the expense of regional shortening along faults like fault B. The direction of the plate vector determined from the GPS study by Dixon et al. (1998) indicates that the majority of the interplate motion is strike-slip in a roughly east-west direction (Fig. 8).
STRATIGRAPHY AND MICROPALEONTOLOGYOF THE CHARCO LARGO-1 WELL
Charco Largo-1 well The Charco Largo-1 well, drilled on the Mella structural block (Fig. 8), provides a complete stratigraphic record of sedimentation and tectonics in the central part of the Enriquillo basin. The well was drilled to a depth of 4830 m by Superior Oil Company (now Mobil) in 1980 and penetrated five formations: Sombrerito, Trinchera, Angostura, Las Salinas, and Jimanf (Fig. 15). Lithologic descriptions are taken from unpublished company reports used with the permission of Mobil. Lithologic records allowed us to compile a histogram of the thickness of halite beds in the well (Fig. 16). Micropaleontologic analysis for ostracodes (by Bold) and foraminifera (by McLaughlin) was undertaken for 143 cuttings samples representing 9.15 m (30 ft) composites at intervals of 30.5 m (100 ft),
61o-
2
1219-
4
ii
BRACK, SH
1829i 6
2438
"~,
8
L I M E S T O N E,~
~
VERY SHALLOW MARINE
,,.....,..,., .'1".,;, ,T. -"2g. 9 ,'l', -"
3048-
,.,-o, ,- 9 ,- 9 ,-., . . . . . . . . .................
10
-.y.=.y.,.y.=y. HYPERSALINE ,.'.,~'A'A'.," "r
3658- 12
,.'A"A'A'.,~ ........ ..-.......-..........j% =_%=_%=..'~=~ EARLIEST PLIOCENE
4267-
14' !
LATE !MIOCENE
"V'.'~.'V'.V, ''''";';';'""
' TRINCHERA
i = 1 / MIDDLE il / M I O C E N E ![
j4877.
16 J
1
'=~'==""'=........ "-.--'.'.'-,
OPEN
........ "
SOMBRERITO
"
MARINE
~=
=
Fig. 15. General stratigraphic section of the Charco Largo-1 well of the central Enriquillo basin. Lithologies are based on a detailed well log and an accompanying report provided to us by Mobil; ages and environments of deposition are based on micropaleontologic work by McLaughlin and Bold using sample material taken from well cuttings. Key to symbols: brick pattern = deep-water limestone; dashed pattern = shale; v pattern -massive halite; stipple = sandstone and shale; heavy black dots -- conglomerate. where possible. The biostratigraphic findings in this study, which are summarized on Fig. 17, support the ages and water depths proposed in previous studies by Bold (1975) and McLaughlin et al. (1991) using surface samples mainly from the eastern part of the Enriquillo basin along the flanks of the Sierra Martfn Garcfa. A summary of the formation names and ages from previous surface studies is given in Fig. 7.
Lithology and age of the Sombrerito Formation in the Charco Largo-1 well The Sombrerito Formation in the Charco Largo-1 well extends from 4508 m to the bottom of the hole at 4830 m (Fig. 15). As in outcrop, the well samples consist of a light tan to light buff, cryptocrystalline
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN limestone rich in planktonic foraminifera (Fig. 17). The lowest part of the Sombrerito Formation, from 4783.5 to 4830 m, is placed in Middle Miocene planktonic foraminiferal zone M11 of Berggren et al. (1995) based on the highest occurrence of the nominate taxa (Fig. 17). Higher in the well from 4570.1 to 4656.1 m, the Sombrerito Formation is characterized by the presence of Sphaeroidinellopsis multiloba in the absence of Neogloboquadrina mayeri, which places this interval at or just above the Middle Miocene-Upper Miocene boundary, in the M12 or M13a zone (Fig. 17). The uppermost sample in this formation (4509.1-4512.2 m) is interpreted to lie in the Upper Miocene M13b zone based on the presence of Sphaeroidinellopsis seminulina and the absence of S. multiloba. Benthic microfauna recovered from the Sombrerito Formation are indicative of a deep-water setting. The presence of Cibicides wuellerstorfi at 4692-4695 m indicates middle bathyal or deeper depths. Higher in the formation, the presence of
Cibicidoides mundulus, Cibicidoides bradyi, Sigmoilopsis schlumbergeri, Anomalina flintii, and Oridorsalis umbonatus indicates an upper bathyal or deeper setting. Similar ages are described from outcrops of the Sombrerito Formation in the Sierra de Neiba along the northern edge of the basin (McLaughlin et al., 1991).
Lithology and age of the Trinchera Formation in the Charco Largo-1 well The Trinchera Formation in the Charco Largo-1 well was encountered from 4237 to 4508 m and consists of 271 m of greenish shale containing Late Miocene foraminifera (Fig. 15) (Table 1). We propose that this interval of Trinchera Formation is equivalent to the Trinchera Formation and possibly the overlying shales of the Quita Coraza Formation of the Azua basin and northern flank of the Sierra Martfn Garcfa (Figs. 7 and 8) (McLaughlin et al., 1991). The 271 m thickness in the well is significantly less than the 1300 to 2200 m of interbedded deep-marine mudstones and turbidite sandstones present in outcrops 30 km to the east in the Azua basin (McLaughlin et al., 1991). In outcrops in both the Azua and San Juan basins, a progression may be traced from outer turbidite-fan facies in the lower part of the Trinchera Formation to inner-fan facies in the upper part. This relation records the north-tosouth or northeast-to-southwest progradation of the Trinchera Formation over the deep-water carbonate environment recorded by the underlying Sombrerito Formation. The Trinchera Formation samples in the well contain an impoverished microfauna (Fig. 17). Planktonic species present are limited to Globigerina spp.
305
and Orbulina universa. The only benthic foraminifer recovered was Siphonina sp. This poor fauna contrasts with the generally more abundant foraminifera reported from the Trinchera Formation to the northeast in the Azua basin and on the northern flanks of the Sierra Martfn Garcfa (McLaughlin and Sen Gupta, 1994). This relation suggests the possibility of a barrier between the Azua area, where up to 6 km of open-water deposits occur, and the basin-central area of the Charco Largo-1 well, where a the thin, impoverished section is found. Based on regional stratigraphic correlation to these outcrops, this interval in the Charco Largo-1 well is interpreted as Upper Miocene to basal Pliocene and therefore correlates well to previous dating of surface exposures (McLaughlin et al., 1991; Fig. 7).
Lithology and age of the Angostura Formation in the Charco Largo-1 well The Angostura Formation in the Charco Largo-1 well consists of 1562 m of clear to milky, moderately hard, anhedral halite interbedded in greenish-gray, indurated, silty, slightly calcareous shale (Fig. 15) (Table 1). The base of the formation is taken as the lowermost bed of halite of approximately 7.5 m thickness at a depth of 4192 m; the top of the formation is taken as a hard, buff-white limestone of approximately 4.6 m thickness at a depth of 2630 m. The Angostura Formation is interpreted to unconformably overlie the Trinchera Formation in this well. Because of its marginal marine environment and impoverished microfauna, there is little fossil control on the age of the Angostura Formation in the Charco Largo-1 well. However, surface samples from the Angostura Formation in the Loma Sal y Yeso contain an impoverished, brackish water or hypersaline ostracode fauna, considered indicative of Late Miocene to Early Pliocene in age by Bold (1975). The ostracode fauna is interpreted by Bold (1975) to represent a brackish or hypersaline environment consistent with the evaporitic lithology. We consider the lower part of the Angostura Formation to be equivalent to gypsum-bearing deposits of the Arroyo Blanco Formation exposed along the northeastern flank of the basin (Fig. 13). The base of the Arroyo Blanco Formation in this area is well constrained biostratigraphically to lie near the top of the Lower Pliocene (McLaughlin et al., 1991). The halite section of the Angostura Formation appears massive on the generalized well lithologic column of Fig. 15 but the Charco Largo-1 well log shows individual halite beds of halite ranging in thickness from 1.5 to 76 m and averaging 3 m (Fig. 16). The thickest beds of halite are generally confined to the lower one-third and upper one-third
306
P. M A N N et al.
Fig. 16. Histogram showing thickness of halite beds in the Charco Largo-1 well as reported on the detailed well log. See text for discussion.
~
9 [] . . . .
-~ .
.
.
.
.
.
.
.[]
[]
l 9 m
II
.
~
~
~
~
[][][]
-- -- --
l .
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
i . . . . .
.
l []
--
9 []
[]
[]
9 -- i
l
[] [] [] _ ~ - [] 9 []
salebrosa - - C~::o~,on,
II
m _~ _ _~ m _ P . , ~ = c ~ . , ~ . a ~=..~a [] [] [] [] Cyonce,ss,mOis
m m
m
m--mm-
m m
iiimmmmmmmmmmmmm
m
[] i
m
mmmi mmmm I-,~ll immm i iN
m
m
m m m
[]
m mm
mmm
mmmmmmmmmmmmmm
mm
[_mmmmmmmmmmmmm-mnnmm-nmmmmm-
mm
I,
m-m
:m -m
12 -i o,;o
o, = o, ~ -~ N
====
c,
u
(3) 9
1~
(~
l,~
.~
oo
9r
~_
I
9-, p i ' - ; p
pI
cn o
o
; .~ I'~ (JO
J~ 0
clO U1 .4~ "-,4 CO P,J C. O1
~
0~ ~
~
o
=
0= ~ '
9
:=~
.F,
.~ m
,
SOMBRERITO
~9
.
-,4 -,4
,,
CypnCe~s mexlcana cyom~oos~s vidua Hem~c'/ons reticulata Penssocythendea so. A r.Jmnoc./lhere stap/inl ~arwmula so. ~leterocyons sO. :~ensscc'/thendea ex~lis Cative/la haws CypdcYe~sso. ~edssocythendea sO. Doleroc/pna so. ~udana gatunensis Ammonia sO. E!ohidium sO. Archaias so. Siphon9 sO. Ondorsalis unbonatus Sigmoiioosis schlumergen Anomalina flintii Cibicicloides brady9 CibiG~doidesmundulus Cibic/de$ wuellerstorfl GloOigenna sO. OrOulina umversa 31oOorotalia rnenradii HastigerYna siphonifera 3pharoidine//opsis seminu/ina Glotjigennoides trY/oOus 31obigennoides obliquus obliquus 31oDoquadtfna altispira 31obigerma venezuelana 31obigenna nepenthes SplTaeroidinallopsi$ multiloDa Globorotalia mayed
m
mmm
=.
Pensscc/lhendeasuDrugosa
~ ~ ~ ~
TRINcHER
I ~= ~
UPPER MIOCENE
ANGOSTURA
PLiOCENE
L.AS SALINAS
UPPER PLJOcENE
JIMANI
PLIOPLEIST
QUATERNARY LACUSTRINE DEPOSITS
QUA TERNARY
FORMATION
STAGE
Fig. 17. List of foraminiferal and ostracode fauna recovered from well cuttings of the Charco Largo-1 well by McLaughlin and Bold.
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN of the formation. In the middle one-third of the formation, the halite bed are generally thinner and are interbedded with siltstones that make up 5 to 10% of the formation's thickness. The interbedded shale comprises approximately 25% of the formation, with shaley zones 3 to 30 m in thickness. The thinner halite beds of the middle one-third of the formation are interbedded with siltstone. The Angostura Formation is virtually devoid of sandstone. There is little gypsum and anhydrite reported from the halite interval in the Charco Largo-1 well. There are, however, two sills of hornblende-diorite intruded into the central part of the Angostura Formation. The sills range from 3 to 9 m in thickness and probably are related to the same Late PliocenePleistocene episode of alkalic basaltic volcanism that created a number of volcanic bodies, whose rocks extend across south-central Hispaniola (Wadge and Wooden, 1982; Mann et al., 1991a). The Angostura Formation in this well contains a microfauna that reflects deposition in a marginal marine setting during the Pliocene. The association of the ostracodes Cyprideis salebrosa and Cyprideis mexicana indicates a chronostratigraphic position no lower than the upper part of the Lower Pliocene. This is slightly younger than previous reports from outcrops of the formation; Bold (1975) recorded Cyprideis subquadraregularis and Cyprideis pascagoulaensis, which together were considered to suggest a position near the Miocene-Pliocene boundary. Cyprideis salebrosa is indicative of low-salinity environments. The foraminifera Elphidium and Ammonia are present in some of the samples and could indicate brackish-water or hypersaline conditions. In addition, Archaias was found in one sample, suggesting the presence of shallow-marine Thalassia sea-grass environments (Sen Gupta and Schaefer, 1973). Although the formation includes a great thickness of evaporite, reflecting deposition in an evaporative lake system, the varying representation of brackish-water and shallow-marine faunas in a succession suggests that significant variations in salinity occurred in the overall evaporative system.
Lithology and age of the Las Salinas Formation in the Charco Largo-1 well The Las Salinas Formation consists of a 2182m-thick, coarsening-upward section of shale, clay, sandstone, conglomerate, and minor anhydrite (Fig. 15). The formation is present at a depth of 448-2630 m in the Charco Largo-1 well (Table 1). The conglomerate contains clasts of quartz and various types of igneous and metamorphic rocks derived from the erosion of Cretaceous to Eocene arc rocks in the Cordillera Central (McLaughlin et al., 1991). The single evaporite bed in the Las Salinas For-
307
mation is a 4.5-m-thick anhydrite at approximately 2595 m depth; this is approximately 535 m above the highest Angostura Formation halites that occur near 2630 m. The base of the Las Salinas Formation in the well is marked by a 4.6-m-thick, hard buffwhite limestone (Table 1). In outcrops at Loma de Sal y Yeso on the southwest side of the basin, this basal limestone is correlative to a 2-m-thick mollusc-shell-rich sandstone bed with symmetrical wave ripples (McLaughlin et al., 1991) (unit 4 on map in Fig. 18A and on section in Fig. 19A). Because of its resistance to weathering, vertical to overturned bedding attitude, and topographic prominence among more erodible rocks of the Angostura Formation and basal Las Salinas Formation, we informally named this unit 'Razorback Ridge'. Overlying this unit in the well, the lower 100 m of Las Salinas Formation consists of 10-20-cm-thick medium-grained sandstone beds alternating with 5 10-cm-thick siltstone beds. The section above this is characterized by 15-25-m-thick cycles of interbedded, trough-cross-bedded pebbly conglomerate, sandstone, and wave-rippled maroon and green siltstone. In outcrop, sandstone and conglomerate beds are generally more resistant to erosion and produce the prominent hogback topography seen in the aerial photograph of the Loma Sal y Yeso area in Fig. 18B. The Las Salinas Formation includes a more diverse association of ostracodes than the underlying Angostura Formation, but occurrences are similarly scattered. The faunas recovered are mostly composed of marginal marine species of Cyprideis and Perissocytheridea. The foraminifera Elphidium and Ammonia were also recovered. Both these forms are tolerant of variable salinity conditions. Cyprideis salebrosa occurs throughout the formation and Cyprideis similis occurs near the top, indicating a mid-Pliocene or higher position. The above ostracode species are consistent with placement of this formation by Bold (1975) in the Pliocene to Holocene Cyprideis salebrosa zone. The Angostura Formation in the Charco Largo-1 well and in outcrops of the Loma Sal y Yeso is considered Early Pliocene in age.
Lithology and age of the Jimani Formation in the Charco Largo-1 well The Jimanf Formation comprises 210 m of interbedded shales, sandstones, conglomerates and limestones in the well, similar to the facies described for the type section near Jimanf in the westernmost Enriquillo basin (McLaughlin et al., 1991) (Fig. 15). The Jimanf Formation is not present in the exposed section to the northeast of the Loma Sal y Yeso (Fig. 18). Samples of Jimanf Formation from the well have scattered occurrences of ostracodes simi-
308
R M A N N et al.
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN lar to those of the underlying Las Salinas Formation. Diversity is slightly higher in the Jimanf Formation, with as many as seven species in a single sample. Cyprideis salebrosa and Perissocytheridea spp. are present in all ostracode-bearing samples, suggesting a reduced-salinity environment where samples were recovered. None of the samples produced species indicated by Bold (1975) as typical of the Jimanf Formation (e.g., Cyprideis portuprospectuensis, 'Campylocythere' perieri). However, they do include species considered typical of the Jimanf or upper Las Salinas Formations, including Cyprideis similis, C. salebrosa, Perissocytheridea subrugosa, and Limnocythere staplini. The Jimanf Formation is therefore considered Late Pliocene to Pleistocene, in the Cyprideis salebrosa zone, based on these occurrences and on the presence of the Quaternary form 'Campylocythere' perieri in outcrop samples. Lithology and age of the undifferentiated Quaternary rocks in the Charco Largo-1 well
The Jimanf Formation is overlain by Quaternary cover (Fig. 15). The ostracode fauna of the Quaternary section is similar to that of the Jimanf but more samples are fossiliferous and exhibit slightly higher diversities, with up to eight species in a sample. Cyprideis spp. and Perissocytheridea spp. are the most notable faunal elements and suggest continued marginal marine conditions. This unit includes many of the same Plio-Pleistocene species listed by Bold (1975) as typical of the Jimanf and Upper Las Salinas, including Cyprideis similis, C. salebrosa, Perissocytheridea subrugosa, and Limnocythere staplini. It also includes the only occurrences in the well of Cyclocypris sp., Cypridopsis vidua, Hemicypris reticulata, and Darwinula sp.
SURFACE GEOLOGY OF THE LOMA DE SAL Y YESO AREA AND STRATIGRAPHIC CORRELATION WITH THE CHARCO LARGO-1 WELL Stratigraphy of L o m a de Sal y Yeso
Loma de Sal y Yeso forms a low range of hills elevated by a combination of thrust faulting and
309
salt diapirism between the Sierra de Bahoruco and the topographic depression of the Enriquillo basin (Fig. 18). Thrust faults of the Bahoruco fault zone extend in a narrow belt along the mountain front to the Haitian border (Fig. 6). Loma de Sal y Yeso is the only exposed salt-cored fold in the belt along this mountain front and supports subsurface data showing that evaporites are restricted to a small area of the eastern Enriquillo basin (Fig. 4). Mapping in the quarry areas of Loma Sal y Yeso was carried out to determine the fault structure of this section of the mountain front and to improve correlations between rocks exposed in the diapir and those penetrated by the Charco Largo- 1 well. Mapping shows three major lithologically distinctive fault-bounded lithologic units in the central part of the Loma de Sal y Yeso that form a composite section at least 350 m thick (Fig. 19A). Deformation associated with folding and diapirism on Loma de Sal y Yeso appears to have eradicated the primary fabrics, sedimentary structures, and most of the bedding characteristics of both the halite and gypsum. This makes interpretation of evaporative depositional environments difficult.
Unit 1 This unit is composed of a minimum of 200 m of the massive halite of the Angostura Formation that occurs as an elongate strip along the northeastern edge of the hill (Fig. 19A). The halite exhibits a recrystallized fabric composed of interlocking, irregular halite crystals 1-2 cm across. Fresh exposures of halite reveal gray to black flow bands in the halite composed of thin layers of concentrated organic debris. This banding probably formed by salt flowage from depth as suggested by seismic line 111 that is adjacent to the Loma Sal y Yeso (Fig. 9). Flow bands vary from 1 to 3 mm in thickness, and can be traced over distances of several meters across single outcrops. These bands commonly exhibit small-scale open folds, and dip either vertically or steeply to the south (Fig. 19C). Locally, the massive flow-banded halite contains intervals of interbedded gray mudstone. We interpret these rocks as shallow-water halite deposits in an enclosed basin subject to periodic influxes of terrigenous rocks.
Fig. 18. (A) Geologic map of Loma Sal y Yeso in the area of the open-pit salt mine and measured section in Arroyo del Pozo. Key to numbered horizons: 1 - recrystallized, flow-banded halite and overlying shale of the Angostura Formation; 2 = recrystallized massive to thinly bedded gypsum and interbedded shale of the Angostura Formation; 3 -- finely laminated gypsum and unfossiliferous, massive, gray silty mudstone overthrust by halite of the Angostura Formation; 4 = 'Razorback Ridge', a resistant bed of shallow-marine grainstone taken as the base of the Las Salinas Formation; 5 -- ripple-marked sandstone and interbedded siltstone; 6 = trough-cross bedded sandstone and interbedded siltstone; 7 = coquina beds interbedded in siltstone; 8 -- growth position coral reefs and reworked coral debris interbedded in siltstone; 9 = recrystallized micrite and calcarenite interbedded in siltstone. (B) Aerial photograph is Marena series (1983), roll 43, flight 21, no. 212 (original scale: 1:40,000).
Z
o
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN
Unit 2 This unit makes up most of the hill and is composed of a minimum of 350 m of massive and bedded gypsum and interbedded siltstone of the Angostura Formation (Fig. 19A). The bedding of the gypsum of unit 2 is locally visible both on 1 : 40,000 scale aerial photographs and in the field (Fig. 18A). Packages of gypsum up to 20 m in thickness are more resistant and preferentially weather out of thinner, poorly exposed siltstone beds. Individual gypsum beds appear to range from 1 cm to 2 0 - 3 0 cm in thickness. The beds are always weathered to nodular alabastrine gypsum as defined by Schreiber et al. (1976) which are composed of centimeterscale masses of white gypsum surrounded by gray shrinkage cracks, and fine-grained gray gypsum and anhydrite (?), and organic inclusions. We interpret these rocks as shallow-water gypsum deposits in an enclosed basin subject to periodic influxes of terrigenous rocks. Unit 3 This unit is a 30-m-thick section of interbedded, gray, finely laminated mudstone ('paper shale'), and interbedded siltstone, and thin (1-cm-thick) gypsum beds (Fig. 19A). The paper shale section is faulted against the halite, which itself contains thin interbeds of a similar paper shale. The paper shales exhibit very fine-grained, light and dark laminations which form parting surfaces. Although extensively sampled in previous studies (e.g., McLaughlin et al., 1991), unit 3 is barren of microfossils. We interpret the rocks of unit 3 as anoxic lacustrine deposits. Gypsum is the most resistant unit and forms the highest elevations of the hill (Fig. 19A). Where unmined, the halite forms a low depression along the crest of the hill. Where mined along the northeastern edge of the hill, large sinkholes caused by salt dissolution have formed at the lowest points of the excavated areas (Fig. 19B). The basinward edge of the salt area is delineated by a 3-m-thick vertical bed of highly resistant limestone that marks the base of the Las Salinas Formation ('Razorback Ridge', or unit 4 on Fig. 19A). The flat area northeast of Loma Sal y Yeso is underlain by outcrops of the Las Salinas Formation (Figs. 18 and 19A).
311
Structure of Loma Sal y Yeso Mapping in the gypsum unit indicates a synform along the top of Loma de Sal y Yeso (Fig. 19A). The contact between the halite of unit 1 and the locally horizontally dipping gypsum of unit 2 is a fault that is either vertical or dipping steeply to the south where exposed (Fig. 19A). The syncline may be a local drag effect caused by reverse motion along this fault contact. In other areas to the east, gypsum of unit 2 appears to have an anticlinal structure at the crest of the hill (Llinas, 1972a,b). A large-scale, transverse right-lateral fault visible on aerial photographs offsets the hill to the southeast of the open-pit salt mine (Fig. 18). Mann et al. (1991a) named this fault the Las Salinas fault zone after the town in the fault valley. At the outcrop scale, a well exposed 1-2-m-wide shear zone separates the unit 1 halite from the interbedded siltstone and gypsum of unit 3 (Fig. 19A). This shear zone strikes parallel to the trend of Loma Sal y Yeso and dips vertically or steeply to the southwest. Small-scale folds within the shear zone are asymmetric and indicate up-to-the-southwest reverse motion across the zone (Fig. 19C). Flow banding indicative of ductile deformation is present within the sheared halite and in brown mudstone interbeds adjacent to the shear zone. Bedding in unit 3 and the overlying strata of the basal Las Salinas Formation is vertical or slightly overturned to the northeast. Bedding dips in the Las Salinas Formation gradually decrease in a basinward direction (Fig. 19A). No other major faults were mapped in the monoclinally dipping section of the Las Salinas Formation north of Loma Sal y Yeso (Fig. 19A). For this reason, these rocks would project downdip into the Charco Largo-1 well (Fig. 13 and Table 1).
Economic significance of salt and gypsum mined from Loma Sal y Yeso The salt and gypsum mined in the Loma Sal y Yeso has been exploited for the past 30 years and is of local economic importance. Halite mined at the quarry has a 96 to 98% purity. The mechanized government salt mine, which formerly exported salt
Fig. 19. (A) Cross-section of the Loma Sal y Yeso in the area of the open-pit salt mine. Tadpoles indicate measured dips of bedding. Numbered horizons are the same as Fig. 18. The downdip extension of the halite and thrust faults is speculative but consistent with the geometry of halite seen on seismic line 111 north of Loma Sal y Yeso (compare to line in Fig. 10). (B) View to the northwest along the open-pit salt mine (see approximate location in A). The deep valley is formed by extraction of halite by mining and by solution of halite by rainwater (note natural sinkholes formed in the bottom of the mined valley). The width of the halite deposit corresponds to the approximately 50-m-wide valley. The outcrop to the left is almost flat-lying recrystallized gypsum; the steep ridge to the right corresponds to a steep reverse fault which thrusts halite on massive silty mudstone cropping out on the vegetated hill to the right (compare this view to the cross-section shown in A). (C) View to northwest of recrystallized halite along the edge of the open-pit salt quarry. Note hammer for scale in the center of the photo; small-scale open folding indicates thrusting with southwest-side up. Note flow banding within sheared halite, and brown mudstone interbeds within halite to left of photo. Color photographs at http://www.elsevier.nl/locate/caribas/
312
E MANN et al.
to the northeastern U.S. for use on winter streets, was closed in 1963. Subsequent salt mining reoccupied the former quarry but uses only manual labor. The salt is broken into gravel-sized pieces, filled in burlap bags, and loaded by hand onto trucks. The salt is used for domestic consumption. The gypsum quarry on the southeastern side of Loma Sal y Yeso is being actively mined by Cooperaci6n Dominicana de Empresas Estatales (CORDE) using fully mechanized methods. Most of the gypsum is exported by ship through the port of Barahona to Santo Domingo, other Caribbean countries, and northern South America for use as a hardening agent in cement. Four different commercial grades of gypsum are recognized by CORDE geologists on the basis of colors ranging from white to dark brown or black. The color variations reflect the amount of siliciclastic impurities within the gypsum.
Stratigraphy and sedimentology of the Las Salinas Formation in the Arroyo del Pozo The three evaporitic units of the Loma Sal y Yeso are overlain by a 1820-m-thick section (Las Salinas Formation) of monoclinally basinward-dipping sandstone, shale, and conglomerate of the Las Salinas Formation. These strata form a relatively flat area along the southern edge of the Enriquillo basin northeast of the Lome Sal y Yeso (Fig. 19A). Five subdivisions of the formation (units 4-9) are based on distinct assemblages of lithologies and sedimentary structures. The distribution of lithologic units 1-9 are shown on the map of Fig. 18 and crosssection of Fig. 19A. Fig. 20 provides a key for all lithologies and sedimentary structures of the five units (units 4-9) we identified from the Las Salinas Formation (Fig. 21). Fig. 22 provides an expanded, more detailed section of the lower, well exposed 27 m of unit 4 and the base of unit 5 at the head of the Arroyo del Pozo. Because of long covered intervals between most units, the boundaries and thicknesses of each unit are approximate. Approximately 50% of sections 5-9 of the Las Salinas Formation consist of easily erodible gray to bluish siltstone that is covered or poorly exposed in most places (Fig. 21). Limestone, sandstone, and conglomerate form the most resistant beds which are exposed in outcrops along the Arroyo del Pozo and along narrow, heavily vegetated strike ridges extending to the east and west of the stream (Fig. 18A).
Unit 4 The base of this unit and of the Las Salinas Formation is marked by the 3-m-thick, resistant bed of shallow-marine grainstone and coquina which forms 'Razorback Ridge' (unit 4 in Fig. 22A).
Fig. 20. Key for lithology and sedimentary structures for Fig. 21 and all subsequent measured sections in this paper. The limestone bed appears to conformably overlie approximately 50 m of gray massive siltstone and 'paper shale' with interbeds of gypsum (unit 3 of Loma Sal y Yeso). The limestone bed is made up of multiple beds of medium- to coarse-grained calcarenite 5-10 cm thick, with pervasive larger shell material. This includes numerous intact and broken bivalves and gastropods (Fig. 22E). Bedding units are defined by alternations in grain size of coarse- and fine-grained shell hash. The limestone of unit 4 is interpreted as a basal transgressive deposit marking a marine or brackish water incursion over the Angostura evaporite basin of the central Enriquillo basin. We correlate this bed with the hard, buff-white, 4.6-m-thick limestone penetrated at 2625 m in the Charco Largo-1 well at the transition between the Angostura and Las Salinas Formations (Fig. 15 and Table 1).
Unit 5 This unit is a ~ 180-m-thick interval of symmetrical, wave-ripple-marked sandstone and interbedded siltstone recording low-energy, nearshore conditions
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN in a shallow-marine and/or lacustrine environment (Fig. 21). The section consists mainly of 0.51.0-m-thick packages of gray to greenish, friable quartz-rich litharenite in beds ranging from 1-2 cm up to 25-30 cm (Fig. 21). The sandstones are fine- and medium-grained, with quartz and lithic fragments showing a moderately well-rounded character. These sandstone packages alternate with 1.0 to 0.25-1.0-m-thick horizons of bluish to greenish siltstone and silty mudstone. The sandstone-rich intervals in the lower part of the section shown in Fig. 22A show packages arranged into two thinningand fining-upward cycles over intervals of 4 to 9 m. Bedding contacts are planar. Sand size is fine- and medium-grained. Symmetrical wave ripple-marks are found on most bedding planes in the sandstone-rich packages (Fig. 22D). The vertical or slightly overturned attitude of bedding at the base of the section allows excellent views of extensive ripple marks on the upper surfaces of the bedding planes. Sand is concentrated in the troughs of the ripples marks. Paleocurrent measurements on 24 ripple marks restored to horizontal indicate a paleowave direction of ENE-WSW (Fig. 22A). This is the approximate present-day direction of prevailing winds which blow westward from the Caribbean Sea into the Enriquillo Valley. Vertical worm burrows (Skolithos?) characteristic of intertidal or beach environments are found in several of the ripple-marked sandstone beds (Fig. 22C). Interbedded within the sandstone beds near the base of unit 5 is a resistant, ripple-marked bed of dense, fractured, buff-white chert, 1 m thick, with oil seeps and small-scale soft-sediment 'diapirs' protruding through the bed into overlying beds (Fig. 22B). The 'diapirs' are 10 to 15 cm in height and are rooted in an underlying bed of the same composition. This rock is similar to descriptions of 'Magadi-type chert' of Jurassic, Eocene, and Pleistocene age in Wyoming (Surdam et al., 1972), of Paleozoic age in Scotland (Parnell, 1988), and Cambrian age in Australia (White and Youngs, 1980). Magadi-type chert is precipitated from alkaline lakes which store large amounts of silica in solution. Conversion to chert involves loss of sodium and water, and a volume loss of at least 25% that causes soft-sediment deformation. These sediments form in stratified bodies of standing water, under conditions conducive to the preservation of organic matter. As a result, they commonly exhibit oil seeps and are associated with petroleum source rocks. The section becomes more sand-rich with thicker (1 m) cross-bedded sandstone beds 25 m above the base of the Arroyo del Pozo section (Fig. 22A). A single coquina bed composed of broken and whole marine shells is present at 21.5 m above the base of the section and may indicate an increasingly
313
open-marine influence (Fig. 22A). Where exposed, the upper 100 m of unit 5 consists of bluish silty mudstone. The top of unit 5 is defined by the base of a sand-rich section of trough cross-bedded sandstone at 180 m (Fig. 21). Unit 6 This unit consists of 390 m of trough cross-bedded sandstone, ripple-marked, channelized sandstone and interbedded siltstone that record higher-energy, deeper-water shallow-marine conditions than unit 5. Unit 6 is distinguishable from overlying and underlying units by its thick sandstone beds and its lack of marine shell horizons. Sandstones form packets ranging in thickness from 2 to 10 m. These sandstone packages are slightly more resistant than intervening siltstone beds and form low ridges which control the direction of Arroyo del Pozo (Fig. 18A). Each sandstone package consists of a thinning- and finingupward cycle marked by a thicker, trough cross-bedded sandstone at the base overlain by thinner-bedded, ripple-marked sandstone beds, passing upwards into siltstone and silty mudstone at the top of the cycle. The basal sandstone beds generally exhibit pronounced basal scour or channelization into the siltstones of the upper part of the underlying cycle. Individual sandstone beds in the packets range in thickness from 10 to 50 cm. Sandstone is brown to greenish in color, and is composed of friable, fine- to medium-grained, moderately well-rounded litharenite. Cement is a dusty-appearing matrix of clay minerals derived from the breakdown of unstable rock fragments. The thicker sandstone beds are planar to broadly lens shaped, especially those with prominent trough cross-bedding. The sandstone beds appear continuous for tens of meters in the field and on aerial photographs (Fig. 18A) and are assumed to be equally extensive in the downdip direction (Fig. 18B). Paleocurrents were measured from oblique streambed exposures of 24 trough cross-beds in unit 6 using the method of DeCelles et al. (1983) (Fig. 21). This method allows more precise estimates of the paleocurrent direction for two dimensional outcrop exposures of trough cross-beds. The paleocurrent direction is E N E - W S W and is consistent with the paleocurrent of tipple marks in unit 5. Unit 7 This unit consists of a largely covered 100-mthick section of siltstone with two 20-cm-thick beds of marine shell fragments identified (Fig. 21). The interval is distinguishable from unit 6 by its lack of thick sandstone beds. Unit 8 Unit 8 consists of a 225-m-thick, coral-rich limestone and siliciclastic section. The base is marked
314
R M A N N et al.
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN by section beginning with a l 1-m-thick coral reef in growth position. The coral reef contains delicate branches of coral, clearly in growth position. The reef unit is overlain by 212 m of bluish siltstone with a few 10-20-cm-thick beds of litharenite.
Unit 9 This is the highest unit measured in the Arroyo del Pozo section and is largely covered. Limited exposures suggest that it seems to consist mainly of bluish siltstone with infrequent, 5-20-cm-thick beds of litharenite. The unit is distinguishable from unit 8 by its lack of coral beds. The unit also includes 0.21.3-m-thick beds of recrystallized micrite and calcarenite containing marine mollusc-shell fragments. These resistant beds form strike ridges which can be followed for several kilometers on aerial photographs (Fig. 18A). Above this unit, Quaternary alluvium of the Enriquillo Valley onlaps the rocks at the top of the section. Facies interpretation of Las Salinas Formation in the Arroyo del Pozo The progression of facies through this section reflects an overall increase in marine influence. An initial marine transgression recorded by the highenergy limestone of unit 4 ends the hypersaline restricted environment recorded by the evaporitic facies of Loma Sal y Yeso; this transgression terminated evaporitic deposition throughout the Enriquillo basin. High-energy wave-dominated shoreline deposition marked by well sorted, rippled sandstones at the base of unit 5 pass upwards into a more restricted, possibly alkaline and lacustrine environment marked by the inferred Magadi-type chert. This restriction of the basin may have been short-lived and was superseded by more open-marine conditions as supported by the reappearance of marine macrofossils 10 m above the chert (Fig. 22). A deeper-water marine environment is marked by the alternating sandstone ridges and intervening shaley horizons dominates sections 5 and 6. We interpret this facies as a shallow-marine, fluvial-deltaic setting. The lower part of the section with numerous sandstone ridges is interpreted as a coastal deltaic complex characterized by thick, sandstone bodies showing thinning-upward profiles (Fig. 21). Upward in the section in units 7, 8, and 9, coral becomes more prominent and suggests the presence of a siliciclastic-dominated shelf setting with scattered coral reefs (Fig. 21). The absence of growth position corals in
315
unit 9 suggests that marine influence and water depth may have varied through time as it has done in the Holocene (Taylor et al., 1985).
Correlation between Angostura and Las Salinas Formations exposed in Loma de Sal y Yeso and in the Charco Largo-1 well Broad similarities in the nine distinct lithologic units exposed in outcrops near Loma Sal y Yeso and in lithologies described in the well log and reports from the Charco Largo-1 well allow correlations to be made between the two sections (Table 1) (Fig. 21). A significantly greater thickness of Angostura Formation evaporites occurs in the Charco Largo-1 well (1562 m, Fig. 15) than in the Loma de Sal y Yeso outcrops (350 m, Fig. 19A). Evaporitic units 1, 2, and 3 are probably thinner than evaporitic units in the Charco Largo-1 well because of local structural complexities and salt flowage effects as seen on seismic line 111 (Fig. 9) and as seen in flow textures of evaporites cropping out on Loma Sal y Yeso (Fig. 19C). For example, the 200 m of massive halite of unit 1 contrasts with 1539 m of interbedded halite and shale in the Charco Largo-1 well. Unit 1 may be significantly thinner than the basin-central deposits because it represents a thin diapiric 'offshoot' of the parental evaporite body as suggested on seismic line 111 in Fig. 9. Massive gypsum of the type and thickness (350 m) seen in Loma Sal y Yeso (unit 2) is not present in the Charco Largo-1 well. The evaporite section in the well consists entirely of halite. It is possible that the basin-center facies were more saline than more basin-edge areas like Loma Sal y Yeso. The 50-m-thick massive gray mudstone and gypsum horizon (unit 3) that is faulted against the halite along the northern edge of Loma Sal y Yeso (Fig. 19A) may correlate to a 23-m-thick shale and gypsum interval in the well that separates the top of the halite at 2652 m from the base of a 4.6-m-thick, dense, white limestone bed at 2630 m in the well. The 3-m-thick shallow-marine limestone of 'Razorback Ridge' (unit 4), which marks the base of the Las Salinas Formation in the Arroyo del Pozo, may correlate to a 4.6-m-thick bed of limestone in the well (Table 1). Both units mark the end of significant halite formation in the Enriquillo basin. The 150-m-thick interval of ripple-marked and burrowed sandstone and interbedded siltstone at the
Fig. 21. Measured section from the Las Salinas Formation in the Arroyo del Pozo north of Loma Sal y Yeso (see Figs. 18A and 19A for locations). Base of section (unit 4) and top of section (unit 9) are shown on the map in Fig. 18A. A detailed measured section of the lower part of the section is shown in Fig. 15. Paleocurrents are measured from trough cross-beds. See Fig. 20 for explanation of lithology and sedimentary structures.
316
E MANN et al.
Fig. 22. (A) Detailed measured section of the lower 28 m of the Las Salinas Formation in the Arroyo del Pozo section (see Figs. 18A and 19A for locations). Paleocurrents are from ripple marks. See Fig. 20 for explanation of lithology and sedimentary structures. Locations of photographs that follow are indicated by their letter to left of column. (B) Magadi-type chert bed interbedded with dark siltstone near the base of the Las Salinas Formation (bed is vertical with top to the right or northeast). Note two dark stains (at arrows) which are oil seeps derived from the chert bed. Oil seeps and soft-sediment deformation features are typical of this type of chert. (C) Vertical burrow of Skolithos (?) with structureless fill. This bed is both overlain and underlain by beds with symmetrical ripple marks. (D) Symmetrical wave ripple marks with continuous crestlines on a vertical bed at the base of shallow-marine beds of the Las Salinas Formation 2 m above the limestone of 'Razorback Ridge' in Arroyo del Pozo. Note that the troughs of the ripples are filled with sand and the round exit hole of Skolithos (?) in right-center of photo. (E) Resistant grainstone forming 'Razorback Ridge', a highly resistant bed forming the approximate boundary between the evaporitic beds of the Angostura Formation and shallow-marine beds of the Las Salinas Formation (see Figs. 18A and 19A for locations). Color photographs at http://www.elsevier.nl/locate/caribas/
base of the lower part of the Las Salinas Formation in outcrop (unit 5) may correlate to a 30-m-thick interval of shale capped by a 6-m-thick bed of anhydrite at 2595 m in the Charco Largo-1 well. In their study of the Charco Largo-1 cuttings, McLaughlin and Bold identified shallow-marine microfauna from unit 4. The anhydrite may mark a drying cycle in the basin at the close of the marine interval of marked by unit 4. The 390-m-thick section of trough-cross bedded sandstone and interbedded siltstone of unit 6 in the
lower half of the Las Salinas Formation correlates to a 4387-m-thick section of 8-60-m-thick beds of sandstone interbedded in shale between 2157 and 2595 m in the well. Some of the thicker sandstone beds may form the prominent reflectors visible in the Las Salinas Formation on seismic line 111 across the area near the well (Fig. 9). Correlations between the upper part of the Las Salinas Formation in outcrop and in the section above the sandstone interval in the Charco Largo-1
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN well become more tenuous because of the basinward thickening of units, as is evident from the seismic data, and loss of exposures at the northern end of the Arroyo del Pozo outcrop. In general the upper part of the Las Salinas Formation (units 7, 8, 9) fines upward in the Arroyo del Pozo while the Las Salinas Formation above 1220 m depth in the well coarsens upward. It is possible that a 3-m-thick limestone bed reported in the well at a depth of 2157 m may correlate to one of the two 1-m-thick coquina beds in unit 7. Because coral clasts are not distinguished on the well log, it is difficult to attempt correlations between unit 8 and the well.
SURFACE GEOLOGY OF THE NORTHEASTERN MARGIN OF THE ENRIQUILLO BASIN Stratigraphy
The northeastern margin of the Enriquillo basin in the valley of the Rio Yaque del Sur consists of northto northwest-dipping beds of four sedimentary rock formations: the Trinchera, Quita Coraza, Arroyo Blanco, and Arroyo Seco (Fig. 23). The formations crop out on the southern limb of the Los Guiros syncline, a major northeast to east-west-trending syncline separating the anticlinal Sierra Martin Garcfa to the south from the anticlinal Sierra de Neiba to the north (Mann et al., 1991a) (Fig. 6). The four formations have been previously correlated by McLaughlin et al. (1991) with outcrops in the Loma Sal y Yeso and in the Azua and San Juan basins (Fig. 7). The facies and micropaleontology of outcrops of the Trinchera and Quita Coraza Formations in the valley of the Rio Yaque del Sur have been studied in detail by McLaughlin et al. (1991). This study concluded that: (1) the two formations represent an Upper Miocene-Lower Pliocene marine sequence which shallows upward from a middle bathyal to inner neritic environment; (2) coarsergrained facies in the Trinchera Formation represent a southwest-prograding submarine fan derived from the erosion of Upper Cretaceous-Eocene island-arc rocks of the Cordillera Central; and (3) finer-grained facies in the Quita Coraza Formation represent shelf and upper slope deposits. Sedimentary facies in the overlying shallowmarine and non-marine rocks of the Arroyo Blanco and Arroyo Seco Formations have only been mapped in a reconnaissance manner by Cooper (1983) and form the basis of the detailed description that follows. Studies of the sedimentary structures, sandstone petrography and conglomerate clast counts of these two formations have been carried out 50 to 100 km to the north in the San Juan basin by Harms (1989). The main results of the Harms (1989) study
317
are consistent with the study by McLaughlin et al. (1991): (1) Upper Miocene-Lower Pliocene sedimentary rocks record a marine-continental transition that is slightly older than the marine-continental transition seen further south in the valley of the Rio Yaque del Sur; and (2) sand and gravel compositions record erosion of Upper Cretaceous-Eocene island-arc rocks of the Cordillera Central. Structure
Resistant beds of Upper Miocene sandstone and conglomerate of the upper Trinchera Formation form the eastern edge of the meander belt of the Rio Yaque del Sur (Figs. 6, 23). Rocks of the Pliocene Quita Coraza Formation occupy the meander belt of the river because they are easily erodible and lack resistant sandstone or conglomerate interbeds. Resistant beds of calcarenite and sandstone beds of the lower Arroyo Blanco Formation form the northwestern edge of the meander belt (Fig. 23). Dips on beds in all formations vary between 20 ~ and 45 ~ (Mann et al., 1991a,c). Several minor high-angle faults cut all three formations but none have significant offset (Mann et al., 1991 a). Planar erosion of the Arroyo Blanco Formation in Quaternary time by the meandering Rio Yaque del Sur has produced excellent exposures of strike ridges formed by resistant gypsum, sandstone, and conglomerates (Fig. 23). In a few areas, graveland cobble-covered river terraces less than 5 m in thickness cover the strike ridges. Because the rocks of the Quita Coraza and Arroyo Blanco Formations are tilted north and northeast, they present a natural outcrop section across the Early Pliocene northeastern shallow-marine margin of the Enriquillo basin (Fig. 23). Detailed interpretations of largescale trends in strike ridges can be integrated with surficial mapping and measured sections to reconstruct the geometry of the margin. Less resistant lithologies which form strike valleys separating the resistant ridges are poorly exposed or are covered. Facies associations of the Arroyo Blanco and Arroyo Seco Formations
For the purpose of this study, sedimentary rocks of the Arroyo Blanco Formation and the overlying Arroyo Seco Formation are grouped into four facies associations shown on the map in Fig. 23A: (1) a deep-marine facies association; (2) a shallow-marine facies association characterized by the presence of gypsum beds; (3) a shallow-marine facies association characterized by the presence of oolitic limestone; and (4) a non-marine facies association. The first three facies associations alternate in the 700-930-m-thick Arroyo Blanco Formation
Fig. 23A. Schematic map of major groups of sedimentary facies along the eastern margin of the Enriquillo basin based on aerial photograph interpretation and field mapping. Key to numbered facies groups" unit 1 -- massive marine siltstone of Lower Pliocene Quita Coraza Formation; units 2 through 8 = shallow-marine oolite and gypsum facies associations of the Lower Pliocene Arroyo Blanco Formation; unit 9 = non-marine facies association of the Upper Pliocene-Pleistocene (?) Arroyo Seco Formation. Letters identify resistant ridges of gypsum, oolitic limestone, or calcarenite-coral debris flows that are used to correlate the measured sections which are identified by name.
OO
Fig. 23B. Example of aerial photograph used as a base map for the map shown in (A) (photograph number 721, ICM-I(58)-13-V14 12-29 December 1958, original scale: 1 "60,000).
Z
',D
Z
>
u~
0
~Z
ZZ
:Z
O
O
> -]
> Z
O
O C~
320 while the fourth association composes the entire 1000-m-thick Arroyo Seco Formation (Fig. 23).
Water depths The use of the terms 'shallow-marine' and 'deepmarine' are relative. There are few constraints on the absolute water depth in which each of the four facies associations shown on the map in Fig. 23A was deposited. McLaughlin et al. (1991) determined that benthic foraminifera identified in samples taken from the base to the top of the Quita Coraza Formation in the map area shown on Fig. 23 indicate a transition from an outer to inner neritic setting. Benthic foraminifera in the lower part of the formation suggest depths greater than 100 m; those in the upper part suggest depths less than 100 m. Microfauna are rare in the Arroyo Blanco Formation and estimates of water depth must rely on analogies to modem environments of deposition for oolites and gypsum. The main field criteria for deep-marine facies association are the presence of marine fossils, gravity-driven sedimentary structures, such as graded bedding, and the absence of gypsum and oolitic limestone.
E MANN et al. The prominent change in strike of bedding from northeast to northwest near the Enriquillo basin is related to bending about a vertical axis associated with the formation of the Los Guiros syncline (Mann et al., 1991a) (Fig. 6). The Quita Coraza Formation (unit 1) and oolite and gypsum shallow-marine facies associations (units 2, 4, 5, 7, and 8) thin towards the northeast. The lower deep-marine facies association (unit 3) thickens to the northeast. The wedge geometry of the upper deep-marine facies association (unit 6) is unclear because of the thinness of the unit and poor exposure. The lower deep-marine facies association (unit 3) contains resistant calcarenite and coral debris flows which form prominent southwest-dipping clinoforms (Fig. 23A). The clinoforms prograde upward in the section from northeast to southwest. To the west, the coral debris flows and calcarenite horizons lose their clinoform shape and become planar. The contacts between facies associations appear parallel-sided in all but one case. On detailed aerial photographs, the eastern part of unit 6 clearly truncates the eastern part of unit 5.
Measured sections in the Arroyo Blanco Formation
Subdivision of the Arroyo Blanco Formation In order to facilitate description and interpretation, the Arroyo Blanco Formation in the map area shown in Fig. 23 is subdivided into six units which are numbered 2-7 on Fig. 23A (unit 1 is the marine Quita Coraza Formation and unit 8 is the non-marine Arroyo Seco Formation). Units 3 and 6 are deep-marine facies associations characterized by the presence of marine fossils, units 2, 5, and 7 are shallow-marine facies associations characterized by the presence of gypsum, and units 4 and 7 are shallow-marine facies associations characterized by the presence of oolitic limestone. The units alternate in three shoaling-upwards cycles: (1) deep-marine facies association of the Quita Coraza Formation (unit 1) - - shallow-marine gypsum facies association (unit 2); (2) deep-marine facies association (unit 3) shallow-marine oolite facies association (unit 4) shallow-marine gypsum facies association (unit 5); (3) deep-marine facies association (unit 6) shallow-marine oolite facies association (unit 7) shallow-marine gypsum facies association (unit 8) non-marine facies association (unit 9, Arroyo Seco Formation).
Stratal geometry Mapping and interpretation of aerial photographs indicates that units 1-8 are wedge-shaped (Fig. 23).
In order to document the nature of the seven units of the Arroyo Blanco Formation, five measured sections were measured over an along-strike distance of 20 km (locations of sections shown on Fig. 23A). Sections were measured using a Brunton compass mounted on a Jacob's staff. Strike and dip of bedding were frequently measured to insure that the true thickness of the section was measured. Where bedding was not exposed, regional strike was estimated using the strike of resistant ridges on aerial photographs. Most of the sections generally follow a major dry stream bed. The areas outside the stream beds are covered by a semiarid, dry forest-type of vegetation characterized by mesquite, low trees, and cactus. In the E1 Granado, Barrero, and Honduras-Facolina sections, vegetation or cover in the stream bed or parallelism of the stream bed to regional strike required that the section be measured out of the stream bed. All of the sections follow a single direction across strike, except for the Honduras section which is offset 1 km along-strike to the northeast to the Facolina section (Fig. 23A). The section was offset because Arroyo Honduras forms a heavily vegetated 'box canyon' with few outcrops. Well-used trails follow all of the dry stream beds and penetrate many of the densely vegetated areas between the streams. Correlation between stream sections is based on the identification of lithologically distinct and resistant beds of gypsum, calcarenite, and oolitic limestone which can be traced on aerial photographs
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN and walked out in the field (beds A - J labelled on Fig. 23A). Distinctive gypsum types in the sections are correlated in the chart on Fig. 24 over alongstrike distances of about 13 km. The tops of the gypsum ridges are not covered with soil or vegetation and form elevated walkways through otherwise heavily vegetated areas. In the lower deep-marine facies association (unit 3), three distinct and resistant beds of calcarenite named X1, X2, and X3 could be correlated over a distance of 5 km between the Barranca and Honduras-Facolina sections (Fig. 23A). In the shallow-marine gypsum and oolite facies associations, distinct and resistant beds of gypsum and oolite could be traced for 10 km from the E1 Granado section to the HondurasFacolina section where the units pinch out on the underlying marine section (unit 3) (Fig. 24).
El Granado section Location The E1Granado section is 143 m thick (Fig. 25A) and lies within unit 3 of the deep-marine facies association, unit 4 of the shallow-marine oolite facies association, and unit 5 of the shallow-marine gypsum facies association (Fig. 23A). The E1 Granado section can be reached by driving north from Santana on the road to E1 Granado. Because there is no named stream nearby, the section was named after E1 Granado, the nearest village.
321
Marker beds A and B Roadcuts at the base of the section expose 2 5-cm-thick beds of gray-brown mudstone with interbeds of 2-3-cm-thick beds of siltstone or very fine-grained sandstone and 2-cm-thick white clay layers. Coarsely crystalline gypsum layers up to 3 cm thick occur as fibrous, secondary overgrowths along bedding planes. The 35-m-thick section above the roadcut forms a grass-covered slope. We infer that this section forms the top of unit 3 of the deep-marine facies association (Fig. 23A). The first resistant units in the E1 Granado section are two marker beds (A and B in Fig. 23A) of oolitic limestone ranging from 2 to 3 m in thickness (Fig. 25B). The lower 60 cm of bed A consists of massive oolitic limestone which is buff-white on broken surfaces and weathers to a gray color. Parallel laminations are seen at the base of this unit. The overlying, 60-cm-thick unit consists of flaggy, 4-6-cm-thick beds of calcarenite with multiple sets of symmetrical, straight-crested ripples with heights of 5 mm and trough-to-trough lengths of 3 - 6 cm. The base of this unit is marked by a 2-cm-thick layer of marine shell hash. The outcrop is pockmarked with cavities, or oomolds, formed by the solution of oolites which are preferentially deposited in the troughs of the ripple marks (Fig. 25B). Close inspection reveals the presence of thin remnant coatings of undissolved and uncemented oolites within some of the cavities. The orientations of 30 ripple marks in both beds A and B suggest a N E - S W wave direction (Fig. 25A). The uppermost unit of bed A consists of a 140-cm-thick bed of massive, weathered fine-grained calcarenite
Fig. 24. Along-strike thickness and textural variations in single horizons of gypsum in the Rfo Yaque area of the eastern Enriquillo basin. Over a distance of 13 km, gypsum beds thin and pinch out to the northeast and thicken toward the center of the Enriquillo basin to the southwest (location of this area relative to basin center is shown on Fig. 13).
322
R M A N N et al.
Fig. 25. (A) E1 Granado measured section. See Fig. 20 for key to lithology and sedimentary structures. Vertical scale is in meters. Locations of photographs that follow are indicated to left of column. (B) Oolitic limestone 40 m above the base of the E1 Granado section at the E1 Granado-Santana road (marker bed A on map in Fig. 23A). Lenticular dark cavities in the center of the outcrop are formed by solution of oolites which preferentially fill troughs of cross-beds. Scale is 10 cm long. (C) Photomicrograph of oolitic limestone from horizon A in the E1 Granado section (outcrop is shown in A). The rock is an oomoldic sparite with few of the original ooids remaining. Note the increase in size of the spar cement from the ooid molds into the primary pore space. Dimensions of the field of view is approximately 4 mm by 6 mm. Color photographs at http://www.elsevier.nl/locate/caribas/
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN with horizontal burrows on the tops of some bedding planes. The rock is massive and has a mottled appearance which appears to be related to pervasive burrowing. Thin sections reveal that the massive parts of the rock are now supported by secondary dolomite cement which formed between the oolites prior to their dissolution (Fig. 25C). Limestone marker bed B is 10.2 m above the top of bed A and consists of 1 m of ripple-marked oolitic limestone (Fig. 25A). The interval between marker beds A and B is covered and assumed to be gray-brown litharenite of bed B 1, which is seen in outcrop and in float in other areas along strike.
Contact between marker beds B1 and C The first gypsum bed (bed C) occurs 28 m above the upper oolitic limestone of bed B (Fig. 25A). Sandstone of bed B1, which separates limestone of beds A and B from gypsum of bed C, is covered in the line of the measured section but is well exposed in a stream cut of the Arroyo de la Salvia near the crossing of the Santana-E1 Granado road 200 m to the west of the line of the measured section. A detailed section was measured at the contact on the road (Fig. 26A). Sandstone-gypsum contact in Arroyo de la Salvia The Arroyo de la Salvia outcrop can be divided into seven distinct lithologic units which are well exposed in the streamcut (Fig. 26A). The first two units are part of bed B while the upper five units are part of bed C. Marker bed B2, unit 1. This consists of parallel-sided, 5-20-cm-thick, poorly sorted, brown, calcite-cemented litharenite separated by interbeds of brownish-gray siltstone (Fig. 26A,B). The sandstone layers are intensely bioturbated and mixed with the over- and underlying siltstone. The burrows are stained red with hematite cement. Sandy beds thicken to 10-20 cm near the top of the unit and consist of 1-2-cm-thick beds of ledge-forming sandstone alternating with mudstone. Sandy beds are clearly truncated by a channel of the overlying unit 2. Sandy layers of unit 1 contain minor fine-grained fossil hash and some well-preserved, very small (--~2 mm) gastropod and pelecypod shells along with one internal mold of a gastropod. Marker bed B2, unit 2. This consists of a 2-m-thick channel which is 2.4 m deep and 5 m wide (Fig. 26B). The base of the channel is filled by 5-10 cm of a poorly sorted muddy grit with pebbles averaging 5 mm in diameter and up to 4 cm in maximum diameter. The lag is thickest at the base of the channel and thins to the walls of the channel. The upper part of the channel is filled with a brown fine- to medium-grained litharenite. The
323
channel is overlain by beds of fine-grained litharenite exhibiting hummocky cross-bedding (Fig. 26A). Marker bed C, unit 1. This consists of a 2-mthick section of gypsum and sandstone. The lower gypsum unit consists of three 4-10-cm-thick beds of primary, coarsely crystalline gypsum layers of the grass-type interbedded in gray mudstone containing well-preserved leaf fragments (Fig. 26E). Marker bed C, unit 2. This consists of a 1-mthick section of massive, spheroidally weathering gray mudstone with molds of gypsum crystals which resemble three-pronged stars. Marker bed C, unit 3. This consists of a 1-mthick ledge-forming bed of gypsum (Fig. 26A,B) composed of 2-3-cm-thick beds of gypsum growing on laminae of a slightly muddier gypsum crystal hash (average crystal length is ~ 1 cm) (Fig. 26C,D). The upper 50 cm of the bed contains at least 20 horizons of well-organized and continuous layers of grass-type gypsum (term of Schreiber et al., 1976) (Fig. 26C). Bed C1. Bed C1 consists of fine- to mediumgrained brown litharenite which is 11.9 m above the top of the gypsum of bed C. The sandstone forms a low cliff on the side of a ridge but does not form a prominent strike ridge similar to the gypsum and limestone marker beds B and C. The total exposed thickness is 5.1 m. The geometry of the entire outcrop suggest a 20-m-long and 5-m-high channel incised into covered rocks assumed to be mudstone or siltstone. The lower 3 m of bed C 1 consists of two fining-upward cycles composed of basal units of grit overlain by large trough cross-beds (20 cm high and 80 cm wide) in medium- to fine-grained litharenite which are, in turn, overlain by better-sorted, medium-grained litharenite with plane-parallel laminations. Measurements of the long axis of the trough cross-beds suggests N E - S W paleoflow. The top 2 m of the outcrop consists of medium- to fine-grained litharenite with large trough cross-beds. Marker bed D. This bed is 5 m above bed C 1 and consists of a 10-cm-thick bed of coarsely crystalline gypsum overlying fine-grained muddy sandstone. In views parallel to bedding, the gypsum consists of a mosaic of sutured stylolitic contacts between square crystals 3 • 3 cm in size. The gypsum is similar to facies 3 (massive selenite) described by Vai and Ricchi Lucchi (1977; their fig. 13) in the Messinian Vena del Gesso basin of the northern Apennines. A similar, 5-cm-thick bed is present 2 m above the lower gypsum horizon. This gypsum horizon is overlain by a 1-m-thick section of platy, black mudstone, or 'paper shales'. Marker bed E. This bed is a 4.91-m-thick gypsum bed 23 m above the top of the upper gypsum horizon of horizon D. The gypsum forms a prominent ridge and is composed of crystals ranging in length from 1 to 2 cm.
324
R M A N N et al.
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN
Interpretation of El Granado section The section is interpreted as a shoaling-upward cycle which begins with shallow-marine oolitic limestone (beds A, B, B1), passes upwards into a more restricted evaporitic interval (beds C, D, E) (Fig. 23A). The sandstone bed B2 marks the boundary between the shallow-marine oolite and gypsum association and probably reflects the regression responsible for this transition. The C 1 sandstone bed within the gypsum section probably records a shorter-lived transition from marine to evaporitic environments.
Arroyo Barrero section Location The Arroyo Barrero section is 750 m thick and crosses unit 4 of the shallow-marine oolite facies association, unit 5 of the shallow-marine gypsum facies association, unit 6 of the deep-marine facies association, unit 7 of the shallow-marine oolite facies association, unit 8 of the shallow-marine gypsum facies association, and unit 9 of the non-marine facies association (units 4 - 8 comprise the Arroyo Blanco Formation while unit 9 belongs to the Arroyo Seco Formation) (Fig. 23A). The Arroyo Barrero section can be reached by driving north on dirt access roads through sugar cane fields north of Santana. The section follows Arroyo Barrero for most of its length. The main features of the complete measured section shown in Lamar (1990) are summarized below and shown on the gypsum correlation chart in Fig. 24.
Description of the Arroyo Blanco Formation Base of the section. The base of the section is marked by a low ridge composed of a 10-m-thick section of reworked coral reef and shell rubble. The unit is well exposed in a cut made for an irrigation canal near the mouth of Arroyo Barrero. Bedding in the unit thickens upward from 6-8-cm-thick beds near the base to 2-m-thick beds in the upper part of the outcrop. The beds are made up of 9 0 - 9 5 % branching coral fragments which average 3-5 cm in length and 1 cm in diameter. No corals in growth position were observed, although several articulated bivalves were found. Calcite is present in the pore spaces of shells. The matrix is a poorly sorted,
325
muddy siltstone with finely broken, 1-3-mm-size shell hash. Paleocurrent measurements were made in two 1-m-thick debris flows where the long axis of branching coral clasts show a consistent orientation. The paleoflow was from northwest to southeast. The top of the coral debris horizon is gradationally overlain by 6 m of massive siltstone, containing scattered bivalves, some of which appear in place. Marker bed A. Bed A of unit 4 is 132 m above the top of the coral debris horizon and consists of 60-cm-thick beds of buff-white limestone. Bed A can be traced on air photos to the oolitic limestone outcrops of bed A in the E1 Granado section (Fig. 24). The limestone is flaggy, composed of 5 15-cm-thick beds, and has a gray, micritic matrix containing 1-5-mm-long gastropods and bivalves. Marker bed B. Bed B of unit 4 is 4.9 m above the top of bed A. The basal unit of consists of 40 cm of recrystallized, flaggy oolitic limestone. Ghosts of oomolds similar to those observed in the E1 Granado section are present (Fig. 25B). Individual limestone beds are 2 cm thick. Marker bed C. The limestone bed is directly overlain by 3.2 m of a poorly exposed and altered gypsum bed with remnant crystals up to 2 cm in length that is correlated to bed C. Most of the gypsum is altered to a fine, white powder. The gypsum is overlain by 2 m of mudstone which is in turn overlain by a 1-cm-thick limestone bed and a 10-cm-thick bed composed of gypsum crystals disseminated in mudstone. Solution of the crystals produces a 'worm-eaten' texture in the mudstone. Marker bed C1. Bed C1 of unit 5 consists of a poorly exposed layer of greenish-brown litharenite 25 m above bed C. The exact thickness of this layer is unknown because it crops out as loose float blocks of sandstone on a slope maintained by the overlying resistant gypsum bed D. The lithology of the float blocks is very similar to the bed C1 in the E1 Granado section (Fig. 25). None of the float blocks of sandstone is greater than 30 cm in thickness. Marker bed D. Bed D of unit 5 consists of a 3-m-thick bed of massive, altered selenite. Alteration to a white powder is most intense on the upper surface of the bed and along joints and fractures in the lower part of the bed. The selenite is composed of 2 x 2 cm to 3 x 3 cm square
Fig. 26. (A) Detailed measured section of the contact between marker horizon B 1 and horizon C of the outcrop shown in (B). See text for description of the seven lithologic units. Locations of photographs that follow are indicated to the left of column. (B) Photograph of outcrop showing contact between horizon B1 and horizon C at outcrop near the crossing of Arroyo de la Salvia and the E1 Granado-Santana road at approximately the same scale as the measured section of this outcrop in (A). (C) View looking perpendicular to bedding in grass-type gypsum from horizon C in the E1 Granado section (unit 5 on measured section in A). Top of bed is towards top of photograph. Note the even banks of crystals which can be traced across the outcrop. (D) View looking down on a bedding plane surface of grass-type gypsum from same horizon as the bed shown in (C). Note fine needles of gypsum suspended in mud matrix. Diameter of coin is 17 mm. (E) Well-preserved leaf fragment found in siltstone bed (unit 3). Leaf and other plant fragments are probably well-preserved because of high salinity of the brine pools in which they were deposited. Color photographs at http://www.elsevier.nl/locate/caribas/
326 crystals when viewed perpendicular to the bedding plane. Immediately overlying the top of bed D are at least two beds of poorly exposed limestone which form a small knob of the resistant gypsum ridge. The two limestone beds which are recrystallized and coarse-grained may have originally been shell hash. Marker bed E. Bed E of unit 5 is 72 m above the top of bed D and consists of a 110-cm-thick bed of massive, altered selenite composed of 1 x 1 cm crystals. The slope beneath bed E is covered but appears to be weathered, massive gray mudstone. Marker bed F. Bed F of unit 6 is 85 m above bed E and consists of a 1-m-thick interval of poorly exposed, flaggy oolitic limestone composed of 2 5-cm-thick beds. The base of the bed appears to be bioturbated. Most of the limestone is recrystallized but casts of oolites about 1 mm across are visible on unweathered surfaces of beds about 1.5 cm in thickness. Also present is a 2-cm-thick bed of micrite containing fossil hash about 2 mm in diameter. An unnamed bed above bed F consists of a 15-cm-thick bed of medium-grained brown litharenite overlain by a 30-cm-thick bed of silty micrite containing fossil hash. These two beds are not traceable on aerial photographs. Marker bed G1. Bed G of unit 7 consists of beds G1 and G2, and, together with bed H, forms a 3-mthick triplet of beds occurring in rapid succession (Fig. 23A). Bed G1 is 88 m above the top of bed F and consists of a 42-cm-thick marly coquina made up of four bedding units. The lower unit of bed G1 is 15-cm-thick and consists of a sandy mudstone containing marine shells. The unit is intensely bioturbated by 75-mm-wide round burrows. The overlying unit is 17 cm thick and contains marine shells and load casts up to 2 cm deep. This unit is overlain by a 2-cm-thick mudstone layer and an 8-cm-thick poorly sorted calcareous siltstone. The shell units appear to be rapidly deposited 'tempestites' or storm deposits which were later subjected to bioturbation. Marker bed G2. Bed G2 of unit 7 is 1 m above bed G1 and consists of a 70-cm-thick bed of very well sorted, brown litharenite. The bed is composed of two 20-30-cm-thick beds which are laminated in 2-cm-thick partings and contain possible trough cross-beds. Marker bed H. Bed H of unit 7 is 1 m above bed G2 and consists of a 60-cm-thick bed of weathered gypsum. The upper part of the bed contains 1-cmthick mud laminations which suggest that the bed was originally a grass-type gypsum. Marker bed I. Bed I of unit 8 is 85 m above bed H and consists of a weathered gypsum horizon about 1-2 m thick which is altered to a white powder. Some loose crystals up to 2-3 cm in diameter are present as float and suggest that bed I was originally a massive selenite.
E MANN et al.
Marker bed J. Bed J of unit 8 is 81 m above bed I and consists of a 2-m-thick of very coarsely crystalline massive selenite with crystals up to 3 x 3 cm in size. The crystals are disorganized and may be reworked as a 'gypsrudite'. Bed J is overlain by brown sandstone and mudstone. Bed J is only 100-200 m below non-marine rocks of the Arroyo Seco Formation (Fig. 23A). Interpretation of Arroyo Barrero section The section is interpreted as two shoaling-upward cycles (Fig. 23A). The base of the section is marked by a coral debris horizon of unit 3 which can be correlated on aerial photographs to resistant clinoforms of calcarenite and coral debris beds to the northeast. The flatness of the horizon in this area (Fig. 23A) suggests that it may have occupied a basinward position relative to the slope clinoforms. This deep-marine facies association of unit 3 shoals rapidly upward into a shallow-marine facies association defined by oolitic limestone in marker beds A and B of unit 4. As in the E1 Granado section, these limestones shoal upward into a shallow-marine gypsum facies association defined by the gypsum beds C, D, and E of unit 5. Marine limestone of bed D and marine siliciclastic rocks of bed C1 suggest that the gypsum association is actually composed of smaller cycles defined by: (1) marine regression (sandstone channels); (2) gypsum and associated mudstone; and (3) marine transgression (oolitic limestone with marine shells). However, the overall trend is shoaling-upwards. The shallow-marine gypsum association of unit 5 is overlain by a 'deep-marine' facies association defined by the dominance of shallow-marine oolitic limestones present in marker bed F of unit 6. The 'deep-marine' interpretation here is used in a relative sense to distinguish this predominantly marine limestone section from the underlying predominantly gypsum section of unit 5. Moreover, gypsum bed E of unit 5 is clearly truncated on aerial photographs by limestone bed F of unit 6 to the east of Arroyo Barrero (Fig. 23A). This is interpreted as a marine transgression of marine limestone bed F of unit 6 over the underlying gypsum of unit 5. The regression is marked by sandstone G2 and the overlying gypsum bed H which marks the base of the overlying shallow-marine gypsum facies association of unit 8 (Fig. 23A).
Arroyo Barranca section Location The Barranca section is 930 m thick and includes three deep-marine facies associations, three shallow-marine gypsum facies associations, and two shallow-marine oolite facies associations of the up-
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN
327
Fig. 27. Arroyo Barranca measured section. See Fig. 20 for key to lithology and sedimentary structures. Vertical scale is in meters. Locations of some of the photographs in Figs. 28 and 29 are indicated in the left margin of the section.
permost Quita Coraza Formation and the Arroyo Blanco Formation (Fig. 27). The base of the section begins at the top of the Quita Coraza Formation at the bluff along the west bank of the Rfo Yaque del Sur and can be reached by driving to the bridge across the Arroyo Barranca on the Tamayo-Vuelta Grande dirt road and walking southeast along the
Arroyo Barranca. The upper part of the section can be reached by walking along the Arroyo Barranca to the northwest.
Upper Quita Coraza Formation The base of the section consists of 15 m of turbiditic sandstone and minor conglomerate over-
328 lying the massive bluish siltstone which makes up most of the Quita Coraza Formation (Fig. 27). The section consists of massive to trough cross-bedded beds, 20 cm to 3 m thick, of fine- to mediumgrained litharenite, grit, and conglomerate. Clasts in conglomerate consist of volcanic, plutonic, and sedimentary rocks derived from Cretaceous-Eocene rocks of the Cordillera Central. Sandstone beds exhibit mm-scale laminations defined by variations in grain size and normal grading. The sandstone appears darker in color than the sandstones seen higher in the Arroyo B lanco and in the Las Salinas Formation to the south because of a much higher content of lithic grains (up to approximately 50% as estimated visually in the field).
Description of the Arroyo Blanco Formation Base of section. Marine rocks of the Quita Coraza Formation are overlain by the lowest of the shallowmarine units of the Arroyo Blanco, a 30-cm-thick alabastrine gypsum bed cropping out in a roadcut on the Tamayo-Vuelta Grande road. This outcrop is 160 m above the uppermost exposed beds of the Quita Coraza Formation (Fig. 27). A second gypsum bed crops out 90 m above the lower bed and consists of a 9-m-thick bed of grass- and spear-type selenite. These two gypsum beds and their associated deposits are significant because they mark the first shallow-marine conditions of unit 2 in the Neogene history of the Enriquillo basin (Fig. 23A). Both gypsum beds are seen to pinch out on aerial photographs and are not present in streams to the northeast of Arroyo Barranca (Fig. 23A). The deep-marine facies association of unit 3 overlying the gypsum association of unit 2 consists of poorly exposed massive siltstone, thick trough cross-bedded sandstone (Fig. 28A), calcarenite and conglomerate composed mainly of coral and shell fragments (Fig. 27). The sandstone consists of litharenite composed of fine- to medium-grained, well-rounded quartz, volcanic, and carbonate grains. Sandstone beds range from 20 to 50 cm and are interbedded with gritty siltstone. Trough cross-beds were measured in the more massive sandstone beds and indicate N E - S W paleoflow. There is one thick (8 m) conglomerate in the sandy section at 350 m. The conglomerate has a highly scoured, erosional lower contact. The matrix is a fine- to mediumgrained sandstone which supports clasts ranging in size from granules to boulders 20 cm to 30 cm in diameter. The coral debris at 475 m in unit 3 consists of a 3-m-thick bed composed of 90% coral-stick rubble (Fig. 29B). The coral does not appear as abraded as coral rubble in the Barrero section 4.9 km to the west (Fig. 23A). The deposit is clast-supported in a sandy mud matrix and occurs in beds 5-10
E MANN et al. cm in thickness which are similar to those seen at Barrero. Average coral size is 5-10 cm long and 2-3 mm thick. Cementation of the uppermost 1 m of the bed gives a coarsening-upward appearance to the outcrop. Coral fragments in this bed showed a preferred alignment, presumed to be the result of transport. Paleocurrent studies on this bed suggests north-to-south paleoflow. Marker beds X1, X2, and X3. Beds X1, X2, and X3 of unit 3 are flaggy, graded, well cemented grainstones which occur 50-90 m above the coral conglomerates described above. The three beds range in thickness from 40 to 60 cm but form resistant strike ridges which can be traced on aerial photographs from Arroyo Barranca for several kilometers to the east and west (Figs. 23A and 29A). Handsample examination of the grainstones shows bioclastic and limestone grains. Lower contacts may have been slightly scoured but show no extreme incision (Fig. 29A). Marker bed C. The overlying shallow-marine interval of unit 5 includes gypsum marker beds C and D (Figs. 23A and 24). Gypsum marker bed E of unit 5 from Arroyo Barrero is not present and is seen to be truncated by the limestone bed F of marine unit 6 on aerial photographs (Fig. 23A). The gypsum at marker bed C consists of 4.32 m of grass-type (Fig. 28B) and bladed selenite (Fig. 28C). Beds of the grass-type gypsum are typically 0.5-1 m thick and are found near the base of the outcrop. The vertically oriented crystals are 1-3 cm in length (Fig. 28B). Beds of bladed selenite overlie the grass-type gypsum, are 1-2 m thick, and are composed of stacks of smaller beds with the thickness of the vertically oriented crystals, 20-40 cm in length (Fig. 28C). These are the longest crystals of gypsum observed anywhere in the study area (Fig. 24). Above gypsum C is a 12-m-thick section of massive brown mudstone and siltstone (Fig. 27). Marker bed D. This unit consists of a 1-m-thick bed of jumbled, 1-3-cm-long gypsum crystals in a buff, muddy matrix which is 40 m above bed C (Fig. 27). As in Arroyo Barrero, gypsum D is abruptly overlain by a flaggy limestone (bed F). Marker bed E As in Arroyo Barrero to the east, the base of the overlying 'deeper-marine' interval of unit 6 is taken as the oolitic limestone of bed F as in the Arroyo Barrero (Fig. 23A). Gypsum bed E is truncated by limestone bed F between Arroyo Barrero and Arroyo Barranca. Beds F and G, in turn, are not found in streams to the northeast of Arroyo Barranca because the 'deeper-marine' interval onlaps the underlying shallow-marine interval and pinches out (Fig. 23A). This onlap is consistent with a thinning of the 'deeper-marine' interval between Barrero and Barranca. Bed F is a flaggy limestone about 3 m thick which shows oomolds in hand sample.
N E O G E N E EVAPORITIC AND S I L I C I C L A S T I C D E P O S I T I O N IN THE E N R I Q U I L L O BASIN
329
Fig. 28. Photographs of different siliciclastic and evaporitic facies at 380 m in the Arroyo Barranca measured section (see Fig. 27 for section and locations). (A) View of light-colored, resistant coral limestone conglomerate bed overlying less resistant, massive trough cross-bedded sandstone in marine unit 3 in the Arroyo Barranca section (compare with regional map of units in Fig. 23A). (B) Grass-type gypsum of marker bed C in evaporitic unit 4 in the Arroyo Barranca section. Top of bed is towards top of photograph and pencil is 14 cm long. (C) Spear-type bladed and twinned selenite of evaporitic unit 4 in the Arroyo Barranca section. Top of bed is towards top of photograph. Length of crystals represents water depth in modern settings of bladed selenite such as Lake MacLeod, Australia. (D) Bed of unsorted gypsrudite of evaporitic unit 8 in the Arroyo Barranca section suggesting nearby erosion of a coarse-grained selenite. Pencil for scale is 14 cm long. (E) Poorly sorted conglomerate composed of well rounded pebbles and cobbles and devoid of marine shells. This unit is taken as the base of the non-marine section (unit 9, Arroyo Seco Formation) in the Arroyo Barranca section. Pencil is 14 cm long. Color photographs at http://www.elsevier.nl/locate/caribas/
330
R MANN et al.
Marker bed G. Bed G is a 1.5-m-thick grainstone containing small intraclasts of broken shell material. Marker bed I. The highest gypsum bed (I) in the overlying shallow-marine interval of unit 8 contains a 15-cm-thick bed of gypsrudite (Fig. 28D) similar to that described by Schreiber et al. (1976) and Vai and Ricchi Lucchi (1977) in the Messinian of Sicily. This is the only gypsrudite seen in the area. Gypsum crystals are 2 cm to 5 cm long and show no preferred orientation within a fine matrix of gypsum powder. Description of the Arroyo Seco Formation The top of the Arroyo Barranca section is marked by non-marine conglomerates which are completely lacking in gypsum or marine fossils (Fig. 27). Clasts within the conglomerates are less than 5 cm in diameter with the majority only 1 cm to 2 cm in diameter (Fig. 28E). Clasts are supported by a fineto medium-grained sandy matrix and are primarily volcanic and metamorphic fragments.
Interpretation of Arroyo Barranca section The section is interpreted as the result of three shoaling-upwards cycles numbered 1-3 on Fig. 23A. The lowest cycle consists of marine sandstone, siltstone, and conglomerate of the upper Quita Coraza Formation (unit 1) which was deposited at inner neritic water depths (McLaughlin et al., 1991). This marine section shoals rapidly upward into the shallow-marine gypsum facies association marked by the two gypsum horizons near the Tamayo-Vuelta Grande road (unit 2 on Fig. 23A). The gypsum section is overlain by a deep-marine facies association of unit 3 marked by horizons of coral conglomerate, trough cross-bedded sandstone, and grainstone (beds X1, X2, and X3) (Fig. 23A). This section shallows upward into an unexposed but inferred shallowmarine oolite facies of unit 4 and an association of gypsum beds C and D of unit 5 is in turn overlain by the 'deep-marine' facies association of beds F and G of units 6 and 7 (Fig. 23A). These marine beds are overlain by the highest gypsum beds (bed I) of unit 8 which show some evidence for reworking (Fig. 28D).
Honduras-Facolina section Fig. 29. Photographs of different marine siliciclastic facies from several measured sections (see Fig. 27 for section and locations). (A) Turbiditic calcarenite of bed X1 in marine unit 3 in the Arroyo Barranca. (B) Coral reef material in debris flow of marine unit 3 in Arroyo Barranca section. In-situ coral reefs are present 12 km to the northeast (shelfward and along strike) of this outcrop (cf. Fig. 23A). Length of pen shown is about 7 cm. Color photographs at http://www.elsevier.nl/locate/caribas/
Location The Honduras-Facolina composite section is 850 m thick and includes a deep-marine unit and a shallow-marine gypsum unit near the area where the gypsum pinches out above the deep-marine unit (Fig. 23A). A complete measured section is presented in Lamar (1990).
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN
Description of the Arroyo Blanco Formation Base of the section. The base of the section begins near the bridge over the Arroyo Honduras on the dirt road from Tamayo to Vuelta Grande. The measured section is offset because of lack of exposure in the upper part of the Arroyo Honduras. The base of the section is marked by a resistant debris flow deposit of reworked coral and shell debris which is 3 m thick. The debris flow deposit exhibits local grading and is intensely bioturbated by Ophiomorpha. Matrix consists of silty to finegrained carbonate sand. Fossils include both coral and bivalve fragments. On aerial photographs, this bed and several others to the east form prominent southeast-dipping clinoforms which climb the section in an eastward direction and suggest westward progradation of carbonate-mantled slope deposits (Fig. 23A). The overlying marine rocks of unit 3 consist of turbiditic calcarenite beds similar to those in the Arroyo Barranca (Fig. 29A), coquinas, and trough cross-bedded sandstone beds. The section is three times thicker than the equivalent section in the Arroyo Barranca (Fig. 23A). A thick siliciclastic section begins 12 m above the grainstone beds near the base of the section and is composed of: (1) 15 m of bedded (beds a few cm thick) bioturbated mediumgrained sandstone; (2) fine-grained calcarenite 1 m thick; and (3) 80 m of massive, poorly exposed, fine- to medium-grained, dark quartz litharenite. Marker beds X1, X2, and X3. These three beds occur over an interval of 235 m of largely covered sandstone and siltstone of unit 3. The interval is much thicker than the correlative interval in the Arroyo Barranca to the west (Fig. 23A). The three beds are similar and consist of 1-1.25 m of turbiditic limestone composed of fine-grained carbonate grains and 1-2 mm fragments of shell hash and occasional 2-5 cm fragments of oyster shells. Bedding thickens from 5-10-cm-thick beds at the base of the bed to 50-cm-thick beds at the top of the bed. The three beds are interbedded with poorly exposed, massive, medium-grained litharenite in beds from 2 to 3 m in thickness which contain trough cross-beds. Coral debris beds and shell beds containing Turritella are interbedded in the sandstone. Marker bed C. Gypsum beds C and D of unit 5 occupy a condensed shallow-marine section which onlaps the underlying marine unit containing beds X1, X2, and X3 (Fig. 23A). Bed C consists of a 15-cm-thick bed of greenish, finely crystalline, grass-type gypsum which contrasts significantly to the thickly bedded, bladed selenite in Arroyo Barranca (Fig. 24). The abrupt change in crystal size may reflect the proximity of an influx of non-marine water (B.C. Schreiber, pers. commun., 1989).
331
Marker bed D. Bed D of unit 5 is a poorly exposed 30-cm-thick bed of recrystallized gypsum. The remainder of this section is covered until the 820 m level, where a thin recrystallized gypsum bed marks the end of subaqueous deposition and the beginning of deposition of non-marine sandstones and conglomerates of the Arroyo Seco Formation (Fig. 23A). Interpretation of Arroyo Honduras-Facolina section This section mainly (unit 3) consists of a thick deep-marine facies association composed of sandstone, coral debris beds, and grainstone. The upper part of the section shoals rapidly into a shallowmarine gypsum association of unit 5 marked by two gypsum beds (beds C and D) which can be seen onlapping the marine section on aerial photographs. The shallow-marine oolitic limestone beds of unit 4 appear to pinch out before reaching the Arroyo Facolina (Fig. 23A). Rio Yaque section Location The Rio Yaque section (Fig. 30A) is a shallow-marine section which is laterally equivalent to deeper-marine rocks of unit 3 described in the Arroyo Honduras-Facolina and Barranca sections (Fig. 23A). Resistant beds can be followed directly between the two areas on aerial photographs. The section crops out in the Arroyo Boca de los Guiros and can be easily reached by walking northward from the village of Sierracita just off the AzuaBarahona highway. Description of the Arroyo Blanco Formation Base of the section. The base of the section is a 10-m-thick interval of marine of siltstone and sandstone that represents the top of the Quita Coraza Formation. The 10-m-thick interval consists of 4-mthick debris flows of coral and gray siltstone fragments in a gray silty matrix; gritty, coarse-grained volcanic sandstone; and 10-30-cm-thick beds of fine-grained sandstone interbedded with blue siltstone (Fig. 30A). Growth position coral reef. The base of the Arroyo B lanco Formation is marked by a 10-m-thick coral reef composed of branching and head corals (Fig. 30A). The upper 1-2 m of the reef is tightly cemented and the reef forms a prominent, resistant ridge which can be traced eastwards on aerial photographs (Fig. 23A). The base of the reef is transitional with the underlying siltstone section of the Quita Coraza section (unit 1). One meter above the base of the reef is the first appearance of head corals in growth position and delicate, upright-branching corals (Fig. 30B).
332
P. MANN et al.
Fig. 30. (A) Rfo Yaque measured section. See Fig. 20 for key to lithology and sedimentary structures. Vertical scale is in meters. Locations of photograph (B) is indicated in the left margin of the section. (B) Growth-position reef forming base of marine unit 3 (compare to map in Fig. 28). Note upright, growth-position coral heads amongst reef rubble. Color photograph at http://www.elsevier.nl/locate/caribas/
Above the reef is a 20-m-thick section of 1050-cm-thick coquina and coral debris flows interbedded with brown silty mudstone. This reef and associated coral debris beds probably supplied the coral clasts and finer-grained carbonate material to grainstones including beds X1, X2, and X3 observed in measured sections to the southwest (Fig. 23A). Sandy, marine section. Above the coral-rich basal part of the section, there is a 290-m-thick interval of mainly litharenite and interbedded siltstone which are associated with intraclasts and layers of marine fossils (Fig. 30A). This interval is thought to represent the shelfward equivalent of more basinward deep-marine facies of unit 3 which were measured in sections to the southeast of this section. The sandstone occurs in characteristic 3-10-mthick cycles composed of: (1) a resistant 3-5-mthick basal massive bed of medium- to coarsegrained litharenite which contains pebble conglomerate in a coarse-grained sandstone matrix; the sand-
stone layers exhibit very low-angle to hummocky to trough cross-bedding; the sand-conglomerate unit tends to fine upwards and is typically scoured into the underlying thinly bedded siltstone and sandstone or exhibits load casts at its base; (2) a thin 5 10-cm-thick coquina bed commonly occurs at the top of the resistant conglomerate-sandstone unit; (3) the upper part of the cycle consists of 5-10 m of thinly bedded fine-grained sandstone and siltstone and is usually covered.
Sandy to conglomeratic, non-marine section. The upper 390 m of the section consists of alternating ridges of conglomerate-sandstone ranging from 3 to 10 m in thickness and intervening thinly bedded sandstone and siltstone which exhibits distinct red and green alterations in color (Fig. 30A). Grain size in the lower, coarser-grained beds increases upsection from medium-grained sandstone to conglomerate. There are no marine shells associated with this section and this observation, along with the distinct
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN red and green color of the intervening finer-grained layers suggests that the section is non-marine. Interpretation of the Rio Yaque section. The lower half of the section is interpreted as a shallowmarine siliciclastic section formed in a shelf setting. The presence of a growth position coral reef shown in Fig. 30A establishes the water depth at the base of the section at several meters below sea level. The presence of marine shells in 290 m of section above the reef indicates a prolonged marine influence. The lack of marine fossils and the red-green color in the upper part of the section suggests a non-marine environment of deposition.
Environment of deposition of the Arroyo Blanco Formation The environment of deposition of the Arroyo Blanco Formation consisted of a shelf edge on the northern margin of the basin which faced southwest towards the center of the Enriquillo basin (Fig. 31). The edge of the shelf is marked by the reef at the base of the Rfo Yaque section. The slope is marked by clinoforms in marine interval 3. The clinoforms flatten in a basinward direction as the grade of the slope decreases in deeper water of the Enriquillo basin. The clinoforms are composed of carbonate material which was derived from the reef and mantled the slope. Paleocurrents were mainly directed from northeast to southwest down the slope. The shallow-marine facies associations are confined to the basin and slope whereas the earlier deeper-marine facies extends across both the slope and shelf. As in the Messinian section of Sicily, the evaporite rocks are sandwiched between marine rocks with the higher evaporite beds overstepping the underlying evaporite units (Fig. 31). In contrast to the lower marine unit, which progrades basinward, the evaporite units prograde landward.
Textural variations in gypsum deposits of the Arroyo Blanco Formation Grass-type and bladed selenite in the gypsum layers and oolitic limestone suggest water depths of probably less than a few meters (Schreiber, 1988). In Fig. 24, we have compiled the measured thicknesses and primary textures of all gypsum horizons observed in the four measured sections. All of the gypsum layers probably formed in shallow brine pools on broad mud flats. It is possible that the length of the larger crystals in the selenite beds may reflect the water depth at the time of deposition as is observed in modem environments of deposition like Lake MacLeod in Australia (Logan, 1987; Schreiber, 1988). If this is the case, the deepest part of the tidal flat (<50 cm) during the deposition of bed C was in the Arroyo Barranca
333
section (Fig. 24). The length of these crystals and the thickness of this horizon rapidly decreased to the northeast and probably reflect the onlap of the gypsum horizon onto the marine rocks of the underlying section (Fig. 23A). Evaporite deposition was probably suddenly terminated by influxes or 'freshenings' of marine water as seen today at Lake MacLeod, Australia (Logan, 1987).
DISCUSSION Application of sequence stratigraphic terminology to the Enriquillo basin Sarg (1988) has proposed two settings for evaporites in the framework of sequence stratigraphy: (1) onlapping lowstand and shelf-margin wedges; and (2) lagoonal-sabkha facies in the interior or back-shelf positions of carbonate banks. The evaporites of the northeastern margin of the Enriquillo basin appear to fall into the first category because: (1) they are deposited basinward of the shelf break suggested by prograding clinoforms in the underlying deep-marine facies association (unit 3 in Fig. 31); and (2) the evaporite units exhibit wedge-shaped geometries which onlap underlying marine rocks in both map view (Fig. 23A) and in cross-section (Fig. 31). The onlapping geometry of the evaporites is similar to the Triassic Dolomita Principale peritidal complex of the Dolomites of Italy (Bosselini, 1984) and the Permian Salado Formation of west Texas (Sarg, 1988). In the terminology of sequence stratigraphy, the basal contact of the evaporite horizon would form a sequence boundary (for example, boundaries between units 1 and 2, units 4 and 5, and units 7 and 8 in Fig. 31). The deep-marine units would represent highstand deposits which prograded out onto a downlap surface. This surface is the boundary between units 1 and 3 (Fig. 31). The geometry and spacing of the evaporite intervals suggest three discrete drops in sea level affected the northeastern margin of the basin and presumably the basin-central evaporite deposit less than 30 km to the southwest (Fig. 13). The first two falls are followed by rises in sea level and transgression of the evaporites by a marine horizon (Fig. 31). This transgression produced an angular truncated contact between evaporite unit 5 and marine unit 6 (Fig. 23A). The toe-of-slope position of the evaporite units suggests that sea level dropped significantly to the base of the slope (~ 100-200 m) (Fig. 31).
Correlation to the Vail sea-level curve It is interesting to compare the timing of the three transgressive deep-marine units of the Quita
334
R MANN et al.
Fig. 31. Summary of measured sections and schematic cross-sections from the northeastern margin of the Enriquillo basin. Pattern of facies association is the same as shown in map view on Fig. 23. Letters identify key marker horizons discussed in text. Vertical scale is in kilometers. The locations of the five long sections described in the text are indicated by boxes. Location of shelf, slope, and basin are based on the geometry of clinoforms in the lower marine facies of the Arroyo Blanco Formation.
Coraza and Arroyo Blanco Formations to a eustatic sea-level curve (Haq et al., 1987) (Fig. 32). Generally, low sea levels in the Late Miocene coincide with coarsening-upward submarine fan rocks of the Trinchera Formation (McLaughlin et al., 1991). A highstand of sea level in the Early Pliocene coincides with the Quita Coraza Formation, which would be interpreted as a highstand condensed section. This highstand is also the same age as the halite of the central Enriquillo Valley, which may have been fed by marine waters spilling over a sill at the mouth of the valley (Fig. 13). A similar shaley unit is present .at the same time in the Cibao Valley of the northern Dominican Republic (Evans, 1986) (Fig. 3). In the Pliocene, there are two significant drops of sea level which may coincide with the regressive episodes in the Arroyo Blanco Formation. Imprecise dating in the Arroyo B lanco Formation, however, makes correlations difficult. An important question to answer is whether the basin was connected to the sea and was responsive to
eustatic sea-level fluctuations during the deposition of the Arroyo Blanco evaporites or whether the basin was isolated from the sea and was unresponsive to sea-level changes. Kendall (1989) has emphasized the point that all saline giants must be completely isolated from the sea (Fig. 2). In the Enriquillo basin, the correlation of evaporites to Early Pliocene highstands (Fig. 32) suggests that preiodic connections to the open ocean were achieved. Three distinct evaporite cycles at the basin edge (Fig. 31) suggest that these connections may have been shortlived, perhaps as a result of wide eustatic sea level fluctuations during the Pliocene.
Models for shallow-water evaporites There are two popular facies models for deposition of shallow-water evaporites. Hardie and Eugster (1971) proposed that deposition of coarsely crystalline selenite in the Messinian of Sicily occurred in the quiet waters of a shallow lagoon or gulf adjacent
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN
335
Fig. 32. Comparison of sea-level curve of Haq et al. (1987) to stratigraphy of the Enriquillo and Cibao basins of the Dominican Republic. Ages from the Enriquillo basin are from McLaughlin et al. (1991) and this study. Ages from the Cibao basin are from Evans (1986). to a littoral belt of laminated gypsum. Gypsum in the laminites and in associated gypsum sand bodies was derived from the area of selenite deposition and transported shorewards onto marginal evaporitic flats during storms. Vai and Ricchi Lucchi (1977) worked on a sequence in the Messinian of Italy that lacked gypsum laminites and suggested transport of gypsum toward the basin center. Gypsum was reworked from older, emergent beds of selenitic gypsum by ephemeral slope-controlled agents such as torrential streams and debris flows which built up shallow alluvial cones that encroached the basin. As Kendall (1984) points out, the models may be complementary and apply to a transgressiveregressive cycle: the first model applies to times of transgression and highstand or when the regression occurs entirely as a consequence of sediment outbuilding (stream gradients are low); the second model may apply to times of regression and lowstand when older evaporites are exposed and reworked in marginal areas (stream gradients are high). In the marginal evaporites of the Enriquillo basin, there is little evidence for reworking of older, emergent beds of selenitic gypsum. Only one isolated example was identified in the upper part of the Barranca section (Fig. 28D). On vertical scales of 10 to 20 m, a characteristic pattern shows a bed of channeled marine sandstone overlain by a gypsum
bed. This suggests that each gypsum bed represents a short regressive episode during a much longer shallow-marine episode. The stratigraphy is best explained by three shoaling-upward cycles composed of a lower, 'deep-marine' interval, an intermediate shallow-marine interval characterized by oolitic limestone, and an upper shallow-marine interval characterized by gypsum (cycles number 1-3 on Fig. 23A).
Relation of basin-margin evaporites to basin-central evaporites For reasons discussed at length by Kendall (1988), it is often difficult to relate basin-margin evaporites to basin-central evaporites. Based on limited dating of the basin-central evaporites, it appears that the basin-central evaporites (Angostura Formation) are coeval with the Quita Coraza or lowermost Arroyo B lanco Formation in the eastern Enriquillo basin (Fig. 7). The basin-central evaporites may correlate with the lowermost of the three shallowmarine intervals or, alternatively, they may correlate with all three of the intervals.
Subsidence analysis of the Charco Largo-1 well Backstripping, or removal of sediment load, and compaction effects is one approach that can be used
336
E MANN et al.
Fig. 33. Subsidence history of the Charco Largo-1 well. See text for discussion. to isolate the tectonic control on subsidence in the central part of the Enriquillo basin (Mella structural block in Fig. 8). A subsidence plot of the Charco Largo-1 well, based on integration of water depth information gained from this study along with backstripping, is shown in Fig. 33. Pulses of maximum tectonic subsidence occur in the latest Miocene, coeval with the deposition of the Trinchera and lower Angostura Formations, and in the mid-Pliocene to Recent time, coeval with the deposition of the upper Angostura, Arroyo B lanco, Las Salinas, Arroyo Seco and Jimanf Formations (Fig. 7). The Early Pliocene is characterized by a pulse of tectonic uplift of the basin. A simple interpretation involves an early period of latest Miocene thrusting and crustal loading related to the oblique collision of the Bahamas Platform with Hispaniola (Fig. 3). The uplift phase in the Early Pliocene may be related to fold-related topographic uplift in the area of the Mella salt-cored anticlines seen on seismic line 111 in Fig. 9. The subsequent mid-Pliocene through Recent subsidence is consistent with rapid siliciclastic sedimentation corresponding to the Las Salinas, Arroyo B lanco and Arroyo Seco Formations (Fig. 7). Burial of the faults bounding the Mella folds in the center of the basin is consistent with the end of tectonic uplift and the renewal of tectonic subsidence (Fig. 33). Comparison of Dominican evaporites to other areas Messinian of the Mediterranean Some of the best studied and most controversial evaporite deposits occur in Messinian (Late Miocene) sections of the Mediterranean. Evapor-
ites range from 0.2 to 2 km in thickness. They consist of largely shallow-water and 'desiccated' deposits interbedded with open-marine sedimentary rocks. Intermediate- and some deep-water evaporites are also recognized (Schreiber, 1988). Most of the gypsum in subaerially exposed basins of Messinian age in the Mediterranean has not been buried deep enough to convert the primary depositional fabrics to nodular anhydrite. The alternation of kilometerthick open-marine turbidites with Messinian gypsum without halite in the Apennines of northern Italy (Vai and Ricchi Lucchi, 1977) is similar to the northeast margin of the Enriquillo basin (Fig. 31). According to Schreiber (1988), seismic profiles in the Mediterranean show a similar 'stratigraphic sandwich' consisting of deep-marine marls passing into evaporites (with a thick halite section) and back into a deep-marine section. Studies of the Messinian in Sicily by Schreiber and Friedman (1976) and Schreiber et al. (1976) have shown that the lowermost part of the basinal evaporite section, which directly overlies deepmarine marls, begins with massive layers of crystalline gypsum. This succession is similar to gypsum occurrence along the margin of the Enriquillo basin both in its lateral continuity, primary textures, and in its intercalated shallow-marine carbonate rocks. According to Schreiber (1988) and Schreiber, pers. commun. (1989), continuous beds of bladed gypsum (or 'ferro-di-lancia' or 'spears of iron') in this lower unit formed in channels or sub-basins where there was active water movement that supplied a continuous ionic feed. These 'ferro-di-lancia' of Sicily look exactly like the long, bladed selenite (marker bed C of unit 5) of the Arroyo Barranca section (Fig. 28C). Schreiber further observes that the length of these
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN crystals (20 cm to 50 cm) probably reflects minimum water depths at the time of deposition (Fig. 24). In areas where greater amounts of freshwater entered the basin as shown by higher clay and sand content, the beds of gypsum crystals become irregular and pass into gypsum-cemented mud and sand layers with some thin stringers or lenses of larger, nearly vertically oriented crystals (2 cm to 5 cm high). The deposits from these water bodies have great lateral continuity and can be followed for kilometers. As in the Dominican Republic, the onset of evaporite deposition in many areas of the Mediterranean was preceded by the appearance of coral reefs, which were subsequently buried beneath the evaporites and associated siliciclastic sedimentary rocks (Warren, 1988; Schreiber, 1988). The lower level of crystalline gypsum of the Messinian of Sicily is overlain by: (1) halite and potassic salts; these contain halite crusts and desiccation polygons and probably formed in a shallow-water to dried-basin setting; and (2) an upper evaporite layer consisting of seven or eight cycles beginning with a brackish-water marl and passing upward into shallow-water gypsum which oversteps and onlaps the underlying evaporites. Although the upper layer resembles the basin-margin evaporites of the Dominican Republic, there is no three-part division of either the basin-margin or basin-central evaporite deposit. In the Charco Largo-1 well, the halite interval directly overlies marine siltstone and a single limestone bed and is terminated by a single anhydrite bed (Fig. 15). The marginal evaporites overlie marine siltstone but are not directly associated with halite (Fig. 31). The cycles in the basin-margin evaporite of the Dominican Republic exhibit the same pattern of onlapping the underlying evaporite cycles of the Messinian in Sicily.
Lac Assal, Djoubouti, Persian Gulf According to B.C. Schreiber (pers. commun., 1989), continuous beds of bladed gypsum as those observed in bed C of unit 5 (Arroyo Barranca section) in the basin-margin evaporites presently grow in evaporative water bodies which make up Lac Assal. The gypsum is fed by marine influx, the crystals are uniform within the bed, and pass laterally into muddy, finer gypsum beds where the basin is fed by freshwater. This lateral change in character is also seen as one passes along bed C from Arroyo Barranca to Arroyo Facolina (Fig. 24). Crystal size decreases dramatically while the muddy fraction increases. Delaware basin, Texas The Permian Castile Formation of the Delaware basin consists of 550 m of gypsum, dark limestone, and halite deposited in four depositional
337
cycles (Lowenstein, 1988). The Castile Formation overlies deep-marine turbiditic rocks and appears to have formed as a sharp drop in sea level affected the silled basin. As in the Quita Coraza Formation of the Enriquillo basin, marine fossil assemblages of the Permian marine turbidite units reflect increasing restriction of the basin. Varve layers in the evaporites, which may record annual changes in circulation, were deposited at a rate of 1.4 mm/year, can be traced great distances of 70-110 km, and have been used as an argument for deep-water deposition (cf. Kendall, 1984, for a review). The geometry and thickness of the Castile Formation suggests that it filled topographic depression of the Delaware basin of 600 m depth in a period of about 260,000 years. In this respect the Castile Formation may be similar to the basin-central evaporites of the Enriquillo basin which also appear to have filled a deep topographic depression in the center of a silled basin (Fig. 2).
Lake MacLeod, western Australia Lake MacLeod is a halite-filled salina that is similar in dimensions and sub-sea-level setting to the present-day Enriquillo Valley. The bed of the lake lies 2.8 m to 4.0 m below sea level and is fed by subsurface flow generated from the nearby Indian Ocean by evaporative drawdown from (Logan, 1987). The lake is isolated by tectonic folds which form a low sill along the coast. Much of the lake is filled by a massive gypsum bed 5-7 m thick, or halite layers. These evaporite deposits mark areas of brine ponding, whereas areas of inflow of marine water are marked by carbonate rocks. The dominance and preservation of thick halite in Lake MacLeod is a result of the extreme aridity as compared to more temperate, gypsum-dominated lakes in south Australia (Warren, 1988). The lack of substantial meteoric input during the history of the lake allowed the lake waters to remain in the halite precipitation field for an extended period of time. Likewise, formation and preservation of halite in the central part of the Enriquillo basin may result from the extreme aridity of this area and a lack of meteoric input. Dead Sea of Israel The Dead Sea of Israel is similar to the present-day slip-strike setting of the Enriquillo basin and contains modern basinal sediments which are similar to anoxic paper shales associated with halite of Loma Sal y Yeso (Fig. 19A). The water surface of the Dead Sea is 403 m BSL (as compared to 42 m BSL for Lago Enriquillo) and the MioceneHolocene sedimentary fill is 8-10 km thick (as compared to approximately 5 km for the Enriquillo basin). The modern lake sediments of the Dead Sea are laminated muds with dark detrital-rich and light aragonite-rich laminae similar to the paper shales of
338 Loma Sal y Yeso. The laminated bottom sediment contains deep-water accumulations of halite which make up less than 10% of the bottom sediment and occur as tiny clear cubes (Warren, 1986, 1988). The Dead Sea is often cited as a likely environment for the accumulation of laminar or bedded deep-water halite and may be analogous to halite accumulation in the central Enriquillo basin. In the Charco Largo-1 well (Fig. 15), much of the halite occurs in beds ranging in thickness up to several meters which are interbedded with unfossiliferous dark shale (Fig. 16). It is possible that the halite accumulated in a basinal environment similar to the Dead Sea.
Paleogeographic evolution of the Enriquillo basin Fig. 34 presents an interpretation of depositional events in the Enriquillo basin and its northeast margin which is slightly modified from McLaughlin et al. (1991) using the results of this paper. The Middle Miocene was a time of deep-water, normal-marine conditions recorded by the deposition of the pelagic limestone of the Sombrerito Formation (Fig. 34A,B). Generally lower stands of sea level in the early-Late Miocene through earliest Pliocene and tectonic uplift of the Cordillera Central coincided with submarine siliciclastic deposition of the Trinchera Formation (Fig. 34C-E). Rising sea level in the Early Pliocene (Fig. 32) flooded the area and ended siliciclastic deposition of the Trinchera Formation. This transgression is marked by siltstone of the Quita Coraza Formation, and seems to have been the flooding event with which the following cycle of flooding and desiccation of the central Enriquillo basin began. The late part of the Early Pliocene saw the deposition a massive halite section in the basin center during these series of marine influx and evaporation events (Fig. 34F). Later regressions are recorded along the northern margin of the Enriquillo basin. During the late-Early Pliocene, while gypsum and interbedded siliciclastics of the Angostura Formation were being deposited in the basin center, similar sediments of the Arroyo B lanco Formation were deposited on the northeast basin margin. In these formations, times of near-normal marine conditions often saw influx of sediments from the northeast.
Relevance of silled basin models to the Enriquillo basin Kendall (1984) presents three models for silled basins which may be the depositional settings of subaqueous evaporites found in 'saline giants' (Fig. 2). Could the Enriquillo basin be considered a late Neogene analog for one of these conceptual models?
E MANN et al.
Deep water-deep basin model In the 'deep water-deep basin' model, a deep restricted basin is filled with seawater passing over a shallow sill (Fig. 2, upper panel). The water in the basin attains a level of hypersaline equilibrium, perhaps through stratification or brine reflux, and evaporites are deposited at either a brine-air or brine-basin floor contact. Many authors have evoked this model to explain 'saline giants' such as the Permian Castile Formation of west Texas, the Jurassic Louann Salt of the Gulf of Mexico, and the Permian Zechstein of northwest Europe (cf. Kendall, 1984, for review). Intercalated black shales, lateral bedding continuity, and the suggestion of great topographic relief within basins have been taken as evidence supporting this model. A major weakness of this model is the difficulty of maintaining a deep-water body at or near halite saturation. Water saturated with salts is hygroscopic and will dilute itself by attracting moisture from an atmosphere with a relative humidity of 65% or greater (Schreiber, 1988). In deep water with low ratios of surface area to volume, it is extremely difficult to reach and maintain halite saturation. Shallow water-shallow basin model Models for 'shallow water-shallow basin' settings involve influx of marine water across a sill into a rapidly subsiding basin filled with sediment (Fig. 2, middle panel). This concept is obviously not applicable to evaporites such as those of the western Mediterranean which clearly formed in a deep basinal setting (Ryan and Cita, 1978).
Shallow water-deep basin model The 'shallow water-deep basin' model was invoked by Hsti et al. (1973) to explain the Messinian evaporites of the Mediterranean (Fig. 2, bottom panel). A shallow-water environment facilitates rapid evaporation per unit volume of water, more efficient solar heating of the water, and a quicker return to evaporite saturation levels after freshwater influx or climate change. The deep setting of the basin promotes periodic spills of seawater during highstands of sea level which replenish brines within the basin. Deep basins below sea level such as the modem Enriquillo Valley and the Dead Sea are fairly common in tectonically active areas and are likely settings for seawater spills.
Preferred model for Enriquillo evaporites The 'shallow water-deep basin' model (Fig. 2, bottom panel) is the preferred model for explaining both the basin central and basin edge evaporites of the Enriquillo basin for the following reasons: (1) the present-day 'ramp basin' structure of the basin characterized by a subsea-level depression about 80
NEOGENE EVAPORITIC AND SILICICLASTIC DEPOSITION IN THE ENRIQUILLO BASIN
339
Fig. 34. Schematic diagram of paleoenvironments in Miocene and Pliocene time of the Enriquillo basin based on McLaughlin et al. (1991) and this study. (A) middle Middle Miocene; (B) late Middle Miocene; (C) early Late Miocene; (D) earliest Pliocene; (F) late Early Pliocene; (G) middle Pliocene; (H) Present.
m BSL and isolated from the sea by a shallow coastal sill (Fig. 6) approximates the basinal conditions suggested by the shallow water-deep basin model; subsidence history of the Charco Largo-1 well (Fig. 33) and regional stratigraphic studies (Fig. 34) indicate that these 'ramp basin' conditions
also existed during the Pliocene evaporite deposition; (2) a shallow water setting would promote more efficient evaporation and rapid formation of 1500 m thick basin evaporites documented by the Charco Largo-1 well (Fig. 15); and (3) a 'deep' basin tectonically depressed by thrust faults and isolated
340 from the sea by a coastal sill would provide a likely site for multiple seawater spills that would replenish brines and sustain the formation of the types of thick basin central halite (Fig. 16) and the three cycles of basin margin evaporites observed (Fig. 31).
CONCLUSIONS (1) There are two distinct facies of Early Pliocene evaporites which formed in the Enriquillo basin: (a) a basin-central evaporite deposit composed mainly of halite (Angostura Formation), approximately 1500 m thick, which is known from the Superior Charco Largo-1 well in the center of the basin and from the diapiric anticline of Loma Sal y Yeso along the southern edge of the basin; (b) a basin-margin evaporite composed mainly of gypsum up to 5 m thick interbedded with shallow-marine siliciclastic and carbonate rocks (Arroyo Blanco Formation) (Fig. 13). (2) Seismic reflection and well data show that the Early Pliocene depositional basin was a major, fault-controlled deep in the center of the valley separated from the Caribbean Sea to the east by a shallow sill (Figs. 13, 14). Siliciclastic source areas lay to the north in the Cordillera Central. The depositional and climatic setting was probably similar to the modem Enriquillo Valley, which is 80 m BSL at its lowest point, separated by a shallow sill from the Caribbean Sea, and receiving siliciclastic sediment derived from the Cordillera Central by the Rio Yaque del Sur (Fig. 5). (3) The structure of the Loma Sal y Yeso consists of a faulted diapir. Halite of the Angostura Formation is extruded as a narrow, 100-m-wide strip along a reverse fault separating the Angostura and Las Salinas Formations (Fig. 19). (4) Primary textures in halite and gypsum exposed in Loma Sal y Yeso are overprinted by alteration and tectonic flowage related to diapirism and folding (Fig. 19). Interbedded sedimentary rocks are finegrained mudstone lacking sedimentary structures. (5) The Las Salinas Formation consists of a lower shallow-marine siliciclastic interval 180 m in thickness which is overlain by a brackish interval 1670 m in thickness containing both siliciclastic and minor carbonate rocks (Fig. 21). The lower shallow-marine interval terminated deposition of the basin-central halite of the Angostura Formation. (6) Correlations between distinctive lithologies in the Angostura and Las Salinas Formations can be made to the lithologies described by geologists of Superior Oil Company in the Charco Largo-1 well (Table 1). A notable difference is the lack of gypsum in the well and the presence of gypsum in the Loma Sal y Yeso.
E MANN et al. (7) The northeast Early Pliocene margin of the Enriquillo basin consists of a southeast-facing shelfslope which was tilted to the north and northeast 2040 ~ in post-Early Pliocene times by tectonic folding. The tilted beds now form a natural cross-section of the Early Pliocene margin (Fig. 23). (8) The sedimentary rocks of the shelf-slope can be divided into gypsum and oolite-bearing shallowmarine facies and a deeper-marine facies (Figs. 23, 31). Nine distinct lithologic groups are recognized including overlying non-marine rocks of the Arroyo Seco Formation. The primary textures of gypsum suggests that these were deposited in several meters or less of water in a tidal-flat setting (Fig. 24). (9) The nine lithologic groups can be interpreted as three shoaling-upwards cycles consisting of: (a) a basal deep-marine section characterized by basinward-prograding clinoforms composed of grainstones and coral debris; (b) an intermediate oolitic limestone section; (c) an upper gypsum section (Fig. 31). (10) The shallow-marine gypsum and oolitic limestone sections are wedge-shaped and occupy a position at the base of the slope (Fig. 23). Interbedded marine units onlap and truncate the underlying gypsum units. The three shallow-marine sections are interpreted as being deposited during three drops in sea level during the Early Pliocene. These sea-level drops correlate reasonably well to the Cibao basin of the northern Dominican Republic (Evans, 1986) (Fig. 32). (11) Correlation between the basin-margin and basin-central evaporites in the Enriquillo basin is difficult because of lack of accuracy in the age of the units. Despite this limitation, we favor a 'shallow water-deep basin' model (Fig. 2) for both areas of evaporites for three reasons: (1) the 'ramp basin' setting of the basin would promote a 'deep' sub-sea level depression in a coastal setting; (2) shallow water within this 'deep' basin would promote efficient evaporation of seawater spilled over the coastal sill; and (3) multiple spills during times of higher Pliocene sea level would replenish brines to sustain the formation of thick basin central evaporites and explain the three cycles of basin margin evaporites observed
ACKNOWLEDGEMENTS
This work represents a distillation of a 15-yearlong effort by the authors. Initial field studies were conducted as part of Mann and McLaughlin's Ph.D. dissertations in the 1980s and were followed up by a 1990 masters study by Mike Lamar that was carfled out at the University of Texas at Austin. Our 1980s work followed up on ostracode and strati-
N E O G E N E EVAPORITIC AND S I L I C I C L A S T I C D E P O S I T I O N IN THE E N R I Q U I L L O BASIN
graphic studies of Bold in the 1960s and 1970s which attempted to integrate outcrop data with early exploration wells. Lawrence worked on the subsurface data as part of the Norconsult group in the early 1980s (Norconsult, 1983). Mann, McLaughlin, and Bold were individually supported by NSF, the Petroleum Research Fund and Mobil Exploration and Production Company. We thank Mobil for providing Mann and Lamar with funds for field work in the Dominican Republic and for providing the subsurface data. We also thank Mobil for allowing McLaughlin and Bold access to the core of the Charco Largo-1 well in their core storage facility in Dallas, and for granting us permission to publish this combined data set. We also thank the Direcci6n General de Minerfa in Santo Domingo for their continued cooperation and permission to publish these data. The careful reviews of Christoph Heubeck, Steve Pierce, and Charlotte Schreiber substantially improved this paper. UTIG contribution number 1424.
REFERENCES
Berggren, W.A., Kent, D.V. and Aubry, M.-E, 1995. A revised Cenozoic geochronology and chronostratigraphy. In: W.A. Berggren, D.V. Kent, et al. (Editors), Geochronology, Time Scales, and Global Stratigraphic Correlation. Soc. Econ. Paleontol. Mineral. Spec. Publ., 54:129-212. Bold, W. van den, 1975. Neogene biostratigraphy (Ostracoda) of southern Hispaniola. Bull. Am. Paleontol., 66, No. 286, 39 pp. Bold, W.A. van den, 1990. Late Holocene ostracoda in and around Lake Enriquillo, Dominican Republic. In: D.K. Larue and G. Draper (Editors), Transactions 12th Caribbean Geological Conference, St. Croix, U.S. Virgin Islands. Miami Geological Society, Miami, FL, pp. 163-173. Bosselini, A., 1984. Progradation geometries of carbonate platforms: examples from the Triassic of the Dolomites, northern Italy. Sedimentology, 31: 1-24. Bowin, C.O., 1975. The geology of Hispaniola. In: A.E.M. Nairn and G.G. Stehli (Editors), The Ocean Basins and Margins: The Gulf of Mexico and the Caribbean. The Ocean Basins and Margins, Vol. 3, Plenum, New York, pp. 501-552. Burke, K. and Seng6r, A.M.C., 1988. Ten meter global sea-level change associated with South Atlantic Aptian salt deposition. Mar. Geol., 83:309-312. Cooper, J.C., 1983. Geology of the Fondo Negro Region, Dominican Republic. M.S. Thesis, State University of New York, Albany, 128 pp. DeCelles, EG., Langford, R.E and Schwartz, R.K., 1983. Two new methods of paleocurrent determination from trough crossstratification. J. Sediment. Petrol., 53: 629-642. De Leon, R., 1983. Aspectos geologicos e hidrogeologicos de la region suroeste. Publicationes Especiales, Museo de Historia Natural, Santo Domingo, 25 pp. DeMets, C., Gordon, R.G., Argus, D.E and Stein, S., 1990. Current plate motions. Geophys. J. Int., 101: 425-478. Dixon, T., Farina, F., DeMets, C., Jansma, E, Mann, E and Calais, E., 1998. Relative motion between the Caribbean and North American plates and related plate boundary zone deformation based on a decade of GPS measurements. J. Geophys. Res., 103:15,157-15,182.
341
Dolan, J.E, Mann, E, de Zoeten, R., Heubeck, C., Shiroma, J. and Monechi, S., 1991. Sedimentologic, stratigraphic, and tectonic synthesis of Eocene-Miocene sedimentary basins, Hispaniola and Puerto Rico. In: E Mann, G. Draper and J. Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary Zone in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 217-263. Dolan, J.E, Mullins, H. and Wald, D., 1998. Active tectonics of the north-central Caribbean region: oblique collision, strain partitioning, and opposing slabs. In: J. Dolan and E Mann (Editors), Active Strike-Slip and Collisional Tectonics of the Northern Caribbean Plate Boundary Zone. Geol. Soc. Am. Spec. Pap., 326: 1-61. Evans, C.C., 1986. Facies Evolution in a Neogene Transpressional Basin: Cibao Valley, Dominican Republic. Ph.D. Thesis, University of Miami, Miami, FL, 103 pp. Friedman, G.M., 1982. Evaporites as source rock for petroleum. In: Depositional and Diagenetic Spectra of Evaporites; a Core Work-Shop. SEPM Core Workshop 3, Calgary, pp. 385-395. Garcfa, S., 1976. Geograffa Dominicana. Amigo del Hogar, Santo Domingo, 272 pp. Guerra Pena, F., 1956. Las principales cuencas sedimentarias de la Republica Dominicana y sus posibilidades petroliferas. Symposium sobre Yacimientos de Petroleo y Gas, Tomo IV, XX Congreso Geologico International, Mexico City, pp. 141159. Haq, B.U., Hardenbol, J. and Vail, ER., 1987. Chronology of fluctuating sea levels since the Triassic. Science, 235:11561167. Hardie, L.A. and Eugster, H.E, 1971. The depositional environment of marine evaporites: a case for shallow, siliciclastic accumulation. Sedimentology, 16:187-220. Harms, F.J., 1989. Konglomeratisches Jungterti~ir im Valle de San Juan (stidwestliche Dominikanische Republik, Grosse Antillen). Zusammensetzung, Herkunft, und Alter der Ger611e, Fazies. Ph.D. Thesis, University of Hannover, Hannover, 141 PP. Heubeck, C., Mann, E, Dolan, J. and Monechi, S., 1991. Diachronous uplift and recycling of sedimentary basins during Cenozoic tectonic transpression, northeastern Caribbean plate margin. Sediment. Geol., 70: 1-32. HsiJ, K.J., Cita, M.B. and Ryan, W.B.F., 1973. The origin of the Mediterranean Evaporites. Init. Rep. DSDP, 13:1203-1231. Kendall, A.C., 1984. Evaporites. In: R.G. Walker (Editor), Facies Models (2nd ed.). Geosci. Canada Repr. Ser., 1: 259-296. Kendall, A.C., 1988. Aspects of evaporite basin stratigraphy. In: B.C. Schreiber (Editor), Evaporites and Hydrocarbons. Columbia University Press, New York, pp. 11-65. Kendall, A.C., 1989. Brine mixing in the Middle Devonian of western Canada and its possible significance to regional dolomitization. Sediment. Geol., 64: 271-285. Kirkland, D.W. and Evans, R., 1981. Source-rock potential of evaporite environment. Am. Assoc. Pet. Geol. Bull., 65: 81190. Lamar, M.E., 1990. Geology of Pliocene Evaporitic Rocks, Enriquillo Valley, Dominican Republic. Unpublished MA Thesis, University of Texas at Austin, 162 pp. Llinas, R.A., 1972a. Geologia del area Polo-Duverg6, Cuenca de Enriquillo, Codia, Part 1. Publication of Colegio Dominicano de Ingenieros, Arquitectosa, y Agrimensores, Santo Domingo, No. 31, pp. 55-65. Llinas, R.A., 1972b. Geologia del area Polo-Duverg6, Cuenca de Enriquillo, Codia, Part 2. Publication of Colegio Dominicano de Ingenieros, Arquitectosa, y Agrimensores, Santo Domingo, No. 32, pp. 40-53. Logan, B.W., 1987. The MacLeod evaporite basin, western Australia. Am. Assoc. Pet. Geol. Mem., 44, 140 pp. Lowenstein, T.K., 1988. Origin of depositional cycles in a Per-
342 mian 'saline giant': the Salado (McNutt zone) evaporites of New Mexico and Texas. Geol. Soc. Am. Bull., 100: 592-608. Mann, P., Taylor, E W., Burke, K. and Kulstad, R., 1984. Subaerially exposed Holocene coral reef, Enriquillo Valley, Dominican Republic. Geol. Soc. Am. Bull., 95: 1084-1092. Mann, P., McLaughlin, P.P., Jr. and Cooper, J.C., 1991a. Geology of the Enriquillo-Azua basins, Dominican Republic, 2. Structure and tectonics. In: P. Mann, G. Draper and J.E Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 367-389. Mann, P., Draper, G. and Lewis, J.E, 1991b. An overview of the geologic and tectonic development of Hispaniola. In: P. Mann, G. Draper and J.E Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 128. Mann, P., Lebr6n, M., Rodriguez, J. and Heubeck, C., 1991c. Geologic maps of the southern Dominican Republic. In: P. Mann, G. Draper and J.E Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary in Hispaniola. Geol. Soc. Am. Spec. Pap., 262, Plates 4a, 4b, and 4c, scale: 1 : 150,000. Mann, P., Taylor, E W., Edwards, R.L. and Ku, T.L., 1995. Actively evolving microplate formation by oblique collision and sideways motion along strike-slip faults: an example from the northern Caribbean plate margin. Tectonophysics, 246: 169. McLaughlin, P.P., Jr., 1989. Neogene planktonic foraminiferal biostratigraphy of the southwestern Dominican Republic. J. Foraminiferal Res., 19:294-310. McLaughlin, P.P., Jr., 1991. Migration of Neogene marine environments, southwestern Dominican Republic. Geology, 19: 222-225. McLaughlin, P.P., Jr. and Sen Gupta, B.K., 1994. Benthic foraminiferal record in the Miocene-Pliocene sequence of the Azua basin, Dominican Republic. J. Foraminiferal Res., 24: 75-109. McLaughlin, P.P., Jr., Bold, W.A. and Mann, P., 1991. Geology of the Enriquillo-Azua basins, Dominican Republic, 2. Foraminiferal/ostracode biostratigraphy and depositional history. In: P. Mann, G. Draper and J.E Lewis (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary in Hispaniola. Geol. Soc. Am. Spec. Pap., 262: 337-366. Norconsult, 1983. Dominican Republic Petroleum Exploration Appraisal (Vol. 1), Norconsult Internal Report, Sandvika, Norway, 79 pp. Parnell, J., 1988. Significance of lacustrine cherts for the environment of source-rock deposition in the Orcadian basin, Scotland. In: A.J. Fleet, K. Kelts and M.R. Talbot (Editors), Lacustrine Petroleum Source Rocks. Geol. Soc. London Spec. Publ., 40:205-217. Ryan, W.B.E and Cita, M.B., 1978. The nature and distribution of Messinian erosional surfaces - - indicators of a several-
E M A N N et al. kilometer-deep Mediterranean in the Miocene. Mar. Geol., 27: 193-230. Sarg, J.E, 1988. Carbonate sequence stratigraphy. In: C.K. Wilgus, B.S. Hastings, C.G.St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. Van Wagoner (Editors), Sea-Level Changes: An Integrated Approach. Soc. Econ. Paleontol. Mineral. Spec. Publ., 42: 155-181. Schreiber, B.C., 1988. Subaqueous evaporite deposition. In: B.C. Schreiber (Editor), Evaporites and Hydrocarbons. Columbia University Press, New York, pp. 182-255. Schreiber, B.C. and Friedman, G.M., 1976. Depositional environments of upper Miocene (Messinian) evaporites of Sicily as determined from analysis of intercalated carbonates. Sedimentology, 23: 255-270. Schreiber, B.C., Friedman, G.M., Decima, A. and Schreiber, E., 1976. Depositional environments of Upper Miocene (Messinian) evaporite deposits of the Sicilian Basin. Sedimentology, 23: 729-760. Sen Gupta, B.K. and Schaefer, C.T., 1973. Holocene benthonic foraminifera in leeward bays of St. Lucia, West Indies. Micropaleontology, 19: 341-365. Surdam, R.C., Eugster, H.P. and Mariner, R.H., 1972. Magadi-type chert in Jurassic and Eocene to Pleistocene rocks, Wyoming. Geol. Soc. Am. Bull., 83: 2261-2266. Taylor, E W., Mann, P., Valastro, S., Jr. and Burke, K., 1985. Stratigraphy and radiocarbon chronology of a subaerially exposed Holocene coral reef, Dominican Republic. J. Geol., 93: 311-332. Vai, G.B. and Ricchi Lucchi, E, 1977. Algal crusts, autochthonous and siliciclastic gypsum in a cannibalistic evaporite basin: A case history from the Messinian of northern Apennines. Sedimentology, 24:211-244. van den Berghe, B., 1983. Evolution s6dimentaire et structurale depuis le Pal6oc~ne du secteur 'Massif de la Selle' (Haiti), Baoruco (R6publique Dominicaine), 'Nord de la Ride Beata' dans l'Orog~ne Nord Caribe (Hispaniola, Grandes Antilles). Ph.D. Thesis, l'Universit6 Pierre et Marie Curie, Paris, 205 pp. Vann, I.R., Graham, R.H. and Hayward, A.B., 1986. The structure of mountain fronts. J. Struct. Geol., 8:215-227. Wadge, G. and Wooden, J.L., 1982. Late Cenozoic alkaline volcanism in the northwestern Caribbean: tectonic setting and Sr isotopic characteristics. Earth Planet. Sci. Lett., 57: 35-46. Warren, J.K., 1986. Perspectives: shallow-water evaporitic environments and their source rock potential. J. Sediment. Petrol., 56: 442-454. Warren, J.K., 1988. Evaporite sedimentology: Importance in hydrocarbon accumulations. Prentice-Hall/IHRDC, 304 p. White, A.H. and Youngs, B.C., 1980. Cambrian alkali playalacustrine sequence in the northeastern Officer basin, South Australia. J. Sediment. Petrol., 50:1279-1286. Winker, C.D. and Buffler, R.T., 1988. Paleographic evolution of early deep-water Gulf of Mexico and margins, Jurassic to Middle Cretaceous (Comanchean). Am. Assoc. Pet. Geol. Bull., 72:318-346.
Chapter 13
Evolution of the Neogene Kingshill Basin of St. Croix, U.S. Virgin Islands
I V A N G I L L , P E T E R P. M C L A U G H L I N ,
JR. a n d D E N N I S K. H U B B A R D
The sedimentary rocks of the Neogene Kingshill basin of St. Croix record part of the evolution of the tectonically complex region at the eastern edge of the North American-Caribbean plate boundary zone. The Kingshill basin is a northeasterly oriented graben or half-graben that contains a thick section of Neogene carbonates bounded by fault blocks of Cretaceous siliciclastic and intrusive rocks. Significant details of basin development have been added by the inclusion of data from a drilling program that included fourteen test holes with cumulative footage exceeding 533 m and a maximum depth of 91 m. Additional information came from outcrop sampling over the ca. 80 km 2 basin, subsurface records, and samples from engineering and water wells donated to the project. Previous models of basin development suggest that the carbonate rocks of the Kingshill basin were deposited (1) in shallow water or (2) entirely within the confines of an insular graben system. These models assume an isolated insular basin with self-contained sediment source. Instead, subsurface evidence suggests that early Kingshill basin sedimentation started in deep marine conditions prior to faulting on the basin margins and includes incursions of coarse, reef-derived sediment from a nearby source. The period of pre-rift sedimentation is documented to extend into the early Middle Miocene, but probably extends into the Oligocene or earlier. The faulting that formed the basin margins was initiated no earlier than the late Middle Miocene. After rifting, the Kingshill basin underwent significant shallowing and uplift in Late Miocene to Early Pliocene time. Basin development culminated in the establishment of a Pliocene reef tract and several episodes of subaerial exposure. The Jealousy Formation, the lowest formation described, is an entirely subsurface Middle Miocene unit of dark marls deposited at middle bathyal depths. The Kingshill Limestone conformably and diachronously overlies the Jealousy Formation and is divided into two members. The La Reine Member is characterized by buff pelagic limestones and marls with an upward increasing proportion of intercalated shelf-derived sediment flows. It ranges from basal Middle Miocene to uppermost Miocene and exhibits a transition from middle bathyal to upper bathyal environments. The Mannings Bay Member is composed of skeletal debris-rich carbonate slope deposits and lies near the Miocene-Pliocene boundary. The Blessing Formation overlies the Kingshill Limestone and represents a reef system that existed on the south coast of St. Croix during the Early Pliocene. Stratal relations on the basin margins indicate that the Jealousy Formation and at least the lower part of the Kingshill Limestone were deposited prior to graben formation near the end of the Middle Miocene. Subsidence analysis of the Neogene section indicates that 400 m of vertical uplift occurred on St. Croix between 10.5 and 3.5 Ma. A right-lateral model of movement between St. Croix and the Puerto Rico platform has been suggested by several recent workers. This model is consistent with the geomorphology of the Virgin Islands Trough and the Anegada Passage with right-lateral strike-slip motion in the Anegada Passage opening the Virgin Islands Trough as a pull-apart basin. However, an older left-lateral model of island movement is consistent with the northeasterly orientation of the normal fault system of St. Croix and the St. Croix Ridge. In addition, left-lateral motion would locate pre-rift St. Croix south of the known extra-basinal sources of Cretaceous and Tertiary shelf sediment required by the timing of Kingshill basin sedimentation. In this model, the Puerto Rico platform area could act to disperse slip between the North American and Caribbean plates. A variety of models are possible, but each should take into account geologic details of the Kingshill basin development.
INTRODUCTION AND GEOLOGIC SETTING
d i s t i n c t f r o m the m a j o r i t y o f the p r i m a r i l y i g n e o u s i s l a n d s o f the L e s s e r A n t i l l e s . A t h i c k s e c t i o n o f
St. C r o i x is the s o u t h e r n m o s t o f the U.S. V i r g i n I s l a n d s , l o c a t e d at the e a s t e n d o f the G r e a t e r A n t i l l e s
Neogene
c a r b o n a t e s o c c u p i e s a c e n t r a l g r a b e n or
h a l f - g r a b e n , h e r e r e f e r r e d to as the K i n g s h i l l b a s i n
a n d the n o r t h w e s t e d g e o f the L e s s e r A n t i l l e s arc
(Fig. 2). T h i s b a s i n lies b e t w e e n fault b l o c k s o f
(Fig. 1). T h e i s l a n d is t e c t o n i c a l l y a n d g e o l o g i c a l l y
C r e t a c e o u s s i l i c i c l a s t i c a n d i n t r u s i v e r o c k s o f the
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 343-366. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
344
I. GILL et al.
Fig. 1. Location map of St. Croix, the Virgin Islands platform, and the Virgin Islands basin. Bahymetry and structure are after Houlgatte (1983). Inset: NOAM = North American plate; SOAM -- South American plate; CARIB = Caribbean plate (after Burke et al., 1984).
Fig. 2. Generalized geologic map of St. Croix from Whetten (1974). Exposed strata mapped as Jealousy Formation by Whetten (1966) are re-mapped as Kingshill Limestone in this paper.
Mt. Eagle Group that comprise the mountainous East End and Northside Ranges. The Neogene carbonate section, which is the focus of this paper, is divided into three forma-
tions (Fig. 3): the blue-gray marls of the Jealousy Formation; marls and limestones of the Kingshill Limestone; and reef limestones of the Blessing Formation (Gill et al., in press). The stratigraphy of this
EVOLUTION OF THE NEOGENE KINGSHILL BASIN OF ST. CROIX, U.S. VIRGIN ISLANDS basin provides clues to the tectonic evolution of the eastern end of the North American-Caribbean plate boundary zone (PBZ). The purpose of this paper is to trace the evolution of the Kingshill basin from the Miocene to Recent based on observations from test holes and outcrops in the central plain of the island. This investigation also evaluates the implications of these findings, tied to marine geology studies in the area, for plate tectonic models of the northeastern part of the Caribbean region. Up to this point, there have been few integrated biostratigraphic-stratigraphic studies that have incorporated subsurface information. Previous studies of the geology of St. Croix have, for the most part, been based solely on outcrop data. These studies considered the carbonates of the Kingshill basin to record deposition in an isolated Oligocene-Miocene graben system (Multer et al., 1977; Gerhard et al., 1978; Lidz, 1984a). Whetten (1966) produced a detailed geologic map of St. Croix and a particularly detailed description of its Cretaceous section; some of this work is summarized in Fig. 2. Multer et al. (1977) and Gerhard et al. (1978) provided modern models for the carbonate sedimentation, including the structural and sedimentological model of the basin. Gerhard et al. (1978) designated the type section and provided detailed petrologic descriptions of the Kingshill Limestone that are still pertinent today. Lidz (1982, 1984a, 1988) defined the biostratigraphic relationships within the basin, tied them to the basin model of Gerhard et al. (1978) and suggested ties to global eustasy. In the only work to include subsurface information, Cederstrom (1950) provided an early geologic map and a detailed description of early deep drilling work on the island. This work includes the type-section description of the Jealousy Formation. Most of these studies either did not address a wider tectonic framework, or have related the late Cenozoic tectonic evolution of St. Croix solely to vertical uplift. However, cores from a drilling program conducted in the 1980's furnish new subsurface data that, in conjunction with outcrop data, help to more clearly delineate the sedimentary and structural evolution of the Kingshill basin during the Neogene. The drilling program included fourteen test wells drilled to depths of up to 91 m as well as data from a number of private wells donated to the project (Fig. 4). These wells were logged during drilling, and samples were collected for sedimentologic, micropaleontologic, and geochemical analysis at intervals of 1.5 or 3 m in the wells (Gill and Hubbard, 1986, 1987; McLaughlin et al., 1995). Core material and logs from pre-existing wells provided additional data on the carbonate units underlying the southeastern portion of the central plain.
345
STRATIGRAPHY AND DEPOSITIONAL SETTING OF THE NEOGENE FORMATIONS OF THE KINGSHILL BASIN Jealousy Formation Lithology and distribution
The Jealousy Formation is a unit of blue-gray marls that underlies much of the central plain of St. Croix. The top of the formation is marked and abrupt in the subsurface; water-well drillers treat the top of this 'blue clay' as hydrologic basement and generally stop drilling when it is reached. The blue-gray marls are rich in planktonic foraminifera and other deep-water microfauna. The formation includes a number of conglomeratic limestone and thin limestone layers in the deep subsurface, below the reach of the drilling conducted for this study (Cederstrom, 1950). These coarse-grained beds are bracketed above and below by pelagic blue marls, so are considered allochthonous deposits of down-slope transported debris. Test well data indicate the Jealousy Formation is present in the subsurface throughout the central plain region, both inside and outside the fault boundaries of the Kingshill basin graben. The type section was defined by Cederstrom (1950) in the deepest of several test wells (Test Well 41) drilled by the Civilian Conservation Corps (CCC) in 1939, where a thickness of more than 426 m was encountered (Fig. 4). It is present in wells M1, M2, and M10 of this study, all west of the eastern fault boundary of the graben, and was reported by Cederstrom (1950) at 18 m below sea level in CCC Test Well C26, approximately 1 km east of this fault (Fig. 4). A maximum thickness of 450 m has thus far been recognized for this unit (Cederstrom, 1950). Although its base has never been reached in the center of the basin, gravity surveys indicate that more than 1800 m of Jealousy and older sedimentary rocks may underlie the central plain (Shurbet et al., 1956; R.C. Speed, written commun., 1994). Although the top of the Jealousy Formation exhibits considerable relief in the subsurface (Fig. 5), no apparent change in bulk mineralogy, microfauna, or grain size is observed across the Jealousy/Kingshill boundary. No hiatus or missing section is evident within the resolution of available biostratigraphic control. Extensive areas of Jealousy Formation outcrop exposures have been mapped in some previous studies (Cederstrom, 1950; Whetten, 1966). However, we recognize the Jealousy Formation as an exclusively subsurface unit and suggest that these outcrops are more correctly mapped as Kingshill Limestone, following Gerhard et al. (1978).
Fig. 3. Stratigraphic column, chronostratigraphic framework, and paleoenvironments of the Neogene section of the Kingshill basin (after McLaughlin et al., 1995). Planktonic foraminiferal zonation based on Bolli and Saunders (1985). Chronostratigraphy, coastal onlap curve, and eustatic cycles after Haq et al. (1988).
t" t"*
4~
EVOLUTION OF THE NEOGENE KINGSHILL BASIN OF ST. CROIX, U.S. VIRGIN ISLANDS
347
Fig. 4. Locations of outcrops, test wells and water wells used in the stratigraphic cross-sections. A/P = Airport/Penitentiary; AQ -Airport Quarry; FC = Five Corners; HC = Hess Cut; MS = Morningstar; R/B -- Rattan/Belvedere; SR --- Salt River valley; VR = Villa La Reine and Fredensburg Quarry; WR = Work and Rest. Core sample locations designated with M are test holes drilled by Gill, the others are from previous studies. Cutting sample locations are noted for water well cuttings studied by the authors. Test holes drilled in 1939 by the Civilian Conservation Corps are designated by C.
Age and paleoenvironment Our recent micropaleontologic studies of borehole samples from the Kingshill graben place the Jealousy Formation in the lower part of the Middle Miocene, ranging from the P r a e o r b u l i n a g l o m e r o s a zone to the G l o b o r o t a l i a f o h s i f o h s i zone (McLaughlin et al., 1995). The Jealousy Formation has previously been referred to as Oligocene (Cushman, 1946; Cederstrom, 1950) and even as low as Middle Eocene (Lidz, 1984b). However, the Oligocene citations are based on older notions of the age significance of certain benthic macrofauna and larger foraminifera. The Middle Eocene citation is based on the planktonic foraminiferal fauna found by Lidz (1984b) in an allochthonous shale clast in the Kingshill Limestone that is presumed to have been derived from the Jealousy Formation. Although no in-place paleontologic evidence exists for an age any older than Miocene for the Jealousy Formation, the estimated 1800 m thickness of sedimentary fill in the Kingshill basin (Shurbet et al., 1956) leaves open the possibility that the Jealousy Formation and any underlying units could extend as far as the Oligocene or lower.
The Jealousy Formation is a dominantly hemipelagic, deep-water unit. The benthic foraminiferal fauna indicates deposition at 600 to 800 m water depth (McLaughlin et al., 1995). Most species recovered are generally associated with middle and upper bathyal environments; several species are present that indicate an environment no shallower than the middle bathyal zone. This differs from previous interpretations of the Jealousy Formation as an estuarine deposit (Van den Bold, 1970; Multer et al., 1977) based on outcrop samples previously mapped as Jealousy.
Source area and paleocurrents The coarse, shelf-derived carbonates in the Jealousy Formation are sandwiched between large intervals of foram-rich basinal sediments. The coarse carbonates are therefore allochthonous, and require a nearby shelf source. The thickness of Jealousy strata makes the uplifted horst blocks of St. Croix an unlikely source area (Gerhard et al., 1978) even if they had existed during Jealousy deposition. No data exist on presumed paleocurrent directions in the Jealousy Formation.
348
I. GILL et al. 146 m
M2
M8
South
North
A'
A 30 m -
SL-
,--30 m
M3 /~
-N
i---_----------_~
N
SL
;,.o.O'.o.. o'.,~~;; -30 m--
o . ; O * Oo-,
-30 m
O* oo-o
;-o-O"o.
o'.o:. ;; ~z.',
o:o..Z"
c~; o * o d
;'o%~ o . p ~
i ,,i
/ I
I !
~-~;;/'-~-I ~ '
I
'-"'
i
i j,_
I
,/
7-'--I"
TEST HOLE
~ v-'-r:'r
1000 rn I~.| ~ I
~eL-l-,.J~/-]~ ! . l .I I ./
~ iI I
'
J I
I I
I
i
i
I
I
t
Y'-t-:-'i-
I/-I'i--I
I~
l --I--
I/l--/
i
I -- I-- I I--I I--
-I
|
FE~
i
VERT. EXAG. 50 X
--
I--I--
I:1
/
~
I I-
DOLOMITE ALLUVIUM
KINGSHILL FM (MANNINGS BAY MBR) & BLESSING FM KINGSHILL FM (LA REINE MEMBER) JEALOUSY FORMATION
r~i
CRETACEOUS SlLIClCLASTICS
West
East
B
B'
Fairplain Fault
30 m
WD272
M5
PA7 PAl0
M4 B7
SL _
:
I
I
i
i
M9 "~,~W~ &~'
Ii~-."
&~.':." ~ "o-
o...%
I~-.'-~ g * . / l - - - - - I ~
,,,, ~ - ; : ~ h ~ ~ ~ , ~ 1,, ~.o....:~ I
__~
./
/
TEST HOLE
1000 m
' 1
z
I~ i
I
J i
I
I~/ I
I
~
~'~ '
3 9
~
"]
I
DOLOMITE ALLUVIUM
FE~ VERT. EXAG. 50 X
.
J
,,. I I..I .I,.[ i
,~
~
i~:..q.~
~.~;~" ~.~"
9-
-30
1-30
KINGSHILL FM (MANNINGS BAY MBR) & BLESSING FM KINGSHILL FM (LA REINE MEMBER) JEALOUSY FORMATION
. . . . . . . . . . . .
,R
.
CRETACEOUS SILICICLASTICS
Fig. 5. Geological crossections through the Kingshill basin. (A) North-south cross-section A-A'. Note that the Jealousy Formation surface roughly follows the topography of the Kingshill Limestone. (B) East-west cross-section B-B'. A normal fault (the Fairplain fault) forms the western boundary of the small graben on the south coast occurs between test wells M 1 and M4. The Jealousy Formation was not reached to the east of this fault.
Kingshill Limestone The blue-gray marls of the Jealousy Formation are succeeded upward by the more carbonate-rich succession of the Kingshill Limestone. The Kingshill Limestone crops out over large areas of the central
plain and can also be mapped in the subsurface based on well data. It is composed of limestones and buff pelagic marls with an upward-increasing proportion of shelf-derived sediment gravity-flows. Lithologic variations permit it to be divided into two members (Gill et al., in press): the interbedded
EVOLUTION OF THE NEOGENE KINGSHILL BASIN OF ST. CROIX, U.S. VIRGIN ISLANDS
Fig. 6. Interbedded planktonic foraminiferal packstones and sediment gravity flows of shallow-water debris at Villa La Reine outcrop, type section of the Kingshill Limestone (La Reine Member). Biogeographic control places this outcrop near the boundary of the Middle Miocene and Upper Miocene. This is denoted as location VR in Fig. 4.
marls and limestones of the La Reine Member and the benthic-foram-rich, burrowed limestone of the Mannings Bay Member.
Lower Kingshill Limestone- La Reine Member Lithology and distribution The La Reine Member is composed of interbedded planktonic foraminifera-rich marls and shallowmarine limestone debris beds with increasing proportions of downslope-transported material upsection. Typical lithologies of the La Reine Member are exposed at the type section of the Kingshill Limestone at Villa La Reine (Fig. 6). The outcrop is a rhythmically bedded alternation of polymictic packstones (Gerhard et al., 1978), some with boulder-sized coral heads, and deep-water planktonic foraminiferal chalks and marls. Similar lithologies occur at Fredensburg Quarry and Estate Work and Rest, but breccia beds at the latter include terrigenous material presumably derived from the Cretaceous Mt. Eagle Series. The lower part of the La Reine Member is similar but includes less transported debris (Five Corners,
349
Rattan/Belvedere, and Morningstar sections). In the subsurface, this interval is dominated by planktonic foraminiferal packstone, with less common lithic-pebble or foraminifera-rich wackestones. The boundary between it and the underlying Jealousy Formation is marked by a distinct color change from tan above to blue-gray below. However, the significance of this change is unclear; sedimentological and micropaleontological evidence reveal no notable change in lithology, mineralogy, or depositional environment, nor is any hiatus resolvable. The stratigraphically highest part of the La Reine Member is exposed in the Airport/Penitentiary section along the Melvin Evans Highway, where it is disconformably overlain by the Mannings Bay Member (Fig. 7). This interval is characterized by regularly bedded intercalations of softer, planktonic foraminifera-rich beds and more indurated, graded beds of shelf-derived debris. The quantity of shelfderived sand is greater than lower in the member, and burrowing appears to be more pervasive. The La Reine Member in the St. John's/Judith Fancy area includes beds of calcareous conglomerate composed of rounded terrigenous gravel and a fauna of shallow-water echinoids and benthic foraminifera. Previously, these and nearby outcrops in the Northside Range have been considered to be shelf and lagoon deposits or part of the Jealousy Formation, based on their similarity to conglomerates encountered in the CCC Test Well 39 (Gerhard et al., 1978; Lidz, 1982; Andreieff et al., 1986). However, because the conglomerate beds are overlain in outcrop, and underlain in Well M10 (Fig. 4), by planktonic foraminiferal packstones, we interpret them as allochthonous beds occurring within a succession of typical La Reine Member deep-water strata. In addition, structural relations support inclusion of these strata in the Kingshill Limestone rather than the Jealousy Formation. Because these exposures occur at elevations similar to outcrops of the La Reine Member only a few kilometers away (Fig. 4), faulting within the graben would be required to raise the stratigraphically lower Jealousy beds to the same elevation as the nearby Kingshill. However, there is no evidence of such faulting. The maximum thickness of the La Reine Member encountered in the test wells drilled for this study is approximately 140 m (Fig. 5A). Cederstrom (1950) reported a thickness range from 0 to 180 m for the Kingshill Limestone, the larger figure referring to extrapolated thickness in the carbonate highlands of the Rattan Hill area. Isopach patterns reveal thinning of the La Reine Member toward the north and northwest margins of the basin. The formation shows a pronounced thickening in the carbonate highlands close to the northern coast of St. Croix and less pronounced thickening toward the south, interrupted
350
I. GILL et al.
Fig. 7. Disconformable contact between the La Reine and Mannings Bay Members of the Kingshill Limestone at the Airport/Penetentiary outcrop, along Melvin Evans Highway, southern St. Croix. Disconformity occurs midway up the outcrop, approximately 7 m above road level. This is denoted as location A/P in Fig. 4. by post-depositional faulting along the south coast (Fig. 5). In general, these isopach patterns follow the trends of the top of the Jealousy Formation. If deformation is ignored, the Kingshill Limestone isopach patterns imply a basin opening to the south but with a deep area under the modem carbonate highlands in the north.
Age and paleoenvironment The La Reine Member extends from basal Middle Miocene to approximately the Miocene-Pliocene boundary (McLaughlin et al., 1995). The subsurface section sampled in our drilling program is the stratigraphically lowest, ranging from basal Middle Miocene (Praeorbulina glomerosa zone) to the medial Middle Miocene (Globorotalia fohsi robusta zone). The stratigraphically lowest outcrops are in the northern part of the island (Salt River Valley and Five Comers, Fig. 4), where the La Reine Member is placed in the lower part of the Middle Miocene (Globorotaliafohsifohsi zone, possibly to Praeorbulina glomerosa zone). The type section of the Kingshill Formation at Villa La Reine represents the middle part of the formation, with faunas indicative of the upper part of the Middle Miocene (Globorotalia mayeri zone, possibly to the Globorotalia menardii zone). The top of the member crops out on the south side of the island (Airport/Penitentiary section), where it is placed near the top of the Miocene (upper part of Globorotalia humerosa zone). The outcrops in the St. John area previously mapped as Jealousy Formation fall biostratigraph-
ically within the range of the La Reine Member, supporting the lithologic and structural arguments against their inclusion in the Jealousy. Van den Bold (1970, in Gill, 1989, and in McLaughlin et al., 1995) considers the ostracode fauna indicative of a position near the Lower Miocene-Middle Miocene boundary. The fauna is completely different from that of the subsurface Jealousy Formation, but contains several species in common with the lower part of the La Reine Member in the subsurface. Middle Eocene to Early Miocene foraminifera were described by Lidz (1984b) from a mud clast in the La Reine Member. Its presence in these deepwater strata indicates that older Tertiary sediments were being eroded and transported by sediment gravity flows during the deposition of the La Reine Member. This is consistent with the occurrence of pebbles in both the Jealousy Formation and Kingshill Limestone that are assumed to be derived from the Cretaceous strata presently exposed in the highlands of the basin-bounding fault blocks. The benthic foraminifera of the La Reine Member in Wells M1, M2, and M10 comprise a middle bathyal fauna (600-800 m water-depth) that differs little from that of the underlying Jealousy Formation (McLaughlin et al., 1995). No significant paleoenvironmental shift is evident at the boundary. Although there are significant numbers of shallower-water species in some of the stratigraphically higher outcrop samples from the La Reine Member, the faunas in the finer-grained beds that over- and underlie these samples are middle bathyal types, indicating that the
EVOLUTION OF THE NEOGENE KINGSHILL BASIN OF ST. CROIX, U.S. VIRGIN ISLANDS shallow-water forms are present due to downslope transport.
Source area and paleocurrents Paleocurrent indicators in the lower Kingshill Limestone at the Villa La Reine type section show west-southwest flow. Four measurements included a pebble halo around a boulder and orientations from cross-lamination and oscillation ripples (Gerhard et al., 1978). Clasts larger than 4 mm are concentrated in the northeast portion of the basin (Gerhard et al., 1978). None of the reef-derived material, including boulder-sized coral heads, is in-place, and this material must therefore be derived from a nearby shelf area.
Upper Kingshill Limestone: Mannings Bay Member Lithology and distribution The Mannings Bay Member of the Kingshill Limestone is characterized by channelized beds of grainstones rich in shelf debris, interbedded with softer wackestones and packstones. The grainstones contain abundant Operculinoides cojimarensis and Paraspiroclypeus chawneri (Behrens, 1976; Gerhard et al., 1978; S. Frost, oral commun., 1986). Many specimens show signs of transport or reworking, such as fracturing, abrasion, and imbrication. The foraminiferal wackestones and the softer packstones also include significant quantities of planktonic foraminifera. These lithologies are well exposed in the type section at the quarry on the southeast side of Mannings Hill (Gill et al., in press) and along
351
parts of Evans Highway, notably the upper part of the Airport/Penitentiary roadcut (Figs. 4 and 7). It was also examined in several of the test wells drilled for this project. The Mannings Bay Member rests disconformably on the La Reine Member, from which it was distinguished based on higher abundance of shallow-water carbonate material (Gill, 1989; Gill et al., in press). At the Airport/Penitentiary section, the disconformity is evident as a scour surface with more than 1 m of relief. This surface appears to have been scoured by partly channelized submarine flows of shelf-derived sediment (Lidz, 1984a; Gill, 1989; Gill et al., 1990). The Mannings Bay Member includes the strata referred to as a "benthic foraminiferal wackestone and grainstone facies" in the Kingshill Limestone by Gerhard et al. (1978) and those strata separated from the Kingshill Limestone as "postKingshill" limestones by Lidz (1982) and Andreieff et al. (1986). The Mannings Bay Member and the overlying Blessing Formation are best developed in a small graben on the south coast of St. Croix (Fig. 8). They can be difficult to differentiate from one another in core but together total more than 50 m thickness in some of the wells (Gill and Hubbard, 1986, 1987). The westernmost documented extent of the Mannings Bay Member is near the western fault boundary of this graben near Fairplain. The eastern edge of this graben is perhaps indicated by where the stream flow makes an abrupt southerly turn to the coast against the exposed Pliocene reef complex (Fig. 9).
Fig. 8. Disconformable contact between the Mannings Bay Member of the Kingshill Limestone (below) and the Blessing Formation (above) at the Airport Quarry outcrop. This is denoted as location AQ in Fig. 4.
352
I. GILL et al.
Fig. 9. Facies map for south coast industrial area. Dolomite in the vadose zone or exposed in outcrop is patchily distributed in an arcuate region following the Pliocene reef trend. Dolomite presently in the phreatic zone is found in off-shore facies. The western boundary of the small, south coast graben is well-defined by a normal fault. The northern and eastern boundaries are poorly known.
Age and paleoenvironment Planktonic foraminifera are generally uncommon and poor in the Mannings Bay Member. The basal part of the member, in the Airport/Penitentiary section, contains faunas suggestive of the lower Pliocene Globorotalia margaritae zone (McLaughlin et al., 1995), consistent with Lower Pliocene findings for this same section by Andreieff et al. (1986, 1987) and Lidz (1982). A subsurface sample from Well M4 yields a fauna indicative of an interval near the Miocene-Pliocene boundary, between the Globorotalia humerosa zone and the Globigerinoides trilobus fistulosus zone (McLaughlin et al., 1995). The foraminiferal control from above and below the disconformity at the base of the Mannings Bay Member suggests a general chronostratigraphic position at or near the terminal-Miocene (5.5 Ma) eustatic fall, but does not provide sufficient resolution to tie it exactly to this event, as has been proposed by Lidz (1982). The biostratigraphic control is also insufficient to determine whether a significant chronostratigraphic interval is missing at the disconformity. Smaller benthic foraminifera recovered from the Airport/Penitentiary section are a mix of outer ner-
itic species and forms associated with shallow-water carbonate environments (McLaughlin et al., 1995). In outcrop and core samples, the larger benthic foraminifera Operculinoides and Paraspiroclypeus are abundant; these forms were likely photic-zone inhabitants (S. Frost, oral commun., 1986). Other bioclasts that contribute to the facies are coralline algal crusts, rhodoliths, echinoid spines and plate fragments, coral fragments, and molluscan debris. These forms, the poorly developed planktonic fauna, and the evidence for transport in the larger foraminiferarich beds together indicate that the shallow-water material was carried into a deeper shelfal setting of approximately 100 m depth.
Source area and paleocurrents The beds of larger benthic forams show sorting and imbrication, the result of extensive current working. Sediment transport direction in the Kingshill basin was dominantly to the west-northwest, and ranged from southwest to north-northeast. This is based on 42 measurements of imbricated large benthic foraminifers on Mannings Hill, probably the same outcrop referred to here as the Airport/Penitentiary outcrop (Gerhard et al., 1978).
EVOLUTION OF THE NEOGENE KINGSHILL BASIN OF ST. CROIX, U.S. VIRGIN ISLANDS
Blessing Formation Lithology and distribution The highest stratigraphic unit, the Blessing Formation, represents a Pliocene reef tract that extended across the south and west coastlines of St. Croix. The reef tract consisted of interspersed reefs and shelf systems similar to the arrangement of reefs around the modem south coastline of St. Croix. The classic reef model with flanking fore- and backreef facies does not appear to apply here. Reef systems on St. Croix apparently formed planar deposits with little topographic relief. This planar geometry is apparently common in Caribbean Tertiary reef deposits (S. Frost, oral commun., 1986). The Blessing Formation was sampled in outcrops near the south coast and in cores. Based on subsurface data, its greatest thickness is in a small graben, just east of the Fairplain fault, where it may be up to 30 m thick. Scattered exposures of reef facies also occur west of the Fairplain fault along Evans Highway, a location near Fredericksted, and at an exposure described by Gerhard et al. (1978); its maximum thickness in that area is estimated to be between 10 and 20 m. The best exposures of this unit are in its type section at a road cut next to the Hess Oil refinery (Gill et al., 1990, in press) (Figs. 4 and 10). Reefal facies are predominantly composed of coralline boundstones characterized by external molds of scleractinians, gastropods and pelecypods, as well as skeletal debris. The scleractinians include species of extant genera such as Agaricia, Diploria, Montastrea, and Siderastrea, as well as the extinct forms Stylophora, Teliophyllia, and Thysanus (Behrens, 1976). Lagoonal facies include skeletal wackestones composed of shallowwater foraminifera, coralline algae and a wide variety of shallow-water invertebrates. The Hess outcrop is marked by several wellcemented undulatory layers (Fig. 10) distinguished by abrupt light stable-isotopic excursions, an onlap surface (Gill, 1989) and karstification (Lidz, 1984a; Gill, 1989). Nearby exposures were apparently marked by terra rosa beds within the Blessing Formation (Behrens, 1976) and underneath it (S. Frost, pers. commun., 1986; L. Gerhard, written commun., 1997). These surfaces indicate that the south coast reef trend of St. Croix was exposed several times during the Pliocene (Gill, 1989; Gill et al., 1990). Age and paleoenvironment The Blessing Formation is loosely placed by biostratigraphic data in the interval between the upper part of Lower Pliocene and the top of the Pliocene. This assignment is based on the occurrence of Globigerina nepenthes in the Hess refinery outcrop
353
(Lidz, 1982), its position above the Lower Pliocene Mannings Bay Member (Andreieff et al., 1986; McLaughlin et al., 1995), and the occurrence of prePleistocene species of the scleractinian corals Teliophyllia, Stylophora, and Thysanus (Behrens, 1976). No planktonic foraminifera were recovered from our samples. The macrofaunal assemblages within the Blessing Formation represent co-existing reef, forereef, and lagoon environments that extended along the south and west coastlines of St. Croix.
Source area and paleocurrents Much of the exposed Blessing Formation appears to be in-place. The morphology of the reef tract mimics the present shoreline, and the reefs would have been affected by open-ocean conditions to the south and to the west.
EVOLUTION OF THE KINGSHILLBASIN In previous studies, the Kingshill basin was thought to have formed in the Oligocene as a result of vertical tectonic movement. Whetten (1966) characterized the carbonate section of the Kingshill basin as reefal and estuarine deposits that accumulated in a graben in the central part of the island following a period of low-rank metamorphism, faulting, folding, igneous intrusion, and uplift. He concluded that there was no significant evidence for strike-slip motion north of St. Croix, and that the Neogene section was affected only by vertical tectonics. More recent studies (e.g. Multer et al., 1977; Gerhard et al., 1978; Lidz, 1982) have accepted this structural framework; for example, Multer et al. (1977) envisioned the Northside Range and the East End Range as subaerially exposed horst blocks that provided both terrigenous and shelf-derived carbonate debris to the basin. However, these studies recognized that deposition took place in a deep-marine setting, which they envisioned as a seaway opening to the northeast and southwest in a basinal setting similar to modem deep basins north of St. Croix.
Basin-margin faulting and basin formation Subsurface stratigraphic evidence indicates that the Kingshill basin graben began to form no earlier than the latest part of the Middle Miocene, during deposition of the upper part of the La Reine Member. The Jealousy Formation and the lower part of the La Reine Member show no evidence of tectonic activity. These strata were deposited at water depths of approximately 600 m during the early part of the Middle Miocene (McLaughlin et al., 1995). Previous studies have suggested the existence of an active graben at this time with subaerially exposed mar-
354
I. GILL et al.
Fig. 10. Reefal facies of the Blessing Formation at the Hess Cut outcrop, southern St. Croix. HG = coral reef hardground, with onlap onto exposure surfaces (ES). CAL -- undulatory caliche layer under karst cavities. Backpack and hammer for scale at arrow. This is denoted as location HC in Fig. 4. gins. However, the existence of tentatively identified deep-water marls of the Jealousy Formation outside the graben in CCC Test Well C-26 (Cederstrom, 1950) indicates that faulting on the graben margins must post-date these early Middle Miocene deposits. In addition, exposed graben margins would require a marginal slope exceeding 45 ~ based on the measured distance from the test well samples in the basin to the hypothetical exposed margin on the northwest side of the graben. This angle is comparable to that of the modem slope north of St. Croix, which is dominated sedimentologically by input of shelf-derived material (Hubbard et al., 1981; Gill, 1983), including reef foraminifera such as Amphistegina (B. Sen Gupta, oral commun., 1984). Our samples from the Jealousy Formation and the lower part of the La Reine Member near this boundary do not show the major input of shelfal material that would be expected with deposition at the foot of a similarly steep slope. If a steep-sided Kingshill basin did not form before the middle Miocene, then the scattered conglomerates and shelf-derived sediments that do occur in the Jealousy Formation and lower part of the Kingshill Limestone must be derived from somewhere other than St. Croix. These lithologies suggest that St. Croix was close to a land mass capable of supporting reef growth and supplying clastic materials to the deep-water environment during the Middle Miocene. Puerto Rico and the Virgin Islands platform, to the northwest of St. Croix, and Anguilla and Saba to the northeast, are possible source areas;
either requires significant lateral translation of the St. Croix platform. We suggest that initiation of the St. Croix fault system occurred no earlier than in the latest part of the Middle Miocene (Fig. 11). The contacts between the lower Kingshill Limestone and the Cretaceous rocks on the east and west sides of the graben have been interpreted as faults by Multer et al. (1977) although they were mapped simply as stratigraphic contacts on the geologic map by Whetten (1966). The western contact, along the edge of the Northside Range, is mostly obscured by alluvial cover (Fig. 4). Gerhard et al. (1978) have suggested that displacement along the eastern fault boundary of the graben was greater than that along the western boundary. The eastern fault contact is sharp and characterized by offset of both Cretaceous and Kingshill strata. This age of graben formation is constrained by fault relations between the La Reine Member and the Cretaceous strata along the eastern boundary of the graben, which indicates that at least the lower part of the La Reine Member was deposited prior to basin faulting (Gill and Hubbard, 1986, 1987). Beds of coral debris and rounded pebble conglomerate exist low in the exposed section near Judith's Fancy and St. John and were penetrated by test hole M10. These materials were interpreted as in-place deposits by Gerhard et al. (1978) and Lidz (1982), but are interpreted as allochthonous deposits here (Gill, 1989). Similarly, the type-section at Villa La Reine (Gerhard et al., 1978) records the influx of shallow-marine and terrigenous debris in the basin.
EVOLUTION OF THE NEOGENE KINGSHILL BASIN OF ST. CROIX, U.S. VIRGIN ISLANDS
355
Fig. 11. Block models of the evolution of the Kingshill basin of St. Croix from the Early Miocene to Recent. The model reflects overall tectonic quiescence until near the end of the Middle Miocene when significant tectonism, uplift, and erosion of uplifted Jealousy Formation sediments apparently began. The Late Miocene to Plio-Pleistocene diagrams trace shoaling in the basin from bathyal to shallow-marine and reefal deposition. This section is placed near the Middle-Late Miocene boundary (Andreieff et al., 1986; McLaughlin et al., 1995), and includes a number of beds that contain large coral heads and siliciclastic material probably derived from exposed and eroded rock outside of the graben. Together these data suggest a shift from relative tectonic quiescence to uplift activity near the end of the Middle Miocene. Breccia beds occur in strata of approximately the same age at Estate Work and Rest along the eastern edge of the basin. These beds have been interpreted as syntectonic breccia by Gerhard et al. (1978) in an outcrop that could not be located for this study. In nearby exposures, including the Estate Work and Rest sec-
tion (Fig. 4), angular clasts of typical Cretaceous Mt. Eagle Group lithologies form breccia layers between beds of the La Reine Member. The Estate Work and Rest section is placed biostratigraphically near the Middle Miocene-Late Miocene boundary, between the Globorotalia mayeri and Globorotalia acostaensis zones (McLaughlin et al., 1995). This breccia may provide the earliest evidence for fault activity on the eastern boundary of the graben.
Eustatic events and basinal shallowing The effects of eustasy superimposed on tectonic uplift produced the disconformity separating the La
356 Reine and Mannings Bay Members at the Airport/ Penitentiary outcrop. This disconformity was interpreted by Lidz (1984a) as evidence of the Messinian eustatic fall of Haq et al. (1988). Our sedimentologic and paleontologic data indicate shoaling across this surface, from depths of approximately 200-300 m in the La Reine Member to 100 m in the Mannings Bay Member. There is no evidence of soil formation, dissolution, or karsting to indicate subaerial exposure. Based on this evidence, we believe this disconformity represents submarine erosion during the uplift of the island in the Late Miocene. The global sea-level fall associated with the Messinian event is of approximately the same age and could have triggered erosion via sediment gravity flows. However, biostratigraphic control (McLaughlin et al., 1995) cannot precisely tie the timing of this unconformity to the Messinian event, nor confirm a single event as its cause. The strata above the unconformity record the development of extensive foraminiferal-coralline algal bank environments on St. Croix during the early part of the Pliocene (Fig. 11). The environments served as shallow-water sources of the larger benthic foraminifera present in the Mannings Bay Member, in particular Operculinoides cojimarensis and Paraspiroclypeus chawneri (Gerhard et al., 1978). This environment is not present in modem St. Croix, where coral ecosystems predominate. It is likely that foram-algal ecosystems were supplanted by corals with the extinction of many larger foram groups in the Neogene (Frost, 1977). It is also possible that the relative lack of coral debris could be the result of (1) the upslope storage of coral reef sediments, with minor deposition only at sporadic intervals (e.g. Moore et al., 1976), (2) changes in circulation, nutrient or temperature conditions yielding competitive advantage to the foraminifera-algal community, or (3) rapid uplift and eustatic variation suppressing the establishment of coral reef systems. Fairplain fault
The tectonism that produced the shallowing across the boundary between the La Reine and Mannings May members is also reflected in faulting within the Kingshill basin. The Fairplain fault, which marks the western margin of the small graben on the south coast, cuts through the Mannings Bay Member and Blessing Formation, indicating that motion occurred on this fault during the Late Pliocene or later. The orientation of the Fairplain fault is roughly parallel to the orientation of the northeast-trending main basin boundary faults. It dips at least 20~ to the east, as indicated by the depth to the Jealousy Formation at the fault contact in several well sections: 29 m below sea level at Well M1 (Fig. 13); 53 m
I. GILL et al. below sea level in CCC Test Well 45a, located 60 m to the east of Well M1 (Fig. 4); and deeper than the 80 m below sea level bottom-hole depth of Well M4 less than 180 m to the east of M1. Minimum vertical fault displacement is 68 m, based on the occurrence of the La Reine Member-Mannings Bay Member contact at 24 m above sea level on Mannings Hill west of the fault, and approximately 44 m below sea level in Well M4, east of the fault. The presence of a fault at this location is also supported by surface features. Strata in nearby outcrops along Evan's Highway (Fig. 4) dip toward the fault line and ephemeral streams locally approximate the trend of the fault near the coastline (Fig. 9). The eastern edge of the presumed graben is not well marked. Stream flow north of Well M5 is from west to east, oddly parallel with the coastline. The eastern edge of this graben may be marked by the ephemeral stream drainage turning abruptly south where it meets the exposed Plio-Pleistocene reef trend (Fig. 9). Surface features also suggest that a northern hinge-line exists for this small graben, just north of the industrial areas along the coast. Together, these fault orientations suggest that normal faulting on the margins of this graben was produced by the same extensional tectonic regime that initiated faulting on the margins of the Kingshill basin during the latest part of the Middle Miocene. The stratal relations across the Fairplain fault suggest that the south-coast subsidiary graben existed as an entity during and after deposition of the Mannings Bay Member foraminiferal-algal facies. The fact that the greatest thickness of these deposits is preserved in the graben has two possible explanations: the graben formed a marine embayment along the south coast where these facies accumulated; or the strata were preferentially preserved within the subsiding graben during island uplift. The former alternative suggests that the faulting produced topographic relief prior to and during deposition of the foraminiferal-algal facies, whereas the latter alternative requires only post-depositional faulting. We believe that both processes are likely to have occurred. Pliocene reef tract and subaerial exposure
Continued tectonic uplift and shoaling of the Kingshill basin resulted in deposition of the Blessing Formation reef tract (Fig. 11). The greatest thickness of reef growth is found in the Krause Lagoon area on the south-central coastline where the arcuate distribution of reef and lagoonal facies suggest the existence of an embayment (Fig. 9). The size and shape of this embayment was probably controlled by faulting along the margins of the south coast graben. The Blessing Formation contains indications of more than one period of Pliocene subaerial exposure
EVOLUTION OF THE NEOGENE KINGSHILL BASIN OF ST. CROIX, U.S. VIRGIN ISLANDS along its southern coastline (described in a previous section), as well as an onlap surface near the Hess Oil Refinery. We suggest that eustatic variations superimposed on the overall tectonic-uplift trend account for the evidence of exposure noted in our field studies in St. Croix, the same as Lidz (1984a). The accumulation of the Pliocene sediments only on the south coast, the apparent extent of erosion/non-deposition in the northern central plain at the same time, and the general southerly dip of Neogene strata in the Kingshill basin suggest that Late Pliocene uplift preferentially raised the northern part of the island. Subsidence analysis
Subsidence analysis (Fig. 12) indicates that the majority of the Neogene shallowing in the Kingshill basin is due to tectonic uplift. The Kingshill basin shallowed from as much as 800 m of water depth in the Middle Miocene (ca. 10.5 Ma) to approximately 100 m in the Early Pliocene (ca. 3.5 Ma). Given a modem stratigraphic thickness of 250 m for the study interval, and correcting for the effects of sediment loading, water loading, and compaction, we calculate 400 m of tectonic uplift in this interval, which translates into a rate of 57 m/Ma.
TECTONIC EVOLUTION OF ST. CROIX AND THE NORTHEASTERN CARIBBEAN
357
and coastline of the island are less steep and less suggestive of recent faulting. On submersible dives in DSRV Alvin along the south wall of the Virgin Islands basin off the north coast of the island, Dill (1977) encountered structures he interpreted as fault gouge in St. Croix basement rocks. In similar dives off the northwest coast, a vertical escarpment greater than 10 m in height was observed at a depth of greater than 2600 m (Hubbard et al., 1981; Gill, 1983). The escarpment was composed of dark, terrigenous rock similar in appearance to the Cretaceous Mt. Eagle Group rocks that form the east and west hills of St. Croix. If in-place, the face of the escarpment suggests the role of fault-offset in creating the north slope of St. Croix. Like the southern margin of the Virgin Islands basin off St. Croix, the northern margin off Vieques and St. Thomas is also characterized by very steep slopes (Fig. 13). If the Virgin Islands basin formed as a result of rifting, these slopes may represent scarps formed during the rifting event. If these scarps were initially juxtaposed, the present position of St. Croix could have been achieved by movement of the island south and east relative to its initial position in the Virgin Islands platform. This motion would require a combination of left-lateral movement and tensional separation, which is consistent with the oblique left-lateral strike-slip motion for the formation of the Virgin Islands basin. Fault orientation
Left-lateral oblique motion models
Stratigraphic and structural evidence indicates that the Kingshill basin graben began to form no earlier than the late Middle Miocene. St. Croix may have began to rift away from Puerto Rico during this period by oblique left-lateral faulting, movement that could have opened the Virgin Islands basin (Fig. 13). Similar left-lateral faulting may have also occurred between St. Croix and the Saba Bank area to the northeast, but structural and bathymetric relations in the intervening St. Croix basin (Fig. 1) are less clear. A left-lateral tectonic model is consistent with several lines of evidence. Bathymetry
Bathymetric profiles along the north coast of St. Croix are rugged and steep, sloping between 23 and 45 ~ to the center of the Virgin Islands basin and dropping off at nearly vertical angles near the shelf edge (Fig. 13). The northwestern shoreline is carved from cliffs of the Northside Range, prompting Meyerhoff (1927) to suggest relatively recent faulting and uplift for the northern coast, probably no older than the Pliocene. Gradients of the southern shelf
All documented Neogene faults on St. Croix are normal faults, including the graben-bounding faults, and strike in a northeasterly direction, oblique to the south margin of the Virgin Islands basin. On the St. Croix Ridge, seismic profiles and GLORIA imagery (Masson and Scanlon, 1991) indicate that the ridge is broken into a series of block-fault 'piano key' structures with the same northeasterly orientation as the St. Croix faults (Fig. 13). These structures are interpreted to be the products of normal faulting similar to those in the Neogene section of St. Croix (Holcombe, 1977). The northeasterly orientation of the apparently continuous set of tensional fractures is consistent with the type of deformation expected in a left-lateral wrench-fault zone aligned along the Anegada Passage. Such fractures tend to form parallel to the short axis of the strain ellipse in clay models (Wilcox et al., 1973). The northeasterly fault orientation is inconsistent with right-lateral strike-slip movement, which would produce normal faulting of the opposite orientation. Alternately, a model similar to that of Geist et al. (1988) may apply here. The consistent orientation of the St. Croix and St. Croix Ridge fracture system implies a tensional
358
I. GILL et al.
Fig. 12. Subsidence analysis for the Neogene of the Kingshill basin. Upward trend of tectonic subsidence curve indicates uplift beginning at approximately 10.5 Ma, coincident with initiation of formation of the Kingshill basin boundary faults. Tectonic subsidence is calculated based on stratigraphic thicknesses, age, and paleobathymetry of each unit, decompacted thickness of the strata based on lithology, and removing the effects of sediment and water loading. The sea floor line represents the position of the sea bottom based on paleobathymetry with eustatic contributions removed. origin under a consistent tectonic regime (Fig. 13). The sharply defined walls of the Virgin Islands basin, including the north slope of St. Croix, suggest relatively recent tectonic activity along this area. Assuming that the origin of this faulting is connected to movement in the Virgin Islands basin/Anegada Passage, the orientation of the fault system on St. Croix suggests left-lateral wrench faulting north of St. Croix beginning in the late part of the Middle Miocene. To the north of the Virgin islands basin, the islands of St. Thomas and St. John both show extensive faulting. Donnelly (1966) mapped a graben structure on both islands that also strikes northeasterly and displaces Cretaceous rocks. Left-lateral strike-slip displacement is apparent within the graben. Although these faults may pre-date St. Croix strata, they show a similar orientation.
Sediment sources for the Kingshill basin Because no uplifted horst blocks were available as sediment sources before the Middle Miocene formation of the Kingshill basin graben, an external source of coarse clastics is required to explain the significant quantities of conglomeratic deposits present in the type section of the Jealousy Formation (Cederstrom, 1950). The southeastern part of Puerto Rico contains exposures of Tertiary carbonates that extend eastward of Puerto Rico only as far as the southern coastline of Vieques Island (Khudoley and
Meyerhoff, 1971). This area is presently more than 60 km to the northwest of St. Croix. Given these distances, St. Croix was probably much closer to these potential sources of coarse clastics during the Middle Miocene than it is today. Such a reconstruction would require tens of kilometers of left-lateral movement to place St. Croix in its present position. A second potential sediment source is the Anguilla/Saba Bank area to the east. Shelf carbonates of the same age as the Jealousy Formation and the Kingshill Limestone exist in Anguilla (Van den Bold, 1970) and contain very similar ostracode faunas. Saba Bank is underlain by rocks interpreted to be early Tertiary carbonates and fluvio-deltaics (Nemec, 1980; Warner, 1989) with a possible sediment source on Puerto Rico. If St. Croix was originally juxtaposed with either of these areas, fault motion would also be left lateral, assuming movement crudely parallel to the Anegada Passage. Lateral movement from Anguilla to the present location of St. Croix would require a greater travel distance than would movement from Puerto Rico. Unfortunately, information on the basins between St. Croix and Saba Bank is too sparse to allow more thorough evaluation. Interestingly, the problem of a sediment source exists for Cretaceous St. Croix as well. Conglomeratic deposits of rudistid bivalves are found within deep marine sedimentary rocks on St. Croix (Whetten, 1966). The closest documented source for these
E V O L U T I O N OF THE N E O G E N E K I N G S H I L L BASIN OF ST. CROIX, U.S. V I R G I N I S L A N D S
359
Fig. 13. Migration model for Tertiary St. Croix and opening of the Virgin Islands basin showing hypothetical positions for St. Croix between the Early Miocene and Recent, assuming an initial position south of Vieques. The bottom diagram contrasts this model with an alternative model that assumes more of a north-south relative motion across the Virgin Islands basin. Position numbers for the same time periods are indicated with italics.
360 rudists of this type is Puerto Rico (H. Santos, pers. commun., 1998).
Seismicity Seismic activity today is detectable only in the shallow zones in the north wall from 0 to 50 km deep. These seismic events occur in swarms and are generally less than magnitude 3.2 (Frankel et al., 1980). Historic records indicate that the potential exists for much larger earthquakes in the Virgin Islands basin/Anegada Passage area. Two major earthquakes caused damage in both St. Croix and St. Thomas in 1867 and were calculated on the basis of tsunami arrival times to have originated in the north wall of the Virgin Islands basin (Reid and Tabor, 1920, cited in Frankel et al., 1980). In general, the Virgin Island basin and Anegada Passage are no more active than the Puerto Rico Platform, and do not show seismic patterns correlated with their bathymetry. To our knowledge, no fault-plane solutions have been calculated for the seismic events currently taking place in the Virgin Islands basin. For these reasons, present seismicity patterns do not support any one tectonic model or significant movement along the Anegada Passage today (J. Joyce, pers. commun., 1998). Tectonic context The Caribbean plate is interpreted to have an eastward present-day motion relative to the North American plate. This plate motion is generally manifested by sinistral slip along the northern Caribbean plate boundary, and by dextral slip zones along the southern boundary (Stephan et al., 1986). Active documented subduction in the northeastern Caribbean is presently taking place only along the Lesser Antilles arc. In the area near the Puerto Rico trench, the northern boundary of the Caribbean plate is characterized overall by slip but with some evidence for compression (Frankel et al., 1980; Burke et al., 1984). Given the modem left-lateral motion of the northern Caribbean plate boundary, it is reasonable to predict that the opening of the Virgin Islands basin between the Virgin Islands platform and St. Croix also reflects left-lateral movement. The consistency and simplicity of this model is perhaps the reason that left-lateral motion in the Anegada Passage was suggested by Hess (1933, cited in Whetten, 1966), Hess (1966), Burke et al. (1984, table 7), and in Case et al. (1984). Estimation of rifling rate For a model in which St. Croix moves from a position south of Vieques, we estimate that the Virgin Islands basin has opened with rate of lateral motion of approximately 8 mm/yr since the late Middle Miocene. This rate is based on an assumed
I. GILL et al. age of 11 Ma for initiation of motion and a distance of lateral movement of 91 km. This rate is somewhat slower than the estimated 20 m m / y r or greater rate of movement between the North American and Caribbean plates cited by Rosencrantz and Mann (1991). This slower rate is consistent with the idea of Puerto Rico and the Virgin Islands platform being decoupled from the Caribbean plate (McCann et al., 1987) and moving eastward with a slower relative motion. The resulting differential motion between the Virgin Islands platform and the main body of the Caribbean plate, including St. Croix, may have caused the opening of the Virgin Islands basin.
Rotating platelet models An alternative group of models call for counterclockwise rotation of a Puerto Rico microplate or terrane. The idea of a separate Puerto Rico platelet was proposed by McCann et al. (1987), who suggested that this platelet had a westward (left-lateral) motion relative to the main Caribbean plate. Lithgow et al. (1987) suggested that the Virgin Islands basin formed as a pull-apart in response to these relative motion differences.
Seismic profiling and sidescan sonar data Based on seismic profiles and GLORIA sidescan sonar surveys (EEZ Scan Scientific Staff, 1987), Scanlon and Masson (1988) proposed that the Puerto Rico microplate has undergone counterclockwise rotation, with a pole of rotation south of Puerto Rico at the juncture between the St. Croix Ridge and the Muertos Trough (Fig. 14C). This model does not address relative motion between the microplate and the Caribbean and North American plates. Based on paleomagnetic studies in Puerto Rico, Reid et al. (1991) have documented approximately 25 ~ of counterclockwise rotation of Puerto Rico relative to North America between 11 and 4.5 Ma. They concluded that the Puerto Rico microplate behaved as a 'roller beating' between the North American and Caribbean plates during this period as it either became uncoupled from Hispaniola or responded to changes in relative plate motions. They suggested that this rotation ceased as the Puerto Rico microplate detached from the Caribbean plate and transferred extensional stress to the Anegada Passage and Mona Canyon. Speed and Larue (1991) proposed that the northern Caribbean PBZ has been in left-lateral transtension for the last 15 to 20 m.y., with much of the motion dispersed by counterclockwise rotation of terranes within the PBZ such as the Puerto RicoVirgin Islands terrane. In their model, as in the others, this rotation caused extension in the Anegada Passage/Virgin Islands basin area.
E V O L U T I O N OF THE N E O G E N E K I N G S H I L L BASIN OF ST. CROIX, U.S. V I R G I N ISLANDS
361
Fig. 14. Three models for the tectonic evolution of the northern Caribbean. (A) Plate tectonic model for the northern Caribbean with dextral slip in the Virgin Islands basin/Anegada Passage via a 'Puerto Rico Festoon' (after Stephan et al., 1986). Note that documented sinistral faults through Puerto Rico are not indicated in this diagram. The inset shows a mechanical analog for the formation of a 'Puerto Rico Festoon' with dextral slip to the east, i.e. Virgin Islands basin/Anegada Passage, and sinistral slip to the west (from Stephan et al., 1986). (B) Plate tectonic model for the northern Caribbean with dextral slip in the Virgin Islands basin/Anegada Passage. Note proposed triple junction to the southeast of Puerto Rico (after Jany et al., 1990). (C) Rotating microplate model of Scanlon and Masson (1988).
362
Discussion With an axis of rotation situated to the southwest of St. Croix (Scanlon and Masson, 1988), counterclockwise rotation of a separate Puerto Rico platelet would produce a zone of extension in the Anegada fault zone that widens to the northeast. In fact, the Anegada Passage has narrow, subparallel walls until it empties into the Sombrero basin to the northeast and into the Virgin Islands basin to the southwest. For this reason, the bathymetry of the region does not support simple rotation alone, and if rotation did occur, it may have been coupled to other motion. In any case, rotational models do not preclude either right or left-lateral movement between St. Croix and the Virgin Island platform. The direction of slip would be dependent on relative motion between the rotating Puerto Rico microplate and the Caribbean plate. If Puerto Rico rotated counterclockwise relative to a fixed or slowly eastward-moving Caribbean plate, this motion would most likely produce rightlateral slip in the Anegada Passage and Virgin Islands basin. However, if the rotation of Puerto Rico is accommodating part of the left-lateral movement between the North American and Caribbean plates, as proposed by Speed and Lame (1991), then the Caribbean plate would have a faster eastward relative motion. The result would be a left-lateral sense of motion in the Anegada Passage and Virgin Islands basin. The extensional nature of the northeasterly trending structures in that area supports the latter model.
Right-lateral models An alternative model for the origin of the Virgin Islands basin proposes that motion in the Virgin Islands basin and Anegada Passage is fight-lateral (Fig. 14A) (Houlgatte, 1983; Stephan et al., 1986). Mauffret et al. (1986) and Jany et al. (1987) propose that motion along the Anegada Passage was originally sinistral, but reversed during the Pliocene or later.
Basin morphology Jany et al. (1990) suggested that the Virgin Islands basin shows a rhomboidal 'lazy Z' shape (Fig. 14B), a shape indicated by Mann et al. (1983) to be diagnostic of fight-lateral shear zones. If strikeslip motion is parallel to the Anegada Passage, the oblique orientation of the Virgin Island basin is consistent with a fight-lateral pull-apart basin. However, although the geometry of the Virgin Islands basin and Anegada Passage is clear, the extension of the southwestern part the basin is not. Depending on the placement of the strike-slip zone to the southwest of the basin, the shape of the Virgin Islands basin could suggest opening of the basin by left-lateral motion.
I. GILL et al.
Seismic profiling Based on north-northeasterly oriented seismic profiles, Mauffret et al. (1986) have interpreted a structure on the north side of the Virgin Islands basin as a northward-verging reverse-fault zone. They cite this faulting as evidence for fight-lateral slip during the formation of the basin. However, if this structure trends normal (east-southeasterly) to this seismic profile, fight-lateral slip should produce extension rather than compression. Therefore, this structure could actually represent sinistral rather than dextral slip. Jany et al. (1990) have interpreted a similar structure - - perhaps the same one on an intersecting north-northwesterly oriented seismic line as a 'flower structure' indicative of strike-slip faulting. This structure overlies what is interpreted as the Late Miocene sedimentary surface. Although strikeslip motion may be a reasonable interpretation for the structures on these two lines, there is no way to assess accurately the sense of strike-slip motion for structures normal to such generally northerly oriented transects. Several studies have noted the predominance of generally easterly to northeasterly oriented normal faulting in the marine basins off St. Croix. Jany et al. (1987) identified northeasterly oriented normal faults on a northwesterly oriented seismic cross-section across the St. Croix basin east of St. Croix. In addition, several other seismic traverses published by Houlgatte (1983) similarly document normal faulting within the Virgin Islands basin. The sense of movement on these faults appears to be dominantly dip-slip; the degree of strike-slip motion is difficult to document. Although these data were related to a model for right-lateral movement in the plate boundary zone, the occurrence of northeasterly oriented extensional features in an easterly oriented strike slip zone is more consistent with left-lateral movement.
GPS data Global Positioning data exist for stations on Puerto Rico and St. Croix. Present data, although still preliminary, indicates that Puerto Rico may be moving to the east-northeast slightly faster than St. Croix (Dixon et al., 1998). These data support fightlateral displacement between St. Croix and Puerto Rico. However, the measurement error is presently too large to make reliable conclusions (P. Jansma, pers. commun., 1998).
Discussion The dextral slip model suffers from several problems when incorporated into a regional model of tectonics. Primary among these is the difficulty in reconciling a dextral strike-slip fault in the Vir-
EVOLUTION OF THE NEOGENE KINGSHILL BASIN OF ST. CROIX, U.S. VIRGIN ISLANDS gin Islands basin with the compression established along the length of the Muertos Trough (McCann et al., 1987). This would require an extension of the Anegada Passage/Virgin Islands basin fault zone westward of its present termination, and there is no seismic or bathymetric evidence to support this. Jany et al. (1990) suggest a triple-junction south of Puerto Rico to accommodate right-lateral plate movements (Fig. 14B). However, the deformation front along the Muertos Trough can be seen in the GLORIA data to extend nearly to the longitude of St. Croix, well east of where Jany's triple junction is shown (K. Scanlon, written commun., 1997). Dextral motion along the Anegada fault zone would also require Puerto Rico to move eastward faster than the Caribbean plate, and a driving mechanism for this movement requires a complicated model. Similarly, it is difficult to reconcile right-lateral motion in the Anegada Passage with established left-lateral motion for the northern Caribbean plate boundary zone (Stephan et al., 1986; Burke et al., 1984). The structural geology of the islands in the area is more consistent with left-lateral than rightlateral motion; easterly oriented terrestrial faults of Puerto Rico and the northern Virgin Islands are mapped with left-lateral displacement and the normal faults on St. Croix and on the St. Croix Ridge are oriented northeasterly. If dextral faulting is occurring in the Anegada fault zone, it must post-date the faulting on the Puerto Rico/Virgin Islands platform and on St. Croix; such a change in the latest Neogene or Quaternary would require a major reversal of plate motion in the northeastern Caribbean. If such a reversal did occur, it apparently left no trace in the rocks exposed on St. Croix, which record deposition and faulting through at least the Early Pliocene. Discussion of tectonic models
An oblique sinistral model for the opening of the Virgin Islands basin and the Anegada Passage satisfies structural evidence on St. Croix such as the northeasterly orientation of the normal fault system. It also permits the Kingshill basin to be paleogeographically situated near a known extra-basinal sediment source. Left-lateral motion, if oblique, provides a mechanism for the opening of the Virgin Islands basin. Left-lateral transtension across the Virgin Islands basin may disperse the slip between the North American and Caribbean plates. Our estimate of 8 mm/yr of movement across the Virgin Islands basin is slower than the 20 mm/yr rates of motion between the North American and Caribbean plates estimated by Rosencrantz and Mann (1991). Dextral motion is consistent with most interpretations of present basin morphology assuming
363
strike-slip movement directly along the Anegada Passage. In addition, recent GPS data may be more supportive of present fight-lateral displacement than with left-lateral models. In contrast, rotating platelet models are supported by paleomagnetic data, and are consistent with either right- or left-lateral movement in the Virgin Islands basin. We feel left-lateral motion is most consistent with the long-term evolution of the Virgin Islands basin area, but all the models have serious weaknesses. The only documented type of faulting between St. Croix and the Virgin Islands platform is extensional. Subsidence analysis in the Kingshill basin indicates that a period of rapid tectonic uplift took place between 10.5 and 3.5 Ma. The timing of this uplift is consistent with the relative plate motion changes between the Puerto Rico platelet and the Caribbean plate cited by McCann et al. (1987), the age of extension in the Virgin Islands basin noted by Lithgow et al. (1987), and tectonic changes north of Hispaniola (Dillon et al., 1992). The lithologies of the lower Kingshill Limestone suggest that uplift of the margins of the Kingshill basin graben may have begun within this time interval, and St. Croix may have rifted away from shelf areas of the Virgin Islands platform and Saba Bank. Further discussion of tectonic models involving the Virgin Islands platform is contained in van Gestel et al. (1998).
CONCLUSIONS
(1) St. Croix and the Virgin Islands basin were produced by transtension and associated uplift and subsidence along the northern Caribbean plate boundary zone during the Neogene. The stratigraphic and structural evolution of the Neogene Kingshill basin record a major tectonic event in this plate boundary zone near the Middle Miocene-Late Miocene transition. (2) The Jealousy Formation is a strictly subsurface unit of blue-gray marls deposited in a deepwater (600-800 m) environment before the onset of tectonism. All documented occurrences from well samples place this unit in the lower part of the Middle Miocene, although the formation may extend lower, perhaps into the Oligocene, below current depths of well penetration. (3) The lithologic transition between the Jealousy Formation and the Kingshill Limestone is abrupt and distinct, but is time-transgressive and does not appear to indicate any major mineralogic, faunal, or environmental change. (4) The Kingshill Limestone records the shallowing of the Kingshill basin and the initiation of uplift of basin-bounding horst blocks near the end of the Middle Miocene. There is an increase in
364 the abundance of coarse-grained debris in the middle of the formation and a progressive change to shallower-water microfaunas. Calculated uplift rates (backstripped, decompacted) for the Late Miocene to Early Pliocene are 57 m / M a . (5) St. Croix acquired its present shoreline configuration by the Pliocene, and an extensive reef and lagoon tract had established itself along the west and south shorelines of the island. (6) Structural control of the coastline in the form of a small graben allowed the accumulation and preservation of reef and platform Pliocene sediments along the south coast. Normal faulting has continued at least into the Pliocene. (7) The Virgin Islands basin is a structure formed by extension and sinistral faulting that rifted St. Croix away from Puerto Rico, the Virgin Islands platform and perhaps Saba Bank. Separation rates are estimated to have been 8 m m / y r . (8) Extension related to left-lateral plate motion is most consistent with the long-term evolution of the Virgin Islands basin area and St. Croix, but none of the regional tectonic models discussed is consistent with all the available data. The only solid tectonic evidence for formation of the Virgin Islands basin is extensive normal faulting. No definitive evidence for strike-slip motion in the Virgin Islands basin or Anegada Passage exists.
ACKNOWLEDGEMENTS
The authors thank Lee Gerhard, Nancy Grindlay, Barbara Lidz, Kathryn Scanlon, and Robert Speed for helpful, critical reviews of this paper, and Paul Mann for his constructive and patient editing. Pamela Jansma and James Joyce generously discussed various aspects of Caribbean tectonics and provided unpublished data. Clyde Moore served as major professor to the senior author throughout field and laboratory work. The aid and cooperation of the West Indies Laboratory, the U.S. Geological Survey Water Resources Division in San Juan, ER., and K. Eastman and the staff of Caribbean Geological Services in undertaking this study is gratefully acknowledged. The assistance of Marc Lowman with illustrations is also appreciated. Funding for St. Croix drilling and field work was provided to Gill by: the Virgin Islands Water Resource Center; SOHIO, Chevron, and Shell field research grants; Geological Society of America and American Association of Petroleum Geologists student grants; D. Eby and Champlin Petroleum; and the Applied Carbonate Research Program, the Department of Geology, and the Basin Research Institute at Louisiana State University. This work is part of a dissertation project of the senior author at Louisiana State University.
I. GILL et al. REFERENCES
Andreieff, R, Mascle, A., Mathieu, Y. and Muller, C., 1986. Les carbonates n6og~nes de Sainte Croix (Iles Vierges): 6tude stratigraphique et p6trophysique. Rev. Inst. Fr. Pet., 41" 336350. Andreieff, E, Bouysse, E and Westercamp, D., 1987. G6ologie de l'arc insulaire des Petites Antilles, et 6volution g6odynamique de l'est-Caraibe. Th~se de Doctorat d'Etat es Sciences, Universit6 de Bordeaux I, 465 pp. Behrens, G.K., 1976. Stratigraphy, Sedimentology and Paleoecology of a Pliocene Reef Tract: St. Croix, U.S. Virgin Islands. Unpubl. Masters thesis, Northern Illinois University, 93 pp. Bold, W.A. van den, 1970. Ostracoda of the lower and middle Miocene of St. Croix, St. Martin, and Anguilla. Caribbean J. Sci., 10:35-61. Bolli, H.M. and Saunders, J.B., 1985. Oligocene to Holocene low latitude planktic foraminifera. In: H.M. Bolli, J.B. Saunders and K. Perch-Nielsen, (Editors), Plankton Stratigraphy. Cambridge University Press, Cambridge, pp. 155-262. Burke, K., Cooper, C., Dewey, J.E, Mann, E and Pindell, J.L., 1984. Caribbean tectonics and relative plate motions. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162:31-63. Case, J.E., Holcombe, T.L. and Martin, R.G., 1984. Map of geologic provinces in the Caribbean region. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The CaribbeanSouth American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162: 1-30. Cederstrom, D.J., 1950. Geology and groundwater resources of St. Croix, U.S. Virgin Islands. U.S. Geolog. Surv. Water Supply Pap., 1067, 117 pp. Cushman, J.A., 1946. Tertiary foraminifera from St. Croix, Virgin Islands. U.S. Geol. Surv. Prof. Pap. 210-A, 17 pp. Dill, R.E, 1977. Deep water erosional feature, bedrock of St. Croix, U.S. Virgin Islands, as seen from the research submerible Alvin. 7th Caribbean Geol. Conf., Abstr., Addendum. Dillon, W.E, Austin, J.A., Jr., Scanlon, K.M., Edgar, N.T. and Parson, L.M., 1992. Accretionary margin of north-western Hispaniola: morphology, structure, and development of the northern Caribbean plate boundary. Mar. Pet. Geol., 9: 70-88. Dixon, T.H., Farina, E, Demets, C., Jansma, E, Mann, E and Calais, E., 1998. Relative motion between the Caribbean and North American plates and related boundary zone deformation from a decade of GPS observations. J. Geophys. Res. B, 103: 15,157-15,182. Donnelly, T.W., 1966. Geology of St. Thomas and St. John, U.S. Virgin Islands. In: H.H. Hess (Editor), Caribbean Geological Investigations. Geol. Soc. Am. Mem., 98: 85-176. EEZ Scan Scientific Staff, 1987. Atlas of the U.S. Exclusive Econonic Zone, Gulf of Mexico and Eastern Caribbean Areas. U.S. Geol. Surv. Misc. Invest. Ser., I- 1864-A, B, 171 pp. Frankel, A., McCann, W.R. and Murphy, A.J., 1980. Observations from a seismic network in the Virgin Islands region: tectonic structures and earthquake swarms. J. Geophys. Res., 85 (B5): 2669-2678. Frost, S.H., 1977. Cenozoic reef systems of the Caribbean m prospects for paleoecologic synthesis. In: S.H. Frost, M.E Weiss and J.B. Saunders (Editors), Reefs and Related Carbonates - - Ecology and Sedimentology. Am. Assoc. Pet. Geol., Stud. Geol., 4: 93-110. Geist, E.L., Childs, J.R. and Scholl, D.W., 1988. The origin of summit basins of the Aleutian Ridge: implications for block rotation of an arc massif. Tectonics, 7:327-341. Gerhard, L.C., Frost, S.H. and Curth, EJ., 1978. Stratigraphy and depositional setting, Kingshill Limestone, Miocene, St. Croix, U.S. Virgin Islands. Am. Assoc. Pet. Geol. Bull., 62:403-418.
E V O L U T I O N OF THE N E O G E N E K I N G S H I L L BASIN OF ST. CROIX, U.S. V I R G I N ISLANDS Gill, I.E, 1983. The Sedimentological Controls on Organic Carbon in the Virgin Islands Trough, U.S. Virgin Islands. Unpubl. MS thesis, Univ. of Rochester, 70 pp. Gill, I.E, 1989. The Evolution of Tertiary St. Croix. Unpubl. Ph.D. dissertation, Louisiana State University, 287 pp. Gill, I.E and Hubbard, D.K., 1986. Subsurface geology of the St. Croix carbonate rock system. Caribbean Research Institute Technical Report 26, College of the Virgin Islands, 86 pp. Gill, I.E and Hubbard, D.K., 1987. Subsurface geology of the St. Croix carbonate rock system, phase II. Technical Report No. 28, Water Resources Research Center, College of the Virgin Islands, St. Thomas, U.S. Virgin Islands, 79 pp. Gill, I.E, Hubbard, D., McLaughlin, EE and Moore, C., 1990. Sedimentological and tectonic evolution of Tertiary St. Croix. In: D.K. Hubbard (Editor), Terrestrial and Marine Geology of St. Croix, U.S. Virgin Islands. Special Publication No. 8, West Indies Laboratory, Teague Bay, St. Croix, U.S. Virgin Islands, pp. 49-72. Gill, I.E, Hubbard, D., McLaughlin, EE and Moore, C.H., in press. Geology and hydrogeology of the St. Croix carbonate aquifer system. In: R. Renken (Editor), Aquifers of the U.S. Virgin Islands and Puerto. U.S. Geol. Surv., Prof. Pap., PP1419-A (30 ms. pages, 10 figs). Haq, B.U., Hardenbol, J. and Vail, ER., 1988. Mesozoic and chronostratigraphy and cycles of sea level change. In: C.K. Wilgus, B.S. Hastings, C.G.St.C. Kendall, H.W. Posamentier, C.A. Ross and J.C. Van Wagoner (Editors), Sea-Level Changes: An Integrated Approach. Soc. Econ. Paleontol. Mineral., Spec. Publ., 42: 71-108. Hess, H.H., 1933. Interpretation of geological and geophysical observations. U.S. Hydrological Office Navy-Princeton gravity expedition to the West Indies in 1932, pp. 27-54. Hess, H.H., 1966. Caribbean research project, 1965, and bathymetric chart. In: H.H. Hess (Editor), Caribbean Geological Investigations. Geol. Soc. Am. Mem., 98: 1-10. Holcombe, T.L., 1977. Geomorphology and subsurface geology west of St. Croix, U.S. Virgin Islands. Am. Assoc. Pet. Geol. Mere., 29: 353-362. Houlgatte, E., 1983. Etude d'une partie de la fronti6re nord-est de la plaque Caraibe. Unpubl. Masters thesis, L'Universite de Bretagne Occidentale, 69 pp. Hubbard, D.K., Suchanek, T.H., Gill, I.E, Cowper, S., Ogden, J.C., Westerfield, J.R. and Bayes, J., 1981. Preliminary studies of the fate of shallow-water detritus in the basin north of St. Croix, USVI. Proc. 4th Int. Coral Reef Symp., Manila, 1, pp. 383-387. Jany, I., Mauffret, A., Bouysse, E, Mascle, A., Mercier de Lepiany, B., Renard, V. and Stephan, E, 1987. Releve bathymetrique Seabeam et tectonique en decrochements au sud des Iles Vierges (Nord-Est Caraibes). C.R. Acad. Sc. Paris, 304, Ser. II, 10: 527-532. Jany, I., Scanlon, K.M. and Mauffret, A., 1990. Geological interpretation of combined Seabeam, GLORIA and seismic data from Anegada Passage (Virgin Islands, North Caribbean). Mar. Geophys. Res., 12: 173-196. Khudoley, K.M. and Meyerhoff, A.A., 1971. Paleogeography and geological history of Greater Antilles. Geol. Soc. Am. Mem., 129, 199 pp. Lidz, B.H., 1982. Biostratigraphy and paleoenvironment of Miocene-Pliocene hemipelagic limestone, Kingshill Seaway, St. Croix, U.S. Virgin Islands. J. Foram. Res., 12: 205-233. Lidz, B.H., 1984a. Neogene sea-level change and emergence, St. Croix, Virgin Islands: evidence from basinal carbonate accumulations. Geol. Soc. Am. Bull., 95:1268-1279. Lidz, B.H., 1984b. Oldest (early Tertiary) subsurface carbonate rocks of St. Croix, USVI, revealed in a turbidite mudball. J. Foram. Res., 14:213-227. Lidz, B.H., 1988. Upper Cretaceous (Campanian) and Cenozoic
365
stratigraphic sequence, northeast Caribbean (St. Croix, U.S. Virgin Islands). Geol. Soc. Am. Bull. 100: 282-298. Lithgow, C., McCann, W.R. and Joyce, J., 1987. Extensional tectonics at the eastern edge of the Puerto Rico Platelet (abstr.). Eos, 68 (44): 1483. Mann, E, Hempton, M.R., Bradley, D.C. and Burke, K., 1983. Development of pull-apart basins. J. Geol. 91: 529-554. Masson, D.G. and Scanlon, K.M., 1991. The neotectonic setting of Puerto Rico. Geol. Soc. Am. Bull., 103:144-154. Mauffret, A., Jany, I., Mercier de Lepinay, B., Bouysse, E, Mascle, A., Renard, V. and Stephan, J.-F., 1986. Releve au sondeur multifaisceaux du bassin des Iles Vierges (extremit6 orientale des Grandes Antilles): role de l'extension et des ddcrochements. C.R. Acad. Sci. Paris, 303, Ser. II, 10: 923928. McCann, W.R., Joyce, J. and Lithgow, C., 1987. The Puerto Rico Platelet at the northeastern edge of the Caribbean Plate. Eos, 68 (44): 1483. McLaughlin, EE, Gill, I.E and Bold, W.A. van den, 1995. Biostratigraphy and paleoenvironments of the Neogene of St. Croix, U.S. Virgin Islands: implications for stratigraphic evolution. Micropaleontology, 41 (4): 293-320. Meyerhoff, H.A., 1927. The physiography of the Virgin Islands, Culebra and Vieques. New York Academy of Science, Scientific Survey of Porto Rico and the Virgin Islands, Vol. 4, no. 2, pp. 152-156. Moore, C.H., Jr., Graham, E.A. and Land, L.S., 1976. Sediment transport and dispersal across the deep fore-reef and island slope ( - 5 5 to -305 m), Discovery Bay, Jamaica. J. Sediment. Petrol., 46:174-187. Multer, H.G., Frost, S.H. and Gerhard, L.C., 1977. Miocene 'Kingshill Seaway' - - a dynamic carbonate basin and shelf model, St. Croix, U.S. Virgin Islands. In: S.H. Frost, M.E Weiss and J.B. Saunders (Editors), Reefs and Related Carbonates--Ecology and Sedimentology. Am. Assoc. Pet. Geol., Stud. Geol., 4: 329-352. Nemec, M.C., 1980. A two-phase model for the tectonic evolution of the Caribbean. 9th Caribbean Geol. Congr. Trans., 2: 23-34. Reid, H. and Tabor, S., 1920. The Virgin Islands earthquakes of 1867-1868. Bull. Seismol. Soc. Am., 10: 9-30. Reid, J.A., Plumley, EW. and Schellekens, J.H., 1991. Paleomagnetic evidence for late Miocene counterclockwise rotation of north coast carbonate sequence, Puerto Rico. Geophys. Res. Lett., 18 (3): 565-568. Rosencrantz, E. and Mann, E, 1991. SeaMARC II mapping of transform faults in the Cayman Trough, Caribbean Sea. Geology, 19 (7): 690-693. Scanlon, K.M. and Masson, D., 1988. Seafloor deformation at the northern Caribbean Plate boundary and rotation of a Puerto Rico microplate (abstr.). Eos, 69 (16): 462. Shurbet, G.L., Worzel, J.L. and Ewing, M., 1956. Gravity measurements in the Virgin Islands. Geol. Soc. Am. Bull., 67: 1529-1536. Speed, R.C. and Larue, D.K., 1991. Extension and transtension in the plate boundary zone of the NE Caribbean. Geophys. Res. Lett., 18: 573-576. Stephan, J.F., Blanchet, R. and Mercier de Lepinay, B., 1986. Northern and southern Caribbean festoons (Panama, Colombia-Venezuela and Hispaniola-Puerto Rico), interpreted as pseudo subductions induced by the east-west shortening of the peri-Caribbean continental frame. In: E-C. Wezel (Editor), The Origin of Arcs. Elsevier, New York, pp. 401422. van Gestel, J.E, Mann, E, Dolan, J. and Grindlay, N.R., 1998. Structure and tectonics of the upper Cenozoic Puerto RicoVirgin Islands carbonate platform as determined from seismic reflection studies. J. Geophys. Res., 103: 30,505-30,530.
366 Warner, A.J., 1989. The Cretaceous age sediments of the Saba Bank and their petroleum potential. Trans. 12th Caribbean Geol. Conf., pp. 341-354. Whetten, J.T., 1966. The geology of St. Croix, U.S. Virgin Islands. Geol. Soc. Am. Mem., 98: 177-239. Whetten, J.T., 1974. Field guide to the geology of St. Croix.
I. G I L L et al. In: H.G. Multer and L.C. Gerhard (Editors), Guidebook to the Geology and Ecology of Some Marine and Terrestrial Environments, St. Croix, U.S. Virgin Islands. West Indies Lab. Spec. Publ., 5: 129-143. Wilcox, R.E., Harding, T.E and Seely, D.R., 1973. Basic wrench tectonics. Am. Assoc. Pet. Geol. Bull., 57 (1): 96.
C h a p t e r 14
Review of the Tectonic Controls and Sedimentary Patterns in Late Neogene Piggyback Basins on the Barbados Ridge Complex
PASCALE HUYGHE, JEAN-LOUIS MUGNIER, ROGER GRIBOULARD, YANN DENIAUD, E L I A N E G O N T H I E R and J E A N - C L A U D E F A U G E R E S
A review of the tectonic control and sedimentary patterns of the late Neogene piggyback basins of the Barbados Ridge Complex is proposed, mainly based on the analysis of seismic reflection profiles and sidescan sonar images. The geometrical relationships between the thrust faults and the syntectonic deposits show that the frontal piggyback basins are controlled by active thrusting. Deformations, with distinct scales, control their evolution: (1) a rapid tilting, strictly localized at the back limb of anticlines, and attributed to migration of active axial surfaces during fault-bend fold propagation; (2) a complex activation of the major thrust system, at a scale of a few kilometers: blind thrusts corresponding to frontal propagation develop seaward whereas motion along thrusts occurs backward. At about 50 km back of the deformation front, abundant muddy material raises up through pre-existing faults and disturbs arcward piggyback basins. The superimposition of the diapiric structures upon deformations linked to tectonic accretion (development of backthrusts and reactivating of forward verging thrusts) and inherited oceanic basement ridges leads to the individualized development of sub-basins bounded by steep topographic features. Most of the sediments of the South Barbados piggyback basins originate from the South American continent and are massively transported to the abyssal plain through canyons. Their course is driven by the main regional structures and their morphology reflects the tectonic activity of the features where they run. Oceanic ridges, by damming and collecting turbidite material control the thickness of sediments added to the complex and then the depth of the decollement and size and filling of the piggyback basins. Tectonics, by generating routes along the faults and excess fluid pressures control the main location and importance of mud diapirs and authigenic deposits. These stiff carbonate crusts preferentially develop on diapiric domes or on the back limb of anticlines up to the edge of supra-prism basins. Clayey diapiric material may be found within the basins where they form important mud flows as well as sliding masses. Tectonic and diapiric structures control gravity reworking processes, whereas structural relieves locally disturb bottom currents and hence control some erosion processes. The development of the piggyback basins of the BRC is closely linked to the evolution of a thrust wedge. The formation of frontal basins is mainly controlled by a forward-verging thrust system that forms a brittle wedge, whereas the development of the arcward basins is mainly controlled by subcretion of deep muds that induces mud diapirism, ductile deformation in the lowermost part of the wedge, a regional gentle slope, and the occurrence of both backward and forward verging thrusts. The evolution of the piggyback basins of the BRC also reflects the north to south changes in width and thickness of the wedge which are mainly related to variations of the increase of the sediment supply. Piggyback basins then evolve from minor depressions filled with very poor sediments in the north to about 10 km long overfilled basins in the south of the complex where abundant sedimentation occurs.
INTRODUCTION
P i g g y b a c k basins d e v e l o p in the external part of o r o g e n i c belts, at the e d g e of the f o r e d e e p basin. Ori and F r i e n d (1984) defined t h e m as basins that f o r m and fill w h i l e b e i n g carried on the b a c k of m o v i n g thrust sheets (Fig. 1). P i g g y b a c k basins are characterized by a s y n t e c t o n i c filling and their g e o m e t r y evolves with g r o w t h of the a c c r e t i o n a r y w e d g e . Migration of the d e p o c e n t e r s of the s y n t e c t o n i c filling furnishes a r e c o r d of the tilting of the thrust sheets
and p r o p a g a t i o n of n e w thrusts ( R o u r e et al., 1990; Z o e t e m e i j e r and Sassi, 1992). T h e s o u t h e r n L e s s e r A n t i l l e s (Fig. 2) is a g o o d place to study the e v o l u t i o n of p i g g y b a c k basins as h i g h s e d i m e n t influx and active faulting o c c u r in the B a r b a d o s R i d g e C o m p l e x (BRC). H i g h s e d i m e n t influx p r o v i d e s an a c c u r a t e stratigraphic r e c o r d of the history of basins d u r i n g late N e o g e n e t i m e and g e o p h y s i c a l t e c h n i q u e s p e r m i t the o b s e r v a t i o n of the p i g g y b a c k basin f o r m a t i o n . This study m a i n l y c o n c e r n s the basins that develop, f r o m the toe of
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 369-388. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
370
R HUYGHE et al. PIGGYBACK BASIN
POSSIBLE BY PASSING
I
ONLAP
SHEETS ~ < ~ ~ ~ - - -
-
MAIN BASIN
" .....
THRUST
SEQUENCECONTEMPORANEOUS
WITHTHRUSTING
NOT TO SCALE
Fig. 1. A sketch of piggyback basin (modified from Ori and Friend, 1984). the accretionary complex to the forearc basin (Fig. 3 adapted from Westbrook et al., 1988) and which were recognized as piggyback basins (Mascle et al., 1990; Huyghe et al., 1996). The aim of this paper is to present the sedimentary pattern of the piggyback basins in the Southern Barbados Ridge Complex (SBRC). We present here single-channel seismic (SCS), 3.5 kHz and sidescan sonar (SAR) data which were acquired in 1987 and 1990 during French surveys with the NO Nadir and Suroit (Faugbres et al., 1991; Griboulard et al., 1991; Faug~res et al., 1993; Huyghe et al., 1996). As the evolution of piggyback basins is representative of the history of thrust-sheet movements, these
data also provide constraints upon the growth of the accretionary wedge.
GEOLOGIC SETTING The BRC (Fig. 2) develops as a result of the subduction of the North American (HA) and South American (SA) plates beneath the Caribbean plate (CA) since the Middle Eocene (Westbrook, 1982; Stephan et al., 1990). The movement of the Caribbean plate relative to the South American plate is complex: obliquity of convergence (Jordan, 1975; Stein et al., 1982; Speed, 1985; DeMets et al., 1990;
Fig. 2. Geodynamic setting of the eastern Caribbean. Solid arrows represent the Orinoco delta and major continental influxes. Open arrows show relative motion of the Caribbean plate with respect to the South American plate (from DeMets et al., 1990). Hatched area: collision and/or strike-slip motion between the Caribbean plate and South American margin (from Biju-Duval et al., 1982); dashed line: limit between continental crust and crust of uncertain nature (from Bouysse and Westercamp, 1988).
LATE NEOGENE PIGGYBACK BASINS ON THE BARBADOS RIDGE COMPLEX
c~
,.o (D 9
E 9
z ,.o
(D
E 9
9 ...., ,.o (D
(D ~
9
(D
9
9 ~ (D ~
~
9
~
~D ~
.1 e4 ~
371
Russo et al., 1993; DeMets et al., 1994), transtension (Sykes et al., 1982) and intraplate deformations (Weber et al., 1995) are still under discussion. In the southern part of the prism, these movements may lead to the development of additional structures such as transcurrent faults (Robertson and Burke, 1989, 1991; Griboulard et al., 1991). Moreover, the boundary between continental and oceanic crusts is not firmly established in the southern part of the complex (Bouysse and Westercamp, 1988). Large volumes of sediments coming mainly from the Orinoco and Amazon systems (Damuth, 1975; Leonard, 1983) are involved in the frontal accretion at the toe of the Barbados wedge. As a result of the southern location of these sources, the B RC narrows from 450 km at its southern extremity to 100 km north of 18~ This variation in width is associated with a change of sediment thickness on the Atlantic ocean floor from 7000 m south of l l~ to 200 m in the north at 19~ (Ewing et al., 1973; Peter and Westbrook, 1976; Mascle and Moore, 1988; Moore and Mascle, 1990). This decrease is also locally controlled by the occurrence of N 140~ ridges which act as barrage to the sediment transport proceeding from south to north (Westbrook et al., 1984; Wright, 1984). These structures have been evidenced by gravity, bathymetry and seismic reflection data (Birsh, 1970; Peter and Westbrook, 1976) and probably represent remnant transform faults in the oceanic basement (Fig. 3) that generate troughs and asymmetrical flanking ridges (Westbrook et al., 1984). Such features induce lateral ramps and lead to the global eastward shift of the deformation front (Calassou et al., 1993; Huyghe et al., 1996). The best known of these are the Barracuda Ridge and the Tiburon Rise (Fig. 2) which still affect the abyssal plain. Mud diapirism is a widespread phenomenon on the SBRC (Fig. 4) (Fontas et al., 1985; Brown and Westbrook, 1988; Langseth et al., 1988; Le Pichon et al., 1990). It is linked to high pore fluid pressures favored by the rapid tectonic burial of undercompacted sediment at the deformation front (Bangs and Westbrook, 1990), the high horizontal compression occurring in a prism and by subcretion that releases fluids into the accretionary complex (Brown and Westbrook, 1988). In the SBRC, most of the diapiric features seem to be controlled by tectonic structures such as faults that provide routes to the surface for mud (Biju-Duval et al., 1982; Brown and Westbrook, 1988). Depending on the viscosity of mud, diapiric forms vary from mud volcanoes fed by low-viscosity mud and generating flows, to higher-viscosity mud ridges or domes forming prominent relieves (Shih, 1967; Brown and Westbrook, 1988). This mud is probably provided by Miocene beds as onshore in Trinidad (Higgins and Saunders, 1974). Propagation
372
E HUYGHE et al.
Fig. 4. Regional map (adapted from Brown and Westbrook, 1987 and Fontas et al., 1985). Numbers refer to piggyback basins shown in Figs. 8, 10 and 16. Color version at http://www.elsevier.nl/locate/caribas/ of excess pore pressure from the decollement below the accretionary complex seaward into the undeformed ocean basin sediments also generates mud diapir fields and venting seaward of the deformation front (Westbrook and Smith, 1983; Silver et al., 1986; Langseth et al., 1988; Henry et al., 1990; Lallemant et al., 1990). In the SBRC, the thick sedimentary series lying above the decollement are intersected by a network of submarine canyons feeding the Orinoco deep-seafan (Belderson et al., 1984; Faug~res et al., 1993). The canyons generally run northward (Embley and Langseth, 1977; Mascle et al., 1990) but their course is also driven by the main regional tectonic features (Fig. 4) and changes from NE to ENE and ESE. At a smaller scale, the course of the canyons is strongly influenced by structures having prominent relief on the seafloor (Mascle et al., 1990), such as anticlines and mud diapirs so that the course of the canyons appears sinuous (Fig. 5). The coarsest turbiditic material (medium sands and gravels) coming from the Guyana Margin is funnelled down to the Atlantic Abyssal Plain through major canyons (Fig. 4). The finest material (fine sands and clay)
overflows the channels and deposits mainly in sedimentary basins contemporaneous with the growth of the SBRC (Biju-Duval et al., 1982; Mascle et al., 1990). As a consequence the late Neogene filling of basins consists of turbiditic layers intercalated with fine hemipelagic muds (Faug~res et al., 1993). The stratigraphic pattern of these syntectonic deposits and their geometrical relationships with faults permit to distinguish two main types of basins that are described below.
FRONTAL BASINS
Frontal basins develop in the zone of initial accretion (Fig. 3) defined by Westbrook et al. (1984). They form between major anticlinal ridges from the wedge toe to 50 km west of the deformation front. The folds delimiting the frontal basins develop above major thrust structures (Biju-Duval et al., 1982) and may be considered as thrust fault related folds above a major decollement (Fig. 6) (Mascle et al., 1990). On the SBRC, spacing of the folds evolve from 5-6 km at 12~ (Fig. 7) to 8-10 km
LATE N E O G E N E P I G G Y B A C K BASINS ON THE BARBADOS RIDGE C O M P L E X
~-~
BASIN~ANYON
"~'~
~THE
---
.~
~_1 ~r
~-~BASlN
~ "
373
BACK LiMB OF THE ANTICLINE--)~
CANYON |~-".,.~'-.~"~'~
"
I:',.-:,~'.:-:~.~
'
"" ,~"1,O~ ~~~ d ~ ~ ' ~~ - ~ )~"
,~'fl' ~ {.., c...
"
',,~'-~r,,,,' ~ ~2~'\ ~I~--.~-~SlEDIMENTARY LEVEES
':': " , " ;,~:~_'~~ --II~~.~,: o,~ ~STRUCTURAL ~ \ .~':-!:-;:]/~!.-"t~ !~ LINEATIONS """
~"
:.-..~:, ',
4'z
'
.
S_s___ "~ ~
,....,
. . . .
I
.
-
6oo M
-
i
";"-"
9
u
'
"
~'
9
",v"
r
-,
.g,:.::.~"~
:,'""
,,..:,.'.~,,
-
:/'.'~."
,':.'~
,4",:
~ 2 . ~
"
.._~.
.,u~,,;"~/ .r~".'~-'~ ~ .-~tt.-~',: ~ - 9 ~.* " ~* ' ~ ~ S Z l F F ""Y
CRUSTS
Fig. 5. Line drawing of SAR images representing the back limb of a fault related fold and the edge of the associated piggyback basin. Near the anticline, the canyon changes direction. Its meanderings are disturbed by tectonic structures which also affect the back limb of the fold covered with stiff crusts (dotted areas).
KM W -4
E FOLDSPACING I-.1KM--.I
5 0 0
I
I
I
I
2
4
6
8
15~ N
I
10
KM NW .2
KM SE
I
-4 -6 -8
0
5
:10
15
20
KMI NW
FOLDSPACING
- 2 -t
I
8 KM
2'5
;30
3,5
z~0 KM SE 10~ N
I
"4- J
-6 -8
0
5
10
15
20
25
30
35
40 KM
Fig. 6. Balanced cross-sections of the Barbados Ridge Complex (from Endignoux and Mugnier, 1990): (a) at the latitude 15~ (CRV 128, Leg 110 transect, rate of shortening 25%); (b) at latitude 10~ (CRV 108 A, rate of shortening 15%); and (c) at latitude 10~ (CRV105, rate of shortening 5%).
374
E HUYGHE et al.
Fig. 7. (a) Diapicar UV water-gun seismic profile (non migrated) at 12~ sediments onlapping the frontal sheet. at 10~ (Fig. 8) and controls the width of basins (Endignoux and Mugnier, 1990; Mascle et al., 1990) respectively varying from a few thousand meters to 6-8 km. The frontal piggyback basins show a mainly N10~176 direction parallel to the frontal thrusts. Their filling reaches up to 1 s TWT at about 10~ (Fig. 8) and shows a wedge-shaped geometry characteristic of syntectonic deposits of piggyback basins undergoing active thrusting.
Stratigraphic pattern at the edge of the basins Inner edge The inner edge of the piggyback basins is formed by the thrust fault bounding the nearby backward
(b) Detail of (a) showing that thrust fault 1 cuts through
tectonic sheet. Relationships between this thrust fault and syntectonic filling is difficult to observe because of 'artefacts' or slumps linked to the slope. Fig. 7 shows the relationships between a major thrust fault (thrust 1) and the sediments deposited at its footwall. Recent sediments onlap the frontal thrust sheet and postdate its tilting. As thickness of these sediments is continuous and regular over 1 kin, their origin has not to be linked to gravity sliding from the above westside fold related fault. Thrust 1 cuts through the sediments onlapping the frontal sheet. Its recent propagation is consequently subsequent to that of the frontal thrust and shows that reactivation occurs backward from the deformation front.
LATE NEOGENE PIGGYBACK BASINS ON THE BARBADOS RIDGE COMPLEX
375
Fig. 8. Diapicar KL water-gun seismic profile (non migrated) at 10~ Horizontal scales are the same on the seismic profile and line drawing. Note the presence and gentle size of a piggyback basin in the zone of initial accretion, whereas a smaller piggyback basin is observed at 12~ (Fig. 7). Dots represent depocenters.
Outer edge The outer edge of the piggyback basin is formed by folds that have frequently kink-shaped geometries with rounded hinges. Interlimb angles of about 160~ have locally been evaluated in the stratigraphically higher strata (Fig. 9). Numerous minor tectonic features affect the limb of the anticlines but there is no major cliff of the sea floor delimiting the edge of the basins as shown by S AR images. The basin edge is only marked by a change of sediment reflectivity (Fig. 5) linked to stiff crusts covering the back limb of the fold and soft sediments within the basin. 3.5 kHz seismic profiles have a resolution better than 5 m and evidence the characteristics of recent strata. The stratigraphic pattern at the edge of the basin shows the following (Fig. 9): (1) a succession of five sedimentary bodies with a wedge-shaped geometry deposited on the 15~ back limb of the anticline (circled numbers in Fig. 9b); (2) the individual body itself thickens towards the center of the piggyback basin; (3) an increase in dip from stratigraphically higher to lower strata is seen in the syntectonic sedimentary bodies depositionally above the back limb; (4) angular disconformities occur between sedimentary bodies in the thinned domain at the back limb of the anticlinal ridge; (5) surfacial
minor thrusts localized at the edge of the basin cut through recent sediments. The pattern described above is characteristic of the growth of fault related folds (Suppe et al., 1992) and the syntectonic strata record the evolution of the fold relative to the history of sedimentation (Burbank and Verg6s, 1994). Growth folding implies axial surfaces migration (Suppe, 1983) and involves progressive rotation of the back limb of the anticlines. From the angular relationships (Fig. 9) and the average of Quaternary sedimentation rate, a mean tilting rate of 1~176 yrs has been estimated (Huyghe et al., 1996). Such tilting gives birth to angular unconformities which develop in growth strata localized on the back limb of the fold and reflects the distinct phases of fold growth. Sedimentary bodies 1, 2, 3 and 5 onlap progressively onto the back limb whereas stratigraphic offlap occurs for unit 4, a stratigraphic pattern that reveals that sediment accumulation exceeds uplift of the anticline as long as onlap occurs but that uplift suddenly and shortly exceeds accumulation during deposition of unit 4. This abrupt stage of fold growing may be responsible for minor displacement along thrusts in the syncline and slumping of unconsolidated material on the back limb of the rising anticline (Fig. 9).
376
E HUYGHE et al.
Fig. 9. (a) Diapisar 9036 seismic profile (3.5 kHz) showing a thrust fault related fold and the recent filling of the associated piggyback basin. Vertical exaggeration 5 x. Location in Fig. 5. (b) Line drawing and detail of (a). Note the presence of a thrust in the syncline. Grey dotted areas indicate slumps. Disconformities between the different sedimentary units 1 to 5 are related to the successive tilting of the ridge. Dips drawn under seismic interpretation take into account 5 x vertical exaggeration of seismic line. (c) Kink-shaped construction of the thrust fault related fold from the 3.5-kHz seismic data. Horizontal scale is the same as in (a). Interlimb angle is about 160~ Location of the western axial surface has not been determined. Occurrence of thrusts in the syncline in the piggyback basin is probably evidence for positioning it more westward. No vertical exaggeration.
Global stratigraphic pattern T h e global stratigraphic pattern of the syntectonic deposits of the p i g g y b a c k basins is a n a l y s e d f r o m the
filling of the frontal basins 1, 2 and 3 (see Fig. 4 for location) f r o m east to w e s t and s h o w n in Figs. 8 and 10. In every basin, the global b a c k w a r d m i g r a t i o n of d e p o c e n t e r s p r o v i d e s e v i d e n c e for a w e s t w a r d
LATE NEOGENE PIGGYBACK BASINS ON THE BARBADOS RIDGE COMPLEX
377
Fig. 10. Diapicar OP water-gun seismic profile (non migrated) and its interpretation showing the backward migration of depocenters in piggyback basins 2 and 3. Dots represent depocenters. Horizontal scales are the same on the seismic profile and line drawing.
tilting of the 7-10-km-long underlying thrust sheets. The geometrical relationships between syntectonic deposits and thrusts display a complex activation of the major thrust system: (a) development of the piggyback basin 2 (seismic units 2.1 and 2.2) linked to the growth of the eastern fault-bend fold (T2); (b) development of another sub-basin as a result of the propagation of a second ramp anticline (T2') west of the latter. In the main basin, depocenter 2.3 is temporally shifted eastward as a result of the formation of the fault related fold (T2'); (c) growing of the western anticline (T2') ceased. The depocenter of unit 2.4 is located east of the blind ramp anticline (T2'), as a result of its asymmetry. A large-scale backward tilting of the whole basin leads to the westward thickening of units 2.3 and 2.4. The successive onlap positions of the sedimentary bodies reflect changes in the evolution of the folds relative to sedimentation rate. This is clearly recorded in the thick filling of piggyback basin 3: seismic units 3.0 and 3.3 are associated with stages of an uplift rate higher than sedimentation rate.
the movement of nearby thrust faults inducing fold growing and the tilting of individual thrust sheets. The successive positions of the depocenters and onlap of the distinct sedimentary bodies indicate a complex activation of the major thrust system relative to sedimentation history. Stages of higher displacement rate take turn with increase of sedimentation rate and indicate out-off-sequence reactivation whereas blind thrusts corresponding to frontal propagation develop seaward under the abyssal plain (Mascle et al., 1990). Close-to-the-surface-motion results either in a whole ramp reactivation from decollement to the surface or in a passive re-adjustment linked to the imbrication of new frontal sheets (Huyghe et al., 1996). Finally, the stratigraphic pattern of frontal piggyback basins reveals a strong tectonic control: (1) a rapid tilting strictly localized at the back limb of anticlinal ridges; (2) a global backward migration of depocenters linked to the tilting of the whole sheet that they are carried on; (3) a complex activation of the major thrust system relative to the history of sedimentation.
Tectonic control of frontal piggyback basins
ARCWARD BASINS
Both the detailed and global patterns show that basin depocenters are dynamically controlled by
Piggyback basins develop in a more arcward position, at about 50 km in back of the deformation
378 front, where the bathymetry rises gently to the west, from water depths of 2000 m to 1400 m (Fig. 4). Numerous mud intrusions occur and the number of diapirs increases both southward and eastward in relation respectively with the extent and thickness of the Orinoco submarine fan on the abyssal plain and the onset of subcretion (Brown and Westbrook, 1988).
Stratigraphic pattern of arcward piggyback basins The sedimentary filling of these basins is thicker than that of frontal basins (1.5 to 2 s TWT). At about 12~ late Neogene sedimentary bodies show westward or eastward wedge-shaped geometries above steep faults (Fig. 11) recognized as thrusts and back-thrusts (Biju-Duval et al., 1982; Brown and Westbrook, 1987). The youngest sediments observed in Fig. 11 and also with 3.5-kHz seismic profiles locally overlap and seal the faults indicating that displacement occurs episodically. Southward, piggyback basins of about 150 km 2 show a complex arcuate shape (Fig. 4) and mud diapirism becomes more intensive. As fault planes form suitable routes to drain clay to the surface, arcward piggyback basins are bounded with diapiric structures (Figs. 12 and 13). Sedimentary bodies sometimes show wedge-shaped geometries thinning towards the borders (Fig. 12), but it is here difficult to assess if it is linked to the displacement along a thrust or back-thrust fault or to the progressive up-rising of mud. The last 0.5 s TWT of the filling is rather tabular and its local and moderate deformation seems to be related to the growing of the mud features. Recent mud intrusions through pre-existing or new forming faults lead to the individualisation of sub-basins between the diapiric edifices (Fig. 13). As a result, their sedimentary filling does not seem to result from tectonic subsidence but rather from borders uprising linked to fault tilting and to mud
E HUYGHE et al. intrusions. Globally, the sedimentary filling is poorly deformed except at the edges of the basins where sediments are locally overturned because of mud intrusions (Fig. 14). Mud flows with an extend of about 3 km are interbedded within tabular sediments (Fig. 13). High pore fluid pressure and steep slopes associated with diapiric structures promote mass movements from the flanks of the diapiric edifices to the basins. Downslope accumulation is very chaotic (Fig. 15) showing the heterogeneous nature of these sliding masses. The mean thickness of the displaced sediments is about 20 m which yields a total volume of more than 120 x 106 m 3. These sliding masses are strictly localized at the edge of the basins as shown by Fig. 15 where the sliding chaotic sheet changes to undeformed sediments when the slope smoothes and shows again the N40~ direction characteristic of that of the frontal piggyback basins. Diapiric edifices undergo short-lived catastrophic events and recur every few tens of years without any steady cyclicity and the volume of material averages 3 x 106 m 3 (Higgins and Saunders, 1974). These observations could explain the occurrence of superimposed slides at the toes of diapirs rather than at the foot of tectonic structures whose reactivation in the case of a thrust fault related fold is far less frequent and less catastrophic (Fig. 9).
Evolution of the arcward piggyback basins Mud rises up through NS- or N40~ faults (Figs. 14 and 15) parallel to the direction of the deformation front. These faults show a recent normal downthrow probably due to late mud up-rising (Huyghe et al., 1996). The intrusion of mud into N40~ faults bounding the arcward basins raises the question whether they are inherited frontal piggyback basins which suffered a long displacement above the major thrust system and which were deformed by mud
Fig. 11. CEPM seismic profile from Biju-Duval et al. (1982). Basins equally form back of thrusts or back-thrusts.
LATE N E O G E N E P I G G Y B A C K B A S I N S ON T H E B A R B A D O S R I D G E C O M P L E X
379
Fig. 12. Diapicar MN water-gun seismic profile (non migrated). Horizontal scales are the same on the seismic profile and line drawing.
intrusions. On the other hand, the influence of N W SE- and NE-SW-trending structures oblique to the deformation front has also been observed or invoked (Biju-Duval et al., 1982; Griboulard et al., 1989; Huyghe et al., 1996) in the arcward zone of the wedge. These structures have been related either to inherited oceanic ridges (Peter and Westbrook, 1976), basement faults extending from the Venezuelan margin to the wedge (Valery et al., 1985), or transcurrent faults associated with the E1 Pilar fault system in the southern part of the prism (Griboulard et al., 1991). These structures may be reactivated as transtensive or transpressive faults so that they generate local uplifting or sinking and may be then responsible for bulges alignment and for the complex arcuate shape of the arcward basins. As mud intrudes pre-exiting and new forming faults, it is hard to assess the origin of the piggyback basins occurring in the arcward areas.
PIGGYBACK BASINS AND LARGE SCALE EVOLUTION OF THE WEDGE
Piggyback basins show a major tectonic control in the zone of initial accretion. Deformation of distinct length scales (described above) controls their sedimentary pattern. At a regional scale, numerical (Chalaron et al., 1995) and analogic modelling (Storti and McClay, 1995; Mugnier et al., 1997) suggest that their repartition reflects the dynamic of the accretionary prism.
Evolution of the piggyback basins from north to south The Barbados wedge shows north to south variations. The most obvious one is its width, that varies from less than 100 km in the northern part to more than 200 km in the southern part. A similar increase
380
E HUYGHE et al.
Fig. 13. Diapicar HI water-gun seismic profile (non migrated). Horizontal scales are the same on the seismic profile and line drawing. Grey-shaded area: argilokinesis clay. Mud structures intrude poorly deformed sediments and determine sub-basins. East of the section, sediments are interbedded with mud flows. of the width and extent of the piggyback basins and of the sedimentation rate occurs (see Table 1). The north-south distribution of the basins and the evolution of the accretionary complex are controlled by the volume balance between sediments entering the wedge (sediments accreted at the toe, deposited upon the prism and incorporated to the wedge by subcretion) and those lost by subduction and surficial erosion (Le Pichon et al., 1990; Speed, 1990). Sediment infilling of the piggyback basins mainly originates from the South American shelf. Therefore sediment supply above the wedge decreases from south to north, and basin geometry evolves from nearly overfilled basins to the south where recent sediments frequently overlap all the structures, to northern areas with very poor sediments trapped between thrusts sheets. Thickness of the sediment column incorporated within the wedge varies with the distance to the southern location of continental sources. The increasing thickness, from few hundred meters to nearly 7000 m, favors the forward propagation of the deformation front and generates a general eastward shift of the deformation front from 16~ south to 1 l~ Furthermore, the increase of the stratigraphic thickness of the thrust sheets from north to south induces an increase of the fold spacing (Fig. 6) and an increase of the width of the basins
from few thousand meters at 12~ to 6-8 km at 10ON. Shortening rate velocities vary from north to south due to the plate kinematics: earthquake mechanisms (Russo and Speed, 1992) and NUVEL-1 plate model (DeMets et al., 1990, 1994) provide strong support for the fight oblique collision model of Caribbean-South American plate interaction but the prolongation towards the east of the CaribbeanSouth American plate boundary zone (CA-SA PBZ) is still debated (Speed et al., 1991). A reasonable hypothesis is to infer that the shortening rate at the back of the wedge decreases from the latitude of the CA-SA PBZ towards the south, leading to the arcuate westward shape of the deformation front and occurrence of tear faults in the displaced wedge. The arcward basins located between 11~ and 10~ are frequently characterized by a crescent-shaped pattern (Fig. 4) and presumably develop above dextral transcurrent shear zones (Gfiboulard et al., 1991). Evolution of the piggyback basins from east to west The above study shows that the sedimentary pattern of frontal basins is controlled by out-offsequence reactivation along forward verging thrusts.
Fig. 14. (a) Sidescan sonar (SAR) images corresponding to the Diapisar 9020 seismic profile shown below. The size of mud volcano is about 8 km 2. Horizontal scales are the same for (a), (b) and (c). (b) Diapisar 9020 seismic profile (3.5 kHz) across the northeastern end of a piggyback basin. The N40~ fault that crosses the section, controls the depression and the mud uprising. South of the mud volcano, sediments of the basin have been turned up by the mud intrusion. Vertical exaggeration of seismic document x5. (c) Line drawing of (b).
X
0
C~
0
>.
>
0 Z ,H
Z
>
C~
Z
O
Z
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
Fig. 15 (continued).
383
a)
MUDVOLCANO~
I
o
0
0
~
'
1
1000 O0
I
,ooooo
I
10 0000
400000
'
'
'
200000
I
~ooooo
I
200000
200 000
I
'
'
'
'
,
300 000
I
~ooooo
I
3 00000
I
"300000 t
I
'
'
''
'
I
I
400000
.|
,
400000
I
400000
I
,
'
'
|
500 000
I
~ooooo
I
=500000
I
6 00000
MUD VOLCANO ~
'
'
EI
)
!
Ie ~1 NW
I
g)
NW
b)
~:i ~
j
I
I
/
DEEP MUD
SE I
(RE)ACTIVATION
~
NO (RE)ACTIVATION
RISINGFLUIDS (1~
/
/
S E]
~
~,~
FRONTAL BASINS
/
I
IL~ ,~J"'
/
voLcANo
MUD
BASINS
PIGGYBACK BASIN3
0.2 ~
ARCWARD
MUD VOLCANO
)
oo, o
f) ~
CARIBBEAN
oo oo
'
'
o~
PLATE
,,oo
.
(~ /
(
O9
~
~
~ ,..a
I~.
0
0-~
I
l"
PIGGYBACK BASIN1 ~
............P . ~ T E ...........
-.
~-o
I
k o =,,~L
x
~, , ~
:..-l~e,
~0 ~ l
,,,,, •
'~
o~oo~o %
~O~::~:ER,CXN:.ii!~.iliiilii~'~~'%"..4"~?~ "~,~
12~
"14~
.
c)
-J >.. m
o_
Z
SI::l
Fig. 16. Distinction between frontal and arcward piggyback basins. (a) Line drawing of the Diapicar HI seismic profile (Fig. 13). Grey shaded area: argilokinesis clay. Mud structures intrude poorly deformed sediments and determine sub-basins. (b) Line drawing of the Diapicar FG seismic profile (from Griboulard et al., 1991). Mud (indicated by shading) rises up through pre-existing faults and intrudes the cores of the anticlinal structures. It may generate thick mud flows that feed piggyback basins. (c) Location of bathymetric profiles 1 to 5. (d) Bathymetric profiles across the SBRC realized from GEBCO bathymetric map (Fontas et al., 1985). The flattening of the topography in the arcward zone of the wedge fits with the occurrence of numerous mud diapirs. (e) Line drawing of the Diapicar OP seismic profile (Fig. 10) showing the backward migration of depocenters in piggyback-back basins. (f) Line drawing of the Diapicar KL seismic profile (Fig. 8) showing the backward migration of depocenters in piggyback basins. (g) Schematic evolution of the wedge and piggyback basins. Frontal basins are mainly controlled by forward-verging thrust system that form a brittle wedge, whereas arcward basin development is mainly controlled by subcretion of deep muds that induces mud diapirism, a gentle regional slope and the occurrence of both backward and forward verging thrusts.
-4000
-4000
-4000
0
0
-4000~,"~'----
O
d)
Iw
~Z
Or
385
LATE NEOGENE PIGGYBACK BASINS ON THE B A R B A D O S RIDGE COMPLEX
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.
~D
9~
=
9
~D r
o A
.A . ,...~
~D ~D c'-I
C',l
-~-
~
PIGGYBACK BASINS OF THE BARBADOS RIDGE
r162
COMPLEX COMPARED TO OTHER PIGGYBACK BASINS
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
k9 .,.,~ 9 > 9
C o
~D
k9
C'-I 9
< ~D
=
"~ tt')
== 9 0
tt')
~
Z
~
<
<
<
386
E HUYGHE et al.
I antetectonic sediments
0
5em
t
I
Fig. 17. Analogue models showing the influence of syntectonic sedimentation on thrust wedge evolution (adapted from Storti and McClay, 1995). The three analogue experiments have been performed with the same experimental conditions in a sandbox apparatus. The three final geometries result from a horizontal shortening applied to the same initial thickness of the antetectonic sediments (drawn in white), but the syntectonic sediments supply (grey shaded) varies: (a) experiment F8 was performed without sedimentation and with 57% contraction; (b) experiment F4 was performed with low syntectonic sedimentation and 43% contraction; syntectonic sediments form piggyback basins above thrust sheets; (c) experiment F7 was performed with high syntectonic sedimentation and 17% contraction. Thrusts dip very steep through the syntectonic sediments. Faug~res et al., 1991) consisting of fine material transported as suspension within the water masses. Rate of d e v e l o p m e n t of the p i g g y b a c k basins
Piggyback basins development and thrust wedge geometry are linked by a strong feed-back: the width and depth of the depression is controlled by the thrust wedge thickness whereas the abundant sediment supply dramatically changes the shape of the Coulomb thrust wedge, decreasing the number of thrust slices and the critical taper of the wedge (Fig. 17, from Storti and McClay, 1995). Simultaneously, syntectonic sedimentation ahead of the thrust front increases the thickness of the incipient thrust. A key parameter in the understanding of these feedback phenomena is the ratio between the sediment supply efficiency and the back-to-the-wedge velocity (column 3 of Table 1) as inferred from numerical modelling (Leturmy et al., 1995). The parameters listed in Table 1 are crude estimations and are affected by great uncertainties. Nonetheless the sedimentation rate at the toe of the accretionary complex (Se) partly reflects the sediment supply efficiency, whereas the adimensional parameter Se/Sh is useful for a comparison of natural piggyback basins or analog experiments. The large piggyback basins of the Subandean Belt (Baby et al., 1995) or Apennines (Pieri, 1989; Mugnier and Endignoux, 1991; Kruse
and Royden, 1994; Artoni and Casero, 1997) are characterized by a ratio Se/Sh close to 20-40%. The piggyback basins of the SBRC (1, 2 and 3 in Fig. 4) are characterized by a small Se/Sh value (less than 5%). The Se/Sh ratio varies dramatically from north to south in the BRC, and a comparison of geometry of the analogic models (Fig. 17) and schematic cross-sections of the zone of initial accretion of the BRC (Fig. 6) outlines the influence of syntectonic sedimentation in the Barbados prism. The model of Fig. 17a, performed without sedimentation, shows analogy with the 15~ cross-section of Fig. 6a. The model of Fig. 17b performed with a 5% Se/Sh ratio looks like the 1 I~ cross-section also characterized by a 5% Se/Sh ratio, whereas the section at 10~ and the model of Fig. 17c show very steep thrusts linked to high syntectonic sedimentation.
SUMMARY
The development of the SBRC frontal piggyback basins is mainly controlled by a forward-verging thrust system formed above a decollement, whereas the development of arcward basins is mainly controlled by the subcretion of deep muds that induces mud diapirism, a very gentle regional slope and the occurrence of both backward and forward verging thrusts. In both areas, features with prominent relief produce lumpiness in the mean regional slope and consist of either hanging-wall of faults or mud domes. They are locally eroded during reworking of slope sediments and form barriers to terrigenous supply from the South American continent. Therefore these reliefs favor the trapping of sediments in the basins of the southern part of BRC, whereas its northern part is poorly nourished.
ACKNOWLEDGEMENTS
We would like to thank N. Bangs, E. Calais, D. Davis and E Mann for helpful reviews of the initial draft of the paper.
REFERENCES
Artoni, A. and Casero, E, 1997. Sequentialbalancing of growth structures, the late Tertiary example from the central Apennine. Bull. Soc. Geol. Fr., 168: 35-50. Baby, E, Coletta, B. and Zubieta, D., 1995, Etude g6om6trique et exp6rimentaled'un bassin transport6: exemple du synclinarium de l'Alto Beni (Andes Centrales). Bull. Soc. Geol. Fr., 166: 797-813. Bangs, N. and Westbrook, G.K., 1990. Seismic modeling of the decollement zone at the base of the Barbados Ridge Accretionary Complex. J. Geophys. Res., 96: 3853-3866. Bangs, N., Shipley T.H. and Moore, G.F, 1996. Elevated fluid
LATE N E O G E N E P I G G Y B A C K BASINS ON THE BARBADOS RIDGE C O M P L E X pressure and fault zone dilatation inferred from seismic models of the northern Barbados Ridge decollement. J. Geophys. Res., 101: 627-642. Belderson, R.H., Kenyon, N.H., Stride, A.H. and Pelton, C.D., 1984. A 'braided' distributary system on the Orinoco deep-sea fan. Mar. Geol., 56: 195-206. Biju-Duval, B., Le Quellec, E, Mascle, A., Renard, V. and Valery, R, 1982. Multi-beam bathymetric survey and high resolution seismic investigations on the Barbados ridge complex (Eastern Caribbean): a key to the knowledge and interpretation of an accretionary wedge. Tectonophysics, 80: 275-304. Bilham, R., Larson K., Freymueller J. and Project Idylhim members, 1997. GPS measurement of present-day convergence across the Nepal Himalaya. Nature, 386: 61-64. Birsh, ES., 1970. The Barracuda Fault Zone in the Western North Atlantic: geological and geophysical studies. Deep Sea Res., 17: 847-859. Bouysse, E and Westercamp, D., 1988. Effets de la subduction de rides ocdaniques sur l'6volution d'un arc insulaire: l'exemple des Petites Antilles. G6ol. Fr., BRGM Eds., 2-3" 3-38. Brown, K.M. and Westbrook, G.K., 1987. The tectonic fabric of the Barbados ridge accretionary complex. Mar. Pet. Geol., 4: 71-81. Brown, K.M. and Westbrook, G.K., 1988. Mud diapirism and subcretion in the Barbados ridge accretionary complex: the role of fluids in accretionary processes. Tectonics, 7: 613-640. Burbank, D.W. and Verg6s, J., 1994. Reconstruction of topography and related depositional systems during active thrusting. J. Geophys. Res., 99: 20281-20297. Calassou, S., Larroque, C. and Malavielle, J., 1993. Transfer zones of deformation in thrust wedges: an experimental study. Tectonophysics, 221: 325-344. Chalaron, E., Mugnier, J.L. and Mascle, G., 1995. Control on thrust tectonics in the Himalayan foothills: a view from a numerical model. Tectonophysics, 248: 139-163. Dahlen, F.A., 1990. Critical taper model of fold-and thrust belts and accretionary wedges. Annu. Rev. Earth Planet. Sci., 18: 55-99. Damuth, J.E., 1975. Echo-character of the western equatorial atlantic floor and its relationship to the dispersal and distribution of terrigenous sediments. Mar. Geol., 18: 17-45. Davis, D. and Lillie, J.R., 1994. Changes in mechanical response during continental collision: active example from the foreland thrust belts of Pakistan. J. Struct. Geol., 16: 21-34. Davis, D., Suppe, J. and Dahlen F.A., 1983. Mechanics of fold-and-thrust belts and accretionary wedges. J. Geophys. Res., 88:1153-1172. DeMets, C., Gordon, R.G., Argus, D.E and Stein, S., 1990. Current plate motions. Geophys. J. Int., 101: 425-478. DeMets, C., Gordon, R.G. and Vogt, E, 1994. Location of the Africa-Australia-India triple junction and motion between the Australian and Indian plates: results from an aeromagnetic investigation of the Central Indian and Carlsberg ridges. Geophys. J. Int., 119: 893-930. Dia, N.A., Castrec, M. and Boul6gue, J., 1997. Trinidad mud volcanoes: where do the expelled fluids come from? Terra Nova, 9, Abstr. Suppl., 1:319. Embley, R.W. and Langseth, M.G., 1977. Sedimentation processes of the continental rise of north-eastern South-America. Mar. Geol., 25: 279-297. Endignoux, L. and Mugnier, J.L., 1990. The use of a forward kinematic model in the construction of balanced cross-sections. Tectonics, 9: 1249-1262. Ewing, M., Carpenter, G., Windisch, C. and Ewing, J., 1973. Sediment distribution in the oceans: the Atlantic. Geol. Soc. Am. Bull., 84:71-88. Faug6res, J.C., Gonthier, E., Mass6, L., Parra, M., Ports, J.C. and Pujol, C., 1991. Quaternary deposits on the South Barbados
387
accretionary prism. Mar. Geol., 96: 247-267. Faugbres, J.C., Gonthier, E., Griboulard, R. and Mass6, L., 1993. Quaternary sandy deposits and canyons on the Venezuelan margin and south Barbados accretionary prism. Mar. Geol., 110:115-142. Fontas, E, Griboulard, R. and Prud'Homme, R., 1985. Morphogdn6se d'un prisme d'accrdtion et variabilit6 du style des ddformations sur une marge active. Apports et spdcificit6 des rdsultats d'une analyse cartographique interprdtative sur la ride de la Barbade. In: A. Mascle (Editor), G6odynamique des Carafbes, Symposium. Technip, Paris, pp. 221-233. Griboulard, R., Faug6res, J.C., Blanc G., Gonthier, E. and Vernette, G., 1989. Nouvelles 6vidences sddimentologiques et gdochimiques de l'activit6 actuelle du prisme Sud Barbade. C.R. Acad. Sci., Paris, 308:75-81. Griboulard, R., Bobier, C., Faug6res, J.C. and Vernette, G., 1991. Clay diapiric structures within the strike-slip margin of the southern leg of the Barbados prism. Tectonophysics, 192: 383-400. Henry, E, Le Pichon, X., Lallemant, S., Foucher, J.E, Westbrook, G.K. and Hobart, M.A., 1990. Mud volcano field seaward of the Barbados accretionary complex: a deep-towed side scan sonar survey. J. Geophys. Res., 95: 8917-8929. Higgins, G.E. and Saunders, J.B., 1974. Mud volcanoes - - their nature and origin. In: Contributions to the Geology and Paleobiology of the Caribbean and Adjacent Areas. Naturforsch. Ges. Verh., Basel, 84: 101-152. Huyghe, E, Griboulard, R., Faug~res, J.C., Gonthier, E. and Bobier, C., 1996. Gdom6trie des bassins du prisme Sud-Barbade. Bull. Soc. G6ol. Fr., 167: 345-359. Jordan, T.H., 1975. The present-day motions of the Caribbean Plate. J. Geophys. Res., 80: 4433-4439. Kruse, S.E. and Royden, L.H., 1994. Bending and unbending of an elastic lithosphere: the Cenozoic history of the Apennine and Dinaride foredeep basins. Tectonics, 13: 278-302. Labaume, E, Henry, E and scientist party of Leg ODP 156, 1995. Circulation et surpression de l'eau interstitielle dans le prisme d'accrdtion nord-Barbade: r6sultats du Leg ODP 156. C.R. Acad. Sci. Paris, 320: 977-984. Lallemant, S.J.C., Henry, E, Le Pichon, X. and Foucher, J.E. 1990, Detailed structure and possible fluid paths at the toe of the Barbados accretionary wedge (ODP Leg 110 area). Geology, 18: 854-857. Langseth, M., Westbrook, G.K. and Hobart, M.A., 1988. Geophysical survey of a mud-volcano seaward of the Barbados ridge accretionary complex. J. Geophys. Res., 93:1049-1061. Leonard, R., 1983. Geology and hydrocarbon accumulation, Colombus Basin, offshore Trinidad. Am. Assoc. Pet. Geol. Bull., 67: 1081-1093. Le Pichon, X., Foucher, J.P., Boul6gue, J., Henry, E, Lallemant, S., Benedetti, M., Avedik, F. and Mariotti, A., 1990. Mud volcano field seaward of the Barbados Accretionary Complex: a submersible survey. J. Geophys. Res., 95:8931-8944. Leturmy, P., Mugnier, J.L. and Chalaron, E., 1995. Erosion et sddimentation au voisinage d'un anticlinal de rampe: apport d'un module numdrique. Bull. Soc. G6ol. Fr., 166: 345-359. Mascle, A. and Moore, J.C., 1988. Proceedings of the Ocean Drilling Project, Initial Reports (Pt. A). College Station, TX, 110, 603 pp. Mascle, A., Endignoux, L. and Chennouf, T., 1990. Frontal accretion and piggyback basin development at the southern edge of the Barbados ridge accretionary complex. Proc. ODE Sci. Results, 110:17-28. Moore, J.C., 1989. Tectonics and hydrogeology of accretionary prisms: role of the ddcollement zone. J. Struct. Geol., 11, 1/2: 95-106. Moore, J.C. and Mascle, A., 1990. Proceedings of the Ocean Drilling Program, Scientific Results. College Station, T X , ,
388 110, 448 pp. Mugnier, J.L. and Endignoux, L., 1991. Cin6matique et vitesse d'6volution des nappes superficielles: une simulation num6rique. Rev. Fr. G6otech., 56: 23-32. Mugnier, J.L., Baby, E, Colletta B., Vinour, E, Bale, E and Leturmy, E, 1997. Thrust geometry controlled by erosion and sedimentation: a view from analogue models. Geology, 25 (5) (in press). Mulugetta, G., 1988. Modelling the geometry of Coulomb thrust wedges. J. Struct. Geol., 10 (8): 847-860. Ori, G.G. and Friend, EE, 1984. Sedimentary basins formed and carried on active thrust sheets. Geology, 12: 475-478. Peter, G. and Westbrook, G.K., 1976. Tectonics of southwestern North Atlantic and Barbados Ridge Complex. Am. Assoc. Pet. Geol. Bull., 60:1078-1106. Pieri, M., 1989. Three seismic profiles through the Po Plain. In: A.W. Bally (Editor), Atlas of Seismic Stratigraphy, Vol. 3. Am. Assoc. Pet. Geol. Stud. Geol., 27:90-110. Robertson, R.E and Burke, K., 1989. Evolution of southern Carribean plate boundary, vicinity of Trinidad and Tobago. Am. Assoc. Pet. Geol. Bull., 73 (1): 490-509. Robertson, E and Burke, K., 1991. Evolution of the southern Caribbean plate boundary, vicinity of Trinidad and Tobago: reply. Am. Assoc. Pet. Geol. Bull., 75 (11): 1795-1796. Roure, E, Howell D.G., Guellec, S. and Casero, E, 1990. Shallow structures induced by deep-seated thrusting. In: J. Letouzey (Editor), Petroleum Tectonics in Mobile Belts. Technip, Paris, pp. 15-30. Russo, R.M. and Speed, R.C., 1992. Oblique collision and tectonic wedging of the South American continent and Caribbean terranes. Geology, 20: 447-450. Russo, R.M., Speed, R.C. and Okal, E.A., 1993. Seismicity and tectonics of the Southeastern Caribbean. J. Geophys. Res., 98: 14299-14319. Shih, T.T., 1967. A survey of the active mud volcanoes in Taiwan and a study of their type and the character of the mud. Pet. Geol. Taiwan, 5: 259-311. Silver, E.A., Breen, N.A. and Prasetyo, H., 1986. Multibeam study of the Flores back arc thrust belt, Indonesia. J. Geophys. Res., 91: 3489-3500. Speed, R.C., 1985. Cenozoic collision of the Lesser Antilles arc and continental South America and the origin of the E1 Pilar Fault. Tectonics, 4 (1): 41-69. Speed, R.C., 1990. Volume loss and defluidization history of Barbados. J. Geophys. Res., 95: 8983-8996. Speed, R.C., Russo, R., Weber, J. and Rowley, K.C., 1991. Evolution of Southern Caribbean Plate Boundary, Vicinity of Trinidad and Tobago: discussion. Am. Assoc. Pet. Geol. Bull., 75: 1789-1794. Stein, S., Engeln, J.F. and Wiens, D.A., 1982. Subduction, seismicity and tectonics in the Lesser Antilles Arc. J. Geophys. Res., 87 (B10): 8642-8664. Stephan, J.E, Mercier, J., De Lepinay, B., Calais, E., Tardy, M., Beck, C., Carfantan, J.Ch., Olivet, J.L., Vila, J.M., Bouysse, E, Mauffret, A., Bourgeois, J., Thery, J.M., Tournon, J., Blanchet,
E H U Y G H E et al. R. and Dercourt, J., 1990. Paleogeodynamic maps of the Carribean: 14 steps from Lias to Present. Bull. Soc. G6ol. Fr., VI: 915-919. Storti, E and McClay, K., 1995. Influence of syntectonic sedimentation on thrust wedges in analogue models. Geology, 23: 999-1002. Suppe, J., 1983. Geometry and kinematics of fault bend folding. Am. J. Sci., 283: 684-721. Suppe, J., Chou, G.T. and Hook, S.C., 1992. Rates of folding and faulting determined from growth strata. In: K.R. McClay (Editor), Thrust Tectonics. Chapman and Hall, London, pp. 105-121. 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: 1065610676. Valery, E, Nely, G., Mascle, A., Biju-Duval, B., Le Quellec, E and Berthon, J.L., 1985. Structure et croissance d'un prisme d'accr6tion tectonique proche d'un continent: la ride de la Barbade au sud de l'arc Antillais. In: A. Mascle (Editor), G6odynamique des Cara'fbes. Technip, Paris, pp. 173-186. Weber, J.C., Ambeh, W.A., Dixon, T.D., Speed, R.C., Lynch, L.L., Saleh, J. and Webb, E, 1995. Historic geodetic constraints on Caribbean-South America relative plate motion, plate boundary zone kinematics and seismic risk. 3rd Geological Conference of the Tobago and 14th Caribbean Geological Conference, Port of Spain, Trinidad, pp. 83-84. Westbrook, G.K., 1982. The Barbados ridge complex: tectonics of a mature forearc system. In: J. Leggett (Editor), Trench and Forearc Geology. Geol. Soc. London Spec. Publ., 10: 275290. Westbrook, G.K. and Smith, M.J., 1983. D6collements and mud volcanoes: evidence from the Barbados Ridge complex for high pore fluid pressure in the development of an accretionary complex, Geology, 11: 279-283. Westbrook, G.K., Mascle, A. and Biju-Duval, B., 1984. Geophysics and the structure of the Lesser Antilles Forearc. Init. Rep. DSDE 78A: 23-38. Westbrook, G.K., Ladd, J.W., Buhl, E, Bangs, N. and Tiley, G.J., 1988. Cross section of an accretionary wedge: Barbados Ridge complex. Geology, 16: 631-635. Williams, C.A., Conners, C., Dahlen, E, Price, E. and Suppe J., 1994. Effect of the brittle-ductile transition on the topography of compressive mountain belts on Earth and Venus. J. Geophys. Res., 99: 19947-19974. Wright, A., 1984. Sediment distribution and depositional processes operating in the Lesser Antilles intraoceanic island arc, Eastern Caribbean. Init. Rep. DSDE 78A: 301-324. Zhao, W.L., Davis, D.M., Dahlen, EA. and Suppe, J., 1986. Origin of convex accretionary wedges: evidence from Barbados. J. Geophys. Res., 91: 10246-10258. Zoetemeijer, R. and Sassi, W., 1992. 2-D reconstruction of thrust evolution using the fault-bend fold method. In: K.R. McClay (Editor), Thrust Tectonics. Chapman and Hall, London, pp. 133-140.
Chapter 15
Tectonic Evolution of the Grenada Basin
DALE E. BIRD, STUART A. HALL, JOHN E CASEY and PATRICK S. M I L L E G A N
Detailed analyses of gravity, seismic reflection and refraction data are integrated with an earlier interpretation of magnetic data to produce a coherent model for the tectonic evolution of the Grenada basin that suggests that the basin formed by near east-west extension. Although the seafloor of the Grenada basin changes from smooth and undisturbed in the south to rugged with relatively high relief in the north, Bouguer anomalies and two-dimensional and three-dimensional gravity models, based upon seismic refraction and reflection data, reveal that the crust gradually thins in an east-west sense towards the center of the basin. Typical back-arc crust is observed in the southern part of the basin, but refraction data are not sufficiently reliable in the northern part to adequately determine the nature of the crust. Several curvilinear discontinuities in magnetic, gravity and bathymetric trends are observed. These discontinuities, when integrated with two-dimensional and three-dimensional modeling and analyses of Bouguer gravity anomalies, are interpreted to be due to late Tertiary compressional forces in the northern part of the region. These compressional forces have resulted in the bifurcation of the Lesser Antilles island arc north of 15~ the westward displacement of part of the Aves Ridge (a remnant island arc), and the crustal deformation observed in the northern Grenada basin. The compressional forces also appear to have sufficiently disrupted the crust in the northern Grenada basin such that earlier magnetic anomaly patterns have been modified to yield the observed magnetic signature.
INTRODUCTION Various orientations ranging from north-south to east-west have been proposed for the direction of extension and subsequent formation of the Grenada basin (Tomblin, 1975; Bouysse, 1988; Pindell and Barrett, 1990). Large-amplitude, long-wavelength magnetic anomalies, oriented generally e a s t - w e s t over the basin, have been used by authors to arrive at these conflicting interpretations of the basin's evolution. The Grenada back-arc basin (Fig. 1) formed in early Tertiary time at approximately 12~ (Duncan and Hargraves, 1984; Ghosh et al., 1984; Pindell et al., 1988; Ross and Scotese, 1988). Anomalies produced by east-west-trending features at low magnetic latitudes have significantly larger amplitudes than anomalies produced by north-south-trending features. The orientation of the Earth's magnetic field at low latitudes is nearly horizontal and its direction is essentially to the north, therefore east-west-oriented features produce large perturbations in the Earth's magnetic field while northsouth-oriented features result in minor anomalies.
If the Grenada basin formed by e a s t - w e s t extension, then magnetic anomalies produced by e a s t west seafloor spreading in the back-arc region would be characterized by small amplitudes. Furthermore, if these north-south spreading centers were offset by east-west-trending features, then anomalies produced by these features would exhibit large amplitudes that would overwhelm the north-south-trending anomalies. We have previously analyzed shiptrack and gridded magnetic data over the Grenada basin, including anomaly trend analyses and three-dimensional forward modeling (Bird et al., 1993). Subtle, subparallel to the Lesser Antilles island arc anomalies have been identified over the southern part of the basin and are interpreted to be due to near e a s t west extension. Amplitudes of these anomalies are approximately 40 nT while amplitudes of e a s t west-oriented anomalies are about 300 nT. This seven-fold increase, the difference between northsouth- and east-west-oriented profiles, was further supported by three-dimensional forward modeling. Although results from our analyses of magnetic data support near e a s t - w e s t extension, and probable
Caribbean Basins. Sedimentary Basins of the World, 4 edited by P. Mann (Series Editor: K.J. Hsti), pp. 389-416. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
390
D . E . B I R D et al.
20"N
0 I
~t
18"N
r
f4~.--~J~
[\
9.o ~ ~_.,,_ ~ ~
0~,~
,3 Saba Bank ' q ~
~~
,,~=:~,
~
200 t
300 ~ I
~5~L
(~Caribbean Plat,
k
~r
,.
I
=
',
"~
I
c
o
2. J
"
:,
c
-~
14"N
A
)
/
i
12"N~
/
I
2"J I
~ ,",..
o/,5 ~
~2 9
~ U
. ,m,,
,o.,-
'
I
.
.
.
.
~- : -. - ~
-
i
64"w
,
"
i
i:~
::.-r.,,~
62ow
4",
i
'".:, '.";..~ ~ ... .~ ~--!.:.!~~South ...'~:" " " " ' ~ American
.....
i ~'w
km
Northh American Plate
._
,
16"N "-"
100 I
i
i Plate
6o'w
t
\
I 58"w
Fig. 1. Physiographyof the eastern Caribbean with 2, 4 and 5 km isobaths contoured (after Bouysse, 1984). The outline of the study area, trace of the subduction zone, and strike-slip fault zones which define the North American/Caribbean and South American/Caribbean plate boundaries are displayed. Heavy dashed lines indicate probable locations for plate boundaries. The inner and outer arcs are represented by dashed and dotted lines, respectively.
back-arc spreading for the formation of the Grenada basin, the results should be integrated with analyses of gravity and seismic data to provide more evidence with which to examine the evolution of the region. In this study, interpretations of seismic and gravity data are combined with our earlier interpretation of magnetic data. Then two-dimensional and threedimensional forward models, incorporating gravity and seismic data (both reflection and refraction), are constructed for the region. Comparison of these results for the Grenada basin with the morphology of the younger Andaman Sea back-arc basin (11 Ma) provide a compelling interpretation for the evolution of the Grenada basin (Mukhopadhyay, 1984).
MODELS FOR THE FORMATION OF THE GRENADA BASIN
Various kinematic models for the formation of the Grenada basin are described by Tomblin (1975), Bouysse (1988), Pindell and Barrett (1990) and Bird et al. (1993). These models outline the formation of the basin by near east-west, northeast-southwest extension, north-south extension and east-west extension, respectively. East-west
extension
Tomblin (1975) describes two possible scenarios for east-west extension. The first involves an early
TECTONIC EVOLUTION OF THE GRENADA BASIN North American Plate .....
o.x, 5
Caribb
(A) South American Plate North American Plate
391 Ridge to the west) requires the formation and subsequent spreading from a north-south-oriented median ridge. He reports that no such ridge is observed. However, since extension is no longer evident in the Grenada basin, the loss of heat at spreading centers would cause the ridge system to cool and subside. Therefore a ridge system may not be obvious in the data, but the center of the basin would still produce a Bouguer gravity high when compared to thicker crust on either side. Intuitively, east-west extension as proposed by Tomblin seems most reasonable, and magnetic anomaly lineations over many back-arc basins support this conclusion (i.e., the South Sandwich, Lau, Havre, North Fiji, Banda Sea, Parece-Vela, Shikoku, and Sea of Japan basins). North-south extension
Caribbean Plate
~
,
(B) South American Plate North American Plate
C a r i b b ~ } Pla
(C) South American Plate 0
300 km
Fig. 2. (A) Possible east-west extension due to a westward shift of the Aves Ridge for the opening of the Grenada basin (Tomblin, 1975). (B) Possible north-south extension for the opening of the basin (Pindell and Barrett, 1990). (C) Possible northeastsouthwest extension for the opening of the basin (Bouysse, 1988). Large arrows indicate the relative motions of the North American, Caribbean, and South American plates and small arrows indicate the directions of extension for the formation of the basin.
Tertiary eastward shift of the subduction zone, and the second involves a westward shift of the Aves Swell (Fig. 2A). In the first model the older rocks of la D6sirade would either have been part of an older orogeny and moved eastward with the subduction zone, or they may represent obduction of part of the Atlantic floor onto the eastward moving Caribbean plate. Tomblin's alternative scenario for the formation of the basin (a shift of the Aves
Pindell and Barrett (1990) feel that the Leeward Antilles were coupled to the northern edge of the South American plate and that north-south spreading in the Grenada basin is a result of continued eastward progression of the Caribbean plate. The basin therefore formed by right lateral shear (Fig. 2B). In this model the Leeward Antilles may have been part of the Aves Ridge prior to the formation of the basin and represents fragmentation of the arc as the Caribbean plate progressed eastward. If the long-wavelength, high-amplitude, eastwest-trending magnetic anomalies over the basin are produced by seafloor spreading, then this model appears to fit the magnetic data. For this model differences in the nature of the crust between the northern and southern parts of the basin are important. Pindell and Barrett (1990) suggest that the northern part of the basin is block faulted with no development of oceanic crust; however, the southern part of the basin may be underlain by oceanic crust. Although Pindell and Barrett's model may appear overly complex, magnetic anomaly patterns over the Andaman Sea basin appear to be oriented at a high angle to the trench line of its subduction zone. Northeast-southwest extension
Similar to Pindell and Barrett's model, Bouysse (1988) describes a possible mechanism for extension in which coupling of the southern part of the Lesser Antilles with the South American plate precedes the opening of the basin (Fig. 2C). He suggests that the Netherlands-Antilles, Lesser Antilles and Greater Antilles formed a continuous Mesozoic arc prior to the injection of the Caribbean plate between the American plates. Bouysse further theorizes that subsequent seafloor spreading was oriented northeastsouthwest at the onset of the Cenozoic in a seg-
392 mented manner such as described by Tamaki (1985) for the Sea of Japan basin. Initial spreading was in the southernmost portion of the basin and gradually progressed northward over time. Bouysse's model provides for contemporaneous formation of the Yucatan and Grenada basins. This development occurred when the northeast-travelling Caribbean plate, with respect to the North American plate, was wedged between the North American and South American plates in Late Cretaceous/early Tertiary time. Subsequent to this collision, the Caribbean plate rotated clockwise and began travelling in an east-west direction.
D.E. BIRD et al. North American
/ i/~JJ~
Caribbean Plat
_J
South AmericanPlate
Plate
J
fJ
~
(A)
Near east-west extension
Bird et al. (1993) interpret the basin to have formed by near east-west extension similar to that described by Tomblin (1975) with late Tertiary tectonic movements disrupting the crust, and magnetic signature, over the northern part of the basin (Fig. 3). These conclusions are supported by forward threedimensional magnetics modeling and identification of subtle anomaly trends over the southern part of the basin. These low-amplitude (about 40 nT) anomalies are interpreted as those produced by a roughly north-south-oriented spreading center(s) near the geomagnetic equator. The chaotic, patchy anomalies over the northern part of the basin are thought to have formed by seafloor spreading also, but later were disrupted by the late Tertiary event responsible for the bifurcation of the Lesser Antilles.
DATA BASE
The data base for this study includes gridded gravity and bathymetry data (Figs. 4 and 5, and WebFigs. 15.1-41), shipboard gravity and bathymetry profile data (Fig. 6), multiple channel seismic reflection sections (Fig. 7A), and reversed and unreversed seismic refraction profiles (Fig. 7B). Grids of total intensity magnetic anomalies (2 km) and free-air gravity (6 km) were compiled in 1987 by the Geological Society of America Decade of North American Geology (DNAG) Committees on the Magnetics and Gravity Maps of North America, respectively. These grids are available from the NOAA/National Geophysical Data Center (NGDC). The bathymetry grid is a portion of the ETOPO5 data. The ETOPO5 data set is a 5-minute grid of topography and bathymetry for the entire world and is available from the NGDC as well. Since the grid spacing is not constant as latitude varies, the data were regridded to 9 km utilizing a moving average least squares method. 1 Available at http://www.elsevier.nl/locate/caribas
0 "
"
I
300 l
km
North AmericanPlate
Caribbean
South AmericanPlate
(B)
Fig. 3. Two-step model for the formation of the Grenada basin by east-west extension. Large arrows indicate directions of relative plate motion and small arrows indicate directions of extension and basin formation. (A) The basin is formed fairly uniformly by early Tertiary seafloor spreading. (B) Late Tertiary compressional tectonism disrupts the northern part of the basin (indicated by long northeast-oriented arrows).
Multi-channel seismic reflection lines were obtained from the Institute Franqais du Petrole (IFP), Lamont-Doherty Geological Observatory (LDGO) of Columbia University, and the University of Texas at Austin (UT) Institute for Geophysics (Fig. 8). The reflection lines obtained from UT were originally acquired by Gulf Oil Company (GULFREX lines). Identification, fold, orientation, length, and year of acquisition of seismic reflection sections used are displayed in Table 1. Seismic reflection line RC1904 (LDGO) was available in digital form as well as its corresponding velocity analyses. Nine horizons of this line were analyzed by LDGO and those within the basin were used to interpret a velocity function for modeling. Refraction information used for the study comes
393
T E C T O N I C E V O L U T I O N OF T H E G R E N A D A BASIN
Fig. 4. Free-air gravity anomalies over the study area. Heavy solid and dashed lines indicate interpreted curvilinear zones of disruption (dashed lines indicate reduced confidence). Contour interval is 10 mGal. Table 1 Parameters of multiple channel reflection seismic data Organization
Line No.
Fold
Direction
Year
Length (km)
LDGO IFP IFP IFP IFP UT (GULFREX) UT (GULFREX) UT (GULFREX) UT (GULFREX) UT (GULFREX) UT (GULFREX) Total
C 1904 A4 118A 131A 217A LS-11 LS- 12 LS- 14 LS-15A VB-11 VB-12
24 48 48 48 48 48 48 24 24 24 24
E SE E SE N NE NE E NE NE SE
1975 1973 1974 1974 1976 1975 1975 1975 1975 1975 1975
870.50 116.66 169.77 222.73 111.36 169.85 143.96 304.84 165.40 138.73 132.89 2546.69
394
D.E. BIRD et al.
Fig. 5. Bathymetryof the study area. Heavy solid and dashed lines indicate interpreted curvilinear zones of disruption (dashed lines also indicate reduced confidence). Contour interval is 100 m.
from published studies (Ewing et al., 1957; Officer et al., 1959; Edgar, 1968; Boynton et al., 1979; Speed and Westbrook, 1984). In general, coverage of refraction data is sparse, particularly in the northern part of the basin where the velocity structure is represented by a single, unreversed profile (Fig. 7B).
TECTONICS OF BACK-ARC BASINS
Models for the formation of back-arc basins are discussed by, among others, Karig (1971), Sleep and Toskoz (1971), Poehls (1978), Uyeda and Kanamori (1979), Dewey (1980), Cross and Pilger (1982),
Taylor and Karner (1983), and Tamaki (1985). A back-arc basin is defined here as an extensional basin located at the edge of an overriding plate of a subduction zone such that the basin formed after initiation of subduction. At least nineteen basins have formed around the world by extension, or seafloor spreading, behind island arcs (Table 2). Back-arc basins formed by organized seafloor spreading should produce magnetic anomalies parallel to spreading centers and those with well defined magnetic anomaly lineations include the South Sandwich, Shikoku, Bismark, South Fiji, Lau, Havre, Banda, and Andaman basins (Taylor, 1979; Weissel, 1980, 1981; Barker and Hill, 1981;
........
_
.
,"
""
..
."
.
...
>,~,
"'***..
.,
'
\~
~
%
'
......
I
o
100
I
,
".
%"
"...
"
j
.~
.
:'
-... i
;
.
i ;
i
\
~....
""':J .... I
I
200
km
\
_,t--
(A)
:..? ....
L . ...~%..~."
,
]
60"W
......... : . ~ , .
;~..
...........
.
,>.
,:.-C .....
"~... .,~
...................
,.. ~ . ~ " :
I'Q
/,,,
!! II
:....................
~-L..
"'
I
-'\ ;~'~I
l~:'. . . . . . .
"~..
) ...............
--.
"
":':':.,,.. ,
"
It
-..~,,.~.....:. ..... _~+"
.
....
/
I
k
,
...""
./
/
"--'"...'-,,-
"..
-....
-.
i
9~ , !
/
. . . . .
/
/
,o.
,...,
62~
,.{..t, . . . . . . "" ..... .. ',,, "".~ .....
~ ~ i
~
I
9
0
" . . . . . . . . ~ .....
9
i
i
9
,
)
,:.:: ..... \
~,~
.:
.............................................
/
o.
"-%
c:'--~:."" 'r""-t . . . . J . . ~ ' . ' : ~ ,-,.':..J,
9,,
......
. . . .
.,::.-:":..............................................
''"
*"
.....-. . . . . . . . . .
........
9
:
.....
~
_
.
%
%.
64:W
Fig. 6. S h i p t r a c k data c o v e r a g e in the study area: (A) gravity, (B) b a t h y m e t r y .
~"
zo
*" i
o
z "-'~"
z
Zo
,,/
/
--
1~~
%
64:W
/
/
.,~
,
!
/
100
.."
/
\
I
200
62,W
km
(B)
I"" I
60~
L/I
Z
>
:>
Z >.
0 Z 0
m < o
0 :Z
m
z
66"W
";
118A
2kin
LS-15A/ ~O
64"W I
-- _
)
~~
i,
62"W I
Z;
i
P
LS-11
LS-14
;-12
,,.
60"W I
0 I
b
100 i
(A)
200 km I
C1904
58"W t
0I
l
i=
50 ,
dl 64"W
100 I
LASP B
.
150 km 1
I
63
"
"-I ql
--
29
62"W
c -=-II
b -=iil
I
d --,,0
'=-0
\
O
B
(B)
60"W
Fig. 7. Multi-channel seismic reflection data and refraction profiles. (A) Reflection data. LS = GULFREX, VB = GULFREX, A = IFP, C = LDGO. (B) Refraction data. For the Lesser Antilles Seismic Project (LASP), circles and triangles indicate shot and receiver locations, respectively, otherwise circles indicate receiver locations for the reversed and unreversed profiles. Velocities for profiles LASP B, 21, 22, 29, a, b, c, d, e, and f are shown in Fig. 15.
m
F_
QD
F_
q,-
I
ta3
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
STD
8.0 8.2 8.4 8.3 8.1 8.0 8.2 8.2 8.5 8.2
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)
2-D and 3-D Water Sediments Layer 2 (basement or upper crust)
1.03 2.30 2.57
2-D only Layer 3 (lower crust) Transition Mantle
2.74 3.05 3.30
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.
4~ 0
TECTONIC EVOLUTION OF THE GRENADA BASIN F
64"W
=,..
407
62"W
I
"--I ~, ~
60"W
I
I
I
q~
%
,, 4)
A
o
.z
qp,
(O
-~(-,)')B 14 (!1) J
G
~ .~
1
3
(15.5)
.,-
~,)
(8) \eqp 9 14 (10) -----.~) G 13(8) 17( l l
..5
H""5)
,,,, C
./=
I J
"",,
i
j
...--
r
,
/
Q
E __-
I
,,A
o
1 loo
I
200
P
km
Fig. l l. Locations for 2-D models of the basin. Posted values coincide with depth (km) to the base of layer 3 (Ludwig et al., 1971) and combined thickness (in parentheses) of both layers 2 and 3 from results of 2-D forward modeling.
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.
408
D.E. BIRD et al. i
i
i
i
i
!
i
i
I
/ s~ ....
I
.~
l 1
1=
GRAVITY (reGal)
- 2 "
1 ~/~'t'/'
~
tr
L \ "~qf
I
I
I
I
,
,
I
lp
~"
I
i
I
Ik =0
I
I
I
I
,
'
l
.I
,
~~
F' = 1.O3
= 2.s _ A= 0 ~~___-~
p = 2.74
k = 8000
p = 3.05
k = 0
p=3.3
k=O
O ~t
I
,
I
I
I
I
100 Distance (km)
I
I
(A)
'
200
0
1
~'2
I
i
~
i
I
I
I
l
GRAVITY (mGal)
, --
, ~~~o
1
(.3 .
t
i
I
I
I
I
I
I
I
i
I --.
i
I
I
i
I
I
I
I
I
p=1.03
I k = 0
I
p =,2.3
k=
G'
G O
O
p = 2.57
k = 8000
p = 2.74
k = 8000
p = 3.05
k = 0
p=3.3
k=O
O
I
i
9
r
_
0
1 O0
200
240
Distance (km) Fig. 12. (A) 2-D model F - F ' oriented east-west in the northern part of the basin. (B) 2-D model G - G f oriented east-west in the central part of the basin. Horizontal and vertical scales of the models are 1" 1.5 and 1 "0.75 million (2" 1 vertical exaggeration); 1 = observed gravity (free-air) profile; 2 -- calculated gravity profile.
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
F
I
yi
Z 9
409
62"W
I
"l~
I
60~
I
I
B
0 F_
F Q
---"
~
I 0
I 100
l 200
km
Fig. 13. Depth to acoustic basement with heavy lines indicating faults (after Speed and Westbrook, 1984). Contour interval is 500 m.
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.
~7
O
TECTONIC EVOLUTION OF THE GRENADA BASIN
~~
.54 .66 2.61
I~ ~ 1 ~ ! 2"15
5.44 ~]
I
411
~\\\-,=
1.69 3.00 3.86 l 4.42 ..~ 4.95
5.95
1~
1.55 to 3 . 5 0 km/s 5 . 5 0 to 6.45 km/s
5.86
21
I
/ 1.54 9 1.67
~x,,\'N ~ . .
1.63 1.83 3.33 9 4.40
6.24 9
,N••
2"37 3.00
J4.00
5.42
22
5.85
I//'//
t . , ~I 2.42
I
1.50
3.76 I
]
5.28
~\\\\\',,i ....... = J I~\\\\~1
I
6.41
1.69 1.93 2.25 2.95
I
1.53 2.41 2.97
_
~
1155 !
, - - - - - ! 1.60 1.79 2.47
,~\\~1300
14.-i
!
15.22
I
I
7.41
29
e
f
i"
'
-
i
T
7.40 ii
<
8.20
LASP B Fig. 15. Refraction velocities in the Grenada basin. Locations of velocity profiles are shown in Fig. 7B. respectively (Table 3). Similarly, lower velocities for profiles LASP B and 29 are 5.30, 6.20, 7.40, 8.2 km/s and 5.28, 6.22, 7.41 km/s, respectively. Velocities 5.28 or 5.30 and 6.22 or 6.30 km/s correlate
with typical oceanic crustal layers 2 (5 km/s) and 3 (6.4 to 7.1 km/s) described by Ludwig et al. (1971). In their models for the formation of the Grenada basin, Bouysse (1988) and Pindell and Barrett (1990)
412
D.E. BIRD et al. (discussed above) could correlate, but the overlying layers do not correlate. Similarly, the 4.98 km/s layer of reversed Profile 22 (Officer et al., 1959) to the south-southwest may be a match for the 4.95 km/s layer, but velocities above and below do not correlate either. Furthermore, to the south-southeast, an unreversed profile displays lower velocities 4.4, 5.42, 5.85, and 6.41 km/s. This area has been suggested to be the northernmost extent of oceanic crust (Bouysse, 1988; Pindell and Barrett, 1990). The 5.42 and 6.41 km/s layers are probably interpreted to coincide with typical oceanic layers 2 and 3, respectively, with the thin 5.85 km/s layer representing a possible intermediate velocity. Since the morphology of the northern part of the basin is rugged, and the refraction data are sparse with no clear correlations to surrounding refraction data, the lone profile near 16~ 63~ is interpreted
suggest that the crust of the northern part of the basin is composed of stretched and thinned arc material rather than accreted oceanic crust. Reversed refraction Profile 21 (Officer et al., 1959) displays lower velocities 5.44, 5.95, and 6.52 km/s. Velocities 5.44 and 6.52 km/s are similar to typical oceanic layers 2 and 3. Although the intermediate 5.95 km/s layer of this profile could be interpreted to represent an intermediate velocity from 5.44 to 6.52 km/s, the location of the profile (over the Aves Ridge) dictates that the crust is composed of island arc material. The only other velocity information for the northern part of the basin is an unreversed profile near 16~ 63~ and oriented approximately east-west (Speed and Westbrook, 1984). The lower velocities from this profile (Fig. 15a), 4.42, 4.95, and 5.86 km/s do not correlate well with other profiles for the region. The 5.95 km/s layer of Profile 21
9
l
Il
I
83"W 1
z
m
(A) (B)
150nTI, 20km
~,i'km
(C) ~D F[_ 62"W
~
Fig. 16. Magnetic anomaly profiles: (A, C) correlated; (B, D) displayed in their acquisition position. Heavy lines in (A) and (B) indicate correlated anomalies. Dots in (B) and (D) correspond to locations of correlated anomalies in (A) and (B). Dashed lines correspond to interpreted curvilinear zones of disruption (Fig. 3).
TECTONIC EVOLUTION OF THE GRENADA BASIN to be less reliable than the rest of the refraction information. Therefore, the nature of the crust for the northern part of the basin remains relatively uncertain. Two-dimensional and three-dimensional gravity models suggest that crustal layers gradually thin to the north in this region. The Bouguer gravity over the Grenada basin has been calculated four times using two methods in this work as well as twice by Kearey (1974; Kearey et al., 1975). Anomaly patterns produced by all these calculations are similar. The resultant long-wavelength high over the central part of the basin diminishes gradually north to the Saba Bank, and south to the continental slope of Venezuela. This indicates that the thickness of the crust of the basin also changes gradually over the length of the basin. In his discussion regarding the magnetic anomalies over the Grenada basin, Bouysse (1988) points out that, "... the present great depth to the oceanic basement due to sedimentary overloading combined with a possible location of the Eastern Caribbean in the vicinity of the geomagnetic equator during the Paleocene may contribute to significantly lowering the anomalies' intensity and to blur the original pattern." Although the geophysical data in the Grenada basin are complex, our interpretation suggests that the for-
%
413 mation of the Grenada basin was by near east-west extension. And though the nature of the crust of the northern part of the basin cannot be determined conclusively, a mechanism for which oceanic crust was later disrupted by tectonism related to the bifurcation of the Lesser Antilles arc is suggested. Strike-slip motion (Speed and Westbrook, 1984) along prominent east-northeast faults in the northern part of the basin may have caused the magnetic signature over this part of the basin to be disrupted. McCann and Sykes (1984) identified magnetic anomalies having amplitudes of over 400 nT produced by fracture zones in oceanic crust, during the Cretaceous quiet period, just north of Hispaniola and Puerto Rico. In order to test this hypothesis, magnetics profiles were re-interpreted with the specific purpose to identify, and correlate displaced anomalies, or packages of anomalies. Fig. 16A,C displays correlated anomaly profiles over the basin and Fig. 16B,D shows the same profiles plotted in their acquisition position. For Fig. 16A,B the northern six profiles may be displaced about 45 km to the southwest with respect to the southern five profiles over this area. The central high (dots) is shown in both figures. Fig. 16C,D is a somewhat less convincing correlation. For this set of profiles, dots follow
%
'U
,Q
4
O
%
q
IAI
J
cBi
j
Fig. 17. East-west reconstruction of the Grenada basin using 2 km isobath: (A) present-day configuration; (B) east-west closure of the basin.
414
D.E. BIRD et al.
the correlated prominent low. These anomalies may be displaced about 25 kin. Fig. 17 shows a simple east-west closing of the Grenada basin along the 2 km isobath. Although the fit is generally good, the basin does not close completely in the south and the isobaths overlap in the north. Knowledge of seafloor spreading geometries and the amount of extension of the arc prior to spreading would allow a better constrained reconstruction. Furthermore, the poor fit in the north is consistent with our interpretation that the island arc has been tectonically modified since the formation of the basin. There are two important aspects of the northsouth (Pindell and Barrett, 1990) and northeastsouthwest (Bouysse, 1988) spreading models suggested for the formation of the Grenada basin. There is no real world analogy for these suggested mecha-
nisms. Although Brooks et al. (1984) have classified the Sulu and Celebes Sea basins as possibly being back-arc basins with the westward dipping Philippine plate related to extension, these basins are probably trapped, small ocean basins (McCabe et al., 1985, 1986; Lee and McCabe, 1986; McCabe and Cole, 1987). The only other back-arc basin in the world which exhibits extension oblique to the trench line of the subduction zone is the Andaman Sea basin. However, there is no counterpart for the Aves Ridge here. That is, a remnant island arc is not bounded by basins on both sides. If oblique subduction resulted in oblique back-arc extension and formation of the Grenada basin, then it is reasonable that the Aves Ridge would be composed of oblique-oriented segmented arc sections. Fig. 18A shows a trace of the subduction zone and ridge/transform sections in the Andaman Sea
A I
\
\
'5\
\ \
L k-,, 1 (A)
(B)
Fig. 18. Modified after Mukhopadhyay (1984). (A) Trace of subduction zone and transform/ridge segments in the Andaman Sea back-arc basin. The large vector (/) represents the relative motion of the Indian plate with respect to the basin. Small arrows and line segments represent projections of ridge segments toward the subduction zone. (B) Upside-down trace of subduction zone and transform/ridge segments in the Andaman Sea back-arc basin. The large vector (A) represents the relative motion between the North/South American plates and the Lesser Antilles subduction zone included for comparison of Andaman Sea and Grenada Basins.
TECTONIC EVOLUTION OF THE GRENADA BASIN (after Mukhopadhyay, 1984). The large vector (I) in Fig. 18A represents the direction of motion of the Indian plate with respect to the island arc. Oblique to the arc back spreading is apparent; however, if one projects individual ridge segments back to the subduction zone perpendicular to ridge segments (small arrows and connecting segments), then these segments are subparallel to the arc for the most part. In Fig. 18B the previous trace of the Andaman Sea is turned upside down. The second vector (A) represents suggested oblique subduction for the Atlantic plate and the Lesser Antilles (Pindell and Barrett, 1990). The similarity between the Andaman Sea and Grenada basins is obvious. Furthermore, the ridge/transform pattern is similar to the tectonic model for the evolution of the Grenada basin (Fig. 6) and the pattern predicted by Dewey (1980) regarding the mechanism for back-arc spreading. There are at least nineteen basins which have formed via back-arc extension in the world today. All but possibly two of these basins exhibit extension perpendicular to the subduction zone. Therefore, with the physiography of the study area in mind, near e a s t - w e s t extension for the formation of the Grenada basin is favored. Second, it is felt that conditions which cause back-arc spreading are primarily a result of the interaction between the subducting slab and the overriding plate. Both northsouth (Pindell and Barrett, 1990) and northeastsouthwest (Bouysse, 1988) spreading models for the formation of the basin rely on coupling along the Caribbean/South American plate transform boundary for back-arc extension. In the Southwest Pacific, the Fiji Plateau, South Fiji basin, and Lau basin are back-arc regions which have formed juxtaposed to transform faults. Magnetic lineations over the Fiji Plateau are oriented subparallel to the subduction zone (Brooks et al., 1984). Extension in the South Fiji basin is from a r i d g e - r i d g e - r i d g e triple junction (Weissel, 1981). The spreading ridge of the Lau basin is parallel to the Tonga Trench (Taylor and Karner, 1983) and forms a small transformridge-transform triple junction near the bounding transform fault to the north. These back-arc regions have not formed as predicted by the trench transform model. Hence, the coupling along a transform fault may only have a minor influence on the development of back-arc basins.
CONCLUSION Thorough analyses of magnetic, gravity and seismic data support near e a s t - w e s t extension for the formation of the Grenada basin. Interpretation of magnetic data included trend analyses and forward modeling (Bird et al., 1993). We have integrated
415 these results with analyses of gravity and seismic data to develop a more coherent model for the evolution of the Grenada basin. This evolution involves two major tectonic episodes. The basin is interpreted to have formed by near e a s t - w e s t back arc seafloor spreading in early Tertiary time. In late Tertiary time, the northern part of the basin was affected by compressional forces causing the Lesser Antilles island arc to split, displacing a section of the Aves Ridge, and disrupting the crust of Grenada basin, including its magnetic signature.
REFERENCES
Bandy, W.L. and Hilde, T.W.C., 1983. Structural features of the Bonin Arc: implications for its tectonic history. Tectonophysics, 99: 331-353. Barker, E and Hill, I.A., 1981. Back-arc extension in the Scotia Sea. Philos. Trans. R. Soc. London, A300: 249-262. Bibee, L.D., Shor, G.G. and Lu, R.S., 1980. Inter-arc spreading in the Mariana Trough. Mar. Geol., 35: 183-197. Bird, D.E., 1991. An Integrated Geophysical Interpretation of Grenada Basin. M.S. Thesis, Univ. of Houston, 316 pp. Bird, D.E., Hall, S.A., Casey, J.E and Millegan, ES., 1993. Interpretation of magnetic anomalies over the Grenada Basin. Tectonics 12 (5): 1267-1279. Bougault, H., Dmitriev, L., Schilling, J.G., Sobolev, A., Joran, J.L. and Needham, H.D., 1988. Mantle heterogeneity from trace elements: MAR triple junction near 14~ Earth Planet. Sci. Lett., 88: 27-36. Bouysse, E, 1984. The Lesser Antilles arc: structure and geodynamic evolution. Init. Rep. DSDP, 78A: 83-103. Bouysse, E, 1988. Opening of the Grenada back-arc basin and evolution of the Caribbean plate during the Mesozoic and Early Paleocene. Tectonophysics, 149: 121-143. Boynton, C.H., Westbrook, G.K., Bott, M.H.E and Long, R.E., 1979. A seismic refraction investigation of crustal structure beneath the Lesser Antilles island arc. Geophys. J. R. Astron. Soc., 58: 371-393. Brooks, D.A., Carlson, R.L., Harry, D.L., Melia, EJ., Moore, R.E, Rayhorn, J.E. and Tubb, S.G., 1984. Characteristics of back-arc regions. Tectonophysics, 102: 1-16. Clark, T.F., Korgen, B.J. and Best, D.M., 1978. Heat flow in the eastern Caribbean. J. Geophys. Res., 83:5883-5891. Cross, T.A. and Pilger, R.H., Jr., 1982. Controls of subduction geometry, location of magmatic arcs, and tectonics of arc and back-arc regions. Geol. Soc. Am. Bull., 93: 545-562. Curray, J.R., Moore, D.G., Lawver, L.A., Emmel, EJ., Raitt, R.W., Henry, M. and Kieckhefer, R., 1979. Tectonics of the Andaman Sea and Burma: In: J.S. Watkins, L. Montadert and EW. Dickerson (Editors), Geological and Geophysical Investigation of Continental Margins. Am. Assoc. Pet. Geol., pp. 189-198. Dewey, J.E, 1980. Episodicity, sequence, and style at convergent plate boundaries. In: D.W. Strangway (Editor), The Continental Crust and its Mineral Deposits. Geol. Assoc. Can., Spec. Pap., 20: 553-573. Dix, C.H., 1955. Seismic velocities from surface measurements. Geophysics, 20: 68-86. Duncan, R.A. and Hargraves, R.B., 1984. Plate tectonic evolution of the Caribbean region in the mantle reference frame. In: W.E. Bonini, R.B. Hargraves and R. Shagan (Editors), The Caribbean-South American Plate Boundary and Region Tectonics. Geol. Soc. Am. Mem., 62: 81-93.
416 Edgar, N.T., 1968. Seismic Refraction and Reflection in the Caribbean Sea. Ph.D. Thesis, University of Columbia. Ewing, J.I., Officer, C.B., Johnson, H.R. and Edwards, R.S., 1957. Geophysical investigations in the eastern Caribbean: Trinidad Shelf, Tobago Trough, Barbados Ridge, Atlantic Ocean. Bull. Geol. Soc. Am., 68: 897-912. Fink, L.K., 1968. Marine Geology of the Guadeloupe Region, Lesser Antilles Arc. Ph.D. Dissertation, University of Miami, FL. Fink, L.K., 1970. Field guide to the island of La Desirade with notes on the regional history and development of the Lesser Antilles island arc. Int. Field Inst. Guideb. to the Caribbean Island-Arc System, Am. Geol. Inst. N.S.P., pp. 287-302. Fox, P.J. and Heezen, B.C., 1975. Geology of the Caribbean crust. In: A.E.M. Nairn and EG. Stehli (Editors), The Ocean Basins and Margins. Plenum Press, London, 3, pp. 421-466. Ghosh, N., Hall, S.A. and Casey, J.E, 1984. Seafloor spreading magnetic anomalies in the Venezuelan Basin. In: W.E. Bonini, R.B. Hargraves and R. Shagan (Editors), The CaribbeanSouth American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162: 65-80. Hayes, D.E., Houtz, R.E., Jarrard, R.D., Mrozowski, C.L. and Watanabe, T., 1978. Crustal structure. In: D.E. Hayes (Editor), A Geophysical Atlas of East and Southeast Asian Seas. Geol. Soc. Am., Map Chart Ser., MC-25. Hussong, D.M. and Uyeda, S., 1981. Tectonic processes and the history of the Mariana Arc: a synthesis of the results of deep sea drilling Leg 60. Init. Rep. DSDP, 60: 909-929. Karig, D.E., 1971. Origin and development of marginal basins in the western Pacific. J. Geophys. Res., 76: 2542-2561. Kearey, E, 1974. Gravity and seismic reflection investigations into the crustal structure of the Aves Ridge, eastern Caribbean. Geophys. J.R. Astron. Soc., 38: 435-448. Kearey, E, Peter, G. and Westbrook, G.K., 1975. Geophysical maps of the eastern Caribbean. J. Geol. Soc. London, 131: 311-321. Lee, C.S. and McCabe, R.J., 1986. The Banda-Celebes-Sulu Basin: a trapped piece of Cretaceous-Eocene oceanic crust? Nature, 322:51-54. Lee, C.S., Shor, G., Bibee, L.D., Lu, R.S. and Hilde, T.W.C., 1980. Okinawa Trough, origin of a back-arc basin. Mar. Geol., 35: 219-241. Ludwig, W.J., Nafe, J.E. and Drake, C.L., 1971. Seismic refraction. In: A.E. Maxwell (Editor), The Sea, 1. John Wiley, New York, 4, pp. 53-84. McCabe, R.J. and Cole, J.T., 1987. Speculations on the Late Mesozoic and Cenozoic evolution of the southeast Asian margin. Trans. 4th Circum-Pacific Energy and Mineral Resources Conf., 4: 375-394. McCabe, R.J., Lee, C.S. and Hilde, T.W.C., 1985. The SuluCelebes-Banda b a s i n - a trapped piece of oceanic crust (abstr.). Eos, 66:1078. McCabe, R.J., Hilde, T.W.C., Cole, J.T., Sager, W. and Lee, C.S., 1986. Sulu-Celebes-Banda Basins: a trapped piece of Cretaceous to Eocene oceanic crust (abstr.). Bull. Am. Assoc. Pet. Geol., 70: 930. McCann, W.R. and Sykes, L.R., 1984. Subduction of aseismic ridges beneath the Caribbean Plate: implications for the tectonic and seismic potential of the northeastern Caribbean. J.
D.E. BIRD et al. Geophys. Res., 89:4493-4519. Mukhopadhyay, M., 1984. Seismotectonics of subduction and back-arc rifting under the Andaman Sea. Tectonophysics, 108: 229-239. Officer, C.B., Ewing, J.I., Edwards, R.S. and Johnson, H.R., 1957. Geophysical investigations in the eastern Caribbean: Venezuelan Basin, Antilles Island Arc, and Puerto Rico Trench. Bull. Geol. Soc. Am., 68: 359-378. Officer, C.B., Ewing, J.I., Hennion, J.E, Harkrider, D.G. and Miller, D.E., 1959. Geophysical investigations in the eastern Caribbean: summary of 1955 and 1956 cruises. Phys. Chem. Earth, 3: 17-109. Okuma, S., Nakatsuka, T., Makino, M. and Morijini, R., 1990. Aeromagnetic constraints on the basement structure of the Okinawa Trough and East China Sea Basin (abstr.). Soc. Explor. Geophys. Expanded Abstr. Biogr., 1: 594-597. Pindell, J.L. and Barrett, S.E, 1990. Geological evolution of the Caribbean region; a plate tectonic perspective. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. Geol. Soc. N. Am., H, pp. 405-432 Pindell, J.L., Cande, S.C., Pitman III, W.C., Rowley, D.B., Dewey, J.E, LaBrecque, J. and Haxby, W., 1988. A plate-kinematic framework for models of Caribbean evolution. Tectonophysics, 155: 121-138. Poehls, K.A., 1978. Intra-arc basins: a kinematic model. Geophys. Res. Lett., 5: 325-328. Ross, M.I. and Scotese, C.R., 1988. A hierarchical tectonic model of the Gulf of Mexico and Caribbean region. Tectonophysics, 155: 139-168. Sleep, N.H. and Toskoz, M.N., 1971. Evolution of marginal basins. Nature, 233: 548-550. Speed, R.C., Westbrook, G.K., et al., 1984. Lesser Antilles arc and adjacent terranes. Atlas 10, Ocean Margin Drilling Program, Regional Atlas Series, Marine Science International, Woods Hole, Mass., 27 sheets. Tamaki, K., 1985. Two modes of back-arc spreading. Geology, 13: 475-478. Taylor, B., 1979. Bismark Sea: evolution of a back-arc basin. Geology, 7: 171-174. Taylor, B. and Karner, G.D., 1983. On the evolution of marginal basins. Rev. Geophys., 21: 1727-1741. Tomblin, J.E, 1975. The Lesser Antilles and Aves ridge. In: A.E.M. Nairn and EG. Stehli (Editors), The Ocean Basins and Margins, Plenum Press, London, 3, pp. 467-500. Uyeda, S. and Kanamori, H., 1979. Back-arc opening and the mode of subduction. J. Geophys. Res., 84: 1049-1061. Warner, A.J., Jr., 1991. The Cretaceous age sediments of the Saba Bank and their petroleum potential. Trans. 12th Caribbean Geol. Conf., St. Croix, U.S.V.I., Miami Geological Society. Weissel, J.K., 1980. Evolution of the Lau Basin by the growth of small plates. In: M. Talwani and W.C. Pittman III (Editors), Island Arcs, Deep Sea Trenches, and Back-arc Basins. Maurice Ewing Ser., Am. Geophys. Union, 1: 429-436. Weissel, J.K., 1981. Magnetic lineations in marginal basins of the west Pacific. Philos. Trans. R. Soc. London, A300: 223-247. Westbrook, G.K., 1975. The structure of the crust and upper mantle in the region of Barbados and the Lesser Antilles. Geophys. J.R. Astron. Soc., 43: 201-242.
Chapter 16
Sequence Stratigraphy of the Eastern Venezuelan Basin
J. DI CROCE, A.W. BALLY and E VAIL
The Eastern Venezuelan Basin and its offshore continuation is a Neogene foredeep superimposed on a Mesozoic passive margin. The basin contains petroleum reserves amounting to some 35,000 million bbls and is the focus of very active exploration. Four major tectonically controlled Phanerozoic unconformities separate distinct structural regimes, i.e. (1) an ill-defined Paleozoic/pre-Jurassic pre-rift phase, (2) a Jurassic syn-rift phase, (3) Cretaceous to Oligocene passive margin phase, and (4) and a Neogene foredeep phase. The pre-Cretaceous unconformity corresponds to the classical breakup unconformity often found on passive margins. The pre-Neogene unconformity is the basal foredeep unconformity encountered in most foreland basins of the world. The Cretaceous to Paleocene passive margin sequence of eastern Venezuela is subdivided into five second-order transgressiveregressive cycles bounded by a 131 Ma (basal Cretaceous) sequence boundary, four maximum flooding surfaces with the inferred age of early Aptian (111 Ma), late Albian (98 Ma), middle Cenomanian (95 Ma), middle Turonian (91.5 Ma) and a Late Paleocene sequence boundary (58.5 Ma). A Late Paleocene to Eocene second-order cycle (58.5 Ma-36 Ma) is followed by the Oligocene which is subdivided into two third-order cycles bounded by 36 Ma, 30 Ma and 25.5 Ma sequence boundaries. The second-order sequence boundaries reflect sea-level fluctuations that are superimposed on a thermally subsiding passive margin. The purely structurally controlled post-Oligocene to pre-Lower Miocene (25.5 Ma) basal foredeep unconformity is associated with the sudden deepening of the passive margin in response to the incipient emplacement of the ancestral Serranfa del Interior far to the northwest of its present-day position. The Miocene of the eastern Venezuela foredeep is characterized by aggrading and prograding deltaic sequences located in the western part of the basin. These sequences consist of three second-order sequences defined by 25.5 Ma, 16.5 Ma and 10.5 Ma boundaries. Two regional flooding events occurred at 16 Ma and 13.4 Ma. However, the Serravallian (13.4 Ma) event can only be recognized in the onshore. Following this flooding event subsidence of the foredeep continued and second- and higher-order sequences formed reflecting sea-level changes and increased sediment supply from the rising mountain ranges of Venezuela. In the offshore a Late Miocene (5.5 Ma) unconformity is associated with deeply incised submarine canyons. This deep-water unconformity is not obviously related to a structural event; instead it may reflect a worldwide Messinian sea-level lowering. The Plio-Pleistocene is characterized by an overall prograding deltaic front in the Orinoco platform area. About sixteen Neogene third-order sequences are recognized on seismic profiles and are tentatively correlated with Haq et al.'s (1987) sequence boundaries.
INTRODUCTION
The onshore Eastern Venezuelan Basin (Fig. 1) is the second richest hydrocarbon province of South America after the Maracaibo Basin. More than 12,000 wells have been drilled and 35 giant fields and 260 minor fields have been developed during 90 years of exploration and production (Gonzalez de Juana et al., 1980). Since 1980 many other fields have been discovered. The generalized distribution of the major fields is shown in Fig. 2. Excluding the reserves of the Orinoco tar belt which occupies parts of the updip margin of this basin the ultimate reserves of the Eastern Venezuelan Basin amount to
45,000 million bbls of which some 11,000 million bbls have already been produced. The total area (i.e. onshore and offshore) of the Eastern Venezuelan Basin is about 200,000 km 2. Elongate and asymmetric, the basin contains up to 8-km-thick Tertiary sediments. The Precambrian underlying the south flank of the Eastern Venezuelan Basin dips gently northward. To the north of the basin the fold thrust-belts of the Serranfa del Interior (Interior Mountain Range), the Gulf of Paria and Trinidad involve sediments that are of the same age as the basin fill. The onshore Eastern Venezuelan Basin is subdivided into two sub-basins: the Gu~irico sub-basin
Caribbean Basins. Sedimentary Basins of the World, 4 edited by R Mann (Series Editor: K.J. Hsti), pp. 419-476. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
420
J. D I C R O C E et al.
Fig. 1. Northern Venezuela, physiographic provinces and sedimentary basins. The dashed outline is the study area bounded to the north by the front of the Monagas foothills. Abbreviations: A F -- Anaco Fault; B A P = Barbados Accretionary Prism; B F -- Bocon6 Fault; E P F = E1 Pilar Fault; I F = Icotea Fault; L B F : Los Bajos Fault; S F F = San Francisco Fault; SER. I N T = Serranfa del Interior; O F = Oca Fault; O D F -- outer deformation front of Monagas foothills" UF -- Urica Fault.
to the west and the Maturfn sub-basin to the east (Fig. 1). The boundary that separates the Gu~irico sub-basin from the folded belt of the Serranfa del Interior and Maturfn sub-basin is the Utica fault system, i.e. the buried lateral ramp of the Serranfa del Interior (Roure et al., 1994). Farther south the Gmirico and Maturfn sub-basins are separated by the complex structures associated with the Anaco fault system and its associated inversion structures (Villaroel, 1993; Bejarano et al., 1996). The offshore continuation of the Eastern Venezuelan Basin is the offshore Orinoco platform or Plataforma Deltana which to the south merges with the South Atlantic passive margin of Guyana.
P U R P O S E O F T H E STUDY
This study was primarily undertaken to provide an integrated view of the onshore and offshore subsurface stratigraphy of the Maturfn sub-basin (in short the Maturfn Basin!) and its offshore continuation. The area was selected to illustrate the transition from a pristine Atlantic-type passive margin to a transpressional foredeep. This transition may be used as a model for the development of earlier transpressional foredeeps that developed farther west in Venezuela. Another key objective of this study was the application of modem sequence-stratigraphic principles to an area where the biostratigraphic control is limited but where the seismic data permit reasonable resolution into sequences that at least may be compatible
with the sequences proposed by Haq et al. (1987). Combining regional tectonics with sequence stratigraphy we will attempt to illustrate the interaction of structural regimes with eustatic and/or erosional events. The conventional stratigraphy of the area has been reviewed in great detail in Gonzalez de Juana et al. (1980) and the International Stratigraphical Lexicon (1956). A more detailed glossary is also contained in the thesis that forms the basis of this paper (Di Croce, 1995).
DATA BASE
All data for this study were kindly provided by Lagoven S.A. and Corpoven S.A., both affiliates of Petroleos de Venezuela S.A. (ED.V.S.A.). Some 14,000 km of seismic profiles and more than 150 onshore and two offshore wells were used. The uniform geographical distribution and the overall quality of these data provided an adequate grid to support a regional study. Fig. 3a shows the location of eighteen key wells that were selected based on the quality of paleontological data, depth of penetration, availability of synthetic seismograms, time-converted logs and regional position. The composite logs of these wells include spontaneous potential (SP), gamma-ray (GR), resistivity (Res) and conductivity curves. Selected paleontological data from these wells will also be summarized in this study. The seismic reflection profiles used for this paper were acquired between 1969 and
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN
421
Fig. 2. Major oil fields of the Eastern Venezuelan Basin. Excluding the Orinoco tar belt to the south, ultimate reserves for this basin are estimated to be in the order of 45,000 million bbls.
1993. Sixty percent of these seismic lines are located offshore and forty percent are onshore. Records range between 6 and 9 s two-way travel time (twt) and most of the profiles are migrated. Most of the original seismic data in this paper are shown in the form of line drawings. The location of these line drawings is shown in Fig. 3b. The drawings were directly derived from the seismic profiles and represent our subjective interpretation. The correlations shown have been traced with reasonable confidence throughout the study area. Line drawings permit legible reproduction and avoid the inclusion of bulky seismic profiles in this paper. However, to further substantiate our interpretations a few selected segments of the original seismic profiles are included. Fig. 3c shows the location of figures that illustrate these seismic details.
A REVIEW OF PAST WORK Early exploration and geosynclinal concepts
During the 90-year long history of hydrocarbon exploration many regional and local studies have been carried out in the Eastern Venezuelan Basin and the folded belts of eastern Venezuela and Trinidad. Most of the published key references are included in the bibliography of Gonzalez de Juana et al. (1980) and the course notebook of Audemard and Lugo (1997); however, much of the fine work done by the staff of many oil companies remains unpublished. Liddle (1928), in his classic 'Geology of Venezuela and Trinidad' introduced the term 'Orinoco
Geosyncline' for the Eastern Venezuelan Basin and until the late sixties most regional stratigraphic studies were done in the context of geosynclinal theories. During the 30's and 50's, knowledge of the Eastern Venezuelan Basin greatly increased thanks to the work of many geologists (e.g. Hedberg, 1937, 1942; Hedberg and Pyre, 1944; Gonzalez de Juana, 1947; Funkhouser et al., 1948). These authors described the general stratigraphy of the basin and defined most of the sedimentary units and their type localities on the outcrop or in subsurface. Hedberg (1950) provided a remarkable synthesis, outlining the geologic history of the basin. He distinguished the present eastern Venezuelan fold and thrust belt from a series of depositional basins which existed throughout the Cretaceous and early Tertiary. Hedberg (1950) further recognized the profound difference in the depositional history and character of the sedimentary units which was controlled largely by the migrated position of the axis of the 'Eastern Venezuela Geosyncline'. Renz et al. (1958) in an other outstanding contribution describes the evolution of the Eastern Venezuelan Geosyncline beginning with a Late Jurassic or Early Cretaceous orthogeosynclinal phase. The Late Cretaceous orogeny transformed the basin into a geanticlinal welt which forced the migration of the geosynclinal axis to the south and formed an exogeosyncline or foredeep which appeared during the Early Oligocene. The authors also include a series of restored cross-sections, paleogeographic and isopach maps for different time intervals. Salvador and Stainforth (1968) presented an integrated synthesis of the stratigraphy of the Eastern
422
J. DI C R O C E et al.
t~,O
CARIBBEAN
MARGARITAI.
62 ~
61 ~
59 ~
"7
SEA 3P
GULF EPF
o
OF
TRINIDAD 0
.,4..._.. MAT~RIN
liT
,,,, ,\o o
eS
ORINOCO
eB
DELTA
~-9~
eA
Ge eF
7-
eE 0 I
eD
a)
640
60 ~ r~,3o
CARIBBEAN
MAI:K~RITA I.
;2~
80 km |
~ 59 ~
61 ~
590
SEA Fig. gb GULF
EPF
OF
TRINIDAD
Fig. 11a Fig. 7a
MATURIN
Fig. 25a ~TLANTIC OCEAN
DELTA
__
Fig. 6 a
9 9a
80 krn 9
b)
1~
630
CARIBBEAN
MAI:II~RITA I.
2~
61 ~
I
I
60 ~ ~
59 ~
60 ~
590
I
SEA 11b GULF
EPF
TRINIDAD
ATLANTIC
MATURIN 9
Fig.24
0
OCEAN
ORINOCO DELTA
Fig. 20b
t ~ R g . 22a & 22b Fig. 21a
Fig. 16a & 16b Fig. 18a & 18b Fig. 21b & 21c .23a
o
| I
c)
640
620
59 ~
Fig. 3. Index maps for the Eastern Venezuelan Basin. (a) Base map for well data used in this study. Offshore bathymetric contours in meters. Point J is a composite surface section for the Serranfa del Interior (Fig. 17). Abbreviations: AF = Anaco Fault; EPF = E1 Pilar Fault; LBF -- Los Bajos Fault; ODF -- outer deformation front of Monagas foothills; SFF = San Francisco Fault; UF -- Urica Fault. (b) Map showing the location of all the line drawings presented in this paper. (c) Map displaying the location of the segments of seismic profiles included in this paper.
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN Venezuelan Basin and Trinidad based on the planktonic foraminiferal zonation and the interpretation of surface exposures and wells in Trinidad. The most important stratigraphic synthesis was done by Gonzalez de Juana et al. (1980) who summarized the stratigraphy of all of Venezuala in great detail. The reader is referred to this publication for the definition of formation names and their type localities. Arnstein et al. (1985) updated the stratigraphic and structural framework of the Eastern Venezuelan Basin in a petroleum geologic context.
The advent of plate tectonics Beginning in the midsixties plate tectonics brought about a basic change in the understanding of eastern Venezuela culminating with the fine reconstructions of Pindell and Dewey (1982), Pindell (1985), Ross and Scotese (1988), Pindell and Barrett (1990), Stephan et al. (1990) and Dercourt et al. (1993). In a plate tectonic context the Eastern Venezuelan Basin formed as the result of the complex interaction between the South America, North America and Caribbean lithospheric plates. Three major tectonic stages control the evolution of the basin. These stages are linked to the postJurassic relative motions of South America with respect to North America and to the Tertiary eastward motion of the Caribbean Plate. The first stage from Middle Jurassic to Late Cretaceous (165 Ma-80 Ma) shows a northwest-southeast divergence. The second stage from Campanian to Eocene (80 Ma-49 Ma) has negligible relative motions between South and North America. However, this stage also marks the inception of the very important relative eastward motion of the Caribbean Plate which dominates from the Eocene to the present (49 Ma-0 Ma). Based on this plate tectonic scenario the following tectonic settings are now differentiated: (1) the Triassic to Late Jurassic rift phase (breakup of Pangea) (Bartok, 1993); (2) the passive margin phase which is Late Jurassic to Late Cretaceous in western Venezuela and Late Jurassic to Oligocene in eastern Venezuela) (e.g. Erikson and Pindell, 1993); (3) the active margin phase (Late Cretaceous to present). This active margin phase is characterized by an eastward-shifting flexural foredeep depression (Audemard, 1991; Lugo, 1991; Lugo and Mann, 1995; Audemard and Lugo, 1996). Fig. 4 shows the Tertiary evolution of the Venezuelan margin based on the reconstructions of Stephan et al. (1990) and the position of these shifting foredeep depocenters. During the Early Paleocene, sea floor spreading between North and South America ceased and the Caribbean Plate began to migrate toward the northeast. Initially this transpressional collision has no effect on the eastern margin of Venezuela, in con-
423
trast to western Venezuela from where since the Late Paleocene (Fig. 4a) transpressional deformation advanced diachronously toward the east along the northern border of South America, in direct response to the eastward migration of the Caribbean Plate (e.g. Pindell and Barrett, 1990; Lugo and Mann, 1995). The main onset of transpression in western Venezuela was during early-Middle Eocene (Fig. 4b) but progressively younger transpression affected eastern Venezuela and Trinidad during Late Oligocene-Middle Miocene (Fig. 4b,c) and continued until today. Other major events recorded during Eocene time also include the initiation of the volcanism in the Lesser Antilles Arc (Pindell and Barrett, 1990). The Oligocene is a tectonically quiet period and perhaps reflects a slowdown of the eastward motion of the plate and eastern Venezuela survived as a passive margin (Stephan et al., 1990). During earlyMiddle Miocene to present the Maturfn Basin and its offshore continuation finally became defined as a foredeep in response to southeastward thrusting of the Serranfa del Interior and foreland loading (Fig. 4c,d). Today the Eastern Venezuelan Basin is located near the junction of the South American Plate, the Caribbean Plate and the Atlantic Plate.
Exploration and stratigraphy in the eighties and nineties The discovery of the giant E1 Furrial fields in 1986 gave a renewed impetus to exploration in eastern Venezuela which was followed by the publication of a new generation of papers that focussed on the hydrocarbon habitat and the stratigraphy of the foothills of the Serranfa del Interior (i.e. the Monagas foothills). Carnevali (1988, 1989) described the hydrocarbon habitat of the new fields in the context of two tectonic phases, i.e. a passive margin phase and a foreland fold-thrust phase (see also Chevalier et al., 1995). Erlich and Barrett (1992) outlined the petroleum geology of the Eastern Venezuela Foreland Basin from a regional perspective. Their work when combined with detailed stratigraphic data compiled during the past four decades, has helped to constrain and refine models of the geohistory of northeastern Venezuela and substantial undiscovered hydrocarbon resources may still be found along the present trend of giant fields. Parnaud et al. (1995) and Gallango and Parnaud (1995) provide an excellent and up-to-date overview over the petroleum geology of the Eastern Venezuelan Basin. For a structural study of the Serranfa del Interior and its foothills (the Monagas foothills) refer to Passalacqua et al. (1995). Erikson and Pindell (1993) based on the subsidence analysis of a composite stratigraphic section of the northern Serranfa del Interior conclude that the Cretaceous-Eocene interval is characterized
424
J. DI CROCE et al. a) LATE PALEOCENE (59 Ma) u
j3
. _ ~ ~ . _.;~
c) EARLY MIOCENE(20 Ma)
d) PRESENT"
o
YB
...._ ...
//~BRK,..x VB
" ~" ~. Stephan et a ~ ~ . : . . ' J
,, ~,,COC__';MB .'"
~.d
SA
J - ~
~
W
.f. JJBR~
ll,, StephanC ~~WICO~3 jc'ACe, al{19,0) CCC
VB
[A] ~
SA rvo "~
.,
Fig. 4. Four stages of the development of the northern and southern Caribbean boundary from Eocene to Present associated with the eastward migration of the Caribbean Plate modified after Stephan et al., 1990. The diagonal pattern shows the center of the foreland foredeep depression. (a) During the Paleocene the Caribbean island arc system was located at the northwestern corner of the northern Venezuela passive margin. The remainder of northern Venezuela was part of a passive/(?)transform margin. (b) During the Middle Eocene compressional deformation due to the Caribbean-South American convergence affected much of northwestern Venezuela where a foredeep formed in the Lake Maracaibo area. (c) Continuation of the oblique convergence of Caribbean Plate and South America. The result is the emplacement of a transpressional folded belt and the associated development of an eastward migrating foredeep. (d) From Late Pliocene to Present the kinematic change in the Eastern Venezuelan Basin is associated with a decrease of the N W - S E contraction and increase in the strike-slip deformation. The center of the foredeep moves into the Orinoco platform area. Abbreviations: A C = Andes Cordillera; AR = A v e s Ridge; BR -- Barbados Ridge; BR = Beata Ridge; CA = Central America; CB =- Colombian Basin; CCC = Central Colombian Cordillera; CP = Cocos Plate; CT -- Cayman Trough; CU = Cuba; GB = Grenada Basin; GYB = Guyana Basin; HI -- Hispaniola; LA = Lesser Antilles; MB = Maracaibo Basin; NP -- Nazca Plate; OCC = Occidental Colombian Cordillera; PR = Puerto Rico; SA = South America; VB -- Venezuelan Basin; YU - Yucatan; YB = Yucatan Basin.
by Atlantic-type passive margin thermal subsidence of the lithosphere. According to these authors the Oligocene-Miocene interval reveals a markedly increased subsidence due to tectonic loading. An unconformity separating the passive margin from the foredeep interval is interpreted as the passage of a peripheral bulge associated with the tectonic load of the Venezuelan Coastal Ranges. In this paper we will present seismic-stratigraphic evidence to add detail to Erikson and Pindell's (1993) concept. Regarding the stratigraphy of our area Azavache et al. (1996) recognize a basal foredeep unconformity that separates the Lower Miocene (Oficina Formation) from the underlying Cretaceous. A deepwater stage is represented by the Lower and Middle Miocene upper Oficina and Freites formations in the foreland, and the Carapita Formation in the Monagas foothills to the north. These are overlain by the
Upper Miocene (lower-middle La Pica Formation) shelf-slope deposits and the fluvial-deltaic PlioPleistocene (upper La Pica, Mesa and Las Piedras formations) deposits. Bejarano et al. (1996) and Crux et al. (1996) emphasize a depocenter switch from the western Late Eocene to Early Miocene Gufirico sub-basin to the eastern Maturfn sub-basin around 12.5 Ma, i.e. considerably later than the obvious basal foredeep unconformity which will be discussed in detail in this paper. Gonzales et al. (1996a,b) also did a very detailed sequence-stratigraphic study in the Maturfn subbasin supplemented with chronostratigraphic charts for parts of the onshore of eastern Venezuala. We will suggest a somewhat different correlation for the late Neogene but perhaps more importantly our paper tries to tie the onshore stratigraphy of east-
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN ern Venezuela with the stratigraphy of the Orinoco offshore. Much less is written about the eastern part of the area, i.e. the offshore Orinoco Delta. Only two papers provide most of the information about this area. Leonard (1983) described the geology and hydrocarbon accumulation of the Columbus Basin, which corresponds to the southeastern Trinidad offshore. He described the structure and stratigraphy of the basin during the Neogene and the major factors controlling hydrocarbon accumulation. In his thesis Prieto (1987) described the subsurface geology of the Orinoco platform and, based on some of the same data used for this paper, established in some detail the timing and distribution of growth-fault provinces in the area. Numerous authors have discussed the stratigraphy of Trinidad (e.g. Kugler, 1953, 1959; Barr and Saunders, 1965; Carr-Brown and Frampton, 1979; Persad, 1979, 1985). Our own study is best compared with Aden and Bierley (1996) who recognize in Trinidad a Cretaceous to Upper Oligocene deep-water passive margin sequence, which correlates with our passive margin sequence. This sequence is overlain by an Upper Oligocene-Middle Miocene synorogenic deep-water sequence and an Upper Miocene through Pliocene shelfal sequence. In contrast, Babb et al. (1996) single out a Paleogene megasequence that
425
separates an Upper Jurassic to Lower Cretaceous passive margin megasequence from the overlying Neogene megasequence.
REGIONAL TRANSECTS ACROSS THE EASTERN VENEZUELAN BASIN I n t r o d u c t i o n and structural setting
Although this paper is mainly a sequence-stratigraphic analysis of the Eastern Venezuelan Basin the regional setting will first be illustrated with line drawings of selected regional transects. The structural setting for the regional transects across the Eastern Venezuelan Basin is shown in Fig. 5 which shows the end product of an oblique convergence of the relatively eastward-moving Caribbean Plate with the northwestward-moving South American Plate. The Eastern Venezuelan Basin is the foredeep (or foreland basin) directly associated with the transpressional Coast Ranges of northern Venezuela which are a product of this collision. The Precambrian to (?) Paleozoic basement underlying the foredeep dips gently underneath the frontal portions on the foothills of the Serranfa del Interior and their eastern continuation in southern Trinidad. The Eastern Venezuelan Basin merges to the southeast with the Atlantic passive mar-
Fig. 5. Block diagram of the southern Caribbean showing the major tectonic elements as related to the Eastern Venezuelan Basin. The Atlantic lithosphere dips to the west and the Caribbean lithosphere dips to the south. The relationship between the two subducted lithospheric slabs is poorly understood.
426
J. DI CROCE et al.
FRONTAL THRUST FAULT
NW
a) o
SE A LINE 7a
B
LINE 9b m
,
............
z._.
.............
. ......
[...~. . . . . . . . . . .
;
. .................
2
LOWER MIOCENE-PLEISTOCENE
2 A (,) (/'}
I.-
4
4
- ::,:,.
51 ~ U N I T V
10 km i;,,.;i~.~; .... _
L E G E N D
LOCATION
S TRA TIGRAPH Y
41o' ~ r~o\ A
KEY WELL
=--=
A
WELL PROJECTED
.......... BASE OF PLIOCENE
(~
CROSS POINT WITH THE SECTION ( ~
' ~ , ~
= - , T O P OF MIDDLE MIOCENE ....
~-,SB,d-~ SEQUENCE BOUNDARY
~
var~A~~ ~
TOP OF PLIOCENE
~--TOP
REFLECTION TERMINATION
EVERSFEAULT
NORMAL FAULT
~"
~-TOP OF LOWER MIOCENE OF OLIGOCENE TOP OF PALEOCENE-EOCENE
......
TOP OF CRETACEOUS INTRA-CRETACEOUS
SHALE DIAPIR/MUD VOLCANO
K-2
EPF
S E A
"~
~
j.,/~-." p'
",
o
l
I
=
ONSHORE SERRA~N~ OFFSHORE i
FORELAND
QUAT PLEISTOCENE
"
PLIOCENE
J
I I ]~~MESNLAS
SEQUENCE STRATIGRAPHY
UNITS
PIEDRAS Fms
VIII
i
UJ a
FOREDEEP '.~
Fm OFICINA Fm. \~~MERECURE
LM Oligocene Eocene
'
! '
Paleocene
! ~ j l
POST-RIFT
CRETACEOUS-
!
VII
LU LU r'r'~
FREITES
MIOCENE MM
O!
~. n'<
Fm.
Sequence Boundary
SE. 13_
VI
LL
I
I L,')~Os j~w ~4~ARATAS Fm. ~ i / VIDONO Fm. i L/ SAN JUAN Fm. i SAN ANTONIO Fm. 7 / QUERECUAL Fm. I CHIMANA F m . . - - ~ '
V IV III
L i
i"r
'~
Maximun Floodingi Surface
JURASSIC
RIFTING
PALEOZOIC
PRE-RIFTING
;
I ~
'
"lm"l~
SB-3
2.3m 2.2 - - -
--
2.1
O1 - -
_ _ - - P 1 - -
UJ
--
"/Vk/k/Vk/Vk/Vk/k/k/V~ 'v
~Vk~
I
K4
--
-- K3 ---K2--
Fm.
i
[~
K1--
SB-2
'"
RIFT
'--
w
~ =
~3.16 3.--~~' 3.14 3 . - ~ 3.12 3.-~-5_ ,3.10 3.9 -3.8 -3.7 - 3.6 -3.5 -- -- M2 3.4 -3.3 - - - M1-3.2 - 3.1 - -
' < ~ TEMBLADOR Gr. EL CANTIL Fm. ~ N Q U I N
PRECAMBRIAN
~o ~
TOP OF BASEMENT
PHASE
U
10~
9o' j~."
NW
rr
11~
~''"-.~__-c'~Z ....i
STRATIG RAPHY
UM
5~o - -
/
o
K-1
b) AGE
c~o
..__,,,~
A
Pz PALEOZOIC .....
6~o
~k ~F ~TRINIDAD
~
Jr JURASSIC
BASEMENT OCEANIC CRUST
i
K-3
C A R I B B E A N
~'~-~
MAP
6,~o
.10
....
..............
~o
i-?- J-: 7 L
CRYSTALLINE
1,
1.1
--
~ J w v w v w v w ~ SB-1
BASEMENT
Fig. 6. The superposition of the eastern Venezuela foredeep sequence on the passive margin of South America. (a) NW-SE regional transect across the offshore Orinoco platform. The onlap on the basement is the SB-1 surface (shown on the diagram below), which here is merged with the SB-2 surface. The basal foredeep unconformity, SB-3 in (b), separates the underlying passive margin sequence from the overlying foredeep sequence. (b) Stratigraphic chart showing general chronostratigraphy, tectonic phases and major sequence-stratigraphic boundaries of the Eastern Venezuelan Basin. Common onshore formation names (e.g. Gonzalez de Juana et al., 1980) are shown in their correct position. There probably is a significant hiatus separating the pre-rift Paleozoic from the syn-rift Jurassic.
Zd
IN31a3SV8 dO d O l OlOZO3lVd Jr
~-)I £-)I ................
DISSVhlNr L-)I
SIqOBOVIBIdO-VEIINI SNO3OVI3UO dO d O l . . . . . .
3N3009130 30 d O I - - - -
3NBOOB-BNBOOBIVd 40 d O l . . . .
3N3001V~ U3MO7 dO d O l . . . . . . 3N3001~ 3qQuI~ 9 0 a 0 1 . . . .
~.,,,
I
009~
,.-.
ISUUO OINV3OO IN3~3SV8 UNIAI/UIdVla 3-1VHS
(~
NO,IO=tS =IHI
H I I M I N l O d SSOEIO
,
4
~
#
--~ o s~~~~so~
S70BMIIS
]-]am AaN
o a ± o a r o u a qqBM
-~,,z
'
0,~9
~ ~ I --
~
(]--
A
kZ)
~
o~
0 7
~
17.
£. ~1~' .
/ "
£.
~t q6 9 N I 7
g
BN
V
V
I
I
7
7
9
9
(2
IliA/IN
(2
0
0
eZ 9 N I 7
0
N
Ms[.
I
/:/
0
O
'A IINN
0
0
IliA l l N f l
N
0
I
J Cml cj
INqlAIqSV~ NVlhlglAIVOqldd
tu~t 01.
IN31AI3SV8 NVIhlSIAIVO31do,
IIA
&l
qz 9NI7
I
e9 9NI7
(qz
(ez
}.s
-t,
o.9..
-I
"OlUoJd oql jo pJ!ql uJoqlJOtl oql uo u~oqs SlOlgO]jOl
[email protected][olpp!lAI pug Jo~oq gu!ddeluo oql ql!~ u!seq oql jo aOlOlglgt]o doopoJoj oql soz!seqdtuo 19OStlgll S - N oq~L (q) "ut.~jetu JlOqs OllgtloqJ133 9.~ozosolA[ poumoJp pug popoJo ug oq o1 poloJdJOltlt, st ltlOttldl139so s~.q~L "SlUOtu!pos gJe!IJO~L £q poddeluo s! qa!q~ 'dJeaSOOled poaunouoJd e s[ odols igluot~!luoa Joa¢oI Oql gu!~lJopUFl "sa!lst.JolaeJeqa u[gJetu oa[ssgd Jeoi a s~oqs laosugJ1 A%S-~N oqj~ (e) "tuJojleld oaou!J O oJoqsjjo oql ssoJae slaosueJ1 IeUO!gON "L "~!~I
~-~ I."
O"
o~
"'
~o~C~7 o~L
......J vs,~v~
-'~
N V :7 @ 8 / ht V 0
op VO
V 9 S
N 0 I1
(3VOINIUI
~
~}.~~//~/%
-~.~
AIdVUNNO8 3 0 N B n 0 3 s "=~SS "-'-~
NOllVNIV~U31 NOIIOBqJBu
•~o~.~o~ #
1 3 A 3 l INB~IBqqOO3Q
ONVOIOA
/
0N9997
aNBOOI7d 30 d O l . . . .
3NBOOIqd 90 3SV8 ..........
AHdVlt~911VI:tl S
063
~:~
N 'e" 9 0 0
-..~
~,
OIZNV71M
-o6
oo~
"oLL
d VIAl
LN31/~35V8 -I0 dO£ ...........
Olsswnr
,r'
3ldI/-Inlhl
~
"0! pue 6 "s~!ct u! un~oqs sllnej Ftuaou apls!i aql jo S,~Z!A Z~I!JlS aae (IA l!UCl) auaaO!lA[ zlPP!lAI aq] u!qI!A~ sllne~ z~t!l-dooas zq~L "lJedJmunoa uJa:lsa,~ sl! o] J,el[.tu[.s s[ ]3asueJl EIS-A~N ~aoqsuo uamsea aqj. (q) "sllneJ Ftuaou se pmaJdamu! aq osie pInoa uo!la~s aql jo pua uJ~qlaou aql uo u~oqs sllneJ aSJZAaJ aq&, "aau~nbas m~Jetu zMssed aql jo l!tu!i p.~eA~pUeI lSOtU papoJa aq] s! q3!qA~ a~pzA~ ,Qmumu!p~s snoz3elaa D u.[q] X.[aA e UO pm!sodap sa!lse[a d~zp~aoj ~Ia!ql sn~oqs laasueal EIS-A~N uJalsa~ aqd, (e) "snoaael&~3 aql pmeJlauad suopaas qloq uo UA~OqSsip,v, aql jo lSOlN "u!se~t u e p n z z u % uamse~I zaoqsuo aql ssoa3e sl3~suea] ieuo!~z~i "8 "~!cI
~"
zsnklo OINV3OO J~>~:i:[':':
ONVO7oAanw~lavla a7VHS
~'* "~ ~'~ ................ sno3ovI3EIO-VEIZNI £no3ov13EIO 40 <:101. . . . . .
w~! 9
<°"t:
9,
.........
I
£.
~
-...... : ' ' ~
"
" ~ ~ ! ! ! r ~ : :
..............
.LN=IIAi3SV8 N V I E I S ~ V O 3 E l d
.s
oq31HS VNVAf7D
O"
,~^~,~.~,,oo~o--O--
I#,
zlnV:l 7VI/~EION 19r1v--I 39EIBAEIEI
3N3OO3-3N3OO37Vc140 dO1 . . . .
,8
NOIIVNIIAIE131 NOIlOBq=IBEI
3N=1009170 =10 c101- - - - - - -
AWVONn08 30N317039 ~ ( ~ N011038 3H1
3N3001~ EI3M07 90 dO1 . . . . . .
~N3OOiV~ ~700i~ ~O aO± . . . . EIN3OOl7d 40 dO1 . . . .
~ I I I A
g,
88""
k~
H±IN~±NiOa SSOUO
o~±o~r'oua 77~a
~N~OOna ~o 3sv~ ..........
J.INA ~__
L"
~
773M ,,k3H
i 0
iii
~-6 3NI"I AHd'dldt~l.I.
~1:11S
MN
$70~1/111S
,N
S "7 - 1 1 H . L 0 0 _ - I S V t g V N O ~
ON.~Dg'I
I
3S
S
V
Et
Et
n
s
N
I
IE
n
J.
v
141 Z
(e8
~ ~
9-
\
,9
~ ~ ,
11
o.
".....o~
o~
~
~
o~ ~
o~ .
~
'
\ m8 "o!d
~' ¢" ~.,,,,~:z--.~.2"~_.~'~..~ ~
/
>l~
o~
~
oO/
tY~9
op
d Vi/V
NV~8~IIdVO
op
NOIJ.
o,~ VO
~
.
.
~
"
~
~
~
..........~.,,.
1N:I~EISVB
~
~,o,~'~ o~.
" "~
I '¢ ~
NVIEISIAIVO3EId
~
~
".Z,
........................ : ' " : ' " ~ ' : " ' : ' "
W
0
e6 3NI"I
0-1
MN
N
I
S'
V
8
8
n
S'
N
/
E/
n
.L
V
/4/
as
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN gin of South America. On the eastern margin it is also shown in Fig. 5 that the oceanic Atlantic lithosphere is subducted towards the west underneath the Antillean island arc. Using the nomenclature of Bally and Snelson (1980) the limited northward continental lithospheric A-subduction of South America merges with the westward B-subduction of the oceanic lithosphere of the central Atlantic. The southward B-subduction of the Caribbean Plate underneath the South American Plate is conjugate to the A-subduction of northern Venezuela. Thus, the transpressional Neogene folded belts of eastern Venezuela, Trinidad and the dominantly transtensional northern Caribbean offshore are formed at a triple junction of three subduction zones. Structural studies of the collisional zone associated with this triple subduction junction have been made by numerous authors (e.g. Audemard et al., 1985; Rossi, 1985; Lilliu, 1990; Chevalier, 1993; Roure et al., 1994; Passalacqua et al., 1995; Chevalier and Spano, 1996; Av6 Lallemant, 1997) and several tectonic maps have been published (Bellizia et al., 1978; Case et al., 1984; Speed et al., 1984; Mascle and Letouzey, 1990). Comments on foredeeps
The Eastern Venezuelan Basin is a foredeep (or foreland) basin. The term 'foredeep' or 'foreland basin' as re-defined by Bally and Snelson (1980) corresponds to the class of perisutural basins that are associated with and adjacent to A-subduction zones (i.e. subduction zones where limited amounts of continental lithosphere are subducted). These basins develop on the continental or transitional lithosphere that underlies fold and thrust belts and their adjacent foreland. Subsidence of these flexural basins occurs in response to loading and is controlled by the rheology of the underlying lithosphere and rates of deformation in the adjacent folded belt (e.g. Beaumont, 1981; Flemings and Jordan, 1990; Jordan, 1995; Miall, 1995). Foreland basins remain underfilled if the rate of subsidence exceeds the rate of sediment supply, a situation that is often realized in the early stages of the foredeep development. The foredeep stratigraphy is modulated by eustasy which particularly affects the upper shallowwater portions of these basins. Bally (1989) showed an idealized cross-section that sums up some of the main characteristics of foreland basins. From bottom to top the following unconformity-bounded tectono-stratigraphic units can be differentiated: (1) the top basement unconformity; (2) a syn-rift sequence topped by a breakup unconformity; (3) a tilted passive margin sequence topped by a basal foredeep unconformity; (4) a deep-water ('flysch') phase that corresponds to the inception of the foredeep; (5) a prograding ('transitional') phase corn-
435
posed of delta and prodelta sediments and associated lowstand wedges; (6) a dominantly alluvial-deltaic ('molasse') phase. Three important and structurally controlled unconformity types separate three different structural regimes, i.e. the pre-rift unconformity, the breakup unconformity (both related to the earlier passive margin development) and the all - important basal foredeep unconformity which marks the inception of the foredeep. The basal foredeep unconformity was defined by Bally (1989) as "a complex unconformity underlying the foredeep clastic wedge (phases 4, 5, and 6) and overlying the preceding platform sequence (phase 3)". The basal foredeep unconformity is often characterized by regional updip stratigraphic truncation below and by the progressive onlap of deep-water sequences, and/or condensed downlap surfaces. This downdip portion of the foredeep sequence represents onlap on a flexed and suddenly deepened former platform top. However, in the updip portion erosion of the basal unconformity is due to eustasy-modulated exposure of the peripheral bulge that is associated with flexural basins (e.g. Beaumont, 1981; Tankard, 1985). In its dynamic evolution, the downdip onlap surface and the updip erosional surface are often linked. The many shelfal unconformities that separate the alluvial, deltaic and prodelta sequences (phases 5 and 6) are third-order to fourth-order unconformities that respond mainly to sea-level changes. These unconformities may frequently be enhanced when they intercept growing compressional anticlines. In its simplest form the Neogene foredeep associated with the East Venezuelan A-subduction boundary is best illustrated by Fig. 6a across the Orinoco platform where a Lower Cretaceous to Oligocene passive margin section is separated by a regional basal foredeep unconformity from the overlying Neogene foredeep sequence. An over 7000-m-thick wedge-shaped Neogene section fills the foredeep which unconformably overlies the Cretaceous to Paleogene passive margin platform both onshore and offshore. The Neogene section thickens toward the orogenic belt of Trinidad. In our area the passive margin stratigraphy is a mixed siliciclasticcarbonate margin, and the Neogene foredeep is entirely siliciclastic and characterized by an overall longitudinal progradration from west to the Orinoco Delta. Fig. 6b summarizes the tectonic context of the deposition of the ' sequence-stratigraphic units that will be discussed in the main part of this paper. Regional transects
In the following we will first discuss two offshore transects across the Orinoco platform (Fig. 7) to contrast them with two onshore north-south tran-
i
|
/
ii
_
lOkm
SHELF BREAK /
|
9,
""
'.,
i
\ \
...
H
.
#
60 ~
,~
"
-
i
=~
(a)
(I
~~
,:,0
~'-~
,
9o.
T I C
11~
8o, km
T L A N
OCEAN On
A
59~ _ ~
.
i
~ ' -
~'"---.-2.4
\ ~
LOWER MIOCENE . " . 9
,,,
.
.
, ,SB-
-SB-3 .
9".-;'.;"Z~:):i. !:~~:;~,.:: . . . . . . . . . . .
.
.
.
__
.!.." ......... :.:;:~-:i:::";-:',
" ,
.
,,__ = ~
,~
.
---
,
i.i.'~;:,.,.~.....
.......Gii.~!....:~:.:.:,.
~~.,oo~.~
" 9ii.i.,i..i...:.i~;;-;~. .
.
.
.......
.
.
.
.
.
---- "_~."..,
i~~':~i~~:~~::~~!~':~~:;~:~;~:'~~~-~
w
~
,.,,,
_
-
W
~
\
//
o~,~.
..,... ~
~,
/
ORI~CO
_
61 ~
.w
Fig.
\
"~~
GULF
62 ~
\
".
o
\
~"
/
9
I
/
j
/Fig. 8a
/
EPF
,
,
/ I I
I
~
9i
o
~
Mmou~ MO:=a~e /
I
/
~
64
.
" ' ~ , ~
(b)
.........
i=
E
,._~
t
E I
Fig. 10. The Mapirito listric normal fault system. (a) Time-structure map showing the location and the general trend of the listric normal fault system in the Mapirito area (redrawn from Daza and Prieto, 1990). (b) Schematic evolution of a west-east cross-section across the Mapirito system showing sequence of faulting and the pronounced truncation of these faults during a major Late Miocene erosional event. The listric faults are younger in the west and older in the east, which contrasts with the classical development of sedimentary load-induced growth fault systems.
I
~
--
,41 o
G)
r
t.aO
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN sects (Fig. 8). Finally, a very long east-west transect (Fig. 9) will illustrate the transition from the onshore Maturfn Basin into the offshore Orinoco platform. The offshore Orinoco transects (Figs. 6a and 7) show a pristine passive margin. The Cretaceous on the passive margin is characterized by a substantial thickness (ranging from about 2 km updip to over 4 km downdip). A thin Paleogene section overlies the Cretaceous. To the east a steep submarine erosional scarp marks the margin of the Orinoco platform (Fig. 7a). The deep-water onlap of Neogene sediments onto this scarp is reminiscent of seismic profiles across the Blake plateau (e.g. Dillon et al., 1985, 1988) and the Florida scarp of the Gulf of Mexico (Wu et al., 1990). In Fig. 7a Neogene subsidence appears to be related to the development of the Atlantic passive margin as it is loaded by the thick Neogene sediments of the Orinoco Delta system. However, in Figs. 6a and 7b the same Neogene infill is wedge-shaped, expanding to the north and onlapping towards the south onto the basal foredeep unconformity in the offshore southern Trinidad area. Two onshore transects in Fig. 8 illustrate the tectonic setting of the onshore Maturfn foredeep. The northern end of both transects corresponds to the leading edge of the foothills of the Serranfa del Interior. The wedge-shaped Neogene directly overlies and onlaps a truncated Cretaceous section that is more deeply eroded and thinner than the Cretaceous of the Orinoco platform. In Fig. 8 the structure of the southern flank of the basin is characterized mostly by southward dipping normal faults often associated with some important oil-bearing structures of the area. Typically, these normal faults trend east-west and southwest-northeast (N60 ~ to 70~ and they involve the basement, the overlying Cretaceous and the Neogene sediments. On seismic dip profiles this fault system displays subvertical throws frequently less than 50 m (Audemard et al., 1985). Farther north in Fig. 8 we show north-dipping reverse faults. The dip of these faults is poorly defined and a normal fault interpretation may be equally acceptable. The margin of the folded belt is shown on the north end in Fig. 8a and b. Note that the structural growth of the southern edge of the folded belt is shown by the updip convergence of Plio-Pleistocene strata on the southern flank of the first major anticline. Thus, very young deformation marks the outer margin of the folded belt. Fig. 9 is a line drawing of a 800 km longitudinal set of seismic profiles corresponds to the onshore Maturfn Basin. A gap of about 70 km separates this profile from its offshore continuation. This profile ties the offshore Orinoco platform profiles to the Maturfn Basin profiles.The relatively thin Cretaceous section onshore and the poorly defined pinchout of the Oligocene contrast with the thick Cretaceous
437
shown in Fig. 7a. The overlying Miocene section progrades from west to east. Also, from west to east several structural features can be seen as follows (1) A major late Neogene reverse fault is related to the Anaco inversion structure trend which consists of a series of elongated northeast-trending and southeast-verging domes. Funkhouser et al. (1948) and Murany (1972) suggested that the dip of the principal fault plane diminishes with depth. The structural history of the Anaco trend was interpreted by Banks and Driver (1957), who considered its origin as a major northwest-dipping syn-sedimentary normal fault that began to form at least in Early Oligocene time. According to Villaroel (1993) this growth fault ceased to be extensional by the end of the Middle Miocene and inverted into a compressional reverse fault during the Plio-Pleistocene. (2) Minor half-grabens of possible Jurassic age. (3) Basement-involved reverse faults that offset the Lower Cretaceous to Oligocene section. Note that these faults could also be interpreted as normal faults as already mentioned for Fig. 8. (4) The listric normal faults shown on the onshore transect (Fig. 9a) are referred to as the Maparito faults (Lilliu, 1990; Daza and Prieto, 1990). On E - W seismic profiles, the basal detachment of this listric normal fault system steps down from 3.5 s in the west to 5.0 s in the east and involves mostly Lower Miocene sediments (Fig. 9a). The listric faults trend northwest-southeast and dip toward the northeast. The stratigraphic relationship of sequence boundaries associated with this extensional system suggests a backstepping listric normal fault system with no perceptible growth. Note the pronounced truncation of the beds. Fig. 10 shows the location, structure and evolution of this system. The key point here is that the faults are younger in the west and older in the east. Thus, in contrast to typical growth fault systems the oldest faults are found basinward. This type of structure is generated by gravitational collapse or slides on a gentle submarine slope that is gravitationally unstable. The listric faults were formed during the latest Middle Miocene to Late Miocene. Note also that the extensional system overlies a complex system of apparently compressional reverse faults and folds. The faults are difficult to understand. It may be reasonably inferred that the deeper reverse faults originally were normal basement-involved normal faults that were subsequently compressed either during the Middle Miocene or perhaps later. The offshore section (Fig. 9b) shows from left to right: (1) The west to east progradation of the Pliocene depocenters. (2) A set of pronounced growth-faults and possibly related toe thrusts of the Columbus channel
438 area (Fig. 11). In the eastern Venezuela offshore the Neogene depocenter overlies the Cretaceous passive margin platform and its thin Paleogene cover. Note the syn-sedimentary rotation of the Neogene beds into the fault planes (Fig. 1 l a). We suggest that the growth faults may be synchronous with at least one compressional fold at the toe of the deltaic system developed on the slope (Fig. 1 lb). However, lacking a tighter grid of seismic data to provide a measure of three-dimensional control, it is difficult to differentiate toe thrusts with a strike that parallels the growth faults from the more complex folds of the Barbados accretionary wedge. (3) At the eastern end is an oblique cross section (Fig. 9b) across the north-south-trending Barbados accretionary wedge (e.g. Speed et al., 1984; Bouysse and Westercamp, 1989). These folds can be easily mapped as part of that accretionary wedge, but seen on a single section they falsely mimic common toe thrusts associated with growth fault systems. The updip convergent reflectors of these folds indicate continuous growth which occasionally is intercepted by deep-water erosional events, that form unconformities.
J. DI CROCE et al. STRATIGRAPHIC OVERVIEW
Our interpretation of the sequence stratigraphy of eastern Venezuela is based mostly on well and seismic correlations and will be organized on the base of sequences as defined by well and seismic data. Fig. 6b shows the three major tectono-stratigraphic regimes that control the stratigraphy of the Eastern Venezuelan Basin, i.e. a poorly defined rift regime, a passive margin and a foredeep sequence regime. A systematic study of regional seismic profiles correlated with well logs permits to identify and date numerous sequence boundaries for a chronostratigraphic framework. These are also summarized in Fig. 6b. A top basement unconformity (SB-1) commonly underlies the mixed siliciclastic-carbonate passive margin megasequence. Only in the presence of rather vaguely defined Paleozoic and/or Jurassic reflectors an intermediate unconformity (SB-2) separates the older earlier sediments from the overlying Cretaceous. The basal foredeep unconformity (SB-3) separates a foredeep siliciclastic from the underlying mixed siliciclastic-carbonate passive margin package (Fig. 6b). A line drawing (Fig. 6a) of a profile to
Fig. 11. Schematic illustration of growth-fault related toe-thrusts and folds and thrust related to the Barbados accretionary wedge. (a) SW-NE close-up of growth faults. For location see Fig. 3b. (b) Schematic model of the offshore Orinoco platform-Barbados accretionary complex transect. See Fig. 3c for location. Note that the toe-thrust shown in the center of the diagram is not related and probably strikes at a right angle to the strike of the Barbados accretionary wedge, folds here shown on a very oblique profile. The question mark indicates in a strike view the area that separates the pristine passive margin from the area underlain by the subducting Atlantic slab.
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN
(1977) and summarized by Gonzalez de Juana et al. (1980). Feo-Codecido et al. (1984) emphasize that the extent of the Precambrian basement to the north of the Maturfn Basin is not known and that one cannot determine whether undeformed Paleozoic and Paleozoic metamorphic belts shown in Fig. 12 as underlying much of the Gufirico sub-basin to west would project underneath the Serranfa del Interior. Offshore, to the east, the Precambrian basement extends eastward underneath the Orinoco platform and the Atlantic passive margin, where it eventually is involved in the transition to Atlantic oceanic crust (Fig. 7a). Onshore the top of the basement is observed dipping to the north. Underneath the frontal folds of the Serranfa del Interior (Fig. 13) the downdip continuation of the basement is poorly defined on seismic profiles and mostly deeper than the 6 s (twt) penetration provided by most seismic profiles (Lilliu, 1990; Chevalier, 1993; Passalacqua et al., 1995). Fig. 13a is a time-structure map of the top basement and Fig. 13b is a depth-converted basement map that encompasses a much larger area, including the Guyana offshore. The general trend is characterized
the south of Trinidad that displays the significance of these two major basin-forming sequence boundaries has been discussed earlier.
THE BASEMENT On most seismic profiles the 'top of the basement' is the deepest 'strong' reflector which is characterized by its high amplitude. The top of basement is the sequence boundary SB-1. Over much of the area, Cretaceous sediments onlap on this surface and record the beginning of the Mesozoic encroachment cycle of this area. More than 50 wells in the area confirm that the basement is the crystalline Precambrian of the Guyana Shield, composed mainly of meta-sedimentary and meta-igneous rocks in an amphibolite to granulite facies intruded by granite intrusions (Feo-Codecido et al., 1984; Fig. 12). The reported age for these crystalline rocks ranges from 3600 Ma to 800 Ma. Complete descriptions and discussions of these Precambrian rocks have been provided by Martfn-Bellizia (1974), Mendoza
N
660
64~Margarita [ . 63~
65~
C A R I B B E A N
439
6~~ TobagoI. ll~
]
S E A
EPF II' II~ %,,,~
9 . "
%
9
~ '-'-'0"... . . . 9* ' % 9 1 4 9 1 4 9 1 4 9
9 9
9
%
9
9
%
ar
9
9
ar
9
9
~
J
9
%%
9
S
9 -" 9 "
9
"
-" "
9
9
9
.,
Stratioraohic C o l l ) m n CTertiary I ~ (2800 m) retaceous~ Eguivale_nt ]~asalticFlow ~ La QuintaFm. (162 Ma) (1645 m) Carboniferous " (740 m) Sandstones ~
9
9
9
9 ar
9
"
%%',,,,
9
%%% J
9
9
9
%
%
9
9
% .."
9
."
9
%% J
9
..*
%
9
,."
9
.#
J J i," J J ~" .*" 9 1 4 9 91 49 9 9 % 9 9 .~ 9 #, .s l 9
l 9
19
DELTA
%
9
#" 9
9
s
.,1'
9*'
-
9 %"
-
"
",
/
~"
SHIELD
8o
WELL
PzMETAMORPHICFy---[ THRUSTFAULT FOLDEDBELT UNDEFoRMED PALEOZO CI ~
NORMALT A U L F
ESPINO GRABEN
~
OUTERDEFOR MATIONFRONT-
~
PALEOZOIC THRUST SYSTEM
~PRECAMBRIAN BASEMENT
%
ORINOCO
9 9 9
9
Precambdan % Basement %
%
9
LEGEND
[~~
%
9
aP
CarrizalFm. ~iiii!iiiii!i'ii!~i!i!!ii!! (1,830 m) %
%
9 9 9 1 9 9 9 9 9 / 9 9 9 9 9, % ~ I I V / /V KI& -~II/-~I
..., 9 9 9 9 9 9149
'
%
.9 J J J .# .*" 9 .~ J ,*" 9 9 9 9 9 9 9 9 9 9 J S J J #t J S ~ .~ ~ 9 % 9 9 9 9 9 9 9 9 9 ,,. 9 9 .r .# ,s 9 9
9
9
9
9 j ,9 j 9 #" . "lb% 9 % 9 "*. 9 j 1 p , f,=- j .~ j .., 9 % .=,~..~-.. % % 9 9
0
100 km I
Fig. 12. Generalized pre-Mesozoic subcrop map including a stratigraphic column illustrating the presence of Paleozoic and Jurassic rocks in the Anaco trend. Redrawn from Feo-Codecido et al., 1984. Jurassic half-grabens are not differentiated in any detail within the study area because their distribution is limited and poorly defined on seismic reflection profiles.
440
J. D I C R O C E et al.
62~
63 ~
T O P OF B A S E M E N T
SEA
CARIBBEAN
9
3.0 =' 2.5 2.0 ~.0
o..~
/
'-"'-.',
....- ,_~,;; ....- '..",'-i-:' ,... ', -.- . '~ -. - - 6= ~ ', ,. - - '. . . . .
~ ~~ . : !~ ~ ~ : : ]~1 ~v ~
I
I b~,,I
r
'..'-. -,'~ -.-
Outer margin of folded belt
I~1
. . . . . . /nrus]: raull:
I',',?,'J , ,,,-
64 U
M.,o~
"," ' - " " ' '~ "-" " ' - " ' - ' ",'" ".-: ' - " " ' '" "-" "'-" ".', ',"
I >'~
62 o
~ .. -- ,. , -."."" 6oo - . . . . . 1 o , , -.- - ___
Stdke-slin fault I
51
" Exposure 9
Precambrian
basement
8 9. . . . . . . . .
Contour oceanic
/'SEA
58 ~
\
,~, ",
Contour crystalline
~
60 ~
CARIBBEAN
.s
sso
\
\
58o
Contour units sec ( T W T )
} I
Contour interval 0.5 sec ( T W T )
I
56 ~
BARBADOS
T O P OF B A S E M E N T
ACCRETIONARY COMPLEX
TRINIDAD
4 r ,q4,be
IIDSDP-1441
OOE4,, v ,., .,...,,..,v %.,.,"-#%/%g% ,% ,%/%g%.,.~%g"*/%g%/%r .
....,,
,.-".,,",,
.
.;.;.;.;.;.:.:.:.i.;.;.;.;.;.;.;.. .,,,,, ",'.:~.:,.,.,,..% .% ,:,',,.,.,,,..,.. ,.GUYANA.;,;
VENEZUELA
b)
0 I
50 I Km
:.-.-. 100 I
"-,',. ] .~-.',.',,',,',',.;.,'..
~
. . . .
SURINAM',',,',','/.,;,'
""~'n-.
". ' ~ 4
Fig. 13. Structure maps on top of basement. (a) Time-structure map of the top of basement derived from regional seismic profiles. For abbreviations see caption of Fig. 3. (b) Depth to the top basement of Eastern Venezuelan Basin, extended to include the basement of the Guyana Basin (based on Petroconsultants, 1989). The Soledad-1 well (Point 'S') is the deepest well to reach basement in the study area.
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
_ Z ~ -~=
..~
~.
-
~
...... _
-
" z ' ~
~UNIT
_
.
~
.
.
.
_
+_+ . . . . . . . . .
~
:
i
_ ' _
-!
~
: PIlO-Pleistoce~
:+'-,"+,,
-~
=
+ - - ' ~ ~ . i-..
,--, ,,~",,--,'-".,-, '~ ~-. ,;y~,,, ~ ~ ~ ~ ~ - ~ , ~ ,-,,,,-;'~,,-,y! . , . . . . . ~....... :~ ....
U N I T I I ~j_
VIii ~
.
~
.~I:.,i"A"~ WELL_
,
SCALE
92(X)O
Tinaj#as Mb.
912000
4200.
Marg=~ I.
911"~,(~
~
L~ J A. Well Offshore Orinoco Delta JJ. Composite Stratigraphic J Column Interior Range
/
0
~
6'1~
100 Km
I
6()0 '"'
Barremian
Aptian
Albian
__Inial
Santonian
MAP CA~ISaEAN SEA
6'3~
6'2~
, , ~
LOCATION
Barranquin Fm.
Garcia Mb
El Cant# Fm.
Chimana Fm.
Querecual Fm.
l,." I vEN~zuE~
.14000
3900. .13000
3600.
3300- .11000
3000- ,100(X)
2700- ,9OOO
2400- 84
2100- 7000
1800- 6000
1500. 5000
.
Maestrichtian
~ampani~n
V~aleeeeeeee~c~e San Juan Fm.
. ~ ~ ~Fm.~ ~ w e r
,?,
Carapita Fro. + i ++!i++++++ i +?i+!:++ i+ i ++++ i +++' +::+,i Naricual Fm. .;.-_-_-.T-_ ,, Areo Fm. ~i::!!:i.:!::!:;..:.:~ Oligocene Los Jabillos Fm. i:..-:..:.-:..:.; I L ' , , . ' , , _ - , ,_+r " - " I
1200. 94000 San Antonio Fm.
900, .3000
600,
300. .1000
(m)o (feet)
. ,
s
LEGEND
UNIT II (CRETACEOUS)
UNIT III ( P A L E O C E N E - E O C E N E )
UNIT IV O L I G O C E N E
SB-3= 25.5 Ma SB-2.2= 36 Ma SB-2.1= 58.5 Ma
mfs= 111 Ma
mfs= 98.25 Ma
mfs= 95.75 Ma
mfs= 91.5 Ma
Limestone
Shale
Sandstone
Aptian
Albian
Datum: Base Miocene
UNIT II
J
UNIT
Fig. 14. Correlations of the Orinoco platform stratigraphy with the stratigraphy of the Serranfa del Interior. (a) Schematic stratigraphic column of offshore wells A and B (tied to a line drawing of Fig. 6a) showing the calibration used in this study. See Fig. 3c for location. (b) Proposed correlation of passive margin stratigraphy of offshore well 'A' with a simplified composite stratigraphic column of the Serranfa del Interior (Chevalier, 1993).
"++"=
~
--+,~:
(a)_ ___+~+~WELL"B"
(b)
0
b~
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
J. DI CROCE et al. 64 ~
9~ 62
63 ~
Margarita.l.,M~
C A R I B B E A N
e-~o S E A
_~\~
eo~
s9o
Tobagol.
Cumana
J
,8.0 1
~.
'~
J .
/ . - ~ ~ ; , , ~ . . . - . ' : ~ .
-., : ' : -; -,-~: :.:. : "64: -~: :, ;:.:. .' ,: - - .'_ 9
"
"
II
'
"
-
I
'
"
-
"I
; ,:
: : : - ~ o -':' -
" ' ""'"
.... : ' I 0
'. -'"
40
80
120 K m r~. -. ..... _-: , ~.
. .... -. .. '. . ' - , ,-_. ~ ': " 62 . . '~ '~ .. .- . .61
I
I
"
II
I
I
:.::,--, ':
'"
6o o
%
8~
':
58o
59 ~
I
o
,9"
9
!
.3
l~r162 I.
.10
C A R IB B E A N
S E A
C.ma,,a 9
~
9
Tob~o J
TRINIDAD
~
i
) La Crm
~
2..
_ ~ o o
i
F ~
Outer margin of folded belt
I ~
i j~si
Thrust fault
~
Strike-slip fault
~ 21
9
.o
.
so
Contour
l,oo~,.
I
Exposure Precambdan
Contour units sec (TWT)
crystalline
Contour
basement
.
interval 0.5 sec (TWT)
I
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)
O0
I
m 0
C~
I
r.~
I.~:.~
.~~
~-~
I ~-~
E'~
~
~1
~_
~
ft.
r2 .~
I
~-~
c~
rO rO
I
I
~~-~ ~.~
.~
~ ~~
-
I
.~
I
C
~_~
0
.~
~!
~1
~ ~ I
~: .~.-_ ~ o
I I~~
I N ~
~
I
I
~'~-~
I
~ .~
I
.~-~
~ o~ 1
I
A
H
V
I
1
~1
3
..
©
0 ,=.-
__
0
~ eq
t~ 0 ;~
I
ua
o.~
O
"~,
Zl
auaao~,s!ald-O!ld
~o
~_
~
=
~~
0
- g
~:¢ .~o,
Z
~y ©
:12 o :~7aa,~, kFIddN IT" >Z o ~ u31no 0
q~ ~
:. ! ::~.::i!i:i~!!.17:!!!i i !.!:::/:i:~:/:::.!:.!
i~ ~
i:??<:::i:!::;!ii:~i i:: ::i~i:S:i!!::i:!!:!:!ii!:ii:i:!:iiii::ii:!~i~ :i :i~:i?!ii2:ii!'{:"!~i~~{:ii":?ii2!:i:!}ff!!i2:~i!:~ :il;::.F!i:i~;%.i?ii?~:E::iS.:::i;i!~:~;!L ::::i:;::!:i~i:i:!!??!.i!!:i:/:i:5:i?!s:i!i. i:i:~i?i/:?~i i::.:.il!~i~2::i71}i?
0
~-~
#
"~
,.C
J
IN~NI.LNOO
,x::
'696L'MO,O
°
NOIIVNOZ
z z
o~
(~pn~s s!ql) XUVaNnO8 3ON3n03s
eueoo~s!eld
NOIIVIAIWO:I
0 .,.-,
VS31AI
I
, ~
eueoO!ld
~"
~::~!: :~!!:ii!ii!::!ii!i~.i!i~!!!iii!i~,i:i:i{}.-~iiii~:i.~:~:.~ i:::ii~:ii:~.
Z
:~ :i co ~ ~~~ i!i :! : {::!!ii~
co,
~~ ~
m O9
I
auaoo!lAI
auaoo!lAI
Jaddlq
SVWO31d SV9
z
elPP!IAI-JeMo7
ZZ"
.... :!:.: :!~!~i!~:!:i~;:~:!<:~!~!~!;:i}~::i!~! ! :! i!~! :!~i:!:,::!;i ~i<@~: i i !!!}!i~y:~! !i:~i~!~:!
gb~I:!:~f!!::~i~:Eiitl~ii i i i!iljii i [~I1
eueooS!lO ~addn
VIIdV~VO
VOid V9
a2~
!i!!ili!!:~:,~!!!ii i!:~!}:!!2!2 ~
0
-~.~_
~
0
~'~
0 ,..o
~~
3wno3w3~N
~g
i !i
: ~:!iI i::!!,!!~, !i!!ii!i ,~!i~ii!i:!i!i~i:i!~!ii: :,ili:=ii!,i~i i:!i~~!ii~ i!i!i!! ~:ii!!ii ~i'~i!il~ii::iil:~:~!i~ '~ i'~:~,~¢~:i!!i~:i~~: :i! ::~:::!~ii~i~:~i!i!~ii~
=~
..~
iiii~I!iii!iliii!iiiiiii!iii!!ii!ili!il i}i!~!iiiiiiiiii!i!!~i!iii~:iii!i !ii ili ii!il,,~i !ii!!!ii i i!i!i~~:~i~ii!iii i~
°:tq,,
:":'~' ! i iii iill !ii~i;~i!ii!ii~i'!i!ii i!!ii!!iiii~?iii i4iiiii~i~i i '~!!i~::~!t! i i ~i i !i i!~,~!)i ~!i;i!iiiil ,~;ili ii~/: ! i~i
©
m
©
=~
~
>
u ~ ~
t,..
"13 -..,.., 0 ,-~
u~
0 ?~-
--
,,
__
~ 3N30019d d e a p e J o :1
~-
> 3N3OOIIAI
.,.~
W3ddn m
~
,.o
to3
>m .~_
!
~
,©
©
©
Z~ Z
~'~ ~
'~
euaoO!ld
euaoo!lAI Jaddn
A
~1
V
aueoo!lAI
I
1
El
elPP!IAI
3
1
.
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,
456
J. DI CROCE et al.
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
457
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.
0
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN
459
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
460
J. DI CROCE et al.
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,
461
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
462
J. DI CROCE et al.
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'.
463
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.
464
J. DI CROCE et al.
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
w a~!
0
R
I
N
0
C
LINE 6a
0
D
465
E
T
A
LINE 7b
0
UNIT VII -! -i
1_.3"
~K4~
t~
4-
l
lOKm
I
~
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
CHRONOSTR~TIGRAPHY SYSTEM '
SERIES STAGES(1) (2) ! (3) I
QUATERNARY
9
'
~ PLBSTOCEm
9 ,,, (~
"O.Ol1
,
3.4 - -
3.4--
5.3----5.2--
' "~.'!~-_
U
(S) l (s)
.z-
-3.4
-.
3.8--3.5-
~
o.
LLI Z
I
5.1 -
- 5.3-- 5.3--5.2-
..~_
o.s
BOUNDARY
- ~.~-
TORTOMAN
-10.4-
lZl
---11.2--
-
,
*
_L . . . . . . .
SB-3.1 0 "s." - ' "
, ~ . , ,., ~ ~u-:J u."
i -113"
_14.2_ ,14.4.::
>-
-
.-"-X
-16.0
-16.3-
-
-16.1
-
,
i - 20.5-
~.,
IMW_
"T.-.
rm,laa,
........
(2.2.5P-.- - 22 . . . . H .
-23.5- "23.8-25.2-
c''" Iii -
TR J
i2050i
k
23.7__.'23.3-
0 o
r--"
.......... ,,,,_. ,,.,_z~.~
,..y
-24.6-
o
I-2-- L
-16a?,-
--21.8--21.5AQUITANLAN
!11
. . . .
i:~.-
:
.I0.2.
mm20.0-
....
,
-14.5 - -14.8 -
L
LLI I._ + III Z
:;_
~- . .%. . . . ~
SB-3.2'-"~._
<:C p-
,"
-152-
- - 1 6 . 6 - -
BU~.JAN
m
AG[
;-11-0--11.2-
---15.1-I.ANOHIAN
-i,
I
-~.~-
-u-
LLJ
o
SB-3.SB-3"' 514 . . . . . . . . . .: . - ' , : ~ ' ~ ~ ~ : ~ ~ ~ SB-3.13 "-.~-~ ~: ~ ~ .-'" ~ ~,'-- ""~s" S B 3 1 2 9 9, ~ - - - : = +,~_ ~ ! ~
SB-3.8-"""
LU
I SEQUENCE
BOUNDARYI
'
.6 -"-1.64-' 1 9 "Pu~,sTo"1.65 1 T/-t.85"
1--
El
(4)
, 0 . O l , ~O.Ol, =,,,+ ~'
SEQUENCE S(QU~NC~ S T R A T I G R A P H Y
29.3- -29.0-
i
:,,.-:,,,
-28.5-
rao.o-
i
,
i
i
i -33.0 -
o
L.U
.
i,,I.
. . . . . . . . . - 35.4
_ _ + : - -_ _
-
---~8.6---
~.0~ Fig. 26. Variations on the Neogene and Oligocene chronostratigraphy according to different authors. On the far right is the inferred position of our sequence boundaries on the Haq et al. (1987) cycle chart. Column (6) shows the ages assigned to various stages by Haq et al. (1987). Column (1) shows ages proposed by the Geological Society of America (DNAG 1989). Column (2) the same after Harland et al. (1990). Column (3) the same after Bolli et al. (1991). Column (4) the same after Vail and Wornardt (1990). Column (5) the same after Hardenbol et al. (in press). plain sandstone and shale (upper Merecure Formation) merge with shallow-water platformal shale and fine- to coarse-grained sandstone littoral bar facies (lower Oficina Formation). To the north, i.e. closer to the present deformation front and within the folded belt, deep-water shales with thin layers of turbidites (lower Carapita Formation) onlap on the basal foredeep unconformity. To the east and southeast limited terrigenous input to the shelf favored the development of a thin carbonate rim at the edge of the platform. Beyond the shelf-break, deposition occurs in upper bathyal water depths of muddy slope fans. The age of the basal foredeep unconformity which forms the base of the Lower Miocene suggests the presence of a structural load in the area of the present-day northern Venezuela offshore. The location and dimensions of that load remain largely hypothetical and are not shown in Fig. 26c.
The Serravallian paleogeographic map of Fig. 27c shows that the sedimentation was strongly controlled by subsidence of the eastward shifting foredeep. As a result the shelf margin, with its coastal facies, moved landward, i.e. toward the Guyana Shield. This period includes maximum flooding in the foredeep domain. The basin shows a narrow trough geometry to the east. To the west sediments are deposited in a broad coastal plain to littoral zone, while in the central part and towards the east deep-water sediments are deposited in an open-sea environment. The symbol coveting the Serranfa del Interior suggests that an accretionary wedge composed mostly of Lower Miocene imbricates was emplaced in that area by Serravallian time. Fig. 27d is a paleogeographic map of the Lower Pliocene just after the SB-3.10 (5.5 Ma) erosional event. The Serranfa and the Monagas foothills are close to their present-day position near the edge of
Fig. 15c.
Fig. 27. Maps showing the paleogeographic evolution of the Eastern Venezuelan basin. (a) Paleogeographic map of the Upper Oligocene just after 30 Ma. This map marks the cessation of the passive margin regime. The present-day area of the Serranfa del Interior is shown as an undeformed extension of the Atlantic passive margin. (b) Paleogeographic map of the Lower Miocene, after inception of the foredeep (25 Ma). The area of the present-day Serranfa del Interior remains undeformed. Presumably the load that caused the initiation of the foredeep was positioned much farther northwest somewhere on the margin of the present-day Caribbean shelf. Compare this figure with the isopach in time shown in Fig. 15b. (c) Paleogeographic map of the Middle Miocene (15 Ma) which shows that much of the Serranfa del Interior is now emplaced while southern Trinidad is not yet deformed. (d) Paleogeographic map of the Early Pliocene just after the SB-3.10 (5.5 Ma). The Serranfa del Interior and Trinidad are now mostly in place. The delta front of eastern Venezuela has migrated from west to east. Compare this figure with the time isopach shown in
Fig. 28. Eastern Venezuelan Basin chronostratigraphic charts. (a) A W-E-oriented chart that sums up the onshore to offshore chronostratigraphy of the longitudinal section shown in Fig. 9. The passive margin Cretaceous section is thin to the west due to onlap below and truncation at the top. Also onshore a very large hiatus is associated with the basal foredeep unconformity. Farther east this hiatus is replaced by condensed sequences of the Paleogene. In the west the Miocene is represented by relatively thick Miocene aggrading and prograding sequences which shale out to become a more condensed section towards the east. The dashed and dotted lines in the second-order cycle column are the sequence boundaries and the maximum flooding surfaces used to separate individual second-order cycles. (b) A SSE-NNW-oriented chronostratigraphic chart that displays the main sequence-stratigraphic units in a section across the onshore tbredeep. (c) A SSE-NNW-oriented chronostratigraphic chart that displays the main sequence-stratigraphic units in a section across the offshore foredeep.
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN the folded belt. The east Venezuela foredeep and it offshore extension are now almost completely developed.
DISCUSSION OF THE CHRONOSTRATIGRAPHIC MODEL Introduction
Our conclusions are summarized in the form of several chronostratigraphic charts that correspond to key transects across the area. Chronostratigraphic charts link the spatial distribution of depositional sequences to geological time. Thus, chronostratigraphic charts display genetically related strata that were deposited during a given time period. They summarize a wide range of information such as: (a) relations of sequences to bounding unconformities, showing onlap, downlap, toplap and truncation; (b) correlation of sequences to standard geochronologic subdivisions; (c) hiatal gaps along unconformities; (d) distribution of facies and environment; and (e) the relation and correlation of named lithostratigraphic units such as groups and formations within the sequences even though such units may be timetransgressive and do not show as a distinct unit on the chronostratigraphic chart. Except for the symbols used for condensed sequences, the legend for the chronostratigraphic charts is mostly self-explanatory. Ideally, a condensed sequence no matter how thin ought to be represented by a symbol that fills the whole time represented by the condensed section. Such a procedure is not altogether desirable because large time segments of a chronostratigraphic chart are often dominated by rather thin condensed sequences. In the eastern Venezuela offshore the problem is further compounded by the observation that much of the downdip Oligocene appears to be missing, suggesting widespread starved sedimentation that is not easy to explain. On the chronostratigraphic charts of this paper, condensed sequences are represented by widely spaced dashed lines with the understanding that particularly for much of the Paleogene it remains questionable whether we deal with very condensed sedimentation or else with a form of sediment starvation. First-order cycle
Overall, the Cretaceous-Cenozoic of the Eastern Venezuelan Basin was deposited during a long-term transgressive-regressive cycle related to the opening of the northern Atlantic Ocean that began in the Middle to Late Jurassic. The cycle reached its peak in the Turonian (Vail et al., 1977; Hallam et al., 1985; Haq et al., 1987) and ended with a post-Turonian overall
471
regressive phase that is still active today. Such a long-term global transgression-regression is often referred to as a major 'Continental Encroachment Cycle'. Onshore the transgressive encroachment of the Lower Cretaceous is well displayed as a thin wedge of Cretaceous clastics (Figs. 9 and 28a). This sedimentary package was deposited during landward displacement of the shoreline with a retrogradational geometry corresponding to a 'Backstepping Transgressive Phase' (Cramez and Vail, 1990). On all chronostratigraphic charts the Turonian K4 surface corresponds to the maximum flooding event. Thus, the K4 surface separates the first-order overall transgressive phase from the overlying first-order regressive phase and is marked by a major downlap surface corresponding to the 91.5-Ma (middle Turonian) event of Haq et al. (1987). The overlying Upper Cretaceous and Cenozoic package is characterized by a progressive seaward displacement of the shoreline, with progradational geometries that correspond to the 'Forestepping Regressive Phase' (Cramez and Vail, 1990). On the southern passive margin of Venezuela sediments included in this overall regressive unit reach a maximum thickness exceeding 5-6 km on the outer shelf and thinning again towards the Atlantic Ocean. The first-order passive margin encroachment cycle is strongly modified within the Eastern Venezuela foredeep and its offshore extension as shown on all chronostratigraphic charts (Fig. 28). On these charts and the corresponding seismic sections, it can be seen that the passive margin regime terminates with the deposition of thin Paleogene cycles characterized by narrow and thin siliciclastic northeastward-prograding wedges and extensive condensed sequences and/or starved sedimentation. These cycles are truncated updip by the basal foredeep unconformity. The main characteristic of the basal foredeep unconformity is the widespread deep-water onlap of Lower Miocene shales on the condensed or starved passive margin of the Paleogene as best illustrated by Fig. 28b and c. It is concluded that the Continental Encroachment Cycle of the Eastern Venezuelan Basin is strongly modified by a superimposed Neogene foredeep phase that began with the Early Miocene and radically changed the geometry and stratigraphic expression of the subsequent higher-order cycles. This change is directly related to the eastward displacement of the transpressive orogen due to the oblique collision of the Caribbean Plate with the South American Plate. Second-order and third-order cycles
A complex series of second-order cycles of different duration and magnitude is superposed on the
472 previously described first-order cycle. These major transgressive-regressive (T-R) facies cycles (second-order cycles) are the most important cycles controlling distribution of facies and are the stratigraphic signature for combined eustatic and tectonic second-order cycles. Second-order tectono-eustatic cycles typically cover a wide time span ranging from 3 m.y. to 50 m.y. Third-order cycles range from 0.5 m.y. to 3 m.y. and are superposed on these second order cycles. At least ten second-order cycles control the facies distribution of the study area. Seven of these cycles occurred during the passive margin phase and three cycles correspond to the foredeep phase of eastern Venezuela (Fig. 28a and b). The Cretaceous to Lower Paleocene (Unit II) is bounded below by SB- 1/SB-2, the pre-Cretaceous unconformity with its lower Cretaceous onlap. The Cretaceous to Paleocene is subdivided into five second-order T/R packages that are best defined by maximum flooding (i.e. downlap surfaces) in the lower Albian (111 Ma), the uppermost Albian (98.5 Ma), the middle Cenomanian (98.75 Ma), and the middle Turonian (91.5 Ma). Sequence boundary SB-2.1, the Late Paleocene (58.5 Ma) unconformity and its conformable continuation, forms the upper boundary of the uppermost second-order cycle. The sequences underlying the Late Paleocene SB-2.1 (58.5 Ma) unconformity are defined in terms of T/R cycles because seismic permits only to identify downlap surfaces which were tentatively correlated with the base of major transgressive shale packages recognized in the stratigraphy of the Serranfa del Interior. Overlying SB-2.1, most second- and thirdorder cycles are defined by unconformities which are best recognizable on the seismic profiles. Over much of the offshore the overlying Late PaleoceneEocene second-order cycle (Unit III) includes dominantly pelagic units and is bounded by SB-2.1 (58.5 Ma) and SB-2.2 (36 Ma). Note that SB-2.1 corresponds in places to a significant hiatus with omission of the Lower Paleocene and the Maastrichtian. A maximum flooding surface perhaps corresponding to the 56.5-Ma maximum flooding event of Haq et al. (1987) permits to tentatively split this secondorder cycle into two third-order T - R cycles. The Oligocene second-order cycle (Unit IV) has only limited distribution in the area and is subdivided into two third-order cycles bounded, respectively, by SB-2.2 (36 Ma), SB-2.3 (30 Ma) and SB-3 (25.5 Ma). The inception of the Neogene foredeep due to the initial emplacement of the transpressional folded belt of Venezuela and Trinidad is heralded by the basal foredeep unconformity SB-3 (25.5 Ma) which modifies the Cretaceous-Present first-order cycle over much of the Eastern Venezuelan Basin. The Lower Miocene (Unit V) as a whole represents another sec-
J. DI CROCE et al. ond-order cycle bounded by SB-3 (25.5 Ma) at its base and SB-3.3 (16.5 Ma) at the top. Based mostly on offshore data this second-order cycle is subdivided into three units, i.e. the SB-3-SB-3.1 (25.521 Ma), the SB-3.1-SB-3.2 (21-17.5 Ma) and the SB-3.2-SB-3.3 (17.5-16.5 Ma) third-order cycles. The Middle Miocene (Unit VI) as a whole may be viewed as a short second-order cycle defined by SB-3.3 (16.5 Ma) at its base and SB-3.7 (10.5 Ma) at its top. It may be subdivided into four third-order cycles, i.e. SB-3.3-SB-3.4 (16.5-15.5 Ma), SB-3.4SB-3.5 (15.5-13.8 Ma), SB-3.5-SB-3.6 (13.8-12.5 Ma) and SB-3.6-SB-3.7 (12.5-10.5 Ma). Note that the base of this second-order cycle is directly overlain by the 16-Ma maximum flooding event. Another major flooding event is represented onshore by the Serravallian 13.4-Ma flooding event which, however, is not recognized in offshore eastern Venezuela. The Upper Miocene to Recent (Units VII and VIII) represents the youngest second-order cycle. Its base is bounded by SB-3.7 (10.5 Ma), its top is the land surface and the seaftoor. This unit is an overall regressive prograding cycle. Using overall regression as a theme it could perhaps be argued that the base of this second-order cycle should be placed earlier to coincide with SB-3.5 (3.8 Ma) and its overlying 13.4-Ma flooding surface. Because paleontogical control is scarce, for the sake of simplicity it was decided to place the boundary of the last secondorder cycle at the base of the Upper Miocene. A major canyon-forming erosional event separates an Upper Miocene from a Pliocene-Pleistocene stack of sequences. The Upper Miocene is split into three thirdorder cycles, i.e. the SB-3.7-SB-3.8 (10.5-8.2 Ma), the SB-3.8-Sb-3.9 (8.2-6.3 Ma) and the SB-3.9SB-3.10 (6.3-5.5 Ma). The canyon-forming erosional event correlated with a worldwide Messinian unconformity which in eastern Venezuela may be enhanced by erosional processes associated with uplifts of the Venezuelan orogen. The Plio-Pleistocene (Unit VIII) is bounded at its base by SB-3.10, i.e. the 5.5-Ma erosional event. Lack of adequate faunal control permits only a very tentative age assignment to six third-order cycles that are separated respectively by SB-3.11 (4.2 Ma), SB-3.12 (3.8 Ma), SB 3.13 (3.0 Ma), SB-3.14 (2.4 Ma), SB-3.15 (1.6 Ma) and SB-3.16 (0.8 Ma).
CONCLUSIONS The Eastern Venezuelan Basin and its offshore continuation is a Neogene foredeep superposed on a Mesozoic passive margin. There is limited evidence suggesting that Jurassic rifting affected the South American Precambrian craton and some of its Paleo-
SEQUENCE STRATIGRAPHY OF THE EASTERN VENEZUELAN BASIN zoic cover. This Jurassic rifting event has also been reported from the Espino Graben in Venezuela and the Tacutu Graben of northeastern Brazil. The stratigraphy of eastern Venezuela and its offshore is encompassed by a major CretaceousCenozoic first-order transgressive-regressive cycle that begins with the Early Cretaceous, culminates with the middle Turonian 91.5-Ma maximum flooding event and ends with an overall Turonian to Recent regression. In the Venezuela foredeep this first-order cycle is interrupted by the inception of the foredeep phase with a 25.5-Ma basal foredeep unconformity. The Cretaceous to Paleogene passive margin stratigraphy is displayed on seismic profiles of the Venezuelan Atlantic offshore and can be correlated with outcrop sections in the Serranfa del Interior and in northern Trinidad. The Cretaceous to Paleocene is best subdivided into five second-order transgressiveregressive cycles bounded by a 131-Ma (basal Cretaceous) sequence boundary, four maximum flooding surfaces with inferred ages of early Aptian (111 Ma), late Albian (98 Ma), middle Cenomanian (95 Ma), middle Turonian (91.5 Ma) and a Late Paleocene sequence boundary (58.5 Ma). An Late Paleocene to Eocene second-order cycle (58.5 Ma-36 Ma) observed in the offshore area includes a dominantly pelagic unit. The Oligocene second-order cycle is characterized by narrow prograding siliciclastic systems which correlate with the Los Jabillos, Areo and Naricual formations of the Serranfa del Interior and some of the reservoirs of the giant E1 Furial-Carito oil fields. The Oligocene can be subdivided in two third-order cycles bounded respectively by sequence boundaries of 36 Ma, 30 Ma and 25.5 Ma. The Oligocene pinches out toward the craton, so that onshore the Oligocene is partially absent. Offshore the downdip equivalent of the prograding sequence appears to be absent in wells suggesting either sediment starvation and/or excessive condensation. A 25.5-Ma basal foredeep unconformity is related to the sudden deepening of the passive margin in response to the incipient emplacement of the folds and thrust sheets of the Serranfa del Interior and Trinidad. These in turn are related to the easterly displacement of the Venezuelan orogen in response to the oblique convergence of the Caribbean and South American plates. The Neogene foredeep fill of eastern Venezuela is best characterized by three second-order sequences bounded by sequence boundaries with inferred ages of 25.5 Ma, 16.5 Ma and 10.5 Ma. In the offshore a 5.5-Ma unconformity is associated with deeply incised submarine canyons. Prominent maximum flooding surfaces are at 16 Ma and 13.6 Ma. About sixteen Neogene third-order sequences can be recognized on seismic and are tentatively correlated with sequence boundaries proposed by Haq et al. (1987).
473
In many cases paleontological control in eastern Venezuela is not adequate to provide a substantial calibration of the postulated cycles. Future work should aim at refining the paleontological control and reconciling detailed log interpretation with seismic profiles that provide a higher resolution. Structural deformation in the Eastern Venezuelan Basin includes: (a) pre-Early Miocene and Neogene normal faulting involving the basement; (b) late Neogene reverse faulting involving the sediments and the basement; (c) extensive d6collement folding and thrust faulting associated with the Barbados accretionary complex; (d) N-S-trending down-tothe-basin pre-Pliocene listric normal faulting in the onshore area; and (e) extensive Neogene growth faulting associated with the Orinoco depocenter in the offshore area that is superposed on deeper basement-involved compressional folds.
ACKNOWLEDGEMENTS
We thank Intevep S.A., Lagoven S.A. and Corpoven S.A., all subsidiaries of Petroleos de Venezuela (ED.V.S.A.) for making the data for this study available for publication and for generously supporting this project. At Lagoven we particularly thank Carlos Sanchez, Yves Chevalier, Max Furrer and Antonello Lilliu. At Corpoven we are indebted to Emira Cabrera. Very special thanks go to our colleagues at Intevep S.A., Felipe Audermard, Pablo Klar, Mojtava Taheri and Irene MacQuahe for continuous encouragement during this project. Our reviewers Amos Salvador, Keith James and Paul Mann made many constructive comments which led to a profound restructuring of the original manuscript. We particularly appreciate Paul Mann's friendly perseverance and Mrs Elaine Bally's patient help in editing this paper.
REFERENCES
Aden, L.J. and Bierley, R.E., 1996. Structural Development of the Southern Basin, Onshore Trinidad: Implications for Hydrocarbon Entrapment (abstr.). Am. Assoc. Pet. Geol. Bull., 80: 1267. Arnstein, R., Cabrera, E., Russomano, E and Sanchez, H., 1985. Revisi6n estratigr~ficade la Cuenca de Venezuela Oriental. VI Conf. Geol. Venez., Mem., 1: 41-69. Audemard, EE., 1991. Tectonics of Western Venezuela. Ph.D. Thesis, Rice University,Houston, Texas (unpubl.). Audemard, EE. and Lugo, J., 1996. The Venezuelan Foredeeps: Structure (abstr.). Am. Assoc. Pet. Geol. Annu. Conv., A7. Audemard, F.E. and Lugo, J., 1997. Petroleum geology of Venezuela. Am. Assoc. Pet. Geol. Annu. Conv., Short Course Notes. Audemard, EE., Azpiritxaga, I., Bauman, E, Isea, A. and Latreille, M., 1985. Marco geol6gico del Terciario de la Faja
474 Petrolffera del Orinoco de Venezuela. VI Congr. Geof., Venez., 1: 70-109. Av6 Lallemant, H.G., 1997. Displacement partitioning and exhumation in the eastern Caribbean/South American plate boundary northeastern Venezuela. Tectonics, 16: 272-289. Azavache, A., Flinch, J., Giffuni, R., Martinez, R., Matta, S., Mitacchione, V. and Pereira, J.G., 1996. Tectono-stratigraphic evolution of the eastern Maturin Foreland Basin (Eastern Venezuela) (abstr.). Am. Assoc. Pet. Geol. Bull., 80 (8): 1270. Babb, S., Mann, E and Buffler, R.T., 1996. Three-stage slope reversal in northern and central Trinidad and its effect on megasequences in Cretaceous-Cenozoic basins (abstr.). Am. Assoc. Pet. Geol. Bull., 80 (8): 1271. Bally, A.W., 1989. Phanerozoic basins of North America. In: A.W. Bally and A.R. Palmer (Editors), The Geology of North America--An Overview. Geol. Soc. Am., Boulder, CO, Vol. A, pp. 397-446. Bally, A.W. and Snelson, S., 1980. Realms of subsidence. In: A.D. Miall (Editor), Facts and Principles of World Petroleum Occurrence. Can. Soc. Pet. Geol. Mem., 6: 1-94. Banks, L.M. and Driver, E.S., 1957. Geologic history of Santa Ana structure, Anaco structural trend, Anzo~itegui, Venezuela. Am. Assoc. Pet. Geol. Bull., 41: 308-325. Barr, K.W. and Saunders, J.B., 1965. An outline of the geology of Trinidad. Trans., 4th Caribbean Geol. Conf. Trinidad and Tobago, pp. 1-10. Bartok, E, 1993. Prebreakup geology of the Gulf of MexicoCaribbean: its relation to Triassic and Jurassic rift systems of the region. Tectonics, 12 (2): 441-459. Beaumont, C., 1981. Foreland basins. Geophys. J.R. Astron. Soc., 65: 291-329. Bejarano, C., Funes, D., Sarzalejo, S., Audemard, E and Flores, G., 1996. Sedimentary sequence evolution in a foredeep basin: Eastern Venezuela (abstr.). Am. Assoc. Pet. Geol. Bull., 80: 1273. Bellizia, G.A., Pimentel, M.N. and Bajo, O.R., 1978. Mapa geologico-estructural de Venezuela. Ministerio de Minas e Hidrocarburos, scale 1:500,000. Blow, W.H., 1969. Late middle Eocene to Recent planktonic foraminiferal biostratigraphy. 1967 Proc., Int. Conf. Planktonic Microfossils, Geneva, pp. 199-422. Bolli, H.M., Beckmann, J.E and Saunders, J.B., 1991. Benthic Foraminiferal Biostratigraphy of the Southern Caribbean. Cambridge University Press, Cambridge, 1230 pp. Bouysse, E and Westercamp, D., 1989. Subduction of Atlantic aseismic ridges and Late Cenozoic evolution of the lesser Antilles Island Arc. Tectonophysics, 175:349 -380. Carnevali, J.O., 1988. Venezuela nor-oriental: exploraci6n en frente de montafia. III Simp. Bolivariano, Soc. Ven. Geol., Caracas, pp. 69-89. Carnevali, J.O., 1989. Geology of new giant oil fields in mountain front of northeastern Venezuela. Proc., 28th Int. Geol. Congr., Washington, 9-19 July, 1, pp. 1240-1241. Carr-Brown, B. and Frampton, J., 1979. An outline of the stratigraphy of Trinidad. 4th Latin Am. Geol. Con. Field Guide, pp. 7-9. Case, J.E., Holcombe, T.L. and Martin, R.G., 1984. Map of geologic provinces in the Caribbean region. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The CaribbeanSouth American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162: 1-30. Chevalier, Y., 1993. A cross-section from the oil-rich Maturin sub-basin to Margarita Island. The geodynamic relations between South American and Caribbean plates. Tectonics and stratigraphy. Am. Assoc. Pet. Geol. Int. Conf., Caracas, March 11-14, Field Trip, 165 pp. Chevalier, Y. and Spano, E, 1996. Tect6nica del transecto E1 Pilar-Delta Amacuro, Cuenca Oriental de Venezuela (abstr.).
J. DI CROCE et al. Am. Assoc. Pet. Geol. Bull., 80:1281. Chevalier, Y., Gonzales, G., Mata, S., Santiago, N. and Spano, F., 1995. Estratigraffa secuencial y tect6nica del transecto E1 Pilar-Cerro Negro, Cuenca Oriental de Venezuela. VI Congr. Colombiano del Petr61eo, pp. 115-125. Cramez, C.D. and Vail, ER., 1990, Stratigraphy Cycles and Major Marine Source Rocks, Short Course, September 5, 1994, Houston, Texas, 82 pp. Crux, J., Sarzalejo, S., Audemard, E, Bejarano, C., Cabrera, S. and Funes, D., 1996. Timing of sequence deposition related to the development of the Eastern Venezuelan basin (abstr.). Am. Assoc. Pet. Geol. Bull., 80: 1284. Daza, J. and Prieto, R., 1990. Fallas de crecimiento en el ~ea de Mapirito-Monagas Central. V Congr. Venezolano de Geofisica, Caracas, 21-25 Oct., pp. 142-149. Dercourt, J., Ricou, L.E. and Vrielynck, B. (Editors), 1993. Atlas Thetys Paleoenvironmental Maps. Gauthier-Villars, Paris, 14 maps, 1 pl., 307 pp. Di Croce, J., 1995. Eastern Venezuelan Basin: Sequence Stratigraphy and Structural Evolution. Ph.D. Thesis, Rice University, Houston, Texas (unpubl.). Dillon, W.E, Paul, C.K. and Gilbert, L.E., 1985. History of the Atlantic continental margin off Florida: the Blake Plateau Basin. In: Geologic Evolution of the United States Atlantic Margin. Van Nostrand Reinhold, New York, pp. 189-215. Dillon, W.E, Scanlon, K.M., Valentine, EG. and Huggett, Q., 1988. Deep-sea erosion of a carbonate platform margin - - the Blake Escarpment (abstr.). Geol. Soc. Am. Bull., 20: 130. Eiras, J.E and Kinoshita, E.M., 1989. Soluqao de um Problema Estratigr~ifico atrav6s da Integra~o Regional, Um Exemplo da Bacia do Tacutu, Norte do Brasil. I Seminario de Interpretacao Exploratoria, Petrobras, 1989, Rio de Janeiro, pp. 243-251. Erikson, J.E and Pindell, J.L., 1993. Analysis of subsidence in northeastern Venezuela as adiscriminator of tectonic models for northern South America. Geology, 21:945 -948. Erlich, R.N. and Barrett, S.E, 1992. Petroleum Geology of the Eastern Venezuela Foreland Basin. In: R.W. Macqueen and D.A. Leckie (Editors), Foreland Basins and Fold Belts. Am. Assoc. Pet. Geol. Mem., 55: 341-362. Feo-Codecido, G., Smith, ED., Aboud, N. and Di Giacomo, E., 1984. Basement and Paleozoic rocks of the Venezuelan Llanos Basin. In: W.E. Bonini et al. (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162: 175-187. Flemings, EB. and Jordan, T.E., 1990. Stratigraphic modeling of foreland basin: interpreting thrust deformation and lithosphere rheology. Geology, 18: 430-434. Fox, EJ., Heezen, B.C. and Johnson, G.L., 1970. Jurassic sandstone from the tropical Atlantic. Science, 170: 1402-1404. Funkhouser, H.J., Sass, L.S. and Hedberg, H.D., 1948. Santa Ana, San Joaqufn, Guario and Santa Rosa oilfields (Anaco fields), central Anzo~itegui, Venezuela. Am. Assoc. Pet. Geol. Bull., 32:1851-1908. Furrer, M.A., 1979. Paleontology of well A and B. Lagoven, S.A., Internal report. Galea-Alvarez, F., Coriano, M. and Aponte, A., 1996. Facies, age and sedimentary environment of the E1 Cantil, Chimana and Querecual formations, Pertigalete, northeastern Anzoategui, Venezuela (abstr.). Am. Assoc. Pet. Geol. Bull., 80: 1291. Gallango, O. and Parnaud, E, 1995. Two-dimensional computer modeling of oil generation and migration in a transect of the Eastern Venezuelan basin. In: Tankard et al. (Editors), Petroleum Basins of South America. Am. Assoc. Pet. Geol. Mem., 62: 727-740. George, R.E, Jr. and Sams R.H., 1993. Eastern Venezuelan basin post-Jurassic evolution as a passive transform margin basin (abstr.). Am. Assoc. Petr. Geol. Bull., 77: 320. Giffuni, G. and Castro-Mora, M., 1996. Biostratigraphic se-
S E Q U E N C E S T R A T I G R A P H Y OF THE E A S T E R N V E N E Z U E L A N BASIN quence analysis of Oligocene-lower Miocene sections in the Orocual Field, Eastern Venezuelan basin (abstr.). Am. Assoc. Pet. Geol. Bull., 80: 1294. Gonzales, G., Mata, S. and Santiago, N., 1996a. Sequence stratigraphy of the Maturin Subbasin (Eastern Venezuela) (abstr.). Am. Assoc. Pet. Geol. Bull., 80: 1295. Gonzales, G., Mata, S., Odehnal, M. and Escandon, M., 1996b. Chronostratigraphic Charts of the Eastern Venezuelan basin (abstr.). Am. Assoc. Pet. Geol. Bull., 80:: 1296. Gonzalez de Juana, C., 1947. Elements of diastrophic history of northeastern Venezuela. Geol. Soc. Am. Bull., 58: 689-702. Gonzalez de Juana, C., Arozena, J. and Picard-Cadillat, X., 1980. Geologfa de Venezuela y de sus Cuencas Petrolfferas. Foninves Ed., Caracas, 1 and 2, 1051 pp. Gouyet, S., Unternehr, E and Mascle, A., 1992. The Demerara Rise and French Guyana, western equatorial Atlantic Ocean (abstr.). 4th Eur. Assoc. Pet. Geol. Conf., Paris, pp. 131-132. Hallam, J.M., Hancock, J.M., LaBraque, J.L., Lowrie, W. and Channel, J.E., 1985. Jurassic and Cretaceous geochronology and Jurassic to Paleogene magnetostratigraphy. In: N.J. Snelling (Editor), The Chronology of the Geological Record. Geol. Soc. London Mem., 10:118-140. Haq, B.U., Hardenbol, J. and Vail, ER., 1987. Chronology of fluctuating sea levels since the Triassic (250 million years ago to the present). Science, 235:1156-1167. Hardenbol, J., Thierry, J., Farley, M.B., Jaquin, T., de Graziansky, P.C. and Vail, ER., in press. The Mesozoic-Cenozoic chronostratigraphic framework. In: EC. De Graziansky, J. Hardenbol, T. Jacquin, ER. Vail and M.B. Farley (Editors), Sequence Stratigraphy of European Basins. Spec. Publ. Soc. Econ. Paleontol. Mineral. Harland, W.B., Cox, EG., Craig, L.E., Smith, A.G. and Smith, D.G., 1990. A geologic time scale 1989. Cambridge University Press, New York, 265 pp. Hedberg, H.D., 1937. Stratigraphy of the rfo Querecual section of northeastern Venezuela. Geol. Soc. Am. Bull., 48: 19712024. Hedberg, H.D., 1942. Mesozoic stratigraphy of northern South America. Am. Sci. Congr. VIII, U.S.A., 1940, Proc. IV, pp. 195-227. Hedberg, H.D., 1950. Geology of the Eastern Venezuelan basin (Anzo~tegui-Monagas-Sucre-Eastern Gufirico portion). Geol. Soc. Am. Bull., 61: 1173-1216. Hedberg, H.D. and Pyre, A., 1944. Stratigraphy of northeastern Anzo~tegui, Venezuela. Am. Assoc. Pet. Geol. Bull, 20: 1128. Hedberg, H.D., Sass, L.C. and Funkhouser, H.J., 1947. Oilfields of the Greater Oficina area, central Anzofitegui. Am. Assoc. Pet. Geol. Bull., 31: 2089-2169. Heezen, B.C. and Freeman-Lynde, R.E, 1976. Sheet 22, Atlantic Ocean 1 : 100,000. In: G. Choubert and A. Faure-Muret (Editors), Geological World Atlas. Commission for the Geological Map of the World, UNESCO. Huang, T., 1996. Nannofossil and sequence chronostratigraphy of a marine flooding surface in the Turonian of Trinidad (abstr.). Am. Assoc. Pet. Geol. Bull., 80:1302. International Stratigraphical Lexicon, 1956. Boletfn de Geologfa. Special Publication No. 1, Ministerio de Minas e Hidrocarburos, Editorial Sucre, Caracas, 664 pp Jankowsky, W.J. and Schlapak, G., 1983. Guyana offshore. In: A.W. Bally (Editor), Seismic Expression of Structural Styles. Am. Assoc. Pet. Geol., Stud. Geol., 15 (2): 2.2.3-47-2.2.3-50. Jordan, T.E., 1995. Retroarc foreland and related basins. In: C.J. Busby and R.V. Ingersoll (Editors), Tectonics of Sedimentary Basins. Blackwell, London, pp. 331-362. Kugler, H.G., 1953. Jurassic to Recent Sedimentary Environments in Trinidad. Bull. Assoc. Suisse Geol. Ing. Petrol., 20: 27-60.
475
Kugler, H.G., 1959. Geologic map of Trinidad and geological sections through Trinidad. Orell Ftissli, Zurich, E. Stanford, London, scale 1 : 100,000. Leonard, R., 1983. Geology and hydrocarbon accumulation, Columbus basin, offshore Trinidad. Am. Assoc. Pet. Geol. Bull., 67: 1081-1093. Liddle, R.A., 1928. The Geology of Venezuela and Trinidad. Mac Gowan, Fort Worth, Texas, 552 pp. Lilliu, A.G., 1990. Geophysical Interpretation of Maturfn Foreland, Northeastern Venezuela. MSc. Thesis, University of Houston, Houston, Texas (unpubl.). Lugo, J., 1991. Cretaceous to Neogene Tectonic Control on Sedimentation: Maracaibo Basin, Venezuela. Ph.D. Thesis, University of Texas at Austin (unpubl.). Lugo, J. and Mann, R, 1995. Jurassic-Eocene tectonic evolution of Maracaibo Basin, Venezuela. In: Tankard et al. (Editors), Petroleum Basins of South America. Am. Assoc. Pet. Geol. Mem., 62: 699-725. Marffn-Bellizia, C., 1974. Paleotect6nica del Escudo de Guyana. IX Conf. Geol. Inter-Guyanas, Ciudad Guyana, Venezuela, Memoria, Bol. Geol., Publ. Esp., 6, pp. 251-305. Mascle, A. and Letouzey, R, 1990. The Geological Map of The Caribbean. Technip, Paris, 2 Plates. Mendoza, V., 1977. Evoluci6n tect6nica del Escudo de Guyana. II Congr. Latinoamericano Geol., Caracas, Noviembre, 1973, Memoria, Bol. Geol., Caracas, Publ. Esp., 7, pp. 2237-2270. Miall, A.D., 1995. Collision-related foreland basins. In: C.J. Busby and R.V. Ingersoll (Editors), Tectonics of Sedimentary Basins. Blackwell, London, pp. 393- 424. Mitchum, R.M. and Van Wagoner, J.C., 1991. High frequency sequences and eustatic cycle in the Gulf of Mexico Basin. In: J.M. Armentrout and B.E Perkins (Editors), Sequence Stratigraphy as an Exploration Tool - - Concepts and Practices in the Gulf Coast. Gulf Coast Section, Society of Economic Paleontologists and Mineralogists Foundation, Program and Abstracts, Eleventh Annual Research Conference, Houston, Texas, pp. 257-267. Murany, E.E., 1972. Tectonic basis for the Anaco fault. Am. Assoc. Pet. Geol. Bull., 56: 860-870. Palmer, V.W., 1945. Fossil fresh-water mollusca from the State of Monagas. Bull. Am. Paleontol., 3: 1-34. Parnaud, E, Gou, Y., Pascal, J.C., Truskowski, I., Gallango, O., Passalacqua, H. and Roure, F., 1995. Petroleum Geology of the central part of the Eastern Venezuelan Basin. In: Tankard et al. (Editors), Petroleum Basins of South America. Am. Assoc. Pet. Geol. Mem., 62: 741-756. Passalacqua, H., Fernandez, F., Gou, Y. and Roure, E, 1995. Crustal architecture and strain partitioning in the Eastern Venezuelan Ranges. In: Tankard et al. (Editors), Petroleum Basins of South America. Am. Assoc. Pet. Geol. Mem., 62: 667-679. Persad, K.M., 1979. Outline of the Geology of the Trinidad Area. Trans., 4th Latin American Geol. Conf., Trinidad and Tobago, pp. 738-758. Persad, K.M., 1985. Presentation of a new tectonic map (and accompanying sections) of Trinidad and Tobago. 1st Geol. Conf. Geol. Soc. Trinidad Tobago, Trans., pp. 1-6. Petroconsultants, 1989. Guyana Basin--Geology and Petroleum Potential. Petroconsultants, Report, pp. 1-11. Pindell, J.L., 1985. Alleghenian reconstruction and the subsequent evolution of the Gulf of Mexico, Bahamas and ProtoCaribbean Sea. Tectonics, 4: 1-39. Pindell, J.L. and Barrett, S.F., 1990. Geological evolution of the Caribbean region; a plate-tectonic perspective. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Vol. H, Geol. Soc. Am., pp. 405-432. Pindell, J.L. and Dewey, J.E, 1982. Permo-Triassic reconstruction of western Pangea and the evolution of the Gulf of Mexico/Caribbean region. Tectonics, 1: 179-212.
476 Prieto, R., 1987. Seismic Stratigraphy and Depositional Systems of the Orinoco Platform Area. Ph.D. Thesis, University of Texas, Austin, Texas (unpubl.). Renz, H.H., Alberding, H., Dallmus, K.E, Patterson, J.M., Robir, R.M., Weibord, N.E. and Masvall, J., 1958. The Eastern Venezuela Basin. In: L.G. Weeks (Editor), Habitat of Oil. Am. Assoc. Pet. Geol., Symp., pp. 561-600. Ross, M.I. and Scotese, C.R., 1988. A hierarchical model of the Gulf of Mexico and Caribbean Region. Tectonophysics, 155: 139-168. Rossi, T., 1985. Contribution ~i l'6tude Geologique de la Frontibre Sud-Est de al Plaque Caraibe: la serrania del interior (Venezuela) sur le Transect Cariaco Mathrin. Ph.D. Thesis, Univ. Bretange Occidentale. Brest, 338 pp. Rossi, T., Stephan, J.E, Blanchet, R. and Hernandez, G., 1985. Etude g6ologique de la Serranfa del Interior oriental sur le transect Cariaco-Maturin. Rev. Inst. Fr. Pdt. Paris. Technip, 42 (1): 3-30. Roure, E, Carnevali, J.O., Gou, Y. and Subieta, T., 1994. Geometry and kinematics of the North Monagas thrust belt (Venezuela). Mar. Pet. Geol., 11: 347-362. Salvador, A. and Stainforth, R.M., 1968. Clues in Venezuela to the geology of Trinidad and vice versa. 4th Geol. Conf. Caribe, Trinidad, pp. 31-40. Sams, R.H., 1995. Interpreted sequence stratigraphy of the Los Jabillos, Areo and (subsurface) Naricual formations, northern Monagas (area), Eastern Venezuelan basin. Bol. Soc. Ven. Geol., 20 (1-2): 30-40. Shepard, EE, 1981. Submarine canyons: multiple causes and long-time persistence. Am. Assoc. Pet. Geol. Bull., 65: 10621077. Speed, R.C., Westbrook, G.K., Biju-Duval, B., Ladd, J.W., Mascle, A., Moore, J.C., Saunders, J.B., Schoomaker, J.E. and Stein, S., 1984. Lesser Antilles arc and adjacent terranes. Atlas 10, Ocean Margin Drilling Program, Regional Atlas Series. Marine Science International, Woods Hole, 27 sheets. Sprague, A.R., Lame, D., Faulkner, B., Sykes, M.A. and Aden, L.J., 1996. Sequence stratigraphy, facies architecture and reservoir distribution, Cretaceous lowstand fan reservoirs, Southern Basin, onshore Trinidad (abstr.). Am. Assoc. Pet. Geol. Bull., 80 (8): 1338. Stephan, J.E, De Lepinay, B.M., Calais, E., Tardy, M., Beck, C., Carfantan, J.C., Olivet, J.L., Vila, J.M., Bouysse, E, Mauffret, A., Borgois, 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., 6: 915-919.
J. DI C R O C E et al. Tankard, A.J., 1985. On the depositional response to thrusting and lithospheric flexure: examples from the Appalachian and Rocky Mountain basins. In: P.A. Allen and P. Homewood (Editors), Foreland Basins. Int. Assoc. Sedimentol. Spec. Pub., 8: 369-395. Vail, P.R. and Wornardt, W.W., 1990. Well log-seismic sequence stratigraphy: an integrated tool for the 90's. In: J.M. Armentrout and B.F. Perkins (Editors), Sequence Stratigraphy as an Exploration Tool n Concepts and Practices in the Gulf Coast. Gulf Coast Section, Society of Economic Paleontologists and Mineralogists Foundation, Program and Abstracts, Eleventh Annual Research Conference, Houston, Texas, pp. 307-320. Vail, P.R., Mitchum, R.M. and Thompson, S., III, 1977. Seismic stratigraphy and global changes of sea level, part 3. Relative changes of sea level from coastal onlap. In: C.E. Payton (Editor), Seismic StratigraphymApplication to Hydrocarbon Exploration. Am. Assoc. Pet. Geol. Mem., 26:63-81. Vail, P.R., Audemard, E, Bouwman, S.A., Eisner, P.N. and Perez-Cruz, G., 1991. The stratigraphic signatures of tectonics, eustacy and sedimentology n an overview. In: G. Einsele et al. (Editors), Cycles and Events in Stratigraphy. Springer-Verlag, Berlin, pp. 617-659. Veeken, P.C.H., 1983. Guyana Basin, offshore Northeast South America. Guyana offshore. In: A.W. Bally (Editor), Seismic expression of structural styles. Am. Assoc. Pet. Geol., Stud. Geol., 15 (2): 2.3.3-1-2.3.3-2. Villaroel, V., 1993. Santa Rosa Field, Venezuela, Eastern Venezuela Basin. In: N.H. Foster and E.A. Beaumont (Editors), Treatise of Petroleum Geology, Atlas of Oil and Gas Fields. Am. Assoc. Pet. Geol., Structural Traps, VIII, pp. 193-210. Villeneuve, M. and Komara, S., 1991. Lower Paleozoic transgressions and regressions in the Bove Basin (Guinea and Guinea Bissau, Africa) Stratigraphic, sedimentologic and paleogeographic data. J. Afr. Earth Sci., 12: 67-77. Villeneuve, M., Diallo, M.C., Keleba, F., Kouroma, S., Paris, F. and Racheboeuf, P.R., 1989. New paleontological data from the Paleozoic series of the Bove Basin (Guinea, west Africa): stratigraphical consequence. Compt. Rend. Ser., 309 (16): 1583-1590. Vivas, V., 1986. Contribuci6n al estudio de la Serranfa Oriental (Venezuela), estratigraffa y tect6nica de la regi6n de Berganffn-Santa Ines (Edo Anzo~itegui). Thesis 3me cycle, Univ. Bretagne Occidentale (unpubl.). Wu, S., Bally, A.W. and Cramez, C., 1990. Allochthonous salt, structure and stratigraphy of the north-eastern Gulf of Mexico: structure. Mar. Pet. Geol., 7 (4): 334-370.
Chapter 1 7
Structure of the Gulf of Paria Pull-Apart Basin (Eastern Venezuela-Trinidad)
J.F. FLINCH, V. RAMBARAN, W. ALI, V. DE LISA, G. HERNANDEZ, K. RODRIGUES and R. SAMS
The Gulf of Paria lies entirely within the broad strike-slip plate boundary zone of the southeastern Caribbean. The structure of the Gulf consists of a complex set of transtensional basins superimposed on a fold-and-thrust belt (Serranfa del Interior of eastern Venezuela). This province extends to the east into the Caroni Basin of northern Trinidad and to the west into the San Juan graben and other minor extensional basins of the Guanoco area. The main structural elements of the Gulf of Paria are: the Casanay-Arima fault bounding the transtensional province to the north, the Warm Springs fault to the south and the Domoil and Gopa Highs located in the central part of the Gulf of Paria. Rapid extensional collapse since the late Neogene resulted in deep half-grabens, which are characterized by large offsets along the major extensional faults and by shallow-water sedimentary fill. Transtensional tectonics in the Gulf of Paria was coeval with north-vergent thrusting in the Pedernales region to the south. The north-vergent imbricates constitute a passive-roof duplex with respect to the underlying south-vergent thrusts of the Serranfa del Interior of eastern Venezuela. Late Pleistocene-Holocene compression in the Gulf of Paria has caused minor positive inversions involving the main basin-bounding faults. Seismic data used in this study have enabled us to significantly modify existing wrench models.
REGIONAL SETTING The Gulf of Paria is located east and downplunge of the eastern Venezuelan fold-and-thrust belt (Serranfa del Interior) (Fig. 1). Its northern limit is the Paria Peninsula in Venezuela and the western extension of the Northern Ranges of Trinidad (Fig. 2). The Gulf of Paria is the link between the Serranfa del Interior of Venezuela and the Central and Southern Range of Trinidad. The Paria Peninsula and the Northern Range consist of allochthonous metamorphic terranes, assigned to the Cordillera de la Costa Nappe (Bellizzia and Dengo, 1990). This nappe consists of pre-Mesozoic basement rocks (Sebastopol Complex) unconformably covered by thick Jurassic-Lower Cretaceous metasediments (Caracas Group) and metamorphosed igneous rocks, intruded by large granitic plutons (Bellizzia and Dengo, 1990). The Cordillera de la Costa Nappe has been considered as an accretionary complex associated with oblique tectonic convergence (Case et al., 1990) or a collisional arc complex (Burke, 1988). The extensional faults of the Gulf of Paria extend further west into the Venezue-
lan Cordillera, forming the San Juan graben and the Guanoco area (northern Guarapiche province), where normal faults overprint thrust sheets of the eastern Venezuelan fold-and-thrust belt (Fig. 1). The structure of the Gulf of Paria extends eastward into onshore Trinidad, west of the Barbados accretionary wedge (Fig. 1).
The South Caribbean plate boundary The Gulf of Paria lies in the broad South Caribbean Plate boundary (Burke et al., 1984). The Caribbean Plate is moving in a fight-lateral sense, with respect to the South American Plate. While most of the eastward escape of the Caribbean Plate is accommodated in the north by the Cayman pull-apart and the strike-slip systems of Motagua (Guatemala) and Jamaica, the southern boundary is taken by several strike-slip systems, the F a l c o n Aruba (Macellari, 1995), Cariaco Trough (Schubert, 1982; Mann and Burke, 1984) and the Gulf of Paria being the most significant ones. In the past many geoscientists considered that the E1 Pilar fault represented the Caribbean-South
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 477-494. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
478
J.E FLINCH et al. 64 ~
63 ~
62 ~
61 ~
60 ~
59 ~
11~
lo ~
#
8~ [,
-
.. ] Accretionary metamorphic terranes
t- ---[
Barbados Accretionay wedge
[~il
Gulf of Paria Extensional province
l
Deformed South-American Passive Margin
[~_~]
Carapita-Nariva Accretionary Wedge
[+++++ 1 GuayanaShield
..1
~
Maturfn Foreland Basin Undeformed Passive Margin
Thrust Fault
Normal Fault ~,,
Strike-slip fault
...................................Shale Ridge
Fig. 1. Structural sketch map of the southeastern Caribbean region. The boundaries of the passive margin and the outline of the Barbados accretionary wedge are traced after Di Croce (1995), the structure of the Maturfn Basin is based on LAGOVEN Internal Reports and the structure of the Gulf of Paria is based on this study. Notice the location of the study area. A F = Arima fault; C F = Casanay fault; C R F -Central Range fault; L B F = Los Bajos fault; S F F = San Francisco fault; S J G = San Juan graben; U F = Urica fault.
American Plate boundary (Molnar and Sykes, 1969; Schubert, 1982; Soulas, 1986). According to Vierbuchen (1984), fight-lateral displacement of at least 150-300 km is required to explain the gravimetric field distribution. Nevertheless, the E1 Pilar fault has an ambiguous field expression, and shows no evidence for significant right-lateral displacement. According to Ball et al. (1971) and Gonzalez de Juana et al. (1980) the E1 Pilar fault is a steeply dipping normal fault that constitutes the southern boundary of the Cariaco Trough, while the Mor6n fault would be its northern limit. Seismicity data do not support continuity of the E1 Pilar fault through Trinidad (Speed, 1985). Burke (1988) considers the plate boundary to be at least 200 km or more wide. The strike-slip motion is distributed among several faults, one of them being the North Coast fault zone, located north of Trinidad (Robertson and Burke, 1989).
Previous work Several models have been proposed to explain the complex structure of the Gulf of Paria (Fig. 2).
Many geoscientists as an example of a strike-slip basin have cited the Gulf of Paria. In most of these models the E1 Pilar fault plays a major role as the main northern bounding fault of the strike-slip system (Fig. 2a,b); according to all models (Fig. 2 a c), the Los Bajos fault as well as other N W SE-trending faults, represent strike-slip faults related to right-lateral transcurrent systems. Under these schemes, the Los Bajos fault joins the E1 Pilar fault with the Southern Range of Trinidad. The data presented in this study (Fig. 2d) significantly constrain the current models on the Gulf of Paria and the entire southeastern Caribbean region. Most of the models shown on Fig. 2 are based on much less data than our model which integrates data from Venezuela and Trinidad. A common feature is that they all emphasize the role for the Los Bajos fault as a link between the E1 Pilar fault and the South Coast fault (Salvador and Stainforth, 1968; Munro and Smith, 1985; Algar, 1995). Most of these models were based on a limited seismic data set. The structural relationships presented by these models are not consistent with the new data set based on data from Trinidad and Venezuela presented in this study.
GULF
100 km
Northern
100 km
%
. . . . .
%
9.....
San Juan
Grabe~
El Pilar Fault
CARIBBEAN
(b)
Maturin
El Pilar Fault
CARIBBEAN
Casana,
GULF
PARI
ULF
SEA
SEA Ran(
100 km
Arima
Spain
Fig. 2. Four tectonic models of the Gulf of Paria in a time perspective: (a) Salvador and Stainforth (1968); (b) Munro and Smith (1985); (c) Algar (1995); (d) this study. Notice main differences regarding the role of Los Bajos fault and the continuity of E] Pilar fault through Trinidad.
El Pilar Fault
CARIBBEAN
SEA
LF
El Pilar Fault
(a)
SEA
CARIBBEAN
480
J.E F L I N C H et al.
STRATIGRAPHY OF THE GULF OF PARIA
STRUCTURE
The stratigraphy of the Gulf of Paria is strongly controlled by its structural evolution. The stratigraphy of this region can be subdivided into a lower fold-and-thrust belt stratigraphy and an upper Northem Transtensional Basin fill and the Southern Compressional Zone basin successions (Fig. 3). With respect to the folded belt, this is the area where the classical Venezuelan and Trinidadian stratigraphic units meet. The superimposed transtensional fill has been stratigraphically subdivided in Trinidad, where most of the units are widely exposed on the island. In this work we will use the classical Trinidad stratigraphic names for the Northern Transtensional Basin. The Southern Compressional Zone is occupied by sedimentary basins that have equivalent stratigraphy in the eastern Venezuelan Basin and in southern Trinidad. Table 1 is based on twelve wells in the area (see Fig. 5) and summarizes the stratigraphy of the region.
A regional cross-section through the Gulf of Paria (see Fig. 1 for location) reveals the presence of two distinct structural provinces: a Southern Compressional Zone and a Northern Transtensional Basin (Fig. 4). Both units are superimposed on an underlying fold-and-thrust belt. These units are separated by a major fault system: the Warm Springs fault (Flinch and Rambaran, 1996) (labeled WSF in the sections). Fig. 5 is a detailed structural map integrating data from the Venezuelan and Trinidadian part of the Gulf of Paria and the island of Trinidad. The map shows the most relevant present-day structural features of the region and the main tectonic provinces.
Southern Compressional Zone AGE
0 Z
0 -Q
The structure of the Gulf of Paria is illustrated by a series of regional seismic profiles (Fig. 6). The
Northern Transtensional Basra
[
S
N : .'_."2.'2.'2:- "2..2.'2 - f...2 : ~ ~ - " " . ~
'';''~"~-'~'"~-''~;-':':::":':::"::"-'~
LLI P L E I S T o . (.9
Regional seismic sections
.9 " .9 .'.'.':..'..--.
--
.-~
- . % .
--
" .# .':.-'.IL'..-
-
11D.il.
" ,dl . k T . . ' . 1
"-.*: ( . " : : - : - - : ~ : - ' ~ ' L : 2 , i . f , ' 3 ; , ' ~ ;
'
Central
Campana
UPPER
~--
Highs
High
Paria Peninsula
MIDDLE
LOWER OLIGOCENE uJ
(.9 O
EOCENE
n
PALEOCENE
iii
L ......
~,
O3
8Iii
UPPER
O < I
LOWER
..
.
/ / / ~ /
/~ A A A A A A A A A A A A A A A A A A A, Ak\ , A A A A A A A A A A A A A A A A A A A A ~.~ ~ /x /k /k A. A. /k /k /k A.
/
I
"~
,#k,,
O
LEGEND
~'~"~'1 Marls
Deep-water sandstones
A n Evaporites
Fig. 3. Schematic litho-stratigraphic north to south section across the Gulf of Paria.
,,
Fluvial-alluvial sandstones I ', 2 "1 Metamorphic Rocks
I
ABSENT
I /.,atus or erosion/
\~
\
conglomerates
Deep-water shales 1------ ~1 Shallow-water shales and claystones
-
Ii~o~ ol Fan-delta
Platform limestones t i : ! : ! I Shallow-water sandstones
/
X-
' \
,
i
S T R U C T U R E OF THE G U L F OF PARIA P U L L - A P A R T B A S I N
481
Table 1 Description of the main stratigraphic units of the Gulf of P a r i a - correlation between eastern Venezuela and Trinidad
Formation
Age
Paria/Mesa PLEISTOCENE
Erin/Morne L'Enfer
PLIOCENE
~ ._ 13_
Lithology
Environment
Mudstones,sandstones and occasional conglomerate beds
Thickness Range (meters)
Fluvial-alluvial deposits Delta Plain
900-100
,,~
800-200
o
Springvale ~. Manzanilla ~
UPPER
._1 La Pica/Forest Cruse Lengua
MIDDLE LOWER
9~. 9 grasso ~ O
EOCENE
PALEOCENE
Cipero Nariva
OLIGOCENE
Calcareous silt, glauconitic sandstones
fluvial-littoral
Sandstones and siltstones with interbedded shales
Glauconitic calcareous clays, silt sands and boulder beds Light grey and greenish grey marls, white chalky marls and occasional radiolaria rich marly-clays Pyrite-bearing fossiliferous black shales and mudstones, limely siltstones and glauconitic sandstones Calcareous shale
> Lizard Springs , San Antonio
Chimana
El Cantil iii
garranquin/ Cuche
Pro-delta
S-vergent thrusting
600-0
Outer shelf-slope
500-0
c.-
Platform, coastal 700-0
Nodular and micritic black to grey limestones and occasional black mudstone and siliceous claystone
Deep-water
Micritic and bioclastic limestones, occasional sandstones and calcareous shales
Platform, r e e f a l
Interbedded sandstones, limestones, foraminifera rich black shales and reddish brown clay Black silty micaceous shale with carbonaceous laminae
I NW
> u) 1600-600
(Delta front-Prodelta
13_
700-500 Slope to deep-water
Evaporites, mainly gypsum and anhydrite
Southern Compressional Zone
Transtension N-vergent thrusting
Deep-water
Massive, yellow, fine to medium grained quartzy sandstones Siliceous shales and sandstones with chert levels
Couva
1500-0
Herrera, Retrench and Karamat Turbiditic sandstones
Calcareous clays and marls with abundant foraminifera fauna Sandstones
Inversion
Foredeep
Deep-water deposits
Shales, blue-grey mudstones and occasional sandstones lenses or channels
San Fernando
?
1700-0
Dark greenish grey marl rich in foraminifera
= Querecual/ c~ Naparima Hill
0 w 0 < bu.i
Fluvio-deltaic neritic
Caratas/Navet
San Juan o~ LIJ O~
Clays and glauconitic sandstones
c
Tectonic Events
shallow-water coastal
500-0
Northern Transtensional Basin S
Gopa High
Pedernales Passive-Roof Duplex
N Paria Peninsula 0 1 2 3
4
4 5 6 7 8 9 Extensional Basin Fill Maturin Foreland Basin Pleistocene
Serrania del Interior Fold-and- Thrust Belt Paleogene-
~
Pliocene Upper Miocene
I
AIIochthonous Metamorphic Terranes (Cordillera de la Costa Nappe)
MiddleMiocene
I:,-- ,t
Cretaceous Couva Evaporites
WSF
WarmSprings Fault
CAF
Casanay-Arima Fault
! Pre'CretaceOus
Fig. 4. Generalized cross-section through the Gulf of Paria, showing the main structural provinces. The section is approximately three times vertically exaggerated (V ~ 3H). See Fig. 1 for location. The section displays the relationship between the northern extensional basin with the southern compressional zone and the underlying folded belt units of South America and the allochthonous metamorphic terranes of Caribbean affinity.
482 sections below describe the main structural units and tectonic elements of the region from west to east. Section I This N-S-trending section illustrates the contact between the transtensional system of the central Gulf of Paria (Northern Transtensional Basin) and the Southern Compressional Zone. The contact between the two structural units is a set of northward-dipping normal faults, referred to here as the Warm Springs fault (WSF). To the north, mostly south-dipping listric normal faults account for the rotation of the overlying Neogene succession. North-vergent thrusts and related ramp anticlines are present in the southern part of the profile. The geometry of the piggyback basins suggests thrust propagation from south to north. The relationships between thrusting and folding to normal faulting suggest nearly coeval north-directed thrusting and transtension. The contact between the basin and the allochthonous metamorphic terranes is a steeply dipping south-vergent normal fault referred to as the Casanay fault (see Fig. 5). Section II This N-S section again shows the boundary between the Southern Compressional Zone and the Northern Transtensional Basin, represented by the WSE From south to north this section shows southdipping compressional imbricates of post-Middle Miocene sediments. The seismic data suggest a basal decollement around 2.4 s (TWT). The imbricates were penetrated by well C which shows duplication of the Pliocene section. Thrust imbricates are offset by northward-dipping normal faults that constitute the boundary of the Northern Transtensional Basin. The Gopa High (see Fig. 5 for location) occurs in the central part of this large extensional basin. Well E penetrated the Gopa High where the Miocene section is most shallow. South- and northdipping normal faults delineate this high. Thus, the Northern Transtensional Basin consists of a southern half-graben, the central Gopa High and a northern half-graben. To the north, southward-dipping normal faults separate the Domoil High from the northern half-graben of the basin. Section III This N-S-trending section shows that the Domoil and Gulf Highs are separated by a narrow half-graben. In map view (Fig. 5) this half-graben has a NW-SE trend. South of the Domoil High the section offers a dip view of the normal faulting style. Listric normal faults sole out into a shaly Lower Miocene section that constitutes the main extensional detachment. Shale ridges bounded by steeply dipping normal faults occupy the central part of the extensional system. The overlying Upper Miocene-
J.E FLINCH et al. Pleistocene section is rotated due to normal faulting, but the amount of growth is not significant. Flat reflectors onlapping onto folded beds suggest that the northern half-graben has undergone positive Pleistocene inversion on its southern flank (the top of the Pliocene is deformed). Pleistocene sediments directly onlap the Domoil and Gulf Highs dating the end of tectonic activity. Section IV This N-S- to NE-SW-trending section shows the contact of the allochthonous metamorphic terranes with the eastern Venezuelan fold-and-thrust belt that underlies the present-day basin. This northernmost area is occupied by a south-vergent thrust system. Thrust faults are offset by a narrow half-graben previously described in section III, to the south is the Domoil High. Seismic data suggest southward thrusting within the Gulf and Domoil Highs (see sections III and IV of Fig. 6). Pleistocene sediments onlap onto the Lower Cretaceous strata, which belonged to the underlying folded belt. This is one of the few seismic evidences of south-verging thrusting underneath the present-day basin. South of the Domoil High, down-to-the-basin listric to steeply dipping normal faults offset the Cretaceous to Pliocene section. Miocene shale constitutes an important extensional decollement level (Brasso and Nariva Formations). A shale ridge (well H) occupies the central part of the extensional basin. South of these wells the southern branch of the extensional system is represented by normal faults that separate a thick Upper Miocene-Pliocene section to the north from a compression anticline constituted by Oligocene-Cretaceous strata. Section V This NE-SW to NW-SE zigzag section shows the western part of the Avocado-Couva High. This high consists of a WNW-ESE-trending anticline related to south-vergent thrusting. According to well data the core of this high is occupied by the Couva evaporite. The prominent shale ridge, already seen in the previous section, occupies the lower part of the half-graben fill. Notice the large offset along the Warm Springs fault in the southern part of the profile. This fault can be pictured as the master fault of the transtensional system. Section VI This roughly N-S-trending section shows the structure of the Northern Transtensional Basin. The Warm Springs fault system bounds the basin to the south, constituting the master fault of the extensional system. Southward-dipping listric normal faults offset the Cunapo Conglomerate in the north and the Manzanilla and Springvale Formations in
/ s e q ! . 12.3/O]1200[/|U'J~!AOSIWA~A~A~//:dl]q 112 IJOlSJ~A JOlO~) "UO!112301 JOJ l "8!cl oaS "£pnls slql u! p~sn SllOA~ ~ql j o uo!112oo I ~q] pu12 slu~m~l~ 112Jn13nJls Ut121.I.t ~4:1 s~u!lmO 12!J12d j o j[n O oql j o d e m 112Jnl:mJlS oqj, " ( I 9 6 I ' £ g 6 I ) Jol~nN uJoJj p ~ g ! p o u J pue p~U!lduJ!s s! pu121s! P12P!U!-LL ~ql j o ,~8OlO~8 ~qj, "p12p!u!Ji pu12 e!J12d j o j l n O ~ql j o deuJ |eJnlonJ] S "g "~!cI
£[oJ!lUO st
deuJ ~qj. "£pnls s!q] uo pas12q
line=! sGu!Jds uJJeAA _,ISAA llne=l so!eEl soq 4 8 q llne_,l/£eueseo _-!0 IIne-I eU.l!JV -IV ------..~_-. . . . . .
euilouAs eU!lO!luv
~"
Sl,lne~ ],snJql ],ueGJeA pJeMqlnos
~,
seoueoloA pnlA! se6p!EI eleqs ~ '
eueoo!iAI Jaddn eseEi oueoO!ld eseEt eueools!elcl eseEt
.
.
.
.
.
eGp!l=l eleqs llnej d!!s-e>l!JlS ~ . ~r
euilOllUV +
sllnel lsnJqx lueBJe^ pJe~qlnos ou!IouXs
sqG!H q sMoq e
llnel IeLUJON f u!seE! leUO!SUelSUeJl uJaqlJoN
I-':.-:-t
o!qdJoLuelelAi snouoqiqoOllV
.
>-
]
, .
.
eullou/[s
eu!l,')u/~s / eu!lo!luv UO!SJeAUl ~ / "
s e ! 7! w J oj u o o u n J O[e lN SOU18JJE)_L
"
[
~, +%
eu!lo!luv
v
sllneJ lsnJq/IUE)6JeA pJeMqlJON
J
uenp' ,
-
~i-ii~:;i~i!!i!~ii!~iiiiiii!~;iii!!i~i!~iii!i!iiii~ii~iiiiii~iiii~;iii~!~!~i~i!~!!~ii~i!~!i!~!!~!i~ii!~!~i! i:i~!iii!iiii!iiiiiiii!i:iii!i;i!ii
~iiTi!ii!!!iT!iT!i!!iiiii!iiiili~i~!ii!7~iiii!iT~i!iiiii!ii!iii!i!:i~iii!~
4SAA
S/V, 0 .._.I
! !/7:
:: !
(
/
i
tl ti li
,
.'"
,:
I t
l :,:~: ......
I#
===========================
# u o a e o
•
--
lit
,~
S ~P
•"
. . . . . .
tt
,,,s
I
JO ,UOd
U I l ~e ¢ CU U
.
o I
kxC_
.
.
.
.
.
.
.
.
.
~,
.. ..,
.
~
....
I
....
.-
I
/
"1
x
/
\
/
~
t
I
/
i
., x
i
"
]
\
7
~
.I
i
....... .-"
\
t
,
..
7
...
i
--.
"
I
_ "~
"-
-,,
.
.
.
.
/
~
/
I
.
~
j.
.
I
i
i~.
/
;,
.
i
~
/
z
1
.
\
.
I
-
-
l
-.
. ""
'"
~
s
.
'
".
.
/
~ -
•
,
,
..
i
.
"
"
,
s-..
/
/
."
"
,
. '
i . -£ a " "
~ ~"I"~-
."
"
"
"
-.
I
.
."
"
,
,
-.I
/
<
-"
,
.
.
.
.
.
.
.
.
.
.
- . .
""
~
.
.
.
.
I
.
.
.
.
. ,.," . ....
I
.
.
.
.
.
.
/'
.
/.
. .,.
/
. /
.
I
. /
I
.
i
""
.
%
..
/
,
t
_
\
.
. ~,
~
-+"
!
. ""
\
. "-~
~
;
. .;
~
t
I
I
~
/
i
I
t
..
/
~
t
.
/
i
/
...
/
I
t
..-
/
i
/
i
I
/
i
/
I
i
..
i .....
"
x
-
~>
"" -
\
.
/
.
I
"
x
i
'. .
/
I
l
i
~
x
i
.
I
~
"
- ..---
/
i
l
"
/
-
/
. . . . . ..,, . --..
i
/
~
j
<"
-
i
" .-.
/
"
.
. ./
_
I ;
.
.....
....
i
--
.
.
_
'1
-
. /"
I
i
--
"_~"
. --
¢
s
'
:"
~
!
I"
,7-(-
j /
.
..
/ . " ~~. ~
/-
!
.~
,,
/
':::=======================================:::::::
s
s
~-
_ ......
-
-
."
,
-
/
~
"
"
i
"
/
/seq!Jl~3/olp,3Ol/lU'JO!AOSlO'~,~//:duq m uo~sJo~ JOlO~) "sou!I o~]!JlS o.re IIIA pu~ IIA suo!13os pu~ sou!I d!p oJe IA o] I suo!]3oS ,~pms s!ql u! u~oqs uo!13os 3~tus~os ll~UO!~Oj pul~ SliO~ jo uo!1~3o I oql SA~OqSd~uJ II~Jnl3nJls o q i "~!J~c[ jo JlnD oql q~noJql suo!13os 3[tustos [~uo!8oJ jo s~u!/~eJp-ou!"1 "9 "8!~
-t
£
.,.
--! 8 m
0
-,=$9
IliA
I
~
~I'
II
I
a
~
('[oJd) V ~ " A/IS
N
O
0 9SS q6lH
a
q6!H euedweo
I^
opuewe4 ues
m
i.
i. o
0
I
q6lH ,owoa
9
AI ~
,14
I II
9s~
I I
w>i O~
' ,M NA/I
III
IIA
fined , { e u e s e o 9 0
ouooo!~ Joddn
lined sl~u!JdS LuJelH 9S/1/1 eueoo!~
snoeoemJO -o,sse~nr
elPp!~
-snoeoe:leJO
:~i~:,~ ~.,.,~-~i
SHNVHM9..L OIHcl~Ol/llYJ..31,41
0
17
173S ISR~IHI-GNV-GTOJ
, w
A/IN I I A
MS
v ]
q61H opeoo^v-e^noo
eueools~elcl
O
v N
_,'3'N
.
-a
[~:~-,?,
~
f;;;;;;;;;;;;:;;]
eueoo~ld~
o ~-
'
17 i~':~¸~¸~ii~'!!i~'i ~¸ ~:-~¸I~!~
"~~ ~'
SNOII O g S 9}II~IIS
........
i~
'
~ :.
'
•
'
, -:~
~, . . . .
; ,
• :.i_
.:i~.i
m
o
0
-=,S
U6!H
I L
opueuJe=l ues
;,
~ !'~'
,~",~ :~',~ : i : ~ ;
'
:;-:
::.;~
~, ~;
,:,
~:
.
~ :~
;
, ,
• .:;
:..~;,.
:,
;
~r~
17
77/=/NISVEt
SNOIIOgS-SSOI:IO 0 N 9 9 9 7 oo m IoAIuoo flaA/1
uol.sJaAul aAfl!SOcl lined le~.~ON
~
o
sa6p!ld aleqS
I i
lsnJql
~
auoz ieuo!ssa~dtuoo uJaq;no S
I
79Nb'HO
sq6!H / s~4o7 u,se~#edV-llnct
I:. . . . . . . I
S~lOOJo!qd~ow~lal41 snouoq~qoollV
0
'"
I
'" ii'1
qDH
9NN seue.ue_L o!qdJoLuelelAI
,,
V,,(bJd) l f '
q6!H ['~o,o~r'
• JlnE)
I!OWOO
I IA
H
r
qDH
0
'
opePlOS" A/IS
AI
,MSS I 9N
pla!:lo~SaleUJe~~o ~°
Snm41n7oo
or) m UJ~l 13 I-
0
q6!a
qDH
N ,linE) qB!H
o
opeplos
0
I,Ol.UOC]([°'d)r([o,d) I
)4
"
!11
-"
q6!H eleid-eued~ueo ---I
~ 0
"
0 N
('[oJd) :t
Ilia
edo9
O S 9N
aS
IIA
oo m
o
II II
1,
O3
.- .... . - ~ - ~ . .
•
i. 0
I
~
""
"
m
i..._Qo ,
0
q6!H eueduJeo
IIA
d~Ml N O U V O 0 7
IliA
SNOI.I.O_=,S dIG
e6p!EI eleqs selewepecl
STRUCTURE OF THE GULF OF PARIA PULL-APART BASIN the south. Normal faults commonly sole out into a Miocene shaly section but locally cut into Cretaceous evaporite (Couva Formation). The Couva evaporites constitute the core of the anticline located underneath the Avocado-Couva High. This high coincides with a change in vergence of the overlying extensional system. This lithologic unit has been also recognized onshore Trinidad and the Paria Peninsula. From a stratigraphic point of view, the most interesting feature provided by this section, is the lateral facies change between the conglomerate of the Cunapo Formation and sandstone and shale of the Manzanilla and Springvale Formations. Normal faulting and the lack of additional well data obscure the geometry of this contact. Section VII This E-W-trending profile across the entire Gulf of Paria is located in the central part of the study area. The section illustrates the lateral ramps of the extensional system and some accommodation zones associated with vergence changes. Eastwardand westward-dipping listric normal faults are connected and combined, constituting extensional basin fill with Upper Miocene-Pliocene strata. Notice significant rotation but no growth of strata. In the central part of the section is the Domoil High, where Pleistocene sediments unconformably overlie Lower Cretaceous rocks. The different stratigraphy of the wells M, I and J along with the seismic expression suggest thrusting within the Domoil High. West of the Domoil High, Upper Miocene sediments unconformably overlie Upper Cretaceous marls, as revealed by well I. Section VIII This E - W to N E - S W section shows the E W-trending Campana High. The easternmost part of the transect shows a broad fold linked with a deep decollement level. Data from well A suggest a Lower Cretaceous decollement system. The structure of the high is characterized by a set of folds involving Miocene to Upper Cretaceous sediments. Thrusting is evidenced by the presence of two repeated sections in well B. Those thrusts are probably lateral ramps of a northward-vergent thrust system. Transtensional basins bound to the north and to the south by the Campana High. The relationship between wells A and B indicates that well A is located on the down-thrown block of a westward-dipping extensional fault. The hangingwall of this extensional detachment is characterized by domino-type normal faulting. East of well B, a synclinal feature located underneath Pleistocene strata may suggest the presence of a relict compressional basin of the underlying fold-and-thrust belt.
487
Structural domains
The structure of the Gulf of Paria can be subdivided into two main structural provinces or domains: a Northern Transtensional Basin and a Southern Compressional Zone. Both units are superimposed on an underlying fold-and-thrust belt that represents the eastern prolongation of the Serranfa del Interior of eastern Venezuela. Regional surface data in eastern Venezuela and Trinidad as well as seismic and well log data presented in this study support southward thrusting of this underlying orogenic belt. The description of the Gulf of Paria will proceed from north to south. Northern Transtensional Basin This northern domain is bounded to the north by the Avocado-Couva High, the Gulf High and the Paria Peninsula to the north and by the Warm Springs fault to the south. The structure of this province is characterized by down-to-the-basin normal faults (Figs. 5 and 6). In the northernmost part of the Gulf of Paria, compressional structures, mainly south-directed thrust and folds, can be recognized underneath the extensional features. Thus, local extensional faults overprint previous compressional structures (Fig. 6). The Northern Transtensional Basin is characterized by a thick late Neogene section (Upper Miocene-Pleistocene) and a variable, but often incomplete, pre-Neogene section. The pre-basin-fill stratigraphy consists of Oligocene-Middle Miocene deep-water shales (i.e. the Brasso, Cipero and Lower La Pica Formations) unconformably overlain by the shallow-water Manzanilla, Springvale and Upper La Pica Formations. Several wells pierced the shale ridges (shale-cored footwall blocks) revealing the presence of deep-water shales and turbidites of the Cipero and Brasso Formations within the core. Structurally low areas, represented by extensional half-grabens, preserved more pre-extensional stratigraphic section (below the top of the Upper Miocene, 10.5 Ma) reaching the Oligocene and locally the Paleogene or Cretaceous. Structural highs (i.e. Domoil, Gulf, and Posa Highs) did not preserve any Upper Cretaceous to Middle Miocene sediments (see sections II, III, IV, and VII of Fig. 6). In the Domoil area Pliocene sediments were also not preserved (sections III, IV and VII of Fig. 6) and Pleistocene and Upper Miocene sediments unconformably overlie Lower Cretaceous rocks. Wells M and I located on the edge of the Domoil High (see Fig. 5) indicate the presence of Cunapo Conglomerate associated with fault scarps. The NNW-SSE-trending Avocado-Couva High constitutes the most prominent structural feature in the northern area. The hinge zone of this anticline
488 coincides with a change in the dip of the overlying extensional detachment and related normal faults. According to well and seismic data, the extensional decollement coincides with the Cretaceous-Neogene contact. Obvious thrust structures are not seismically clear south of the Avocado-Couva and Gulf Highs, since they are crosscut and obscured by extensional faults. The main decollement level of the thrust system is probably the Lower Cretaceous Couva evaporite and Cuche shale (well exposed in Trinidad). The northernmost part of the Northern Transtensional Basin (i.e. the Avocado-Couva and Gulf Highs) has a thin and very incomplete stratigraphic section. The Cunapo Conglomerate unconformably overlies Lower Cretaceous shales (Cuche Formation). A thick section of Lower Cretaceous evaporite (Couva Formation) is only present in this area and in the Paria Peninsula. The Cunapo Conglomerate represents fan deltas and slope deposits associated with fault scarps and were predominantly sourced from the north and are coeval with uplift of the Northern Ranges and the Paria Peninsula. The wells N, P and Q were drilled through these stratigraphic units. To the south, the conglomerate of the Cunapo Formation interfingers with the Manzanilla and Springvale Formations, suggesting a Late Miocene to Pliocene age for the conglomerate. The southern boundary of the Avocado-Couva and Gulf Highs is a southward-dipping system of anastomosing normal faults (Figs. 5 and 6). The Gulf High plunges to the west and disappears. On the Venezuelan side of the Gulf of Paria, a major basin-bounding fault (Casanay fault) separates the rocks of the Paria Peninsula from the Northern Transtensional Basin. The Casanay fault is a major normal fault with minor Pleistocene positive inversion and probably some strike-slip component, which is difficult to estimate (Fig. 6). The southern boundary of the Northern Transtensional Basin is the NE-SW-trending Warm Springs fault (WSF) which accounts for an important offset of the late Neogene and Cretaceous-Paleogene section (Fig. 6). The WSF does not consist of a single fault but of a set of north-dipping normal faults and related lateral ramps. We assume, based on regional data, that the WSF has been active since Late Miocene time. The structure of the transtensional basin is characterized by E-Wand NW-SE-trending normal faults (Fig. 5). Northsouth-trending lateral and oblique ramps that join the Warm Springs fault are observed on the Trinidad side of the Gulf of Paria. The Northern Transtensional Basin (NTB) is tilted to the south against the Warm Springs fault, which constitutes the master fault of the transtensional system. Roll-overs associated with basin-bounding faults are the most conspicuous structures of this domain.
J.E FLINCH et al. The main decollement of the extensional system is located within Oligocene-Lower Miocene shales (i.e. Cipero and Brasso Formations). The southeastern part of the NTB (Trinidad side) is occupied by a complex pattern of NE-SW-trending shale ridges that core footwall blocks. Well L (see Fig. 6) penetrated a shale ridge in this area (see sections IV and V of Fig. 6). Two main WNW-ESE-trending structural highs are located within the Northern Transtensional Basin: the Domoil-Gupe and GopaPosa Highs (Fig. 5). These structural highs are bounded by down-to-the-basin normal faults and are cored by Lower Cretaceous rocks. Anticlines developed against the main normal faults and the thickness distribution of the Pleistocene section suggest Pleistocene positive inversion. Basin-bounding normal faults were reactivated as reverse faults.
Southern Compressional Zone The Southern Compressional Zone (SCZ) is located south of the Warm Springs fault system. This domain is characterized by a thin Plio-Pleistocene stratigraphic section and predominantly north-vergent Late Miocene to Pliocene thrusting (Figs. 5 and 6). The Southern Compressional Zone represents the continuation of the structures of the Central and Southern Ranges of Trinidad described by Kugler (1961). Although the dominant vergence of thrusts and folds is reported to be towards the south in the island of Trinidad, north-vergent back-thrusts can be recognized in the Southern Range of Trinidad (Fig. 5). The stratigraphic section of the Southern Compressional Zone consists of basinal Upper Miocene deep-water shale and turbiditic sandstone (Lower La Pica Formation) overlain by a very thin PlioPleistocene section. The Manzanilla Formation is absent in the east on the San Fernando High (section V of Fig. 6). The overlying Springvale and Talparo Formations are affected by folding and thrusting. The pre-Neogene section is better known in the south because of the lack of thick superimposed transtensional basins. Upper Cretaceous sediments, absent in the Northern Transtensional Basin, were encountered in wells located in the San Fernando Bay and the Soldado High (see Fig. 6). The seismic expression of the Los Bajos fault is that of a continuous NW-SE-trending shale ridge, only interrupted by the WSF (Fig. 5). In the Morro area, where the Posa field is located, the regional decollement may be located above the Upper Miocene unconformity (10.5 Ma sequence boundary) (see section II of Fig. 6). This area consists of arcuate-shaped approximately E-W-trending and south-vergent thrusts that constitute a thrust-imbricate system (Fig. 5). The NE-SW-trending piggyback basins related to late Neogene thrusting are
STRUCTURE OF THE GULF OF PARIA PULL-APART BASIN characterized by the merging of unconformities and the presence of progressive unconformities. Progressive unconformities consist of folded unconformities in the lower part of the basin fill that are crosscut by less steep unconformities in the upper part of the section. This type of unconformity records the timing of thrust emplacement. Shale-cored doubly verging anticlines associated with thrusting are present in the Pedernales area and extend to Trinidad; mud volcanoes are common along this structural zone. One of the most prominent structural features in this area is the Campana High (Fig. 5). Wells A and B were drilled on structural highs constituted mostly by Upper Cretaceous and Paleogene sediments (section VIII of Fig. 6). This NE-SW-trending anticlinorium is interpreted to be related to a deep-seated thrust rooted in the Cretaceous section. Nevertheless, seismic data are not conclusive regarding the vergence of the underlying thrust system. The seismic data display N-S-trending lateral ramps associated with north-vergent thrusts south of the Campana High (see section VIII of Fig. 6). In the Southern Compressional Zone, transtensional structures are subordinated to thrusts, but locally thrust sheets are overprinted by steeply dipping normal faults. Westward-dipping normal faults offset the western edge of the Campana High (see section VIII of Fig. 6).
Implications for the structure of Trinidad The knowledge of the structure of the Gulf of Paria has important implications for the geology of the island of Trinidad. The Caroni Basin is the onshore continuation of the transtensional basins of the Gulf of Paria. The Warm Springs fault constitutes the southern boundary of this transtensional province. This apparent normal contact has been subsequently modified by compression, like most of the major normal faults of this northern domain. The northern flank of the Central Range of Trinidad constitutes the field expression of the positive inversion overcome by the Caroni Basin. Northward-dipping faults and strata, on the hangingwall of the Warm Springs fault System, opposite to what would be expected for a normal fault, reveal compressional reactivation after extension. The Guatapajaro Anticline, located in the center of the Caroni Basin (Fig. 5), is evidence of this late Neogene inversion that we recognized offshore, in the Gulf of Paria. The extensional character of the Casanay fault indicates to us that the nature of the E1 Pilar fault system as the northern boundary of the Northern Transtensional Basin should be reviewed. The seismic data, as well as most of the surface data along the southern boundary of the Northern Range suggest that the present-day contact between the allochthonous
489
metamorphic terranes and the folded sediments of the South American Plate is a northward-dipping extensional contact in the study area. Fission-track data in the Northern Range (Algar, 1995) suggest Late Miocene uplift, which was coeval with extensional tectonics in the Gulf of Paria and Caroni Basin. The Northern Range of Trinidad, as well as the Paria Peninsula of eastern Venezuela can be pictured as core complexes bounded by the extensional basins of the Caribbean Sea to the north and by the transtensional basin of the Gulf of Paria in the south. The Central and Southern Ranges of Trinidad belong to the Southern Compressional Zone (Fig. 5). This structural trend extends towards the west into the Gulf of Paria, and can be traced further west into the eastern Venezuelan Serranfa del Interior fold-and-thrust belt. This single folded belt involves the deformed northern passive-margin sediments of the South American Plate. The so-called Nariva fold-and-thrust belt of southern Trinidad represents, like the Carapita overpressure shale belt of eastern Venezuela, the westward continuation of the Barbados accretionary wedge (Fig. 1). The continuity between both is only interrupted by the Neogene transtensional features of the Gulf of Paria. The transtensional basins crosscut and are superimposed onto the underlying and otherwise continuous folded belt. This relationship not only applies to thrusts and folds but also to the strike-slip faults that bound thrust units, i.e. tear faults like the Los Bajos fault. This crosscutting relationship was observed by early exploration works (Wilson, 1968). The Southern Basin and the Erin Basin represent piggyback basins in the sense of Ori and Friend (1984) related to the emplacement of the Southern Range. These piggyback basins can be traced in the subsurface towards the west until the Campana High, where the Neogene sediments onlap onto Cretaceous-Paleocene rocks. They are equivalent to the piggyback basins of the Carapita Formation in eastern Venezuela. Mud volcanoes are associated with thrust faults of this southern tectonic province (Fig. 5). Careful analysis of the Southern Range geological maps reveals the presence of doubly vergent thrusts and folds (Fig. 5). In the southern part of the central Gulf of Paria the westward extension of this unit is represented by the Morro northward-vergent imbricates (backthrusts). From a geometric point of view the contact between south-vergent thrusts and north-vergent imbricates, can be interpreted as a 'passive-roof duplex' in the sense of Banks and Warburton (1986). This uppermost compressional system (backthrust zone) involves slope and basinal deposits. The Pedernales shale ridge located in the southern part of the SCZ (Fig. 1, Fig. 5, and section I of Fig. 6) constitutes an ENE-WSW-trending doubly verging shale-cored anticline linked the basal decollement
490
J.F. FLINCH et al.
Fig. 7. Block diagrams showing the Neogene structural evolution of the Gulf of Paria. (a) Fold-and-thrust belt stage characterized by south-vergent thrusting and related piggyback basins. (b) Transtension in the north coeval with north-vergent thrusting in the south. (c) Positive inversion of transtensional basins and north-vergent thrusting. Color version at http://www.elsevier.nl/locate/caribas/
STRUCTURE OF THE GULF OF PARIA PULL-APART BASIN
491
Fig. 7 continued. of the passive-roof duplex. The shale ridge extends from Maturfn in the eastern Venezuelan Basin to the Southern Range of Trinidad.
Tectonic evolution The data presented here permit establishment of an evolutionary model for the Gulf of Paria. The relative timing of the structures is reasonably well known in the area, even though the absolute timing of the main tectonic events is not fully determined. The Neogene tectonic evolution of the Gulf of Paria can be subdivided into three main stages (Fig. 7): (1) fold-and-thrust belt (south-directed thrusting); (2) transtension and north-directed thrusting; (3) positive inversion. Table 1 summarizes the timing and nature of the main tectonic events within each structural domain.
Fold and thrust belt (Oligocene-Lower Miocene) Before Late Miocene time the Gulf of Paria was occupied by a peneplained (eroded) fold-and-thrust belt. The block diagram of Fig. 7a displays the hypothetical geometry of the fold-and-thrust belt that occupied the region before transtension took place.
The structures underlying the Gulf of Paria represent the offshore continuation of the Central Range of Trinidad and the eastern Venezuelan fold belt. The Lower Cretaceous Couva evaporites (AvocadoCouva High) may represent part of a much more widespread evaporitic nappe that could occupy most of the basement of the Gulf of Paria. In fact, these evaporites could represent Triassic or Jurassic evaporites emplaced within Lower Cretaceous deep-water sediments like in the Gulf of Mexico. Broad synclinoria are located south of the thrust that separates Lower Cretaceous from Tertiary sediments. Brasso shales and Nariva turbidites occupy the core of these synclinoria. The westward continuation of these synclinoria is occupied by the Carapita Formation of eastern Venezuela. We interpret these synclinoria as relict piggyback basins of the eastern Venezuelan fold-and-thrust belt. The presence of south-vergent thrusts underneath the basin fill can be inferred from seismic data (Fig. 6). The Campana High was developed before 10.5 Ma (initiation of half-graben tectonics) and is probably associated with south-vergent thrusting (sections I, II and VIII of Fig. 6). A well-developed shelf-margin occurs in the southernmost part of the Gulf of Paria (i.e. Erin Basin, Pedernales and Posa
492 areas) and represents the northern edge of the eastern Maturfn foreland basin.
Transtension and northward thrusting (Middle Miocene-Pliocene) From Late Miocene time on, half-grabens overprinted the underlying thrust sheets in the northern part of the Gulf of Paria. This tectonic stage was characterized by the development of transtensional basins controlled by steeply dipping normal faults (Fig. 7b). As transtensional collapse progressed, the relief of lows and highs increased. Shale ridges form shale-cored footwall blocks located in the central part of the transtensional basin. The main structural feature active during this time was the Warm Springs fault that separates the northern transtensional province from the southern compressional province. The Casanay and Arima faults constitute the boundary between the NTB with the allochthonous metamorphic terranes exposed along the Paria Peninsula and the Northern Range of Trinidad. Extensional collapse induced by pull-apart extended from the Caroni Basin of Trinidad to the Guanoco area and the San Juan graben in eastern Venezuela (see Fig. 1). Fission-track ages in northern Trinidad (Algar, 1995; D.A. Farrell, pers. commun., 1996) suggest that extensional collapse in the Gulf of Paria was coeval with the uplift of Trinidad's Northern Range. High tectonic subsidence rates were compensated by huge sediment supply, resulting in shallow-water Upper Miocene-Pliocene deposition (i.e. the Manzanilla and Springvale Formations). Fan deltas, represented by the Cunapo Conglomerate, were shed from the uplifted Northern Range, Paria Peninsula and Domoil, Gulf and Avocado-Couva Highs. During Pliocene time, turbiditic wedges (developed in front of pre-existing south-vergent thrusts) were thrusted and folded by north-vergent imbricates (Fig. 7b). Turbidites were encountered by drilling in the Pedernales and Posa oil fields (Fig. 5). Piggyback basins developed associated with this north-vergent thrusting. Positive inversion (Pleistocene) During Late Pliocene-Early Pleistocene time, inversion of pre-existing normal faults took place in the northern area (Fig. 7c). Normal faults were reactivated with reverse displacement and anticlinal features developed in the hangingwall of the major pre-existing normal faults (see sections III, IV and VI of Fig. 6). The orientation of the inversion structures is predominantly NW-SE to E-W coinciding with the trend of major extensional faults (Fig. 5). N-S-trending normal faults were not inverted, suggesting that inversion is mainly related to a N-S-directed compression. Inversion structures
J.E FLINCH et al. are widespread within the Caroni Basin of Trinidad and can be recognized in surface exposures along the northern border of the Central Range. North-vergent thrusting was still active in the south (Pedernales area) until Late Pleistocene-Holocene time.
DISCUSSION A valid model for the Gulf of Paria should explain: (1) the orientation of the Domoil, Gulf and Posa Highs and the nature of their border faults; (2) the fact that the Los Bajos fault is crosscut and overprinted by the Warm Springs fault system, and therefore cannot be carried through the entire Gulf of Paria; (3) that transtension in the northern Gulf of Paria (half-graben development) was coeval with north-vergent thrusting in the south, that is, extension took place in an overall compressional scenario; (4) the Pleistocene positive inversion of major NWSE- to E-W-trending bounding faults, suggesting widespread post-Pliocene N-S compression.
The pull-apart model From a geometric point of view fault patterns in the study area resemble pull-apart basins generated by experimental models (McClay and Dooley, 1995) (Fig. 8). Following a comparison with an experimental pull-apart model, the E1 Pilar fault represents the westward side of the Principal Displacement Zone (PDF), characterized by fight-lateral strikeslip displacement (Fig. 8). Steeply dipping normal faults such as the ones that bound the Domoil, Posa and Gulf Highs would represent oblique-slip faults with normal and fight-lateral component (crossbasin fault zones in the sense of McClay and Dooley, 1995). The lack of reference lithologic or structural markers makes it difficult to demonstrate the magnitude of transcurrent displacement on most of these faults. The northern border of the pull-apart system is represented by the Casanay and Arima faults 0, while the southern border is the Warm Springs fault (Southern Basin sidewall). The eastward side of the strike-slip system would be represented by the Trinidad Central Range. The so-called Central Ranges fault may represent the eastward PDZ of the presumed pull-apart system that generated the Gulf of Paria (Fig. 8). The pull-apart model also explains the eastward narrowing of the Caroni Basin and the orientation of the San Juan graben of eastern Venezuela, as well as the normal faulting along the eastern Venezuelan Serranfa del Interior. In our model the Los Bajos fault does not have the same significance as in the other models where it plays the role of a major strike-slip fault across the entire Gulf of Paria (Munro and Smith, 1985; Algar, 1995).
STRUCTURE OF THE GULF OF PARIA PULL-APART BASIN
493
Fig. 8. Comparison of the structure of the Gulf of Paria (lower panel) with the pull-apart experimental analogue model of McClay and Dooley (1995) (upper panel). The westward PDZ is represented by the E1 Pilar fault while the eastward PDZ is the Central Range fault of Trinidad. The Southern Basin sidewall is the Warm Springs fault and the northern the Arima-Casanay fault. Cross-basin fault zones define the Domoil, Posa, Gulf and Avocado-Couva Highs.
The Los Bajos fault, like the Urica fault or San Francisco-Quiriquire faults (Fig. 1) is not related to the pull-apart system of the Gulf of Paria. They are tear faults related to the differential thrust transport displacement between thrust sheets of the eastern Venezuelan fold-and-thrust belt (Fig. 1). The pullapart of the Gulf of Paria can be explained by stepping over between the E1 Pilar fault system and the Central Range fault of Trinidad (Fig. 8). The nature of this releasing step-over could be an ancient tear fault like the ones of the Serranfa del Interior. Epicenters of earthquakes and related focal mechanisms indicate a change from a single E - W strike-slip fault (El Pilar fault) to a distributed zone of deformation along NW-SE-trending strike-slip faults (P6rez and Aggarwal, 1981; Villasefior et al., 1992). High subsidence rates and parallel filling (no growth) with clastic wedges (fan deltas) associated with fault scarps (Cunapo Conglomerate) are typical characteristics of pull-apart basins (Link et al., 1985). The presence of south-vergent thrust sheets, underlying the present-day basin, implies additional
complication to simple pull-apart generated above non-deformed strata and differs from a conventional basement-involved pull-apart basin. Pull-apart in the Gulf of Paria overprinted an underlying fold-andthrust belt.
CONCLUSIONS The Gulf of Paria is characterized by NW-SEto E-W-trending transtensional basins. These basins developed during late Neogene time in the northern part of the Gulf of Paria, and were coeval with compression in the southern part. This permits the definition of two tectonic provinces: the Northern Transtensional Basin, dominated by a complex set of half-grabens, and the Southern Compressional Zone, consisting of doubly vergent thrust imbricates. The Caroni Basin is the eastward continuation of the Northern Transtensional Basin, while the Central Range and Southern Ranges fall into the Southern Compressional Zone. The extensional structures of
494 the Gulf of Paria extend towards the west and are superimposed to Cretaceous thrust sheets of the eastern Venezuelan Serranfa del Interior. The structural evolution of the Gulf of Paria, can be summarized into three main stages: (1) Fold-and-thrust belt (south-vergent thrusting). (2) Transtension and northward thrusting. (3) Positive inversion. The structure of the Gulf of Paria can be interpreted as a fight-lateral pull-apart caused by the step-over of the E1 Pilar and the Central Range faults.
ACKNOWLEDGEMENTS
We thank CVP (Corporaci6n Venezolana de Petroleo), PETROTRIN and LAGOVEN (today PDVSA Oil and Gas) for permission to use the data presented here. Thanks also go to V. Hunter, who provided biostratigraphic support. The article has benefitted from the suggestions of Albert W. Bally and Max Furrer. We want also to thank Kevin Burke, Paul Mann and Kris Meisling for helpful and constructive reviews.
REFERENCES
Algar, S., 1995. Interaction of the Caribbean and South American Plates as revealed in the Northern Range of Trinidad. Field Trip Guide, 14th Caribbean Geological Conference. Port of Spain, Trinidad, pp. 38-60. Ball, M.M., Harrison, C.G.A., Supko, ER., Bock, W.D. and Maloney, N.J., 1971. Marine geophysical measurement on the southern boundary of the Caribbean Sea. In: T.W. Donnelly, (Editor), Caribbean Geophysical, Tectonic and Petrological Studies. Geol. Soc. Am. Mem., 130, 224 pp. Banks, C.J. and Warburton, J., 1986. 'Passive-roof' duplex geometry in the frontal structures of the Kirthar and Sulaiman mountain belts, Pakistan. J. Struct. Geol., 8: 229-237. Bellizzia, A. and Dengo, G., 1990. The Caribbean mountain system, northern South America. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, H, Geological Society of America, Boulder, CO, pp. 167-175. Burke, K., 1988. Tectonic evolution of the Caribbean. Annu. Rev. Earth Planet. Sci., 16:201-230. Burke, K., Cooper, C., Dewey, J.E, Mann, E and Pindell, J.L., 1984. Caribbean tectonics and relative plate motions. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162:31-63. Case, J.E., MacDonald, W.D. and Fox, EJ., 1990. Caribbean crustal provinces; seismic and gravity evidence. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, H, Geological Society of America, Boulder, CO, pp. 5-36. Di Croce, J., 1995. Eastern Venezuela Basin: Sequence Stratigraphy and Structural Evolution. Unpublished Ph.D. Thesis, Rice University, Houston, TX, 225 pp.
J.E F L I N C H et al. Flinch, J.E and Rambaran, V., 1996. Neogene extensional tectonics of the Gulf of Paria (eastern Venezuela-western Trinidad). Am. Assoc. Pet. Geol. Bull., 80 (8): 1290. Gonzalez de Juana, C., Iturralde, J. and Picard, X., 1980. Geologia de Venezuela y de sus cuencas petrolfferas. Ediciones Foninves, Caracas, 1031 pp. Kugler, H.G., 1953. Jurassic to recent sedimentary environments in Trinidad. Bull. Assoc. Suisse Geol. Ing. Pet., 20: 27-60. Kugler, H.G., 1961. Geological Map of Trinidad, scale 1:100,000. Edited by Orell Fussli Arts Graphiques S.A. Zurich, copyright of The Petroleum Association of Trinidad. Link, M.H., Roberts, M.T. and Newton, M.S., 1985. Walter Lake Basin, Nevada: an example of Late Tertiary (?) to Recent sedimentation in a basin adjacent to an active strike-slip fault. Soc. Econ. Paleontol. Mineral. Spec. Publ., 37: 105-125. Macellari, C., 1995. Cenozoic sedimentation and tectonics of the southwestern Caribbean pull-apart basin, Venezuela and Colombia. In: A.J. Tankard, R. Suarez-Soruco and H.J. Welsink (Editors), Petroleum Basins of South America. Am. Assoc. Pet. Geol. Mem., 62: 757-780. Mann, E and Burke, K., 1984. Neotectonics of the Caribbean. Rev. Geophys. Space Phys., 22 (4): 309-362. McClay, K. and Dooley, T., 1995. Analogue models of pull-apart basins. Geology, 23:711-714. Molnar, E and Sykes, L.R., 1969. Tectonics of the Caribbean and Middle America regions from focal mechanisms and seismicity. Geol. Soc. Am. Bull., 80: 1629-1684. Munro, S.E. and Smith, ED., Jr., 1985. The Urica fault zone, northeastern Venezuela. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162: 213-215. Ori, G.G. and Friend, EE, 1984. Sedimentary basins, formed and carried piggy-back on active thrust sheets. Geology, 12: 475-478. P6rez, O. and Aggarwal, Y.E, 1981. Present-day tectonics of the southeastern Caribbean and northeastern Venezuela. J. Geophys. Res., 86 (Bll): 10,791-10,804. Robertson, E and Burke, K., 1989. Evolution of southern Caribbean plate boundary, vicinity of Trinidad and Tobago. Am. Assoc. Pet. Geol. Bull., 73 (4): 490-509. Salvador, A. and Stainforth, R.M., 1968. Clues in Venezuela to the Geology of Trinidad, and vice versa. 4th Caribbean Geol. Conference Transactions, pp. 31-40. Schubert, C., 1982. Origin of Cariaco Basin, Southern Caribbean Sea. Mar. Geol., 47: 345-360. Soulas, J.E, 1986. Neotect6nica y tect6nica activa en Venezuela y regiones vecinas. VI Congreso Geol6gico Venezolano, Tomo X, pp. 6639-6656. Speed, R.C., 1985. Cenozoic collision of the Lesser Antilles and Continental South America and the origin of E1 Pilar fault. Tectonics, 4:41-69. Vierbuchen, R.C., 1984. The geology of the E1 Pilar fault zone and adjacent areas in northeastern Venezuela. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The CaribbeanSouth American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162: 189-212. Villasefior, A., Mufioz, M.I., Franke, M. and Gajardo, E., 1992. Estudios de microsismicidad en el norte de Venezuela, 3. Zona nororiental. VI Congreso Venezolano de Geoffsica, pp. 529535. Wilson, C.C., 1968. The Los Bajos fault. Transactions of the 4th Caribbean Geological Conference, Port of Spain, Trinidad, pp. 87-90.
Chapter 18
Structural and Sedimentary Development of a Neogene Transpressional Plate Boundary between the Caribbean and South America Plates in Trinidad and the Gulf of Paria
STEPHEN
BABB and PAUL MANN
Trinidad, the Gulf of Paria, and Eastern Venezuela lie in a diffuse and complex zone of Cenozoic tectonic interaction between the Caribbean and South America plates. Numerous models have being proposed to explain the complex tectono-stratigraphic evolution of the area. In this paper, we interpret an integrated data base consisting of well logs and seismic reflection profiles to document five sequences of the Gulf of Paria-Northern basin of Trinidad. Sequences 1 and 2 consist of a Late Jurassic-early Valanginian carbonate megaplatform with overlying carbonate bank buildups that together form the basement to overlying siliciclastic sedimentary sequences 3 through 5 of mainly Neogene age. Sequence 3 is a Late Miocene-Early Pliocene shallow marine to brackish-water conglomerate and sandstone that represents a southward-fining and thinning and eastward-thickening siliciclastic wedge derived from the Late Miocene uplift and erosion of Trinidad's Northern Range and records early activity along the E1 Pilar strike-slip fault at the mountain front of that range. Sequence 3 and overlying sequences 4 and 5 fill in topographic relief created by the growth of the carbonate banks of sequences 1 and 2 and by space created by tectonic extension in the Gulf of Paria and Northern basin. Sequence 4 is an Early to Middle Pliocene inner neritic to shallow marine conglomerate, sandstone, silt, and clay that represents a northward-fining and thinning siliciclastic wedge derived primarily from the Late Miocene-Early Pliocene uplift and erosion of the Central Range with some input from the northwest and north. Sequence 5 is a Late Pliocene to Pleistocene marine to brackish-water sand, silt, clay and minor conglomerate that represents continued siliciclastic deposition in an increasingly restricted basin between the uplifted Northern and Central Ranges. Comparison of two-way travel time structural contour maps of each boundary between sequences 3-5 (totaling 2-3 km in thickness) with two-way travel time isochron maps of the seismic sequence immediately overlying each structural surface allows a better visualization of the space available at tectonically significant times in the history of the Gulf of Paria-Northern basin and the subsequent sedimentary infill onto that structural surface and sequence stratigraphic boundary. This comparison allows the identification of three Neogene deformational phases that have affected the area and closely control the proposed sequence stratigraphic boundaries: Phase one, Late Miocene-Early Pliocene strike-slip motion along the E1 Pilar fault and north-to-south filling of the Gulf of Paria and Northern basins; Phase two, Middle to Late Pliocene strike-slip motion along the Warm Springs-Central Range fault zone and south-to-north filling of the Gulf of Paria and Northern basins; and Phase three, Late Pliocene to Pleistocene strike-slip motion along the Warm Springs-Central Range fault zone and continued filling of the Gulf of Paria and Northern basins. Paleocurrents and environments of deposition of the Gulf of Paria and Northern basins have been closely controlled by structural events during the three deformational phases. Paleocurrents for deformational phase one (deposition of sequence 3) are oriented orthogonal (north-south) and parallel (east-west) to the axis of the basins and reflect strong control of the bounding E1 Pilar fault zone along the northern edge of the basin. Paleocurrents for deformational phases two (sequence 4) and three (sequence 5) become increasingly dominated by a west-to-east paleoflow along with a southeast-to-northwest flow in the area of the Goodrich sub-basin (southern Gulf of Paria basin). Environments of deposition reflect the Late Miocene (~ 10 Ma) uplift and erosion of the Northern Range during phase one followed by subsequent uplift and erosion of the Central Range and the gradual decrease in the importance of the Northern Range siliciclastic source area during phases two and three. Environments generally remain proximal, shallow, and brackish along the E1 Pilar fault zone on the northern edge of the basin and distal, deeper, and less restricted along the central and southern edges of the basin. Through time the basin has become increasingly brackish and more restricted. Seismic lines are used to illustrate the variation in the structural style and geomorphic expression of the E1 Pilar fault zone along the northern edges of the Gulf of Paria-Northern basin. In the western Gulf of Paria, the fault exhibits a transpressional structure consistent with its east-northeast strike. A slight change in strike to a more east-southeast strike in the eastern Gulf of
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 495-557. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
496
S. B A B B and E M A N N
Paria is consistent with the more transtensional structure of the fault and its association with the Puerto Grande sub-basin in the northern Gulf of Paria basin. The trace of the E1 Pilar fault zone is poorly expressed onland in Trinidad and may reflect its waning activity in recent time in an eastward direction and/or cultural obliteration of its tectonic geomorphology in this highly urbanized part of Trinidad. Seismic lines are used to illustrate the structural style of previously undescribed oblique-slip faults in the Gulf of Paria. These styles include: (1) gravity-related detachment faults possibly localized along the unconformity between Cretaceous carbonate and overlying siliciclastic rocks; and (2) transtensional faults formed in the stepover area between the E1 Pilar fault to the north and the Warm Springs fault to the south. Seismic lines are used to illustrate the previously undescribed structural styles and geomorphic expression of the Warm Springs-Central Range-Caigual fault zone. In the Gulf of Paria, the Warm Springs fault zone exhibits a transtensional structure consistent with its east-southeast strike. An abrupt change in strike to a more east-northwest strike near the western shoreline of Trinidad is consistent with the more transpressional structure of the continuation of the Warm Springs fault zone in the Central Range (Central Range-Caigual fault zones). We reinterpret the overall structure of the Central Range as a transpressional uplift bounded by inwardly dipping reverse faults and bisected by the Central Range-Caigual strike-slip fault system rather than as the southeast-verging fold-thrust belt proposed by previous workers. Integration of data by previous workers in the Southern basin of Trinidad allows us to make a comparison between the style and age of deformation in the Southern basin and that described by us in the Gulf of Paria-Northern basin. The style of deformation in the Southern basin is predominantly south-vergent and older than the deformation in the Gulf of Paria and Northern basins because Middle Miocene (~11.4 m.y. horizon) units onlap folded and faulted older Middle Miocene units. The unconformity between the Late Miocene-Early Pliocene sequence 3 and the middle to Late Pliocene sequence 4 in the Northern-Gulf of Paria basins correlates with a less prominent unconformity between the Upper Cruse and Lower Forest Formations of southern Trinidad. The deformation in the Southern basin is therefore older than observed in the Gulf of Paria-Northern basin. Compilation of regional geologic data and integration of the Southern basin data with our results from the Gulf of PariaNorthern basins allows the following events to be constrained in the Middle Miocene through recent evolution of the Trinidad region. Middle Miocene ('.~10 Ma): high-angle faulting affects the Southern basin; we do not regard this event as a regional fold and thrusting event because we do not observe widespread fold-thrust sturctures in the pre-Late Miocene rocks of the Gulf of Paria-Northern basin. Late Middle Miocene: uplift of the Central Range at this time may be related to the Middle Miocene thrusting observed in the Southern basin. Early Late Miocene: during this period a lobe of conglomeratic sediment was shed from the Northern Range into the Gulf of Paria-Northern basin and is interpreted to reflect the first significant lateral movement along the right-lateral E1 Pilar fault zone. We propose that the fault was propagating from west to east during this time as the accretionary wedge and forearc of the Lesser Antilles arc moved past the South American passive margin. The Los Bajos fault is thought to have propagated southeastward at this time. Termination of that strike-slip fault on the Southern Range may have led to a widening of the Neogene belt of transpressional thrusting during this time. Late Miocene-Early Pliocene: during this time, there was a rapid spread of strike-slip and oblique-slip faulting to the south of the E1 Pilar fault zone along the Warm Springs fault zone and within the Goodrich sub-basin. This southward shift is interpreted as recording the development of a pull-apart basin at the stepover area between the E1 Pilar and Warm Springs fault zones. Formation of the stepover may have deactivated slip along the eastern continuation of the E1 Pilar fault zone at the northern margin of the Northern basin. Early to middle Pliocene: this period saw continued development of the subsidence and faults of the previous period in the stepover area between the E1 Pilar and Warm Springs faults. Late Pliocene to Pleistocene: a major wedge of sediment derived from erosion of the Central Range and sources to the north and northwest filled the Goodrich sub-basin and continued southeastward propagation of the Los Bajos fault offset fold axes in the Southern basin by 10.5 km. The E1 Pilar fault zone appeared to undergo transpressional reactivation in the Northern basin but most offset seems to have shifted to the Warm Springs-Central Range-Caigual fault zone in the Central Range. We compare this sequence of events and the structural development of Trinidad with the events and structure of the Eastern Venezuelan basin in order to infer large-scale tectonic controls on the broad zone of deformation between the South America and Caribbean plates. The Eastern Venezuelan basin to the west and along-strike of Trinidad is a classic foreland basin associated with an adjacent, southeast-verging fold-thrust belt (Serranfa del Interior) formed by ~ 4 5 - 9 0 km of shortening. Trinidad exhibits a slightly younger record of fault-related sedimentation, lacks evidence for significant shortening of the Cretaceous-Pleistocene sequences 1-5 documented in this paper, and exhibits transpressional-type structures with no uniform sense of vergence in the Northern Range, Gulf of Paria-Northern basin and Central Range. In order to explain these differences between the pattern of deformation in the Eastern Venezuelan basin and Trinidad, we propose that Precambrian crust of the Guyana shield adjacent to the Eastern Venezuelan basin acted as an effective backstop and did not allow continued southeastward migration of the obliquely colliding Caribbean crust. In contrast, Trinidad appears to occupy the site of a rifted passive margin whose thinner crust acted as less of a backstop than the Precambrian crust underlying the Eastern Venezuelan basin. This less confined 'free face' to the southeast of Trinidad allowed strike-slip movement of Caribbean and passive margin crust to the southeast in a direction roughly parallel to the trend of the former passive margin.
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY INTRODUCTION
Tectonic problems in Trinidad The island of Trinidad lies in a diffuse and c o m p l e x zone of N e o g e n e interaction b e t w e e n the Caribbean and South A m e r i c a plates (Fig. 1A). The geologic and tectonic evolution of the plate bound-
497
ary zone in Trinidad and adjacent eastern Venezuela has r e m a i n e d controversial and this controversy has led to n u m e r o u s p u b l i s h e d m o d e l s put forward to explain the c o m p l e x tectono-stratigraphic evolution of the area. T h e s e tectonic m o d e l s predict different late N e o g e n e tectonic and structural styles in the Trinidad area that include oblique collision and thrusting, right-lateral strike-slip faulting, and
Fig. 1. (A) Present-day plate structure of the Caribbean region. Directions and rates of plate motion relative to a fixed Caribbean plate are from DeMets et al. (1994) and Dixon et al. (1998). (B) Main tectonic elements of the Caribbean plate and its margins include a Late Cretaceous oceanic plateau beneath the Caribbean Sea and cropping out in uplifted areas at its margins; a Cretaceous-Recent island arc or 'Great Arc of the Caribbean'; Late Cretaceous-Eocene back-arc basins; arrows indicate inferred direction of opening in the Yucatfin and Grenada back-arc basins; and Late Jurassic-Cenozoic passive margins on rifted continental crust in northern South America, Chiapas and the Yucatfin Peninsula and the Bahama platform. Tectonic features of boxed area are shown in Fig. 2.
498 transpressional models involving both thrusting and strike-slip. The Paleogene history of the island, that is thought by most previous workers to represent a thermally subsiding passive margin, remains even less understood than the Neogene active margin history because of the strong late Neogene tectonic overprint on these older rocks and because of the limited surface exposure and well penetration of Paleogene sedimentary rocks. These conflicting tectonic models by previous workers in Trinidad reflect varying emphasis on methods, data types, and local study areas in Trinidad. For example, previous models have emphasized surface structural, stratigraphic and geochronologic data from older rocks exposed in the three ranges of Trinidad, subsurface well and seismic reflection data made available from oil companies exploring the region, and earthquake data collected from worldwide or local seismic networks. Surface geologic mapping has mainly focussed on the Northern Range which offers the most complete and best exposed section of Jurassic and Cretaceous age. Published industry data are most plentiful from the southern part of the island and its offshore area where most hydrocarbons have been found to date (Persad, 1984, 1985).
Objectives of this study This study was undertaken to conduct a detailed and systematic regional analysis of the geologic evolution of northern Trinidad within the Gulf of Paria and Northern basins. We used an integrated data base consisting of well logs and seismic reflection profiles. These data provide a new set of geologic constraints on the sedimentary and tectonic history of the area. The Neogene basins of northern Trinidad were chosen for investigation in this tectonically-oriented study for the following reasons: (1) these basins record the nature and timing of major tectonic events because they are adjacent to key tectonic features of Trinidad including the Northern Range, the E1 Pilar fault zone, the Warm Springs fault zone, the Central Range-Caigual fault zone, and the Central Range; (2) these basins are relatively undeformed and contain a fairly complete marine sedimentary record which can be accurately dated using marine microfauna and pollen; and (3) the basins cover a significant part of Trinidad and its offshore shelf areas and can be used to test regional structural, stratigraphic and tectonic predictions made by previous workers for the Trinidad region. The general focus of the study is to constrain the timing of tectonic activity in the basins of northern Trinidad with emphasis on dating the movement on late Cenozoic basin bounding faults and the
S. BABB and E MANN effects of tectonic activity on sedimentation and depositional processes. Specific sedimentary goals of the study include: (1) integration of seismic and well data to enhance and refine knowledge of the depositional systems and stratigraphic architecture; (2) determination of the structural style and age of major fault zones that act as important controls on basinal sedimentation; and (3) determination of tectonic controls on the formation of major sequence boundaries that separate the five identified sequences in the basins of northern Trinidad.
TECTONIC AND GEOLOGIC SETTING OF TRINIDAD
Plate tectonic setting The Caribbean plate is presently bounded to the east by the Lesser Antilles subduction zone, to the west by the Middle America subduction zone, to the north by left-lateral strike-slip faults, and to the south by fight-lateral strike-slip faults (Mann et al., 1990) (Fig. 1A). Rates of relative motion of the North and South America plates, relative to a fixed Caribbean plate, range from about 1 to 2 cm/yr in a general east-west direction (DeMets et al., 1994; Dixon et al., 1998). The crust of the Caribbean plate is made up of three main components: (1) The Caribbean oceanic plateau formed about 88 Ma ago probably as the result of the ascent of a mantle plume into oceanic crust of Jurassic to Early Cretaceous age (Burke, 1988; Diebold and Driscoll, Chapter 19; Driscoll and Diebold, Chapter 20). The plateau forms a large area of the stable central part of the plate beneath the Caribbean Sea (Fig. 1B). (2) Passive margins formed above rifted crust of Precambrian to Paleozoic age in the ChiapasYucatfin area of southern Mexico, the Bahamas platform near Cuba and Hispaniola, and the northern margin of South America in Colombia, Venezuela, Trinidad and Guyana (Fig. 1B). These margins reached their approximate present-day positions following the breakup of North and South America during Late Jurassic and Early Cretaceous times (Burke, 1988; Pindell and Barrett, 1990). During the Early to Late Cretaceous, these margins subsided thermally and became the sites of both carbonate and siliciclastic sedimentation. (3) The Great Arc of the Caribbean of Burke (1988) ranges in age from Early Cretaceous to Recent in the northern, eastern and southern quadrants of the circum-Caribbean (Fig. 1B). The Great Arc generally moved from west to east during Cenozoic time and formed arc-collision zones in areas where the east to northeastward-facing arc collided with continental passive margins like those of the Ba-
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY hamas platform or northern South America (Pindell and Barrett, 1990). (4) Prior to these collisions, the Great Arc experienced a Paleogene back-arc rifting phase that led to the formation of the Yucatfin (Rosencrantz, 1990) and Grenada (Bird et al., Chapter 15) back-arc basins leaving remnant arcs at the Cayman and Aves ridges (Fig. 1B). The only active segment of the arc is found in the Lesser Antilles (Fig. 1B). This segment of the arc did not encounter continental material and continues to subduct oceanic crust of the Atlantic Ocean. Trinidad and eastern Venezuela lie within a complex transpressional plate boundary zone separating the Mesozoic passive margin of South America from the southern ends of the active (Lesser Antilles) and inactive (Aves Ridge) Great Arc of the Caribbean (Fig. 1B). Active tectonics and seismicity of the Trinidad region
Earthquake epicenters at depths from 0 to 200 km are densely concentrated in the Trinidad area (Fig. 2A). Events greater than 20 km are generally related to the Benioff zone of the subducted oceanic crust of the Atlantic Ocean that passes to a depth of about 150 km to the northwest of Trinidad (Perez and Aggarwal, 1981; Wadge and Shepherd, 1984; Russo et al., 1993) (Fig. 2B). The concentration of shallow to deep events is related to tear faulting at the junction between the subducting oceanic part of the South America plate and the unsubducted continental part of the plate as shown in Fig. 2B (Molnar and Sykes, 1969, 1971). Russo et al. (1993) compiled earthquake focal mechanisms from this region of tear faulting using both local network results and worldwide databases. They grouped focal mechanisms into three main groups: (1) right-lateral strike-slip events within a 60-km-wide east-west zone parallel to the Mor6nE1 Pilar fault system; (2) shallow thrust faults with east-northeast striking fault planes between the Araya Peninsula of Venezuela and Gulf of Paria to the east of Trinidad; and (3) shallow normal faulting events east and northeast of Trinidad which are interpreted to be expressions of bending of the oceanic part of the South America plate as its subducted at the Lesser Antilles subduction zone (Fig. 2B). As in other subduction to strike-slip transition areas of the world, the amount of seismicity falls off markedly from the Lesser Antilles subduction zone to the strike-slip zone parallel to the Mor6n-E1 Pilar strike-slip faults (Fig. 2A). Earthquake generation and strain release is less efficient along strike-slip boundaries in comparison to subduction boundaries. Stored seismic energy along strike-slip
499
faults is released in a few large events, perhaps separated by centuries, rather than the smaller more frequent events characteristic of most subduction settings (Mann et al., 1990). Known surface rupture associated with an earthquake along the E1 Pilar fault zone occurred in 1931 (Mann et al., 1990) and in 1996 (R.M. Russo et al., pers. comm., 1997). Perez and Aggarwal (1981), Speed (1985), and Russo and Speed (1992) have pointed out that the low levels of strike-slip earthquake activity on the E1 Pilar fault zone east of its intersection point with the Los Bajos fault near the Paria Peninsula and the absence of large historical strike-slip events in the Trinidad area are consistent with the interpretation that the E1 Pilar fault zone in the Trinidad area has become inactive.
REGIONAL GEOLOGIC SETTING OF NORTHERN SOUTH AMERICA AND TRINIDAD Offsets of strike-slip faults in northern South America and Trinidad
Right-lateral strike-slip faults form prominent morphologic features that can be divided into two groups with difering fault strikes in northern Venezuela and Trinidad: (1) east-west-striking faults in the coastal areas of northern Venezuela and northern Trinidad (Mor6n-E1 Pilar; E1 Coche-North Coast); and 2) east-southeastward- to southeastwardstriking faults in more inland areas of northern Venezuela and in southern Trinidad (unnamed faults of the Foothills fold-thrust belt, Urica, San Francisco, E1 Soldado, Los Bajos, and Warm SpringsCentral Range) (Fig. 3). As can be seen on the radar image in Fig. 3, the first group of faults have much longer fault traces that are straighter and more continuous over distance of hundreds of kilometers, while faults of the second group are more curved and extend over distances less than 100 km. Known strike-slip offsets of both groups of faults total from 10 to 70 km (cf. compilation in Mann et al., 1990) (Fig. 4A). East-west strike-slip faults in Venezuela and Trinidad M o r 6 n - E l Pilar fault zone This fault is traditionally identified as the main plate boundary fault of the South AmericaCaribbean plate boundary zone because of its geomorphic prominence, its east-west straight character over distances of hundreds of kilometers, and the strike-slip ruptures of the fault trace during the earthquakes of 1931 and 1996 (Schubert, 1984; R.M. Russo, pers. commun., 1997) (Figs. 3, 4A). The turning of the E1 Pilar fault zone northeast of Trinidad
500
S. BABB and E MANN
Fig. 2. (A) Relationship of earthquakes to faults and lithologic belts along the Caribbean-South America plate boundary zone. Earthquake epicenters have magnitudes between 4.0 and 6.2 and are from the International Seismological Centre (ISC) catalogue. Direction of South America-Caribbean plate motion is from DeMets et al. (1994). Area of vertical slab of subducted Atlantic crust is from Wadge and Shepherd (1984). (B) Block diagram showing a three-dimensional view of the tectonic features seen on the gravity map of the southeastern Caribbean shown by Mann et al. (Chapter 1). Black lines on the North and South America plates are fracture zone highs that pass under the accretionary wedge of the Barbados Ridge complex and are subducted at the Lesser Antilles subduction zone. The dense concentration of intermediate and deep earthquakes at the junction of the transform and subduction boundaries near Trinidad are related to a tear fault along which the South America plate is constantly torn as the Caribbean plate moves eastward over it.
into a northeasterly direction toward the lithospheric trace of the Lesser Antilles subduction zone supports its interpretation as a major plate boundary zone fault of the southeastern Caribbean plate (Robertson and Burke, 1989; Figs. 2, 4). Right-lateral offsets of the M o r 6 n - E 1 Pilar fault zone based on field mapping in the A r a y a - P a r i a area range from 20 k m (offset of B a r r e m i a n - A p t i a n rocks, Vierbuchen, 1984) to 70 k m (amount of displacement needed to form the Cariaco pull-apart basin b e t w e e n the offset
strands of the Mor6n and E1 Pilar fault zones over the past 2 million years, Schubert, 1984) (Fig. 4A). In the Trinidad area, Robertson and Burke (1989) postulated 3 5 - 4 0 k m of offset on the E1 Pilar fault by realigning the truncated Late M i o c e n e synclinal axis of the Northern basin and an u n n a m e d offshore, Late Miocene synclinal basin to the east of Trinidad (Fig. 4A). An offset of 3 5 - 4 0 km over the past 6 - 8 m.y. would yield an average rate of ~ 5 m m / y r on the E1 Pilar fault zone.
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY El Coche-North Coast fault zone This high-angle, submarine right-lateral strikeslip fault extends parallel to the coast of northern Venezuela and Trinidad and is best described from seismic reflection profiles presented by Robertson and Burke (1989) in the area north of Trinidad (Fig. 4A). Its strike-slip offset is unknown. East-southeast- to southeastward-striking faults in Venezuela and Trinidad Faults in the Foothills fold-thrust belt Two east-southeastward-striking right-lateral faults offset the frontal thrust fault of this Paleogene fold-thrust belt in central Venezuela by 1030 km (Fig. 4A). This fold-thrust belt formed mainly in Eocene and Oligocene time and is responsible for the formation and subsidence of the Gu~irico sub-basin of the Eastern Venezuelan basin (Parnaud et al., 1995) (Fig. 4B). Urica fault zone This southeastward-striking fault zone exhibits 35 km of right-lateral offset and forms the western edge of the Serranfa del Interior (Fig. 4A). Munro and Smith (1984) constrained the 35-km right-lateral offset amount using gravity data. This fold-thrust belt formed mainly in Oliogocene-Miocene time and was responsible for the formation and subsidence of the Maturfn sub-basin of the eastern Venezuelan foreland basin (Erlich and Barrett, 1992; Roure et al., 1995; Parnaud et al., 1995; di Croce et al., Chapter 16) (Fig. 4B). Shortening is now inactive as suggested by the distribution of earthquake epicenters (Fig. 2A) and unconformity relationships (Roure et al., 1995). San Francisco fault zone and unnamed parallel faults This fault exhibits 18-40 km of right-lateral offset that is readily apparent on geologic maps of the Cretaceous and Paleogene passive margin units in the Serranfa del Interior (e.g., Parnaud et al., 1995) (Fig. 4A). The radar image shown in Fig. 3 shows three other unnamed faults between the San Francisco and the Gulf of Paria (Nos. 4, 5, and 6) with similar strikes, tectonic geomorphology, and possibly strike-slip offsets. El Soldado fault zone This fault is present in the eastern Gulf of Paria and is believed to be a fight-lateral strike-slip fault (Fig. 4A). Its offset is unknown. Los Bajos fault zone The Los Bajos fault zone is continuous with the northwestern part of the Warm Springs fault zone
501
(Fig. 4A). Onshore offset of the axis of the ErinSiparia syncline indicates 10.5 km of fight-lateral offset for the Los Bajos (Wilson, 1940). The Los Bajos and Warm Springs faults strike southeast and are oblique to the trend of the E1 Pilar fault zone (Fig. 4A). Tyson (1989) mapped transpressional and transtensional features along the fault trace and showed that they are closely dependent on fault terminations and slight changes in the direction of the fault trace. The Los Bajos fault zone curves to the northeast in southern Trinidad with strike-slip offset converted to thrust offset on the Galeota Point thrust fault (Fig. 4A) that bounds the steep, southern coast of the island and is adjacent to the Columbus foredeep basin (Fig. 4B). Warm Springs-Central Range fault zone The Central Range fault zone was originally mapped by Kugler (1953) as a northwest-dipping thrust fault that bisects the topographically higher elevations of the Central Range and bounds the oldest outcrops of the lower Cretaceous Cuche Formation (Fig. 4A). Based on new data presented in this paper, we propose that the Central Range fault is the onshore continuation of the offshore Warm Springs fault zone in the Gulf of Paria and that both faults are sub-vertical right-lateral strike-slip faults. A component of transtension on the Warm Springs and transpression on the Central Range fault can account for the differences in the structural and topographic character of the Gulf of Paria and Central Range. Cenozoic sedimentary basins of northern South America and Trinidad and their relationship to major strike-slip and thrust faults
Cenozoic basins of this area can be divided into two groups of similar shape and tectonic origin: (1) basins elongate in an east-west direction and controlled by roughly east-west strike-slip fault zones in the coastal areas of northern Venezuela and northern Trinidad (Mor6n-E1 Pilar; E1 CocheNorth Coast); and (2) foreland basins elongate in an east-northeast direction and bounded by roughly east-northeast-striking frontal thrusts of the Foothills belt and Serranfa del Interior (Gu~irico and Maturfn sub-basins of the Eastern Venezuelan foreland basin) (Fig. 4B). The first group of strike-slip basins are much smaller in area than the second group and have much thinner sedimentary fills (1-3 km) than basins of the second group (5-10 km). North Coast basin This offshore basin filled with marine sedimentary rocks forms the northern flank of the Northern Range of Trinidad and was described by Robertson
502
S. BABB and R MANN
o
~
~
o
~0
~'~
~=~ =
I1.1
~
~-~ . ~
~~
.~"~
~)
-,~
~ "~
~
{o ~
~~
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY and Burke (1989) (Fig. 4B). Basal units are Middle to Late Miocene in age, and overlie Pliocene and Pleistocene rocks that are mainly claystone and sandstone. The total thickness of the basinal rocks is about 2 km. The North Coast basin appears to have formed along the obliquely downthrown side of the North Coast-E1 Coche fault zone (Fig. 4A).
Gulf of Paria-Northern basin This off- and onshore, 2- to 3-km-thick basin occurs south of the Paria Peninsula and Northern Range and forms the focus of this paper (Fig. 4B). Basin subsidence is closely related to down-to-thesouth and lateral movements on the right-lateral E1 Pilaf fault zone and down-to-the-north and lateral movements on the Warm Springs-Central Range fault zone (Fig. 4A). The Southern basin in southern Trinidad formed in a post-Early Miocene low south of the Central Range and was affected by an early phase of Middle Miocene folding and reverse faulting (Dyer and Cosgrove, 1992) along with Late Miocene to Recent transpression.
Columbus basin The Columbus basin is the offshore basin directly adjacent to and southeast of the Galeota Point thrust fault, that terminates right-lateral slip along the Los Bajos fault (Fig. 4B). Michelson (1976), Leonard (1983) and di Croce et al. (Chapter 16) document a thick fill of sediments mainly derived from Late Miocene to Holocene progradation of the Orinoco delta of eastern Venezuela (Prieto, 1987) into nonmarine to shelf-slope environments of the offshore area separating eastern Venezuela and Trinidad.
Maturin and Gufirico sub-basins of the Eastern Venezuelan basin These sub-basins of the Eastern Venezuelan basin formed as foreland basins mainly in Oligocene to Pliocene by thrust movement along the frontal thrusts of the Foothills belt and Serran/a del Interior, respectively (Gonzalez de Juana et al., 1980; Rohr, 1991; Erlich and Barrett, 1992; di Croce et al., Chapter 16) (Fig. 4B). The Urica fault forms the boundary between the Matur/n and Gufirico sub-basins (Erlich and Barrett, 1992) (Fig. 4A).
Morochito piggyback basin This basin formed on top of the Serrania del Interior fold-thrust belt in response to out-of-sequence thrust movement on the Pirital thrust fault (Route et al., 1995) (Fig. 4B). It is filled with a tilted sequence of shallow-water and continental deposits of Late Miocene to Pliocene age.
503
Lithologic belts of northern South America and their extensions in the Trinidad area
The South America-Caribbean plate boundary zone can be subdivided into several subparallel lithologic belts defined on the basis of similar lithologies and ages of deformation and metamorphism (Fig. 4C). Some of these belts appear to extend from eastern Venezuela into Trinidad. From north to south the main features of these belts are summarized below. Caribbean volcanic arc These mildly deformed and metamorphosed rocks represent the leading edge of the exotic Caribbean plate that is presumably moving eastward from its original Cretaceous location in the eastern Pacific Ocean (Burke, 1988; Pindell and Barrett, 1990) (Fig. 1B). This belt contains well exposed Early Cretaceous to Paleogene arc rocks on the island of Tobago northeast of Trinidad (Snoke et al., 1990), Cretaceous-Paleogene granitic and basaltic intrusive rocks on Margarita Island (Av6 Lallemant, 1997), and Cretaceous-Eocene volcanic rocks from wells and outcrops in the Leeward Antilles (Speed et al., 1985) (Fig. 4C). The precise boundary between the Caribbean volcanic arc and the Cordillera de la Costa belt to the south is poorly defined because elements of both belts are found across a broad zone on the island of Margarita (Av6 Lallemant, 1997). Cordillera de la Costa belt This highly deformed metamorphic belt includes lithologies mainly affiliated with both the Mesozoic passive margin of South America and exotic oceanic lithologies presumably obducted during the oblique collision between the Caribbean volcanic arc and the passive margin (Fig. 4C). Av6 Lallemant (1997) subdivided this belt into the Coastal Range/Margarita belt consisting of rocks of oceanic affinity and the Cordillera de la Costa belt sensu stricto consisting of continentally derived rocks. The former belt contains serpentinites and other ophiolitic rocks while the latter belt contains mainly low-grade metamorphic rocks of passive margin affinity and a few metaigneous and serpentinite bodies. The age of the protolith of the metasedimentary rocks is known from a few places like the Araya-Paria Peninsula of Venezuela (Av6 Lallemant, 1997) and the Northern Range of Trinidad (Potter, 1976) to be of Jurassic or Early Cretaceous age and to have been deposited in a passive margin setting (Fig. 4C). Deformation of these rocks is complex with one or two metamorphic phases of deformation and several brittle phases. Av6 Lallemant (1997) proposed that the high-grade part of the belt on the Araya Peninsula experienced pure and simple shear in its
504
S. B A B B and R M A N N
Fig. 4. (A) Offsets of strike-slip faults and thrust belts in northern South America. Offsets of both thrust and strike-slip faults along the plate boundary are controversial. Some tectonic models call for large offsets >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
S. BABB and E MANN (D
r (D
9 (D
9 .,..,
(D o,..~ ~D r/3 9 ~D
Z ~0
~
.=. 9
9 ~ (D 9 ~D .,..~
r/3 (D
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.
9
?: I 9 .,.., ~9 (D
O
9 .,..~ 9
~
r
~9
:< I
< ~
oK
~
Passive to active margin transition Transitional units between the deeper, passive margin Paleogene units and the shallower-water active margin Neogene units include the Brasso and Tamana Formations (Fig. 8). During the Early Miocene, calcareous clay, silt, and conglomerate of the Brasso Formation were deposited across the Northern Basin and Gulf of Paria (Fig. 7). Foraminiferal and macrofaunal assemblages in the Brasso Formation indicate an outer neritic environment of deposition. The Tamana Formation is a reefal limestone unit developed in the Central Range during the latest Early Miocene and middle part of the Middle Miocene (Erlich et al., 1993) (Fig. 8). The Tamana Formation is interpreted as a shallowing-upward sequence, where outer shelf and slope shales grade stratigraphically upsection into highenergy shoreline and intertidal limestone and shale (Erlich et al., 1993). The occurrence of the Tamana limestone indicates the first appearance of shallow-water conditions in the Central Range area (Carr-Brown and Frampton, 1979). Robertson and Burke (1989) and Erlich et al.
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
511
S T R A T I G R A P H Y OF THE N O R T H E R N BASIN AND GULF OF PARIA, TRINIDAD TIME (Ma) I AGE" POL.',,,N FORMATION S LITHOLOGY O ......
Z ,,,
Cedros
h" rr 1.65 9
.~
Upper Talparo
~,?,8 . 9~ ~ -
"
.
~ E
.
-
.
DEP. ENVlR.
SEQ.
.
Sst, Cly
5
Brackish water
I
I
Central Range transpression, 9 Gulf of Paria transtension
I
~. ~ =
Sst, Cly
Open marine
Sst, Sits, Cly
Inner neritic
Lower Talparo Springvale
~;;.% %=%"
' Northem Basin receives sediment 9 from Central Range; W a r m Springs, El Pilar ii faults active and pull=apart initiates
=i/.-,p=
5.2
"""'" " i~.~.~:~. Manzan,lla Cgl, Sst, Sh
Brackish water to inner neritic
Late ~ t..~..~_-,~_,.~, M ioc. ~ ~~~~_~.=. 10.2 -
"~
Early- = Middle ' Mioc.
252 --
66.5
96-
TECTONICS
Coastal terraces
n
~. I 3.5-
Sst, Cly, Cgl
I Olig Eoc Pal' , L. C'ret.
. '
,
Early Cret. 131 Late Jurassic
Tamana Brasso_
, Sst, Sh, , Nanva'--Cgl (rare)--
Cipero C , Lst, S a n F emando.., - - ==.,~1. i-...,.,Sst, oh N.. .. ... .. .. . . . . . . . . . . .Sst ......
Pointc-c~Pieffe Chaudiere
Guyaguayare ..... NaDadma H i l = - Gautier Cuche.
Clavs Ma. .-.. ==~m~-ffallow water~nentic " Deep water
Sst, Sh
i
~
~
II ~
, Not .~north; thrusting i n ~ -included-Southern basin --in - First tectonic ----" this ~activity on passive_ " study margin? --
D~ep water
Deep wate, ~
" Siltstone/Claystone: Sh, Sst Lst~ Sst, Sh
Deep neritic
Slope
L
1 ,
' ' Lst, Evp, Sh ~Carbonate platform~ '--Couva Evaporite----' Evaporite, Sh ,------Restricted ,
Lst, Evp, Dol, Sh
I
,~~ "~iFirstshedding of --'clastic rocks from--
_Carbonate_.--'Shallow marine , reefs ,
~['
2------Passive margin :subsidence
Carbonate platform?
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
512
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.
514
S. B ABB and R M A N N (D ,.~ ,..a .,..~ ..~
%...~ %...~ O 9
~,
.,...,
,..a
.,..~
O
. ,...~ . ,...,
Z ~
.,.~
O
~~
9
,-..,
tt'3
~D
.,zZ~ 9
~3
t,..,
..~
~
.,..,
.,..,
.~
o
.,..a
+..a
~o~
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
515
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.
516
S. BABB and E MANN
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
517
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
S. BABB and E MANN
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.
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
519
EAST
WEST WELL E ~
I
' I
6 KM
WELL V
16.5 KM
WELL J
R
~<~
~
0
1-
SEQUENCE 5 w
_J
I
I ,
< "r
,
DURHAM SAND
03
.----=SB
Z
I.U 7' O 103 a r
MAXIMUM FLOODING SURFACE
SEQUENCE 4
UNIT IV
z < 03
SB
14.1 c3
T a BRACKISH mu.i
I
~
n.. LU I--
UNIT V
UNIT III
'SEQUENCE 3
z_
MIDDLE MIOCENE SHALE OF BRASSO FM.
UNIT II DETAILEDLOG OF MIDDLE MIOCENE UNCONFORMITYAT WELL E
SHELF MIDDLE MIOCENE UNCONFORMITY SURFACE
SB L,=18GO
MIDDLE MIOCENE, SHALE OF BRASSO FM.
Fig. 15. East-west stratigraphic correlation between electric logs of wells E and J 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. Paleoenvironments based on benthic foraminifera are listed to the left. Inset shows detail of electric log of well E. Paleodepth work in well E reveals a major environmental change from middle-outer shelf in the Middle Miocene to brackish in the Late Miocene. Lithologic data from the well indicate that this abrupt transition is represented by an unconformity with possible subaerial exposure indicated by red shale horizon.
Conglomerates of units I and II are correlated with the Cunapo Formation known mainly from drillholes including well F of the Northern basin (Payne, 1991; Tyson et al., 1991) (Fig. 9A). Equivalent and interfingering sandy units to the south are correlated with the Manzanilla Formation which has been well described from outcrops along both flanks of the Northern basin (Saunders, 1968; Carr-Brown and Frampton, 1979). These workers have subdivided the
Manzanilla Formation into three, shallowing-upward members: (1) the oldest San Jose member consists of dark, calcareous silts, rich in small molluscs and foraminifera deposited in an inner neritic environment; (2) the Montserrat glauconitic sandstone overlies the San Jose and was deposited on a shallow shelf; and (3) the Telemaque member records shallowing and infilling of the Northern basin as indicated by arenaceous foraminiferal faunas.
520
S. BABB and R MANN
Fig. 16. East-west stratigraphic cross-section between wells B and S showing the correlation between sequences 3, 4, and 5. The datum is the base of the Durham sand at the base of sequence 5. The lithology of unit IV is described in detail in the text. The top of the Lower Cretaceous platform in wells B and S is indicated.
Lithologic variations of sequence 3 in an east-west direction Sequence 3 can be divided into units II and III described in detail above on two approximately east-west basin profiles using wells E, V and J in the Goodrich sub-basin and the southern part of the Northern basin (Fig. 15) and wells B, H, T1, P, and S on the Avocado high and Northern basin (Fig. 16). On the first profile in Fig. 15, the lower boundary of sequence 3 is not penetrated in well V but is recognized in both wells E and J. Unit II of sequence 3 appears to be missing in well E, which is located in the Goodrich sub-basin (Fig. 9B). The east-west profile along the northern part of the Gulf of Paria and Northern basins using wells B, H, T1, P, and S in Fig. 16 indicates two areas of predominantly conglomeratic fill of the Cunapo Formation in the
west at well B and in the east at well S. The conglomeratic wedge thickens eastward from 1230 m at well B on the Avocado high to 2340 m at well S. The boundaries between conglomerate, sandstone, and shale units I, II and III of sequence 3 are not well defined on this profile (Fig. 16).
Geologic interpretation of sequence 3 in an east-west direction The southward transition from massive conglomerate along the E1 Pilar fault zone to sandstone and shale of the basin axis remains a relatively straight, east-west boundary. The conglomerate wedge appears to have been about 20 km wide in the west and narrowed to a width of 10 km or less in the Northern basin.
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
Age of sequence 3 (Cunapo and Manzanilla Formations) The age of sequence 3 and its equivalent lithologic formations shown in chart form in Fig. 8 is based on paleontology and palynology of well cuttings tied to seismic lines. The source of these biostratigraphic data was internal company reports of the Petroleum Company of Trinidad and Tobago. Sequence 3 is a Late Miocene-Early Pliocene, mainly shallow-marine to brackish-water conglomerate and sandstone unit whose age is based on recognition of palynozones UMI and UMII in well V (Fig. 15) which has been tied to the seismic line shown in Fig. 13. Sequence 3 is bounded at its base by the postMiddle Miocene unconformity surface identified using well data. The age of the lower boundary of the unconformity is based on recognition of the Globorotalia mayeri zone (top Middle Miocene) in well E (Fig. 15). Truncation beneath the Middle Miocene unconformity is not clear on lines shown in Figs. 12 and 13 because of poor seismic imaging at this depth. The base of the siliciclastic wedge of sequence 3 (Cunapo Formation) was deposited in an outer neritic environment based on paleobathymetric data from well B that is tied to the seismic line shown in Fig. 12 (internal report of the Petroleum Company of Trinidad and Tobago, 1991).
Seismic character and definition of sequence 4: Early to middle Pliocene Manzanilla and Springvale Formations Distribution of sequence 4 Using the seismic lines and wells shown in Fig. 6A, sequence 4 was mapped in detail in the E1 Pilar, Puerto Grande, and Goodrich sub-basins of the Gulf of Paria basin. Sequence 4 is also inferred to extend eastward and underlie much of the Northern basin (Fig. 7). Sequence 4 is thickest in the Goodrich sub-basin adjacent to the Warm Springs right-lateral strike-slip fault zone and thins to the north in the E1 Pilar and Puerto Grande sub-basins (Fig. 9A).
Base reflector of sequence 4 The base of sequence 4 on a northwest-southeast line crossing the Goodrich sub-basin (Fig. 9A) is defined by downlap and onlap of a northwestward-directed sedimentary wedge south of well V (Fig. 13). Onlap of the base of sequence 4 across a deformed sequence 3 is also observed on a northsouth line in the E1 Pilar sub-basin (Fig. 18). Sequence 4 undergoes pronounced thinning over these underlying structural highs.
521
Top of sequence 4 The upper boundary of this sequence is an erosional truncation. The underlying sequence 3 is disrupted by what appears to be listric normal faults (Figs. 17 and 18). Seismic facies within sequence 4 A variety of seismic facies are present within sequence 4 including prograding oblique clinoforms and associated mounds, low-amplitude continuous reflectors, and very low-amplitude, continuous reflectors. East of the Goodrich sub-basin, high- to mediumamplitude, relatively continuous, oblique clinoforms prograde eastward and downlap apparent mounds at the base of sequence 4 (Fig. 19). In the Goodrich sub-basin, sequence 4 is characterized by a zone of very low-amplitude and continuous reflectors (Fig. 20). The low-amplitude seismic facies appears transitional to higher-amplitude reflectors on the more proximal northern and southern ends of this northwest-southeast line. In the E1 Pilar and Goodrich sub-basins, a zone of two to five, very high-amplitude continuous reflectors occurs near the upper contact of unit 4 (Fig. 11). This seismic facies is found in the northern and southern parts of the Gulf of Paria basin. Correlation of seismic facies of sequence 4 with lithologies from electric logs and cores from wells Cross-sections were made in both an approximately north-south (Fig. 14) and east-west direction (Figs. 15 and 16) to provide a better understanding of how seismic/sedimentary facies of sequence 4 vary on a basin-wide scale. Spacing between wells along these sections varies from 5 to 16 km. Lithologic variations of sequence 4 in a north-south direction Sequence 4 can be divided into two 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 and sandy rocks of the Cunapo Formation of Late Miocene to Early Pliocene age (Fig. 8). Unit IV of this conglomeratic section is a 50-m-thick retrogradational stack of massive conglomerate, sandstone, and shale that is transitional to equivalent sandstone and shale at wells C1 and C2 10 km to the south (Fig. 14). In well V, unit IV consists of poorly organized sand overlain by a well developed shale interval interpreted as a maximum flooding surface. The position of this maximum flooding surface is defined by the turnaround from a retrogradational stacking pattern of unit IV to the
522
S. BABB and E MANN
Fig. 17. North-south seismic section in northern Gulf of Paria basin showing the boundary between sequences 4 and 5 defined by erosional truncation. See Fig. 9B for location. progradational stacking pattern of unit V (Fig. 14). The east-west cross-section of the basin using wells E, V, and J (Fig. 15) and B, H, T1, R and (Fig. 16) S shows a similar shale-dominated unit V above the maximum flooding surface.
Geologic interpretation of sequence 4 in a north-south direction Very low-amplitude continuous reflectors seen on a north-south line through the Goodrich sub-basin are interpreted as a uniform shaley lithology in the central part of the Goodrich sub-basin (Fig. 20). This lithology is transitional with higher-amplitude reflectors to the north and south. These high-amplitude reflectors at the basin edges are interpreted as representative of higher-impedance sand and conglomerate at the northern and southern edge of the basin (Fig. 12). Retrogradational stacking of unit IV in well B indicates basin deepening and overall transgression
(Fig. 14). The very high-amplitude, continuous reflectors in the upper part of sequence 4 in the northern and southern parts of the basin are interpreted to represent very uniform, quiet sedimentary conditions during the period of maximum flooding recorded by shaley horizons of unit V in Figs. 14 and 15. This maximum flooding surface indicates a change from retrogradation to progradation within sequence 4 and defines a major reorganization in basin deposition.
Geologic interpretation of sequence 4 in an east-west direction Mounded reflectors at the base of sequence 4 on the line in Fig. 19 suggest the occurrence of gravity flow processes and formation of lobes at the base of a slope during eastward progradation of sequence 4. Clinoform geometry also seen in Fig. 19 indicates an eastward increase in water depth. Seismic characteristics and age of sequence 4 indicate that it can be correlated with the Springvale
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
523
Fig. 18. North-south seismic section in northern Gulf of Paria basin showing onlap and thinning of sequence 4 onto folded and tilted sequence 3. The section also illustrates syn-faulting growth along young faults affecting sequence 5. See Fig. 9B for location.
Formation, which is interpreted by Carr-Brown and Frampton (1985) to represent a marine incursion into the Northern basin (Fig. 8). Faunas from the Springvale Formation indicate its deposition in an inner neritic environment. Despite the deepening trend in sequence 4, brackish water conditions persisted along the margins of the Gulf of Paria and Northern basins adjacent to the Northern and Central Ranges (Fig. 6A). Paleobathymetry data from well L in the southwestern Northern basin (Fig. 6A) indicate that rocks of sequence 4 at this locality record a coastal environment with marginal marine influence. Uplift of the Central Range during this time may have resulted in
narrowing of the Northern basin and consequent development of restricted, brackish water environments (Fig. 6A). Deepening in the central part of the basin may have resulted in the transgressive and retrogradational character of sequence 4. In the northern Gulf of Paria basin, deposition of sequence 4 was localized in the Puerto Grande sub-basin (Fig. 9A) which indicates an eastward shift in the depocenter from the E1 Pilar sub-basin of sequence 3 time.
Age of sequence 4 The age of sequence 4 and its equivalent lithologic formations is based on paleontology and palynology of well cuttings tied to seismic lines. The
524
S. BABB and R MANN
Fig. 19. Northwest-southeast seismic section in Goodrich sub-basin showing eastward progradation and downlap of clinoforms of sequence 4 over those of sequence 3 and mounded reflectors at base of sequence 4. See Fig. 9B for location.
source of these biostratigraphic data was internal company files of the Petroleum Company of Trinidad and Tobago. Sequence 4 is a mid-Pliocene, mainly shallow-marine to brackish-water sandstone and shale unit whose age is based on recognition of palynozones PI and PII (Fig. 8).
Base reflector of sequence 5 The base of sequence 5 is defined by an erosional unconformity in the eastern Gulf of Paria basin (Fig. 17). In other parts of the Gulf of Paria basin, the base of sequence 5 is conformable with sequence 4 (Figs. 18-20).
Seismic character and definition of sequence 5" Late Pliocene-Pleistocene Talparo and Cedros Formations
Top of sequence 5 For the purpose of this study, the top of sequence 5 is assumed to be the seafloor in the Gulf of Paria basin or the ground surface in the Northern basin. Payne (1991) identified two additional unconformity surfaces within the Early to Middle Pleistocene of sequence 5 that are not recognized in this study.
Distribution of sequence 5 Using the seismic lines and wells shown in Fig. 6A, sequence 5 was mapped in the E1 Pilar, Puerto Grande, and Goodrich sub-basins of the Gulf of Paria basin. Sequence 5 is also inferred to extend eastward and underlie much of the Northern basin (Fig. 6A). Sequence 5 is thickest in the Goodrich sub-basins adjacent to the Warm Springs fight-lateral strike-slip fault zone and thins to the north in the E1 Pilar and Puerto Grande sub-basin (Fig. 9A).
Seismic facies within sequence 5 Three main seismic facies are present within sequence 5: prograding oblique clinoforms and associated mounds, low-amplitude continuous reflectors, and continuous and parallel reflectors.
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
525
Fig. 20. Northwest-southeast seismic section in the Goodrich sub-basin showing low-amplitude seismic facies of sequence 4 in the central part of the basin and clinoforms showing southeastward progradation in sequence 5. See Fig. 9B for location.
In the southeastern Goodrich sub-basin, highto medium-amplitude, relatively continuous, oblique clinoforms of sequence 5 cyclically prograde northwestward (Fig. 13) and southeastward (Fig. 20). These clinoforms are continuous with relatively parallel reflectors interpreted as basinal, deeper-water equivalents of strata represented by the clinoforms.
Correlation of seismic facies of sequence 5 with lithologies from electric logs and cores from wells Cross-sections were made in both an approximately north-south (Fig. 14) and in an east-west direction (Figs. 15, 16) to provide a better un-
derstanding of how seismic/sedimentary facies of sequence 5 vary on a basin-wide scale. Spacing between wells along these sections varies from 5 to 16 km.
Lithologic variations of sequence 5 in a north-south direction Sequence 5 can be correlated as a relatively uniform tabular body on wells in a north-south direction (Fig. 14). A 24-m-thick sandstone unit, the Durham sand, is present near the base of sequence 5 in all wells in the profile (Fig. 14).
526
Geologic interpretation of sequence 5 in a north-south direction Clinoforms in the southeastern part of the Goodrich sub-basin indicate northwestward progradation of sequence 5 from source areas in the Central Range (Fig. 13). The presence of southeastward progradation of sequence 5 (Fig. 20) into the same basin suggests that the Goodrich sub-basin was a closed basin surrounded by source areas to the north (Northern Range) and south (Central Range). Paleoenvironmental data (Internal company report, 1991) from wells B (Fig. 14) and E and V (Fig. 15) indicate a shallow-marine to brackish-water depositional environment during the deposition of sequence 5 in the area of the Goodrich sub-basin and Avocado high (Fig. 9A). This type of environment had been recognized in the Lower and Upper Talparo Formations using outcrop samples from the northern and southern edges of the Northern basin (Kugler, 1953; Carr-Brown and Frampton, 1979) (Fig. 8). Geologic interpretation of sequence 5 in an east-west direction The area of thickest deposition of sequence 5 appears closely controlled by faults bounding the Goodrich sub-basin (Fig. 9A). For this reason, the sediment thickness decreases abruptly to the east in the Northern basin as seen from the thinner sections of sequence 5 in wells T1, P, and S in the interior parts of the Northern basin (Fig. 16). Age of sequence 5 (Talparo and Cedros Formations) The age of sequence 5 and its equivalent lithologic formations is based on paleontology and palynology of well cuttings tied to seismic lines. The source of this biostratigraphic data was internal company files of the Petroleum Company of Trinidad and Tobago. Sequence 5 is a Late Pliocene to Pleistocene, mainly shallow-marine to brackish water sandstone and shale unit (Fig. 8). The inferred environments of deposition for sequence 5 present in wells B, E, and V range from shallow marine to brackish water (Figs. 15, 16). Sequence 5 is correlated with the Upper Talparo and Cedros Formations (Fig. 8). The Lower Talparo Formation represents a marine environment while the Upper Talparo is indicative of brackish water (Carr-Brown and Frampton, 1979). In coastal exposures of the Talparo Formation along the east coast of Trinidad, Saunders (1968) identified lagoon, beach, and tidal fiat environments. The Pleistocene Cedros Formation is correlative with the upper part of sequence 5 and consists of poorly consolidated clay, fine to coarse sandstone with fragments of leaves and other carbonaceous matter (Kugler, 1953; Donovan, 1994).
S. BABB and E MANN INTEGRATION OF SEQUENCE IDENTIFICATIONWITH STRUCTURAL MAPPING
Introduction In order to best illustrate the tectonic controls on the three Neogene sequences defined above in the Gulf of Paria, we present structural contour maps of the boundaries between the three sequences. In addition, we present three isochron maps of the sequences overlying each of these surfaces. Comparison of the structural, contour and isochron maps can reveal correlations between the geometry of the surface preceding the deposition of a sequence and the thickness of that sequence. The comparison of these map pairs also allows better visualization of the space available at tectonically significant times in the history of the Gulf of Paria and the sedimentary infill onto that sequence stratigraphic boundary. These comparisons indicate source areas and paleogeography of land areas surrounding the basin. Limits of the approach are due to the effects of tectonic deformation, subsidence, and differential compaction.
Neogene deformation phase one: Late Miocene-Early Pliocene strike-slip motion along the El Pilar fault Structure contour map at top of sequence 2 (latest Early Cretaceous carbonate platform) This map (Fig. 21A) reveals the following features at the top of the Cretaceous platform horizon. (1) The Gulf Domoil and Avocado highs. These highs formed by the morphology of the latest Early Cretaceous bank margin that was later modified by Neogene faulting (Babb, 1997). The Cretaceous highs form the main east- and northwest-trending, relief-forming features on the seafloor at the time of initial deposition of the siliciclastic sediments in sequence 3 (Fig. 21B). (2) Lows. Prominent lows include the area northeast of the Gulf high and north and northwest of the Avocado high (El Pilar sub-basin) and an area of several isolated lows in the area between the Gulf and Domoil highs (Goodrich sub-basin). These lows in the Gulf of Paria appear to have formed along normal faults related to pull-apart formation. (3) Faults. Because faulting is predominantly of Late Miocene-Pleistocene age in the Gulf of Paria and Northern basins, the fault pattern on the top of the Cretaceous platform resembles that mapped in the late Neogene sequences. Isochron map of sequence 3 (Late Miocene-Early Pliocene) The isochron map of sequence 3 (Fig. 21B) reveals two lobate sediment bodies composed mainly
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
527
Fig. 21. (A) Structural contour map in two-way travel time to top of sequence 2 (latest Early Cretaceous horizon). The location and history of the latest Early Cretaceous edge of carbonate platform is indicated by heavy line. (B) Isochron map of Late Miocene-Early Pliocene sequence 3 in the Gulf of Paria and location of lines used to constrain isochrons. Source of sequence 3 is from the E1 Pilar fault zone and Northern Range.
528 of conglomerate of the Cunapo Formation. The body is about 0.8 s in two-way time and fills the low between the Gulf, Domoil and Avocado highs (El Pilar sub-basin). Southward thinning indicates a northern source for sequence 3 in the Northern Range as observed in the north-south profile shown in Fig. 14. A second, smaller lobe of sequence 3 is present south of a narrow saddle separating the Gulf and Avocado highs. This lobe appears to consist of thinner, 0.2 s two-way time material from the larger lobe to the north that was spilled southward through the saddle area. Because of this saddle, the Goodrich sub-basin remained partially isolated from an influx of sequence 3 from the northern source area during this period. Structural control on sedimentation within the basin itself during this phase was not pronounced. Gravity-driven coarse siliciclastic sedimentation appears to have filled existing lows adjacent to topographic highs created along the Early Cretaceous carbonate passive margin. The erosion of the Northern Range is interpreted here as reflecting down-tothe-south, oblique-slip motions on the E1 Pilar fault zone. We interpret this Late Miocene vertical motion as the first manifestation of strike-slip displacement on the E1 Pilar fault zone along which the Northern Range or equivalent rocks of the Coastal Fringe/Margarita lithologic belt of Margarita and Trinidad move eastward along the northern margin of South America (Fig. 4C).
Neogene deformation phase two: middle to Late Pliocene strike-slip motion along the Warm Springs-Central Range fault zone Structural contour map to top of sequence 3 (Early Pliocene horizon) This map (Fig. 22A) reveals the following features at the top of sequence 3 (Early Pliocene horizon). (1) The Gulf Domoil and Avocado highs. These highs still formed prominent highs despite some filling of lows during structural phase one. (2) Lows. The prominent low during this period was within the Goodrich sub-basin north of the Warm Springs fault zone. (3) Faults. The fault pattern at the top of sequence 3 reflects increased oblique opening of the Goodrich sub-basin. Movement on these faults is generally down toward the northwest-trending axis of the basin and acted to create additional accommodation space for the sediments of sequence 4. The deeper structure or topographic edges of the northwest-trending, underlying Cretaceous carbonate rocks may have acted to control the faults of the Goodrich sub-basin that parallel the edges of the banks.
S. BABB and E MANN
Isochron map of sequence 4 (middle-Late Pliocene) The isochron map of sequence 4 (Fig. 22B) reveals a major northwest-trending lobe that is confined by the faults bounding the Goodrich sub-basin. Northward thinning indicates a mainly southerly source in the Central Range with additional sources i n the west and northwest. Structural control on sedimentation during this phase is pronounced in the Goodrich sub-basin but less pronounced on more discontinuous faults in the area south of the E1 Pilar fault zone (Fig. 22A). We interpret this pattern as reflecting gradual abandonment of the E1 Pilar fault zone and transfer of slip to the Warm Springs fault zone across the oblique-slip faults in the Goodrich sub-basin. Neogene deformation phase three: Late Pliocene to Pleistocene strike-slip motion along the Warm Springs-Central Range fault zone Structural contour map at top of sequence 4 (Late Pliocene horizon) This map (Fig. 23A) reveals the following features at the top of sequence 4. (1) The Gulf Domoil and Avocado highs. The carbonate banks still form prominent seafloor highs despite the filling of their intervening lows during structural phases one and two. (2) Lows. The prominent low during this period remains the Goodrich sub-basin north of the Warm Springs fault zone. Continued fault movement on the oblique-slip faults may be have depressed the basin. (3) Faults. The fault pattern at the top of sequence 4 reflects increased oblique opening of the Goodrich sub-basin. The faults of this structural phase affect a larger area and are wider spaced than the faults of the early phase in the Goodrich sub-basin. This change in fault pattern suggests increased opening of the basin. Movement on these faults will act to create additional space to be filled by sediments of sequence 5 because of the increased rate of displacement along faults initiated in phase two. Isochron map of sequence 5 (Late Pliocene-Pleistocene) The isochron map of sequence 5 (Fig. 23B) reveals a widening and lengthening of the major northwest-trending lobe that is confined by the faults bounding the Goodrich sub-basin. Northward thinning indicates a continuation of the southern source in the Central Range and possibly the Orinoco delta that began in structural phase two (Fig. 22B). Structural control on sedimentation during this phase was pronounced in the Goodrich basin but not so obvious in the area south of the E1 Pilar
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
529
Fig. 22. (A) Structural contour map in two-way travel time to top of sequence 3 (Early Pliocene horizon). (B) Isochron map of Middle-Late Pliocene sequence 4 in the Gulf of Paria and location of lines used to constrain isochrons. Source of sequence 4 is mainly from the Warm Springs-Central Range fault zone and the Central Range, with additional sources from the west and northwest.
530
S. BABB and E M A N N
Fig. 23. (A) Structural contour map in two-way travel time to top of sequence 4 (Late Pliocene horizon). (B) Isochron map of Late Pliocene-Pleistocene sequence 5 in the Gulf of Paria and location of lines used to constrain isochrons. Source of sequence 5 is mainly from the Warm Springs-Central Range fault zone and the Central Range with possible input from the Orinoco River. Dark gray indicates thickest accumulation of sequence 5.
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY fault zone. We interpret this pattern as reflecting gradual abandonment of the E1 Pilar fault zone and transfer of slip to the Warm Springs fault zone across the oblique-slip faults in the Goodrich sub-basin. These faults open the Goodrich pull-apart basin and increases its accommodation space.
Comparison of paleocurrents and paleogeography produced during structural phases one through three Introduction Fig. 24 combines the isochron map of Figs. 2123 with paleocurrent data derived from seismic sections and from east-to-west and north-to-south stratigraphic sections based on alignments of wells (Figs. 14, 15, 16).
Paleocurrents Paleocurrents for structural phase one (deposition of sequence 3) are oriented parallel to and normal to the axis of the Gulf of Paria-Northern basin (Fig. 24A). North to south paleocurrents interpreted from clinoforms and thinning in the Late Miocene Cunapo conglomerate (Fig. 8) are consistent with the erosion of the Northern Range. Basin axis clinoforms such as one near well B (Fig. 9A) indicate a west to east flow along the floor of the basin. Paleocurrents for structure phases two (sequence 4) and three (sequence 5) become increasingly dominated by a west to east paleoflow along with a southeast to northwest flow in the area of the Goodrich sub-basin (Figs. 24B, C).
Environments of deposition Environments of deposition reflect the early uplift and erosion of the Northern Range during structure phase one followed by later uplift of the Central Range and the gradual decrease in the importance of the Northern Range and possibley Orinoco siliciclastic sources during structure phases two and three. Environments remained shallow and brackish along the E1 Pilar fault scarp and its fan apron based on well samples from wells shown in Figs. 14 and 16. To the south in the Gulf of Paria and Northern basins, early environments in phase one are brackish but become deeper water in phase two and return to brackish conditions in phase three. The deepest-water area in phases two and three remains the fault-controlled Goodrich sub-basin (Fig. 24B, C).
531
STRUCTURAL STYLES AND GEOMORPHIC EXPRESSION OF THE EL PILAR FAULT ZONE Introduction Here we illustrate the structural styles and geomorphic expressions of the E1 Pilar fault zone. The character, age of motion, and tectonic role of the E1 Pilar fault zone has been controversial (Fig. 5). This controversy has been in large part because there has been little seismic data available across the fault trace to constrain its structure and age of movement. We suggest on the basis of new seismic reflection data presented here that the strike-slip style seen along the E1 Pilar fault and its apparent inactivity in the Quaternary as one moves west along its trace supports a general model for progressive west-toeast abandonment of the E1 Pilar fault zone. This idea is consistent with our interpretation of the structural and isochron maps summarized in Fig. 24A-C in the previous section.
Transpressional style of the El Pilar fault zone in the Gulf of Paria Two lines across the E1 Pilar fault zone in the northern Gulf of Paria both suggest that the fault is a high-angle strike-slip fault with a transpressional component (both locations shown on map in Fig. 23). A seismic line east of Patos Island shows four fault strands of which the central strand breaks the seafloor and appears active (Fig. 25). Down-tothe-south components of movement on the southern strands have allowed a thick accumulation of sequences 3 and 4. The fault has an east-northeast strike in this area which appears to be a transpressional fault orientation throughout Trinidad (Robertson and Burke, 1989; Payne, 1991; Tyson, 1989). The upthrown northern side of the fault may have uplifted Patos Island to the west of the location of the line shown in Fig. 25. On the seismic line in Fig. 25, overstepping of the southern two strands by sequence 5 of Late Pliocene-Pleistocene age is consistent with decreased activity (or at least focussing) of active movement on the E1 Pilar fault zone. The line in Fig. 26 near the northeast coast of the Gulf of Paria (Fig. 23) shows three fault strands that are onlapped by sequences 3, 4 and 5. This onlap is consistent with the abandonment of the active traces. Sequence 3 exhibits northward growth consistent with the isochron map in Fig. 2lB. Sequence 5 is affected by a late period of transpressional reactivation. As in the case of the line in Fig. 25 near Patos Island, the fault trace has a slight east-northeast orientation that might be responsible for the late phase of transpressional reactivation.
532
S. BABB and R MANN
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY Transtensional style in diffuse extensional zone south of the El Pilar fault zone in the Gulf of Paria
The line in Fig. 27 is located across the trend of the Puerto Grande sub-basin (location in Fig. 23). The east-west-striking rift or negative flower structure reflects local extension responsible for the formation of this local depocenter and thickening seen in sequence 5. The faults bounding this sub-basin cannot be traced as continuous features to the west across the Gulf of Paria. This localized zone of extension may therefore represent a southward step in the main trace of the E1 Pilar fault zone that places it in a more basinward position within the onland Northern basin. This southward step would lead to a decrease in activity of the E1 Pilar fault along the mountain front of the Northern Range (Fig. 9A). Onland geomorphic expression of the El Pilar fault zone
Fig. 28 is an aerial photograph showing the Arima fault zone and onland extension of the E1 Pilar fault zone. In this area, the onland projection of the E1 Pilar fault zone is covered by alluvium and appears to have no geomorphic expression. Some or all of its geomorphic expression may have been lost during the intensive development of this heavily populated and urbanized area south of Port-of-Spain, the capital and largest city of Trinidad and Tobago. The Arima fault appears to have had mainly down-tothe-south oblique-normal throw followed by late up-to-the-south oblique-reverse reactivation as seen on the line shown in Fig. 26 (Algar and Pindell, 1993).
STRUCTURAL STYLE OF OBLIQUE SLIP FAULTS IN THE GULF OF PARIA Detachment off flank of Gulf high
The line in Fig. 29 illustrates the style of detachment of lower Tertiary or Upper Cretaceous to Pleistocene siliciclastic rocks off the Gulf high in the northern part of the Goodrich basin (Fig. 23). This structure is believed to result when the large lithologic contrast across the unconformity at the
533
top of the Early Cretaceous platform and the overlying siliciclastic sediments is subjected to slightly transtensional opening at the time of sequences 4 and 5. Downslope slumping off of the carbonate passive margin highs may contribute to the rapid infilling of the Goodrich sub-basin as seen on the isochron maps of sequences 4 (Fig. 22B) and 5 (Fig. 23B). Transtension between the Gulf and Domoil highs
The line in Fig. 30 illustrates transtension between these two highs with an early phase of transpression that has resulted in folding (Fig. 23). Thickening of sequence 5 shows that the later phase of movement has been oblique opening. This type of oblique opening and widening of the fault pattern can be seen by comparing the structural maps of sequences 4 (Fig. 22B) and 5 (Fig. 23B).
STRUCTURAL STYLES AND GEOMORPHIC EXPRESSION OF THE WARM SPRINGS-CENTRAL RANGE FAULT ZONE Warm Springs fault zone
The seismic lines in Figs. 31 and 32 show the negative flower structure profile of the active, rightlateral Warm Springs fault zone. The syn-tectonic thickening of Plio-Pleistocene sequence 5 north of the Warm Springs fault seen on the isochron map of Fig. 23B is also apparent on the seismic lines in Figs. 31 and 32. Transition between offshore Warm Springs and onshore Central Range fault zone
The geologic map in Fig. 33 and the aerial photo in Fig. 34 shows the abrupt transition at the western coast of Trinidad between the offshore east-west-striking and transtensional Warm Springs fault zone and the onshore, northeast-striking and transpressional Central Range-Caigual fault zone. Thick Early Pliocene-Pleistocene rocks of sequences 3-5 deposited by northward progradation along the eastern margin of offshore Goodrich sub-basin are exposed by uplift and folding along the northern edge of the Warm Springs-Central Range fault zone (Fig. 34). Widening of the out-
Fig. 24. (A) Paleogeographic map of the Gulf of Paria and Northern basins and basin flanks during the Late Miocene-Early Pliocene (sequence 3); (B) Paleogeographic map of the Gulf of Paria and Northern basins and basin flanks during the Early and Middle Pliocene (sequence 4). (C) Paleogeographic map of the Gulf of Paria and Northern basins and basin flanks during the Late Pliocene-Pleistocene (sequence 5). These maps were made by combining the isochron data shown in Figs. 21-23 with the well logs shown in Figs. 14-16 and with outcrop data published by previous workers and discussed in the text.
534
S. BABB and R MANN
Fig. 25. North-northeast-striking seismic line showing structure of east-northeast-striking, transpressional part of the E1 Pilar fault zone in the marine strait (Boca Grande) east of Los Patos Island in the Gulf of Paria (see Fig. 23A for location). Note onlap of southern and northern strands by sequence 5 of Late Pliocene-Pleistocene age and folding of reflectors between the fault strands. Seafloor scarp indicates continued activity on the central fault strand.
crop pattern of these rocks along the edge of the Northern basin is consistent with the transpressional nature of the uplift adjacent to the northeast-trending Central Range and Caigual fault zones (Fig. 33).
Reinterpretation of the structure of the Central Range There are two possible interpretations of the n o r t h - s o u t h section of the southwestern end of the
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
535
Fig. 26. North-south-striking seismic line from the northeastern Gulf of Paria basin showing structure of east-west segment of the E1 Pilar fault zone in the northeastern Gulf of Paria (see Fig. 23A for location). Note onlap of southern traces of the fault by sequences 3 and 4 of Early to Late Pliocene age and reactivation of the fault to form fold in sequence 5 of Late Pliocene-Pleistocene age.
Central Range adjacent to the Gulf of Paria (line of section shown in Fig. 6B). The first by Kugler (1953, 1959) and Speed (1985) proposed that the Central Range is an asymmetric, south-southeast-verging Late Neogene fold-thrust belt based on outcrop mapping. In the Kugler (1953, 1959) interpretation, shown in Fig. 35A, the thrust faults along
the northern flank of the Central Range are problematic because younger rocks are thrust over older rocks. A reinterpretation of the outcrop pattern is proposed here based on offshore profiles of the Warm Springs fault zone shown in Figs. 31 and 32. In this interpretation shown in Fig. 35B the Central
536
S. BABB and E MANN
Fig. 27. North-south-striking seismic line showing structure of the discontinuous zone of transtensional faults south of the E1 Pilar fault zone that form the Puerto Grande sub-basin in the northeastern part of the Northern basin (see Fig. 23A for location). Note growth of Late Pliocene-Pleistocene sequence 5 along faults. Range is a symmetrical, late Neogene restraining bend along the transpressional Central Range and Caigual strike-slip faults. The topographic uplift of Mount Harris, the highest point of the Central Range at 265 m, occurs at a slight left step between these two transpressional faults (Fig. 33). The high-angle Central Range-Caigual fault zone thrusts older rocks of the Central Range over younger rocks along the southern flank of the Northern basin. The predicted depth to the top of the Mesozoic carbonate platform is closer to the 3.5 km depth known from offshore and onshore seismic lines than the much greater depth predicted in the Kugler section (Fig. 35B).
This reinterpretaton is consistent with the interpretation of the line shown in Fig. 36 from the central part of the onland Northern basin (location shown in Fig. 33). In this line, the southeastern extension of the Fishing Pond fault zone is expressed as a 5-km-wide sub-vertical strike-slip fault or 'flower structure' along which the Guatapajaro anticline has been folded as a transpressional feature. The dip on the thrust faults beneath the anticline is southeastward and not northwestward as depicted by Kugler (1953, 1959) and Speed (1985). This transpressional interpretation of the Central and Southern Ranges uplifts bounded by thrusts and reverse faults dipping
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
537
Fig. 28. Air photograph of the Port-of-Spain area showing the scarp of the Arima fault separating the Mesozoic Chancellor and Lavantille Formations along the southern edge of the Northern Range (see Fig. 23A for location). Note the onlap of alluvium on the lower mountain front along the northern margin of the Northern basin. The projection of the offshore E1 Pilar fault zone seen on seismic lines projects into the southernmost strand of the E1 Pilar fault zone shown in alluvium of Northern basin by Kugler (1959).
538
S. BABB and E MANN
Fig. 29. Seismic line showing detachment surface between flat-lying Cretaceous rocks along the southwestern margin of the Gulf high and undifferentiated Lower Tertiary and Upper Cretaceous rocks and overlying Miocene-Pleistocene rocks of sequences 3-5 (see Fig. 23A for location). Movement may be a gravitational response to Tertiary sediment loading or tectonic dip-slip movement along the Gulf fault zone, a transtensional fault between the E1Pilar and Warm Springs strike-slip fault zones. inward beneath the ranges is also consistent with the regional cross-sections of Persad (1984) (Fig. 7).
STRUCTURAL STYLES AND AGE OF DEFORMATION IN THE SOUTHERN BASIN Introduction
The Southern basin of Trinidad has been well mapped during petroleum exploration and offers a wealth of structural and stratigraphic information that needs to be reconciled with any regional tectonic model (Fig. 37). The two cross-sections in
Fig. 38 are based on wells drilled in the Barrackpore and Penal oil fields in the Southern basin (Dyer and Cosgrove, 1992) (see locations of lines on map in Fig. 37). The thrust faulted margin in both areas is similar to that interpreted in the section of Fig. 35B along the northern margin of the Central Range. A major difference is that this Southern basin deformation is southward verging and begins earlier than that of the Northern and Gulf of Paria basin deformation, because Middle Miocene ('~ 11.4 m.y. horizon Gg.7) units onlap onto slightly older Middle Miocene units (Dyer and Cosgrove, 1992; Fig. 38). The unconformity between Late Miocene-Early Pliocene sequence 3 and middle to Late Pliocene
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
539
sequence 4 in the Northern and Gulf of Paria basins correlates with a less prominent unconformity between the Upper Cruse and Lower Forest Formations of southern Trinidad (Fig. 8). The two cross-sections in Fig. 38 document the older Middle Miocene tectonic phase in the Southern basin. However, this deformation does not appear to be a regional, low-angle thrust event as envisioned by Kugler (1953, 1959) and Speed (1985). Correlation of unconformity at the base of sequence 3 with an unconformity in the Southern basin of Trinidad
Based on the similarity in age, the shallowing event defining the base of sequence 3 is correlated with the major Middle Miocene unconformity documented in Trinidad's Southern basin (Fig. 8). The structural sections of Dyer and Cosgrove (1992) show the onlap and thinning of Late MiocenePliocene sediments on the Middle Miocene unconformity (defined by the top of the Globorotalia mayeri-Gg 7 biozone). The folding during the Middle Miocene has been interpreted to be the result of convergence in response to arrival of the Lesser Antilles arc and accretionary prism along this part of the South American passive margin (Dyer and Cosgrove, 1992; Algar and Pindell, 1993).
DISCUSSION Summary of Middle to Late Miocene-Pleistocene fault evolution in the Trinidad area
A summary of the Middle Miocene to Late Pliocene-Pleistocene fault evolution of the Gulf of Paria and western Northern basin is given in Fig. 39. The fault and depocenter maps are based on the original structural contour and isochron maps shown in Figs. 21, 22, and 23 and summarized in Fig. 24. Middle Miocene During this interval reverse faults and loading affected the area of the Southern basin (Figs. 38, 39A). This deformation is attributed to the earliest transpressional effects related to the oblique collision of the Caribbean arc with the northern passive margin of South America (Pindell and Barrett, 1990). We do not interpret this event as a regional folding and thrusting event as envisioned by Kugler (1959), Speed (1985), and, more recently, Flinch et al. (Chapter 17), because we do not observe pervasively folded and thrust rocks at the level of sequence 2 in the Gulf of Paria and Northern basins. Instead, we observe a relatively flat-lying basement (defined at the reflector we take to be top of the carbonate
Fig. 30. Seismic line showing east-west-striking transtensional fault trough between the Gulf and Domoil fault zones (see Fig. 23A for location). Note evidence for folding within sequences 3 and 4. megaplatform (cf. Figs. 11, 12; Babb, 1997). Some regional lines such as the one across the northwestern edge of the Goodrich sub-basin do show thrust faulting (Fig. 13), but these faults appear localized like the ones seen in the Southern basin (Fig. 38). Because of these observations, we conclude that the Northern, Gulf of Paria and Southern basins of Trinidad are unaffected by major contractional deformation until the onset of transpressional and transtensional faulting associated with discrete strike-slip fault zones like the E1 Pilar and Warm Springs in the Late Miocene. Prior to a relatively subtle Middle Miocene folding and faulting event (Fig. 38), Trinidad was characterized by passive margin sedimentation and subsidence. Algar and Pindell (1993) have proposed that uplift of the Central and Southern Ranges in the Middle Miocene may have been related to the peripheral bulge formed as the
540
S. BABB and R MANN
Fig. 31. North-northwest-striking seismic line across the Warm Springs fault zone and across the southwestern flank of the Domoil high in the southern Gulf of Paria basin (see Fig. 23A for location). The Warm Springs fault zone exhibits a negative flower structure zone characteristic of a slightly transtensional strike-slip fault.
Caribbean arc approached the subducting passive margin of northeastern South America. We can offer no new observations to support or contradict their bulge hypothesis. The Tamana limestone, a reefal unit formed at sea level in the Middle Miocene (Robertson and Burke, 1989; Erlich et al., 1993), marks the appearance of
the Central Range to sea level. From this time, the Central Range divided the Northern and Southern basins which exhibit differing stratigraphies as seen on the regional cross-section in Fig. 7. The uplift of the Central Range may be linked to an early phase of folding seen by Middle Miocene time in the Southern basin (Dyer and Cosgrove, 1992; Fig. 38)
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
541
Fig. 32. North-south-striking seismic line across the Warm Springs fault zone exhibiting a negative flower structure profile characteristic of a slightly transtensional strike-slip fault (see Fig. 23A for location) and Northern basin (Fig. 13) and/or to movement along the Warm Springs-Central Range-Caigual strike-slip fault zone (Fig. 33).
Early Late Miocene During this period a lobe of conglomeratic sediment was shed from the Northern Range and is interpreted to reflect significant lateral movement along the E1 Pilar fault zone (Fig. 39B). We concur with the proposal of Robertson and Burke (1989) that the fault is propagating from west to east during this time as the accretionary wedge and forearc of the Lesser Antilles moved past the South American
passive margin (Fig. 1B). The Northern Range, a complexly metamorphosed and folded block of passive margin lithologies, was also moving across the area at this time after being deformed and metamorphosed in a setting far to the west (Burke, 1988; Av6 Lallemant, 1997). Basin topography during this time was heavily influenced by the presence of remnant topography inherited from Early Cretaceous, northeast-facing banks formed during the passive margin phase. Strike-slip faulting is largely concentrated on the E1 Pilar zone and has not spread southward to the Warm Springs fault zone or Central Range (Fig. 39B).
542
S. BABB and P. MANN
Fig. 33. Geologic map of the Central Range and Warm Springs-Central Range-Caigual fault zones compiled from a variety of sources discussed in text. The Los Bajos fault is thought to have propagated southeastward in Late Miocene time as marked by the subsidence of the North Soldado sub-basin at its junction with the Warm Springs fault zone (Radovsky and Iqbal, 1979; Tyson et al., 1991). Termination of this fault by thrusting on the Southern Range widens the zone of Neogene deformation. The Orinoco delta begins to fill the Southern basin and offshore Columbus basin of Trinidad at this time (Leonard, 1983).
Late Miocene-Early Pliocene During this interval of deposition of sequence 3 in the Northern and Gulf of Paria basins, there was rapid increase in strike-slip and oblique-slip faulting to the south along the Warm Springs fault zone and in the Goodrich sub-basin (Fig. 39C). This southward shift is interpreted as the development of a pull-apart basin at the stepover area between the E1 Pilar and Warm Springs fault zones in the Gulf of Paria area. Transtensional opening of the pull-apart basin transfers slip from the E1 Pilar to the Warm Springs fight-lateral faults and deactivates the previously active, eastern continuation of the E1 Pilar in the Northern basin (Fig. 39C). This conclusion differs from Robertson and Burke's (1989) interpretation that the eastern E1 Pilar fault remains active and is possibly accelerating in slip rate. Transpressional uplift of the Central Range accompanies increased displacement along the Warm Springs-Central Range-Caigual fault zones (Fig. 33). Increased displacement also led to transten-
sional widening of the Goodrich sub-basin and an influx of sequence 3 sediments into the rifted area.
Early to middle Pliocene This interval continued the subsidence and fault pattern of the previous interval as motion continued to be transferred from the E1 Pilar to the Warm Springs system (Fig. 39D). Faults became wider spaced and more extensive in the Goodrich sub-basin and may reflect the increased extension related to pull-apart opening. Subsidence of the North Soldado basin (Radovsky and Iqbal, 1979; Tyson et al., 1991) reflected southeastward propagation of the Los Bajos fault beyond its junction with the Warm Springs fault zone. We believe that the primary sediment supply to the Goodrich pull-apart basin during this interval is the Central Range because clinoforms generally dip to the northwest away from that feature and because the area of south-central Trinidad formed a long-lived, broad high which restricts the inflow of south-derived Orinoco River sediments (Fig. 7). However, it is possible that the Orinoco River delta and sources to the north and northwest also are contributing to the infilling of the basin. Late Pliocene-Pleistocene A major wedge of sediment derived from uplift and erosion of the Central Range along with sediment from the north, northwest, and possibly the
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
543
Fig. 34. (A) Aerial photograph of the Warm Springs fault zone area at the western coast of Trinidad (see Fig. 23A for location). (B) Interpretation of photograph in (A) showing transition between offshore transtensional Warm Springs fault zone and onshore transpressional Warm Springs-Central Range-Caigual fault zone. Thick Early Pliocene-Pleistocene rocks of sequences 3-5 deposited by erosion and northward progradation along the eastern margin of offshore Goodrich sub-basin are exposed by monoclinal uplift and folding along the northern edge of the Warm Springs-Central Range-Caigual fault zone.
544
S. BABB and E MANN
Fig. 35. Two contrasting interpretations of a north-south cross-section of the southwestern end of the Central Range adjacent to the Gulf of Paria. Key to numbered formations the same as shown on compilation map by Kugler (1953): 1 - 5 -- Cretaceous; 6 = Cretaceous fragments; 7-12 = Tertiary. (A) Kugler (1953) proposed that the Central Range was an asymmetric, south-southeast-verging Late Neogene fold-thrust belt based on outcrop mapping. Note that his interpretation of thrust faults along the northern flank of the Central Range is problematic because younger rocks are thrust over older rocks. (B) Reinterpretation of the outcrop pattern based on offshore profiles of the Warm Springs fault zone shown in Figs. 31 and 32 suggests that the Central Range is a symmetrical, late Neogene restraining bend structure along transpressional strike-slip faults. Note that the high-angle Warm Springs fault zone thrusts older rocks of the Central Range over younger rocks along the northern flank of the Northern basin and that the predicted depth to the top of the Mesozoic carbonate platform is closer to the 3.5 km depth known from offshore and onshore seismic lines than the deeper depth depicted in the Kugler (1953) section. Orinoco River delta continues to infill the Goodrich sub-basin (Fig. 39E). Tyson et al. (1991) suggested that the North Soldado sub-basin continued to be active and recorded southeastward propagation of the Los Bajos fault which attained a total lateral offset of 10.5 k m on fold axes of the Southern basin (Fig. 37). Activity on the E1 Pilar system to the east in the Northern basin consisted of transpressional reactivation (Fig. 26) perhaps related to the slightly north of east strike of the fault in this area (Fig. 23A). Activity on the part of the east-west-striking W a r m Springs fault in the offshore area continued to be
transtensional as seen by the negative flower structure in Fig. 32. Continued offset on the Los Bajos fault (Wilson, 1940) and linked thrusting in the Central Range helped the deformation belt to widen in southern Trinidad.
How similar is the geology of the Eastern Venezuela basin and Trinidad? A fundamental question about the tectonic evolution of Trinidad is whether its structure, stratigraphy and age of deformation is identical to that seen in
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
545
Fig. 36. North-south-striking, onland seismic line across the east-central part of the Northern basin (see Fig. 33 for location) showing a sub-vertical strike-slip fault zone as postulated in the interpretation of Fig. 35B and transpression-related folding of Late PleistoceneRecent age that affects Late Miocene to Pleistocene sequences 3-5. Note evidence for Late Pleistocene-Recent transpressional reactivation of the E1 Pilar fault zone along the northern edge of the basin. A similar style of transpressional reactivation of the E1 Pilar fault zone is seen in the seismic line from the northern Gulf of Paria that is shown in Fig. 26.
the Gufirico and Maturfn sub-basins of the Eastern Venezuelan basin to the west (Fig. 4B). This question was originally posed in the title of the 1965 paper by Salvador and Stainforth entitled 'Clues in Venezuela to the geology of Trinidad, and vice versa' and deserves to be reexamined in the light of the new data presented in this paper. Some authors like Speed (1985) and Russo and Speed (1992) have emphasized that the same set
of thrust-related structures of the Maturfn sub-basin extend to the east-northeast across the Gulf of Paria and into Trinidad, albeit with a slightly younger age of deformation in Trinidad (Fig. 4A). Other workers like Robertson and Burke (1989) and Erlich and Barrett (1992) have emphasized that faults in Trinidad are dominantly strike-slip faults and therefore differ in character from the predominant thrust faults seen in the fold-thrust belts and associated foredeep
546
S. BABB and E MANN
Fig. 37. Geologic map of the Southern Range compiled from Kugler (1953) and other sources discussed in the text.
basins of the Eastern Venezuelan basin (Erlich and Barrett, 1992; Passalacqua et al., 1995; di Croce et al., Chapter 16). The purpose of this section is to point out some differences in the geologic character of the two areas derived from comparison of regional cross-sections shown in Fig. 40. The structure and stratigraphy of the Maturfn sub-basin is summarized in the form of a regional cross-section modified from Parnaud et al. (1995) (Fig. 40B) while the structure and stratigraphy of Trinidad is summarized in the form of a regional cross-section modified from Persad (1984) and Robertson and Burke (1989) (Fig. 40A). The horizontal and vertical cross-sections of the two sections are approximately the same to facilitate comparison.
Comparison of the main stratigraphic features of the Eastern Venezuelan basin and Trinidad area in cross-section The regional cross-sections of eastern Venezuela (Fig. 40B) and Trinidad (Fig. 40A) reveal the following stratigraphic information about this part of the passive margin of northern South America.
Precambrian basement Basement rocks of the Precambrian Guyana shield can be seen dipping northward beneath the Maturin foreland basin and the Serranfa del Interior fold-thrust belt (di Croce et al., Chapter 16) (Fig. 40B). A peripheral bulge related to its flexure beneath the overthrust belts is located near the northern margin of the Guyana shield (Fig. 40B). In Trinidad pre-Jurassic crystalline basement
rocks are not exposed, have not been encountered in wells, and have not been seen on seismic reflection lines (Fig. 40A). At depth beneath the carbonate passive margin section of Trinidad seen on the regional lines of Figs. 11 and 12, basement is possibly a thinned equivalent of the Precambrian crust cropping out in the Guyana shield (Fig. 40B).
Late Jurassic rift phase Late Jurassic half-grabens are present at depths of 1-2 km in the zone of thin foreland sedimentary cover near the Guyana shield (Fig. 40B). These rifts record a phase of opening between South America and YucaUin Peninsula in Late Jurassic time (Pindell and Barrett, 1990; Mann, Chapter 1). In Trinidad no rifts of Jurassic age have been drilled or observed on seismic reflection lines (Burke, 1988; Fig. 40A). Rifts are possibly present beneath the thick Early Cretaceous carbonate and siliciclastic rocks of the passive margin section described by Babb (1997) beneath the Gulf of Paria and seen on the regional lines of Figs. 11 and 12. Latest Jurassic-Paleogene passive margin phase In the Eastern Venezuelan basin, a thick sequence of passive margin rocks are seen on seismic lines dipping northward beneath the foreland basin (Fig. 40B). More distal equivalents of these Early Cretaceous-Eocene units have been imbricated in the Serranfa fold-thrust belt that is described in detail by Parnaud et al. (1995), Roure et al. (1995) and di Croce and Bally (Chapter 16). In both Venezuela and Trinidad the passive margin history can be subdivided into two phases:
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
547
Fig. 38. Cross-section based on wells of the Barrackpore (A) and Penal (B) oil field within the Southern basin (from Dyer and Cosgrove, 1992; see Fig. 37 for location). Note the steepness of the faulted margin similar to that shown in the line of Fig. 29 along the northern margin of the Central Range and also note onlapping of late Middle Miocene (~11.4 m.y. horizon Gg.7) onto slightly older Middle Miocene units. The unconformity between Early Pliocene sequence 3 and middle to Late Pliocene sequence 4 in the Northern basin would correlate to a less prominent unconformity below the top middle Cruse horizon.
(1) An early shallow carbonate bank phase from latest J u r a s s i c ? - M i d d l e Cretaceous. This shallow water bank was d r o w n e d in both areas by the end of the Early C r e t a c e o u s (Aptian in Trinidad and Albian in eastern Venezuela; Babb, 1997). It is the irregular top
of this b a n k that forms the surface b e n e a t h the G u l f of Paria and Northern basins that was infilled by siliciclastic sediments during the C e n o z o i c (Fig. 3 9 B - E ) . The shallow carbonate bank rocks are not well e x p o s e d in eastern Venezuelan b e c a u s e they have
548
S. BABB and E MANN
Fig. 39. Structural evolution of sub-basins in the Gulf of Paria based on sequence and structural mapping described in this study. (A) Middle Miocene. (B) Early Late Miocene. (C) Late Miocene-Early Pliocene (sequence 3). (D) Early-middle Pliocene (sequence 4); (E) Late Pliocene-Pleistocene (sequence 5). See text for discussion.
been overthrust in the Serranfa del Interior and are now deeply buried beneath the Maturfn basin (Figs. 4B, 40B). Evaporitic rocks are present in the Cretaceous carbonate section beneath the Gulf of Paria basin of Trinidad (Bray and Eva, 1983; Eva et al., 1989; Babb, 1997) and have been described in small sections from the Araya-Paria Peninsula (Gonzalez de Juana et al., 1968). The passive margin section of Trinidad is locally within 1-3 km of the surface and is known through drilling and seismic reflection (Fig. 40A). General rock types and depositional environments are compiled onto the stratigraphic columns shown in Fig. 41. (2) A later deeper-water passive margin phase
from Late Cretaceous through Paleogene characterized by both carbonate and siliciclastic units that are generally thin. The northward thinning of these units is shown on the regional cross-section of Trinidad in Fig. 40A because the overall amount of shortening related to transpressional faulting and uplifted ranges like the Central and Southern Ranges is small. In contrast, the large amount of shortening in the Serranfa del Interior make recognition of a similar pattern of northward thinning of passive margin rocks impossible in eastern Venezuela (Fig. 40B). The most obvious difference between the passive margin sections of Trinidad and eastern Venezuela
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
549
~o ~ o
0 0 0
0
0 .0
=
0
.o
~
oz~ 0
,~
0
~
~
.~ .
> .~ "a "~, ~
~
0
0.-~
0
~
~o~ ~"~
0
~9o ~ -~~
c~"~ ~'~
9=
550
S. BABB and R MANN
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
551
Fig. 42. Map showing major structural features of the northern margin of South America in Venezuela and Trinidad and the ages of convergent deformation along strike of the margin. The development of northwest-striking right-lateral strike-slip faults in the Trinidad area may be related to oceanic free face or lack of a continental backstop between the eastward moving Caribbean plate and arc terranes and the South American continent. See text for discussion. is their degree of deformation (Fig. 40). In eastern Venezuela, the margin is intensely imbricated in a fold-thrust style with 4 5 - 9 0 km of horizontal shortening while the strata in the Gulf of Paria are generally horizontal in attitude and deformed mainly by late Neogene strike-slip faults.
Oblique collision and foreland basin phase In eastern Venezuela a foreland basin (Maturfn sub-basin; Fig. 4B) formed as a result of thrusting in the Serranfa del Interior fold-thrust belt in Late Oligocene through Pliocene time (Parnaud et al., 1995; Roure et al., 1995; di Croce and Bally, Chapter 16). Foreland basin sedimentary rocks are mainly marine in the early stages and shallow marine to non-marine in the later stages of thrusting as the Orinoco River delta prograded eastward (Robertson and Burke, 1989; Erlich and Barrett, 1992) (Fig. 41). In Trinidad the earliest evidence of folding, limited thrusting and shallowing in the Central Range and Southern basin is Early Miocene in age. During the Early to Middle Miocene, the passive margin appears to convert to an active margin characterized by somewhat shallower depositional environments and an increasing amount of structural control on deposition (Figs. 39 and 41).
Possible explanation for structural differences between Trinidad and eastern Venezuela One possible difference to explain the lack of thrusting in Trinidad is the lack of Precambrian continental crust of the Guyana shield which acts as a 'backstop' southeast of the fold-thrust belt (Fig. 42). Instead, Trinidad appears to occupy the site of a rifted passive margin characterized by the northwest-trending carbonate highs formed in latest Jurassic to Early Cretaceous time (Fig. 21). This thinner crust acts as less of a barrier or backstop and provides an unconfined or less confined 'free face' for the strike-slip movement of crustal material to the southeast in a direction roughly parallel to the trend of the former, northeast-facing passive margin (Fig. 42). Transpression between the Caribbean and South America plates was accommodated by now extinct thrusting in the confined area to the west but now can be relieved through southeastward propagation of strike-slip faults without thrusting over the continent-ocean transition in Trinidad and to the south in the Orinoco delta region and along the Guyana margin (Jankowsky and Schlapak, 1983). Northwest-oriented basement fracture zones in the Trinidad passive margin (possibly reflected in northwest-trending carbonate banks as seen on the map of the carbonate platform in
Fig. 41. Chart comparing the Jurassic to Recent stratigraphy of the Northern basin, Gulf of Paria, and Southern basin of Trinidad and the Eastern Venezuelan basin. Striped pattern on older units designates mainly neritic or deep-water marine units deposited in a passive margin setting. Uncolored pattern on younger units designates mainly shallower water units deposited after the formation of an oblique collisional boundary. The darker pattern in the Eastern Venezuelan basin designates the Carapita Formation, whose deep marine environment is thought to have accompanied the flexural formation of the foreland basin (Parnaud et al., 1995). Note that Oligocene foreland basin subsidence in the Eastern Venezuelan basin precedes the main pulse of Late Miocene to Recent subsidence in the Northern basin and Gulf of Paria in Trinidad and is therefore consistent with the eastward motion of the Caribbean plate along the northern passive margin of South America.
552 Fig. 21) may also act to reorient strike-slip faults from east-west to more east-southeast or southeast striking. Movement of material along these northwest-striking faults that possibly follow deep-seated basement fracture zones results in the thick-skinned deformational style of Trinidad that is distinctive from the thin-skinned deformational style seen in eastern Venezuela (Fig. 40). Thicker crust to the east fronting the Foothills and Serranfa del Interior fold-thrust belts is a backstop that does not allow southeastward translation of crust to the southeast (Fig. 42). The subsequent space problem that resulted as the edge of the Caribbean plate attempted to move southeastward created the thin-skinned Foothills and Serranfa del Interior fold-thrust belts in Oligocene to Pleistocene time. As a result, there is an extensive foreland basin in eastern Venezuela that formed as the continental crust flexed under the load of
S. BABB and E MANN the southeastward-moving thrust sheets (Fig. 4B). In contrast, basins of Trinidad are much less extensive and confined as steep-sided basins between high-angle strike-slip or transpressional fault zones (Fig. 40). The driving force for this southeastward movement of material along faults like the Warm Springs, Los Bajos, and E1 Soldado (Fig. 4A) is the long-term eastward translation of the Caribbean arc along the passive margin of South America (Fig. 43). This oblique collisional process now active in Trinidad and active as recently as the Pleistocene in eastern Venezuela has progressively affected the passive margins of North and South America with Late Paleocene thrust deformation documented in Lake Maracaibo of western Venezuela (Lugo and Mann, 1995) and Late Paleocene-Early Eocene deformation affecting western Cuba (Gordon et al., 1997) (Fig. 43A-F).
Fig. 43. Map showing the west to east migration during Late Cretaceous to Recent time of the eastern margin of the Caribbean plate into the space formed by the Mesozoic opening between North and South America. (A) Campanian-Maestrichtian. (B) Latest Paleoceneearliest Eocene. (C)Latest Early Eocene. (D) Middle Eocene'Middle Miocene. (E) Middle Miocene. (F) Post-Middle Miocene. The position of the first interaction of the Caribbean plate in the southeastern Caribbean is taken as Oligocene in the Eastern Venezuelan basin and Late Miocene in Trinidad. See text for discussion.
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY
Rocks of the Cordillera de la Costa belt These rocks are present in both the Araya-Paria Peninsula north of the Serrania del Interior foldthrust belt and in the Northern Range of Trinidad. The rocks of the Araya-Paria area exhibit both high- and low-grade metamorphic grades and ductile structures including southeast-verging isoclinal folds and thrust faults (Fig. 40B). Av6 Lallemant (1997) interpreted these structures to have formed in a trench setting far to the west of their present position (Fig. 43A). Rocks of the Northern Range are all low metamorphic grade and exhibit northward vergence (Algar and Pindell, 1993) that is in the opposite direction to that predicted by Speed (1985) for late Neogene fold-thrust belts (Fig. 40A). As in the case of the Araya-Paria Peninsula, it is likely that the Northern Range was juxtaposed with the areas of Trinidad's basins by strike-slip faulting and originally formed in a location far to the west of its present position as shown on the plate reconstruction of Fig. 43. Tectonic implications of Venezuela-Trinidad comparison Differences in the age and style of deformation seen in Trinidad and eastern Venezuela provide helpful constraints on the large-scale tectonic processes affecting the plate boundary zone.
Eastward movement of Caribbean plate First, the transition from passive to active margin tectonics, marked by abrupt shallowing in sedimentary environments and the formation of foreland basins flanking thrust belts (Fig. 4B), is about 17-20 m.y. later in Trinidad than in eastern Venezuela. Units on the stratigraphic chart in Fig. 41 illustrating the transition between passive and active margin tectonics include the following: (1) In Trinidad, the main transition is proposed to occur during the deposition of the conglomeratic wedge from the Northern Range (sequence 3, Late Miocene-Early Pliocene). While there are other anomalous episodes of shallowing (e.g., Middle Miocene shallowing of the Central Range), the deposition of sequence 3 marks the first radical departure from passive margin sedimentation on a north-facing slope. (2) In eastern Venezuela, the main transition is proposed to occur during the deposition of marine shales in the foreland basin (Oligocene; Fig. 41). Foredeep subsidence continued through the Miocene and only becomes inactive in the Pleistocene (Fig. 4B). The delay between the Late Miocene Trinidad event and the Oligocene Eastern Venezuela event is consistent with the slow (1-2 cm/yr) but steady
553
west-to-east motion of the Caribbean plate along the northern margin of Trinidad as shown in Fig. 42. Complications in the age of deformation along the margin may relate to irregularities in the shape of the Jurassic-Cretaceous passive margin which we have shown to have significant vertical relief in the subsurface beneath the Gulf of Paria and Northern basins. The age of deformation could be varying in a complex way as the arc encounters such a margin with variable topographic relief and structural trends.
CONCLUSIONS The main conclusions of this study can be summarized as follows. (1) An integrated data base consisting of well logs and seismic reflection profiles mainly from the Gulf of Paria and Northern basins was used to document major tectonic and sedimentary events in the Neogene geologic history of Trinidad. (2) Previous models proposed to explain the tectono-stratigraphic evolution of the area can be divided into three categories: (a) models in which oblique collision and thrusting are the dominant process; (b) models in which strike-slip faulting and related normal faulting, folding, and thrusting are the dominant tectonic process; and (c) transpressional models which combine elements of thrusting and strike-slip. The conflicting tectonic models reflect varying emphasis on methods, data types, and local areas of Trinidadian geology. (3) For this study, the term 'Gulf of Paria basin' is used to describe all the sedimentary rocks in the offshore Gulf of Paria that overlie latest Early Cretaceous rocks of the passive margin of northern South America. The 'Northern basin' refers to the equivalent sedimentary succession of sedimentary rocks with the onshore area of northern Trinidad ('Caroni basin' of Robertson and Burke, 1989). Mapping of seismic sequences on seismic reflection lines tied to wells revealed the presence of four fault-controlled sub-basins within the Gulf of Paria basin and western Northern basin. These are the E1 Pilar, Puerto Grande, and Goodrich sub-basins which contain Neogene sedimentary sequences ranging in two-way travel time from 2.0 to 3.3 s. (4) Five seismic sequences were recognized in the Gulf of Paria and Northern basins based on the recognition of onlap, toplap, downlap and truncation surfaces, and seismic reflection characteristics and the correlation of seismic reflection characteristics with well logs in the Gulf of Paria and Northern basins. Sequences 1 and 2 consist of a Late Jurassic to Middle Cretaceous carbonate megaplatform with overlying carbonate banks that form the basement to sequences 3 through 5 of Neogene age. Sequence
554 3 is a Late Miocene-Early Pliocene shallow-marine to brackish-water conglomerate and sandstone that represents a southward-fining and thinning and eastward-thickening siliciclastic wedge derived from the Late Miocene uplift and erosion of the Northern Range and activity along the E1 Pilar strike-slip fault close to the mountain front of that range. Sequence 3 and the overlying sequences 4 and 5 fill in topographic relief created by the growth of the carbonate banks of sequences 1 and 2 that was later modified by extensional structures formed in the Gulf of Paria pull-apart basin. Sequence 4 is an Early to middle Pliocene inner neritic to shallow-marine conglomerate, sandstone, silt, and clay that represents a northward-fining and thinning siliciclastic wedge derived from the Late Miocene-Early Pliocene uplift and erosion of the Central Range with additional contributions from the north and west and possibly the Orinoco River delta. Sequence 5 is a Late Pliocene to Pleistocene marine to brackish-water sand, silt, clay and minor conglomerate that represents continued siliciclastic deposition in an increasingly restricted basin between the Northern and Central Ranges. (5) Comparison of structure contour maps of each boundary between sequences 3-5 with isochron maps of sequences 3-5 allowed a better visualization of the space available at tectonically significant times in the history of the Gulf of PariaNorthern basin and the nature of the subsequent sedimentary deposition onto those structural surfaces and sequence stratigraphic boundaries. This crosscomparison identified three Neogene deformational phases that have affected the area: phase one, Late Miocene-Early Pliocene strike-slip motion along the E1 Pilar fault and north-to-south filling of the Gulf of Paria and Northern basins; phase two, middle to Late Pliocene strike-slip motion along the Warm Springs-Central Range fault zone and south-to-north filling of the Gulf of Paria and Northern basins; and phase three, Late Pliocene to Pleistocene strike-slip motion along the Warm Springs-Central Range fault zone and continued south-to-north filling of the Gulf of Paria and Northern basins. (6) Paleocurrents and environments of deposition of the Gulf of Paria and Northern basins were closely controlled by structural events during the three deformational phases. Paleocurrents for deformational phase one (deposition of sequence 3) are oriented orthogonal and parallel to the axis of the basins and reflect the strong control of the bounding E1 Pilar fault zone along the northern edge of the basin. Paleocurrents for deformational phases two (sequence 4) and three (sequence 5) become increasingly dominated by a west-to-east paleoflow along with a southeast-to-northwest flow in the area of the Goodrich basin. Environments of deposition reflect the uplift and erosion of the Northern Range during
S. BABB and E MANN phase one followed by uplift and erosion of the Central Range and the gradual decrease in the Northern Range siliciclastic sources during phases two and three. Environments generally remained proximal, shallow and brackish along the E1 Pilar fault zone on the northern edge of the basin and distal, deeper and less restricted along the central and southern edges of the basin. Through time the basin becomes increasingly brackish and more restricted. (7) Seismic lines were used to illustrate the variation in the structural style and geomorphic expression of the E1 Pilar fault zone along the northern edges of the Gulf of Paria-Northern basin. In the western Gulf of Paria, the fault exhibits a transpressional structure consistent with its local, slightly east-northeast strike. A slight change in strike to a more east-southeast strike in the eastern Gulf of Paria is consistent with the more transtensional structure of the fault and its association with the Puerto Grande sub-basin. The trace of the E1 Pilar fault zone is poorly expressed onland in Trinidad and may reflect its waning activity in recent time in an eastward direction and/or the overprinting cultural effects on low scarps in this densely populated region. (8) Seismic lines were also used to illustrate the structural style of oblique-slip faults in the Gulf of Paria. These styles include gravity-related detachment faults possibly localized along the unconformity between Cretaceous carbonate and overlying siliciclastic rocks and tectonic transtensional faults formed in the stepover area between the E1 Pilar fault to the north and the Warm Springs fault to the south. (9) Seismic lines were used to illustrate the structural styles and geomorphic expression of the Warm Springs-Central Range-Caigual fault zone. In the Gulf of Paria, the Warm Springs fault zone exhibits a transtensional structure consistent with its slightly east-southeast strike. An abrupt change in strike to a more east-northeast strike near the western shoreline of Trinidad is consistent with the more transpressional structure of the continuation of the Warm Springs fault zone in the Central Range (Central Range-Caigual fault zones). On the basis of seismic lines and existing geologic maps, we reinterpreted the overall structure of the Central Range as a transpressional uplift bounded by inwardly dipping reverse faults and bisected by a strike-slip fault system rather than as the southeast-verging fold-thrust belt proposed by previous workers. (10) Compilation of data by previous workers in the Southern basin of Trinidad allows us to make a comparison between the style and age of deformation in the Southern basin and that described by us in the Gulf of Paria-Northern basin. The style of deformation in the Southern basin is dominantly south-vergent and starts earlier than in the Gulf
DEVELOPMENT OF A NEOGENE TRANSPRESSIONAL PLATE BOUNDARY of Paria-Northern basins because Middle Miocene (~ 11.4 m.y. horizon) units onlap slightly older Middle Miocene units (Fig. 38). The unconformity between the Late Miocene-Early Pliocene sequence 3 and the middle to Late Pliocene sequence 4 in the Northern-Gulf of Paria basin correlates with a less prominent unconformity between the Upper Cruse and Lower Forest Formations of southern Trinidad. The deformation in the Southern basin begins earlier than most of that observed in the Gulf of PariaNorthern basin. (11) Compilation of regional geologic data and integration with our results from the Gulf of PariaNorthern basins allows the following events to be constrained in the Middle Miocene through Recent evolution of the area. Middle Miocene: high-angle faulting affects the Southern basin; we do not regard this event as a regional fold and thrusting event because we do not observe widespread foldthrust structures in the pre-Late Miocene rocks of the Gulf of Paria-Northern basin. Instead, we conclude that most deformation observed in Trinidad is transpressional or transtensional in style, related to movements on the E1 Pilar, Warm Springs-Central Range-Caigual, and Los Bajos fault zones, and is Late Miocene to recent in age. Late Middle Miocene: the uplift of the Central Range at this time may be related to the Middle Miocene thrusting observed in the Southern basin. Early Late Miocene: during this period a lobe of conglomeratic sediment is shed from the Northern Range into the Gulf of PariaNorthern basins and is interpreted to reflect the first significant lateral movement along the E1 Pilaf fault zone. We suggest that the fault propagated from west to east during this time as the accretionary wedge and forearc of the Lesser Antilles arc moved past the South American passive margin. The Los Bajos fault is thought to have propagated south by this time. Termination of this fault on the Southern Range may have led to a widening of the Neogene belt of transpressional thrusting. Late Miocene-Early Pliocene: during this time, there was a rapid spread of strikeslip and oblique-slip faulting to the south of the E1 Pilar fault zone along the Warm Springs fault zone and within the Goodrich sub-basin. This southward shift is interpreted as the development of a pull-apart basin at the stepover area between the E1 Pilar and Warm Springs fault zones. Formation of the stepover may reduce slip along the eastern continuation of the E1 Pilar fault zone at the northern margin of the Northern basin. Early to middle Pliocene: this interval saw the continued subsidence and fault pattern of the previous interval in the stepover area between the two faults. Late Pliocene to Pleistocene: a major wedge of sediment derived from erosion of the Central Range with probable contributions from the north and west and possibly from the Orinoco
555
delta filled the Goodrich basin and continued southeastward propagation of the Los Bajos fault offsets fold axes in the Southern basin by 10.5 kin. The E1 Pilar fault zone appears to undergo transpressional reactivation in the Northern basin but most offset seems to have shifted to the Warm Springs-Central Range-Caigual fault zone in the Central Range. (12) We compared the above sequence of events and structure of Trinidad with the events and structure of the Eastern Venezuelan basin in order to infer large-scale tectonic controls on regional deformation. The Eastern Venezuelan basin is a foreland basin associated with an adjacent, southeast-verging fold-thrust belt (Serranfa del Interior) associated with 45-90 km of shortening. Trinidad exhibits a slightly younger record of fault-related sedimentation, lacks evidence for significant shortening of the Cretaceous-Pleistocene sequences 1-5 documented in this paper, and exhibits transpressional-type structures with no uniform sense of vergence. We propose that Precambrian crust adjacent to the Eastern Venezuelan basin acted as an effective backstop and did not allow continued southeastward migration of obliquely colliding Caribbean crust. In contrast, Trinidad appears to occupy the site of a rifted passive margin whose thinner crust acts as less of a backstop than the Precambrian crust to the southwest. This less confined 'free face' allowed strike-slip movement of Caribbean and passive margin crust to the southeast in a direction roughly parallel to the trend of the former passive margin.
ACKNOWLEDGEMENTS
This paper is a partial summary of a Ph.D. dissertation conducted by the first author from 1992 through 1997 at the University of Texas at Austin and supported by the Southern Basin Consortium, Petroleum Company of Trinidad and Tobago (Petrotrin), and the Institute for Geophysics of the University of Texas at Austin. We thank Petrotrin for release of these data for the dissertation study. Special thanks to Ph.D. committee members R. Buftier (co-chair), E. McBride (co-chair), W. Galloway, and N. Tyler for their suggestions on improving this study. UTIG contribution number 1425.
REFERENCES
Algar, S.T. and Pindell, J., 1993. Structure and deformation history of the Northern Range of Trinidad and adjacent areas. Tectonics, 12: 814-829. Av6 Lallemant, H., 1997. Transpression, displacement partitioning, and exhumation in the Caribbean/South American plate boundary zone. Tectonics, 16: 272-289. Babb, S., 1997. Tectonics and Sedimentation of the Gulf of Paria
556 and Northern Basin, Trinidad. Unpublished Ph.D. dissertation, University of Texas at Austin, 200 pp. Barr, K.W., 1952. Limestone blocks in the Lower Cretaceous Cuche Formation of the Central Range, Trinidad, B.W.I. Geol. Mag., 89:417-425. Barr, K.W. and Saunders, J.B., 1968. An outline of the geology of Trinidad. Trans. Fourth Caribbean Geol. Conf., Port of Spain, pp. 1-10. Bartenstein, H., Bettenstaedt, F. and Bolli, H.M., 1957. Die foraminiferan der Unterkreide von Trinidad, B.W.I., Erster Teil: Cuche- und Toco Formation. Eclogae Geol. Helv., 50: 5-67. Bartenstein, H., Bettenstaedt, F. and Bolli, H.M., 1966. Die foraminiferen der Unterkreide von Trinidad, B.W.I., Zweiter Teil: Maridale Formation (Typlokalitat). Eclogae Geol. Helv., 59: 129-178. Bray, R. and Eva, A., 1983. Age, depositional history, and tectonic significance of the Couva Marine evaporite, offshore Trinidad. Tenth Caribbean Geol. Conf., Cartagena, Colombia (Abstr. Vol.), p. 372. Burke, K., 1988. Tectonic evolution of the Caribbean. Annu. Rev. Earth Planet. Sci., 16:201-230. Carr-Brown, B. and Frampton, J., 1979. An outline of the stratigraphy of Trinidad. In: Field Guide to the Fourth Latin American Geological Congress, Port of Spain. Key Caribbean Publications, Trinidad, pp. 7-19. DeMets, C., Gordon, R.G., Argus, D.E and Stein, S., 1994. Effects of recent revisions to the geomagnetic time scale on estimates of current plate motions. Geophys. Res. Lett., 21: 2191-2194. Dixon, T.H., Farina, E, DeMets, C., Jansma, R, Mann, E and Calais, E., 1998. Relative motion between the Caribbean and North America plates and related boundary zone deformation based on a decade of GPS observations. J. Geophys. Res., 103: 15,157-15,182. Donovan, S.K., 1994. Trinidad. In: S.K. Donovan and T.A. Jackson (Editors), Caribbean Geology: An Introduction. University of West Indies, Kingston, pp. 209-228. Dyer, B.L. and Cosgrove, R, 1992. Penal/Barrackpore field West Indies, South Trinidad basin, Trinidad. In: E.A. Beaumont and N.H. Foster (Editors), Structural Traps VII. American Association of Petroleum Geologists, Tulsa, Okla., pp. 139-157. Erlich, R.N. and Barrett, S.E, 1992. Petroleum geology of the Eastern Venezuelan foreland basin. In: R.W. Macqueen and D.A. Leckie (Editors), Foreland Basins and Foldbelts. American Association of Petroleum Geologists, Tulsa, Okla., pp. 341-362. Erlich, R.N., Farfan, EE and Hallock, E, 1993. Biostratigraphy, depositional environments, and diagenesis of the Tamana Formation, Trinidad: A tectonic marker horizon. Sedimentology, 40: 743-768. Eva, A.N., Burke, K., Mann, R and Wadge, G., 1989. Four phase tectonostratigraphic development of the southeastern Caribbean. Mar. Pet. Geol., 6: 9-21. Figueroa de Sanchez, L. and Hern~indez, L., 1990. Exploraci6n geoffsica-geol6gica del ~irea de Guarumen. Trans. 5th Venezuelan Geophys. Congr., pp. 219-227. Frey, M., Saunders, J. and Schwander, H., 1988. The mineralogy and metamorphic geology of low-grade metasediments, Northern Range, Trinidad. J. Geol. Soc., London, 145: 563575. Furrer, M.A., 1968. Paleontology of some limestones and calcphyllites of the Northern Range of Trinidad, West Indies. Trans. Fourth Caribbean Geol. Conf., Port of Spain, pp. 2124. Gonzalez de Juana, C., Munoz, N.G. and Vignali, M., 1968. On the geology of the eastern Paria (Venezuela). Trans. Fourth
S. BABB and R M A N N Caribbean Geol. Conf., Port of Spain, pp. 25-29. Gonzalez de Juana, C., Iturralde de Arozena, J.M. and Picard, X., 1980. Geologia de Venezuela y de sus cuencas petroliferas. Ediciones Foninves, Caracas, 1031 pp. Gordon, M.B., Mann, P., C~iceres, D. and Flores, R., 1997. Cenozoic tectonic history of the North America-Caribbean plate boundary zone in Cuba. J. Geophys. Res., 102: 10,05510,082. Haq, B.U., Hardenbol, J. and Vail, P.R., 1987. The chronology of fluctuating sea level since the Triassic. Science, 235:11561167. Higgins, G.E. and Saunders, J.B., 1968. Excursion number 4, Western Central Range. Trans. Fourth Caribbean Geol. Conf., Port of Spain, pp. 435-437. Jankowsky, W.J. and Schlapak, G., 1983. Guyana offshore. In: A. Bally (Editor), Atlas of Seismic Stratigraphy. Vol. 2. American Association of Petroleum Geologists, Tulas, OK, pp. 2.2.3-47. Koutsoukos, E.A.M. and Merrick, K.A., 1985. Foraminiferal paleoenvironments from the Barremian to Maestrichtian of Trinidad, W.I. Trans. First Geol. Conf. Geol. Soc. Trinidad and Tobago, Port of Spain, pp. 85-101. Kugler, H.G., 1953. Jurassic to recent sedimentary environments in Trinidad. Bull. Assoc. Suisse Geol. Ing. Pet., 20: 27-60. Kugler, H.G., 1959. Geological map and sections of Trinidad (scale 1:100,000). The Petroleum Association of Trinidad, Port of Spain. Leonard, R., 1983. Geology and hydrocarbon accumulations, Columbus basin, offshore Trinidad. Bull. Am. Assoc. Pet. Geol., 67: 1081-1093. Lugo, J. and Mann, P., 1995. Jurassic-Eocene tectonic evolution of Maracaibo basin, Venezuela. In: A.J. Tankard, R. Su~irez and H.J. Welsink (Editors), Petroleum Basins of South America. American Association of Petroleum Geologists, pp. 699725. Mann, P., Schubert, C. and Burke, K., 1990. Review of Caribbean neotectonics. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Vol. H, Geological Society of America, Boulder, Colo., pp. 307-338. Michelson, J.E., 1976. Miocene deltaic oil habitat, Trinidad. Am. Assoc. Pet. Geol. Bull., 60:1502-1519. Molnar, P. and Sykes, L.R., 1969. Tectonics of the Caribbean and Middle America regions from focal mechanisms and seismicity. Geol. Soc. Am. Bull., 80: 1639-1684. Molnar, P. and Sykes, L.R., 1971. Plate tectonics in the Hispaniola area: discussion. Geol. Soc. Am. Bull., 80:1123-1126. Munro, S.E. and Smith, ED., Jr., 1984. The Utica fault zone, northeastern Venezuela. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162: 213-215. Parnaud, E, Gou, Y., Pascual, J.C., Truskowski, I., Gallango, O. and Passalacqua, H., 1995. Petroleum geology of the central part of the Eastern Venezuelan basin. In: A.J. Tankard, R. Su~irez and H.J. Welsink (Editors), Petroleum Basins of South America. Am. Assoc. Pet. Geol. Mem., 62: 741-756. Passalacqua, H., Fernandez, E, Gou, Y. and Roure, E, 1995. Crustal architecture and strain partitioning in the Eastern Venezuelan Ranges. In: A.J. Tankard, R. Sugtrez and H.J. Welsink (Editors), Petroleum Basins of South America. Am. Assoc. Pet. Geol. Mem., 62: 667-679. Payne, N., 1991. An evaluation of post-Middle Miocene geological sequences, offshore Trinidad. Trans. Second Geol. Conf. Geol. Soc. Trinidad and Tobago, Port of Spain, pp. 70-87. Perez, O.J. and Aggarwal, Y.R, 1981. Present-day tectonics of the southeastern Caribbean and northeastern Venezuela. J. Geophys. Res., 86:10791-10804.
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 Persad, K.M., 1984. Tectonic map and geologic-tectonic sections through Trinidad and Tobago (scale 1: 200,000). Robertson Research (U.S.), Houston, Texas. Persad, K.M., 1985. Outline of the geology of Trinidad. Fourth Latin Am. Geol. Congr., Port of Spain, pp. 738-758. Pindell, J.L. and Barrett, S.E, 1990. Geological evolution of the Caribbean region: a plate-tectonic perspective. In: G. Dengo and J.E. Case (Editors), The Caribbean Region. The Geology of North America, Vol. H, Geological Society of America, Boulder, Colo., pp. 405-431. Potter, H., 1976. Type sections of the Maraval, Maracas, and Chancellor Formations in the Caribbean Group of the Northern Range of Trinidad. Trans. Seventh Caribbean Geol. Conf., Guadeloupe, pp. 505-527. Prieto, R., 1987. Seismic Stratigraphy and Depositional Systems of the Orinoco Platform Area, Northeastern Venezuela. Unpublished Ph.D. Dissertation, University of Texas at Austin, 144 pp. Radovsky, B. and Iqbal, J., 1979. Geology of the North Soldado field. Fourth Latin Am. Geol. Congr., Port of Spain, pp. 759769. Robertson, E and Burke, K., 1989. Evolution of southern Caribbean plate boundary, vicinity of Trinidad and Tobago. Am. Assoc. Pet. Geol. Bull., 73: 490-509. Rohr, G.M., 1991. Paleogeographic maps, Maturfn basin of Eastern Venezuela and Trinidad. Trans. Second Geol. Conf. Geol. Soc. Trinidad and Tobago, Port of Spain, pp. 88-105. Rosencrantz, E., 1990. Structure and tectonics of the Yucatan basin, Caribbean Sea, as determined from seismic reflection studies. Tectonics, 9:1037-1059. Roure, F., Carnevali, J.O., Gou, Y. and Subieta, T., 1995. Geometry and kinematics of the North Monagas thrust belt (Venezuela). Mar. Pet. Geol., 11: 347-362. Russo, R.M. and Speed, R.C., 1992. Oblique collision and tectonic wedging of the South American continent and Caribbean terranes. Geology, 20: 447-450. Russo, R.M., Speed, R.C., Okal, E.A., Shepherd, J. and Rowley, K.C., 1993. Seismicity and tectonics of the southeastern Caribbean. J. Geophys. Res., 98: 14299-14319. Salvador, A. and Stainforth, R.M., 1968. Clues in Venezuela to the geology of Trinidad, and vice versa. Trans. Fourth Caribbean Geol. Conf., Port of Spain, Trinidad, pp. 31-40. Saunders, J.B., 1968. Excursion Number 1: Manzanilla coast.
557
Trans. Fourth Caribbean Geol. Conf., Port of Spain, pp. 427429. Saunders, J.B., 1972. Recent paleontological results from the Northern Range of Trinidad. Trans. Sixth Caribbean Geol. Conf., Margarita Island, pp. 455-460. Schubert, C., 1984. Basin formation along the Bocono-Mor6nE1 Pilar system, Venezuela. J. Geophys. Res., 89:5711-5718. Snoke, A.W., Yule, J.D., Rowe, D.W., Wadge, G. and Sharp, W.D., 1990. Stratigraphic and structural relationships on Tobago, West Indies, and some tectonic implications. In: D.K. Larue and G. Draper (Editors), Transactions of the 12th Caribbean Geological Conference, St. Croix, U.S. Virgin Islands. Miami Geological Society, Coral Gables, Fla., pp. 389402. Speed, R.C., 1985. Cenozoic collision of the Lesser Antilles arc and continental South America and the origin of the E1 Pilar fault. Tectonics, 4:41-69. Trechmann, C.T., 1935. Fossils from the Northern Range of Trinidad. Geol. Mag., 72: 166-175. Tyson, L., 1989. Structural features associated with the Los Bajos fault and their interpretation in light of current theories of strike-slip tectonics. Trans. 12th Caribbean Geol. Conf., St. Croix, U.S. Virgin Islands, pp. 403-414. Tyson, L. and Ali, W., 1990. Cretaceous to Middle Miocene sediments in Trinidad. Trans. Second Geol. Conf. Geol. Soc. Trinidad and Tobago, Port of Spain, pp. 266-277. Tyson, L., Babb, S. and Dyer, B., 1991. Middle Miocene tectonics and its effects on Late Miocene sedimentation in Trinidad. Trans. Second Geol. Conf. Geol. Soc. Trinidad and Tobago, Port of Spain, pp. 26-40. Vierbuchen, R.C., 1984. The geology of the E1 Pilar fault zone and adjacent areas in northeastern Venezuela. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The CaribbeanSouth American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mere., 162: 189-212. Wadge, G. and Macdonald, R., 1985. Cretaceous tholeiites of the northern continental margin of South America: the Sans Souci Formation of Trinidad. J. Geol. Soc. London, 142: 297-308. Wadge, G. and Shepherd, J.B., 1984. Segmentation of the Lesser Antilles subduction zone. Earth and Planetary Science Letters, 71: 297-304. Wilson, C.C., 1940. The Los Bajos fault of south Trinidad, B.W.I. Am. Assoc. Pet. Geol. Bull., 24: 2102-2125.
Chapter 19
New Insights on the Formation of the Caribbean Basalt Province Revealed by Multichannel Seismic Images of Volcanic Structures in the Venezuelan Basin
J O H N D I E B O L D , N E A L D R I S C O L L and EW-9501 S C I E N C E T E A M 1
A regional multichannel seismic survey (EW-9501) was carried out in the Venezuelan Basin aboard R/V Ewing. The resulting reflection profiles image previously unseen structures within the entire thickness of the Caribbean oceanic plateau basalts east of Beata Ridge. Identifiable volcanic products include large amounts of extrusive material, which form two morphologically distinct, vertically stacked sequences. The lower sequence consists of local highs and ridges, flanked by wedges made up of dipping flows. The upper sequence is more homogeneous, featuring widespread flows which fill morphological and extensional lows in the lower sequence. Dipping sequences of volcanic flows are seen with length scales varying from 20 km to over 100 km. In contrast to seaward-dipping reflectors described elsewhere, these flows appear to be of submarine origin, and to have maintained their primary sense of dip. The morphology and scales of the flows are controlled by the volume of source material. The top of the upper 'high-volume' sequence forms the smooth B" horizon sampled by DSDP and ODP drilling. A more detailed survey was made of the southeastern edge of the plateau which includes a 100-kin-wide sequence of flows forming gently dipping wedges overlying thinned oceanic crust. The resulting images show clearly that the present edge is constructional, and that buried beneath it is an earlier, possibly rifted edge. This two-phase development is characteristic of the plateau everywhere in the Venezuelan Basin, and interpretations of the Caribbean's history must take this into account. It is not clear how much time elapsed between the end of the first, and the onset of the later phases, but our results indicate that the later phase is the only one so far sampled by Caribbean drilling.
INTRODUCTION Two-ship refraction profiles carried out in the 1950s yielded velocity functions of a Caribbean crust that was anomalously thick (up to 20 km), with an unusual two-layer structure. Based on early analog single-channel seismic reflection data, two distinctive horizons were designated A" and B", according to their resemblance to A and B in the Atlantic, and A', B' in the Pacific. Sediment cores, carefully located on the basis of those single-channel profiles, showed A" to correlate with an Eocene chert layer, a result later confirmed by D S D P drilling, which indicated that layers of limestone also contributed to the reflective character of A". The makeup of B" (stands for acoustic ' b a s e m e n t ' ) remained unknown for a while longer. Magnetic anomalies in 1Lewis Abrams, Peter Buhl, Thomas Donnelly, Edward Laine, Sylvie Leroy, Adrienne Toy
the Venezuelan Basin seem to form linear N E S W trends, which correspond with those of minor faults and grabens i m a g e d in reflection profiles and m a p p e d by Case and H o l c o m b e (1980). A n u m b e r of efforts (Christofferson, 1973; Ghosh et al., 1984) have been m a d e to fit the magnetic 'lineations' in the Venezuelan and C o l o m b i a n Basins to known seafloor spreading sequences, but the results have been unpersuasive. Drilling, during D S D P Leg 15, sampled B" for the first time, and found it to be the top of a Cretaceous volcanic sequence. The igneous rocks of the Caribbean basalt province, as sampled by drilling at five sites during D S D P Leg 15, Site 1001 during O D P Leg 165, and in on-land occurrences, are virtually all basalt. Their character has been s u m m a r i z e d by Donnelly et al. (1990; Fig. 1). The bulk of the basaltic material occurs in three petrochemical varieties: transitional tholeiitic basalt, highly magnesian basalt, and mildly alkalic basalts. The transitional
Caribbean Basins. Sedimentary Basins of the World, 4 edited by P. Mann (Series Editor: K.J. Hsti), pp. 561-589. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
562
J. DIEBOLD et al.
Fig. 1. EW-9501 track, with line numbers and sonobuoylocations on bathymetry; 500 m contours. DSDP and ODP sites where volcanic basement was sampled by drilling are indicated. tholeiitic basalts are the most widely distributed and the most voluminous of the basalt types sampled in the Caribbean region. The highly magnesian basalts have been termed 'komatiite' on Isla Gorgona by Echeverria (1980) and in Colombia by Spadea et al. (1989), 'picrite' on Curaqao by Beets et al. (1982) and Klaver (1987), and 'low-Ti-basalts' in the Santa Elena Peninsula of Costa Rica by Wildberg (1984). The mildly alkalic basalts are best known from the Beata Ridge (DSDP Site 151) and from the Dumisseau Formation of southern Haiti. They are also represented in the younger (latest Cretaceous or Paleogene) basaltic centers from Costa Rica (Wildberg, 1983; Frisch et al., 1992). These basalts are enriched in K, Th, and several minor elements. They resemble well-known, mildly alkalic basalts of Hawaii, and are characteristic of ocean island basalts. Where there is some local basis for stratigraphic positioning, they seem to occur among the later eruptive products. One of the most unsatisfactory aspects of the previous studies of the Cretaceous flood basalts has been establishing ages of magmatic activity. Conventional K-Ar age determinations that are very numerous in the literature (Donnelly et al., 1990) can be shown in several instances to contradict clear stratigraphic constraints. In general, the ages ascertained from K-Ar are too young because of Ar loss or K uptake during seafloor weathering or during burial metamorphism. Problems with biostratigraphic ages
also occur because it is difficult to assess the duration of the hiatus between basalt emplacement and pelagic accumulation. Even worse, if the basalts are sills then the biostratigraphic age only yields a maximum age of emplacement. Despite the numerous dating problems, there are several firm ages for the termination of magmatic activity, indicated by the locally youngest ages of the Cretaceous basalts. As reviewed in Donnelly et al. (1990) these dates are mainly derived from dating sedimentary deposition on the basalt complex, including DSDP Leg 15 sites. These ages cluster tightly at the latest Turonian, but in a few cases could be as young as Santonian, and even early Campanian (DSDP Site 152 and ODP Leg 165). Recently, a detailed radiometric study of basalts from Gorgona, Costa Rica, Haiti, and Curaqao has been undertaken by Sinton et al. (1998) using 4~ incremental heating methods, which yield dates consistent with the sedimentary ages for the termination of the event. The number and geographic extent of consistent age determinations at 88-90 Ma suggest that a truly vast igneous event occurred at this time. Nevertheless, there remain several circumstantial fossil occurrences of Albian and Aptian age associated with the Cretaceous basalt province. For example, intercalated within the Curaqao basalt complex is a sedimentary succession with six genera of ammonites that indicate a latemiddle Albian age (Weidmann, 1978), which is significantly older than the biostratigraphic and ra-
MULTICHANNEL SEISMIC IMAGES OF VOLCANIC STRUCTURES IN THE VENEZUELAN BASIN
diometric ages obtained elsewhere. There are two interpretations: (1) the ammonite ages are incorrect, or (2) there was an earlier age of magmatic activity of unknown extent, and the 88-90 Ma event was a terminal event for the main Caribbean Cretaceous basalt province, but not the initial event. Early multichannel seismic (MCS) data acquired by IFP (France), UTIG (Texas) and L-DEO (Lamont) occasionally imaged internal layering beneath B", and showed that to the southeast, the volcanic sequence had a distinct edge, beyond which lay a triangular area of deep, rough acoustic basement (called 'rough B"') whose reflective character resembled oceanic crust. This area was found, however, to be practically devoid of magnetic signature (Donnelly, 1973), suggesting that the crust was either created by seafloor spreading during a magnetically quiet period, or that the original magnetic signature was degraded by thermal or tectonic modification. Early MCS surveys, carried out during the 1970s, defined the approximate area of rough basement. The distinction between the rough and smooth varieties of B" was first described by Talwani et al. (1977) and a series of normal faults was shown to coincide with the rough-smooth boundary by BijuDuval et al. (1978). The acquisition of additional crossings of the boundary showed that it had a variety of manifestations, ranging from imperceptibly gradual to abrupt (Diebold et al., 1981). Lamont data (described by Diebold et al., 1981) provided five crossings of the rough-smooth B" boundary, which were not enough to characterize the exact nature of the transition. Additional crossings by French (IFP) and US (UTIG) investigators all converged at the boundary in the same area as Lamont line 108. The resulting coverage showed that the basement faults controlling the rough-smooth boundary had strikes oblique to the general trend of the transition, suggesting that the basement fabric pre-dated the volcanic emplacement. MCS line 108 was unique in that it showed clearly how depth to Moho increases, in a step-wise fashion, northward across the roughsmooth transition, leading Diebold et al. (1981) to interpret smooth B" as the top of Cretaceous flows which had overrun and depressed rough B" crust. Sonobuoy refraction lines shot in 1974 in conjunction with MCS data (Talwani et al., 1977) followed by two-ship multichannel seismic expanding spread profiles (ESPs) carried out in 1976 showed that the rough basement horizon B" was the top of what appeared to be anomalously thin oceanic crust, and that the crust underlying the only ESP located over smooth B" was, as expected, unusually thick, and comprising (in a gross sense) two layers (Diebold et al., 1981). Though it was the first to be discovered, the area of smooth B" has been the most difficult to pene-
563
trate by seismic reflection. This is principally due to the high reflection coefficients of the sedimentbasement interface there, though in early work, large water depths and weak profiling systems played a part. Reflections beneath the smooth B" horizon were detected from time to time in early MCS data (Hopkins, 1973; Ladd and Watkins, 1980; Stoffa et al., 1981; Diebold et al., 1981). In some cases, these reflectors were observed to dip in relation to the B" surface, though their identification was often open to doubt. The strong basement reflection coefficient gives rise to strong intrabed ('peg-leg') multiples that reverberate efficiently within the sedimentary section. The fact that water depths were quite large, compared to the length of receiving arrays used in the early work, made it difficult to discriminate between those multiples and primary arrivals. Intrabed multiples are still a problem, even with the 4 km receiving array of R/V Ewing.
EW-9501
In February and March, 1995, 5200 km of MCS data, shot with a 20-airgun (8415 cubic inch) source array, and a 4-km, 160-channel hydrophone streamer, were acquired aboard R/V Ewing. Underway geophysics included measurements of gravity, magnetics, and swath bathymetry. 104 successful deployments of expendable sonobuoys were made. A track map, identifying the seismic reflection lines by number, with sonobuoy deployment locations, is shown as Fig. 1. The purpose of this NSF-funded cruise was to better define the boundary between smooth and rough B", to image structures within and beneath the Caribbean volcanic complex, and to provide new data on the structure and development of the Beata Ridge. The new data meet these goals, and also reveal several unexpected features of the sedimentary column which provide important new constraints on the tectonic and oceanographic development of the Caribbean region. The motivation behind the cruise was to obtain deep reflection seismic data in the Venezuelan Basin and on the Beata Ridge crest and its eastern flank. In the Venezuelan Basin, the survey track was laid out to complement MCS data previously acquired by R/V CONRAD in the 1970s (Talwani et al., 1977; Stoffa et al., 1981; Diebold et al., 1981) specifically where these lines crossed the rough-smooth B" boundary. The general NNW-SSE orientation of the lines, and their northwestward extent, reflected the desire to image deep structure (if any) related to previously mapped basement topography (Case and Holcombe, 1980) and magnetic lineations (Donnelly, 1973) which trend ESE-WSW. In retrospect, the Venezuelan Basin lines should have extended farther
564 to the north-northwest, and more cross-lines should have been shot, but to do this extra work would have required an impossibly long cruise. 104 Navy-supplied expendable sonobuoys were deployed during the survey. Data from most of these were recorded to maximum offsets of between 30 and 40 km, providing refraction and wide-angle reflection data for crustal velocity analysis. Onedimensional velocity analysis, using interactive ray trace modeling has been attempted for all of the sonobuoys. Velocity functions were obtained for 85 buoys. Results from this analysis, combined with earlier wide-angle reflection and refraction work, provide several important constraints on the structure and mode of emplacement of the Caribbean LIP (Large Igneous Province). The 4-km-long hydrophone array used in acquiring the reflection profiles, along with the powerful seismic source array of R/V Ewing, has produced the best MCS data ever shot in the deep-water basins of the Caribbean. The streamer length, in particular, allows high resolution of hyperbolic stacking velocity, which, if accurately determined, results in amazingly detailed images of what now can be seen as intricate and convoluted sub-basement structures in the Venezuelan Basin. The other side of this coin is that stacking velocity analysis is a painstaking and time-consuming process. The high velocity contrasts at the sediment-basement interface, and the large range of source-receiver offsets present in the data (and required, to ameliorate the effects of intrabed multiples), cause unusually extreme lateral changes in the stacking velocities, despite the large water depths of the Caribbean, which normally could be expected to minimize these changes. As a result, most of the reflection data we present here are in the form of 'brute' stacks. In processing, normal moveout corrections were made based on a velocity model created by extrapolation from sonobuoy results, and the stacking velocity analyses manually carried out for line 1293, the only line whose stacking velocities have so far been completely determined. Therefore, the foldout section for line 1293 (Fig. 2) is very important in evaluating the discussions presented below. While the other, brute-stack sections show features never before seen, the 1293 profile gives a feeling for the level of fine detail that will eventually be imaged by these data. Excepting velocity analysis, all of the profiles were processed identically. The field data were filtered and decimated to 4-ms sample size. Water depth-dependent outer and inner mutes were applied (the effect of outer muting is to make imaging of sedimentary structures less sensitive to errors in stacking velocities; the inner mute reduces the effects of peg-leg and water column multiples). After stacking, the data were filtered again and AGC was
J. DIEBOLD et al. applied. The reflection data plots presented here have been horizontally compressed so that they can be shown in their entirety. This has resulted in a vertical exaggeration of about 20.8 in the water column. A special dip-adaptive trace mixing process was employed, to enhance the appearance of coherent reflectors, so that when only every fifth trace is plotted, the result is to retain most of the structural information in the data. All of the reflection profiles are plotted with north and west to the left, south and east to the right. CDP numbers are in chronological sequence in each line. CDP spacing is 12.5 m, so 800 CDPs equals 10 km horizontal distance.
CRUSTAL ELEMENTS OF THE COLOMBIAN AND VENEZUELAN BASINS Thin crust
Despite 20 years accumulation of evidence to the contrary, the concept that the Colombian and Venezuelan Basins are uniformly capped by a Cretaceous igneous body persists. In fact, the crustal structure of these basins is much more complex, and is made up of at least four categories of lithological elements: (1) thin ocean-like crust; (2) volcanic extrusive sequences with dipping reflectors; (3) massive bodies of volcanic extrusives and intrusives; and (4) extrusive volcanic mounds. While it might be argued that the last three categories listed above are simply different manifestations of the same crust-thickening Cretaceous volcanic event, the Caribbean Sea contains a significant amount of crust whose thickness varies from normal (6-8 km) to abnormally thin (3-5 km). Although analysis of many of the early two-ship refraction profiles indicated the presence of anomalously thick crust, several, particularly west of Beata Ridge, detected crust of normal oceanic thickness (Ewing et al., 1960; Edgar et al., 1971). Ludwig et al. (1975) described the basement reflector in the Colombian Basin as 'rough' in comparison to that in the Venezuelan Basin. In 1977, Houtz and Ludwig mapped a number of areas in the Colombian Basin where crust was deep, rough and thin. Bowland and Rosencrantz (1988) re-mapped a roughsmooth boundary in one of the areas outlined by Houtz and Ludwig (1977) and mapped several others. To the east, Talwani et al. (1977) provided MCS profiles showing rough basement in the Venezuelan Basin. Simultaneously recorded sonobuoys showed that crust in those areas was thin, compared to normal oceanic crust. Biju-Duval et al. (1978) also observed rough, deep basement in the southeastern comer of the Venezuelan Basin and suggested that it corresponded to oceanic crust, upon which was su-
MULTICHANNEL SEISMIC IMAGES OF VOLCANIC STRUCTURES IN THE VENEZUELAN BASIN perimposed the volcanics to the northwest. Diebold et al. (1981) interpreted additional two-ship velocity profiles and obtained velocity-depth functions showing that the rough B" crust resembled unusually thin oceanic crust. They also extended mapping of the rough-smooth B" transition around the Venezuelan Basin and imaged Moho in a few places beneath both rough and smooth crustal types. A compendium map of crustal thickness, throughout the Caribbean, was presented by Leroy (1995). The Moho reflection can be seen beneath rough B" on all of the EW-9501 profiles, and on a few of the older RC- 1904 and RC-2103 lines (Diebold et al., 1981). Interval two-way time of the rough crust is typically 1.2 s, with a corresponding sonobuoyderived thickness of 3800 m. In a few places, crustal thickness varies, thinning to as little as 0.9 s, and thickening to as much as 1.6 s. Velocity seems to vary also, though no systematic changes have yet been detected. Our results agree with those of Talwani et al. (1977) and Diebold et al. (1981). Upper crustal velocities are commonly around 5.5 kin/s, though occasionally, a thin upper layer is detected with velocities from 4.2 to 4.9 km/s. Mid-crustal velocities of 6.6 km/s are typical, and often velocities around 7.1-7.2 km are seen below these. 'Normal' Moho velocities are never seen, but instead, critically refracted energy with velocities between 7.4 and 7.8 km/s appears, corresponding to rays turning at depths below reflection Moho. These crustal thicknesses are well below the normal range, as defined by White et al. (1992), who also demonstrated a historical bias towards the underestimation of oceanic crustal thickness. White et al. (1992) compared inversion results based on forward synthetic seismogram modeling to those obtained by traditional slope-intercept methods and found that they differed (typically) by 20%. Even if such a bias is present in our methods (neither slope-intercept, nor synthetics-based modeling) the thicknesses we (and Talwani et al., 1977) determine for rough B" crust are abnormally thin. EW-9501 MCS profiles in the southeastern Venezuelan Basin indicate that the seismically fast (3.5-4 km/s) 'transparent' sediments identified by Biju-Duval et al. (1978) are, in our view, laminated turbidite sequences directly overlying rough basement there. The stratal relationships that make this obvious (filling bathymetric lows and onlapping highs) are particularly well represented in line 1298 (Fig. 3). The interval two-way time between the easily identifiable A" and underlying 'basement' increases dramatically, from 0.2 s at CDP 14,500 to about 0.5 s at CDP 12,500. The number of reflectors within the A"-B" interval increases dramatically across the rough-smooth B" transition, and the reflections become more distinct. Similar sediments
565
have been imaged (though not always recognized for what they are) in the Colombian Basin (Lu and McMillen, 1982; Kolla et al., 1984; Bowland and Rosencrantz, 1988; Bowland, 1993). The presence of these sediments appears to be diagnostic of deep, rough, thin crust, and identifying them in older data can further extend the mapped areas of thin crust. Several EW-9501 MCS profiles show that in a few places the anomalously thin rough B" crust of the SE Venezuelan Basin produces low-angle, upward-concave reflectors which appear to be the surfaces of listric faults extending from the seafloor to the Moho; see for example line 1293 (Fig. 2) CDPs 24,000-25,000. As a rule, these do not correlate with crustal offsets, even where the reflections are seen to intersect the top of crust. Similar reflections have been imaged in several areas, including marginal NW Atlantic crust of Jurassic age (McCarthy et al., 1988; Mutter and Karson, 1992; Rosendahl et al., 1992; Zehnder Mutter, 1992) and are, apparently, the products of some aspect of seafloor spreading. A recent study by Kent et al. (1997) has demonstrated, however, that the same pattern of reflections may arise due to sideswipe from a sub-parallel basement ridge, and it is certainly possible that some of the 'faults' we see are artifacts of this kind. Although velocity analysis (not yet complete) indicates that these are intra-crustal events, Kent et al. (1997) have shown that streamer feathering can be the cause of high stacking velocities for sideswipe events, which would otherwise be better imaged with lower velocities. The question remains: was Venezuelan Basin rough crust formed thin, or was it thinned by some tectonic process, presumably, extension. Unusually thin oceanic crust has been associated with the development of passive margins, e.g. the Iberian margin (Whitmarsh et al., 1990) and in the Labrador Sea (Hinz et al., 1979; Srivastava and Keen, 1995) and occasionally in ocean basins (Jackson et al., 1982; Muller et al., 1997) presumably created at slowspreading ridges (Reid and Jackson, 1981; Chen, 1992; White et al., 1992). Mantle material beneath this thin crust often exhibits abnormally low velocity, presumably from serpentinization, resulting from invasion of seawater along fault planes. Similar velocities are observed in the southeastern Venezuelan Basin (Talwani et al., 1977). We note that the 'originally thin' oceanic crust of the Iberian margin forms a narrow strip (Whitmarsh et al., 1990), while the rough Venezuelan Basin crust is much wider in all directions. The southeastern extent of thin, rough B" crust is entirely unknown; it can be seen (for example in EW-9501 line 1319, Fig. 4) being subducted beneath the accretionary wedge of the Curacao Ridge. The destruction of the crust by normal faulting, apparent in this profile, is typical of
566 that seen in other MCS profiles crossing this zone. The amount of thin crust consumed by subduction and obduction is difficult to estimate, and its original southern boundary is difficult to reconstruct. The Iberian margin also features a narrow zone (25-30 km wide) of greatly thinned continental crust whose abnormally high basal velocities are thought to result from underplating (Whitmarsh et al., 1990). Both of these Iberian scenarios (thinned continental crust, thin transitional oceanic crust) produce crust having a thickness similar to that of the southeast Venezuelan Basin, and also having little or no magnetic signature. The variation of apparent velocities in rough Venezuelan Basin crust is such that seismic velocity alone cannot be used to make the distinction between crustal types described by Whitmarsh et al. (1990). As in the case of its thin, early-stage oceanic crust, the radically thinned continental crust of the Iberian margin is found only in a narrow band. Rosendahl et al. (1992) have determined a wider (not specified, but shown as at least 65 km) zone of attenuated continental crust off the African margin in the Gulf of Guinea. This crust has also, apparently, been thinned by faulting, which causes a rough 'pseudo-oceanic' appearancepseudo-oceanic appearance in reflection profiles. This crust is of comparable thinness to that of Caribbean rough B", but this thinness is defined by an elevated 'equilibrium' Moho whose smooth reflection character is quite dissimilar to that of Moho imaged anywhere in the Caribbean. Next, we present evidence that crustal extension was taking place during the emplacement of Venezuelan Basin smooth B" volcanics. It is likely that such extension may have reactivated preexisting faults in rough B" crust.
Rough-smooth B" transition It is apparent from the EW-9501 data that at least some of the area of thin rough crust pre-dates the emplacement of smooth B" volcanic material, and that the area of 'originally' rough crust is, therefore, significantly larger than previously realized. Both the reflection profiles and the sonobuoy data show that the thin crust underlies the Cretaceous volcanics, and is present at least 50 km northwest of the roughsmooth boundary along some profiles. Ewing profiles 1300-1316 provide 3-D coverage dense enough to show that the model proposed by Diebold et al. (1981) for the rough-smooth boundary as an overrunning of en-echelon normal fault blocks by lava flows encroaching from the northwest is basically correct. The uppermost, and presumably youngest, sequence of extrusives included highly mobile flows, which were able to spread in thin fingers, along the valleys separating gently tilted fault blocks in the
J. DIEBOLD et al. rough B" crust. The strike of those faults which can be accurately mapped is ENE-WSW, and none seems to extend for more than 100 km m most are shorter. Their en-echelon pattern may be indicative of wrench faulting. Some of the extrusive fingers can be seen as bright, smooth spots as far as 20-30 km from the principal rough-smooth boundary (for example, line 1298, CDP 11,000; line 1293, CDP 20,350). When successive flows filled and finally buried the valleys, overlying flows were able to spill over into the next valley, thus spreading quickly to the east-northeast, more slowly south-southeast. In several places, the tilting of the normally faulted blocks continued after a thin veneer of volcanic material had been emplaced, creating the smoothtopped tilted blocks ('ski-jumps') characteristic of the Central Venezuelan Fault Zone (so named by Biju-Duval et al., 1978), and seen here in Ewing lines 1293 (Fig. 2, CDP 19,100), 1298 (Fig. 3, CDP 14,000), and 1300 (Fig. 5, CDP 15,000). Lines 1300 and 1302, though only 35 km apart, show different manifestations of the rough-smooth B" transition. On line 1300 the distal edge of the smooth B" volcanics is tilted back into a 'skijump' form, probably caused by post-emplacement extensional rotation of underlying blocks of rough basement. The fault system that controls the geometry and structure of the 'ski-jump' must be relatively surficial and antithetical to a larger fault system that dips north-northwest and controls the location and geometry of the large divergent wedge observed beneath B" (CDPs 9000-13,000; in line 1293, CDPs 10,500-19,000) and the large step in Moho topography (CDP 15,000; in line 1293, CDP 19,000). In line 1302, the transition (Fig. 6, CDP 5200) is quite gradual, the flows having apparently run around the end of line 1300's rotating block. In both cases, thin, upper flow units appear to have traveled in and out of the profile, leaving thin edges facing both north (CDP 10,900) and south (CDP 6900). The rough-smooth transition as seen in line 1302 is gradual, as the expression of the normal faults is only slight there. As a result, the rough B" surface can be seen beneath the distal edge of the smooth B" flows. A pair of the previously described throughgoing listric crustal 'faults' characteristic of rough B" crust can also be seen (CDPs 6500 and 7500). Once it is covered by smooth B" volcanics, the reflection from the top of the thin crust is quite weak, which is probably the result of three things, i.e. a small velocity contrast, a rough interface, and the absence of an intervening sedimentary layer. Some sonobuoys from this and previous studies have produced complete crustal sections, particularly in an area where sonobuoy records (e.g. Fig. 7) show that strong postcritically refracted energy is frequently returned from the Moho. This area is crossed by
~o
0 0 cO
0 0
0 0
0 0 ¢0
0 0
0 0
0 0
0 0 0
0 0
13.. T O
o
CO
oes 'ewp,/~eN~-OM± o~
0
iI ,
.
•
. . . . . .
,.
0 0
0 0 co
0
.
..
2 - " - _- -.
. ~ . N
...
....
-..~.
.-..r
.-
.
•
.. .
~.
7._-. ~. >. . . . _-. _
.
.... ...,- ::. ....
.~ .: . . . . -_
-<.:C.
. . .
.!'. f.-.:
.
.. -~-..:~ :
.'..~--:_ ".
.::.. : s : ..
,:,,
O4
~:~L.:-C~::ii~:~ ::i..<~.
.
.
:
.
.
_-
.
,
..
1 . . ~ - .
.
•
04
..
, _:-: .:.!:i.) ::: J:.::":: .
-~.
.
.-; .:,-.!.:~..... ....... .
1--
i
0 0 0
0 0 0
0 0 0 CO
0 0 0
0 0 0 ¢0
0 0 0
0 0 0 GO
o o ' r ' -
8 0
0 0
o
o
0 0 0
0 0 0 co
0 0 0 o~
0 0 0 0 0~1
0 0 0 ~!" 04
0 0 0 CO 04
0 0 0
-. . . .
0 0 0
04
0 0 0
0~1
':.<~:Y~<:~i<:>:i:
0 0 0
:
.
0 0 0
--..-.,
0 0 0
0
0 0 0 ¢0
0 0
0 0
~0oO
0
~o
.~ _1 0 1.0
~o LU
®o
'eLl.ill/~eM-OM.L
o4
00 ° 03
CO 08£
E uD
O
t
c~ ~D
o o o
c~
.,--~
> 0 ¢j i
*r
,..0
o q
.o
© ©
? ©
©
©
C'-I
© ,..o rm
©
t¢'3
©
. ,...,
MULTICHANNEL SEISMIC IMAGES OF VOLCANIC STRUCTURES IN THE V E N E Z U E L A N BASIN
571 {D
9
9
. ,...~
. ,...~
N
y. 9
O
.,..~ +..a
O
r,.)
e~t) ....~
. ,...~ o o tt-)
.,..~
o
~" ~
t--,~
572
J. D I E B O L D et al.
o<
O
.,..~
~D
~[--, ~c5 ~D
~D
O
~D
~D
.,..~ . ,...~
~D
. ,...~
oo==~ bq
9
9
,.~
O
r~
,.~
0
d~
0
MULTICHANNEL SEISMIC IMAGES OF VOLCANIC STRUCTURES IN THE VENEZUELAN BASIN lines 1300 (CDPs 5000-14,000) and 1302 (CDPs 6000-12,000). The velocity functions determined for these buoys consistently include a pair of deep layers whose velocities (high 6 over low 7) are consistent with a 'buried' thin crustal layer. A final argument for identifying rough B" beneath smooth B" crust is the similarity of the Moho reflection character, as seen on EW-9501 profiles throughout the Venezuelan Basin. North-northeast of the rough-smooth B" boundary, the reflectors immediately beneath the smooth B" basement exhibit a divergent pattern (Figs. 2, 3, 5 and 6) and have interval velocities on the order of 5 km/s. The divergent geometries and spatial distribution of the reflector packages observed across the Venezuelan Basin suggest multiple volcanic extrusive events, with the earlier events being less voluminous and more spatially restricted. The multiple magmatic events inferred from the stratal geometry suggests that some of the older dates for magmatic activity (e.g. Aptian) may actually be recording real events and are not artifacts of the dating technique (Donnelly et al., 1990). The rotation and divergence of the onlapping seismic reflectors diminishes up section and is indicative of differential subsidence and block rotation (Fig. 2). Note that the spatial extent of each successive divergent reflector package increases within the wedge. Symmetry of the dipping packages is also often observed locally towards seamounts and away from highs in the underlying volcanic sequence.
Smooth B"; extrusive volcanics EW-9501 seismic profiles show extensive layering within the volcanic extrusive sequences making up smooth B" crust. Patterns of layering are consistent from line to line in the Venezuelan Basin, and along with sonobuoy results, define several distinguishable sequences. Distinctions can be most easily made in the southern half of the basin. An upper sequence thickens dramatically in the direction of the rough-smooth B" boundary. Sub-units of this upper sequence also thicken towards the southeast, and seem to be individually thick and voluminous. Distinct and characteristic refracted arrivals in sonobuoy records show that velocity within individual layers tends to increase gradationally. Sonobuoy 33 (Fig. 7) deployed at the location of CDP 10,100 of line 1300 (Fig. 5) is an excellent example. Refracted arrivals from the volcanic sequence are seen to emerge from the seafloor reflection at 8.2-km source-receiver offset and 7 s, reduced traveltime. This record also features a strong Moho critical reflection. The base of the uppermost layer is clearly imaged as a reflection in MCS line 1293. The layer increases in thickness from 0.3 s (780 m) at CDP
573
13,500 to 0.56 s (1455 m) at CDP 1800, and has a velocity (varying from 5.0 to 5.2 km/s) which is distinctly lower than that of most underlying volcanic layers, in which velocities typically increase to 5.4-5.6 km/s. This velocity contrast is responsible for the clear break in slope of the first arrivals at 12 km of sonobuoy 33. The reduced velocities of this upper layer may be the result of some compositional or physical variation (vesicles or brecciation, for example), weathering, or interbedded (but non-reflective) sediments. Planke (1994) found systematic velocity variations within individual flows in seaward-dipping wedges, sampled at ODP Site 642 on the V~ring volcanic margin. Velocities decreased steadily from the interior to about 3 km/s at the upper and lower boundaries of all logged flows. If flows were thick enough (more than 9 m) velocities in the body of the flows would stabilize at about 5.2 km/s, with an upper velocity gradient zone 7 m thick, and a bottom zone 2 m thick. The marginal velocity decreases are attributed by Planke to brecciation and vesicularity. Similar patterns at a larger scale are seen in Colombia River basalt flows (Waters et al., 1981). Planke's observation predicts that the average velocity of a sequence of flows will depend on the average thickness of individual component flows; a sequence with a larger number of thinner flows will have an overall velocity lower than another sequence made up of a smaller number of thicker flows. It well may be that the uppermost Caribbean flow sequence is made of thinner flows than the lower sequence. On the other hand, such systematic velocity layering might also produce an anisotropic substance, whose true average velocity would not be determined from sonobuoy or MCS data. In some places, a thin (ca. 40 m) layer can be seen in the reflection profiles, forming the uppermost element of smooth B". EW-9501 acquisition parameters (source and receiver towed at 7.5 m) dictate that this is the thinnest basaltic layer that can be detected. Drilling of seaward-dipping reflector sequences has rarely sampled a flow thicker than 10 m, but we note that the morphology of the smooth B" sequences is more reminiscent of continental flood basalts, in which individual flows may have thicknesses of 100 m or more (Waters et al., 1981). The observation of simple, edge-like terminations of these thin upper layers suggests that they might in fact be individual flows. An example, referred to above, is seen at CDP 6900 of line 1302 (Fig. 6). Similar flow sequences form the thin, mobile 'fingers' filling basement lows southeast of the principal rough-smooth B" boundary. Their identification as flows, and not turbidite lenses is based on the strong reflection amplitude of their upper surfaces, their distribution over thin B" crust, and on detailed analysis of stacking velocities in a few areas. Many
574
J. D I E B O L D et al.
tt~
~D
~D
9
O
~D
o,..,
I ~D ~D
~D 9
E o,.., O o,.., o~
O 9
O ~D
tt')
d0
o,..,
cr ~D ~ 9 ~D
r162 9
9 eq r ~D o,,.~
~5
MULTICHANNEL SEISMIC IMAGES OF VOLCANIC STRUCTURES IN THE VENEZUELAN BASIN
575
Fig. 7. Sonobuoy33, located along line 1300 (Fig. 5) returned an unusually fine record. Precritically reflected, postcritically reflected and refracted arrivals were all used to find a detailed 1-D solution. sonobuoy records contain weak, wispy arrivals, apparently refracted in thin flows (or sills) separated from the underlying body of volcanics by sedimentary layers. These observations suggest that the reduction in overall upper-sequence velocity arises from the inclusion of intercalated sediments. Such sediments (marbleized limestones) were sampled at DSDP Site 152 (Edgar et al., 1971). A sonobuoy deployed as closely as possible to the location of DSDP Site 150 features an early refracted arrival characteristic of a thin (ca. 100 m) sill (or sequence of sills) whose top is about 300 m above the massive volcanic body. Similar events have been observed in sonobuoy records elsewhere in the Venezuelan Basin and in the western Pacific (Abrams et al., 1993). In many other sonobuoy records from the zone of smooth B", the refracted arrivals similarly show that what appears in the reflection profiles to be the upper surface of the LIP is in fact a sill (or flows), or a sequence of sills (or flows) and sedimentary layers, whose overall thickness varies from zero to 300 m. Similar sills were drilled during DSDP Leg 15. The bottom of this uppermost layer (or sequence of layers) is marked in much of line 1293, by a visible reflector, indicating the transition to material with higher seismic velocity. The fact that this layer, with its characteristic slightly lower velocity, is seen in many sonobuoy records, suggests that it exists in most of the area, and that its base will be seen as a
reflector when more of the lines are re-stacked after detailed velocity analysis. The upper layer described in the preceding paragraphs is the uppermost member of a sequence of similar layers. Like the uppermost layer, these underlying layers thicken towards the rough-smooth B" boundary. Between CDPs 11,000 and 16,000 of line 1293 (Fig. 2), the base of the resulting wedge is marked by a bright reflector, whose continuity is broken only by scattering from a structural high in the upper smooth B" surface at CDP 14,400. Interval two-way time for this entire wedge increases from about 150 ms at CDP 11,000 to 1700 ms at CDP 16,000. Sonobuoy-derived velocities indicate that this represents a thickening from 400 m to 5 km, which takes place across a distance of 62.5 km. Similar bright reflectors can be seen in the less-painstakingly processed sections, and are particularly apparent around CDP 17,500 of line 1298 and CDP 11,500 of line 1300 (Fig. 5). The strength of these reflectors suggests three characteristics of the corresponding horizon: smoothness, a single interface with extremely high impedance contrast, and/or superposition of a sequence of thin layers with moderately high impedance contrast. The sonobuoy results indicate a slight overall increase in seismic velocity, and therefore, impedance, but the resulting contrast alone is not sufficient to explain the reflection strengths. The smooth appearance of the
576 reflections may result from the inclusion of an appreciable amount of metamorphosed sedimentary layers between volcanic extrusives, which would produce reflection-enhancing fine-scale layering. Noting also the patterns of less-strong but still clear reflectors between this basal horizon and B" it is worthwhile to re-examine the sequences mapped in southern Haiti by Maurrasse et al. (1979) who interpreted them as the products of intermittent but "extensive magmatic outpouring - - at least throughout the late Cretaceous, and possibly throughout the early Cretaceous as well." This ca. 2000-m-thick series of diabase and basalt flows and sills, some featuring columnar jointing, and others pillows, typically separated by layers of pelagic sediments, would provide a reasonable model for parts of the superior wedge of Venezuelan Basin volcanics. The most remarkable feature of the upper sequence is its massive, wedge-like southeastward thickening toward the rough-smooth B" boundary. In the eastern part of the basin, as exemplified in line 1293 (Fig. 2) the wedge is about 100 km long, but north of the narrow point of the wedge (CDP 11,000) the upper sequence is relatively thin, though its thickness seems to increase in the vicinity of the seamounts at CDPs 2100 and 6000. Examining the profiles, one by one, progressively to the west, the flat part of the upper sequence is (usually) clearly defined, and can be seen to thicken steadily. Fortunate velocity contrasts perhaps coupled with the presence of intervening layers of sediments, make the basal reflector of the upper sequence quite apparent between CDPs 300 and 7000 on line 1300 (Fig. 5). Interval two-way time of the sequence is
J. DIEBOLD et al. more or less constant at 0.8 s, or ca. 2 km thickness. Moving farther west, to lines 1304-1317, the upper-sequence thickness increases, and begins to vary, as two sub-basement highs begin to appear (Figs. 8 and 9). The upper-sequence thickness at CDP 5000 is about 4 km, nearly equal to the maximum 5 km of the wedge of line 1293. In general, the upper surface of the upper sequence is far smoother than its bottom. The lower surface also tends to dip away from local basement highs. Such dips are strikingly apparent between CDPs 11,000 and 16,000, line 1293 (Fig. 2), but they are seen at various places on all of the other lines as well. In one area, sonobuoy records consistently show that this upper sequence includes at its base a low-velocity zone with locally substantial thickness. The most persuasive example of the seismic arrivals characteristic of this zone are seen in sonobuoy 65. Located on N-S line 1320, near the W S W - E N E line of volcanic edifices whose magnetic anomalies were previously mapped, this record (Fig. 8) shows clear evidence for a significant low-velocity zone within the volcanic sequence. Based on analysis of the adjacent sonobuoys, this zone thins rapidly to the north and south and is, we judge, likely to comprise volcaniclastics from eruptive events at the nearby volcanic mounds, similar to those which have also been imaged along the northern and northwestern edges of the Venezuelan Basin. Directly beneath the bright reflector marking the base of this extensive, smooth-surfaced upper, wedge-forming sequence is another sequence of dipping reflectors whose character is entirely different. As it appears between CDPs 11,500 and 13,000 of
Fig. 8. Sonobuoy65, though not yielding as high a quality signal, nor lasting as long as sonobuoy 33, returned a remarkable diagnostic pattern of arrivals defining a prominent low-velocityzone beneath the upper volcanic layer.
MULTICHANNEL SEISMIC IMAGES OF VOLCANIC STRUCTURES IN THE VENEZUELAN BASIN line 1293, this sequence resembles the 'seawarddipping reflectors' (SDRs) observed at many passive continental margins (Hinz, 1981; Mutter et al., 1982) but also those seen in the Kerguelen Plateau (Ramsay et al., 1986). In comparison to the major reflectors of the upper wedge, which are typically well separated, and some of which can be followed continuously for 75-100 kin, these 'SDR-like' reflections are downward-concave and are densely spaced. Although the elements of the upper wedge can be clearly seen to terminate at and overflow the tops of rotated normal fault blocks, the basal extent of these 'SDR' reflectors cannot (so far) be reliably determined. In nearly every case, dipping reflectors of both kinds dip away from a basement high. This kind of high is best developed at CDP 10,000, line 1293. Here, as on some other lines, wedges of reflectors can be seen dipping away from the highs in both directions, though in every case, the larger and most clearly developed set dips toward the southeast and the rough-smooth B" boundary. Massive volcanic bodies
Between the zones of dipping reflectors, we find basement highs, in the form of gentle domes or ridges. These highs are generally about 30-40 km across and feature few, if any, internal reflectors. In the Venezuelan Basin, the principal highs form two linear trends, essentially coincident with linear magnetic anomalies mapped in the region by Donnelly (1973). The two lineations are somewhat oblique, and trends mapped by Case and Holcombe (1980) indicate that they may intersect near CDP 10,000 of line 1293 (Fig. 2). This might be a factor in producing the prominence of the high in this particular crossing. The dipping reflector sequences appear to originate in these highs which suggests that they are the locus of volcanic sources. Small knolls are evident at a few of the highs, made particularly apparent in the horizontally compressed sections shown here. Examples can be seen at CDP 1200, line 1298 (Fig. 3) and CDP 13,000, line 1302 (Fig. 6). These may be the peaks of (or the flanks of nearby) extrusive mounds. The northernmost of the two ridge-like lineations is crossed completely by only two reflection profiles: 1293 (Fig. 2) and 1320 (Fig. 10). ODP Site 150 was located near CDP 5500 of line 1320, on the flank of this high, whose relief, in conjunction with ocean currents, has produced a dramatically thinned sedimentary section there (see Driscoll and Diebold, Chapter 20). In both profiles, the Moho can be seen continuously beneath the high, roughly maintaining a constant two-way time, indicating that the ridges are supported by a thickened crust. Between these two profiles, four EW-9501 profiles
577
end just at (apparently) the crest of the ridge-like high. Brute-stack processing of the connecting lines 1299 and 1303 shows that there is little relief along this ridge, and that Moho remains flat beneath it. The southern ridge, though well developed at CDP 12,500 on line 1319, becomes less distinct, and possibly bifurcates as it is mapped to the E N E WSW, along crossings of Conrad MCS lines 119 and 137 (Diebold et al., 1981) and EW-9501 lines 1304, 1302, and 1300 (Fig. 11). In line 1300 (Fig. 5) the basement high is actually a shoulder, at CDP 9000, but a small but perceptible high can be clearly seen in the lower volcanic sequence there. Sixty-five km farther east-northeast, no high can be seen; the magnetic lineation has disappeared, and the upper, dipping volcanic sequence leading to the roughsmooth B" transition has shifted NNW-SSE. Volcanic mounds
Volcanic edifices, some completely covered by sediment, others seen as protruding seamounts, are numerous in the Venezuelan and Colombian Basins. Some of these appear to fall along linear trends. The seamount shown in fig. 4 of Ladd and Watkins (1980) is one of an E-W-trending chain, which includes the large seamount imaged at CDP 2100, line 1293. Other volcanic cones are found parallel to the NE-SW-trending magnetic anomalies mapped by Donnelly (1973) and Watkins and Cavanaugh (1974). As in the case of the seamount at CDP 5900, line 1293, many of these cones (the shape of this one is well determined by multibeam echo sounding and side scan sonar) appear to be surrounded by a now submerged moat, or perhaps more accurately, a thickening of the upper volcanic sequence described above. This thickening might be responsible for some of the elevation of the northeastward basement high. The fact that most of the sonobuoys along this N E - S W trend show clear evidence of a low-velocity zone (Fig. 11) suggests that these low velocities correspond to volcaniclastic material causing the thickening of the upper sequence. These low-velocity zones are seen only along this trend, and they are always found just below the upper, slow extrusive layer. A possible explanation is that the culminative volcanic sequence began with production of volcaniclastics, followed by extensive, mobile flows, whose final expression is the thin flows seen to overlie smooth and rough B" basement alike. Fig. 12 shows the SE end of EW-9501 line 1323, on the flank of the Beata Ridge. Completely buried by sediments, a feature we interpret as a volcanic mound is located at CDP 23,800. The upper volcanic sequence is clearly defined by horizon B" and the distinctive basal reflector. A deeper (7.5 s, two-way time) bright reflector, truncated around CDP 19,700,
578
J. D I E B O L D et al.
o ca o
~
~D
.o
,.r o 9
T
,.r
o ~D
~D
~ ~D
~ "O cD
O r162 ~D
c5
o,I
,,..~
9
,..-, o
rz~r162 , ~
r162
~D
r~
~D
ca
tt3 G'~ er
M U L T I C H A N N E L SEISMIC IMAGES OF V O L C A N I C STRUCTURES IN THE V E N E Z U E L A N BASIN
579
Fig. 10. MCS line 1320, which runs N-S, through DSDP Sites 146/149 and 150. A broad basement high is seen at CDP 5000, and a volcanic edifice appears at CDP 12,500. Sonobuoy 65 (Fig. 8) was deployed near DSDP Site 146/149.
Fig. 11. Map of basement highs, dipping reflectors, the rough-smooth B" boundary and related normal faults. Locations of sonobuoys whose records show low-velocity zones (cf. sonobuoy 65, Fig. 8).
580
J. DIEBOLD et al.
Fig. 12. ESE end of MCS line 1323, Beata Ridge flank. The upper volcanic sequence thickens dramatically toward what is interpreted as a buried seamount. To the west-northwest, flexural rebound has elevated the formerly flat B" horizon.
is interpreted as the reflection from the top of a sill. The post-volcanic deformation and uplift seen in the WNW end of the figure is related to the extensional flexure of the Beata Ridge flank (Driscoll and Diebold, Chapter 20). The sill-like reflector is, essentially, parallel to overlying B", and would therefore (according to our reconstruction) have been horizontal when originally emplaced. The thickness of the upper volcanic sequence doubles towards the volcanic mound. This pattern is frequently observed throughout the Venezuelan Basin, suggesting that the eruptions feeding the mounds started before, or simultaneously with, the flows forming the upper sequence. The association of these volcanic features with the underlying divergent reflectors indicates that these features are long-lived and not just a late-stage phenomenon sourced from small degree partial melts associated with the extension (e.g. decompressional melting) or differentiated magma from the solidifying basalt province. Plateau rifting and volcanoes are features of other oceanic LIPs (Manihiki, Ontong Java, Kerguelen) and may be an integral part of the evolution of oceanic plateaus. Determination of the age and petrologic character of these events would distinguish between late-stage parasitic volcanism versus a more long-lived source of volcanism throughout the development of the basalt province.
DISCUSSION An increasingly important part of the puzzle of the origin of the Caribbean plate is the zone of thin, rough basement, described by Talwani et al. (1977). The EW-9501 profiles show clearly that this crust pre-dates at least the voluminous late-stage volcanism that characterizes the Caribbean oceanic plateau, and may, therefore, typify 'original' Caribbean ocean crust. Rough mapping of this zone (Diebold et al., 1981) was improved by Leroy (1995) and continues in this paper. EW-9501 results indicate that this thin, rough basement extends beneath smooth B" flows at least as far as the leading edge of the underlying early-stage series of dipping reflectors (Figs. 13 and 14). This interpretation is based on the continuity of reflections from the top of the rough crust, on the identifiable character of reflections from the base of the upper volcanic sequence, and on sonobuoy velocities. Sonobuoy velocity analyses do not rule out the possibility that thin, rough crust is sandwiched between the Moho and all of the volcanics of the Venezuelan Basin, but neither do they prove it. The new profiles have also led to an important reinterpretation of the sediments immediately overlying rough B". The thickness of the A"-B" interval increases markedly in the zone of rough basement, particularly to the south and east. This increase is
MULTICHANNEL SEISMIC IMAGES OF VOLCANIC STRUCTURES IN THE VENEZUELAN BASIN
581
5000
6000
9~
seafloor
o 7000 ~
E d 8000 E
~
horizon A"
J
to~can~,_ /
dippingreflectors->'~. ~
early stage
thin, rougn B"crust
9000 10000 11000 -
f
12000 -"
'
'
'
I
.
.
.
.
I
.
.
.
.
!
.
.
.
.
I
'
'
'
'
I
'
I
'
'
'
'
I
'
EW9501 MCS line 1293 0 -
5000 -
10000 -
/ 15000 -
"~" 20000
'
0
~ '
V.E. 5.0 '
'
I
5000
'
'
'
'
I
'
'
10000
'
'
I
15000
'
'
'
'
20000
'
'
'
25000
CDP#
Fig. 13. Line drawings showing our interpretation of the major horizons in MCS line 1293 (Fig. 2) in two-way time (top) and converted to thickness (bottom) using averaged sonobuoy velocities.
due to the presence of a body of horizontally layered sediments that is not seen above the higher-standing smooth B" volcanic basement. Biju-Duval et al. (1978) identified these sediments as a 'transparent' sequence not seen northwestward, in the smooth B" zone, and otherwise characterized by relatively high seismic velocities. This well-laminated deeper sequence fills basement lows and onlaps highs, appearing to be the result of turbiditic, rather than pelagic deposition. In agreement with earlier results, we determine an acoustic velocity for the turbidites, ca. 3.2 km/s, which is distinctively higher than that of the overlying pelagic sediments. This sequence reaches thicknesses as great as 800-1000 m, and thus represents a significant volume of material. Sediments with similar onlapping relationships have been observed to the west, in the Colombian Basin (Lu and McMillen, 1982; Kolla et al., 1984; Bowland and Rosencrantz, 1988; Bowland, 1993) and interpreted there as volcaniclastic turbidites. The stratal geometries and areal extent (at least 50,000 km 2) of these sediments in the Venezuelan Basin, however, suggests strongly that they are of terrigenous origin (see Driscoll and Diebold, Chapter 20). It is likely that the rough crust underlying these similar sediments in the southern Colombian Basin is also thin. Examination of better-quality 1970's vintage MCS data on
the Colombian Basin also reveals elevated sequences of dipping reflectors adjacent to this rough crust, similar to the relationships we see at the roughsmooth B" transition in the Venezuelan Basin. It is not easy to explain the large area of abnormally thin oceanic crust in the Caribbean. A zone of thin (3.3 km) crust in the Labrador Sea was associated with slow spreading rates (6 mm/year) by Srivastava and Roest (1995). This thin crust is broken into tilted blocks, exhibiting a basement topography much rougher than that seen in the Venezuelan Basin, however. Similar, thin crust observed at the passive Iberian margin exists only in a narrow zone, whereas the thin, rough crust we see in the Caribbean occupies an extensive area. A scenario like the one presented by Rosendahl et al. (1992) whereby a passive continental margin is flanked by a broad zone of thinned 'oceanized' continental crust is also possible. The interpreted composite section presented as that paper's fig. 2 also includes a wedge of seaward-dipping reflectors over a necked-down crustal section at the seaward boundary of this thinned crustal zone. One of the observations that supports their notion of an elevated, 'equilibrium' Moho is that some of the diamondpattern apparent faults continue through and below reflection Moho. If many of the 'faults' were ac-
582
J. DIEBOLD et al. 6000 = seafloor
7000
8ooo
horizon A"
9000
" ..
10000
thin, rough B" crust
......" " . Moho
11000
'
'
'
'
I
'
'
'
'
I
j
~
9 ".....:!::::::.....i
'
'
'
'
I
'
'
'
I
'
'
'
'
"
"
I
'
EW9501 MCS line 1300
5000 -
~. 10000
15000
V. E. 5.0
20000
'
0
'
'
'
I
'
'
'
'
5000
I
'
10000 CDP #
'
15000
'
'
'
I
'
20000
Fig. 14. Line drawing interpretation of major horizons in MCS line 1300 (Fig. 5) as picked in two-way time (top) and converted to thickness (bottom).
tually sideswipe reflections from line-oblique basement ridges (and this kind of geometry could be expected from most of the ship tracks in the data set of Rosendahl et al., 1992) then the amount of thinning required would be greatly increased. The absence of such Moho-penetrating reflectors from EW-9501 data could be explained in this way. On the other hand, this amount of extension of the thinned oceanic crust should have generated significant decompression melting and underplating, for which there is no evidence. In fact, the Rosendahl et al. (1992) model involves a volcanic-extrusive transition to pure seafloor spreading, which requires a seaward smoothing of the crustal surface. The transition to thin crust in the Caribbean shows the opposite trend. Even thinner crust (3-4 km) has been recorded at slow spreading ridges in the Arctic (Keen and Barrett, 1972; Jackson et al., 1982). We determine that minimum dimensions for the Venezuelan Basin thin crust are roughly 300 km (E-W) and 350 km (N-S). At a 1 cm/year half-spreading rate (and neglecting subsequent extension) it would take at least 30 m.y. to create this much crust, which is unlikely, given the constraint of reasonable rates of pelagic sedimentation (1-2 cm/ka). For the thin B"
crust to have been created at a much slower rate (such as the 0.5 cm/year, full rate, determined for the Arctic mid-ocean ridge) would exacerbate this problem of the lack of pelagic sediments, significant amounts of which cannot be detected between the rough B" crust and the overlying turbidites. A third possibility is that the crust, once of normal thickness, was pervasively thinned by extension, without any associated volcanism. It is difficult to document such a process elsewhere, however. The thin, rough crust has a tendency for brittle failure along E - W trends, as evidenced in its destruction by normal faulting beneath the Curaqao Ridge (Fig. 4). Elsewhere, extensional features are most concentrated near the rough-smooth B" transition. Tilted blocks, typical of normal faulting, are seen from place to place, and EW-9501 results confirm that the en-echelon, ENE-WSW-trending pattern determined for the 'Central Venezuela Fault Zone' by Biju-Duval et al. (1978) is characteristic of the entire rough-smooth boundary. The faulting of the rough basement is not restricted only to the CVFB. These faults appear to be the result of extensional processes which may exploit pre-existing faults which were, perhaps, formed at the time of
MULTICHANNEL SEISMIC IMAGES OF VOLCANIC STRUCTURES IN THE VENEZUELAN BASIN crustal creation. We suggest that the extensional direction is somewhat oblique to that of the strike of the preexisting faults (and hence, perhaps, to an original seafloor spreading direction) resulting in an en-echelon pattern of ' wrench' -like tilted blocks. This oblique relationship, which allows the Cretaceous flows to run around the ends of the en-echelon fault segments, entirely explains the variations seen in the character of the rough-smooth B" transition. In several locations at which the flows appear to be dammed by a rough basement fault block, the distal edges of the upper flows are tilted back, forming characteristic 'ski-jumps'. The likely mechanism for these is that the underlying block continued to rotate after being covered by a thin Cretaceous flow deposit, tipping it back towards the northwest. In some cases, subsequent flows filled in the resulting trough, and in line 1293, an entire buried 'ski-jump' sequence can be seen between CDPs 17,000-18,000, at 8.5 s two-way time. This process suggests that extension and emplacement of the upper volcanic sequence were penecontemporaneous. That some tilted blocks appear to be capped by (presumably originally sub-horizontal) thin flows requires that tilting and extension continued, at least briefly, after volcanic production in the area had stopped. North-northwest of the present-day roughsmooth boundary, an even more deeply buried series of dipping reflectors is seen, exhibiting the smallerscale, steeper dips and concave-downward curvature characteristic of typical SDRs imaged at the Vcring Plateau escarpment and elsewhere. The top surface of the rough, thin basement exposed to the southeast can be followed almost continuously from the rough-smooth transition to the area in which this 'typical' sequence is seen. All of these features suggest (or, rather, require) that at least the southeastern 75 km of Venezuelan Basin volcanics were deposited as a sequence of flows, with little, if any, sill formation. It appears that volcanic sources were 'uphill', to the northwest, and that the southeastward-thickening of the individual flows was the joint result of local relief, the source area having been uplifted, perhaps as White and McKenzie (1995) suggest, by the thermal effects of a mantle plume, and of the steady subsidence of the underlying weak, thin crust. It is conceivable that the small-scale dipping reflectors of the early volcanic sequence, as seen near CDP 12,500 on line 1293 (Fig. 2), mark the initiation of a seafloor spreading episode during which rough B" crust was formed. It is thus tempting to correlate the ski-jumps with landward-dipping 'D' reflectors observed by Keen and Potter (1995) marking the transition between a broad seaward-dipping sequence and oceanic crust southeast of Nova Scotia. At least two observations argue against (but do not preclude) this scenario. First, the en-echelon patterns
583
of faulting and of the rough-smooth boundary suggest some change in spreading direction between the time that rough crust was created and when it was extended and covered by volcanics. The extension and volcanism are definitely contemporaneous, but the oceanic-like crust seems to have existed for at least a short time. Second, we see, in line 1293, an entire rough-smooth boundary sequence, some 25 km northwest of the current boundary, buried beneath at least 5 km of volcanics forming the subsequent sequence. The EW-9501 profiles show that the volcanic structures imaged beneath the smooth B" horizon form two sequences, distinguished by morphology and separated by a prominent reflecting surface. The upper and lower volcanic bodies both contain dipping reflectors, but these too have distinct morphologies. The uppermost dipping reflectors beneath smooth B", first identified by Hopkins (1973), Ladd and Watkins (1980), Stoffa et al. (1981), and Diebold et al. (1981) form systematic sequences that are quite different than those observed elsewhere at passive continental margins (Hinz, 1981; Mutter et al., 1982; Planke and Eldholm, 1994). Dipping reflectors in the deeper, older Caribbean sequence, however, resemble these 'classic' seaward-dipping reflectors (SDRs) much more closely. Similar features are found at some oceanic plateaus, particularly Kerguelen (Ramsay et al., 1986; Coffin et al., 1986; Colwell et al., 1988; Schlich et al., 1993), a plateau whose size is comparable with that of the Caribbean Plateau. As in the case at Kerguelen, the sources of the older dipping sequences in the Venezuelan Basin appear to be basement highs, beneath which reflectors, if any, are horizontal. Similar, too, is the organization of the basement highs, which form several roughly parallel ridges, which lie perpendicular to the direction of extension. The plate tectonic setting of Kerguelen's origin is much clearer than that of the Caribbean, however, since it is flanked by reliably identified seafloor spreading magnetic anomalies in the southeast Indian Ocean. Unlike the examples just cited, the Caribbean Plateau features an overlying thicker, more extensive set of flows, each apparently more voluminous and mobile than the individual flows of the underlying 'classic' sequence. We are prompted, therefore, to envision a two-phase volcanic history, in which an initial period of slow extension and crustal thickening, typified by the construction of local volcanic highs, with flanking dipping flows, is superseded by a more vigorous episode of extension and magmatic emplacement. This second phase could also be subdivided, as indicated by the low-velocity zones found within it, and by low seismic velocities and high fluidity evident in the uppermost flow sequences. The upper series of flows can be seen to thicken towards
584 the west, and this thickening is particularly apparent between the underlying highs, which appear to be sources of the early-stage dipping reflectors. This trend continues in the EW-9501 reflection profiles on the flank of the Beata Ridge, indicating that the volume of second-stage volcanics increases toward the west. An upper sequence of volcanics has also been imaged on the Ontong Java Plateau (Hagen et al., 1993) but those data unfortunately lacked the penetration to reveal underlying structure. Bercovici and Mahoney (1994) proposed, based on Ar-Ar age dating, that the Ontong Java Plateau was emplaced in two phases, the first from 122 to 120 Ma, and a second from 90 to 88 Ma. Therefore, it is conceivable that the upper volcanic sequence there is 30-34 m.y. older than the underlying material. Tarduno et al. (1991), however, hypothesized that the Ontong Java Plateau was emplaced over only 3 m.y., during the early Aptian. Northwest of the rough-smooth boundary, NESW-trending grabens and monoclinic normal faults, previously mapped by Case and Holcombe (1980) and others, may be the result of post-volcanic extension, or perhaps are simply the result of thermal subsidence. It is our interpretation that sonobuoyderived low-velocity zones (Fig. 11) near the flanks of the volcanic cones situated along the northwesterly edge of our Venezuelan Basin seismic grid are volcaniclastic aprons of those volcanic edifices, which were produced early in the final stage of extrusive production, and emplaced on the upper, perhaps sediment-covered surface of the previous sequence. Low-velocity zones just landward of seaward-dipping reflectors along North Atlantic rifted margins are shown in the synthesis by Eldholm and Grue (1994), but their origin was not identified, and no nearby volcanic mounds were imaged. It is frequently observed that around these volcanic edifices (either seen as seamounts or submerged beneath sediments) the upper surface of the lower sequence is depressed. Since it is unlikely that the sheer weight of these relatively minor volcanic cones is sufficient to result in crustal flexure, we conclude that while they have subsided, as the consequence of a combination of thermal contraction and the withdrawal of the underlying magmatic source, it is also likely that the volcanic cones, and other sources of late-stage volcanism, were located in topographic lows between the highs typifying early-stage crustal thickening.
Magmatic and tectonic development of the Caribbean On the basis of the new MCS data acquired in the eastern Caribbean (Fig. 1), together with previous collected data from the region, we have de-
J. DIEBOLD et al. veloped a volcano-tectonic model for the magmatic and tectonic development of the crust that floors the Venezuelan Basin (Fig. 15). This model is consistent with the available age, compositional, and geophysical data, and is compatible with the 'mobilist' model for the development of the Caribbean. We assume that the proto-Caribbean crust was formed by seafloor spreading in Late Jurassic/Early Cretaceous time in the eastern Pacific. An unresolved issue is the initial thickness of the proto-Caribbean crust. We have assumed normal oceanic crust thicknesses of approximately 6 km for the proto-Caribbean because the seismic reflection data indicate that faulting and extension were concomitant with, and subsequent to, the magmatic activity. However, we appreciate the fact that the initial crustal thickness could be substantially less if formed at an ultra-slow spreading ridge, as discussed above. Prior to the Senonian, crustal intrusion and eruption of basaltic flows began (Fig. 15a). The fact that the sub-basement highs we feel were created during this process are relatively transparent in seismic reflection profiles may result from their being composed principally of vertical dikes, which cannot be imaged with conventional profiling techniques. This magmatism was accompanied, at least in its late stages, by extension and crustal thinning, producing some of the structure imaged in the upper surface of this sequence. The extent and volume of magmatic products imaged in the EW-9501 data are consistent with a hotspot or mantle plume source. It is difficult to determine whether extensional deformation predated the impact of the plume head and associated magmatic activity (Galapagos hotspot; Duncan and Hargraves, 1984) or vice versa. Numerous hypotheses concerning the interaction of plume heads with both active and passive rifting of the overlying lithosphere have been advanced (Coffin and Eldholm, 1994). One school of thought is that the impact of the plume head on the mechanical boundary layer at the base of the lithosphere causes large-scale melting and the transfer of stress from the plume to the overlying lithospheric plate (i.e. active rifting). Modifications on this theme have been put forth by White and McKenzie (1995) in which they propose that the magmatic activity is a response to the thermal uplift and associated decompression melting. The alternative hypothesis is that regions of the lithosphere deformed by plate reorganizations act as foci or escaping points for hot regions of the mantle, allowing melts to breach the lithosphere (Anderson et al., 1992). In this scenario, the crustal thinning would be related to subduction along the eastern plate and the associated back-arc extension. Because of the extensive overprinting by the magmatic activity, it is difficult to ascertain whether the plume head or lithospheric extension was the cause or response.
MULTICHANNEL SEISMIC IMAGES OF VOLCANIC STRUCTURES IN THE VENEZUELAN BASIN
585
Fig. 15. Cartoon showing the volcanic and tectonic evolution of the Venezuelan Basin. (a) Original, normal-thickness oceanic crust (right) is thickened by intrusion (vertical dikes) with mound-building extrusion producing sloping flanks, and through underplating by depleted source material. Original oceanic crust is indicated by gray shading. (b) A second phase of volcanism is marked by increased extension, re-intrusion by sills, production of massive extrusives forming the smooth B" horizon, and additional underplating. (c) Further extension is focused in the west, resulting in rifting of the entire thickened crust and flexural uplift of Beata Ridge.
The structure of the Caribbean basalt province, as imaged by EW-9501 data, indicates at least two episodes of magmatic activity which is consistent with the age dating information (Donnelly et al., 1990). The early stage was more localized with steeper-dipping reflectors flanking basement highs (Fig. 15a); the late stage is regionally more extensive with localized brittle deformation (Fig. 15b). Based on seismic reflection and refraction, the plate was thickened further during this second phase by extrusion of lava flows over a vast area, intrusion of dikes and sills, and underplating by residual mantle from the melting event. The rough-smooth B" boundary imaged in seismic profiles is, in this model, the edge of the basalt province overlying the older plate (Fig. 15b). The edge of the basalt province is also coincident with an abrupt shoaling of Moho (Figs. 2 and 13). The divergent wedge that is well imaged below B" in line 1293 (Fig. 2) and the abrupt shoaling of Moho may be controlled by a northwestward
dipping fault system (Driscoll and Diebold, Chapter 2O). The dipping reflector packages observed in both the early- and late-stage volcanic sequences in the Venezuelan Basin exhibit structural similarities to seaward-dipping reflectors (SDRs) observed along other volcanic margins (Mutter et al., 1982; Coffin and Eldholm, 1994). The commonly cited model for emplacement of the SDRs is that thermal dynamic support uplifts the magmatic source above sea level with extrusions flowing away from the source region (Mutter et al., 1982). With time, the source region evolves into a spreading center and when the anomalous heat dissipates, the plate cools, contracts, and subsides. The primary dip of the volcanics is reversed because of the subsidence and the SDRs diverge, or thicken, toward the extinct source. In the case of the eastern Caribbean, however, we know that the adjacent rough crust was present at the time of the late-stage magmatic em-
586 placement forming the dipping wedges northeast of the rough-smooth B" boundary, because of the contemporaneous faulting and overlap with tongues of smooth basement. The thickness and the distribution of the dipping reflectors varies between the earlyand late-stage basalt flows. We propose that these observed variations record changes in the volume of magma supplied to the region through time. The internal geometry of the divergent volcanic wedges and the intercalated limestones indicate that the dipping divergent wedge was emplaced in the marine environment. Consequently, we propose that the formation of ' seaward' -dipping reflectors can occur in the marine environment, and that changes in the flow morphology record variations in magma production. These effects may explain the variety of scales at which continental margin SDRs have more recently been observed (e.g. Skogseid and Eldholm, 1989). Due to the segmented nature of the NNE fault systems forming the rough-smooth boundary in the Venezuelan Basin, tongues of basalt flows have flowed subparallel to the down-dropped hanging walls, overrunning rough basement towards the southeast, generating a complex but predictable pattern between the rough and smooth B" basement. Our hypothesis, that the source of the basalts is either along-strike or up-dip, is consistent with new models for the evolution of similar dipping reflectors on the Kerguelen Plateau (Schlich et al., 1993). Furthermore, recent studies from the Pacific indicate that as in the case of some continental extrusives (Self et al., 1996) marine basalt flows develop an outer carapace (chill margin) that allows the flows to be insulated and consequently flow over great distances (Clague et al., 1990; Abrams et al., 1993). The earlier models proposed that the flows were subaerial because they could not reconcile the vast spatial distribution of the volcanics with inferred high-viscosity submarine flows (Mutter et al., 1982). The location and N E - S W distribution of the seamounts appear to have a structural control. Most of the seamounts that have been mapped are in the northern half of the Venezuelan Basin (Matthews and Holcombe, 1976) but as the EW-9501 data show (cf. Fig. 12) many seamounts in the southern half of the basin have been buried by sediments and are, thus, difficult to map. We hypothesize that seamounts there tend to lie along the same E N E - W S W trends as the basement highs typifying the lower volcanic sequence. Late-stage faulting across the region with the same trends would provide conduits for any molten material. Nevertheless the underlying dipping reflectors beneath the volcanic constructs are indicative of their longevity and preclude a simple late-stage parasitic origin for the seamounts.
J. DIEBOLD et al. CONCLUSIONS The concept of the 'Caribbean Plateau' as a monolithic allochthon of crust thickened by Cretaceous flood basalts is laid to rest by multichannel seismic data acquired aboard the NSF research vessel Ewing in 1995. In the search to define the location and structure of the edge of the Caribbean volcanics, we found two edges, whose stratigraphic relationships show that the volcanism occurred in at least two distinct phases. The upper series of volcanics are the only ones which have been sampled, at least in the Venezuelan Basin. The underlying phase must certainly be chronologically earlier, though how much earlier is difficult to quantify. Some of the crustal elements and their stratigraphic relationships, described above, provide some constraints on the relative sequence of events. Some, if not all, of the thin, rough crust now seen as basement in the southeast part of the basin must have existed before the emplacement of the upper volcanic sequence. This is clear, because thin stringers of the uppermost flows can be followed from where they overlie 'smooth' areas to where they are seen as a thin covering of what is in every other respect thin, 'rough' crust. Where that covering is seen, it is evident that an intervening layer of sediment, if any, is very thin, implying that the age difference is not great. That the thin crust was not being created simultaneously (say, in a rifting margin) is ruled out by the evidence that the thin crust was being mildly extended, via block faulting, while the upper sequence of volcanics was spilling onto it. Similar relationships between thin, rough crust and the lower sequence of volcanics, whose southeasterly edge forms structures very similar to the wedges of seaward-dipping reflectors observed at many passive margins, are not clear in the data. It is certainly possible that the generation of the thin crust and these wedges of dipping reflectors was coincident. We hypothesize that the southeastern edge of the early-stage volcanic sequence in the Venezuelan Basin, with its 'seaward-dipping' reflectors is the rifted margin, either of a pre-existing submarine plateau of modest thickness, or perhaps even of a re-rifted segment of normal oceanic crust. The thin, rough crust would then have been formed, perhaps with normal thickness, by seafloor spreading away from this margin. If at this point the crust in question was still in the Pacific Ocean, the conjugate margin would later have been subducted beneath the western edge of South America. At some time thereafter, the final phase of volcanic extrusion began. Although re-extension was taking place in the new, rough crust, volcanic sources were located further to the northwest, supporting the notion that the upper (and only sampled) volcanic sequence stemmed from the drifting Caribbean plate's
MULTICHANNEL SEISMIC IMAGES OF VOLCANIC STRUCTURES IN THE VENEZUELAN BASIN crossing of the Galapagos hotspot (Duncan and Hargraves, 1984). At some short time after volcanic extrusion had ended (at least in the SE Venezuelan Basin) the Caribbean plate became close enough to a source of voluminous turbidites that the thin, rough crust, along with the lower parts of what is now the r o u g h - s m o o t h B" boundary became covered by these low-filling, stratified sediments. The facts that in places, turbidite layering downlaps on tilted, smooth B"-covered blocks, and that both the turbidites and the thin stringers of smooth B" material seem to directly overlie rough B", provide important constraints on the timing of the formation of thin rough B" basement, its extensional wrench faulting, its overspreading by smooth B" volcanics, and the subsequent infilling of bathymetric lows by (presumably terrigenous) turbidites. Coring, by deep-sea drilling, will provide the sediment and basement samples needed to answer the few fundamental questions that stand between us and full understanding of the tectonic history of the Caribbean basins. What is the age and origin of thin, rough B" basement? Was it formed thin, or has it been subsequently extended? What is the age and provenance of the sediments directly overlying it, i.e. was rough B" formed in an open Pacific Ocean environment? Could it have been produced in a back-arc basin of the proto-Caribbean? What sediments are sandwiched between the smooth B" stringers and the underlying rough B" crust? What are the ages and provenances of the thick turbiditic layers that overran thin, rough B" and appear to have crossed the r o u g h - s m o o t h B" boundary at some points? Our interpretation of the seismic reflection data predicts that the volcanic thickening of the Venezuelan Basin has been due primarily to the emplacement of extrusives (therefore, in bottom-up sequence) and that the final, youngest extrusive event had an extremely large areal extent, spreading as thin, mobile flows even out over the previously uninvaded rough B" crust. Testing of this hypothesis, and the refinement of it, can be accomplished by drilling and sampling the easily accessed upper layer, but this must be done in several places. Much more of a puzzle, and likely to remain so, is the nature and age of the early-stage volcanics, imaged as a clearly definable structural sequence buried by several kilometers of late-stage extrusives and separated from it by sedimentary and/or volcaniclastic material, whose age and makeup is entirely unknown. Since the upper sequence appears to thicken westward, sampling (i.e. drilling) of the earlier structures might be more effectively carried out to the east, far from any locale previously sampled by DSDP or ODP.
587
ACKNOWLEDGEMENTS
This geophysical investigation in the Caribbean region was supported by the National Science Foundation under grant OCE-93-02578. The manuscript benefited greatly from reviews by R. Duncan and M. Coffin. Lamont-Doherty Earth Observatory Contribution no. 5683.
REFERENCES
Abrams, L.J., Larson, R.L., Shipley, T.H. and Lancelot, Y., 1993. Cretaceous volcanic sequences and Jurassic oceanic crust in the East Mariana and Pigafetta basins of the western Pacific. AGU Geophys. Monogr., 77: 77-101. Anderson, D.L., Zhang, Y.-S. and Tanimoto, T., 1992. Plume heads, continental lithosphere, flood basalts, and tomography. In: B.C. Storey, T. Alabaster and R.J. Pankhurst (Editors), Magmatism and the Causes of Continental Break-Up. Geol. Soc. London Spec. Publ., 68: 99-124. Beets, D.J., Klaver, G., Beunk, F.F., Kieft, C. and Maaskant, E, 1982. Picrites as parental magma of MORB-type tholeiites. Nature, 296:341-343. Bercovici, D. and Mahoney, J., 1994. Double flood basalts and plume head separation at the 600-kilometer discontinuity. Science, 266: 1367-1369. Biju-Duval, B., Mascle, A., Montadert, L. and Wanneson, J., 1978. Seismic investigation in the Colombia, Venezuela and Granada Basins, and on the Barbados Ridge for future IPOD drilling. Geol. Mijnbouw, 57:105-116. Bowland, C.L., 1993. Depositional history of the 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 of the western Colombian Basin, Caribbean Sea. Geol. Soc. Am. Bull., 100: 534-546. Case, J.E. and Holcombe, T.L., 1980. Geologic-tectonic map of the Caribbean region, scale 1:2,500,000. In: Miscellaneous Investigations, Ser. Map I-1100, U.S. Geol Survey. Chen, Y.J., 1992. Oceanic crustal thickness versus spreading rate. Geophys. Res. Lett., 19: 753-756. Christofferson, E., 1973. Linear magnetic anomalies in the Colombia Basin, central Caribbean Sea. Geol. Soc. Am. Bull., 84: 3217-3230. Christofferson, E., 1976. Colombian Basin magmatism and Caribbean plate tectonics. Geol. Soc. Am. Bull., 87: 12551258. Clague, D.A., Holcombe, R.T., Sinton, J.M., Detrick, R.S. and Torresan, M.E., 1990. Pliocene and Pleistocene alkalic flood basalts on the seafloor north of the Hawaiian islands. Earth Planet. Sci. Lett., 98: 175-191. Coffin, M.E and Eldholm, O., 1994. Large igneous provinces: crustal structure, dimensions and external consequences. Rev. Geophys., 32: 1-36. Coffin, M.F., Davies, H.L. and Haxby, W.F., 1986. Structure of the Kerguelen Plateau province from Seasat altimetry and seismic reflection data. Nature, 324:134-136. Colwell, J.B., Coffin, M.F., Pigram, C.J., Davies, H.L., Stagg, H.M.J. and Hill, EJ., 1988. Seismic stratigraphy and evolution of the Raggatt Basin, southern Kerguelen Plateau. Mar. Pet. Geol., 5: 75-88. Diebold, J.B., Stoffa, EL., Buhl, E and Truchan, M., 1981. Venezuelan Basin crustal structure. J. Geophys. Res., 86: 7901-7923. Donnelly, T.W., Melson, W., Kay, R. and Rogers, J.J.W., 1973.
588 Basalts and dolerites of Late Cretaceous age from the Central Caribbean. Init. Rep. DSDP, 15:989-1011. Donnelly, T.W. and 10 others, 1990. History and tectonic setting of Caribbean magmatism. Decade of North American Geology, Vol. H, Geological Society of America, Boulder, CO, pp. 339-374. Driscoll, N.W., Hogg, J.R., Karner, G.D. and Christie-Blick, N., 1995. Extensional tectonics in the Jeanne d'Arc basin: implications for the timing of break-up between Grand Banks and Iberia. In: R.A. Scrutton, M.S. Stoker, G.B. Shimmield and A.W. Tudhope (Editors), The Tectonics, Sedimentation and Palaeoceanography of the North Atlantic Region. Geol. Soc. London Spec. Publ., 90: 1-28. Duncan, R.A. and Hargraves, R.B., 1984. Plate tectonic evolution of the Caribbean region in the mantle reference frame. Geol. Soc. Am. Mem., 162:81-94. Echeverria, L.M., 1980. Tertiary or Mesozoic komatiites from Gorgona Island, Colombia: field relations and geochemistry. Contrib. Mineral. Petrol., 73: 253-266. Edgar, N.T., Ewing, J.I. and Hennion, J., 1971. Seismic refraction and reflection in Caribbean Sea. Am. Assoc. Pet. Geol. Bull., 55: 833-870. Eldholm, O. and Grue, K., 1994. North Atlantic volcanic margins: dimensions and production rates. J. Geophys. Res., 99: 2955-2968. Ewing, J., Antoine, J. and Ewing, M., 1960. Geophysical measurements in the western Caribbean Sea and in the Gulf of Mexico. J. Geophys. Res., 65: 4087-4126. Frisch, W., Meschede, M. and Sick, M., 1992. Origin of the Central American ophiolites: evidence from paleomagnetic results. Bull. Geol. Soc. Am., 104:1301-1314. Ghosh, N., Hall, S.A. and Casey, J.E, 1984. Seafloor spreading anomalies in the Venezuelan Basin. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162: 65-80. Hagen, R.A., Mayer, L.A., Mosher, D.C., Kroenke, L.W., Shipley, T.H. and Winterer, E.L., 1993. Basement structure of the northern Ontong Java Plateau. Proc. ODE Sci. Results, 130: 23-31. Hinz, K., 1981. A hypothesis on terrestrial catastrophes. Wedges of very thick oceanward dipping layers beneath passive continental margins. Geol. Jahrb., E22: 3-28. Hinz, K., Schluter, H.-U., Grant, A.C., Srivastava, S.E, Umpleby, D. and Woodside, J., 1979. Geophysical transects of the Labrador Sea: Labrador to southwest Greenland. Tectonophysics, 59:151-183. Hopkins, H.R., 1973. Geology of the Aruba Gap Abyssal Plain near DSDP site 153. Init. Rep. DSDP, 15: 1039-1050. Houtz, R.E. and Ludwig, W.J., 1977. Structure of Colombia Basin, Caribbean Sea, from profiler-sonobuoy measurements. J. Geophys. Res., 82: 4861-4867. Jackson, H.R., Reid, I. and Falconer, R.K.H., 1982. Crustal structure near the Arctic mid-ocean ridge. J. Geophys. Res., 87: 1773-1783. Keen, C.E. and Barrett, D.L., 1972. Seismic refractions studies in Baffin Bay: and example of a developing ocean basin. Geophys. J. R. Astron. Soc., 30: 253-271. Keen, C.E. and Potter, D.E, 1995. The transition from a volcanic to a nonvolcanic rifted margin off eastern Canada. Tectonics, 14: 359-371. Kent, G.M., Detrick, R.S., Swift, S.A. and Collins, J.A., 1997. Evidence from Hole 504B for the origin of dipping events in oceanic crustal reflection profiles as out-of-plane scattering from basement topography. Geology, 25: 131-134. Klaver, G.T., 1987. The Curacao lava formation: an ophiolitic analogue of the anomalous thick layer 2B of the mid-Cretaceous oceanic plateaus in the western Pacific and central
J. D I E B O L D et al. Caribbean. GUA (Univ. of Amsterdam) Papers of Geology, Ser. 1, No. 27, 168 pp. Kolla, V., Buffler, R.T. and Ladd, J.W., 1984. Seismic stratigraphy and sedimentation of Magdalena fan, southern Colombian Basin, Caribbean Sea. Am. Assoc. Pet. Geol. Bull., 68: 316332. Ladd, J.W. and Watkins, J.S., 1980. Seismic stratigraphy of the western Venezuela Basin. Mar. Geol., 35:21-41. Leroy, S., 1995. Structure et origine de la plaque Caraibe; implications Geodynamiques. PhD Thesis, Univ. Paris IV, pp. 95-32. Lu, R.S. and McMillen, K.J., 1982. Multichannel seismic survey of the Colombia Basin and adjacent margins. In: J.S. Watkins and C.L. Drake (Editors), Studies in Continental Margin Geology. Am. Assoc. Pet. Geol. Mem., 34: 395-410. Ludwig, W.J., Houtz, R.E. and Ewing, J.I., 1975. Profilersonobuoy measurements in Colombian and Venezuelan Basins, Caribbean Sea. Am. Assoc. Pet. Geol. Bull., 59:115-123. Matthews, J.E. and Holcombe, T.L., 1976. Regional geological/geophysical study of the Caribbean Sea (Navy ocean area NA 9). Naval Oceanographic Office ref. Publ. RP 3, Washington, DC. Maurrasse, E, Husler, J., Georges, G., Schmitt, R. and Damond, E, 1979. Upraised Caribbean sea floor below acoustic reflector Btt at the southern peninsula of Haiti. Geol. Mijnbouw, 58: 71-83. McCarthy, J., Mutter, J.C., Morton, J.L., Sleep, N.H. and Thompson, G.A., 1988. Relic magma chamber structures preserved within Mesozoic North Atlantic crust? Geol. Soc. Am. Bull., 100: 1423-1436. Muller, M.R., Robinson, C.J., Minshull, T.A., White, R.S. and Bickle, M.J., 1997. Thin crust beneath ocean drilling program borehole 735B at the Southwest Indian Ridge? Earth Planet. Sci. Lett., 148: 93-107. Mutter, J.C. and Karson, J.A., 1992. Structural processes at slow-spreading ridges. Science, 257: 627-634. Mutter, J.C., Talwani, M. and Stoffa, E, 1982. Origin of seawarddipping reflectors in oceanic crust off the Norwegian margin by 'sub-aerial seafloor spreading'. Geology, 10:131-134. Planke, S., 1994. Geophysical response of flood basalts from analysis of wire line logs: Ocean Drilling Program site 642, Vcring volcanic margin. J. Geophys. Res., 99: 9279-9296. Planke, S. and Eldholm, O., 1994. Seismic response and construction of seaward dipping wedges of flood basalts: VCring margin. J. Geophys. Res., 99: 9263-9278. Ramsay, D.C., Colwell, J.B., Coffin, M.E, Davies, H.L., Hill, EJ., Pigram, C.J. and Stagg, H.M.J., 1986. New findings from the Kerguelen Plateau. Geology, 14: 589-593. Reid, I. and Jackson, H.R., 1981. Oceanic spreading rate and crustal thickness. Mar. Geophys. Res., 5: 165-172. Rosendahl, B.R., Meyers, J., Groschel, H. and Scott, D., 1992. Nature of the transition from continental to oceanic crust and the meaning of reflection Moho. Geology, 20:721-724. Schlich, R., Rotstein, Y. and Schaming, M., 1993. Dipping basement reflectors along volcanic passive margins new insight using data from the Kerguelen Plateau. Terra Nova, 5: 157-163. Self, S., Thordarson, T., Keszthelyi, L., Walker, G.EL., Hon, K., Murphy, M.T., Long, E and Finnemore, S., 1996. A new model for the emplacement of Colombia River basalts as large, inflated pahoehoe lava flow fields. Geophys. Res. Lett., 23: 2689-2692. Sinton, C.W., Duncan, R.A., Storey, M., Lewis, J. and Estrada, J.J., 1998. An oceanic flood basalt province within the Caribbean plate. Earth Planet. Sci. Lett., 155:221-235. Skogseid, J. and Eldholm, O., 1989. Vcring Plateau continental margin: seismic interpretation, stratigraphy, and vertical movements. Proc. ODE Sci. Results, 104: 993-1030.
M U L T I C H A N N E L SEISMIC I M A G E S OF V O L C A N I C S T R U C T U R E S IN THE V E N E Z U E L A N BASIN Spadea, E, Espinosa, A. and Orrego, A., 1989. High-Mg extrusive rocks from the Romeral zone ophiolites in the southwestern Colombian Andes. Chem. Geol., 77: 303-321. Srivastava, S.E and Keen, C.E., 1995. A deep seismic reflection profile across the extinct Mid-Labrador Sea spreading center. Tectonics, 14: 372-389. Srivastava, S.E and Roest, W.R., 1995. Nature of thin crust across the southwest Greenland margin and its bearing on the location of the ocean-continent boundary. In: E. Banda, M. Torne and M. Talwani (Editors), Rifted Ocean-Continent Boundaries. NATO ASI Ser. C, Vol. 463, pp. 95-120. Stoffa, EL., Mauffret, A., Truchan, M. and Buhl, E, 1981. Sub-B tl layering in the southern Caribbean: the Aruba Gap and Venezuela Basin. Earth Planet. Sci. Lett., 53:131-146. Talwani, M., Windisch, C., Stoffa, EL., Buhl, E and Houtz, R.E., 1977. Multichannel seismic study in the Venezuelan Basin and the Curacao Ridge. In: Island Arcs, Deep Sea Trenches and Back-Arc Basins. AGU, Maurice Ewing Series, I, pp. 83-98. Tarduno, J.A., Sliter, W.V., Kroenke, L., Leckie, M., Mayer, H., Mahoney, J.J., Musgrave, R., Storey, M. and Winterer, E.L., 1991. Rapid formation of Ontong Java Plateau by Aptian mantle plume volcanism. Science, 254: 399-403. Waters, A.C., Myers, C.W., Brown, D.J. and Ledgerwood, R.K., 1981. Colombia Plateau with special emphasis on the basalt stratigraphy of the Pasco Basin. In: K.V. Subbaro and R.N. Sukheswala (Editors), Deccan Volcanism and Related Basalt
589
Provinces in Other Parts of the World. Geol. Soc. India Mem., 3: 19-44. Watkins, J. and Cavanaugh, T., 1974. Implications of magnetic anomalies in the Venezuelan Basin. Transactions of the 7th Caribbean Geological Conference, Guadalupe, pp. 129-138. Weidmann, J., 1978. Ammonites from the Curacao Lava Formation, Curacao, Caribbean. Geol. Mijnbouw, 57:361-364. Wildberg, H., 1983. Die Magmatite des Nicoya-Komplexes, Costa Rica, Zentralamerika. Dissertation, Wilhelms-Univ., Munster, 168 pp. Wildberg, H., 1984. Die Nicoya-Komplex, Costa Rica, Zentralamerika: Magmatismus und Genese eines polygenetischen Ophiolith-Komplexes. Munster. Forsch. Geol. Paleontol., 62: 123. White, R.S. and McKenzie, D., 1995. Mantle plumes and flood basalts. J. Geophys. Res., 100: 17,543-17,585. White, R.S., McKenzie, D. and O'Nions, R.K., 1992. Oceanic crustal thickness from seismic measurements of rare earth element inversions. J. Geophys. Res., 97: 19,683-19,715. Whitmarsh, R.B., Miles, ER. and Mauffret, A., 1990. The ocean-continent boundary off the western continental margin of Iberia, I. Crustal structure at 40~ Geophys. J. Int., 103: 509-531. Zehnder Mutter, C., 1992. DSDP Hole 504B: are gabbros in sight? JOI/USSAC Newsl., 5: 4, 5, 12.
Chapter 20
Tectonic and Stratigraphic Development of the Eastern Caribbean" New Constraints from Multichannel Seismic Data
N E A L W. D R I S C O L L and J O H N B. D I E B O L D
New high-resolution multichannel seismic reflection data collected in the eastern Caribbean during R/V Ewing cruise 9501 imaged both the crustal structure and overlying stratigraphic successions. On the basis of these new MCS data, we define the geologic development of the Beata Ridge and Venezuelan Basin. The proto-Caribbean crust was formed by seafloor spreading in Late Jurassic-Early Cretaceous time. Prior to the Senonian, widespread and rapid eruption of basaltic flows began in concert with extensional deformation of the proto-Caribbean crust. Large divergent volcanic wedges observed along the rough-smooth B" boundary are coincident with the abrupt shoaling of Moho and appear to be bounded by a large northwest-dipping fault system. The location of the major extensional deformation migrated through time from the Venezuelan Basin to the western flank of the Beata Ridge. Extensional unloading of the Beata Ridge footwall caused uplift and rotation of the ridge and collapse of its hangingwall (i.e., Hess Escarpment). Sediment thicknesses and stratal geometry observed across the Venezuelan Basin and Beata Ridge suggest that the majority of the 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 bending and subduction of the Caribbean Plate along the Muertos Trough south of Puerto Rico to compressional deformation and obduction of the Caribbean Plate south of Hispaniola. We propose that this difference in deformational style is, in part, a consequence of the thicker crust on the Beata Ridge, which is more resistant to subduction. Coincident with the rough-smooth crustal boundary in the Venezuelan Basin is a marked change in sediment thickness. The stratal geometry and spatial distribution of the basal sequence suggest that the Late Cretaceous to Early Miocene sediments are terrigenous deposits derived from South America. We propose that during the Late Cretaceous to Early Miocene the dominant drainage in South America was a northward-flowing axially parallel fluvial system that drained the Maracaibo-Peruvian foreland basin east of the Andes and supplied sediment to the Venezuelan Basin. The middle Tertiary uplift and deformation along the northern South America Plate boundary blocked this axial-parallel fluvial network. As a result of the blockage, the axial-parallel rivers were dammed or diverted eastward. This blockage, together with the renewed uplift of the Central and Eastern Cordillera, supplied abundant sediment to the developing foreland basin and other paleo-structural lows to the east. The consequent regrading of the fluvial systems as the foreland basin was filled allowed drainage systems to flow east across the South American continent (e.g., Amazon and Orinoco) and deliver sediment to the Atlantic Ocean. The seismic reflection data also imaged an Eocene-Early Miocene current-controlled drift deposit which reflects the movement of bottom currents from the eastern Pacific to the Caribbean during this period. The gradual shoaling of the Central American Isthmus in Late Oligocene-Early Miocene time closing the gateway is consistent with the diminished occurrence of current-controlled features upsection and paleo- deep and intermediate water geochemistry in DSDP cores.
T h e h i s t o r y o f the C a r i b b e a n ' P l a t e ' is m u c h deb a t e d w i t h the m a i n c o n t r o v e r s y c e n t e r i n g a r o u n d
N o r t h and S o u t h A m e r i c a ( W i l s o n , 1966; M a l f a i t and D i n k e l m a n , 1972; P i n d e l l and D e w e y , 1982; B u r k e et al., 1984; D u n c a n and H a r g r a v e s , 1984; P i n d e l l and Barrett, 1990). In the 'fixist' m o d e l , the
the t e c t o n i c r e l a t i o n s h i p o f the C a r i b b e a n to the
p r e s e n t plate d e v e l o p e d m a i n l y in situ and experi-
s u r r o u n d i n g N o r t h A m e r i c a n and S o u t h A m e r i c a n
e n c e d o n l y m i n o r C e n o z o i c m o t i o n w i t h r e s p e c t to
plates. In the ' m o b i l i s t ' m o d e l , the C a r i b b e a n P l a t e
the A m e r i c a s ( K l i t g o r d and S c h o u t e n ,
is p u r p o r t e d to h a v e f o r m e d in the e a s t e r n Pacific
nelly, 1989). T h e d e b a t e c o n t i n u e s , in large part,
and s u b s e q u e n t l y m o v e d e a s t w a r d w i t h r e s p e c t to
b e c a u s e o f the structural c o m p l e x i t i e s b o t h o n s h o r e
INTRODUCTION
Caribbean Basins. Sedimentary Basins of the World, 4 edited by E Mann (Series Editor: K.J. Hsti), pp. 591-626. 9 1999 Elsevier Science B.V., Amsterdam. All rights reserved.
1987; D o n -
592 and offshore, the limited high-quality marine data (e.g., seismic, drilling, and magnetic data), and, most importantly, the fact that few studies have integrated the wealth of onshore and offshore data to establish an internally consistent model for the development of the Caribbean region. A related and equally puzzling problem is the crustal thickness variations observed across the Caribbean Plate, ranging from 4 km in the Venezuelan Basin to 12 km across the Beata Ridge. The origin of the thickened oceanic crust in the Caribbean and its relation to the tectonic development of the Caribbean Plate remains poorly understood. Duncan and Hargraves (1984), using a plate tectonic model tied to the fixed hotspot reference frame, proposed that the Caribbean Plate was located over the initiation of the Galapagos mantle plume during the Late Cretaceous. The excess volcanism associated with the mantle plume increased the oceanic crustal thickness and formed the Caribbean Cretaceous basalt province. Subsequent northeastern movement of the Farallon Plate brought the resultant oceanic plateau from the Pacific into the Caribbean. In this model, the remnant hotspot trail was consumed by plate subduction at the Central American arc. However, although the Galapagos hotspot scenario explains the unusual thick crust in the Caribbean, it fails to
N.W. DRISCOLL and J.B. DIEBOLD account for both the origin of the faulted and highly attenuated crust in the Venezuelan Basin and the deformation across the Beata Ridge and surrounding region. Because of the limited constraints, the timing and origin of the tectonic deformation across the Beata Ridge and surrounding region also remains controversial. Recent studies have proposed that a renewed phase of deformation occurred in the Miocene and that deformation is still ongoing in the region of the Beata Ridge (Mauffret and Leroy, 1997; Mauffret and Leroy, Chapter 21). In the Mauffret and Leroy model, the Colombian Plate is presently overthrusting the Venezuelan Plate with the thrust front being located along the eastern flank of the Beata Ridge (Mauffret and Leroy, 1997). Such a two-plate kinematic model for the Caribbean was first proposed by Dewey and Pindell (1985), and requires differential motion between the Venezuelan and Colombian basins. This plate kinematic model predicts that deformation across the Beata Ridge accommodated the differential motion between the eastern and western Caribbean plates (Fig. 1). Earlier studies also proposed that the Beata Ridge was uplifted after the Eocene (i.e., post-A") on the basis of seismic reflection data (Fox et al., 1970). An alternative hypothesis, the single-plate model for the Caribbean,
Fig. 1. Map of the Caribbean region showing the plate boundaries with North and South America and the basalt provinces onshore and offshore. The dark pattern in the Caribbean Sea corresponds to the mapped extent of seismic horizon B" in the eastern half (Venezuelan Basin) with a presumed western extension (Hess Escarpment and southern Nicaraguan Rise). Black patches on land are obducted fragments of the basalt province. For example, Nicoya (Costa Rica), Isla Gorgona (Colombia), Dumisseau Formation (Haiti), and Curaqao. Drilled Sites 146, 150, 151, 152, and 153 of DSDP Leg 15 (solid circles) and ODP Site 999 and 1001 are also shown.
TECTONIC AND STRATIGRAPHIC DEVELOPMENT OF THE EASTERN CARIBBEAN
593
Fig. 2. Track map for EW9501 MCS data in the eastern Caribbean. Gray circles denote locations of sonobuoysalong track. purports that the majority of the deformation observed across the Beata Ridge and Venezuelan Basin occurred early in the history of the Caribbean (i.e., Late Cretaceous) prior to large sediment accumulation. Minor fault reactivation along the eastern Beata Ridge and Venezuelan Basin is inferred to have been caused by the different styles of deformation in response to the north-south compressional stress (Holcombe et al., 1990; Driscoll et al., 1995). In this paper, we use recently acquired high-resolution multichannel seismic reflection data (MCS) collected onboard the R/V Ewing from the eastern Caribbean (Fig. 2), together with onshore geologic information from South America, to constrain the timing and nature of the tectonic deformation across the Beata Ridge and Venezuelan Basin. This analysis provides critical constraints on the history of the relative motion of the Caribbean Plate and enables us to discern whether the azimuthal and deformation data can be better explained by a one- or two-plate model. The emplacement history of the Caribbean Cretaceous basalt province and its relationship to the Caribbean tectonic deformation is discussed in more detail in Diebold et al., Chapter 19. Using this data set, we also are able to define the stratigraphic evolution of the Venezuelan Basin and its relationship to the Caribbean-South American Plate
boundary. Because the high-quality seismic reflection data imaged the preserved stratal geometry of the basal stratigraphic sequences, we were able for the first time to decipher the depositional processes and sediment sources responsible for their formation. Using these new insights, we re-interpreted the older seismic reflection data, in which the basal sequences are either transparent or highly distorted, to establish a more regional framework for the depositional systems in the Venezuelan Basin (Biju-Duval et al., 1978; Ladd and Watkins, 1980; Stoffa et al., 1981; Diebold et al., 1981).
GEOLOGIC DEVELOPMENT OF THE VENEZUELAN BASIN
The Venezuelan Basin is bounded to the east and west by the Aves Rise and Beata Ridge, respectively (Figs. 1 and 2). The internal northeasterly trending grain of the Venezuelan Basin, defined by apparently linear magnetic anomalies (Donnelly, 1973), basement faulting, and a marked boundary between two distinct crustal types (smooth and rough) with little to no gravity signature (Case and Holcombe, 1980; Case et al., 1990a), is sub-parallel to the Beata Ridge. Early seismic refraction results (Officer et
594
N.W. DRISCOLL and J.B. DIEBOLD
9 ,.~ .,..a
azl .,.a .,...
O
z~
.,..a
r.tz
. ,....~
0.~
r.gz
~<
,.t::
O
r
,.t::
o
O
c~o
9
. ,....~
r.~
. ,....
0.)
,...0.,) e~o
.,.~ r~
a9
;~
~ ,...~
r . ,...~
.,..a e,.9
""
,x:
TECTONIC AND STRATIGRAPHIC DEVELOPMENT OF THE EASTERN CARIBBEAN al., 1957, 1959) showed that crustal structure in the Venezuelan Basin is unusual. It comprises two distinct layers, and is thicker than normal oceanic crust but thinner than typical continental crust. Later work (Ewing et al., 1971; Edgar et al., 1971) revealed similar, but even thicker, crustal structure in the Colombian Basin. Analysis of closely spaced airgunsonobuoy wide-angle reflection and refraction profiles suggested that the base layer in both basins was geophysically similar to 'normal' oceanic crust, but that the outer rind of the upper layer had a much smoother appearance than oceanic crust imaged in other reflection profiles (Ludwig et al., 1975; Houtz and Ludwig, 1977). This horizon was named B" and was inferred to be the top of a volcanic sequence. The source of the B" has been attributed to the Galapagos hotspot (Duncan and Hargraves, 1984). By the early 1970s, DSDP drilling in the central Caribbean had penetrated B" and sampled Cretaceous intrusives (Edgar et al., 1973a,b; Donnelly, 1973). Despite the numerous dating problems, there are several firm ages for the locally youngest age of the Cretaceous basalts and thus, the termination of magmatic activity. As reviewed in Donnelly et al. (1990) these dates are mainly derived from dating sedimentary units within and on the basalt complex, including DSDP Leg 15 sites. These ages cluster at the latest Turonian, but in a few cases could be as young as Santonian, and even early Campanian (DSDP Site 152 and ODP Leg 165). Detailed radiometric study of basalts from Gorgona, Costa Rica, Haiti, and Curacao undertaken recently using 4~ incremental heating methods, yielded dates consistent with the sedimentary ages for the termination of the event (Sinton et al., 1993). The number and geographic extent of these age determinations at 90-88 Ma suggests that a truly vast igneous event occurred at this time (i.e., the Caribbean Cretaceous basalt province). Velocity analysis and laboratory experiments suggested that the sub-B" interval observed in the multichannel seismic data is principally igneous (Ludwig et al., 1975). Estimates of sub-B" velocities are variable from place to place, a variation that may be indicative of intercalated volcanic clastic sediments (e.g., Abrams et al., 1993). Smooth B" is not ubiquitous in the Caribbean and the boundary between rough and smooth B" crust in the Venezuelan Basin was discovered by MCS profiling (Biju-Duval et al., 1978; Diebold et al., 1981; Stoffa et al., 1981). In more cases than not, the boundary is an abrupt one in this basin, accompanied by changes in basement elevation and, sometimes, apparent faulting. A highly reflective horizon observed within the sedimentary packages overlying B" was named horizon A". Examination of the seismic reflection and DSDP and ODP drill data indicates that horizon A"
599
correlates with the interface between unconsolidated Miocene and Eocene oozes and consolidated Lower Eocene cherts and chalks (Edgar et al., 1973a,b; Holcombe et al., 1990). However, the DSDP and ODP sites are located on structural highs and thus failed to sample the basal onlapping sedimentary sequence between A" and B". Furthermore, the transparent acoustic character of this basal sequence in the existing seismic reflection data across the Venezuelan Basin makes it difficult to infer the depositional processes responsible for its formation. The A" to B" sediment interval is thickest in the southeastern portion of the Venezuelan Basin and systematically decreases toward the east, north, and west. Similar basal onlapping packages also have been observed in structural lows in the Colombian Basin (Edgar et al., 1971; Lu and McMillen, 1982; Kolla et al., 1984; Bowland and Rosencrai~tz, 1988; Bowland, 1993). Similar to the A" to B" sediment interval, the stratigraphic sequence overlying A" in the Venezuelan Basin thickens dramatically toward the southeast displaying a pronounced increase in thickness across the rough-smooth B" basement boundary. The acoustic character changes markedly across horizon A" in the Venezuelan Basin: from laminated below the horizon to more chaotic and hummocky above it. This change in acoustic character is attributed to the onset of current-controlled deposition in the Caribbean (Driscoll et al., 1995). In light of the complex tectonic and stratigraphic history of the Beata Ridge and Venezuelan Basin and the possibility that more than one magmatic event was responsible for the thickened crust, we collected a series of MCS lines across and along the roughsmooth B" boundary to determine the tectonic, magmatic, and stratigraphic development of the region and to define the complex interplay between South American deformation, evolving drainage patterns, and the consequent sediment supply to the Caribbean Plate through time. During cruise 9501 of R/V Ewing in February and March, 1995, we collected over 5700 km of MCS data and deployed 104 successful sonobuoys in the Caribbean (Fig. 2). A 20-airgun, 8415 in 3 array and a 160-channel hydrophone streamer were used to acquire the MCS data.
RESULTS Acoustic b a s e m e n t
We will present the MCS data from the Venezuelan Basin and Beata Ridge from east to west and discuss the observations beginning with basement structure for each cross-section. Seismic line 1294, located along the eastern edge of the Venezuelan Basin, trends sub-parallel to the
600 Aves Ridge (Fig. 3). The rough-mooth B" basement boundary, which also delineates the abrupt increase in sediment thickness below reflector A", occurs near CDP 9000 (Fig. 3). Continuing farther west, line 1293 again illustrates the abrupt boundary between the rough and smooth B" acoustic basement (Fig. 4). The rough-smooth B" boundary occurs at approximately CDP 19,000 and is associated with a reversal of dip near the boundary and a sharp shoaling of Moho (see fig. 2, Diebold et al., Chapter 19). The change in dip is local (~10 km, Fig. 4). Brittle deformation and fault rotation near the roughsmooth B" basement boundary is the most likely cause for this reversal of dip. The fault system that controls the geometry and structure of the dip reversal must be relatively surficial and antithetic to a larger normal fault system that dips north-northwest and controls the location and geometry of the large divergent wedge observed beneath B" (CDPs 10,500-19,000) and the large step in Moho topography (CDP ~ 19,000). The divergent wedge beneath the smooth B" basement has interval velocities on the order of 5 km/s (Fig. 4; Diebold et al., Chapter 19). Similar reflector geometries and velocities on numerous other rifted margins have been attributed to large igneous events at, or immediately prior to, the cessation of continental rifting and the onset of seafloor spreading and are termed seaward-dipping reflectors (SDRs; Hinz, 1981; Mutter et al., 1982; Schlich et al., 1993). The divergent geometries and spatial distribution of the reflector packages across the Venezuelan Basin suggest multiple igneous events. The rotation and divergence of the onlapping seismic reflectors diminishes upsection and is indicative of differential subsidence and block rotation (Fig. 4). The spatial extent of each successive divergent reflector package increases within the wedge. The rough-smooth B" basement boundary is also delineated by a local reversal in dip on line 1300 (Fig. 6). Horizon B" is offset by a number of small graben structures centered around CDP 1500 and 9500 (Fig. 6). Graben formation appears to have occurred soon after the emplacement of the basalt flows and prior to the onset of significant pelagic sedimentation. Although the overlying pelagic deposits mirror the underlying topography, there is no evidence of growth faulting observed with these structures. Similar to line 1293, the late-stage divergent reflectors are located in close proximity to the rough-smooth B" basement boundary, the local reversal of dip (i.e., fault escarpment), and the shoaling of Moho topography. The rough-smooth B" basement boundary varies dramatically along strike from a well defined escarpment to a more subdued gradual boundary (see Diebold et al., Chapter 19). The rough-smooth
N.W. DRISCOLL and J.B. DIEBOLD boundary is not marked by an escarpment on line 1302. In this region, the smooth B" basement surface is made up of several en-echelon reflectors that commonly downlap onto the underlying surface, suggesting multiple flows (Diebold et al., Chapter 19). Along composite line 1304/1317 the roughsmooth B" boundary occurs across a pronounced basement high (Fig. 7), a high which appears to be a volcanic edifice similar to those observed along line 1293 (Fig. 4). This high, however, is not associated with underlying divergent reflectors. Graben structures that offset smooth B" basement also are observed in this region. Sub-B" divergent reflectors beneath the structural high near DSDP Site 150 indicate that the basalts recovered at Site 150 are only representative of the late-stage evolution of the Cretaceous basalt province. Correlation of seismic reflection line 1320 to DSDP Sites 146 and 150 suggests that B" correlates with Late Cretaceous dolerites (Fig. 9; Edgar et al., 1973a,b). The dolerite recovered at DSDP Site 146 is a sill within Turonian limestone. Beneath the Turonian limestone another 7 m of dolerite was penetrated with fairly high recovery (Edgar et al., 1973a,b). At DSDP Site 150, diabase also was recovered at the base of the hole, with approximately 4 m of penetration and again good recovery. Rough B" basement has yet to be sampled and its relationship to smooth B" basement remains conjectural (Diebold et al., 1981). Analyses of basalts recovered at ODP Site 1001 on the Hess Escarpment indicate that the Caribbean oceanic plateau was active until at least 77-76 Ma (mid-Campanian). Benthic foraminiferal assemblages together with the vesicular nature of the basalts suggests that emplacement on the Hess Escarpment occurred at shallow water depths followed by rapid subsidence (Sigurdsson et al., 1997). The mid-Campanian volcanism is consistent with the Duncan and Hargraves (1984) model that volcanism associated with the Galapagos hotspot continued from 100 to 75 m.y.B.P, in the Caribbean. Because Site 1001 failed to recover Senonian and older basalts (88 Ma), the results from the site cannot be used to determine whether or not rapid outpouring across the Caribbean Plateau occurred during Senonian time (90-88 Ma). Farther west, seismic line 1321 that crosses the western Venezuelan Basin and the Beata Ridge illustrates the highly faulted nature of the region (Fig. 10). The western flank of the Beata Ridge is bounded by a large westward-dipping normal fault. The basement offset across the normal fault is approximately 5 s two-way travel time (TWT; ~3750 m). The Beata Ridge is markedly asymmetric with the western scarp having more offset and a steeper slope. The uplift, rotation and wavelength of the
TECTONIC AND STRATIGRAPHIC DEVELOPMENT OF THE EASTERN CARIBBEAN
601
6000 -
7000
'=
-
8000 -
6 9000 -
10000"
I
I
i
I
i
1293
5000
-
nAo 'ooo ~300 6000
1700 17(~ 1700 +Mo I ~ o Ioo0
17001700 leSO 1o8o
1700 1700 1700 IOOO leOO 1o8o
1700 18~o
2000 2000
2000
20002000
2000 2000 2000
2000
_2___
5 ~.,,. 0
24O0
-
300
~
9
I~9 2000
~
3
17oo, ,, 1880
IToo 1880
2000
2000
,
A ll 3
0
0
0
33003300
3300
3300
311100
I~oo I_~) 2000
Ik~OO
S300
7000
-
8000
-
S300
63O0
9000
-
10000 -16000
!
~
!
21000
I
!
213000 GDP
#
Fig. 5. Velocity model derived from sonobuoy data collected during EW9501 used to process MCS line 1293. The lateral velocity gradient observed within the B" to A" interval coincides with the region where the basal terrigenous sediments pinch out by onlap against the escarpment.
Beata Ridge topography suggest that extensional unloading and the consequent flexural rebound of the lithosphere play an important role in its formation. Along the eastern flank of the Beata Ridge are a series of half-graben structures and the basal reflectors in the graben dip and thicken toward the west. Note the small seamounts/mounds on top of B" toward the southeastern portion of the line. The normal fault located at CDP ~ 12,000 is spatially coincident with a small mound suggesting that these volcanic constructions might be zones of weakness that focus the extensional deformation (Fig. 10). The onlap of B" to A" sequence onto the rotated and deformed hangingwall blocks indicates that deformation in this region occurred soon after magmatic emplacement and prior to substantial sediment accumulation. Together with the basal onlap, the differential thickness of the B" to A" sediment interval across the faults attests to the existence of the fault at the time of deposition (Fig. 10). In regions where the subbottom reflectors are not distorted by the rough seafloor, sub-B" reflectors are well imaged along the eastern flanks of the Beata Ridge (e.g., CDP 21,000). West of the Beata Ridge are thick ponded turbidites with minor growth faults offsetting the lower stratigraphic
sections and their occurrence diminishes upsection (Fig. 10). Extension within what is now the Beata Ridge and along its western flank is focused in the north, and broadens to the south. Neglecting the possible effects of N-S compression along the northern edge of the plate (Muertos Trough, southern Hispaniola) the distribution and number of faults as well as the heave across any one given fault explains the architecture of Beata Ridge: high and narrow in the north, wide and low in the south (Fig. 2). In fact, toward the south, the eastern flank of the ridge is more dissected by N-S-trending grabens. Crust at the foot of the Beata Ridge escarpment, in the Haiti Basin, appears to have been thinned by normal faulting and block tilting (Fig. 10), a thinning which has reached its maximum immediately south of Haiti (Fig. 2).
B" to A" sedimentary sequence The sedimentary succession overlying B" is separated by two prominent reflectors, horizon A" and eM. Where sampled, horizon A" correlates with the interface between unconsolidated Miocene
602 to Eocene oozes and consolidated Lower Eocene cherts and chalks and horizon eM correlates with the boundary between Early Miocene radiolarian ooze and Early to Middle Miocene calcareous ooze (Edgar et al., 1973a,b; Holcombe et al., 1990; Sigurdsson et al., 1997). A marked change in acoustic character also occurs across horizon A", from well laminated below the horizon to more lenticular and hummocky immediately above it. The hummocky character persists throughout the overlying sequence (A"-eM), culminating just above horizon eM. Seismic line 1294, located along the eastern edge of the Venezuelan Basin, trends sub-parallel to the Aves Ridge (Fig. 3). The abrupt increase in sediment thickness below reflector A", near CDP 9000, is spatially coincident with the rough-mooth B" basement boundary (Fig. 3). Farther west, the thickness of the B" to A" sediment interval systematically increases toward the south along line 1293 with a marked increase also occurring across the rough-smooth boundary (Fig. 4). Fig. 5 shows the smoothed interval velocity field derived from sonobuoy data collected during EW9501 and was used in conjunction with stacking velocities to process MCS line 1293. The lateral velocity gradient within the B" to A" interval coincides with the region where the basal terrigenous sediments pinch out by onlap against the escarpment separating smooth and rough B" basement. On the basis of velocity information, the maximum thickness of this sequence is approximately 600 to 800 m and its distribution is roughly coincident with the area of rough basement in the Venezuelan Basin. This lateral velocity gradient displayed in Fig. 5 was also recognized by Talwani et al. (1977). The velocity increase from 2.4 km/s to 3.2 km/s that accompanies the lateral change in thickness also implies the existence of a facies boundary between pelagic and terrigenous sediments (Fig. 5). The higher velocities were not consistent with the velocities determined for A" by DSDP Leg 15 drilling (Edgar et al., 1973a,b), which forced Talwani et al. (1977) to identify a new reflector, 'horizon X'. Despite the velocity difference, they did recognize that horizon X might actually be chronostratigraphically equivalent to horizon A". Nevertheless, at this early stage of interpreting seismic reflection data, they placed more weight on correlating surfaces with similar velocities than trying to determine why certain seismic units exhibited lateral velocity variations (Talwani et al., 1977). Compaction, small readjustments caused by sediment loading, and minor fault reactivation have caused horizon A" to be offset across the roughsmooth B" basement boundary in seismic line 1300 (Fig. 6). A number of small graben structures are observed across the region, two of which are centered around CDP 1500 and 9500 (Fig. 6). The graben
N.W. DRISCOLL and J.B. DIEBOLD formation appears to have been complete soon after the emplacement of the basalt flows and prior to the onset of significant pelagic sedimentation. Although, the overlying pelagic deposits mirror the underlying topography, there is no evidence of growth faulting observed with these structures. The stratigraphy overlying B" in line 1300 exhibits the same evolution from a well laminated interval between horizons B" and A" to a more acoustically hummocky character above A". The B" to A" sediment thickness diminishes toward the west (Figs. 4, 6 and 7). In addition, along the southward-dipping slope separating smooth and rough B" basement, the B" to A" sediment interval pinches out completely near CDP 9500 on line 1304/1317 with thicknesses increasing away from this region to both the south and north. The unit appears to thin predominantly by downlap and onlap rather than by truncation suggesting that the hiatus is recording non-deposition rather than erosion (Fig. 7). Farther west on line 1320 (Fig. 2), DSDP Site 150 is located at CDP 4900 in an area where the B" to A" sediment interval thins updip across the structural high by onlap, again suggesting that the hiatus in the B" to A" sequence is recording non-deposition rather than erosion (Fig. 8). At DSDP Site 146, radiolarian sandy turbidites occur in the Campanian and Maastrichtian and appear to be locally derived. Late Paleocene to Early Eocene zeolitic brown clays and marls recovered at DSDP Site 150 show characteristics of terrigenous mineralogy. For example, the mica, kaolin-chlorite and minor corroded plagioclase could be derived either from South America or from an increase in ash input during this time (Edgar et al., 1973a,b). Minor amounts of Paleogene clays were also recovered from ODP Site 1001 located along the eastern flank of the Hess Escarpment, which were interpreted to be of terrigenous origin on the basis of chromium (Cr) and titanium (Ti) (Sigurdsson et al., 1997). Furthermore, Cr and Ti analysis of the Site 1001 cores suggests that there is a quantifiable change in sediment provenance from the Paleocene-Eocene and the Miocene and Pleistocene (Sigurdsson et al., 1997).
A" to eM sedimentary sequence The hummocky acoustic character above A", which has been recognized by previous researchers and mapped across the Venezuelan Basin (Talwani et al., 1977), is a time transgressive boundary that progressively shoals through time from south to north. For example in line 1294, the boundary between well laminated horizons and the hummocky acoustic character occurs at approximately 6.8 s two-way travel time at CDP 2000 in the south, is
Fig. 6. Interpreted seismic reflection line 1300 showing that the abrupt shoaling of Moho is spatially coincident with the rough-smooth boundary. Northwest is loward the left. See Diebold et al., Chapter 19 for the uninterpreted section.
Fig. 7. Interpreted composite seismic reflection lines 1304 and 1317 illustrating the complex nature of the rough-smooth boundary in this region. Note that the B" to A ~' sedimentary sequence is thinning towards the west away from lines 1293 and 1300. Northwest is toward the left.
Fig. 8. Interpreted seismic reflection line 1320 illustrating the style of sediment thinning (truncation versus onlap-downlap) from Site 146 to Site 150. Note the sub-B" dipping reflectors near Site 150. North is toward the left.
611
TECTONIC AND STRATIGRAPHIC DEVELOPMENT OF THE EASTERN CARIBBEAN
EW9501 Line 1320 CDP # Site 146/149 9_ - - _ - _ . . - -
..--.-
._-_--_---
- - - -
.--_.-._--
- , - - ~
._.- -.-_--
- - - -
. --_-_--..-
. , - - -
. -_-_--
- . - - -
, -
--
' , -
-
- . - - . -
--
-,.-
..-
-
-...-
-..-.-.-
. . . . . . . . . . . . . . . . . .
I
8000 I
7000 I
6000 I
truncation
Site 150
-_-
1- " l r '1- . r 9r T T ' r /-.I-.I-.I- r " r - r "Ir
. . . . 91- '1- '1- ~r 9 9 9 9 I''1-'I'-i- : - - : - ' : " ~ '1" '1- " r 'T LA..A...A. 1- ' I " "1" '1" . . . . '1" '1" " r ' ~ I~K_A..A, "r "r "r T . . . . " r " r - r ~r 9
.
I
.-
----
-
9000
I
.
,
~
I...,_.A_~
-6
'1-'l-'1r'~r 1--r -r "r 9' r " r " r "It
A
1" "It" " r - r - r - r " r ~r ' r "1" " r "t" 91" " r '1- ~r
~..A..A_.~
91- - r - r - r 1" " r "1" T -r "r "It -r 1- " r T " r "r "r -r "f
,_.A_,_.~ L_,~.I_.~
A
Ak
,&
,A ,~
A
A -1r
A
^ ^ ^ ^ , , ^ ,
&k
-7
-8
~
limestone
r ~ oozeSilice~
~
clay
chalk
~
D
basalt
calcareous ooze
Fig. 9. Correlation of seismic reflection line 1320 to DSDP Sites 146 and 150 indicates that Btt correlates with the top of Late Cretaceous dolerites. Horizon Att correlates with the interface between unconsolidated Miocene to Eocene oozes and consolidated Lower Eocene cherts and chalks and horizon eM correlates with the boundary between Early Miocene radiolarian ooze below and Early to Middle Miocene calcareous ooze above. The change from siliceous to carbonate ooze, together with the diminished effects of currents on the depositional regime after the Early Miocene, are consistent with the shoaling of the Panamanian Isthmus.
exposed at the seafloor near CDP 9200, and becomes quite dramatic by CDP 11,500 in the north (Fig. 3). Consequently, this horizon has no time stratigraphic significance and is, in fact, only recording the complex interplay between gravity flows and geostrophic bottom currents. This change in acoustic character from laminated below to acoustically hummocky above is also observed in the Colombian Basin (Lu and McMillen, 1982). Similar observations have been made in other ocean basins where gravity flows interact with abyssal circulation (Driscoll and Laine, 1996). In the region where the hummocky acoustic character is observed on the seafloor in the MCS data, the 3.5 kHz precision depth records indicate that this acoustic character correlates with sediment waves having amplitudes of tens to hundreds of meters and wavelengths on the order of a few kilometers (Fig. 3). As in line 1294 toward the east, the acoustically hummocky chaotic sequence is observed above horizon eM in line 1293 (Fig. 4). The sequence above horizon A" is systematically less well laminated toward the west (from lines 1293
to 1304) and exhibits a more hummocky acoustic character (Fig. 6). The A" to eM sediment interval in line 1300 continuously thickens toward the southeast, in contrast to lines 1293 and 1294, where this interval shows signs of thinning both towards the northwest and southeast. The eastern Caribbean paleo-structure during Eocene time appears to have controlled, to some degree, the thickness variations of the A" to eM interval. For example, the thinning towards the southeast Venezuelan Basin observed in lines 1293, 1294, and 1296 might reflect basement relief at the time of deposition (Fig. 2). Along line 1297 (Fig. 2), which trends northeast-southwest, both the B" to A" and the A" to eM sediment intervals thicken markedly away from the basement high (i.e., western edge of the Aves Rise). As previously mentioned, where the hummocky acoustic character outcrops on the seafloor (i.e., north of CDP 8500), 3.5 kHz precision depth recordings indicate that the acoustic character is the result of sediment waves (Damuth, 1980). In fact, individual sediment waves can be identified in the MCS data (Fig. 6).
612
N.W. D R I S C O L L and J.B. D I E B O L D EW9501 Line 1321 - Beata Ridge - CDP # 25000
20000
~!~:~~
15000
10000
5000
_d,.,~,~
f
=.))~.'~r
+~:7.'..":r;{c'.: ::
:!~ii)~i;:1!i:7{ ,g;!i~:7: ~
;1
"G3~7~-L'I'.~<-"I=~L,:(!
:i~
,7 :! ~;.D.'..f<7.:'.:+%.i:,.)'.'..:"~7,~,:&+.,'~!:.<.,7,1 :.ji~,iii:-~i;~i:~)~'!~.~.,:,~ (";.7L.,.::~E.;~J;iLI;(::~,:( ;.:,7.~.?~'1::.
,]r .:,~17".7'":. :. :L,~,>~ii~:~'..<;s;7~."f,"g;~iq-_'eL 3~1~(.~.s163
i:~i~:
:~:
EW9501 Line 1321 - Beata Ridge - CDP # 25000
20000
15000
10000
5000
sE thickness
I 5
! I
~
~
~
-
onlapoing
,
~ i ~ ; ~ ~~. _
;.... ~]..~ 'Q ,~:~ - !
.- ~ :, __.
package
~
~'~-
'5
-
I
6
.a=
>,
I ~
'~ ,
ii~ '~:'~' .... !'I~4!~:~'
'~
~"
"..~:.-.~
. . . . . . . . . .
~
"
....
, ' ~ ' ~
.7
8
"8
9
"9
lO
"10
11
"11
I
Fig. 10. Interpreted and uninterpreted seismic reflection line 1321 illustrating the highly faulted nature of the Beata Ridge. The western flank of the Beata Ridge is bounded by a large westward-dipping normal fault. The basement offset across the normal fault is approximately 5 s two-way travel time (~3750 m). The Beata Ridge is markedly asymmetric with the western scarp having more offset and a steeper slope. The sediment thickness and stratal geometry observed across the Venezuelan Basin and Beata Ridge suggest that the majority of the deformation in this region occurred soon after the emplacement of the volcanics.
The A" to eM sediment interval displays a similar pattern of thinning as the underlying B" to A" sequence in the region. Nevertheless, in the area of maximum thinning of the B" to A" sequence, this overlying younger unit maintains a finite thickness (Fig. 7). Line 1320 traverses one of the few areas in the study region where the A" to eM sequence shows signs of truncation (Figs. 8 and 9). DSDP Site 150 is located at CDP 4900 in an area where the maximum hiatus occurs due to erosion of the A" to eM sequence and non-deposition within the B" to A" unit. Farther to the north away from this region of erosion-
nondeposition, DSDP Site 146 (CDP 10,200) sampled a more complete sedimentary section. In fact, the erosional and non-depositional unconformities have coalesced between CDPs 4200 and 6000, maximizing the duration of the hiatus in the region. The updip contemporaneous pelagic units also display lateral thickness variations across the volcanic constructions located at CDP 12,500 and 14,000. At DSDP Site 146, iron manganese nodules (up to 5 cm in diameter) occur at the Early Miocene-Late Eocene unconformity at Site 150 suggesting slow accumulation rates (i.e., current-controlled deposition;
TECTONIC AND STRATIGRAPHIC DEVELOPMENT OF THE EASTERN CARIBBEAN Driscoll and Laine, 1996). Because the paleo-water depths between DSDP Sites 146 and 150 are not demonstrably different and some of the thinning in the A" to eM sequence occurs by truncation, Edgar et al. (1973a,b) suggested that dissolution associated with the CCD could not account for the increased hiatus toward DSDP Site 150 (Fig. 9). They postulated that abyssal currents might explain the hiatus and the occurrence of manganese nodules. Similar to the hiatus observed at DSDP Site 150, there is a large hiatus ('-~30 Ma) that separates Early-Middle Eocene chalk from Middle Miocene calcareous ooze along the Hess Escarpment at ODP Site 1001 (Sigurdsson et al., 1997).
613
rough-smooth B" boundary this thinning is much less pronounced than the thinning of the two underlying sequences (e.g., B" to A", and A" to eM). The change from siliceous to carbonate ooze across horizon eM, together with the diminished effects of currents on the depositional regime after the Early Miocene, are consistent with the shoaling of the Panamanian Isthmus in the Early Miocene which would have closed the gateways for the deep currents.
DISCUSSION
Tectonic development of the Caribbean eM to present sedimentary sequence All of the sedimentary wedges observed in the southern portion of the Venezuelan Basin thin toward the north and interfinger with pelagic deposits on the structural highs that were accumulating sediment concomitantly. However, this lateral thinning updip is less pronounced in the upper wedge (Early Miocene-Holocene; Fig. 4). These pelagic sediments uniformly drape the existing topography with thickness variations only occurring across steep topographic features (e.g., seamounts and fault scarps). The height where the pelagic sediment pinches out against seamounts varies dramatically from one side to the other (e.g., line 1293, CDP 2200; Fig. 4). In this example, a marked change in seafloor character accompanies this variation in sediment elevation on either side of the seamount with sediment waves located toward the southeast and smoother seafloor occurring toward the northwest (Fig. 4). Because of the different sedimentary styles, the seamount appears asymmetric in the 1 from the area, with a large slope to the north and small slope to the south. South of the seamount, the seafloor rises gently, and is characterized by sediment waves with amplitudes of tens to hundreds of meters and wavelengths of about 1-2 km. The acoustic character above horizon eM systematically becomes more laminated toward the west (from lines 1293 to 1304). Above horizon eM, thinning along the northern slope still occurs, but the lateral thickness variations are much less pronounced. The more acoustically laminated unit immediately overlying horizon eM thins both toward the south and north away from the depocenter (Fig. 6, CDP 14,500) and is onlapped by the overlying sedimentary package. The upper sequence, the flat-lying well-laminated package, thickens markedly towards the south and also systematically thickens towards the west (Fig. 7). Even though the eM to Present stratigraphic interval thins across the escarpment delineating the
On the basis of the new MCS data acquired in the eastern Caribbean (Fig. 2), together with previous collected data from the region, we have developed the following tectonic model to illustrate the geologic development of the crust that floors the Venezuelan Basin (Fig. 11). Details of the magmatic emplacement and building of the Caribbean oceanic plateau can be found in Diebold et al., Chapter 19. Our tectonic model supports the 'mobilist' model for the development of the Caribbean (Pindell and Barrett, 1990). We assume that the proto-Caribbean crust was formed by seafloor spreading in Late Jurassic-Early Cretaceous time in the eastern Pacific (Fig. 11), with an initial oceanic crust thickness of approximately 6 km (Diebold et al., Chapter 19). Prior to the Senonian, widespread and rapid eruption of basaltic flows began in concert with extension and thinning of the 'old' plate. It is difficult to determine whether extensional deformation pre-dated or post-dated the impact of the plume head and associated magmatic activity (Galapagos hotspot; Duncan and Hargraves, 1984). Nevertheless, extensional deformation had minimal effect on the thickness and distribution of the early-stage basalt flows (Fig. 11). Numerous hypotheses concerning the interaction of plume heads with both active and passive rifting of the overlying lithosphere have been advanced (Anderson et al., 1992; Coffin and Eldholm, 1994; White and McKenzie, 1995) and are discussed in more detail in Diebold et al., Chapter 19). The observed structure of the Caribbean basalt province indicates at least two episodes of magmatic activity (Diebold et al., Chapter 19), which is consistent with the age information reported by Donnelly (1990). The early stage was more localized with steeper-dipping reflectors; the late stage is regionally more extensive with its eastern boundary being faultcontrolled. The rough-smooth B" basement boundary imaged in seismic profiles is, in this model, the edge of the basalt province overlying the older plate (Fig. 11). The divergent wedge imaged below B" in
614
N.W. DRISCOLL and J.B. DIEBOLD
Fig. 11. Schematic illustrating the geologic development of the Caribbean crust. (1) Proto-Caribbean oceanic crust formed by seafloor spreading in Late Jurassic-Early Cretaceous time in the eastern Pacific. (2) Widespread and rapid eruption of basaltic flows in concert with extension and thinning of the 'old' plate. The plate was thickened by at least two stages of basalt flows. The large divergent volcanic wedge observed along the rough-smooth B't boundary, are coincident with the abrupt shoaling of Moho, and appear to be bounded by a large northwest-dipping fault system. (3) Minor extensional deformation across the Venezuelan Basin continued after magmatic thickening of the crust as evidenced by faulted and rotated basalt flows. 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).
line 1293 (Fig. 4) and the abrupt shoaling of Moho are inferred to be controlled by a northwestwarddipping fault system. This fault system was active during the late stage of basalt flows. The high velocities (,-~5 k m / s ; Diebold et al., Chapter 19) of this divergent wedge suggest basaltic material with only minor intercalated volcaniclastics. The dips of the reflectors that comprise the divergent wedge diminish upsection and this evolution is very similar to the
stratal geometry observed across evolving rift basins (Driscoll et al., 1995). If the emplacements of the volcanics are similar to those of syn-rift sediments, then this implies that the source of the basalts is either along strike or toward the northwest, that is structurally updip, and the volcanics progressively infill the deforming basin (Fig. 11). The dipping reflector packages observed in the Venezuelan Basin exhibit many structural similari-
TECTONIC AND STRATIGRAPHIC DEVELOPMENT OF THE EASTERN CARIBBEAN ties to seaward-dipping reflectors (SDRs) observed along other volcanic margins (Mutter et al., 1982; Coffin and Eldholm, 1994). The commonly cited model for emplacement of the SDRs is that thermal dynamic support uplifts the magmatic source above sea level with extrusions flowing away from the source region (Mutter et al., 1982). When the anomalous heat dissipates the plate cools, contracts, and subsides, the primary dip of the volcanics is reversed and the SDRs dip toward the extinct source. In the case of the eastern Caribbean, the fault-controlled eastern extent of the late-stage flows together with the intercalated sills and marine limestones indicate that the divergent wedge was formed in the marine environment. Details concerning the source of the basalts flows and their inferred low viscosity is discussed in Diebold et al., Chapter 19. Minor extension continued after magmatic thickening of the Venezuelan Basin crust as evidenced by faulted and rotated tongues of smooth basement (Fig. 11). However, the location of the major extensional deformation migrated through time from the western portion of the Venezuelan Basin to the western flank of the Beata Ridge. Strain hardening of the lithosphere in the Venezuelan Basin might have caused the deformation to migrate from an area with thin crust to a region with thicker crust. The deformation across the Beata Ridge is the consequence of extension across a westward-dipping fault, with the Hess Escarpment being the collapsed hangingwall block. Antithetic faults along the eastern Hess Escarpment result in a graben-like structure between the Beata Ridge (footwall block) and the Hess Escarpment (hangingwall block). Furthermore, the regional uplift of the Beata Ridge caused minor faulting across the plateau. The onlap of the B" to A" sediment interval onto the rotated and deformed blocks suggests that deformation occurred soon after magmatic emplacement. Furthermore, the currentcontrolled depositional features, which are enhanced around structural highs, suggests that the majority of the deformation responsible for the formation of Beata Ridge was completed prior to any substantial sediment accumulation. This timing of the deformation is also consistent with the stratal patterns of the turbidites that infill the graben between Beata Ridge and the Hess Escarpment and the basal turbidites that infill the Venezuelan Basin. Decompression melting as a result of the extensional deformation may have led to partial melting of the depleted residual mantle, generating some new crust formation in Campanian time (DSDP Site 152 and ODP Site 1001 basalts). The stratal geometry and basement structure are not consistent with the two-plate hypothesis used to explain the origin of the observed deformation across the Beata Ridge and Venezuelan Basin. Mauffret and Leroy (1997) proposed that a renewed phase
615
of deformation occurred in the Miocene and that deformation is still ongoing. They propose that this deformation is associated with the overthrusting of the Colombian Plate onto the Venezuelan Plate along the eastern flank of the Beata Ridge (Fig. 10, CDP 13,000). In our alternative hypothesis, the one-plate model, the majority of the deformation occurred early on in the history of the Caribbean with minor reactivation being caused by the north-south bending of the plate. The onlap above the rotated block, which is the inferred thrust fault of Mauffret and Leroy (1997), indicates that the deformation occurred prior to sedimentation in this region. In summary, much of the topography associated with the formation of the Venezuelan Basin and Beata Ridge pre-dated substantial sediment input into the region (Diebold et al., 1995a,b). We propose that the minor fault reactivation in the Neogene along the eastern flank of the Beata Ridge is associated with the shortening between the North and South American Plates which began in the Eocene (Pindell and Barrett, 1990). The deformation style changes along strike from the Muertos Trough to the east and the Hispaniola margin toward the west. A northsouth-trending transfer zone (i.e., tear fault) accommodates this change in the deformation style from bending and subduction of the Caribbean Plate along the Muertos Trough to the east and compressional deformation and obduction of the Caribbean Plate along Hispaniola to the west (Fig. 2). Onshore Hispaniola at the intersection of the island with the Beata Ridge, NNE-trending normal faults occur along the eastern margin of the Sierra Bahoruco (Pindell et al., 1988). Toward the west of the normal faults, the Beata Ridge becomes subaerial (i.e., Sierra Bahoruco) as a result of the collisional deformation with the Central Cordillera arc of Hispaniola (Biju-Duval et al., 1982; Heubeck and Mann, 1991). Our hypothesis explains the north-south trend of the faults and predicts that fault displacement should increase northward in concert with the relief between the downgoing and obducted slab. We propose that this difference in deformational style along this portion of the Caribbean Plate is, in part, a consequence of the thicker crust on the Beata Ridge, which is more resistant to subduction.
Late Cretaceous to Eocene turbidites deposition Coincident with the rough-smooth B" boundary is a marked change in the thickness of the A" to B" sediment interval (Figs. 3, 4 and 6). Correlation from existing DSDP Sites 146 and 150 in the Caribbean constrains the age of this interval to be older than Middle Eocene (~50 Ma) and younger than Senonian (~88 Ma; Fig. 9). However, because the DSDP and ODP sites are located on structural
616
N.W. D R I S C O L L and J.B. D I E B O L D
Fig. 12. Plate reconstructions of the Caribbean during the Late Cretaceous and early Tertiary illustrating the sediment delivery from the axial-parallel fluvial systems draining the Maracaibo-Peruvian foreland basin to the Venezuelan Basin (modified from Pindell and Barrett, 1990 and Pindell and Tabbutt, 1995). (Top) Uplift and deformation of the Andes that began in the Late Cretaceous led to the formation of the Maracaibo-Peruvian foreland. The northern margin of South America was relatively undeformed at this time allowing the axial-parallel fluvial system access to the Venezuelan Basin. Obduction of the arc and related terranes occurred along the eastward verging Romeral suture-thrust in the Late Cretaceous. (Bottom) Progressive infilling of the Maracaibo-Peruvian foreland basin occurred from south to north from Late Cretaceous to early Tertiary time.
TECTONIC AND STRATIGRAPHIC DEVELOPMENT OF THE EASTERN CARIBBEAN highs, they have failed to sample the basal onlapping sedimentary sequence (Figs. 8 and 9). The A" to B" sediment interval is thickest in the southeastern portion of the Venezuelan Basin and systematically decreases toward the east, north, and west. Similar basal onlapping packages also have been observed in structural lows in the Colombian Basin (Edgar et al., 1971; Lu and McMillen, 1982; Kolla et al., 1984; Bowland and Rosencrantz, 1988; Bowland, 1993). However, their regional extent is much more limited than the onlapping packages infilling the Venezuelan Basin, suggesting that they either have a local source or represent the distal reaches of larger more continuous proximal deposits. The onlapping packages observed in the Colombian Basin may be derived from local fiver networks draining the western flanks of the Central Cordillera (Fig. 12). The thinning of the sequence in the Venezuelan Basin occurs by onlap onto pre-existing structural highs. Their welllaminated acoustic character together with the onlap pattern suggest that these deposits are the result of gravity flows infilling topographic lows (Driscoll et al., 1995). On the basis of velocity information (Fig. 5), the maximum thickness of the B" to A" sequence is approximately 600 to 800 m and its distribution is roughly coincident with the area of rough basement in the Venezuelan Basin (Diebold et al., 1981). We interpret the velocity increase from 2.4 km/s to 3.2 km/s that accompanies the change in thickness as evidence for a facies boundary between pelagic sediments to the north and terrigenous sediments to the south (more silt-sand prone). The thickness and distribution of this sedimentary succession suggests that it represents the distal portion of a terrigenous depositional system. The proximal portions of the deposit to the south may underlie the Curaqao Ridge (Fig. 15; Edgar et al., 1971) and infill a marginal basin, which may explain the large negative gravity anomaly (Bowin, 1976) observed in that region. If our interpretation is correct and these flatlying, acoustically laminated sediments are the distal portion of a large terrigenous system, then these sediments will provide valuable information about the evolution of the Caribbean-South American Plate boundary during the Late Cretaceous and early Tertiary. Based on paleo-reconstructions of the South American Plate, together with the spatial distribution of these deposits, we propose that the source for these terrigenous sediments is a northward-flowing axial-parallel drainage system developed east of the uplifting and deforming Andean Cordillera during the Late Cretaceous and early Tertiary (Fig. 11). We will briefly discuss the tectonic and stratigraphic development of South America, the establishment of drainage systems, and their relationship to the onlapping packages in the Venezuelan Basin. For a
617
more in-depth discussion on the tectonic history of South America the reader is referred to Tankard et al. (1995). Several phases of tectonic uplift and deformation have been documented for the Andes between the Late Cretaceous and Pleistocene (Hoorn et al., 1995). The Maracaibo-Peruvian Trough, which as the name implies extends from southern Peru to the Gulf of Maracaibo, Venezuela, was first isolated from the open ocean to the west by a volcanic barrier in the Central Cordillera of Colombia in the Late Triassic-Early Jurassic (Browning and Walper, 1982; Macelli, 1988; Pindell and Tabbutt, 1995). During the Late Jurassic and Early Cretaceous several marine incursions of the MaracaiboPeruvian Trough occurred from the north and penetrated southward to at least southern Peru (Browning and Walper, 1982). In Colombia, metamorphism and uplift of the Amaime-Chaucha Terrane occurred by an eastward-vergent thrust system (i.e., Romeral Suture) onto the Central Cordillera, which was an active margin during Campanian time (Bellizzia and Dengo, 1990; Case et al., 1990a,b; Pindell and Barrett, 1990; Kellogg and Vega, 1995; Pindell and Tabbutt, 1995). Fission-track data from the region, which are a proxy for denudation, indicate that uplift and erosion began in Late Cretaceous-Paleocene time, but accelerated in the late Tertiary (Kroonenberg et al., 1990). Escalante (1990) proposes compression across the Romeral fault zone during the Late Cretaceous, resulting in the obduction and emergence of Pacific Ocean crust (i.e., the Choc6 Block) against the South American Craton. The tectonic loading associated with the thrusting caused subsidence in the foreland basin to the east of the uplift. The tectonic subsidence was augmented by the contemporaneous long-term eustatic sea level rise (Fig. 12). During the Late Cretaceous through early Tertiary, the Maracaibo-Peruvian foreland basin was progressively infilled with fluvially dominated deposits from south to north (Fig. 12 Browning and Walper, 1982; Macelli, 1988; Pindell and Tabbutt, 1995). The sediments infilling the basin were derived from the Guyana and Central Brazilian shields to the east and the incipient Andes to the west (Fig. 14; Browning and Walper, 1982). The westerly flowing fluvial systems draining the Guyana and Central Brazilian shields exhibited a radial pattern away from the structural highs with no preferred orientation to the drainage networks. The drainage from the Andes and the shields were tributaries to a large-scale northerly flowing drainage system, which was axial parallel to the Maracaibo-Peruvian foreland basin. In actively deforming regions, be they extensional or compressional, the dominant drainage networks are usually axial parallel to the tectonically induced topography (Driscoll et al., 1995). The
618 northern margin of South America was relatively undeformed during this period and thus allowed the axial-parallel fluvial networks access to the Venezuelan Basin (Figs. 12G and 14B). The high-standing Guyana and Central Brazilian shields, as well as the structural barriers (e.g., De Purus and Iquitos arches) across the Marfijo and Solim6es basins, inhibited drainage to the east (i.e., towards the Atlantic Ocean) away from the Maracaibo-Peruvian foreland basin (Browning and Walper, 1982; Hoorn et al., 1995; Pindell and Tabbutt, 1995). Likewise, the Andean Central Cordillera prevented drainage from the Maracaibo-Peruvian foreland basin toward the west into the Pacific Ocean (Fig. 12). Through time, the fluvial networks became graded and the established equilibrium profiles change in concert with base level variations (Schumm, 1991). This equilibrium profile was perturbed by relative sea level changes and tectonic uplift of the hinterland. During the Late Cretaceous and early Tertiary, the long-term eustatic sea level was falling, which further entrenched the axial-parallel drainage system in the MaracaiboPeruvian foreland basin (Fig. 12). The transition from shallow marine to predominantly continental conditions in the Maracaibo-Peruvian foreland basin throughout the Late Cretaceous and early Tertiary records some combination of the following: (1) a decrease in tectonic subsidence; (2) an increase in eustatic sea level fall; and (3) an increase in sediment supply to the evolving basin. On the basis of the existing data, it is difficult to determine which of the above was the main cause for the regional regression across the area. The newly acquired MCS data from the Venezuelan Basin for the first time imaged the stratal geometry of the basal sequences and allowed us to develop a consistent model for the onshore deformation, the development of the South American drainage systems, and the delivery of sediments to the Caribbean. Because the preservation potential in the oceans is much greater than onshore, combining both data types provides for a more complete reconstruction of the tectonic and stratigraphic evolution of the Caribbean-South American Plate boundary. Eastward migration of the Caribbean Plate and the onset of compression between the North and South American plates in the early Tertiary (Pindell and Barrett, 1990) caused progressive deformation along the Caribbean-South American Plate boundary (Fig. 13). Nappes were emplaced southeastwards onto the Venezuelan margin since the Paleocene with the age of the deformation becoming systematically younger towards the east (Pindell et al., 1988; Pindell and Barrett, 1990; Bosch and Rodriguez, 1992). Although, there is controversy over the vergence direction in some of the northern basins (e.g., Maracaibo; Lugo and Mann, 1995), the onset of the deformation and increased uplift and
N.W. DRISCOLL and J.B. DIEBOLD subsidence appears to have occurred in the Late Paleocene to Early Eocene. Obduction of island-arc, frontal basin, and oceanic crust over the northdipping continental crust along the Frontal Thrust segmented the larger Maracaibo-Peruvian foreland basin into a series of smaller basins and associated topographic highs, which led to the formation of the Venezuelan Caribbean Mountains and Venezuelan foredeep (Kasper and Larue, 1986). Kasper and Larue (1986) proposed that the northern-trending Andes foredeep (i.e., Maracaibo-Peruvian foreland basin) and the northeast-trending Venezuelan foredeep acted as conduits for sediment transport to the Caribbean. Furthermore, terrigenous sediment transport through the Venezuelan foredeep might explain the existence of Paleogene sandstones on the Island of Barbados that appear to be derived from the Cordillera and Shield regions as well as the Mirador Formation sandstones that are interpreted as the fluvial deposits of a large Paleogene deltaic complex in Colombia and western Venezuela (Kasper and Larue, 1986). Northeast progradation over the present-day Lake Maracaibo region (e.g., Misoa Formation sandstones) in the Early to Middle Eocene is additional evidence for a southwestwardly derived sediment source at this time. Increased resistance to obduction along the Frontal Thrust may have produced a back thrust to accommodate the north-south strain between the Caribbean and South American plates (Bosch and Rodriguez, 1992). Along the northern boundary of the Venezuelan Caribbean Mountains a transform zone developed between the Caribbean Plate to the north and the obducted terranes to the south, which accommodated the large strike-slip motion. The east-west Oca, Bocon6, and Bocon6-Mor6nE1 Pilar fault systems exhibit major fight-lateral displacement since the Late Oligocene (Fig. 13). The two fault systems that partition the strain, the South Caribbean marginal fault and the North Venezuelan fault belt, define the northern and southern extent of the Bonaire Block (Bosch and Rodriguez, 1992). Continued uplift and deformation along the northern South America Plate boundary (e.g., Venezuelan Caribbean Mountains; Bellizzia and Dengo, 1990) affected the axial-parallel fluvial networks flowing into the Venezuelan Basin. Because the uplift first occurred along the western portion of the margin, the drainage systems were progressively deflected towards the east (Figs. 13 and 14A). The consequent increase in the overall river length (this evolving drainage system will become the Orinoco River) caused an increase in base level and deposition along the upper reaches of the drainage network. The uplift of the Venezuelan Caribbean Mountains that dammed and diverted the fiver systems along the northern boundary of South America in the
TECTONIC AND STRATIGRAPHIC D E V E L O P M E N T OF THE EASTERN C A R I B B E A N
619
Fig. 13. Plate reconstructions of the Caribbean during the Eocene and Early Miocene illustrating the progressive deformation of the northern South American Plate boundary (modified from Pindell and Barrett, 1990 and Pindell and Tabbutt, 1995). (Top) Onset of deformation along the northern boundary of South America during the Paleocene and Eocene, segmenting the larger Maracaibo-Peruvian foreland basin into a series of smaller basins and associated topographic highs. (Bottom) Continued deformation along the northern margin uplifted the Venezuelan Caribbean Mountains along the southward verging Frontal Thrust, which blocked the northward-flowing axial-parallel fluvial systems. This late-stage deformation, in large part, reorganized the drainage of South America and led to the development of the Orinoco and Amazon drainage systems. The east-west Oca, Bocon6, and Bocon6-Mor6n-el Pilar fault systems along northern South America exhibit major right-lateral displacement.
620
N.W. DRISCOLL and J.B. DIEBOLD
Fig. 14. (A) Topographic map of northern South America showing present-day drainage systems. Note that the Orinoco River has been diverted east and flows north along the western boundary of the Guyana Shield. Black bars denote barriers in the Marajo (MB) and Solim6es (SB) basins. WC -- Western Cordillera; CC = Central Cordillera; EC = Eastern Cordillera; MA = M6rida Andes; M = Lake Maracaibo; VCM = Venezuelan Caribbean Mountains. (B) Paleogeographic map for Late Cretaceous to early Tertiary illustrating the development of the Central Cordillera (CC) and the Maracaibo-Peruvian foreland basin. Radial drainage from the Guyana and Brazilian shields as well as the Central Cordillera are tributaries to the northward-flowing axial-parallel drainage. Early Miocene (Figs. 13 and 14A), together with a renewed phase of tectonic activity in the Central and Eastern Cordillera, dramatically changed the drainage pattern of South America. Prior to these events, the major drainage was to the north and supplied sediment to the Venezuelan Basin. The
thickness and spatial distribution of the onlapping basal turbidites and overlying current-controlled deposits indicate a southem source for the terrigenous sediments (Fig. 15), which persisted from Late Cretaceous to Early Miocene time (Figs. 4, 6 and 7). A dramatic change in the sediment supply occurred
TECTONIC AND STRATIGRAPHIC DEVELOPMENT OF THE EASTERN CARIBBEAN -76"
-72 ~
-68 ~
621
-64" O
20"
(7 O
16"
12"
Fig. 15. Regional isopach map for the eastern Caribbean region (modified and updated from Edgar et al., 1971) with the inferred current axis superposed. Note the thick depocenter (> 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.
REFERENCES
Abrams, L.J., Larson, R.L., Shipley, T.H. and Lancelot, Y., 1993. Cretaceous volcanic sequences and Jurassic oceanic crust in the east Mariana and Pigafetta Basins of the western Pacific. The Mesozoic Pacific: geology, tectonics, and volcanism. Geophys. Monogr., 77: 77-101. Anderson, D.L., Zhang, Y.-S. and Tanimoto, T., 1992. Plume heads, continental lithosphere, flood basalts, and tomography. In: B.C. Storey, T. Alabaster and R.J. Pankhurst (Editors), Magmatism and the Causes of Continental Break-Up. Geol. Soc. London Spec. Publ., 68: 99-124.
T E C T O N I C AND STRATIGRAPHIC D E V E L O P M E N T OF THE E A S T E R N C A R I B B E A N Baldwin, S.L., Harrison, T.M. and Burke, K., 1986. Fission track evidence for the source of accreted sandstones, Barbados. Tectonics, 5: 457-468. Bellizzia, A. and Dengo, G., 1990. The Caribbean mountain system, northern south America: a summary. In: G. Dengo and J.E. Case (Editors), The Geology of North America: The Caribbean Region. Decade of North American Geology, Vol. H, Geological Society of America, Boulder, CO, pp. 167-175. Biju-Duval, B., Mascle, A., Montadert, L. and Wanneson, J., 1978. Seismic investigation in the Colombia, Venezuela and Granada Basins, and on the Barbados Ridge for future IPOD drilling. Geol. Mijnbouw, 57:105-116. Biju-Duval, B., Bison, G., Mascle, A. and Muller, C., 1982. Active margin processes: field observations in southern Hispaniola. Am. Assoc. Pet. Geol. Bull., 67: 242-270. Bosch, M. and Rodriguez, I., 1992. North Venezuelan collisional crustal block: the boundary between the Caribbean and South American plates. J. S. Am. Earth Sci., 6: 133-143. Bowin, C., 1976. Caribbean gravity field and plate tectonics. Geol. Soc. Am. Spec. Pap., 169: 1-79. Bowland, C.L., 1993. Depositional history of the 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 of the western Colombian Basin, Caribbean Sea. Geol. Soc. Am. Bull., 100: 534-546. Browning, J.M. and Walper, J.L., 1982. A plate tectonic model for the evolution of the Maracaibo-Peruvian Trough. Pac. Geol., 16: 15-38. Burke, K., Cooper, C., Dewey, J.F., Mann, J.E and Pindell, J., 1984. Caribbean tectonics and relative plate motions. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mere., 162:31-64. Case, J.E. and Holcombe, T.L., 1980. Geologic-tectonic Map of the Caribbean region. U.S. Geol. Surv. Misc. Invest. Map I- 1100, scale 1 : 2,500,000. Case, J.E., Shagam, R. and Giegengack, R.F., 1990a. Geology of the northern Andes; an overview. In: G. Dengo and J.E. Case (Editors), The Geology of North America: The Caribbean Region. Decade of North American Geology, Vol. H, Geological Society of America, Boulder, CO, pp. 177-200. Case, J.E., MacDonald, W.D. and Fox, P.J., 1990b. Caribbean crustal provinces; seismic and gravity evidence. In: G. Dengo and J.E. Case (Editors), The Geology of North America: The Caribbean Region. Decade of North American Geology, Vol. H, Geological Society of America, Boulder, CO, pp. 15-36. Coffin, M.F. and Eldholm, O., 1994. Large igneous provinces: crustal structure, dimensions and external consequences. Rev. Geophys., 32: 1-36. Damuth, J.E., 1980. Use of high frequency (3.5-12 kHz) echograms in the study of near-bottom sedimentation processes in the deep-sea: a review. Mar. Geol., 38:51-75. Damuth, J.E. and Flood, R.D., 1984. Morphology, sedimentation processes, and growth pattern on the Amazon deep-sea fan. Geo-Mar. Lett., 3:109-117. Dewey, J.F. and Pindell, J.L., 1985. Neogene block tectonics of Turkey and northern South America: continental applications of the finite difference method. Tectonics, 4: 71-73. Diebold, J.B., Stoffa, EL., Buhl, E and Truchan, M., 1981. Venezuelan Basin crustal structure. J. Geophys. Res., 86: 7901-7923. Diebold, J.B., Driscoll, N.W., Abrams, L., Donnelly, T.W., Laine, E.E and Leroy, S., 1995a. Tectonic model for the origin and evolution of the rough/smooth basement in the Venezuelan Basin. EOS, 76: F615. Diebold, J.B., Driscoll, N.W., Abrams, L., Donnelly, T.W., Laine, E.E and Leroy, S., 1995b. Stratigraphy and tectonics of the
625
Venezuelan Basin and Beata Ridge; results from a new regional geophysical survey. Geological Society of America, Special Program: International Division and IGCP, Project 364: Caribbean Arcs and Ophiolites: Geology and Tectonics of the Caribbean Region. Donnelly, T.W., 1973. Magnetic anomaly observations in the eastern Caribbean Sea. Init. Rep. DSDP, 15:1023-1030. Donnelly, T.W., 1989. Geologic history of the Caribbean and Central America. In: A.B. Balley and A.R. Palmer (Editors), The Geology of North America, Decade of North American Geology, Vol. A, Geological Society of America, Boulder, CO, pp. 299-321. Donnelly, T.W., 1990. Pelagic sediment, deep water chemistry, and tectonics: an application of the history of biological sediment accumulation on the tectonic history of the Caribbean. Riv. Ital. Paleontol. Stratigr., 96 (2-3): 143-164. Donnelly, T.W. and 10 others, 1990. History and tectonic setting of Caribbean magmatism. In: G. Dengo and J.E. Case (Editors), The Geology of North America: The Caribbean Region. Decade of North American Geology, Vol. H, Geological Society of America, Boulder, CO, pp. 339-374. Driscoll, N.W. and Laine, E.E, 1996. Abyssal current influence on the southwest Bermuda Rise and surrounding region. Mar. Geol., 130:231-263. Driscoll, N.W., Hogg, J.R., Karner, G.D. and Christie-Blick, N., 1995. Extensional tectonics in the Jeanne d'Arc basin: implications for the timing of break-up between Grand Banks and Iberia. In: R.A. Scrutton, M.S. Stoker, G.B. Shimmield and A.W. Tudhope (Editors), The Tectonics, Sedimentation and Palaeoceanography of the North Atlantic Region. Geological Society London Special Publ., 90: 1-28. Duncan, R.A. and Hargraves, R.B., 1984. Plate tectonic evolution of the Caribbean region in the mantle reference frame. In: W.E. Bonini, R.B. Hargraves and R. Shagam (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162:81-94. Edgar, N.T., Ewing, J.I. and Hennion, J., 1971. Seismic refraction and reflection in the Caribbean Sea. Am. Assoc. Pet. Geol. Bull., 55: 833-870. Edgar, N.T., Saunders, J.B., et al., 1973a. Initial Reports of the Deep Sea Drilling Project, Vol. 15. Washington U.S. Government Printing Office, Washington, DC, 1137 pp. Edgar, N.T., Holcombe, T. and Ewing, J., 1973b. Sedimentary hiatuses in the Venezuelan Basin. Init. Rep. DSDP, 15:10511062. Escalante, G., 1990. The geology of southern Central America and western Colombia. In: G. Dengo and J.E. Case (Editors), The Geology of North America: The Caribbean Region. Decade of North American Geology, Vol. H, Geological Society of America, Boulder, CO, pp. 201-230. Ewing, J., Edgar, N.T. and Antoine, J.W., 1971. Structure of the Gulf of Mexico and Caribbean Sea. In: Maxwell (Editor), The Sea, Vol. 4, part 2. Wiley, New York, pp. 321-358. Fox, EJ., Ruddiman, W.F., Ryan, W.B.F. and Heezen, B.C., 1970. The geology of the Caribbean crust, I. Beata Ridge. Tectonophysics, 10:495-513. Heubeck, C. and Mann, E, 1991. Geologic evaluation of plate kinematic models for the North-American-Caribbean plate boundary zone. Tectonophysics, 191: 1-26. Hinz, K., 1981. A hypothesis on terrestrial catastrophes. Wedges of very thick oceanward dipping layers beneath passive continental margins. Geol. Jahrb., E22: 3-28. Holcombe, T.L., Ladd, J., Westbrook, G., Edgar, N.T. and Bowland, C., 1990. Caribbean marine geology; ridges and basins of the plate interior. In: G. Dengo and J.E. Case (Editors), The Geology of North America: The Caribbean Region. Decade of North American Geology, Vol. H, Geological Society of America, Boulder, CO, pp. 231-260.
626 Hoorn, C., Guerrero, J., Sarmiento, G.A. and Lorente, M.A., 1995. Andean tectonics as a cause for changing drainage patterns in Miocene northern South America. Geology, 23: 237-240. Houtz, R.E. and Ludwig, W.J., 1977. Structure of Colombia Basin, Caribbean Sea, from profiler-sonobuoy measurements. J. Geophys. Res., 82: 4861-4867. Kasper, D.C. and Larue, D.K., 1986. Paleogeographic and tectonic implications of quartzose sandstones of Barbados. Tectonics, 5: 837-854. Kellogg, J.N. and Vega, V., 1995. Tectonic development of Panama, Costa Rica, and the Colombian Andes: constraints from Global Positioning System geodetic studies and gravity. In: P. Mann (Editor), Geologic and Tectonic Development of the Caribbean Plate Boundary in Southern Central America. Geol. Soc. Am. Spec. Pap., 295: 75-90. Klitgord, K. and Schouten, H., 1987. Plate kinematics of the Central Atlantic. In: B. Tucholke and P. Vogt (Editors), The Geology of North America: The Western Atlantic Region. Decade of North American Geology, Vol. M, Geological Society of America, Boulder, CO, pp. 351-378. Kolla, V., Buffler, R.T. and Ladd, J.W., 1984. Seismic stratigraphy and sedimentation of Magdalena fan, southern Colombian Basin, Caribbean Sea. Am. Assoc. Pet. Geol. Bull., 68: 316332. Kroonenberg, S.B., Bakker, J.G.M. and van de Wiel, M., 1990. Late Cenozoic uplift and paleogeography of the Colombian Andes: constraints on the development of high-Andean biota. Geol. Mijnbouw, 69: 279-290. Ladd, J.W. and Watkins, J.S., 1980. Seismic stratigraphy of the western Venezuela Basin. Mar. Geol., 35:21-41. Lu, R.S. and McMillen, K.J., 1982. Multichannel seismic survey of the Colombia Basin and adjacent margins. In: J.S. Watkins and C.L. Drake (Editors), Studies in Continental Margin Geology. Am. Assoc. Pet. Geol. Mem., 34: 395-410. Ludwig, W.J., Houtz, R.E. and Ewing, J.I., 1975. Profilersonobuoy measurements in Colombian and Venezuelan Basins, Caribbean Sea. Am. Assoc. Pet. Geol. Bull., 59:115-123. Lugo, J. and Mann, P., 1995. Jurassic-Eocene tectonic evolution of Maracaibo Basin, Venezuela. In: A.J. Tankard, R.S. Soruco and H.J. Welsink (Editors), Petroleum Basins of South America. Am. Assoc. Pet. Geol. Mem., 62: 699-726. Macelli, C.E., 1988. Cretaceous paleogeography and depositional cycles of western South America. J. S. Am. Earth Sci., 1: 373418. Malfait, B.T. and Dinkelman, M.G., 1972. Circum-Caribbean tectonic and igneous activity and the evolution of the Caribbean plate. Geol. Soc. Am. Bull., 83:251-272. Mauffret, A. and Leroy, S., 1997. Seismic stratigraphy and structure of the Caribbean Sea. Tectonophysics, 283:61-104. McCave, I.N. and Tucholke, B.E., 1986. Current controlled sedimentation in the Western North Atlantic. In: P. Vogt and B. Tucholke (Editors), The Geology of North America: The Western Atlantic Region. Decade of North American Geology, Vol. M, Geological Society of America, Boulder, CO, pp. 451-478. Mutter, J.C., Talwani, M. and Stoffa, P.L., 1982. Origin of seaward-dipping reflectors in oceanic crust off the Norwegian margin by subaerial sea-floor spreading. Geology, 10: 353357. Officer, C.B., Ewing, J.I., Edwards, R.S. and Johnson, H.R., 1957. Geophysical investigations in the Eastern Caribbean: Venezuelan Basin, Antilles island arc, and Puerto Rico trench. Bull. Geol. Soc. Am., 68: 359-378.
N.W. D R I S C O L L and J.B. D I E B O L D Officer, C.B., Ewing, J.I., Hennion, J.E, Harkrider, D.G. and Miller, D.E., 1959. Geophysical Investigations in the Eastern Caribbean: summary of 1955 and 1956 cruises. In: L.H. Ahrens et al. (Editors), Physics and Chemistry of the Earth, Vol. 3. Pergamon Press, London, pp. 17-109. Pindell, J.L. and Barrett, S.E, 1990. Geological evolution of the Caribbean region: a plate-tectonic perspective. In: G. Dengo and J.E. Case (Editors), The Geology of North America: The Caribbean Region. Decade of North American Geology, Vol. H, Geological Society of America, Boulder, CO, pp. 405-432. Pindell, J.L. and Dewey, J.E, 1982. Permo-Triassic reconstruction of western Pangea and the evolution of the Gulf of Mexico/Caribbean region. Tectonics, 1: 179-212. Pindell, J.L. and Tabbutt, K.D., 1995. Mesozoic-Cenozoic Andean Paleogeography and regional controls on hydrocarbon systems. In: A.J. Tankard, R.S. Soruco and H.J. Welsink (Editors), Petroleum Basins of South America. Am. Assoc. Pet. Geol. Mem., 62: 101-128. Pindell, J.L., and 6 others, 1988. A plate-kinematic framework for models of Caribbean evolution. Tectonophysics, 155: 121138. Schlich, R., Rotstein, Y. and Schaming, M., 1993. Dipping basement reflectors along volcanic passive margins - - new insight using data from the Kerguelen Plateau. Terra Nova, 5: 157-163. Schumm, S.A., 1991. River response to baselevel change: implications for sequence stratigraphy. J. Geol., 101: 279-294. Sigurdsson, H., Leckie, R.M., Acton, G.D., et al., 1997. Proceedings of the Ocean Drilling Program, Initial Report, 165. College Station, TX, 865 pp. Sinton, C.W., Duncan, R.A. and Storey, M., 1993. 4~ ages from Gorgona Island, Colombia, and the Nicoya Peninsula, Costa Rica. EOS, 74: 553. Stoffa, EL., Mauffret, A., Truchan, M. and Buhl, E, 1981. Sub-B" layering in the southern Caribbean: the Aruba Gap and Venezuela Basin. Earth Planet. Sci. Lett., 53: 131-146. Talwani, M., Windisch, C., Stoffa., EL., Buhl, E and Houtz, R.E., 1977. Multichannel seismic study in the Venezuelan Basin and the Curacao Ridge. In: M. Talwani and W.C. Pitman III (Editors), Island Arcs, Deep Sea Trenches and Back-Arc Basins. Am. Geophys. Union, M. Ewing Symp. Ser., 1: 83-98. Tankard, A.J., Soruco, R.S. and Welsink, H.J., 1995. Petroleum Basins of South America. Am. Assoc. Pet. Geol. Mem., 62, 761 pp. Westall, E, Rossi, S. and Mascle, A., 1993. Current-controlled sedimentation in the Equatorial Atlantic: examples from the southern margin of the Guinea Plateau and the Romanche Fracture Zone. Sediment. Geol., 82: 157-171. White, R.S. and McKenzie, D., 1995. Mantle plumes and flood basalts. J. Geophys. Res., 100: 17,543-17585. Wilson, J.T., 1966. Are the structures of the Caribbean and Scotia arcs analogous to ice rafting? Earth Planet. Sci. Lett., 1: 335-338. Worthington, L.V., 1966. Recent oceanographic measurements in the Caribbean Sea. Deep-Sea Res., 13: 731-739. Worthington, L.V. and Wright, W.R., 1970. North Atlantic ocean atlas of potential temperature and salinity in deep water. The Woods Hole Oceanographic Atlas Series Volume II. Wright, W.R. and Worthington, L.B., 1970. The water masses of the North Atlantic ocean: a volumetric census of temperature and salinity. Serial Atlas of the Marine Environment, Folio 19. American Geographical Society, New York, 8 pp. and 7 plates. Wunsch, C. and Grant, B., 1982. Towards the general circulation of the North Atlantic Ocean. Prog. Oceanogr., 11: 1-59.
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-
r
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
Z'-6 ~,,..~
9
,.~
o
"~
'-~
O
~
~zZ
~
~
~;_v,
~
~
~
~,...~
o "~
~
,-~
.,.
~''-d
9
~
.
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.
642
A. MAUFFRET and S. LEROY 72~
16~
[I
I ~, ~, ~
15~
72"W JI
I I _k'~/t/
p
I,....-"'
71~ 116~
I
15ON
]34 DSDP 31 B3
B2
B1
g 0
72~
9
72W~'
Depth to bucmeat
71~
contour iaterv~: 0.2 km
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.
t-
>
4~ 4~
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
645 == 9
..s=
rae~
,.IZZ
9
O
..s=
9
c.tz .,..~
.,..o .X:Z
. ,...~ ,..o
. ,...,
r.gz c~
,
9
e~
,.o
~ m
m ~
m
r,~
. ,...~ ~
~
r
~
.,-~
<
=
r~
646
A. MAUFFRET and S. LEROY
~ ,...~ k~
.,:..;
9 ,.t:Z
~
~D
,....:. r
r,.) tt~
. ,...~ ;.r..,
~
e~t)
9
[..., ~
9
~
~ ,,,,.~
. ,...~ 9
.=_,
+,.a O
xz
9
t.l:Z
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
647 ..c:
,.o
~D
..o
.x::
~D
,.o
(D
9....~
.,.~
-,~
~'~
""~
o
~~
.,.~
~
~-~
~."
.,.~ ~.~
9
.,~
.~~
.~ <
b
648
A. M A U F F R E T and S. L E R O Y
Fig. 19. A tilted block maybe related to a strike-slip fault. The magnetic profile suggests that this block is a volcano. Note the prominent reflector R visible also in Figs. 16 and 18. See the inset map (depth to basement: B") and Fig. 13 for profile location.
Fig. 20. A seamount has a probable volcanic origin as shown by the magnetic curve. This seamount does not seem to be affected by any deformation. See the inset map (depth to basement: B") and Fig. 13 for profile location.
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
649
Fig. 21. Depth to basement map (B'~; modified from Leroy and Mauffret, 1996). This figure shows the contact between the Colombian basin and the Beata ridge. The seismic profiles are calibrated by the DSDP Site 153 results (Hopkins, 1973). The locations of the seismic profiles shown in Figs. 22 and 23 are indicated. B" reflector is 4.3 km deep on the western side of the Taino ridge, 3.7 km on the top and 4.5 km on the eastern side (Fig. 18B). The uplift of the Taino ridge is about 1 km if we neglect a probable initial topography. The western fault is apparently the master fault, but the cumulate throw of the eastern reverse faults is larger than the western fault and the B" reflector is 0.2 km higher on the western side than on the eastern side (4.3 and 4.5 km, respectively, Fig. 18B). The master fault is probably a former normal listric fault with east-facing throw because we cannot explain the eastwards deepening of the B" reflector with the observed motion of the reverse faults (Fig. 18C). We conclude that the east-
wards regional dip, evident for the B" reflector and the sea floor, and normal faults are initial features formed during the construction of the Beata volcanic plateau (Diebold et al., 1999; Driscoll and Diebold, 1999). However, this dip has been probably accentuated during the Miocene to Present compressional event. The results of the DSDP Site 31 indicate approximately 1 km of uplift during the Miocene (Benson et al., 1970). This uplift of the Beata ridge is regional and unrelated to the Taino ridge because the DSDP Site 31 is located at the base of this structure (Fig. 18A). This regional uplift probably occurs step by step from the undisturbed Venezuelan basin to the Beata ridge. The Taino ridge is one
650
A. MAUFFRET and S. LEROY
Fig. 22. Seismic profile showing the Pecos fault zone (modified from Leroy and Mauffret, 1996). This is a pop-up structure delimited by reverse faults. Note the difference in level of the reflectors A" and B" between the Colombian and the Venezuelan basins. The Early Miocene (eM) is deformed, whereas the recent sedimentary layers onlap the structure. The sub-B" reflector (Stoffa et al., 1981) is prominent. The compression began during Miocene time, but the deformation of the sea floor indicates that this compression is yet active. See the inset map and Fig. 21 for profile location. of these steps but other reverse faults exist east of the ridge (Figs. 11 and 12). The dominant motion is probably a thrust of the Beata ridge over the Venezuelan basin, but our survey is not enough extended to the east. However, the crustal fault shown in Fig. 11C has a clear westwards dip. The combination of N - S to N N W - S S W thrusts, N E - S W fight-lateral and N W - S E left-lateral strike-slip faults (Fig. 13) indicates that the stress is oriented N E SW.
ARUBA GAP
This area, that was already studied (Hopkins, 1973; Stoffa et al., 1981; Leroy, 1995; Leroy et al., 1996; Leroy and Mauffret, 1996), is located between
the Beata ridge and the South American deformed belt and forms a sill between the Colombian and Venezuelan basins (Fig. 1). A Casis seismic detailed survey in addition to previous seismic profiles (Fig. 21) allow us to contour a detailed map of the acoustic basement (Fig. 21). The acoustic basement of the eastern Colombian basin is relatively smooth and we can identify the typical B" reflector in the eastern Colombian basin (Fig. 22), whereas the true oceanic rough crust is localized in the western part of the Colombian basin (Bowland and Rosencrantz, 1988). DSDP Site 153 (Edgar et al., 1973b) allows calibration of the seismic profiles (Hopkins, 1973) and identification (Fig. 22) of a prominent Middle Miocene horizon, an Early Miocene marker at the top of a disturbed layer, the Middle Eocene A" reflector, and the top of the Cretaceous volcanic
N E O G E N E INTRAPLATE D E F O R M A T I O N OF THE C A R I B B E A N PLATE AT THE BEATA RIDGE
plateau (B" reflector). In addition an internal reflector named sub-B" was identified (Stoffa et al., 1981; Leroy et al., 1996). The acoustic basement of the Colombian and Venezuelan basins dips to the south towards the South American accretionary prism, which has been active since the Early to Middle Miocene (Biju Duval et al., 1982b). The Pecos fault zone that trends N W - S W separates the Beata from the Colombian basin (Figs. 21 and 22). Reverse faults, with a throw of 100-200 m, are clearly displayed on the Casis seismic lines (Fig. 22). The reverse faults are facing towards the northeast and southwest and the Pecos fault zone is clearly a compressive feature. Flower structures (Fig. 23) indicate also the presence of strike-slip faults. The identification, by seismic correlation with the western Colombian basin, of the early Miocene reflector in the eastern Colombian basin at the base of a well layered seismic unit (Fig. 22), indicates that the main tectonic event occurred during the Early Miocene (23 Ma) because the post-Early Miocene layers of the Colombian basin onlap the flank of the Pecos fault zone (Fig. 22), but some recent reactivations (Fig. 23) are evident and the sea floor is deformed. In the southern part of the study area, near and below the accretionary prism, the deformation is weak and the Pecos structure is low (less of 0.4 km of throw relative to the Colombian basin; Cas A10, Fig. 24). Towards the north the deformation increases (Cas A07, Fig. 22; Cas A08, Fig. 24) and a new area of deformation appears at the base of the Pecos fault zone (Cas A06, Fig. 24; Figs. 23 and 25). The top of the Pecos fault zone is 1 km higher than the floor of the Colombian basin (5.4 km and 6.4 km deep, respectively; Fig. 21). Nevertheless, we do not know the initial topography in this area. The Early MioceneA" interval (Fig. 22) is much thicker (800 m) in the Colombian basin than in the Venezuelan basin (300 m) and this observation indicates that some topography (500 m?) existed prior to the Miocene deformation. The B"-sub-B" interval and the crust are thinner in the Colombian basin than in the Venezuelan basin and this is again an argument to suppose a different initial level between the two basins. The 1 km step between the Pecos fault zone and the Colombian basin results from an initial topography (500 m?) and a Miocene reactivation (500 m?). It is clear that the deformation along the NW-SE-trending Pecos fault zone increases towards the north and consequently is not related to the subduction of the volcanic plateau beneath the South American deformed belt (Burke et al., 1978). A first-motion solution for an earthquake in the southern Beata ridge was reinterpreted from Molnar and Sykes (1969) by Kafka and Weidner (1981). Strike-slip and reverse motions were determined re-
651
spectively. The main stress is oriented N E - S W in the two interpretations (Fig. 25). This stress is compatible with a compression along the Pecos fault zone and left-lateral faults that trend E - W (Fig. 25). We estimated the shortening of the Pecos structure to 40 km per degree of latitude (Leroy and Mauffret, 1996).
DSDP 151 RIDGE
This ridge forms the backbone of the Beata ridge (Figs. 2 and 3). In the Colombian basin the seismic profile Cas-C03 (Fig. 26) crosses the southern extension of the DSDP 151 ridge. This part of the ridge is not reactivated, the top of the ridge is 4.7 km deep and its high relative to the adjacent basement of the Colombian basin is 1.9 km (Fig. 26C). The 5.6 km depth of the top of the ridge on the depth to basement map (Fig. 26B) is a mistake due to the smoothing effect of the automatic contouring. The buried ridge is clearly offset fight-laterally relatively to the northern ridge by faults that trend N E - S W (Fig. 26B). A prominent seamount, 1.8 km deep (2.4 km on the depth to basement, Fig. 26B) is located in the southern part of the DSDP 151 ridge. This ridge is 130 km long and flanked by a 4.4-km-deep depression to the east (Fig. 27A). The seismic profiles recorded during the Seacarib 1 cruise allow a good control of the shape of the ridge. The Seabeam map (Fig. 27B) shows the northern end of the ridge where is located the DSDP Site 151 (Fig. 28B). The trend of the DSDP 151 ridge that is north-south changes abruptly and is delimited by a N E - S W steep scarp (Figs. 27 and 28B). This scarp bounds a basin that trends also N E - S W (Fig. 27).
W E S T E R N BOUNDARY OF THE BEATA RIDGE
A steep scarp delimits the Beata ridge and the Haiti sub-basin (Fig. 29). The Seabeam map (Fig. 29B) clearly displays the steepness of the slope oriented NE-SW. However, the Tairona ridge, 1 km deep, shows a N-S trend. A deep southern terrace (between 3 and 4 km, Fig. 29B) is progressively accreted to the slope towards the north and finally disappears. The upper part of the slope is less abrupt and is covered by slumped sediments (Fig. 30), but this profile does not show the true dip because it is oblique to the slope. The lower part of sedimentary cover of the Haiti sub-basin is deformed (Fig. 30). We think that this deformation is related to recent compression, although presently inactive, but the quality of the seismic profile is too poor to see the wedge displayed in Fig. 8.
652
A. MAUFFRET and S. LEROY
Fig. 23. Flower structure indicating the presence of strike-slip fault with a compressional component (transpression). See the inset map and Fig. 21 for profile location.
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
653
Fig. 24. Depth sections (modified from Leroy and Mauffret, 1996) showing the evolution of the Pecos fault zone from south to north. In the south (Cas A10) the structure is low and affected by few reverse faults; the sea floor is undeformed. The Cas A08 profile shows the rising of the structure and the formation of a pop-up. The northern section (Cas A06) shows a deformed sea floor and the formation of a second zone of deformation at the base of the Pecos fault zone. See the inset map and Fig. 22 for profile location.
DISCUSSION The western and eastern flanks of the Beata ridge are completely different. On the eastern flank the compressive structures are predominant, but these structures are offset by right-lateral strike-slip faults that trend NE-SW. On the western flank the faults that trend N E - S W are the main structural features and we failed to identify any compressional structure (Fig. 31). However, our seismic profiles do not have a good quality and the DSDP 151 ridge may have been reactivated and uplifted by compression. This ridge and the Tairona ridge trend N - S and were probably the same initial feature; consequently, the faults that trend N E - S W may have a strong fight-lateral strike-slip component. The basin that delimits the DSDP 151 ridge to the north could be a pull-apart basin that trends NE-SW. Recent visual observations from a submersible survey (Mauffret et al., in prep.) indicate that the steep scarp that bounds the Haiti sub-basin is divided in short segments of north-south thrusts and northeast-southwest strikeslip faults. The fault that bounds the Haiti sub-basin
and the Hess escarpment were supposed to have a normal component related to rifting of the Haiti subbasin (Donnelly et al., 1995; Driscoll and Diebold, 1999). However, this basin is closed to the south by the Warao rise (Fig. 31) that formed a high connecting the Beata ridge to the Nicaragua rise (Mauffret and Leroy, 1997). This rise cannot be formed by a rifting of the Haiti sub-basin and we suggest that this basin is an initial feature surrounded by volcanic plateaus (Mauffret and Leroy, 1997). The structural activity that affected the southern part and the Beata plateau has been related to lithospheric flexure accompanying Miocene to Recent convergence along the Curaqao ridge (Holcombe et al., 1990). A lithospheric flexure to the north (Muertos trench) and to the south (Curaqao ridge) is indeed evident in the Venezuelan basin and is surrounded by the 4-km depth contour (Fig. 2). However, the normal faults formed by these flexures trend E - W (Silver et al., 1975; Diebold et al., 1999), whereas the structures described in this paper (Taino ridge, Beata plateau, Pecos fault zone) are oriented N N W SSE. Consequently these structures that show evi-
654
A. MAUFFRET and S. LEROY
Fig. 25. Structural interpretation of the Aruba Gap area (modified from Leroy and Mauffret, 1996). The Pecos fault zone is a pop-up structure delimited by reverse faults. A second zone of deformation exists at the base of the northern part of the Pecos fault zone. The shape of this deformed zone and the flower structure shown in Fig. 23 suggest left-lateral strike-slip faults. These faults and the Pecos fault zone are compatible with a compressive stress oriented NE-SW. Such a stress is also determined by earthquake focal mechanism shown in the top of the figure (Kafka and Weidner, 1981). The locations of the depth sections shown Fig. 24 are indicated.
d e n c e o f c o m p r e s s i o n and strike-slip faulting c a n n o t be r e l a t e d to the l i t h o s p h e r i c flexure of the C a r i b b e a n
T h e c o m p r e s s i v e stress, a c c o r d i n g to our structural study and focal m e c h a n i s m analysis ( K a f k a and
plate.
Weidner, 1981) is o r i e n t e d N E - S W .
T h e s a m e off-
NEOGENE INTRAPLATE DEFORMATION OF THE CARIBBEAN PLATE AT THE BEATA RIDGE
655 .,.., ,.i=:
.,..~
O .,,o O . ,...,
.,.., .xZ
b ~
go . ,...,
~o
r.~
,.o
9
t".-
c,.) ,4 ,.lzZ
9~
.xz
tth
9
r~
..xZ
,..o 9
,..o
~D
,..o
O
ra~ .,..~
eq~
656
A. M A U F F R E T and S. L E R O Y
73~ 15045N -
73~ I
730W
73"15W
72054W 15"45N
SEAeARIB' BEATA IM ER.CATORprojection isobathlcontour: 100 m , ~lid line: 5(X)m
~ ii i
v/
15~
15030N
15015N,
15"15N
Rg. 28 B
it
> Fig. 28 A 15~
150N
DSDP site 151
14054N. 73~
,o 7:3~
,14054N
7:30W
7:3~
72054W
B
~.2~
150N
14~
74~
73~
A
Fig. 27. (A) Depth to basement map: B 't. (B) Seabeam map of the Seacarib 1 cruise. The position of the survey is indicated in Fig. 3. The northern tip of the DSDP 151 ridge, that trends north-south, terminates abruptly towards the north and a steep scarp shows a northeast-southwest orientation. The positions of the seismic profiles illustrated in Fig. 28 are shown.
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
17~
71~
72~
72~
71~
i
:~~:~e~RIE ~BEATAI~. : I i MERCAT(,R :l)rojeet~ I eorlour:lOOm I :Solid line: !
9:,#j/~ ~. . . .
17~
.,,.~"
[
17~
17~
7~
17~
4000
17~
,17ON 2000
116~
16~
16~
a 72~
72~
72~
71~
71~
q
17~
16~
A 73~
72~
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
i/
0.91kin
-2000
Haiti Rateau
-4000 -6OO0 F
-74.00
__~
0.00
1.1 km -2000 L ~
Haiti sub basin
Dominican / sub basin
-40O0 -6O0O -74.00
/
20001 NiRicsearagua
-7:3.00
/
I
-70.00
2.4kin 1.7 km
-4000 -6000 D
o
-74.00 19~ ~
ud
~2~
-Tz00
/
-70.00
-2000 ' ~ / . 7 km | ~
-4000
I VenezuelaBasin
"
"O
18~
-6000 C
r'a ~" 03
-74.00 Haiti sub basin )
-2000 16~
._, s ~ ~ - ~
1.8 km
WaraoRise
BeataPlateauu
_
-74.00
-72.00
-70.00
"1
-2000
4~
4.7 km
14~
tl
-4000
ColombiaBasin
~L
Pecosfault 1.9km
-6000 7.4 km G
TainoRidge 2.2 km
-4000
15~
...... ColombiaBasin \
-70.00
-600O
BeataRidge
3.1o
-72.00
17~
VenezuelaBasin -
3ON
-74.00
-72.00
-70.00 Vertticalexaggeration:X40
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
N E O G E N E I N T R A P L A T E D E F O R M A T I O N OF THE C A R I B B E A N PLATE AT THE BEATA R I D G E of the western Colombian basin. Geol. Soc. Am. Bull., 100: 534-546. Bryne, D., Suarez, G. and McCann, W.R., 1985. Muertos trough subduction; microplate tectonics in the northern Caribbean. Nature, 317: 420-421. Burke, K., Fox, E and Seng6r, A.M.C., 1978. Buoyant ocean floor and the evolution of the Caribbean. J. Geophys. Res., 83, 3949-3954. Calais, E. and Mercier de Lepinay, B., 1993. Semiquantitative modeling of strain and kinematics along the Caribbean/North America strike-slip plate boundary zone. J. Geophys. Res., 98: 8293-8308. Calais, E., Mercier de Lepinay, B., Saint-Marc, E, Butterlin, J. and Schaaf, A., 1992. La limite de plaques d6crochante nord carfffbe en Hispaniola: 6volution paldog6ographique et structure c6nozoique. Bull. Soc. G6ol. Fr., 163" 309-324. Case, J.E. and Holcombe, T.L., 1980. Geologic-tectonic map of the Caribbean region. Scale 1:2,500,000. U.S. Geol. Surv. Misc. Invest. Ser., Map I-II00. Deng, J. and Sykes, L.R., 1995. Determination of Euler pole for contemporary motion of Caribbean and North American plates using slip vectors of interplate earthquakes. Tectonics, 14: 39-53. Dengo, G., 1985. Mid America: tectonic setting for the Pacific Margin from southern Mexico to northwestern Colombia. In: A.E.M. Nairn, EG. Stehli et al. (Editors), The Ocean Margin and Basin, 7A. The Pacific Ocean. Plenum Press, New York, pp. 123-180. Diebold, J.D., Driscoll, N.W. and EW-9501 Science Team, 1999. New insights on the formation of the Caribbean basalt province revealed by multichannel seismic images of volcanic structures in the Venezuelan Basin. In: E Mann (Editor), Caribbean Sedimentary Basins. Elsevier, Amsterdam, pp. 561-589. Dillon, W.E, Austin, J.A., Scanlon, K.M., Edgar, N.T. and Parson, L.M., 1992. Accretionary margin of north-western Hispaniola: morphology, structure and development of part of the northern Caribbean plate boundary. Mar. Pet. Geol., 19: 7088. Donnelly, T.W., Duncan, R.A., Sinton, C., Diebold, J., Driscoll, N., Abrams, L. and Mauffret, A., 1995. Proposal for drilling a major large igneous province: the Caribbean Cretaceous basalt province. Proposal ODP 480, 53 pp. Drewes, H.D., Kaniuth, K., Stber, K., Tremel, H., Kahle, H.G., Straub, C., Hernandez, N., Hoyer, M. and Wildermann, E., 1995. The Casa 93 GPS campaign for crustal deformation research along the south Caribbean plate boundary. J. Geodyn., 20:129-144. Driscoll, N.W. and Diebold, J.D., 1999, Tectonic and stratigraphic development of the Caribbean: new constraints on the Caribbean-South American plate. In: E Mann (Editor), Caribbean Sedimentary Basins. Elsevier, Amsterdam, pp. 591-626. Duncan, R.A. and Hargraves, R.B., 1984, Plate tectonic evolution of the Caribbean region in the mantle reference frame. In: W.D. Bonini, R.B. Hargraves and R. Shagam (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geological Society of America, pp. 81-94. Duque-Caro, H., 1979. Major structural elements and evolution of northwestern Colombia. In: J.S. Watkins, L. Montadert et al. (Editors), Geological and Geophysical Investigations of Continental Margins. Am. Assoc. Pet. Geol. Mem., 29: 329351. Duque-Caro, H., 1984. Structural style, diapirism, and accretionary episodes of the Sinu-San Jacinto terrane, southwestern Caribbean borderland. In: W.E. Bonini, R.B. Hargraves et al. (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162:303-316.
667
Edgar, T.N., Ewing, J. and Hennion, J., 1971. Seismic refraction and reflection in the Caribbean Sea. Am. Assoc. Pet. Geol. Bull., 55, 833-870. Edgar, T.N., Holcombe, T., Ewing, J. and Johnson, W., 1973a. Sedimentary hiatuses in the Venezuela basin. In: T.N. Edgar and J.B. Saunders et al. (Editors), Initial Report of the Deep Sea Drilling Project. U.S. Government Printing Office, Washington, D.C., Leg 15, pp. 1051-1062. Edgar, T.N., Saunders, J.B. et al. (Editors), 1973b. Initial Report of the Deep Sea Drilling Project. U.S. Government Printing Office, Washington, D.C., Leg 15. England, E and Wortel, R., 1980. Some consequences of the subduction of young slab. Earth Planet. Sci. Lett., 47: 403415. Ewing, J., Antoine, J. and Ewing, M., 1960. Geophysical Measurements in the Western Caribbean Sea and in the Gulf of Mexico. J. Geophys. Res., 65: 4087-4125. Farina, F., Calais, E., DeMets, C., Dixon, T., Jansma, E and Mann, E, 1995. GPS measurements across the northern Caribbean plate boundary zone: preliminary results. Eos 76 (17): $94. Fisher, D.M., Gardner, T.W., Marshall, J.S. and Montero, W., 1994. Kinematics associated with late Cenozoic deformation in Central Costa Rica: western boundary of the Panama microplate. Geology, 22, 263-266. Fox, EJ. and Heezen, B.C., 1975. Geology of the Caribbean crust. In: A.E.M. Nairn and T.G. Stehli (Editors), The Ocean Basins and Margins, Vol. 3. The Gulf of Mexico and the Caribbean. Plenum Press, New York, pp. 421-466. Fox, P.J., Ruddiman, W.F., Ryan, W.B.F. and Heezen, B.C., 1970. The geology of the Caribbean crust, 1. Beata ridge. Tectonophysics, 10:495-513. Freymueller, J.T., Kellogg, J.N. and Vega, V., 1993. Plate motions in the North Andean Region. J. Geophys. Res., 98: 21853-21863. Heubeck, C. and Mann, E, 1991. Geologic evaluation of plate kinematic models for the North American-Caribbean plate boundary zone. Tectonophysics, 191: 1-26. Hilst, R. and Mann, E, 1994, Tectonic implications of tomographic images of subducted lithosphere beneath northwestern South America. Geology, 22:451-454. Holcombe, T.H. and Moore, W.S., 1977. Paleocurrents in the eastern Caribbean: geologic evidence and implications. Mar. Geol., 23: 35-56. Holcombe, T.L., Ladd, J.W., Westbrook, G., Edgar, T. and Bowland, C.L., 1990. Caribbean marine geology and basins of the plate interior. In: G. Dengo and J.E. Case (Editors), The Geology of North America, Vol. H. The Caribbean Region. A Decade of North American Geology, Geological Society of America, Boulder, Colo., pp. 231-260. Hopkins, H.R., 1973. Geology of the Aruba Gap abyssal plain near DSDP Site 153. In: T.N. Edgar, J.B. Saunders et al. (Editors), Initial Reports of the Deep Sea Drilling Project. Washington, D.C., Leg 15, pp. 1039-1050. Houtz, R.Z. and Ludwig, W.J., 1977. Structure of the Colombia basin, Caribbean Sea, from profiler sonobuoy measurements. J. Geophys. Res., 82:4861-4867. Jany, I., 1989. Ndotectonique au sud des Grandes Antilles, collision (ride de Beata, Presqu'~le de Bahoruco)-subduction (fosse de Muertos)-transtension (passage d'Anegada). Th6se, Universit6 E et M. Curie, Paris 6, 306 pp. Jany, I., Scanlon, K.M. and Mauffret, A., 1990. Geological interpretation of combined Seabeam, Gloria and seismic data from Anegada Passage (Virgin Islands, north Caribbean). Mar. Geophys. Res., 12: 173-196. Kafka, A.L. and Weidner, D.J., 1981. Earthquake focal mechanism and tectonic processes along the southern boundary of the Caribbean plate. J. Geophys. Res., 86: 2877-2888.
668 Kellogg, J.N. and Bonini, W.E., 1982. Subduction of the Caribbean Plate and basements uplifts in the overriding South America Plate. Tectonics, 1 (3): 251-276. Kellogg, J.N. and Vega, V., 1995. Tectonic development of Panama, Costa Rica, and the Colombian Andes: constraints from global Positioning system geodetic studies and gravity. In: P. Mann (Editor), Geological and Tectonic Development of the Caribbean Plate Boundary in southern Central America. Geological Society of America, Boulder, Colo., pp. 75-90. Klaver, G.T., 1987. The Curagao lava formation: an ophiolitic analogue of the anomalous thick layer 2B of the mid-Cretaceous oceanic plateaus in the western Pacific and central Caribbean. GUA (Univ. Amsterdam) Pap. Geol., Ser. 1, 27: 128. Kolla, V., Buffler, R.T. and Ladd, J.W., 1984. Seismic stratigraphy and sedimentation of Colombian basin. Am. Assoc. Pet. Geol. Bull., 68:316-332. Ladd, J.W. and Watkins, J.S., 1978. Active marine structure within the north slope of the Muertos Trench. Geol. Mijnbouw, 57: 255-260. Ladd, J.W. and Watkins, J.S., 1980. Seismic stratigraphy of the western Venezuela basin. Mar. Geol., 35:21-41. Ladd, J.W., Shih, T.C. and Tsai, C.J., 1981. Cenozoic tectonics of central Hispaniola and adjacent Caribbean Sea. Am. Assoc. Pet. Geol. Bull., 65: 675-689. Ladd, J.W., Truchan, M., Talwani, M., Stoffa, E, Bulh, E, Houtz, R., Mauffret, A. and Westbrook, G., 1984. Seismic reflections profiles across the southern margin of the Caribbean. In: W.E. Bonini, R.B. Hargraves et al. (Editors), The Caribbean-South American Plate Boundary and Regional Tectonics. Geol. Soc. Am. Mem., 162: 153-159. Ladd, J., Holcombe, T.L., Westbrook, G.K. and Edgar, N.T., 1990. Caribbean marine geology; active margins of the plate boundary. In: G. Dengo and J.E. Case (Editors), The Geology of North America, Vol. H. The Caribbean Region. A Decade of North American Geology, Geological Society of America, Boulder, Colo., pp. 261-290. Lebron, M.C. and Perfit, M.R., 1994. Petrochemistry and tectonic significance of Cretaceous island-arc rocks, Cordillera Oriental, Dominican Republic. Tectonophysics, 229: 69-100. Lehner, E, Doust, H., Bakker, G., Allenbach, E and Gueneau, J., 1984. Active Caribbean margin of South America. In: Bally, A.W. (Editor), Seismic Expression of Structural Styles. Am. Assoc. Pet. Geol., Stud. Geol., 15:3.4.1.111-3.4.2.128. Leroy, S., 1995. Structure et origine de la plaque cara'ibe. Implications g6odynamiques. Thbse de l'Universit6 Paris 6, 240 PP. Leroy, S. and Mauffret, A., 1996. Intraplate deformation in the Caribbean region. J. Geodyn., 21: 113-122. Leroy, S., Bitri, A. and Mauffret, A., 1996. Migration velocity analysis based on common shot depth migration applied to the seismic data of the Caribbean oceanic plateau. Geophys. J. Int., 125: 199-213. Lundgren, ER. and Russo, R.M., 1996. Finite element modeling of crustal deformation in the North America-Caribbean plate boundary zone. J. Geophys. Res., 101: 11317-11327. Malav6, G. and Suarez, G., 1995. Intermediate-depth seismicity in northern Colombia and western Venezuela and its relationship to Caribbean plate subduction. Tectonics, 14 (3): 617628. Mann, P. and Lawrence, S.R., 1991. Petroleum potential of Southern Hispaniola. Jour. of Pet. Geol., 14, 291-308. Mann, P., Draper, G. and Lewis, J.E, 1991a. An overview of the geologic and tectonic development of Hispaniola. In: P. Mann, G. Draper et al. (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary in Hispaniola. Geol. Soc. Am., Spec. Pap., 262, 1-28. Mann, P., McLaughin, P.P. and Cooper, J.C., 1991b. Geology
A. M A U F F R E T and S. L E R O Y of the Azua and Enriquillon basins, Dominican Republic; 2, Structure and Tectonics. In: Mann, E, Draper, G. et al. (Editors), Geologic and Tectonic Development of the North America-Caribbean Plate Boundary in Hispaniola. Geol. Soc. Am., Spec. Pap., 262: 367-389. Mann, E, Taylor, EW., Lawrence Edwards, R. and Teh-Lung Ku, 1995. Actively evolving microplate formation by oblique collision and sideways motion along strike-slip faults: an example from the northeastern Caribbean plate margin. Tectonophysics, 246: 1-69. Marshall, J.S. and Anderson, R.S., 1995. Quaternary uplift and seismic cycle deformation, Peninsula de Nicoya, Costa Rica. Geol. Soc. Am. Bull., 107: 463-473. Mascle, A., Cazes, M. and Le Quellec, E, 1985. Structure des marges et bassins cara'fbes: une revue. In: A. Mascle (Editor), Symposium sur la g6odynamique des Cara'ibes. Technip, Paris, pp. 1-20. Masson, D.G. and Scanlon, K.M., 1991. The neotectonic setting of Puerto-Rico. Geol. Soc. Am. Bull., 103: 144-154. Mauffret, A. and Jany, I., 1990. Collision et tectonique d'expulsion le long de la fronti~re nord-Cara'fbe. Oceanol. Act., 10: 97-116. Mauffret, A. and Leroy, S., 1997. Seismic stratigraphy and structure of the Caribbean Sea. Tectonophysics (in press). Mauffret, A., Mercier de Lepinay, B., Leroy, S., Vila, J.M., Campan, A., Green, C., Gorini, C., Marton, G. and Reyes, J.R., 1994. Premiers r6sultats de la campagne Casis dans le bassin Cara'fbe. C.R. Acad. Sci. Paris, 318: 1379-1386. Mauffret, A., Hallot, E., Leroy, S., Mercier de Lepinay, B. and Vila, J.M., in prep. First results of a submersible survey of the Beata ridge. Maurasse, E, Husler, G., Georges, G., Schmitt, R. and Damond, P., 1979. Upraised Caribbean Sea floor below acoustic reflector B :t and the Southern Peninsula of Ha'fti. Geol. Mijnbouw, 8: 71-83. Meijer, P.T., 1992. The dynamics of motion of the South American plate. J. Geophys. Res., 97: 11915-11931. Mercier de Lepinay, B., 1987. Evolution de la bordure nordcara'fbe: l'exemple de la transversale d'Hispaniola. Th6se de Doctorat d'Etat, Universit6 Paris 6, 325 pp. Mercier de Lepinay, B., Mauffret, A., Jany, I., Bouysse, P., Mascle, A., Renard, V., Stephan, J.E and Hernandez, E., 1988. Une collision oblique sur la bordure nord-cara'fbe h la jonction entre la ride de Beata et la fosse de Muertos. C.R. Acad. Sci. Paris, 307:1289-1296. Mocquet, A. and Aggarwal, Y.P., 1983. Seismic slip rates in the Greater and Lesser Antilles; Implications for the present-day motion of the Caribbean plate relative to North America. Eos, 64: 832. Molnar, P. and Sykes, L., 1969. Tectonics of the Caribbean and Middle America regions from focal mechanism and seismicity. Bull. Geol. Soc. Am., 80:1639-1684. Mtiller, R.D. and Smith, W.H.E, 1993. Deformation of the oceanic crust between the North American and South American plates. J. Geophys. Res., 98: 8275-8291. Mtiller, R.D., Cande, S.C., Royer, J.Y., Roest, W.R. and Masschenkov, S., 1996. A quantitative analysis of North American and South American plate motions since chron 34 (83 Ma). In: P. Mann (Editor), Sedimentary Basins of the World, vol. 4, Caribbean Basins. Elsevier, Amsterdam. Pargo-Casas, E and Molnar, P., 1987. Relative motion of the Nazca (Farallon) and south American plates since Late Cretaceous time. Tectonics, 3 (3): 233-248. Pennington, W.D., 1981. Subduction of the Eastern Panama basin and seismotectonics of Northwestern South America. J. Geophys. Res., 86: 753-770. Pindell, J.L. and Barrett, S.E, 1990. Geological evolution of the Caribbean Region. In: G. Dengo and J.E. Case (Editors), The
N E O G E N E INTRAPLATE D E F O R M A T I O N OF THE C A R I B B E A N PLATE AT THE BEATA R I D G E Geology of North America, Vol. H. The Caribbean Region. A Decade of North American Geology, Geological Society of America, Boulder, Colo., pp. 405-432. Protti, S. and Schwartz, Y., 1994. Mechanics of back arc deformation in Costa Rica: evidence from an aftershock study of the April 22, 1991, Valle de la Estrella, Costa Rica, earthquake (M = 7.7). Tectonics, 13: 1093-1107. Pubellier, M., Mauffret, A., Leroy, S. and Vila, J., 1999. Plate boundary readjustment in oblique convergence: example of the Neogene of Hispaniola, Greater Antilles. Tectonics (in press). Rosencrantz, E. and Mann, E, 1991. SeaMarc II mapping of transform faults in the Cayman Trough, Caribbean Sea. Geology, 19: 600-693. Russo, R.M. and Villasenor, A., 1995. The 1946 Hispaniola earthquakes and the tectonics of the North AmericasCaribbean plate boundary zone northeastern Hispaniola. J. Geophys. Res., 100: 6265-6280. Russo, R.M. and Villasenor, A., 1997. Reply. J. Geophys. Res., 102: 793-802. Scientific Party, Leg 165, 1996. Deep sea cores from the Caribbean reveal history of volcanism, tectonic and oceanic changes. Eos, 77 (31): 291. Silver, E.A., Case, J.E. and MacGillavry, H.J., 1975. Geophysical study of the Venezuelan borderland. Geol. Soc. Am. Bull., 86: 213-226. Stein, S., DeMets, C., Gordon, R.G., Brodholt, J., Argus, D., Engeln, J.E, Lundgen, E, Stein, C., Wiens, D.A. and Woods, D.E, 1988. A test of alternative Caribbean Plate relative motion models. J. Geophys. Res., 93: 3041-3050. Stoffa, E, Mauffret, A., Truchan, M. and Buhl, E, 1981. Sub-B" layering in the southern Caribbean: the Aruba Gap and Venezuela basin. Earth Planet. Sci. Lett., 53: 131-146. Sykes, L.R., McCann, W.R. and Kafka, A.L., 1982. Motion of Caribbean Plate during last 7 millions years and implications for earlier Cenozoic movements. J. Geophys. Res., 87: 1065610676. Talwani, M., Ewing, J., Ewing, M. and Saito, T., 1966. Geological and geophysical studies of the submarine escarpments. Geol. Soc. Am. Spec. Pap., 101 : 217-218.
669
Talwani, M., Windisch, C.C., Stoffa, EL., Buhl, E and Houtz, R.E., 1977. Multichannel seismic study in the Venezuelan basin and Curacao Ridge. In: M. Talwani and W.C.I. Pitman (Editors), Islands Arcs, Deep Sea Trenches, and Back-Arc Basins. Am. Geophys. Union, Maurice Ewing Ser., 1: 83-98. Toto, A. and Kellogg, J.N., 1992. Structure of the Sinu-San Jacinto fold belt-- an active accretionary prism in northern Colombia. J. South Am. Earth Sci., 5, 211-222. Van der Hilst, R. and Mann, E, 1994. Tectonic implications of tomographic images of subducted lithosphere beneath northwestern South America. Geology, 22:451-454. Van Fossen, M.C. and Channell, J.E.T., 1988. Paleomagnetism of Late Cretaceous and Eocene limestones and chalks from Haiti: tectonic interpretations. Tectonics, 7:601-612. Vernette, G., Mauffret, A., Bobier, C., Briceno, L. and Gayet, J., 1992. Mud diapirism, fan sedimentation and strike-slip faulting, Caribbean Colombian Margin. Tectonophysics, 202: 335-349. Vila, J.-M., Pubellier, M., Jean-Poix, C., Feinberg, H., Butterlin, J., Boisson, D., Amilcar, H. and Amilcar, H.C., 1988. D6finition de la limite entre les blocs m6ridional et septentrional d'Hispaniola: d6couverte d'un t6moin de la nappe de Macaya dans l'anticlinal de Pierre Payen (centre d'Ha'fti, cha~ne des Matheux, Grandes Antilles); implications g6odynamiques: C.R. Acad. Sci. Paris, 307: 603-608. Vila, J.M., Jany, I., Lepvrier, C., Feinberg, H. and Mauffret, A., 1990. Mise en evidence de l'~ge post-plioc~ne inf6rieur de la collision entre la ride de Beata et l'orog~ne nord-cara'l"oe (Grandes Antilles). C.R. Acad. Sci. Paris, 311: 1359-1366. Vitali, C., 1985. Etude morphostructurale des prismes de Panama, de Colombie et du Venezuela: leurs relations avec les domaines oc6aniques et continentaux proches. Th~se de 3me cycle, Universit6 Paris 6, 225 pp. Vitali, C., Mauffret, A., Kenyon, N., Renard, V. and Belderson, B., 1985. Deformed belts off Panama and Colombia (Caribbean Sea) and plate tectonics in Panama area. In: A. Mascle (Editor), Symposium sur la G6odynamique des Caraibes. Technip, Paris, pp. 451-461.
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 -
-
-
-
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
