Special Paper 442
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THE GEOLOGICAL SOCIETY OF AMERICA®
The Terrane Puzzle: New Perspectives on Paleontology and Stratigraphy from the North American Cordillera
edited by Robert B. Blodgett and George D. Stanley Jr.
The Terrane Puzzle: New Perspectives on Paleontology and Stratigraphy from the North American Cordillera
edited by Robert B. Blodgett U.S. Geological Survey–Contractor 4200 University Drive Anchorage, Alaska 99508 USA George D. Stanley Jr. The University of Montana, Department of Geosciences 32 Campus Drive #1296 Missoula, Montana 59812 USA
Special Paper 442 3300 Penrose Place, P.O. Box 9140
Boulder, Colorado 80301-9140, USA
2008
Copyright © 2008, The Geological Society of America, Inc. (GSA). All rights reserved. GSA grants permission to individual scientists to make unlimited photocopies of one or more items from this volume for noncommercial purposes advancing science or education, including classroom use. For permission to make photocopies of any item in this volume for other noncommercial, nonprofit purposes, contact the Geological Society of America. Written permission is required from GSA for all other forms of capture or reproduction of any item in the volume including, but not limited to, all types of electronic or digital scanning or other digital or manual transformation of articles or any portion thereof, such as abstracts, into computer-readable and/or transmittable form for personal or corporate use, either noncommercial or commercial, for-profit or otherwise. Send permission requests to GSA Copyright Permissions, 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA. Copyright is not claimed on any material prepared wholly by government employees within the scope of their employment. Published by The Geological Society of America, Inc. 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA www.geosociety.org Printed in U.S.A. GSA Books Science Editors: Marion E. Bickford and Donald I. Siegel Library of Congress Cataloging-in-Publication Data The terrane puzzle : new perspectives on paleontology and stratigraphy from the North American cordillera / edited by Robert B. Blodgett, George D. Stanley, Jr. p. cm.—(Special papers (Geological Society of America) ; 442) Includes bibliographical references. ISBN 978-0-8137-2442-3 (pbk.) 1. Paleontology—North American Cordillera. 2. Paleontology, Stratigraphic. 3. Geology, Stratigraphic—North American Cordillera. 4. Paleogeography—North American Cordillera. I. Blodgett, Robert B. II. Stanley, George D. QE745.T47 2008 557—dc22
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Cover: View of the classic Upper Triassic carbonate succession of the Wrangellia terrane exposed on the east face of Green Butte, Wrangell Mountains, south-central Alaska. Lighter, more massive carbonates of the Chitistone Limestone are overlain by slightly darker, thinner beds of the Nizina Limestone. Photo taken in August 2004 by Andrew H. Caruthers.
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Contents Introduction ................................................................................................................................................v 1. Paleogeographic significance of Ediacaran cyclomedusoids within the Antelope Mountain ..........1 Quartzite, Yreka terrane, eastern Klamath Mountains, California Nan Lindsley-Griffin, John R. Griffin, and Jack D. Farmer 2. Silurian-bearing terranes of Alaska ..................................................................................................39 Constance M. Soja 3. Silurian Gastropoda from the Alexander terrane, southeast Alaska ..............................................51 David M. Rohr and Robert B. Blodgett 4. Provenance, depositional setting, and tectonic implications ...........................................................63 of Silurian polymictic conglomerates in Alaska’s Alexander terrane Constance M. Soja and Lena Krutikov 5. Devonian brachiopods of southwesternmost Laurentia: Biogeographic .......................................77 affinities and tectonic significance Arthur J. Boucot, Forrest G. Poole, Ricardo Amaya-Martínez, Anita G. Harris, Charles A. Sandberg, and William R. Page 6. Devonian brachiopods from northeastern Washington: Evidence for a ........................................99 non-allochthonous terrane and Late Devonian biogeographic update Peter E. Isaacson 7. Paleobiogeographic affinities of Emsian (late Early Devonian) gastropods ...............................107 from Farewell terrane (west-central Alaska) Jiří Frýda and Robert B. Blodgett 8. Significance of detrital zircons in Upper Devonian ocean-basin strata .......................................121 of the Sonora allochthon and Lower Permian synorogenic strata of the Mina México foredeep, central Sonora, México Forrest G. Poole, George E. Gehrels, and John H. Stewart 9. The flora, fauna, and sediments of the Mount Dall Conglomerate ..............................................133 (Farewell terrane, Alaska, USA) David Sunderlin 10. Late Triassic silicified shallow-water corals and other marine fossils .........................................151 from Wrangellia and the Alexander terrane, Alaska, and Vancouver Island, British Columbia Andrew H. Caruthers and George D. Stanley Jr.
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Contents 11. Conodont biostratigraphy and facies correlations in a Late Triassic ...........................................181 island arc, Keku Strait, southeast Alaska Erik C. Katvala and George D. Stanley Jr. 12. Stratigraphy of the Triassic Martin Bridge Formation, Wallowa terrane: ..................................227 Stratigraphy and depositional setting George D. Stanley, Jr., Christopher A. McRoberts, and Michael T. Whalen 13. Late Triassic (Carnian-Norian) mixed carbonate-volcaniclastic facies ......................................251 of the Olds Ferry terrane, eastern Oregon and western Idaho Todd A. LaMaskin 14. Early Jurassic bivalves of the Antimonio terrane (Sonora, NW México): ...................................269 Taxonomy, biogeography, and paleogeographic implications Annemarie Scholz, Martin Aberhan, and Carlos M. González-León 15. Dinosaurs of Alaska: Implications for the Cretaceous origin of Beringia ...................................313 Anthony R. Fiorillo
The Geological Society of America Special Paper 442 2008
Introduction Robert B. Blodgett U.S. Geological Survey–Contractor, 4200 University Drive, Anchorage, Alaska 99508, USA George D. Stanley Jr. University of Montana Paleontology Center, 32 Campus Drive, Missoula, Montana 59812, USA
Displaced or tectonostratigraphic terranes have been on the horizon of geology from popular books to technical papers. Terranes are now well recognized and compose a large portion of real estate in the North American Cordillera. They are discrete, fault-bounded blocks of regional “mappable” extent whose rocks and fossils differ to a great degree from those of adjacent blocks. When mapped in detail, the terranes resemble a collage of mixed rock types and tectonic styles, metamorphism, and volcanic origin—each part resembling the pieces of a puzzle. The allochthonous nature of terranes, as compared with the craton and with each other, has been suspected ever since the term was first used in this sense by Irwin (1972). Terrane studies are at the heart of discussions on the geological evolution of western North America from Mexico to the western Great Basin to the blocks, tectonic slices, and faults making up most of Alaska. Since the initiation of the terrane concept (Jones et al., 1977), terrane research continues to proliferate as a major agenda in geology. It also has spawned many new ideas in disparate areas of geology. Research continues to generate new hypotheses to test and solve terrane puzzles, and this research stimulates new directions of geological research. Paleontology and the discovery of shelly fossil remains, combined with the stratigraphic architecture of terrane successions, were the initial instruments in the emerging concepts about terranes. Paleontology and stratigraphy continue at the forefront in this effort. With the discovery in key terranes of “exotic” Permian corals and fusulinids (Stevens, 1979, 1983a, 1983b; Monger and Ross, 1971; Ross and Ross, 1983; Danner, 1997), paleontology and stratigraphy assumed an instrumental role, not only in the recognition of tectonostratigraphic terranes but also in expanding early development of basic concepts of terranes (Jones et al., 1977, 1982; Jones, 1990). Attempts were made to explain the “exotic” signature of the marine faunas found within many Permian to early Mesozoic island arc terranes of the western seaboard within the concept of long-range terrane dispersal (Ross and Ross, 1983; Tozer, 1982; Newton, 1988; Stanley and Yancey, 1990). These papers focused on long-distance travel versus near-craton origin for tropical to subtropical terranes, including Cache Creek, Quesnelia, Wrangellia, and the Wallowa and Stikine terranes. Fossils, especially tropical reef-type taxa, appear useful in assessing terrane movement and reconstructing the paleogeography and paleontology of tectonic elements that coalesced early on (Stanley, 1996). Paleontology has much to contribute in this arena of geology. Fossils occur widely distributed within complex terrane collages of North America, extending from northern Alaska southward. In the Southern Hemisphere, Upper Triassic reef faunas are known from rocks in the Andes of Chile and Peru (Stanley, 1994) and, together with Alaskan occurrences, produce a range in latitude of ~74°. Even lower latitude Southern Hemisphere occurrences are known. Anomalous tropical shallow-water corals, mollusks, crinoids, and other shelly biotas of early Mesozoic age
*E-mails:
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[email protected] Blodgett, R.B., and Stanley, G.D., Jr., 2008, Introduction, in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. v–viii, doi: 10.1130/2008.442(00). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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are known as far south as the Antarctic Peninsula, and this occurrence most certainly is tectonically displaced. Many North American terranes clearly enjoyed enormous mobility relative to the more fixed craton, and models have been presented for their evolution, fragmentation, displacement, and accretion; superterranes have evolved from the amalgamation of originally separate entities (Monger and Nokleberg, 1996). Paleomagnetic evidence (Irving and Wynne, 1991) has helped constrain paleolatitude and aided reconstruction of the paleogeography of various pieces of the North American Cordilleran, currently part of the crazy-quilt pattern comprising the terrane puzzle. Monger (1997) addressed the role of paleontology in resolving conflicts encountered in paleomagnetic results regarding displacements in the Cordillera during the Carboniferous to Jurassic history of terranes, especially with discussions on “exotic Tethyan” versus cratonal origins of paleofaunas in the Canadian Cordillera (Danner, 1997). Paleobiogeographic studies of the Cordilleran terrane faunas and rocks have increased in sophistication with the introduction of more precise taxonomic assessments and quantitative approaches (Aberhan, 1998, 1999; Belasky and Runnegar, 1993, 1994; Blodgett, 1998; Blodgett and Boucot, 1999; Blodgett and Frýda, 2001; Blodgett et al., 2002; Frýda and Blodgett, 2001, 2004; Hoover, 1991; Newton, 1983, 1987; Ross and Ross, 1990; Smith and Tipper, 1986, Soja, 1996; Soja and Antoshkina, 1997; Stevens et al., 1990; Yarnell, 2000). Results of studies employing lithostratigraphy, paleomagnetism, timing of magmatic and tectonic episodes, and other aspects of geology, including stratigraphic approaches utilizing zircon and Nd isotope provenance, have also been brought to bear on terrane problems. Although originally conceived as being far-traveled, relative to the craton of North America, most terranes in the eastern part of the ancient Pacific are now considered to represent island arcs or fragments of rifted continental margins, with some terranes being parautochthonous along strike-slip faults relative to the craton. One interesting aspect of paleobiogeographic studies of fossil Paleozoic faunas from Alaskan terranes during the past decade is that most of them appear to be of Eurasian origin, related to the Siberian craton or peri-cratonic Siberia terranes or the Urals (Blodgett, 1998; Blodgett and Boucot, 1999; Blodgett et al., 2002, 2003; Frýda and Blodgett, 2004, and this volume; Garcia-Alcalde and Blodgett, 2001; Rohr and Blodgett, this volume). Today, more than ever, the study of fossils and strata is particularly relevant to many geologic investigations of rocks and strata extending from Alaska through most of western Canada, western conterminous United States, and regions of Mexico. Only a few general synthesis volumes have been devoted to the history of North American Cordilleran terranes with a focus on fossils, stratigraphic settings, and the integration of tectonics and terrane accretion in North America. Unfortunately, most of these are quite dated (Stevens, 1983a, 1983b; Hashimoto and Uyeda, 1983; Howell, 1985). Since the 1980s, there has been no comprehensive volume or book treating the paleontologicalstratigraphic aspects of the terrane issue, and a volume dedicated
to paleontology and stratigraphy of terranes is long overdue. We are thus pleased in this GSA Special Paper to address a variety of relevant topics on displaced terranes of the Cordillera. The original concept was spawned during a special two-day thematic session titled “New insights from paleontology, stratigraphy, and sedimentology on accreted terranes of western North America.” This session, held in Puerto Vallarta, Mexico, was organized by Robert B. Blodgett and George D. Stanley Jr. for the Geological Society of America Cordilleran section meeting in April 2003. Some of the original participants in that session also present papers in this GSA Special Paper. In May 2006, Robert B. Blodgett and Eric Katvala organized a similar, two-day thematic session entitled “Accreted terranes of western North America: An update on current research on the construction of the Cordillera” at the combined meeting of the Geological Society of America Cordilleran Section, the Pacific Section of the American Association of Petrologists, and the Alaska Section of the Society of Petroleum Engineers in Anchorage, Alaska. Many of the same participants from the Puerto Vallarta conference also spoke or were in attendance at the Anchorage conference. The 15 papers by 28 authors in this GSA volume focus on Cordilleran terranes in general and, in particular, on aspects of stratigraphy and paleontology. Fossils covered range from plants to dinosaurs, conodonts, mollusks, brachiopods, corals, and jellyfish, with sedimentary rocks of Precambrian to Cretaceous age. Techniques vary from paleobiogeographic analyses and sequence stratigraphy to provenance studies using zircons and conglomerates. The papers in this volume are arranged by the age of strata and fossils addressed by the contributing authors. The first paper, by Linsley-Griffin, Griffin, and Farmer, discusses the significance of their recent discovery of Ediacaran cyclomedusids from the Antelope Mountain Quartzite of the Yreka terrane, which is situated in the eastern Klamath Mountains of northern California. The next paper, by C.M. Soja, summarizes current knowledge regarding Silurian strata and fossil fauna from Alaskan terranes. The paper by Rohr and Blodgett documents and discusses the paleogeographic implications of a richly diverse Silurian gastropod fauna found in the Heceta Limestone on Prince of Wales Island, southeast Alaska, in strata belonging to the Alexander terrane. The Heceta Limestone is also the focus of the paper by Soja and Krutikov, which approaches the question of paleogeography, stratigraphy, and setting by using provenance of clasts within conglomerates of this formation. The paper by Boucot, Poole, Amaya-Martínez, Harris, Sandberg, and Page describes three separate brachiopod faunas from the Devonian of Sonora, Mexico. Their finds are exciting, especially in light of our previously poor knowledge of this interval from southwestern Laurentia. The paper by P.E. Isaacson examines the Frasnian brachiopod fauna from Limestone Hill, northeastern Washington State. On the basis of his study, he concludes that the area is parautochthonous with respect to North America and shows strong affinities with craton-bounded fauna known from Idaho, Montana, Utah, and Nevada. The paper by Frýda and Blodgett addresses the paleobiogeographic affinities of Emsian
Introduction
(late Early Devonian) gastropods recovered from Limestone Mountain in the Farewell terrane of west-central Alaska. This study reaffirms the close affinities recognized with Eurasian and Australian Emsian gastropods, as well with those known from the Alexander terrane, rather than with faunas known from cratonic North America. The paper by Poole, Gehrels, and Stewart examines the significance of detrital zircons recovered from Upper Devonian and Lower Permian strata in Sonora, Mexico. The detrital zircon geochronology indicates that most of the detritus in both samples was derived from Laurentia to the north, whereas some detritus in the Permian synorogenic foredeep sequence was derived from an evolving accretionary wedge to the south. The paper by D. Sunderlin addresses the Early Permian flora, fauna, and sediments of the Mount Dall Conglomerate of south-central Alaska, belonging to the Farewell terrane. The paleofloral data indicate a mixed phytogeographic affinity of both Angaran and Euramerican provinces and support the placement of this terrane within a midlatitude climate belt during the Early Permian. The Triassic represents the most highly studied time interval in this volume, with a total of four papers devoted entirely to rocks of this age. Upper Triassic silicified corals are the focus of the paper by Caruthers and Stanley, which is based on the study of corals and associated fauna from the Wrangellia and Alexander terranes of Alaska and Vancouver Island, British Columbia. The paper by Katvala and Stanley helps refine Upper Triassic biostratigraphy and lithostratigraphy within the Alexander terrane, using new data from conodonts retrieved from rocks exposed in the Keku Strait of southeast Alaska. The next two Triassic contributions in the volume deal with the Blue Mountains Province of northeastern Oregon and adjacent Idaho, where a number of volcanic terranes are well recognized. The paper by Stanley, McRoberts, and Whalen treats the Upper Triassic stratigraphy of the Wallowa terrane. These authors present a new regional stratigraphic framework and depositional interpretations for the shallow- to deeper-water settings of the Martin Bridge Formation. The paper by T. LaMaskin describes the sedimentology of Upper Triassic strata in the Olds Ferry terrane and presents a model for arc-flanking mixed carbonate-volcaniclastic sedimentation. The two final papers in the volume cover post-Triassic subjects. The first of these papers concerns Early Jurassic bivalves from the Antimonio terrane in Mexico. Fleischer, Aberhan, and González-León analyze diverse assemblages of these fossils to arrive at conclusions regarding the ancient position of the Sonora block relative to other well-known terranes. Dinosaurs were, until recent time, thought to be uncommon in Alaska. The final paper by Fiorillo explores the record of Cretaceous dinosaurs of northern Alaska in depth and proposes extending the concept of Beringia back in time to recognize its existence during the Cretaceous. The contributions span a time interval of late Precambrian (Vendian) to Cretaceous and address a large number of the Cordilleran terranes distributed from Mexico to Alaska. Together, the authors in this volume provide an intellectually stimulating look at some key terranes, extracting relevant paleogeographic
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issues and developing and testing tectonic and paleogeographic hypotheses. They demonstrate the utility of paleontologic and stratigraphic approaches and demonstrate how these approaches interface with tectonic and geophysical problems. Although the paradigm of Cordilleran geology is far from complete, the papers in this volume go far in elucidating some key pieces of the greater terrane puzzle. We hope this volume serves as a valuable resource to those interested in the terrane issues and the diversity of approaches utilizing fossils and stratigraphy. REFERENCES CITED Aberhan, M., 1998, Paleobiogeographic patterns of pectinoid bivalves and the Early Jurassic tectonic evolution of western Canadian terranes: Palaios, v. 13, p. 129–148, doi: 10.2307/3515485. Aberhan, M., 1999, Terrane history of the Canadian Cordillera: estimating amounts of latitudinal displacement and rotation of Wrangellia and Stikinia: Geological Magazine, v. 136, p. 481–492, doi: 10.1017/ S001675689900299X. Belasky, P., and Runnegar, B., 1993, Biogeographic constraints for tectonic reconstruction of the Pacific region: Geology, v. 21, p. 979–982, doi: 10. 1130/0091-7613(1993)021<0979:BCFTRO>2.3.CO;2. Belasky, P., and Runnegar, B., 1994, Permian longitudes of Wrangellia, Stikinia, and Eastern Klamath terranes based on coral biogeography: Geology, v. 22, p. 1095–1098, doi: 10.1130/0091-7613(1994)022<1095:PLOWSA> 2.3.CO;2. Blodgett, R.B., 1998, Emsian (Late Early Devonian) fossils indicate a Siberian origin for the Farewell terrane: Short Notes on Alaskan Geology 1997: Alaska Division of Geological and Geophysical Surveys Professional Report, v. 118, p. 27–34. Blodgett, R.B., and Boucot, A.J., 1999, Late Early Devonian (late Emsian) eospiriferinid brachiopods from Shellabarger Pass, Talkeetna C-6 quadrangle, south-central Alaska and their biogeographic importance: Further evidence for a Siberian origin of the Farewell and allied Alaskan accreted terranes: Senckenbergiana Lethaea, v. 72, no. 1, p. 209–221. Blodgett, R.B., and Frýda, J., 2001, On the occurrence of Spinidelphinulopsis whaleni [Late Triassic (early Norian) Gastropoda] in the Cornwallis Limestone, Kuiu Island, southeastern Alaska (Alexander terrane) and its paleobiogeographic significance: Bulletin of the Czech Geological Survey, v. 76, no. 4, p. 267–274. Blodgett, R.B., Rohr, D.M., and Boucot, A.J., 2002, Paleozoic links among some Alaskan accreted terranes and Siberia based on megafossils, in Miller, E.L., Grantz, A., and Klemperer, S.L., eds., Tectonic evolution of the Bering Shelf-Chukchi Sea-Arctic margin and adjacent landmasses: Geological Society of America Special Paper, v. 360, p. 273–290. Blodgett, R.B., Rohr, D.M., Karl, S.M., and Baichtal, J.F., 2003, Early Middle Devonian (Eifelian) gastropods from the Wadleigh Limestone in the Alexander terrane of southeastern Alaska demonstrate biogeographic affinities with central Alaska terranes (Farewell and Livengood) and Eurasia, in Galloway, J.P., ed., Studies in Alaska by the U.S. Geological Survey, 2001: U.S. Geological Survey Professional Paper 1678, p. 105–115. Danner, W.R., 1997, Fusulinids and other Paleozoic Foraminifera of accreted terranes, southwestern British Columbia and northwestern Washington, in Ross, C.A., Ross, J.R.P., and Brenckle, P.L., eds., Late Paleozoic Foraminifera: Their biostratigraphy, evolution, and paleoecology, and the mid-Carboniferous boundary. Cushman Foundation for Foraminiferal Research Special Publication, v. 36, p. 21–25. Frýda, J., and Blodgett, R.B., 2001, Chulitnacula, a new paleobiogeographically distinctive gastropod genus from Upper Triassic strata in accreted terranes of southern Alaska: Journal of Czech Geological Society, v. 46, no. 3/4, p. 299–306. Frýda, J., and Blodgett, R.B., 2004, New Emsian (Late Early Devonian) gastropods from Limestone Mountain, Medfra B-4 quadrangle, west-central Alaska (Farewell terrane), and their paleobiogeographic affinities and evolutionary significance: Journal of Paleontology, v. 78, no. 1, p. 111–132, doi: 10.1666/0022-3360(2004)078<0111:NELEDG>2.0.CO;2. Garcia-Alcalde, J., and Blodgett, R.B., 2001, New Lower Devonian (Upper Emsian) Myriospirifer (Brachiopoda, Eospiriferinae) species from Alaska
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and northern Spain and the paleogeographic distribution of the genus Myriospirifer: Journal of the Czech Geological Society, v. 46, no. 3/4, p. 145–154. Hashimoto, M., and Uyeda, S., 1983, editors, Accretion tectonics in the circumPacific region: Proceedings of the Oji International Seminar on Accretion Tectonics, Japan, 1981: Tokyo, Terra Scientific Publishing Co., Advances in Earth and Planetary Sciences, v. 15, 358 p. Hoover, P.R., 1991, Late Triassic cyrtinoid spiriferinacean brachiopods from western North America and their biostratigraphic and biogeographic implications: Bulletins of American Paleontology, v. 100, p. 63–109. Howell, D.G., ed., 1985, Tectonostratigraphic terranes of the circum-Pacific region: Houston, Circum-Pacific Council for Energy and Mineral Resources Earth Science Series, no. 1, 581 p. Irving, E., and Wynne, P.J., 1991, Paleomagnetic evidence for motions of parts of the Canadian Cordillera: Tectonophysics, v. 187, p. 259–275, doi: 10.1016/0040-1951(91)90423-P. Irwin, W.P., 1972, Terranes of the western Paleozoic and Triassic belt in the southern Klamath Mountains, California: U.S. Geological Survey Professional Paper 800-C, p. C103–C111. Jones, D.L., 1990, Synopsis of late Paleozoic and Mesozoic terrane accretion within the Cordillera of western North America: Philosophical Transactions of the Royal Society of London, v. A(331), p. 479–486. Jones, D.L., Silberling, N.J., and Hillhouse, J., 1977, Wrangellia—A displaced terrane in western North America: Canadian Journal of Earth Sciences, v. 14, p. 2565–2577. Jones, D.L., Cox, A., Coney, P., and Beck, M., 1982, The growth of western North America: Scientific American, v. 247, p. 70–84. Monger, J.W.H., 1997, Plate tectonics and Northern Cordilleran geology: An unfinished revolution: Geoscience Canada, v. 24, no. 4, p. 189–198. Monger, J.W.H., and Nokleberg, W.H., 1996, Evolution of the northern North American Cordillera: Generation, fragmentation, displacement and accretion of successive North American plate margin arcs in Coyner, A.R., and Fahey, P.L., eds., Geology and ore deposits of the American Cordillera: Proceedings, Geological Society of Nevada Symposium, Reno/Sparks, Nevada, April 1995, v. 3, p. 1133–1152. Monger, J.W.H., and Ross, C.A., 1971, Distribution of fusulinaceans in the western Canadian Cordillera: Canadian Journal of Earth Sciences, v. 8, p. 259–278. Newton, C.R., 1983, Paleozoogeographic affinities of Norian bivalves from the Wrangellian, Peninsular, and Alexander terranes, western North America, in Stevens, C.H., ed., Pre-Jurassic rocks in Western North American suspect terranes: Los Angeles, Pacific Section, SEPM, p. 37–48. Newton, C.R., 1987, Biogeographic complexity in Triassic bivalves of the Wallowa terrane, northwest United States: Oceanic islands, not continents, provide the best analogues: Geology, v. 15, p. 1126–1129, doi: 10. 1130/0091-7613(1987)15<1126:BCITBO>2.0.CO;2.
Newton, C.R., 1988, Significance of Tethyan fossils in the Cordillera: Science, v. 242, p. 385–391, doi: 10.1126/science.242.4877.385. Ross, C.A., and Ross, J.R.P., 1990. Late Palaeozoic bryozoan biogeography, in McKerrow, W.S., and Scotese, C.R., eds., Palaeozoic palaeogeography and biogeography: Geological Society [London] Memoir 12, p. 353–362. Ross, J.R.P., and Ross, C.A., 1983, Late Paleozoic accreted terranes of western North America, in Stevens, C.H., ed., Pre-Jurassic rocks in western North American suspect terranes: Los Angeles, Pacific Section, SEPM, p. 7–22. Smith, P.L., and Tipper, H.W., 1986, Plate tectonics and paleobiogeography: Early Jurassic (Pliensbachian) endemism and diversity: Palaios, v. 1, p. 399–412, doi: 10.2307/3514477. Soja, C.M., 1996, Island-arc carbonates: Characterization and recognition in the ancient geologic record: Earth-Science Reviews, v. 41, p. 31–65, doi: 10.1016/0012-8252(96)00029-3. Soja, C.M., and Antoshkina, A.I., 1997, Coeval development of Silurian stromatolite reefs in Alaska and the Ural Mountains: Implications for paleogeography of the Alexander terrane: Geology, v. 25, p. 539–542, doi: 10. 1130/0091-7613(1997)025<0539:CDOSSR>2.3.CO;2. Stanley, G.D., Jr., 1994, Late Paleozoic and early Mesozoic reef-building organisms and paleogeography: The Tethyan-North American connection: Courier Forschungsinstitut Senckenberg, v. 172, p. 69–75. Stanley, G.D., Jr., 1996, Confessions of a displaced reefer: Palaios, v. 11, no. 1, p. 1–2, doi: 10.2307/3515111. Stanley, G.D., Jr., and Yancey, T.E., 1990, Paleogeography of the ancient Pacific: Science, v. 249, p. 680–681, doi: 10.1126/science.249.4969.680-a. Stevens, C.H., 1979, Reconstruction of Permian paleogeography based on distribution of Tethyan faunal elements: Congrès International de Stratigraphie et de Géologie du Carbonifière, 9th, v. 5, p. 383–394. Stevens, C.H., 1983a, Corals from a dismembered late Paleozoic paleo-Pacific plateau: Geology, v. 11, p. 603–606, doi: 10.1130/0091-7613(1983)11 <603:CFADLP>2.0.CO;2. Stevens, C.H., 1983b, editor, Pre-Jurassic rocks in Western North America Suspect Terranes: Los Angeles, Pacific Section, SEPM, 141 p. Stevens, C.H., Yancey, T.E., and Hanger, R.A., 1990, Significance of the provincial signature of Early Permian faunas of the eastern Klamath terrane: Geological Society of America Special Paper, v. 25, p. 201–218. Tozer, E.T., 1982, Marine Triassic faunas of North America: Their significance for assessing plate and terrane movements: Geologische Rundschau, v. 71, p. 1077–1104, doi: 10.1007/BF01821119. Yarnell, J.M., 2000, Paleontology of two North American Triassic reef faunas: Implications for terrane paleogeography [M.S. thesis]: Missoula, University of Montana, 141 p. MANUSCRIPT ACCEPTED BY THE SOCIETY 14 DECEMBER 2007
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The Geological Society of America Special Paper 442 2008
Paleogeographic significance of Ediacaran cyclomedusoids within the Antelope Mountain Quartzite, Yreka subterrane, eastern Klamath Mountains, California Nancy Lindsley-Griffin John R. Griffin Department of Geosciences, 214 Bessey Hall, University of Nebraska, Lincoln, Nebraska 68588-0340, USA Jack D. Farmer Department of Geological Sciences, Arizona State University, Tempe, Arizona 85287-1404, USA
ABSTRACT Newly recognized cyclomedusoid fossils in the Antelope Mountain Quartzite confirm that it is latest Neoproterozoic (Ediacaran) in age. Biogeographic affinities of the cyclomedusoid fossils suggest that the Yreka subterrane and its close associate, the Trinity subterrane, formed after the breakup of Rodinia in an ocean basin bordering Australia, northern Canada, Siberia, and Baltica. Reevaluating biogeographic, geological, and paleomagnetic evidence in the context of this starting point, the Yreka subterrane and Trinity subterrane may have been located at either 7°N or 7°S latitude ca. 580–570 Ma, but were not necessarily close to Laurentia. Continental detrital zircons (3.2–1.3 Ga) in the Antelope Mountain Quartzite most likely came from Australia or Siberia rather than Laurentia. The Yreka subterrane and Trinity subterrane record ~180 m.y. of active margin events somewhere in Panthalassa (Proto-Pacific Ocean). Paleozoic biogeographic data, paleomagnetism, and regional relationships indicate that Yreka subterrane and Trinity subterrane were located throughout the early Paleozoic in the part of Panthalassa surrounded by Australia, NW Laurentia, Siberia, China, Baltica, and the Uralian terranes. By the mid-Devonian they were located at 31°N or 31°S in a somewhat isolated location, probably in a Northern Hemisphere oceanic plateau or island chain well outboard of other tectonic elements, and by the Permian they were almost completely isolated from other tectonic elements. The Yreka subterrane, as part of the Klamath superterrane, was not native to North America and did not accrete to it until the Early Cretaceous. Keywords: Ediacaran cyclomedusoids, Antelope Mountain Quartzite, Yreka subterrane, Trinity subterrane, Klamath Mountains.
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[email protected] Lindsley-Griffin, N., Griffin, J.R., and Farmer, J.D., 2008, Paleogeographic significance of Ediacaran cyclomedusoids within the Antelope Mountain Quartzite, Yreka subterrane, eastern Klamath Mountains, California, in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 1–37, doi: 10.1130/2008.442(01). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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INTRODUCTION Most folded mountain belts comprise a collage of tectonostratigraphic terranes developed on different types of basement and assembled by tectonic plate collisions (Helwig, 1974; Coney et al., 1980; Ben-Avraham et al., 1981). In this paper, following the usage recommended by Dover (1990), we use terrane descriptively, to denote a fault-bounded geologic region that differs from adjoining geologic regions by its distinctive stratigraphy, structure, tectonic history, and in some cases biota. We use subterrane for small fault-bounded geologic regions distinct from each other that combine with other geologic regions to form a recognized terrane. Terranes may be either allochthonous or autochthonous, far-traveled or near their point of origin. Terranes may amalgamate together to form a superterrane, and terranes become accreted terranes after they collide with continental crust along an active plate margin. The origin of each accreted terrane is suggested by its combination of basement type, lithology, structure, paleontology, paleomagnetic poles, and regional relationships. However, the distance traveled is far more difficult to determine. Paleomagnetic poles reveal latitude, but not longitude or hemisphere. Thus, we must combine paleomagnetic results with other geologic information to track a terrane’s travels. We can use biotic assemblages that contain provincial species to determine shared basins, but different biotic assemblages do not indicate distance, only barriers to migration (Valentine, 1972; Valentine and Moores, 1970, 1972). Likewise, similarity between biotas does not denote closeness in terms of distance, but merely indicates migration was possible because the biotas shared a basin, or occupied several basins connected by a migration path. In this paper we describe the stratigraphic and structural setting of newly recognized Ediacaran1 fossils from the Yreka subterrane (Fig. 1) and of the Antelope Mountain Quartzite (Fig. 2), which contains them, and we clarify contact relationships and tectonic history of the formation. We formally define the Schulmeyer Gulch Complex, a polygenetic mélange underlying the Antelope Mountain Formation (Fig. 2). Correlating Ediacaran and Paleozoic faunal data with regional geologic relationships and high-quality paleomagnetic data (Mankinen et al., 1989; Irwin and Mankinen, 1998; Mankinen et al., 2002), we place the Klamath superterrane into the context of current paleogeographic reconstructions, beginning with the late Proterozoic breakup of Rodinia and following it through the Middle Devonian paleolatitude obtained from overlapping strata (Mankinen et al., 2002) to the development of the unique Permian McCloud biota (Stevens et al., 1990). REGIONAL SETTING The Klamath superterrane (Fig. 1) of northern California and southern Oregon includes, from west to east, the Western
Figure 1. Principal elements of the Klamath superterrane. Shaded areas comprise the Eastern Klamath terrane, which includes the Yreka, Trinity, and Redding subterranes and the herein-defined Forest Mountain (FM) subterrane. Dots indicate location of Figure 2. After LindsleyGriffin et al. (2006).
Klamath, Rattlesnake Creek, Western Hayfork, Eastern Hayfork, North Fork, Fort Jones, and Central Metamorphic terranes, as well as the four subterranes that constitute the Eastern Klamath terrane. The Eastern Klamath terrane includes the Yreka, Trinity, and Redding subterranes (Fig. 1) and the Forest Mountain subterrane, which we define in this paper. Terranes west of the Eastern Klamath terrane (Fig. 1) consist of westward-younging arcuate belts formed on oceanic crust that amalgamated together in a series of collisions with the Eastern Klamath nucleus from the earliest Paleozoic to the Mesozoic (e.g., Irwin, 1972, 1981, 1985; Irwin and Mankinen, 1998, Lindsley-Griffin et al., 2006). However, the Klamath superterrane accreted to the North American continent only after 144 Ma, based on cessation of rotation from paleomagnetic data (Mankinen et al., 1989), or in the Early Cretaceous, based on the first development of the Great Valley successor basin over the southern part of the superterrane ca. 136 Ma (Mankinen et al., 1989; Irwin, 2003). The post-accretion Hornbrook successor basin overlying the northern part of the Klamath superterrane began developing in middle Cretaceous
Following the latest recommendations of the International Commission on Stratigraphy, we use “Ediacaran” for the latest period of the Neoproterozoic, formerly known as Vendian; numerical ages are from the International Geological Time Scale (Gradstein et al., 2004).
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Figure 2. Geologic map and cross section of the northern part of the Yreka subterrane and the adjoining Forest Mountain, Central Metamorphic, and Fort Jones terranes. Scale on map given by latitude-longitude; scale on cross section given by bar scales. CG—Cram Gulch; PC— Pythian Cave; Sch G—Schulmeyer Gulch; SG—Skookum Gulch. Map units: Yreka subterrane: am—Antelope Mountain Quartzite (shaded); Dd—Duzel Phyllite; Ds—Sissel Gulch Graywacke; Ofd—mélange of Facey-Duzel Rock; Om—Schulmeyer Gulch Complex mélange (light pattern); Oms—Schulmeyer Gulch semischist; OSDm—Gregg Ranch Complex; OSm—Skookum Gulch mélange; Smc—Moffett Creek Formation. Forest Mountain subterrane: Da—amphibolite; Dm—Devonian mélange; FMu—undivided Forest Mountain subterrane. Fort Jones terrane: TrPzg—Triassic-Paleozoic greenstone-chert assemblage; TrPzs—Triassic-Paleozoic Stuart Fork Formation. Other: ∇—blueschist block; Ki—Cretaceous intrusion; Kh—Hornbrook Formation; Tv—Tertiary volcanic rocks; Qa—Quaternary alluvium; Qas—Quaternary alluvium of Shasta Valley. Adapted from Hotz (1977; 1978); Rohr (1978); and Lindsley-Griffin and Griffin (1983).
(Nilsen, 1984; Lindsley-Griffin et al., 1993), suggesting a diachronous accretion proceeding from south to north. This paper combines paleontological, geological, and paleomagnetic data to address the question of where the Klamath superterrane was before it accreted to North America.
Klamath Mountain Terrane Terminology The names applied to terranes of the Klamath Mountains have evolved through time, in some cases with different authors applying the same name to different terranes or applying multiple
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names to the same terrane. Irwin (1960, 1966) initially defined four geologic subprovinces within the Klamath Mountains province, breaking out “plates” based on distinct geologic assemblages separated by major faults, much as we define tectonostratigraphic terranes today. His prototype subprovinces included the Eastern Klamath Belt, Central Metamorphic Belt, Western Paleozoic and Triassic Belt, and Western Jurassic Belt. The “eastern Klamath plate” (Irwin, 1966), “eastern Klamath belt” (Irwin, 1977), or “Eastern Klamath terrane” (Irwin, 1985; Mankinen et al., 1984, 1989) encompassed all or part of the four subterranes discussed herein. Irwin (1977) recognized three Eastern Klamath subterranes: the “Yreka-Callahan strata,” the “Redding section,” and the “Trinity ultramafic sheet,” which separates them. In their seminal work on Cordilleran suspect terranes, Coney et al. (1980) adopted two of Irwin’s subprovince names, breaking the Klamath superterrane into just three terranes: the eastern Klamath Mountains, the Triassic and Paleozoic Klamath Mountains, and the Foothills Klamath Mountains. Irwin’s “Yreka-Callahan strata” were termed the “Yreka terrane” by Silberling et al. (1984), and the “Trinity ultramafic sheet” became the Trinity terrane (e.g., Silberling et al., 1984; Lindsley-Griffin, 1991, 1994). However, disregarding the precedent set by Irwin’s pioneering work, some authors applied the “Eastern Klamath” name to the Redding section only (e.g., Silberling et al., 1984; Miller and Harwood, 1990; Potter et al., 1990a). In this paper, we follow the usage of Irwin (2003) in treating the Redding subterrane, Trinity subterrane, and Yreka subterrane (Fig. 1) as separate subterranes of the Eastern Klamath terrane. To these we add the newly recognized Forest Mountain subterrane, which lies west of the Yreka subterrane (Fig. 1). Terranes of the Klamath Mountains The Eastern Klamath terrane acted as a nucleus against which the more westerly terranes of the Klamath superterrane (Fig. 1) successively collided, beginning in the Middle Devonian and continuing in the early to middle Mesozoic (Irwin and Mankinen, 1998; Irwin, 2003). As each collision added another terrane to the western edge of the nucleus, causing it to grow progressively larger through time, the Redding subterrane developed over the eastern edge of the nucleus from Middle Devonian to Early Jurassic time. From east to west, components of the Klamath superterrane are as follows: 1. Eastern Klamath terrane, discussed in more detail in the following section; 2. Central Metamorphic terrane, consisting of metavolcanic Salmon Hornblende Schist overlain by metasedimentary Grouse Ridge Formation, which collided with the southern part of the Eastern Klamath terrane along the Bully Choop Fault in Middle Devonian, ca. 380 Ma (Davis et al., 1965; Irwin, 1981; Irwin and Mankinen, 1998); 3. Fort Jones terrane (also termed Stuart Fork terrane), consisting of phyllitic quartzites and blueschist-metabasalt-metachert mélange, which collided along unnamed
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5.
6.
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pre-Siskiyou faults in the Triassic (Hotz et al., 1977; Goodge, 1990; Irwin and Mankinen, 1998); North Fork terrane, a subduction complex consisting of Permian–Lower Jurassic ophiolite overlain by metamorphosed mélange of basalt, argillite, limestone, and Lower Jurassic radiolarian chert, which collided along the Siskiyou Fault in the Early Jurassic (Irwin, 1972; Ando et al., 1983; Blome and Irwin, 1983; Irwin and Mankinen, 1998); Eastern Hayfork terrane, a Permo-Triassic argillite mélange containing mafic volcanic and sedimentary rocks, and exotic Upper Permian–Triassic limestone and chert, which collided along the Twin Sisters Fault in the early (?) Middle Jurassic (Irwin, 1972; Wright, 1982; Irwin and Mankinen, 1998); Western Hayfork terrane, a volcanic arc consisting of interbedded metavolcanic (Hayfork Bally Meta-andesite), volcaniclastic, and argillaceous rocks, plus related plutons, which collided along the Wilson Point Fault in the late Middle Jurassic (Irwin, 1972; Wright, 1982; Wright and Fahan, 1988; Irwin and Mankinen, 1998); Rattlesnake Creek terrane, whose high-grade metamorphic equivalent is the “Marble Mountain terrane,” consisting of Triassic–Lower Jurassic mélange containing blocks of ophiolitic rocks (serpentinized harzburgite, amphibolite, metabasalt, metachert), upper Paleozoic– Middle Jurassic limestone, Upper Triassic–Middle Jurassic chert, and Lower Jurassic plutonic rocks, which collided along the Salt Creek Fault in late Middle to middle Late Jurassic time (Irwin, 1972; Blake et al., 1985; Gray, 1986; Donato, 1987; Irwin and Mankinen, 1998); Western Klamath terrane, consisting of the Josephine ophiolite (backarc basin lithosphere) overlain by Galice Formation (forearc basin sediments) and Rogue Formation (volcanic arc), which collided along the Bear Wallow Fault in the Late Jurassic (Harper, 1980; Pinto-Auso and Harper, 1985; Irwin and Mankinen, 1998).
Following accretion of the Klamath superterrane to North America, the Pickett Peak terrane, now part of the Coast Ranges province, collided with its western edge along the South Fork Fault in middle Early Cretaceous time (Irwin and Mankinen, 1998). THE EASTERN KLAMATH TERRANE The cyclomedusoid-bearing Antelope Mountain Quartzite is part of the Yreka subterrane (Figs. 1 and 2), an accretionary complex that structurally overlies the ophiolitic Trinity subterrane. The Yreka and Trinity subterranes comprise rocks of the same ages (Table 1) and have been closely associated in both time and space since the late Neoproterozoic (Lindsley-Griffin et al., 2006). The Yreka subterrane consists of stacked thrust sheets of marine sedimentary and metasedimentary rocks ranging in age from Ediacaran through Early Devonian. Each of these thrust
Rock unit
TABLE 1. SELECTED ISOTOPIC DATA AND AGES FOR EASTERN KLAMATH AND RELATED TERRANES Isotopic age(s) Method (Ma)
A. Precambrian Detrital Zircons In Antelope Mountain Quartzite, YT Archean 3200–3130; 2980–2900; 2750; 2680; 2530 Paleoproterozoic 2450, 2380; 2150–2300; 2000–2008; 1900–1800 Mesoproterozoic 1300 B. Inherited Mesoproterozoic component in Skookum Gulch mélange, YT Trail Gulch monzodiorite block C. Ediacaran rifting In Trinity ophiolite, TT Trinity metagabbro, plagiogranite Trinity basalt
1
U-Pb zircon 1 U-Pb zircon 1 U-Pb zircon 207
206
Pb/ Pb zircon
1033
2
3, 4
556–579 ca. 565–570
U-Pb zircon 5 paleomagnetism
565 r 5; 564 r 7 567 r 78; 566–572
U-Pb zircon 4 U-Pb zircon
E. Early Ordovician plutonism, TT-YT Intrusions, Trinity ophiolite Uplift and partial melting, Trinity peridotite
470–480 472 ± 32
U-Pb zircon 6 Sm-Nd isochron
F. Late Ordovician metamorphism and plutonism, YT Skookum Gulch protolith Skookum Gulch metamorphism
467 ± 45 447 ± 9
Rb-Sr WhR 7 Rb-Sr phengite
ca. 435–420 ca. 440–420, 410 ca. 420–410 ca. 430–400
U-Pb zircon 1 U-Pb zircon 1 U-Pb zircon 1 U-Pb zircon
435 ± 21 431 ± 3 415 ± 3; 412 ± 10
Sm-Nd isochron 9 U-Pb zircon 1, 4 U-Pb zircon
407–398 399 r 12; 391 r 12 392 ± 11; 371 ± 9
K-Ar hornblende 8 K/Ar hornblende 40 39 8 Ar/ Ar WhR
404 ± 3; ca. 398 400 ca. 398–385 Emsian–pre-398
U-Pb zircon 11 U-Pb zircon 5 paleomagnetism 12 fossil fish plate
ca. 380
Rb-Sr WhR, muscovite
214 ± 3; 222 ± 4.4; 222 ± 2.5; 223 ± 3.2
K-Ar, white mica
161 ± 4 170–159
Ar/ Ar 14 K-Ar, U/Pb, Pb/Pb
D. Ediacaran plutonism, Skookum Gulch Mélange, YT Plagiogranite crystallization, Skookum Butte Monzodiorite crystallization, Trail Gulch block
G. Silurian and Devonian detrital zircons, YT Moffett Creek Fm.—Silurian Duzel Phyllite—Silurian, Early Devonian Sissel Gulch Gw.—Siluro-Devonian Gazelle Fm. volcaniclastics—S-D H. Siluro-Devonian plutonism, TT Silurian intrusions S-D intrusions I. Early–Middle Devonian amalgamation, YT-FMT Semischist, Schulmeyer Gulch mélange Amphibolite of FMT J. Early Devonian plutonism and metamorphism, YT-TT-RT Intrusions, TT Schneider Hill stock, TT-YT Pillow basalts overlying YT-TT Balaklala Rhyolite, base of RT K. Late Devonian amalgamation, CMT Metamorphism of Salmon and Abrams Schists L. Triassic amalgamation, FJT Blueschist blocks in mélange below FMT M. Pre-Nevadan magmatism Middle Jurassic sheeted dikes, TT Middle Jurassic plutons, KST N. Nevadan orogeny, KST Nevadan thrusting Nevadan plutons O. Post-Nevadan magmatism Early Cretaceous plutons, KST Cretaceous dike cutting Devonian lavas, YT-TT Cretaceous tonalite dike, YT-TT
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4
7
1
6
10
10
40
39
10
13
11
15
ca. 150 ± 2 14 155–148 14 153–148
K-Ar, U/Pb, Pb/Pb
14
144–141 140 125
K-Ar, U/Pb, Pb/Pb 11 U-Pb zircon 11 U-Pb zircon
14
P. Accretion of KST to western North America 15 Rotation and drift ca. 144 (Early Cretaceous) paleomagnetism 16 Great Valley deposition on Klamath basement ca. 136 (Cretaceous) overlap strata 16 Base of Hornbrook successor basin ca. 100 (mid-Cretaceous) overlap strata Note: gw—graywacke; CMT—Central Metamorphic terrane; FJT—Fort Jones terrane; FMT—Forest Mountain subterrane; KST—Klamath superterrane; RT—Redding subterrane; S-D—Siluro-Devonian; TT—Trinity terrane; YT—Yreka-Trinity composite subterrane; WhR—whole-rock. 1 2 3 4 Data sources: Wallin (1989), Wallin and Gehrels (1995), and Wallin et al. (2000); Wallin (1990); Wallin et al. (1991); Wallin et al. (1988); 5 6 7 8 Mankinen et al. (2002); Jacobsen et al. (1984); Cotkin and Armstrong (1987) and Cotkin et al. (1992); Cashman (1980), and Hotz (1974), 9 10 11 recalculated by Cashman; Wallin and Metcalf (1998); Lanphere et al. (1968) and Kelley et al. (1987); Wallin, Martin, and Lindsley-Griffin, 12 13 14 15 16 unpub. data; Boucot et al. (1974); Hotz et al. (1977); Barnes et al. (1992); Irwin and Mankinen (1998), Irwin (2003); Mankinen et al. (1989), Nilson (1984), and Lindsley-Griffin et al. (1993).
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sheets exhibits a distinct combination of internal structure, composition, and metamorphic grade. Thrusting ended in the Early to Middle Devonian as the southeastern edge of the Yreka subterrane overrode the Ediacaran to Early Devonian mafic-ultramafic basement of the Trinity subterrane (Lindsley-Griffin, 1991, 1994; Lindsley-Griffin et al., 2006). Geophysical evidence for the subsurface relationships of these two subterranes is equivocal but seems to indicate a dense, relatively thin subhorizontal sheet of mafic-ultramafic rocks that extends westward beneath the Yreka subterrane from the Trinity subterrane, as indicated by a positive gravity anomaly over the Yreka subterrane (LaFehr, 1966; Griscom, 1977). However, the positive magnetic anomaly seen over the Trinity subterrane exposures (Irwin and Bath, 1962) decreases where the traverse crosses the Yreka subterrane (Griscom, 1977), suggesting that the Trinity basement extends beneath only part of the Yreka subterrane rather than all the way to its western edge, and therefore is discontinuous with the ultramafic rocks of the Forest Mountain subterrane (Fig. 1). To the east (Fig. 1), the Trinity subterrane appears to underlie the Redding subterrane as well as Cenozoic volcanic rocks of the Cascades Range (Blakely et al., 1985, 1997; Fuis et al., 1987). Modeling based on seismic, gravity, and aeromagnetic data suggests that the Trinity subterrane is 6.5–8 km thick and overlies less dense rocks (Zucca et al., 1986). The Trinity Subterrane The polygenetic ophiolitic Trinity subterrane (Trinity Complex of Lindsley-Griffin, 1991 and 1994) (Fig. 1) consists of two or more late Precambrian to early Paleozoic oceanic fragments sutured together along a pre-Silurian intraoceanic fault, which subsequently formed the basement for a younger ophiolitic crustal sequence built over it (Wallin and Metcalf, 1998; Lindsley-Griffin et al., 2006). The main components of the Trinity subterrane are 1. The late Proterozoic Trinity ophiolite (Lindsley-Griffin, 1994), a sequence of metagabbro, diabase and plagiogranite, and pillow basalts, overlying and faulted against intensely deformed serpentinized harzburgite. The metagabbro and plagiogranite have yielded Ediacaran zircons (C in Table 1; Wallin et al., 1988, 1991) and the pillow basalts have yielded a paleopole that is consistent with an Ediacaran age (Mankinen et al., 2002). Ductile deformation in the Trinity ophiolite ceased before the Early Ordovician, because relatively undeformed Lower Ordovician plagiogranitic dikes and stocks intrude the ophiolite (E in Table 1; Wallin et al., 1988). 2. The Cambro-Ordovician (?) Trinity peridotite is an areally extensive subhorizontal sheet of harzburgite, lherzolite, and dunite with mantle tectonite fabric and ductilely deformed compositional layering. The earliest ductile structures typical of mantle deformation are likely late Ediacaran to Cambrian. Mantle deformation ceased and the Trinity peridotite underwent pressure-release partial
3.
melting in Late Ordovician time (ca. 470 Ma; E in Table 1; Jacobsen et al., 1984), resulting in plagiogranite and gabbro dikes that crosscut and postdate the mantle tectonite fabric. Quick (1981, 1982) interpreted this partial melting as evidence of uplift in a mantle diapir. The Trinity peridotite probably was juxtaposed against the Trinity ophiolite before or during the partial melting event because we have observed relatively undeformed Upper Ordovician intrusions in both oceanic blocks. A nearly undeformed Siluro-Devonian ophiolitic sequence of voluminous plutons that intruded through the oceanic basement, and associated basaltic lavas overlying them, rests on the Trinity ophiolite-Trinity peridotite basement. Zircons from gabbros and plagiogranites of this assemblage have yielded ages of 435–412 Ma (H in Table 1; Jacobsen et al., 1984; Wallin et al., 1988; Wallin, 1989; Wallin and Metcalf, 1998). Petrology and geochemistry of this crustal sequence suggest that it represents a supra-subduction zone (SSZ) ophiolite that formed in the forearc of an active convergent margin (Wallin and Metcalf, 1998).
The Forest Mountain Subterrane West of the Trinity and Yreka subterranes is a linear belt of ultramafic and mafic rocks ~5 km wide that extends 55 km from Yreka to Callahan, which we propose to name the Forest Mountain subterrane (Fig. 1). This oceanic belt collided with the Yreka subterrane in the late Early Devonian, collapsing the Yreka subterrane into a series of thrust sheets and thrusting them eastward over the Trinity subterrane. This collision ca. 400 Ma is marked by plutonism in the Trinity subterrane, metamorphism of the oceanic amphibole gabbro in the Forest Mountain subterrane, and development of semischist in Yreka subterrane rocks along the suture zone (I in Table 1; Lanphere et al., 1968; Hotz, 1974; Cashman, 1980; Kelley et al., 1987). We propose the name “Forest Mountain subterrane” because of the excellent exposures and key radiometric age dates (I in Table 1) along California State Highway 3 where it crosses the Forest Mountain summit a few kilometers west of Yreka (Fig. 1). The western edge of the Forest Mountain subterrane is harzburgite with a mantle tectonite fabric superposed over compositional banding defined by varying proportions of olivine and pyroxene. A thin belt of amphibole gabbro borders the harzburgite on the east. Although Hotz (1977) termed this unit “amphibolite,” it is a typical uralitized layered gabbro of oceanic origin with relict igneous textures and compositional layering, altered to albite-hornblende-epidote by oceanic hydrothermal metamorphism. Lenses of true amphibolite within the Forest Mountain subterrane (mapped by Hotz, 1977, 1978, as ultramafic lenses) probably represent metamorphosed oceanic basalt. Both the eastern edge of the Forest Mountain subterrane and the western edge of the Yreka subterrane are metamorphosed to semischist. Hotz (1977) accurately described these semischists as
Paleogeographic significance of Ediacaran cyclomedusoids cataclastic protomylonites, indicating formation near or slightly above the brittle-ductile transition. We consider the semischist metamorphic age to represent the time when the Forest Mountain subterrane amalgamated with the Yreka and Trinity subterranes, ca. 400 Ma (I in Table 1). Geophysical data (Griscom, 1977) support the interpretation that the Forest Mountain subterrane is discontinuous with the Trinity subterrane. Thus, it should be separated from the older Trinity subterrane, as well as from the Central Metamorphic terrane that lies west of the Forest Mountain subterrane (Fig. 1). Redding Subterrane and the Forest Mountain–Yreka– Trinity–Redding Nucleus The Redding subterrane was deposited over the Trinity subterrane after it amalgamated with the Yreka and Forest Mountain subterranes; these four subterranes constitute the Forest Mountain–Yreka–Trinity–Redding nucleus against which all the younger terranes to the west collided. Although the Redding terrane is “native” to the Klamath superterrane, its paleomagnetic record shows that it rotated in the opposite sense to Laurentia throughout the Paleozoic and Mesozoic, and that it did not become part of North America until the entire Klamath superterrane arrived at the North American continental margin (Mankinen and Irwin, 1990; Irwin and Mankinen, 1998). Situated on a relatively passive edge of the growing superterrane, the Redding subterrane accumulated an ~200 m.y. record of intermittent sedimentation and sparse volcanism that reveals its perambulations through the ancient Pacific (Panthalassic) Ocean. Although the Redding subterrane experienced no collisions until the “docking” of the Klamath superterrane with North America, all the outboard terranes were amalgamated successively to the westward edge of the Forest Mountain–Yreka–Trinity–Redding nucleus (Irwin and Mankinen, 1998). Although the Redding subterrane has been termed an island arc by many authors, a subset of eastern Klamath workers have considered it to represent a forearc or intra-arc basin setting because of its relative paucity of volcanic rocks and long-lived but intermittent development (e.g., Fagin, 1980, 1983; Irwin, 1981; Lindsley-Griffin and Griffin, 1983; Wallin et al., 2000). The Redding subterrane (also termed Redding section and eastern Klamath terrane) consists of marine metasedimentary strata that are in part volcanogenic but which lack the thick metavolcanic and volcanic strata characteristic of true island arcs (cf. Garcia, 1978). The nearly 200-m.y. time span from the early Middle Devonian through the Early Jurassic during which its strata formed is much longer than is likely for relatively shortlived island arcs. The volcanic rocks that form the basal unit of the Redding subterrane, Copley Greenstone and Balaklala Rhyolite, interfinger with each other and grade rapidly upward into marine shales and graywackes of the Kennett Formation. Early Middle Devonian (ca. 390 Ma) basaltic dikes and pillow lavas that “stitched” the Yreka-Trinity subterranes together probably correlate with the oldest basaltic lavas of the Copley Greenstone.
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A Middle Devonian fish plate (Boucot et al., 1974) dates the Balaklala Rhyolite, and the Kennett Formation contains Middle Devonian brachiopods (Boucot and Potter, 1977). Intrusion of the 400 Ma Mule Mountain stock into the Balaklala Rhyolite further constrains its age (Albers et al., 1981). Thus the rifting episode that initiated the Redding subterrane was very short and gave way rapidly to marine sedimentation. Other volcanic units in the Redding subterrane are thin and represent short-lived local volcanic eruptions interspersed with carbonate platform development and volcaniclastic sedimentation into inter-arc basins (cf. Fagin, 1983; Miller and Harwood, 1990; Watkins, 1990). The Yreka Subterrane Wells et al. (1959) defined two formations in what eventually came to be termed the Yreka subterrane: the Upper Ordovician (?) metasedimentary Duzel Formation and the sedimentary Upper Silurian Gazelle Formation. Preston E. Hotz (1977, 1978) subdivided and renamed the Duzel Formation, breaking it into four formal units: Duzel Phyllite, Sissel Gulch Graywacke, Moffett Creek Formation, Antelope Mountain Quartzite, and several informal units: “Schulmeyer Gulch sequence,” “limestone of Duzel Rock,” and “schist of Skookum Gulch.” This paper examines the relationships between three of these units, redefining the “Schulmeyer Gulch sequence” of Hotz (1977) as the Schulmeyer Gulch Complex and discussing it and the Duzel Phyllite in context of the late Neoproterozoic cyclomedusoid fossils recovered from the Antelope Mountain Quartzite. Brief descriptions of other Yreka subterrane units are included for completeness. Antelope Mountain Quartzite Hotz (1977) named the Antelope Mountain Quartzite for the excellent exposures of quartzite on Antelope Mountain south of Yreka. Although Hotz considered it to be a distinctive lithologic unit of Ordovician (?) age, he suggested that it might be a facies of the Duzel Phyllite as is the Sissel Gulch Graywacke. Hotz (1977, 1978) showed the basal contact as depositional in some places and a thrust fault in others, whereas both Klanderman (1978) and Cashman (1980) considered the basal contact a low-angle fault and separated the Antelope Mountain Quartzite from the Duzel Phyllite because of its different composition and structural style. We have reexamined the basal contacts of the unit and conclude that they are all faults because they crosscut bedding; no beds above the contact are conformable with it, as would be necessary in a depositional contact. The age of the Antelope Mountain Quartzite is now known to be late Neoproterozoic–Ediacaran, according to cyclomedusoid fossils we have collected from it (Lindsley-Griffin et al., 1989, 2002, 2003, 2006). The formation is weakly metamorphosed to low greenschist facies, chlorite/ biotite zones. Stratigraphy. Klanderman (1978) recognized three informal members within the formation: (1) a lower quartz wacke unit; (2) a middle quartz wacke unit with somewhat thicker bedding, more modal quartz, and marker beds of granule and
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pebble conglomerate; and (3) an upper unit of bedded chert. His stratigraphy holds up fairly well when examined in the field, although it neglects some important aspects of the formation and may imply a greater simplicity than actually exists. Thick packets of phyllite and siliceous mudstone are interspersed with the quartz arenites of Klanderman’s lower and middle units, but were not well exposed in the late 1970s. Greatly expanded logging operations as well as new and improved roads over the past 10–15 yr have produced new outcrops and exposed sections of phyllites and siliceous mudstones up to several hundred meters thick within Klanderman’s lower and middle “quartzite members.” Good examples of extensive undeformed pelitic sections crop out on Antelope Mountain (Fig. 2) in the ditch of the logging road east of the summit, and on the road running along the north side of Cram Gulch toward Pythian Cave (Fig. 2). Much of Klanderman’s “chert member” is actually siliceous mudstone or argillite with minor phyllite, in some cases interbedded with thin quartzite layers. Composition. Both Hotz (1977) and Klanderman (1978) recognized three major lithologies within the Antelope Mountain Quartzite: quartz arenite with minor conglomerates, phyllite, and bedded chert. Both of them considered the quartz arenite to dominate the formation, although our 2003 fieldwork suggests that quartzite and phyllite are roughly equal in importance, with siliceous mudstone and chert constituting perhaps one-quarter of the unit. Only the resistant quartzites and true cherts form readily observable outcrops, but where roads and logging tracks have cut through the deep soil down to outcrop, sections of pelitic rocks up to 100 m thick are preserved. Quartz arenite and conglomerate. The quartz arenites (Fig. 3A) and the conglomerates interbedded with them (Fig. 3B) range in color through various shades of bluish gray and yellowish brown to grayish red, weathering to shades of brown, yellow, red, and orange due to iron oxide in the matrix. Klanderman (1978, p. 30 and 41), reported from his modal analyses of thin sections that the arenites consist of 60%–80% detrital quartz (both monocrystalline and polycrystalline); less than 2% detrital feldspar, both plagioclase and potassium feldspar, with trace amounts of sphene, zircon, rutile, muscovite, and shale fragments; 10%–27% matrix consisting of sericite, chlorite, hematite, and pyrite; and siliceous cement. Bond and DeVay (1980), who analyzed 18 samples of quartzite from the formation, reported that the quartz arenites consist of 90%–99% quartz, 0%–10% plagioclase, 0% potassic feldspar, and 0%–5% rock fragments, chert, and mica. Bond and DeVay (1980) also determined that the quartz consists of 52% coarsely polycrystalline grains, 37% monocrystalline grains, and 11% finely polycrystalline grains. Grain shapes range from subangular to rounded, and grain sizes range from very fine to very coarse. Hotz (1977, p. 18) also reported tourmaline and “altered ilmeno-magnetite” among the heavy minerals of the quartz arenites, and biotite in addition to the sericite and chlorite in the matrix. Hydrothermal quartz veins are common (Fig. 3A) and probably originated during metamorphism to low greenschist facies.
Bedding is well defined but undulatory, with bed thicknesses ranging widely from centimeter scale to half-meter scale. The undulatory bedding appears to result from thickening and thinning, possibly due to prelithification submarine creep, rather than from load casts, ripple marks, or other sedimentary structures, because the topography of the undulations is expressed on both the tops and bottoms of beds (Fig. 3C). Only rarely are true ripple marks observed. The quartz arenites include interbeds of quartz granule and pebble conglomerates (Fig. 3B). Klanderman’s modal analyses (1978, p. 51) indicate that the composition of the sandy conglomeratic matrix is similar to that of the quartz arenites: 44%–58% quartz, 0%–6% plagioclase, and 30%–38% matrix (clay, hematite, and chlorite), with silica cement (6%–17%). However, he found that for grains >2 mm diameter, the composition varies considerably, with the quartz clasts consisting mainly of polycrystalline quartz, and up to 12% of one or more of the following: microcline, chert, phyllite, and mudstone clasts (Klanderman, 1978, p. 52). Klanderman classified the conglomerates as sandy conglomerates to pebbly sandstones after Krumbein and Sloss (1963). Phyllite and argillite. Interbedded with the quartz arenites are pelites metamorphosed to phyllites and argillites (Fig. 3D), dark gray to greenish or olive gray on fresh surfaces, and weathering tan to greenish brown. Although both Hotz (1977) and Klanderman (1978) referred to these as shales, the bedding surfaces reflect light, and silt grains within the phyllites are flattened parallel to the foliation, clearly indicating that the recrystallized clays define a metamorphic foliation. The very fine grained siliceous argillites locally are interbedded with the phyllite intervals; in some cases phyllite bases grade upward into argillite tops (Fig. 3E). The fossiliferous argillites tend to be moderate reddish orange or brown (10 R 6/6–10 R 3.5/5), whereas the phyllite is slightly darker reddish brown (10 R 3/4) to grayish green (10 G 4/2). The cyclomedusoid fossils occur on the argillite tops of these graded beds (Fig. 3E). Siliceous mudstone and chert. Associated with the quartz arenites and phyllites are rocks that previous workers (Hotz, 1977; Klanderman, 1978) termed bedded cherts, but which are in large part siliceous mudstones (hardness <6) grading into and interbedded with true chert (hardness = 7). Some layers have chert cores and grade in both directions into siliceous mudstone tops and bottoms. Both siliceous mudstones and cherts are rhythmically bedded in layers 3–10 cm thick (Fig. 3F). However, the beds lack the varicolored millimeter-scale or finer internal laminations typical of marine siliceous oozes and the cherts that form from them. Instead, most of the beds consist of massive bluish gray to dusky blue (5 PB 5/2–3/2) cryptocrystalline silica that is translucent on thin edges. In some beds the color changes gradually within the bed, but without a systematic correlation between base, top, and color index. Other beds are marked by abrupt, pronounced color changes (e.g., from dusky blue to orange or pink). The harder layers consist of cryptocrystalline to microcrystalline quartz with thin phyllitic or argillitic partings; the softer layers
Figure 3. Antelope Mountain Quartzite. (A) Thick bed of massive quartz arenite, cut by white hydrothermal quartz veins; scale in cm and in. (B) Granule conglomerate bed, scale in cm. (C) Thin rhythmic beds of quartz arenite, scale in cm and in. (D) Phyllite interbedded with quartz arenite, scale in cm and in. (E) In situ cyclomedusoid fossils on the argillite top of a graded metasiltstone layer. The two large fossils, ~4.5 cm in diameter, resemble forms of the cyclomedusoid genus Beltanella Sprigg 1947. Several smaller, more deeply weathered impressions are also visible. Vertical joint sets and hydrothermal quartz veins cross the surface. (F) Rhythmically bedded black chert near the stratigraphic top of the Antelope Mountain Quartzite. Hammer handle is ~18 cm long.
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consist of very fine siliceous mud and a hammer easily scratches them. Klanderman (1978) reported sparse detrital grains of mica (muscovite?) and silt-sized quartz, as well as authigenic pyrite and hematite. In thin section he observed ghosts of carbonate rhombs replaced with silica, which he interpreted as evidence that the cherts were deposited above the calcium compensation depth. Klanderman (1978) found radial masses of chlorite lining some fractures within the chert, which we interpret as evidence that the siliceous lithologies are the same metamorphic grade as the quartz arenite and phyllite. Some bedding surfaces are undulatory or knobby. Structure. The Antelope Mountain Quartzite constitutes one of the structurally highest thrust sheets in the Yreka subterrane accretionary complex, structurally overlying thrust sheets of Silurian Duzel Phyllite and Ordovician? Schulmeyer Gulch Complex (Fig. 2). Block faults bound the exposures north of Shasta Valley (Fig. 2). In the northeastern exposure (Fig. 2), bedding is subhorizontal to gently dipping, and is folded into broad asymmetrical folds whose F1 fold axes are themselves folded gently about an F2 axis of folding nearly perpendicular to the first. This contrasts sharply with the steeply dipping bedding of both the Duzel Phyllite and the Schulmeyer Gulch Complex. On Antelope Mountain (Fig. 2), bedding is steeply dipping and more tightly folded, but is discordant with bedding of the underlying units. Small normal and reverse faults are common, but internal shearing is minimal and the packets of relatively incompetent pelitic rocks distributed throughout the stratigraphic section retain their original undeformed bedding, indicating that the unit is neither a broken formation nor a mélange. The fold style of the Antelope Mountain Quartzite contrasts markedly with the tight isoclinal folds of the Duzel Phyllite, which are crossed by moderately tight second folding giving it a “crinkled” appearance, as well as with the intensely sheared blocks-in-matrix nature of the Schulmeyer Gulch Complex. Multiple joint sets (Fig. 3E) disrupt bedding in all lithologies, causing outcrops to weather into piles of small blocks and flagstones. Contact relationships. Understanding the contact relationships between the Antelope Mountain Quartzite and adjoining units is essential to understanding its origin and being able to place the cyclomedusoid fossils in context. Previous workers have proposed many different contact relations and corresponding hypotheses for the formation’s origin: (1) partially thrust over, and partially depositional with, the Duzel Phyllite (Hotz, 1977); (2) thrust over the “Schulmeyer Gulch sequence” (Hotz, 1977; Klanderman, 1978); (3) thrust over the Duzel Phyllite (Klanderman, 1978; Cashman, 1980); (4) olistolith or slide block into the “Schulmeyer Gulch sequence” (Wallin et al., 2000); (5) interbedded with the Duzel Phyllite, and the same age as the Grouse Ridge Formation of the Central Metamorphic terrane (Table 1). This last hypothesis, generated by Cashman’s (1980) interpretation of the “Schulmeyer Gulch sequence” of Hotz (1977), is disproved by the presence of Ediacaran cyclomedusoids in the Antelope Mountain Quartzite, as well as by the presence of the Forest Mountain subterrane west of the Yreka subterrane. During
the 2003 field season we revisited most of the basal contacts of the Antelope Mountain Quartzite, and we carefully reexamined the Duzel Phyllite and “Schulmeyer Gulch sequence” to test these hypotheses. In most localities, the Antelope Mountain Quartzite lies structurally above either the Duzel Phyllite or the “Schulmeyer Gulch sequence” of Hotz (1977). Discordant bedding attitudes everywhere along the base of the Antelope Mountain Quartzite confirm that it is fault-bounded, thus disproving hypotheses 1, 4, and 5 and confirming both hypotheses 2 and 3. We interpret the subhorizontal faults along the basal contact as gently folded thrust faults and subvertical faults along the basal contact as younger block faults. We have observed mylonitized quartzite and chert along one basal thrust fault on Antelope Mountain, suggesting that the thrusting may have occurred near or below the brittle-ductile transition. A tectonically sliced, scalyfoliated outcrop of Antelope Mountain quartzite and phyllite marks one well-exposed near-vertical block fault that crosscuts the thrust fault, but the outcrop displays no mylonite, suggesting that the younger block faulting was shallower than the older thrust faulting. Fossils. J.R. Griffin collected the first cyclomedusoid fossils in 1970 on Antelope Mountain (Fig. 1); however, paleontologists working in the area did not consider the disk-shaped impressions to be of biological origin. The features came to our attention again in 1987 when a local citizen donated a sample containing several impressions (Figs. 4A and 4B) to the museum at the Klamath National Forest headquarters in Yreka (Lindsley-Griffin et al., 1989), and we began collecting additional samples for analysis. Over 250 individual fossils have been recovered from five localities within the Antelope Mountain Quartzite. Cyclomedusoid fossils. Farmer has described the cyclomedusoid fossils in detail in another paper (Lindsley-Griffin et al., 2006); results are summarized here. The disk-shaped impressions occur in very fine grained reddish-orange argillites located on the hinges or limbs of folds where axial plane cleavage dies out or is nearly parallel to bedding. The fossiliferous zones are located near the stratigraphic top of the formation, where they comprise intervals a few meters thick of rhythmically bedded argillite, siliceous mudstone, and fine-grained phyllite layers 2–12 cm thick, with rare interbeds of quartz arenite. The layers have recrystallized micaceous grains several millimeters in diameter in their phyllitic bases, but they grade rapidly up into mudstone, and then into very fine grained argillite at their tops. Within the thin fossiliferous zones, cyclomedusoids are preserved as epireliefs on the upper surfaces of multiple fine-grained layers (Figs. 3E and 4A). They appear as round to oblate impressions with raised outer rims and centers depressed convexly into the underlying sediment (Figs. 4B and 4C). In some cases, individual impressions are regularly spaced 3–10 cm apart (Fig. 4A); in other cases groups of two to seven impressions are crowded together in clusters (Fig. 4C). The fossils form convex-downward molds into the thinly laminated argillite bed tops (Fig. 4B), and are covered by convex-upward casts of phyllitic metasiltstone (Figs. 4D and 4E). Thin sections cut through the impressions
Figure 4. Ediacaran cyclomedusoid fossils in the Antelope Mountain Quartzite. (A) Large argillite slab with four cyclomedusoid fossils (arrows) exposed on its top surface, and possibly two additional fossils in the next lower bed represented by circular humps (arrows with question marks). The half disk preserved along the broken left edge of the sample resembles an Ediacaria-type form with a wider central core and a narrower exumbrellar surface. The large complete disk in the center of the sample resembles a Beltanella-type form with a broad median surface. Three joint sets cross the bedding surface, with a train of en echelon tension gashes at the right. Scale in cm. (B) Close-up view of the large Beltanella-like form in the center of (A), showing the tripartite structure characterized by a raised outer ring, a flat median surface, and a central core depressed into the underlying sediment. Although faint concentric furrows mark the median surface, it lacks pronounced radial ornamentation. Top of the central core has been eroded away. Scale in cm. (C) Cluster of in situ cyclomedusoid fossils similar to Ediacaria forms: four are spaced 1 cm or less apart, with a fifth ~2.5 cm away. Diagonal lines crossing bedding surface are joints. Arrows indicate fossils; scale in cm. (D) Side view of Ediacaria-type form showing its upper surface rising 0.5 cm above the bedding surface. The fossil compresses the continuous laminae below; a thin film of sediment overlies it. Tops are indicated by decrease in grain size from silt to clay in direction of white arrow. Scale in cm. (E) Cyclomedusoid fossil on argillite top of a size-graded lower bed, covered by veneer of coarse silty to fine sandy phyllite that forms base of the upper bed. Scale in cm. (F) Cylindrical tube-shaped features (arrows) on chert-siliceous mudstone layers that resemble trace fossils suggest bioturbation. Most are confined to the upper bedding surfaces but the one at right appears to be vertical and concentrically color-zoned. Scale in cm.
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show that the underlying argillite laminae are depressed downward but remain continuous as they pass beneath the fossils. The laminae are not broken by silt pipes or other evidence of dewatering. Thus, these features are biogenic rather than inorganic pseudofossils (see discussion in Lindsley-Griffin et al., 2006). The Antelope Mountain assemblage is dominated by forms allied to the Cyclomedusa complex of Sun (1986). Cyclomedusa davidii (Sprigg, 1947), the type species for the genus, is characterized by a sub-umbrellar surface, which is divided by fine radial grooves and concentric furrows that subdivide the disk into distinct regions. However, not all species of this genus exhibit radial ornamentation (Glaessner, 1984; Farmer et al., 1992). Among cyclomedusoids reported from northwestern Laurentia (e.g., Narbonne and Hofmann, 1987; Narbonne and Aitken, 1990) and from Baltica (Farmer et al., 1992), radial ornamentation is rare to lacking. The Antelope Mountain cyclomedusoids (Fig. 4) lack radial ornamentation and thus resemble specimens from the latter localities. The Antelope Mountain specimens are similar to the cyclomedusoid genus Ediacaria, characterized by a basic tripartite structure in which the ex-umbrellar surface is subdivided by concentric furrows into three regions with a raised marginal rim of variable width (Fig. 4A). The central subdivision may be either raised or depressed. Other specimens resemble the cyclomedusoid genus Beltanella Sprigg 1947 (Fig. 4B). They range in outer diameter from ~1.8 cm to 10 cm and in shape from circular to oval (Fig. 4C). These fossils are typical of the widespread Ediacaran biota, and they fall within the accepted range of variation reported in the literature (e.g., see discussion by Farmer et al., 1992). Thus far, we have recovered Ediacaran cyclomedusoid fossils only from the fine-grained siliceous layers near the stratigraphic top of the formation; none have been observed in the grayish-green phyllites or the siliceous mudstones or cherts. In the quartz arenites, we have observed only one disk-shaped feature that may be a poorly preserved cyclomedusoid. Some of the true chert layers overlying the fossiliferous intervals exhibit cylindrical structures that appear to be burrows, but the fossiliferous argillite layers are undisturbed by bioturbation. Possible fossils in cherts. Above the cyclomedusoid-bearing argillites, and in some cases partially interbedded with them, are siliceous mudstones that grade upward into cherts with siliceous mudstone tops, then into cherts. In thin sections of this lithology, Klanderman (1978) observed tiny spheres filled with polycrystalline quartz, which he interpreted as molds of radiolarians. He dissolved six samples in hydrofluoric acid, yielding “several fragmented and one whole siliceous sponge spicule” (Klanderman, 1978, p. 64), and he interpreted the Antelope Mountain cherts as recrystallized biogenic ooze of radiolaria and spicules. We have observed some beds with bases that are depressed into scourlike shapes and tops that appear to exhibit buildups of thin siliceous laminae that resemble stromatolitic layers. Locally we observed oblate and cylindrical features that resemble borings and bottom markings along bed surfaces (Fig. 4F); these suggest limited bioturbation of the siliceous unit.
Provenance and origin. Bond and DeVay (1980) studied the Antelope Mountain quartz arenites to determine their provenance, finding both monocrystalline and polycrystalline quartz to constitute up to 90%, and plagioclase to be more abundant than potassic feldspar. They concluded that this suggested a source terrane of metamorphic and potassic plutonic rocks. Because of the relatively high proportion of polycrystalline to monocrystalline quartz grains they ruled out largely volcanic sources such as a volcanic arc as well as mature quartz sandstones of the craton. The textural immaturity of the quartz arenites and granule-pebble conglomerates (Klanderman, 1978) suggests a proximal source. Bond and DeVay (1980) interpreted their data on the arenites as supporting deposition in a submarine fan near a passive continental margin. They interpreted their sandstone petrography as indicating that the Antelope Mountain Quartzite was derived from a continental source; however, they found no evidence to link it with the Laurentian craton. They did find the Antelope Mountain Quartzite to be similar to the Shoo Fly Formation of the northern Sierra Nevada, an issue that we will address below. E.T. Wallin and his colleagues have reported a highly diverse suite of detrital zircons derived from the quartz arenites of the Antelope Mountain Quartzite (Wallin, 1989; Wallin and Gehrels, 1995; Wallin et al., 2000). They analyzed zircons that are free of cores and inclusions, which yielded concordant to somewhat discordant ages ranging from 3.0 Ga to 1.32 Ga (Table 1). Wallin divided the detrital zircons into four distinct suites: (1) 1.3 Ga (three grains); (2) 1.8–1.9 Ga concordant ages (most grains); (3) Early Proterozoic (2.0–2.08, 2.15–2.30, 2.38, 2.45, 2.53 Ga); and (4) Archean (2.53, 2.68, 2.75, 2.9–2.98, 3.13–3.20 Ga). Depositional environment. Several conflicting hypotheses about the depositional environment of this unit have been published. Hotz (1977) interpreted the Antelope Mountain Quartzite as a submarine fan complex, but Klanderman (1978) noted that the lack of Bouma sequence sedimentary structures suggests deposition of the arenites by non-turbid gravity flows. Wallin et al. (2000) interpreted the unit as shallow marine on the basis of oscillation ripple marks. We have not observed such features, but we have found asymmetrical current ripple marks in quartzite immediately below the fossiliferous argillite zones. We concur with Klanderman’s (1978) observation that the quartz arenites and granule-pebble conglomerates are compositionally mature but texturally immature. We also agree with Klanderman’s interpretations of the pelitic intervals as hemipelagic deposition between gravity flows and of the cherts as resulting from a major environmental change allowing accumulation of siliceous oozes, accompanied by a sharp decrease in siliciclastic input. However, the subtle graded bedding within the argillite intervals argues for current deposition, perhaps in the form of slightly turbid gravity flows of very fine grained sediment. The cyclomedusoid fossils reveal little about the environment because little is understood about their life cycle; however, their presence is consistent with shallow to moderately deep marine conditions (e.g., Glaessner, 1979; Runnegar, 1992a, 1992b).
Paleogeographic significance of Ediacaran cyclomedusoids These combined observations suggest that the Antelope Mountain quartz arenites were deposited by non-turbid quartzrich gravity flows, interspersed with hemipelagic sedimentation in a moderately deep marine basin near a continent. Over time, the continental source became more distant, siliciclastic input decreased, and slightly turbid, very fine grained gravity flows dominated. The fossiliferous zones of phyllite grading into argillite tops suggest distal gravity flow deposition at irregular intervals, each depositional event burying the existing cyclomedusoid colony under fine-grained silt; then, as the suspended clay-sized particles of silica and dust settled out, another cyclomedusoid group would colonize the new surface. As time passed, the oceanic environment became more silica enriched, and deposition of siliceous ooze and siliceous muds became important. As oceanographic conditions allowed siliceous oozes to dominate, either the cyclomedusoids died out or they ceased to be preserved, perhaps because the bottom scavengers that produced the borings (Fig. 4F) destroyed their remains. Schulmeyer Gulch Complex Informally termed the “Schulmeyer Gulch sequence” by Hotz (1977), this assemblage is herein named the Schulmeyer Gulch Complex, because it meets the criteria for a mélange, namely, it consists of native and exotic blocks in sheared matrix (e.g., Hsu, 1968, 1974; Silver and Beutner, 1980). According to the Code of Stratigraphic Nomenclature (NACSN, 1983) a mélange should be classified as a Complex. Hotz (1977) originally named the unit for exposures in Schulmeyer Gulch (Fig. 2), which he considered typical. However, the unit varies so much in its block composition that it is difficult to define a type locality or even a type area. Table 2 lists some readily accessible exposures that display typical characteristics of the unit. Previous work. Other than the excellent pioneering work by P.E. Hotz (1977, 1978) and the master’s thesis by Klanderman (1978), descriptions of the Schulmeyer Gulch assemblage are sparse. Hotz described the phyllite-siltstone of the matrix, and several block compositions including chert, quartz arenite,
Location
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limestone, and metavolcanic schist. Klanderman (1978) briefly described the “Schulmeyer Gulch sequence” adjacent to the Antelope Mountain Quartzite in his thesis area as consisting of recrystallized limestone, bedded chert, and radiolarian chert conglomerate, but he did not observe any quartzite within the unit. Cashman (1980) considered the Schulmeyer Gulch assemblage to be part of the Duzel Formation, undifferentiated, whereas Noto (2000) considered it to be mélange. Wallin et al. (2000) considered the Antelope Mountain Quartzite to be a block within the Schulmeyer Gulch mélange but did not discuss the Schulmeyer assemblage. Composition and structure. Blocks in the Schulmeyer Gulch Complex include the following lithologies: (1) Limestone: recrystallized impure carbonate, now a coarsely crystalline marble with thin phyllitic folia. It ranges in color from dark dusky blue to nearly white. It is resistant in this area and forms knockers that are readily visible from a distance (Fig. 5A). Hotz mapped these limestone blocks as a separate lithofacies; the largest is over 600 m wide and 3 km long (Hotz, 1977, p. 20). In Figure 5B, phyllitic matrix separates and encases large limestone blocks whose internal structures are discordant to each other. The edges of these large blocks are mylonitized (Fig. 5C), with thin bands of recrystallized calcite separated by trains of graphite and chlorite, but the deformation dies out away from the limestone block margins. A swarm of smaller limestone pieces, sliced off during deformation, surrounds most of the larger limestone blocks. (2) Chert, now recrystallized to cryptocrystalline quartz, with thin (centimeter-scale) bedding. Klanderman (1978) reported float blocks of chert conglomerate containing molds of radiolarians from the Schulmeyer Gulch Complex. The bedded chert ranges in color from dusky blue and dark charcoal gray to lighter grays and off-white. (3) Quartz wacke, dark bluish gray and recrystallized, and metasiltstone. Because these lithologies are quartz-poor relative to the quartz arenites of the Antelope Mountain quartzite, they are unlikely to have been derived from it. We observed no quartz arenites similar to the Antelope Mountain quartzite within the Schulmeyer Gulch Complex. (4) Massive greenish-brown
TABLE 2. TYPICAL EXPOSURES OF THE SCHULMEYER GULCH COMPLEX Matrix Blocks observed
A. Yreka: Trails from College of the Siskiyous campus to Greenhorn Park and water tank, NW ¼ SW ¼ Sec. 34 to NW ¼ SE ¼ Sec. 33, T 45 N, R 7 W (Figs. 5D, 5E, and 5F)
Abundant sheared phyllite and metasiltstone
Chert, hydrothermal quartz, hydrothermal quartz in phyllite, quartz wacke
B. Yreka: NW corner of the intersection of Greenhorn Road and Hiram Pitt Road., NE ¼ NW ¼ Sec 34, T 45 N, R 7 W (Figs. 5B and 5C)
Sparse green phyllite
Recrystallized limestone, metavolcanic schist
C. Yreka: Hillside Park, road cut at intersection of Comstock Road and Pyrite Lane, NE ¼ NW ¼ Sec. 35, T 45 N, R 7 W
Sheared green phyllite
Recrystallized limestone, metasiltstone, quartz wacke
D. Highway 3 southwest of Yreka, SE ¼ NE ¼ Sec. 4, T 44 N, R 7 W (Fig. 5A)
Sheared green phyllite
Recrystallized limestone
E. Between Yreka and Montague: road cuts along Montague Road, just west of railroad crossing, SE ¼ NW ¼ Sec. 24, T 45 N, R 7 W
Abundant sheared phyllite
Chert, hydrothermal quartz, recrystallized limestone, metavolcanic schist
Note: Localities in T 45 N are on the U.S. Geological Survey topographic map of the Hornbrook 15’ quadrangle; locality D, in T 44 N, is on the Yreka 15’ quadrangle.
Figure 5. Schulmeyer Gulch Complex. (A) Ductilely stretched limestone block surrounded by sheared phyllitic matrix, Highway 3 SW of Yreka. The matrix exhibits subvertical N30°W scaly foliation (dashed lines) that is discordant to the long axis of the block (solid line). Shrub in upper center, ~1.5 m high, for scale. (B) Large limestone blocks separated by and encased in phyllitic matrix. Outcrop is in Yreka, at the intersection of Hiram Page Street and Oregon Street off Greenhorn Road. Person for scale. (C) Mylonitized limestone at margin of one of the large blocks in (B) Scale in cm and in. (D) Fractured hydrothermal quartz block cut by ductilely stretched and necked quartz vein (arrows). Outcrop is on hiking trail between College of the Siskiyous campus and Greenhorn Park in Yreka. Scale is 15 cm (6 in). (E) Block composed of intensely fractured and tectonically sliced phyllite invaded and partly replaced by multiple generations of hydrothermal quartz veins. Dashed line outlines one folded vein; it is discordant to other folded veins above and below the line. Outcrop is on hiking trail between College of the Siskiyous campus and Greenhorn Park in Yreka. Scale in cm and in. (F) Sheared phyllite and metasiltstone matrix of the Schulmeyer Gulch Complex with subhorizontal scaly foliation. Outcrop is above College of the Siskiyous campus along the trail from Greenhorn Park to the water tank in Yreka.
Paleogeographic significance of Ediacaran cyclomedusoids to greenish-gray phyllitic schist (sericite + chlorite + epidote ± actinolite) with crinkly foliation that is typically deeply weathered. Hotz (1977, p. 21–22) described these as ranging from massive to phyllitic and containing the metamorphic minerals chlorite, epidote, and (in some blocks) abundant actinolite. We interpret these as metavolcanic schist and greenstone. (5) Intensely fractured hydrothermal quartz, white cryptocrystalline SiO2 that is typically cut by multiple generations of hydrothermal quartz veins (Fig. 5D). Where multiple veins are present, the older veins are ductilely stretched, necked, and in some cases folded. Undeformed or less deformed younger veins crosscut them. The youngest veins, judging from crosscutting relationships, are undeformed. (6) “Wads” or balls of phyllite intensely penetrated by multiple generations of hydrothermal quartz veins that replace up to 80% of the original rock volume (Fig. 5E). These blocks are not part of the phyllitic matrix because shear zones bound them and the veins do not cross into the matrix. The matrix is phyllitic to schistose, and intensely deformed (Fig. 5F). Even though it is recrystallized and metamorphosed to low greenschist facies, some original bedding is still present and its character suggests that the matrix probably originated as interbedded mudstone and siltstone. Most of the metasiltstone in the matrix is noncalcareous, in contrast to that of the Duzel Phyllite, and finer-grained than Duzel metasiltstone. The phyllite in the matrix is not as green as the Duzel Phyllite, ranging in color from greenish gray to reddish brown to moderate or dark red. Metasiltstone beds in the matrix pinch and swell ductilely, and are also brittlely sliced and offset along the foliation. Individual grains of quartz and feldspar within the matrix are locally deformed into pinch-and-swell structures and flattened parallel to the foliation, demonstrating that the matrix has undergone semipenetrative deformation. Beds are tectonically sliced along the scaly foliation (Fig. 5F), which is defined by subparallel flakes of chlorite and sericite. Because some of the matrix deformation is ductile but the metamorphic grade is too low to produce this type of deformation, we suggest that deformation of the matrix began while the sediments were still soft, and presumably the pore fluid pressure was high. Deformation continued while, or resumed after, the matrix sediments attained brittleness. We have observed that this change in mechanical behavior is typical of polygenetic mélanges such as the Franciscan Complex, or the Gregg Ranch Complex of the Yreka subterrane (discussed below). Locally, the Schulmeyer Gulch matrix is folded into decimeter-scale isoclinal folds that are not reflected within the blocks. We have seen such folds in Franciscan Complex matrix near Crescent City, California. The overall structure of blocks-in-matrix is enhanced by the relationships of foliation and bedding along and around block boundaries. Foliation of the matrix strikes into block compositional boundaries, either ending abruptly at the compositional change where foliation is perpendicular to the boundary, or bending and wrapping around the block where foliation is subparallel to the boundary (Fig. 5A). Blocks are sheathed by deformed matrix (Fig. 5B). Foliation and/or bedding within the blocks also
15
is cut by block margins (Fig. 5A), demonstrating that the relationship is structural rather than sedimentary. Some of the blocks appear to be less deformed than the matrix (Fig. 5D), whereas others appear to be more deformed than the matrix (Fig. 5E). The margins of many blocks are more intensely deformed than the block interiors (Fig. 5C). In some cases, block margins are sliced and folded into the deformed matrix; in other cases a thin rind of finely comminuted mylonitized block material surrounds the block (Fig. 5C). In all cases this boundary effect dies out rapidly toward the block interior. Hotz (1977) observed that the Schulmeyer Gulch Complex is cataclasized near the suture zone with the Forest Mountain subterrane, in places being a protomylonite. We concur with this observation. His descriptions and photomicrographs reveal classic deformation textures even more intense than the pervasive shearing of the mélange matrix. Metamorphic minerals in the semischist include epidote or clinozoisite, colorless mica, biotite, actinolite, shredded chlorite, and in one sample, porphyroblastic garnet; the quartz is crystalloblastic (Hotz, 1977, p. 22). The age of the semischist is ca. 400 Ma (I in Table 1), suggesting the minimum age of the Schulmeyer Gulch Complex is Silurian, but no other isotopic data are available. Origin. During the 2003 field season we tested the various hypotheses proposed by previous workers by examining a number of key outcrops and contacts (Table 2). The large blocks of chert and limestone are exotic to the matrix sediment, because the carbonate probably formed in a shallow-marine environment whereas the cherts probably formed in a deep-marine environment rich in siliceous organisms and remote from siliciclastic input. The metavolcanic schist blocks are also exotic to the matrix. The mudstone-siltstone matrix might represent oceanic trench turbidites, although we observed no Bouma structures— such features, even if present originally, would not have survived the intense deformation. The hydrothermal quartz-phyllite balls must have begun forming during earliest metamorphism of the unit, with multiple generations of veins continuing to form during ongoing metamorphism and deformation. The older veins are stretched, pulled apart, and ptygmatically folded, whereas the younger veins that crosscut them are less complexly folded (Fig. 5E), attesting to multiple cycles of deformation under metamorphic conditions. The undeformed quartz veins formed after ductile folding ceased but before the intense shearing that separated the intensely veined phyllite into blocks of hydrothermal quartz and quartz-veined phyllite. Hotz (1977) separated the Schulmeyer Gulch assemblage from other Yreka subterrane units because of its heterogeneity. He noted its lack of internal stratigraphy and suggested that the unit might be a mélange. However, he considered the quartz wackes to be analogous to the Antelope Mountain Quartzite and the phyllites to be analogous to the Duzel Phyllite, an interpretation we dispute, and suggested it might be “a sedimentary facies of the two other formations” (Hotz, 1977, p. 24). This latter interpretation is untenable because the Antelope Mountain Quartzite is Ediacaran and the Duzel Phyllite–Sissel Gulch Graywacke
16
Lindsley-Griffin et al.
sequence is Siluro-Devonian (G in Table 1); both units exhibit different structural styles from the Schulmeyer Gulch Complex and thus have had different histories. Also, the phyllitic siltstone of the mélange matrix is only locally calcareous, whereas the Duzel Phyllite is calcareous almost everywhere, one of the distinguishing characteristics Hotz included as part of its definition. Cashman (1980, p. 458) included the Schulmeyer Gulch mélange, the Duzel Phyllite, and amphibolite of the Forest Mountain subterrane in the same unit on the basis of similar “parental rock types, lithologic associations, isoclinal folding accompanied by greenschist to lower amphibolite facies metamorphism, parallel structural trends, and Devonian metamorphic age.” She interpreted the exotic limestone blocks in both the Schulmeyer Gulch mélange and the semischist along the suture zone as interbeds within the Duzel Phyllite. However, detrital zircons in the Duzel Phyllite (G in Table 1) indicate that it was deposited in the interval from Late Silurian through Early Devonian, whereas the Forest Mountain subterrane was not emplaced against the Yreka subterrane units until mid-Devonian time (I in Table 1), rendering Cashman’s (1980) hypothesis unlikely. Wallin et al. (2000, p.121) stated that the Antelope Mountain Quartzite “is a large, resistant tectonic block in the Schulmeyer Gulch mélange that lies in the footwall of the Mallethead Thrust. …The Antelope Mountain was incorporated into the mélange by the Early to Middle Devonian because it was folded tightly along with the main thrust of the Mallethead thrust system.” The thick pelitic intervals between the quartz arenites of the Antelope Mountain Quartzite lack evidence of shearing, tectonic slicing, and scaly foliation (cf. Figs. 3D, 3E, and 5F). Furthermore, swarms of fragments of the same composition typically surround large blocks in mélange, having been sliced off from them and transported along the scaly foliation during tectonic dismemberment. No such swarms of quartzite, argillite, or chert that might have been derived from the Antelope Mountain Quartzite are present in the Schulmeyer Gulch Complex. Finally, blocks in mélanges typically are sheathed in deformed matrix, which is lacking from the Antelope Mountain Quartzite. Thus the Antelope Mountain Quartzite cannot be a block in the Schulmeyer Gulch Complex. Although Klanderman (1978) did not discuss the phyllitic matrix or the blocks-in-matrix character of the Schulmeyer Gulch Complex, he described the internal contacts as nondepositional and termed the unit a mélange. He considered the Antelope Mountain Quartzite to be thrust over the Schulmeyer Gulch assemblage. We agree with Klanderman’s conclusion; the Antelope Mountain Quartzite is thrust over the Schulmeyer Gulch Complex as well as over the Duzel Phyllite at Antelope Mountain (Fig. 2). Thrusting likely occurred during emplacement of the Forest Mountain subterrane (Figs. 1 and 2) in the mid-Devonian. The age of the Schulmeyer Gulch Complex is poorly constrained. The K-Ar age of 432 Ma from colorless mica in the semischist (Hotz, 1977, p. 24) provides a minimum age of Early Silurian for the unit, whereas the observation by Klanderman
(1978) of radiolarian molds in chert conglomerate float blocks suggests a maximum age of early Paleozoic, probably Ordovician. We conclude that the Schulmeyer Gulch Complex is a polygenetic mélange formed by a combination of sedimentary and tectonic processes, most likely along an ensimatic convergent margin. It may be as old as late Precambrian, or as young as Early Silurian. Other Yreka Subterrane Units Lindsley-Griffin et al. (2006) reviewed evidence that the Yreka terrane is an accretionary wedge-forearc terrane. Other thrust sheets within the Yreka subterrane support repeated cycles of mélange formation, trench fill, and forearc deposition (e.g., Lindsley-Griffin and Griffin, 1983; Lindsley-Griffin et al., 1991, 2006; Wallin et al., 1991, 1995; Wallin and Trabert, 1994; Wallin and Metcalf, 1998). In addition to the Schulmeyer Gulch Complex, mélanges occur in four other thrust sheets within the Yreka subterrane. All are polygenetic, each with a distinctive assemblage of exotic and native blocks, a unique structural style, and other characteristics that distinguish it from the other mélanges. The mélanges are notable for their mixture of limestone, volcanic or volcaniclastic, and radiolarian chert blocks. The limestones suggest temperate- to tropical-latitude reefs, the cherts suggest deep ocean ooze, and the volcanic blocks indicate proximity to a volcanic arc. Interleaved with the mélanges are multiple thrust sheets of two trench-fill sequences, one of Silurian age and the other Siluro-Devonian (F in Table 1). Different zircon populations, structural styles, and metamorphic grades distinguish these trench deposits from each other. A nearly undeformed Early Devonian trench-slope basin sequence overlies mélange of the Gregg Ranch Complex along the southeastern edge of the subterrane. Key points for each are summarized below. Facey-Duzel mélange. The weakly metamorphosed FaceyDuzel mélange (Fig. 2) occupies the structurally highest klippen and probably represents fragments derived from an ocean island chain (Lindsley-Griffin and Griffin, 1983). It contains Early and Late Ordovician limestone blocks (Table 3) as well as bedded chert and chert conglomerate, mudstone-siltstone conglomerate, red shale, volcanic breccia, and pillow basalt (Porter, 1973; Hotz, 1977; Nachman, 1977; Rohr, 1978). The pillow lava is high-titanium alkali olivine basalt (Nachman, 1977) typical of intraplate ocean-island lavas. The pelitic to silty, tectonically sheared matrix is characteristic of fine-grained trench fill common to mélanges formed at convergent margins. Its maximum age is Late Ordovician. Skookum Gulch mélange. The Late Ordovician (F in Table 1) “schist of Skookum Gulch” or Skookum Gulch mélange is a high-grade assemblage containing mainly oceanic blocks: lawsonite- and glaucophane-bearing blueschist, serpentinite, plutonic and metabasaltic rocks, greenschist, metaquartzite, calcareous and dolomitic marble, and metachert in a blueschist-greenschist matrix (Hotz, 1977, Haessig, 1988, 1989; Cotkin, 1987; Cotkin et al., 1992). Cotkin et al. (1992) concluded that the metabasaltic blocks formed at a mid-oceanic ridge, and that most of the schists
Age Neoproterozoic Ediacaran
TABLE 3. AGES AND BIOGEOGRAPHIC AFFINITIES OF YREKA AND REDDING SUBTERRANE BIOTA Unit Fossil Group Affinities Comments Reference Antelope Mountain Quartzite
Cyclomedusoids
Balt, Yuk
Red argillites near stratigraphic top of formation
Lindsley-Griffin et al. 1989, 2002, 2003, 2006
Gregg Ranch Complex
Gastropods
YT only: 5 genera, 19 species; others AK, Sct
Mélange, KCF
Rohr, 1980
Not recognized
N.A.
N.A.
N.A.
N.A.
Facey Rock Ls
Conodonts
Cosmopolitan
Potter et al., 1990a
Llanvirn (Middle) Llanvirn (Middle) Llanvirn (Middle) Llandeilo (Middle) Llandeilo (Middle)
Horseshoe Gulch unit Horseshoe Gulch unit Horseshoe Gulch unit Gregg Ranch Complex Gregg Ranch Complex
Conodonts
Cosmopolitan
Graptolites
Pacific (Aust-Amer); Yuk, NV Inner-slope, E & W North America VA-TN-AL; WM; Girvan, Sct; Kzkh YT; Girvan, Sct; VA-TNAL; WM; Kzkh
Ls block in mélange of Facey-Duzel Rock Mélange—Potter’s GF “member 2” Mélange—Potter’s GF “member 2” Mélange—Potter’s GF “member 2” Mélange—Potter’s GF “member 1” Mélange—Potter’s “GR unit member 1”
Llandeilo (Middle)
Gregg Ranch Complex
Llandeilo (Middle) Middle (undivided) Llandeilo (Middle)
Gregg Ranch Complex Gregg Ranch Complex Gregg Ranch Complex, Horseshoe Gulch unit
Gastropods and conodonts Gastropods, brachiopods Sphinctozoan sponges
W North America
Llandeilo–early Caradoc (Middle)
Horseshoe Gulch unit
Trilobites
Ashgill (Late)
Gregg Ranch Complex Gregg Ranch Complex Gregg Ranch Complex
Brachiopods
Effna Limestone, VA; Chi; Sweden; ScotoAppalachian; Girvan, Sct; Kzkh WM; E-Cent. AK; Yuk; ML NW Eur
Early Paleozoic Middle Ordovician– Early Devonian Cambrian Ordovician Arenig (Early)
Ashgill (Late) Ashgill (Late)
Ashgill (Late)
Ashgill (Late) Silurian Wenlock (Middle) Ludlow (Late)
Ludlow (Late)
Trilobites Brachiopods Brachiopods: Xenambonites, Craspedelia Conodonts
Conodonts Solitary rugose corals
Gr Bas; E App; Bal; Girvan, Sct; SCh
“Old World aspect” YT; Aust; AK-AT; Yuk
Grewingkia penobscotensis Elias, 1982 only in YT, ML, northern ME YT; Aust; AK-AT; Yuk; ML
Mélange—Potter’s “member 1” and “member 3” Red ls clasts in cgl blocks, KCF mélange Gray ls clasts in cgl blocks, KCF mélange Mélange—Potter’s GF “member 1”; KCF “member 3”
Mélange—Potter’s GF “member 2”
Potter et al., 1990a Berry et al., 1973 Potter et al., 1990a Potter, 1990a, 1990b; Potter et al., 1990a Potter, 1990a, 1990b; Potter et al., 1990a Potter et al., 1990a
Rohr, 1980 Rohr, 1980 Rigby and Potter, 1986; Rigby et al., 1988; Potter et al., 1990a; Rigby et al., 2006 Potter et al., 1990a
Mélange—Potter’s GF “member 5” Mélange—Potter’s GF “member 5” Mélange—Potter’s GF “member 5”
Potter, 1990a, 1990b; Potter et al., 1990a Potter et al., 1990a
Mélange—Potter’s GF “member 5”
Potter et al., 1990a; Rigby et al., 2005
Potter et al., 1990a
Gregg Ranch Complex, Horseshoe Gulch unit Gregg Ranch Complex
Sphinctozoan sponges
Tabulate corals
ML; App, Eur, Urals, Sib, Chi, Aust
Mélange—Potter’s GF “member 5”
Potter et al., 1990a
Gregg Ranch Complex Gregg Ranch Complex
Tabulate corals
YT; Alt, Chi, App, Eur, Urals, Sib, Aust “Uralian-Cordilleran region”
Potter et al., 1990a
Gregg Ranch Complex
Tabulate corals
Mélange—Potter’s Gazelle Fm.; KCF Ls olistoliths or blocks in mélange—Potter’s GF “member 1” N.A.
Shelly faunas
App, Eur, Urals, Sib, Alt, Chi, Aust
Potter et al., 1990a
Potter et al., 1990a Continued
18
Age
Lindsley-Griffin et al. TABLE 3. AGES AND BIOGEOGRAPHIC AFFINITIES OF YREKA AND REDDING SUBTERRANE BIOTA (continued) Unit Fossil Group Affinities Comments Reference
Devonian Early
Gregg Ranch Complex Gregg Ranch Complex Gregg Ranch Complex
Conodonts
Early
Gregg Ranch Complex
Brachiopods
Lochkovian (Early)
Gregg Ranch Complex
Early to Middle?
Gregg Ranch Complex Gazelle Fm. of Lindsley-Griffin et al., 1991
Trilobite Warburgella rugulosa subsp. indet. Tetracorals
Early Early
late Siegenian to middle Emsian (Early)
Tabulate corals Tetracorals
4 cosmopolitan; 1 YT and NV only Only in YT or Eureka, NV Gr Bas, W Canada; Yuk
NV, N. BC, Yuk, Canadian Arctic (“Cordilleran”) Uncertain
Conodonts
Gr Bas; W and N Canada; SE Chi N.A.
Ls blocks in mélange Mélange—Potter’s GCU and GF Ls clasts in cgl blocks, in mélange—Potter’s “GF member 3” Mélange—Potter’s “Gazelle member 3” Mélange—Potter’s “Gazelle member 3”
Potter’s “member 1” of the Gazelle Fm. Calc sdst concretion interbedded in argillites near lower contact with GRC Calc sdst lens or concretion IB in nontectonized mudstone and siltstone N.A.
Savage, 1976a; Potter et al., 1990a Potter et al., 1990a Potter et al., 1990a
Boucot and Potter, 1977; Potter et al., 1990a Potter et al., 1990a
Potter et al., 1990a Savage, 1977
Emsian (Early)
Gazelle Fm. of Lindsley-Griffin et al., 1991
Conodonts
“Cordilleran”
Early
Gazelle Fm. of Lindsley-Griffin et al., 1991 Balaklala Rhyolite— RT
Dichotomous branching plants Placoderm fish plate
N.A.
Cosmopolitan
Tuff layer in upper strata
late Eifelian (Middle) late Eifelian (Middle)
Kennett Fm.—RT
Brachiopods
Uncertain
Ls, sdst, mdst
Kennett Fm.—RT
RT only
Limestone
Savage, 1976b; Potter et al., 1990a
late Eifelian (Middle)
Kennett Fm.—RT
Cosmopolitan
Limestone
Middle
Kennett Fm.—RT
Conodont: Polygnathus kennettensis Conodont: Polygnathus cf. trigonicus Tabulate corals
Eureka, NV; Suplee terrane, OR
N.A.
Savage, 1976b; Savage and Boucot, 1978 Potter et al., 1990a
Permian Early
McCloud Ls—RT
Middle
Boucot et al., 1974; Potter et al., 1990a
Wallin and Trabert, 1994 Boucot et al., 1974; Boucot and Potter, 1977 Savage, 1976b
Fusulinids, Accreted terranes, CA, Defines the distinctive Stevens et al., 1990; corals, BC, AK; or endemic to McCloud Province Belasky et al., 2002 brachiopods, RT gastropods Note: Abbreviations (rock names): calc sdst—calcareous sandstone; cgl—conglomerate; GCU—“Grouse Creek Unit” of Potter et al. (1990a); GF—Gazelle Fm. of Potter et al. (1990a); GR—“Gregg Ranch unit” of Potter et al. (1990a); GRC—Gregg Ranch Complex; KCF—“Kangaroo Creek Fm.” of Rohr (1980); IB—interbedded; ls—limestone; mdst—mudstone; sdst—sandstone. Abbreviations (localities): Alt—Altai fold belt of Mongolia; AK—Alaska, no terrane specified; AK-AT—Alaska, Alexander terrane; AK-NFT—Alaska, Nixon Fork subterrane of Farewell terrane; Amer—American; App—Appalachian mountain belt; Aust—Australia; Balt—Baltica; BC—British Columbia; CA—California; Chi—China; Eur— Europe; Gr Bas—Great Basin; Kzkh—Kazakhstan; ME—Maine; ML—Montgomery Ls, northern Sierra Nevada, CA; NV—Nevada; OR— Oregon; RT—Redding subterrane; SCh—South China; Sct—Scotland; Sib—Siberia and NE former USSR; VA-TN-AL—Virginia, Tennessee, Alabama; WM—White Mountains, Alaska; YT—Yreka subterrane; Yuk—Canadian Yukon; N.A.—not applicable or unknown.
formed from oceanic sediments, whereas the tourmaline-bearing quartz-albite schists indicate continental provenance. A monzodiorite block of probable continental origin and a plagiogranite block possibly derived from the Trinity ophiolite (Lindsley-Griffin et al., 2006) are Ediacaran in age (D in Table 1). Erosion of this mélange soon after uplift shed detritus of quartz-albite schist, graphitic schist,
and plagiogranite into Late Silurian conglomerates of the Horseshoe Gulch mélange (Potter, 1982; Haessig, 1989). Mélange of the Callahan area. At the southern end of the Yreka subterrane (Fig. 1) is a weakly metamorphosed mélange containing blocks of Callahan Chert, limestone conglomerate with clasts containing Silurian corals and brachiopods, undated
Paleogeographic significance of Ediacaran cyclomedusoids recrystallized limestone, chert conglomerate and chert breccia, and basaltic pillow breccia (Rohr, 1978; Lindsley-Griffin and Griffin, 1983). The age of the radiolarians in the Callahan Chert is undetermined. Silurian fossils recovered by Rohr (1978) from limestone clasts in the conglomerate blocks give maximum age of those blocks only. The matrix, Rohr’s (1978) “mudstone of Thompson Gulch,” exhibits semi-penetrative scaly foliation that wraps around and encases the Callahan Chert blocks (LindsleyGriffin et al., 2006). This mélange occupies a thrust sheet above the Moffett Creek Formation and below the thrust sheet of FaceyDuzel mélange. Its maximum age is Silurian. Gregg Ranch Complex. The Gregg Ranch Complex (Lindsley-Griffin et al., 1991) is a weakly metamorphosed Early Devonian mélange that occupies one or more thrust sheets along the southeastern edge of the Yreka subterrane (Fig. 1). LindsleyGriffin and her colleagues (Lindsley-Griffin, 1982; LindsleyGriffin et al., 1991, 2006; Sivers, 2002) have recognized at least eight mélange units within this lithologically diverse assemblage, excluding the Horseshoe Gulch assemblage discussed below. Blocks within the Gregg Ranch Complex include a wide variety of shallow- to deep-marine sedimentary rocks that contain most of the early Paleozoic biota in the Yreka subterrane (Table 3), as well as debris derived from the Trinity subterrane (LindsleyGriffin et al., 2006). The matrix typically is mudstone but locally grades into siltstone and fine sandstone; it exhibits semi-penetrative scaly foliation. Kinematic analysis of the Gregg Ranch Complex suggests that it formed by active convergence from the west (present coordinates) during the Early to early Middle Devonian convergence and collision of the Forest Mountain subterrane (Sivers, 2002; Lindsley-Griffin et al., 2006). Mélange of Horseshoe Gulch. When Lindsley-Griffin et al. (1991) defined the Gregg Ranch Complex as including all the lithologically diverse, stratally discontinuous unmetamorphosed rocks in the southern Yreka subterrane, they included the Horseshoe Gulch mélange unit as part of the Gregg Ranch Complex. However, as more information becomes available about these rocks, it seems that they may comprise a distinct mélange more closely linked to the Skookum Gulch blueschist mélange than the Gregg Ranch Complex. Reasons for this interpretation include the following: (1) Only Ordovician fossils are known from Horseshoe Gulch (Table 3). If the Horseshoe Gulch unit formed in the mid-Devonian as did the Gregg Ranch Complex, it logically should have blocks containing Silurian and Early Devonian fossils. (2) The Horseshoe Gulch mélange blocks contain many distinctive fossils that do not appear in other mélange units of the Gregg Ranch Complex, including graptolites, trilobites, and sphinctozoan sponges (Table 3). (3) Late Silurian conglomerates in the Horseshoe Gulch assemblage contain detritus eroded from the Skookum Gulch mélange after it was metamorphosed (Potter, 1982; Haessig, 1989).Thus, the maximum age of the Horseshoe Gulch mélange is Early Silurian. (4) The Horseshoe Gulch mélange appears to occupy a separate thrust sheet at the lowest structural level, just below the Skookum Gulch blueschist mélange (Hotz, 1978; Lindsley-Griffin and Griffin, 1983; Haes-
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sig, 1988). This is the logical position it should occupy, if it is an older mélange formed along the same active margin on which the Skookum Gulch mélange formed. However, the two thrust sheets may not have been juxtaposed until the mid-Devonian collision of the Forest Mountain subterrane. Moffett Creek Formation trench-fill sequence. The Silurian Moffett Creek Formation consists of thick-bedded calcareous metasiltstone and metasandstone exhibiting abundant white mica on weathered surfaces, with thin weakly metamorphosed phyllitic layers at their tops. The silty-sandy lithologies range from quartz wackes with some plagioclase feldspar and rare potassic feldspar to feldspathic wackes (Hotz, 1977; Lindsley-Griffin, 1982). Composite lithic grains of quartzite, volcanic rock, and metachert suggest a variety of sources, ranging from cratonic to volcanic edifice and recycled subduction complex (Hotz, 1977; Gin and Wallin, 1994). Sedimentary structures in the wackes include graded bedding, cross-laminae, and sole markings; these and the rhythmic repetition of sand-shale beds attest to the turbidite origin of the formation (Lindsley-Griffin, 1982). The structure ranges from locally olistostromal (Lindsley-Griffin, 1982) to broken formation near faults (Hotz, 1977), but most of the formation is stratally continuous (Gin and Wallin, 1994). The formation exhibits locally developed slaty cleavage subparallel to bedding, folded into gentle open symmetrical folds (LindsleyGriffin et al., 2006). Detrital zircons (Wallin and Gehrels, 1995) are 435–420 Ma or Silurian (G in Table 1); thus, the maximum age of the Moffett Creek Formation is Silurian. Sissel Gulch Graywacke and Duzel Phyllite trench-fill sequence. The Sissel Gulch Graywacke consists of thick, massive, very fine grained feldspathic metagraywacke beds with thin phyllitic partings (Hotz, 1977). It is texturally immature, consisting of abundant quartz, ~10% fresh plagioclase feldspar (An30–35), no potassium feldspar, many rock fragments of metachert, and siliceous and mafic metavolcanic rock (Hotz, 1977). This composition suggests an intraoceanic active margin provenance. The graywacke interfingers with and grades up into the Duzel Phyllite of Hotz (1977), which consists of green chloritic phyllite with thin interbeds of metagraywacke cut by abundant quartzo-feldspathic segregation bands. Metagraywackes in the Duzel Phyllite consist of 45% detrital quartz (34% monocrystalline and 11% polycrystalline); 45% matrix (clay, chlorite, hematite, mica); 9% calcite cement; and a trace of plagioclase (Klanderman, 1978). The Duzel Phyllite is isoclinally folded with near-vertical axial surfaces, and subvertical to steeply dipping bedding. A second deformation produced open asymmetrical folds with near-vertical axial surfaces roughly perpendicular to axial surfaces of the first folding. Sparse, highly ductile slump folds are contained within isoclinal fold limbs. Both formations are metamorphosed to greenschist facies, low chlorite zone. Ages of detrital zircons (Wallin and Gehrels, 1995) are 440–410 Ma for the Duzel Phyllite and 420–410 Ma for the Sissel Gulch Graywacke (G in Table 1, Siluro-Devonian); thus, the maximum age is Early Devonian. Gazelle Formation trench-slope basin. Lindsley-Griffin et al. (1991) redefined the Gazelle Formation of Wells et al. (1959),
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restricting the formation to the Lower Devonian graywackes and siliceous shales exposed on Gazelle Mountain. The upper Siegenian to middle Emsian age is based on conodonts from calcarenite interbeds within the unit (Table 3). The formation consists mostly of well-bedded and laterally continuous radiolarian-bearing siliceous shale and mudstone, with minor interbeds of organic-rich calcarenite, and radiolarian chert, volcaniclastic sandstone, and conglomerate (Lindsley-Griffin et al., 1991; Wallin and Trabert, 1994). The Gazelle Formation is at least 1.25 km thick and occupies an elongate synclinal basin 22 km by 8 km (Wallin and Trabert, 1994). The formation overlies the Gregg Ranch Complex along a tectonized depositional contact that suggests it was deposited directly onto a substrate of active accretionary complex (Lindsley-Griffin and Fisher, 1989). The rhythmically bedded shales and siliceous mudstones of the Gazelle Formation exhibit dewatering veins and contain abundant radiolaria outlined by pyrite framboids and dead oil, suggesting organic-rich hemipelagic sedimentation in a marine-upwelling regime analogous to modern west-facing convergent margins such as Peru (LindsleyGriffin and Fisher, 1989; Lindsley-Griffin et al., 1991). PALEOGEOGRAPHY OF THE EASTERN KLAMATH TERRANE Geological and Geophysical Evidence To place the Yreka subterrane in its paleogeographic context, we must consider all geological and geophysical evidence. It is widely accepted that the Yreka subterrane has been closely associated with the Trinity subterrane in space and time (e.g., Lindsley-Griffin and Griffin, 1983; Irwin and Mankinen, 1998; Wallin and Metcalf, 1998; Lindsley-Griffin et al., 2006). Thus we can use evidence from the Trinity subterrane to flesh out the story of the Yreka subterrane. Tectonostratigraphic Character The Yreka subterrane comprises an accretionary complexforearc terrane. The turbidite packages are Silurian, Siluro-Devonian, and Early Devonian in age; and the two mélanges whose ages have been established are the Upper Ordovician blueschist-bearing Skookum Gulch schist and the Lower Devonian Gregg Ranch Complex (Tables 1 and 3). At least three, and possibly four, distinctly different mélanges occupy other thrust sheets (Lindsley-Griffin and Griffin, 1983; Lindsley-Griffin et al., 2006); permissible ages for these range from late Precambrian? (Schulmeyer Gulch Complex) through early Paleozoic (Ordovician? mélange of Facey-Duzel Rock; Silurian? mélange of the Callahan area). The Trinity subterrane consists of a late Neoproterozoic ophiolite faulted against a Cambro-Ordovician? oceanic block, which formed the composite basement over which a Siluro-Devonian SSZ ophiolite developed (Wallin and Metcalf, 1998; Lindsley-Griffin et al., 2006). This composite polygenetic subterrane comprises the oceanic basement over which the Yreka subterrane was thrust in the late Early Devonian. Amalgamation of the Forest Mountain subterrane with the Yreka
and Trinity subterranes began the mid-Devonian development of the Redding subterrane (Fig. 1), which continued to evolve until the Late Jurassic. Although no fully developed volcanic arc is preserved, sediments in the Redding subterrane appear to record nearby volcanic activity through Paleozoic and Mesozoic time; thus the Klamath superterrane was associated with ensimatic active margins until its Early Cretaceous accretion to North America. The Antelope Mountain Quartzite is anomalous because it was deposited in a shallow to moderately deep marine basin receiving detritus from a continent. Paleomagnetic Data Paleomagnetic data for 580–570 Ma units of the Trinity ophiolite show that it formed at 7.1° ± 3.2° latitude at some unknown longitude (Mankinen et al., 2002). We assume, from the extensive evidence in the literature (Lindsley-Griffin and Griffin, 1983; Irwin and Mankinen, 1998; Wallin and Metcalf, 1998; Lindsley-Griffin et al., 2006), that the Yreka subterrane was close to the Trinity subterrane in the Neoproterozoic; thus it, too, must have been near 7° latitude in the Neoproterozoic. After the Early Devonian juxtaposition of the two subterranes, paleomagnetic data for the mid-Devonian basaltic lavas that overlap both subterranes show that the composite was at 31.1° ± 5.0° latitude (Mankinen et al., 2002). Although Mankinen et al. (2002) interpreted this evidence as indicating latitudinal concordance with North America, paleomagnetic data do not differentiate between northern and southern paleolatitudes. We shall use the geological and paleontological evidence to select the likeliest hemisphere. Paleomagnetic declinations also provide useful information. Irwin and Mankinen (1998) compared paleopoles derived from the Klamath superterrane with North American (Laurentian) paleopoles, finding that the Klamath superterrane paleopole was oriented ~110° counterclockwise away from North America’s paleopole in the Middle Devonian. As the Klamath superterrane grew by progressive amalgamation, it gradually rotated clockwise while North America rotated counterclockwise (Irwin and Mankinen, 1998); the two could not have accreted together until these opposing rotations ceased. Detrital Zircons The Antelope Mountain Quartzite contains detrital zircons ranging in age from 3.0 Ga to 1.32 Ga, according to analyses of grains that are free of cores and inclusions (Wallin, 1989; Wallin et al., 2000). U-Pb analysis of 46 single grains yielded concordant to slightly discordant results, which are summarized in our Table 1. The turbidite trench deposits of the Yreka subterrane all yield lower Paleozoic detrital zircons (Table 1). Although Wallin assumed a North American, British Columbia, origin for the Antelope Mountain Quartzite zircons (Wallin, 1989; Wallin and Gehrels, 1995; Wallin et al., 2000), such an origin does not explain all the zircon data. For example, British Columbia does not include source rocks for the 1.3 Ga detrital zircons (Wallin et al., 2000). However, composition of the Antelope Mountain Quartzite is similar to that of quartzite blocks in the Sierra City
Paleogeographic significance of Ediacaran cyclomedusoids mélange of the Shoo Fly Complex, northern Sierra Nevada (Bond and DeVay, 1980; Harding et al., 2000) as well as to quartzites of the Roberts Mountains allochthon of Nevada and the Alexander terrane of Alaska (Wright and Wyld, 2006). Detrital zircons in the Antelope Mountain Quartzite most likely originated from the same source as detrital zircons in the Alexander terrane (Wright and Wyld, 2006), an issue we examine in more detail below. Regional Relationships Many authors assume that the Yreka subterrane formed on, or very close to, the continental margin of western Laurentia (e.g., Boucot et al., 1974; Wallin et al., 2000). However, the Klamath superterrane lies outboard of a number of other terranes accreted before its Early Cretaceous collision, including the Roberts Mountains allochthon, Golconda allochthon, Black Rock Desert terrane, and Hells Canyon terrane (Burchfiel et al., 1992). Klamath subterranes have been linked to the Alexander terrane and the Nixon Fork subterrane of the Farewell terrane, which lie outboard of other Alaskan terranes (e.g., Churkin et al., 1982; Csejtey et al., 1982). These inboard tectonic elements must have accreted to North America before the Klamath superterrane
could dock, and they would have acted as barriers to biotic communication and sediment transport. Tectonics and Biogeography To interpret the significance of the fossil faunas correctly, we must map them on the best paleogeographic reconstructions available for their time period. All accreted terranes represent tectonic elements that formed somewhere other than their present location; these accreted terranes range from far-traveled to parautochthonous. It may be difficult to determine just how far a given terrane has traveled, and some may have translated laterally along a continental margin during a long drawn-out accretion process. Accretion is not complete until rotation and deformation end, even if “first contact” is distant in both space and time. A “cosmopolitan” biota is a group of plants and/or animals that is widely distributed over a number of geographic areas worldwide, whereas a “provincial” biota is a group of temporally and spatially associated communities of plants and/or animals (Jackson, 1997). Provinciality was relatively high during most of the Paleozoic, but was minimal during the Late Devonian and Mississippian (Table 4). Controls on faunal diversity
TABLE 4. LATE PRECAMBRIAN AND PALEOZOIC TIME UNITS AND PROVINCIALITY Epoch Stage Age at base Character (Ma) Permian Guadalupian, N.A. 270.6 Highly provincial Lopingian (Late) Cisuralian N.A. 299.0 Highly provincial (Early) Carboniferous Pennsylvanian N.A. 318.1 Somewhat provincial Mississippian N.A. 359.2 Somewhat cosmopolitan Devonian Late Famennian Highly cosmopolitan Frasnian 385.3 Cosmopolitan Middle Givetian Somewhat cosmopolitan Eifelian 397.5 Provincial Early Emsian Very highly provincial Pragian Highly provincial (Siegenian) Lochkovian 416.0 Highly provincial (Gedinnian) Silurian Pridoli N.A. 418.7 Provincial Ludlow Ludfordian/Gorstian 422.9 Provincial Wenlock Homerian/Sheinwoodian 428.2 Provincial Llandovery Telychian Somewhat provincial Aeronian/Rhuddanian 443.7 Cosmopolitan Ordovician Late (Ashgill) 460.9 Somewhat provincial Middle Darriwilian Provincial (Caradoc, Llandeilo) 471.8 Early (Llanvirn, Arenig) Provincial Tremadocian 488.3 Cambrian Furongian N.A. Provincial Paibian 501.0 Middle N.A. 513.0 Highly provincial Early N.A. 542.0 Provincial. Ediacaran N.A. N.A. 600.0 Provincial (Vendian) Cryogenian N.A. N.A. 850.0 N.D. Note: Estimated provinciality based on Jell (1974), Boucot (1988), Jaanusson (1979), Palmer (1979), Ross (1979), Ross and Ross (1979), Blodgett et al. (1988, 1990, 2002), Stevens et al. (1990), Waggoner (1999, 2003), Boucot and Blodgett (2001). Numerical ages and time-stratigraphic terminology after Gradstein et al. (2004); parentheses indicate terms common in the literature that have been dropped by the International Commission on Stratigraphy. N.A.—not applicable; N.D.—not determined. Period
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and provinciality include tectonic, ecologic, and biologic factors that have been extensively discussed in the literature (e.g., Valentine, 1972; Valentine and Moores, 1970, 1972; Johnson and Dasch, 1972; Newton, 1987; Cocks and Torsvik, 2002). Although differences between faunas may also arise from facies changes, Yreka subterrane faunas seem to represent mostly the same facies—shallow reefal communities intermixed with deep marine trench and accretionary-complex lithologies. Distance between different communities may be less important than the connections or barriers between them. High levels of endemism suggest a high degree of isolation; endemism plus a mixed biogeographic signal may be typical of isolated island chains such as Hawaii (Newton, 1987). In interpreting biogeographic relationships, faunal similarity does not equate to geographic closeness, nor does dissimilarity equate to geographic distance. In some cases, dissimilarity may simply indicate a barrier to interaction, and uniqueness may indicate isolation. Both highly cosmopolitan faunas and highly endemic faunas give little information about biogeography (e.g., Cocks and Torsvik, 2002). In the present-day western Pacific Ocean, multiple island arc systems, oceanic plateaus, and other tectonic elements line up along the continental margins. Some are quite close to the continents, whereas others are several thousand kilometers away. Some are amalgamating with each other even though separated from the nearest continent by several intervening terranes. Each of these tectonic elements may be barriers to some biota as well as migration paths for others. Looking into the future, one can predict that each of these western Pacific tectonic elements will eventually accrete onto the Eurasian continent or the Indo-Australian margin, creating a stack of multiple accreted terranes similar to that formed during the Paleozoic and Mesozoic in the North American Cordilleran terranes. Neoproterozoic and Paleozoic biogeography. A single large ocean, Panthalassa, and the supercontinent of Rodinia, which formed ca. 1100 Ma, characterized the late Precambrian world. The arrangement of fragments within Rodinia is subject to debate (cf. Dalziel, 1997; Lawver et al., 2002; Scotese, 2002). We have chosen to begin with the paleogeographic reconstruction used by Waggoner (1999) as our starting point because it fits the Ediacaran biota best (Fig. 6). Paleozoic biogeography (Figs. 7–12) is shown on the paleogeographic maps of Lawver et al. (2002), which seem to accommodate the faunal, paleomagnetic, and regional geologic data best. The reconstruction used by Waggoner (1999) for the breakup of Rodinia places Antarctica and Australia along the western edge of Laurentia (present orientation) as in the SWEAT (Southwest U.S.–East Antarctic) (Moores, 1991) and AUSWUS (Australia–Western U.S.) (Brookfield, 1993) hypotheses. Rodinia rifted and drifted apart to form Laurentia and Gondwana, although the timing and nature of the breakup are controversial. Some authors contend that Rodinia did not finish breaking up until ca. 550 Ma (e.g., Moores, 1991; Dalziel, 1991; Veevers et al., 1997); others contend that western Laurentia had separated from East Gondwana by ca. 700 Ma (e.g., Dalziel, 1997; Unrug, 1997). According to Murphy and
Figure 6. Ediacaran biogeography relative to Rodinia, the Yreka subterrane, and the Trinity subterrane. Symbols indicate biogeographic affinities of Ediacaran biota: squares—White Sea assemblage; triangles—Nama assemblage (darker triangles have cloudiniids only, no soft-bodied biota); circles—Avalon assemblage. Yreka subterrane cyclomedusoids are most similar to the White Sea assemblage, and the 580–570 Ma Trinity ophiolite formed in a post-Rodinia ocean basin. Stars indicate three possible locations of the Yreka and Trinity subterranes at either 7°N or 7°S latitudes (Mankinen et al., 2002) during the rifting of Rodinia. Modified after Waggoner (1999, fig. 3; 2003), Rogers (1996), Veevers et al. (1997), and Dalziel (1997).
Nance (2004), Rodinia broke up ca. 760 Ma. As the fragments drifted away, they left behind a short-lived Southern Hemisphere supercontinent termed Pannotia (Dalziel, 1997). Pannotia began to break up ca. 550 Ma (Murphy and Nance, 2004). Because Trinity ophiolite oceanic crust is 580–570 Ma in age, it formed after Rodinia rifted apart, on an active margin peripheral to Pannotia, in the Panthalassic Ocean (Fig. 7). As Rodinia broke apart, Australia and Antarctica separated first from northern Laurentia, now the miogeocline of western North America, and Siberia broke off from eastern Laurentia, now the Canadian Arctic (Fig. 7A). Laurentia then moved south during the late Ediacaran (Fig. 7B), possibly by rifting off from Siberia (Kirschvink, 1992). Australia and Antarctica likely remained proximal to Siberia through the Early Cambrian, and Baltica probably broke off from Laurentia in late Ediacaran to Early Cambrian time (Kirschvink, 1992). However, biogeographic analysis by Lieberman (1997) suggests that Baltica rifted off from Laurentia before Siberia did. Throughout the early and middle Paleozoic, continental fragments created by the breakup of Rodinia and Pannotia dispersed across the Panthalassic Ocean (Figs. 8–11) before converging
Paleogeographic significance of Ediacaran cyclomedusoids
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Figure 7. Ediacaran biogeography in the late Neoproterozoic at (A) 600 Ma and (B) 570 Ma. Some possible positions of the Yreka subterrane (oval) and the Trinity subterrane (crescent) indicated by symbols at near 7°N and 7°S latitudes (Mankinen et al., 2002). Counterclockwise rotation relative to Laurentia is derived from paleopole of Mankinen et al. (2002). Stars indicate Ediacaran affinities to tectonic elements created by rifting of Rodinia. Paleogeographic reconstructions after Dalziel (1997) and Lawver et al. (2002).
Figure 8. Middle Cambrian (510 Ma) biogeography. Yreka subterrane (oval) has no Cambrian fossils, but was likely located near the Trinity subterrane (crescent) in a position intermediate between Ediacaran and Ordovician locations, and near Alaskan terranes with similar faunal affinities, placing it between Siberia and Laurentia. Paleogeographic reconstructions after Dalziel (1997) and Lawver et al. (2002).
(Fig. 12) to form Pangea in the late Paleozoic (Dalziel, 1997; Lawver et al., 2002; Scotese, 2002). Disparate terranes consisting of island arcs, intraplate island chains, carbonate platforms, and other elements formed; some were consumed or destroyed and others amalgamated or accreted. Some of these terranes became the accreted terranes of the Appalachian-Caledonian orogenic belt in a series of collisions in the early to middle Paleozoic, culminating when Europe and Africa collided. Other terranes, including Kazakhstan, were trapped between Europe and Siberia during a complex series of collisions that lasted some
Figure 9. Middle Ordovician (460 Ma) biogeography, including biogeographic affinities for the Late Ordovician. The most likely location of the Yreka-Trinity subterranes is in the Uralian Seaway between Laurentia, Siberia, Kazakhstan, and Baltica. East-to-west oceanic currents (arrows) in the equatorial belt would have facilitated development of faunal affinities with Australia, N. China, and S. China. At the eastern edge of Laurentia, the equatorial currents would have split into two currents flowing around the continent, facilitating faunal affinities with eastern Appalachia (E. App.) and the north and south margins of Laurentia. Paleogeographic reconstructions after Dalziel (1997) and Lawver et al. (2002).
50 m.y. in the middle to late Paleozoic (Talent et al., 1987), forming the Ural Mountains within the Eurasian supercontinent by the late Carboniferous (Fig. 12). However, the South China block did not accrete to Eurasia until the Triassic (Talent et al., 1987). The Alexander terrane formed adjacent to Siberia or in the Uralian Seaway between Siberia and Laurentia (Soja and
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Figure 10. Late Silurian (420 Ma) biogeography. Yreka-Trinity subterrane locations in the Northern Hemisphere between Kazakhstan, Siberia, and Euramerica are more likely than locations in the Southern Hemisphere because the Yreka subterrane displays no faunal affinities with southern Laurentia and southern Gondwana. Arrows suggest possible ocean currents. Paleogeographic reconstructions after Dalziel (1997) and Lawver et al. (2002).
Antoshkina, 1997) in the early to middle Paleozoic. Most of the Cordilleran terranes remained adrift in the Panthalassic Ocean until the late Mesozoic, avoiding being trapped within the south polar continent of Gondwana (Fig. 12), which formed at the end of the Paleozoic (Dalziel, 1997; Lawver et al., 2002; Scotese, 2002). Table 3 summarizes ages and biogeographic affinities for Yreka subterrane fossils as well as for selected fossils from the Redding subterrane. A review of all Klamath superterrane fossils is beyond the scope of this paper, but Middle Devonian and Permian fossils from the Redding subterrane are included because they are especially significant. In most cases, we use the biogeographic terminology given in the original paleontological description, even where it is difficult to relate these early descriptions to present terrane terminology. For example, where the original source used “Alaska” we give “Alaska—no terrane specified” in Table 3. We give more specific biogeographic information where it was included in the original source or where we could deduce a more specific link from a locality map. Table 4 gives pertinent period and stage names along with their numerical age ranges (Gradstein et al., 2004) and indicates times that were particularly provincial or cosmopolitan. Table 5 summarizes biogeographic subdivisions and the localities included in them for the late Precambrian and the Paleozoic. In descending order of size, subdivisions are termed realms, regions, provinces, and subprovinces; the still evolving Ediacaran terminology uses assemblage. Because biogeographic terminology evolved piecemeal over many decades, beginning with early, pre-plate tectonics work, most biogeographic names are based on present locations of fossils, not locations where the animals originally lived. Thus the “localities included” column in Table 5 indicates the present-day location of the various terrane elements and tectonic components, but not necessarily where the faunas were when they were alive.
Figure 11. Devonian biogeography. Paleogeographic reconstructions after Dalziel (1997) and Lawver et al. (2002). (A) Early Devonian (410 Ma) Yreka and Trinity subterranes shown as outline shape. The Trinity subterrane at this time consisted of two amalgamated oceanic blocks on which a supra-subduction zone ophiolite was developing. The Yreka subterrane lay nearby and was receiving detritus from the Trinity subterrane. (B) Middle Devonian (380 Ma) Yreka and Trinity subterranes are shown as outline shape after collision with the Forest Mountain subterrane. Paleomagnetic data on Middle Devonian overlap strata suggest a location at either 31°N or 31°S (Mankinen et al., 2002) and a 110° counterclockwise rotation relative to Laurentia (Irwin and Mankinen, 1998). Yreka-Trinity subterranes were located outboard of other Cordilleran terranes, and were most likely in the Northern Hemisphere judging from faunal affinities. The Suplee terrane (parallelogram), now part of the Blue Mountains of Oregon, was an outboard carbonate platform and/or island chain with affinities to the Yreka subterrane; its location is entirely arbitrary.
Yreka Subterrane Biogeography Many previous workers have stated that Yreka subterrane faunas are “North American” (e.g., Potter, 1990a, 1990b; Potter et al., 1990a; Wallin et al., 2000). However, careful reexamination of the original literature in the light of current paleogeographic
Paleogeographic significance of Ediacaran cyclomedusoids
Figure 12. Early Permian (290 Ma) biogeography. Yreka, Trinity, Forest Mountain, and Central Metamorphic terranes are shown in outline shape after their amalgamation, with the Redding subterrane developing on their northern edge. By the Permian, faunal isolation of the evolving Klamath superterrane had increased enough to produce the McCloud fauna, which was isolated from the North America craton by a 2000–3000-km-wide ocean basin. The Klamath superterrane was still rotated 110° counterclockwise from Laurentia (Irwin and Mankinen, 1998) and it could have occupied either hemisphere. Judging from Devonian faunal affinities and assuming a conservative amount of motion, it most likely was still in the Northern Hemisphere. Paleogeographic reconstructions after Dalziel (1997) and Lawver et al. (2002).
reconstructions shows that the overall aspect of Yreka subterrane faunas is not North American—that is, Laurentian, the core of North America in late Precambrian through early Paleozoic time—but Pacific Rim. Many Yreka subterrane fossils exhibit biogeographic affinities to regions that once ringed Panthalassa, the proto-Pacific Ocean, and have now dispersed to other localities: Australia, Baltica, Siberia, Kazakhstan, and China (Table 3). Some “North American” affinities are to localities in accreted terranes that formed elsewhere, such as the Sierra Nevadan, Alaskan, and Appalachian terranes. However, many Yreka subterrane fossils also exhibit affinities to the miogeoclinal continental margin of Laurentia that bordered the Panthalassic Ocean (Table 3). The many unique species that are endemic to the Yreka subterrane, such as the early Paleozoic gastropods (Table 3), are more significant than previous authors have credited, as they indicate a high degree of faunal isolation. Talent et al. (1987) estimated that oceanic barriers 750–1000 km wide would have prevented dispersal of more than 50% of marine organisms, and a marginal sea 1500–1700 km wide would have prevented dispersal of almost all marine organisms, provided no migration paths crossed it. Species that are endemic to other accreted terranes as well as the Yreka subterrane, such as Llandeilo and Ashgill sphinctozoan sponges (Table 3), may indicate linear strings of tectonic elements such as island arcs and oceanic plateaus stacked up along active plate margins. Accreted terranes with faunal or geologic links to the Yreka subterrane include the Alexander and Farewell terranes of Alaska, the Sierra City mélange of the northern Sierra terrane in California,
25
and the Roberts Mountains allochthon of Nevada (cf. Soja, 1988; Blodgett et al., 2002; Wright and Wyld, 2006). Most of the Ordovician through Early Devonian fossils of the Yreka subterrane (Table 3) are from blocks within polygenetic mélange (Lindsley-Griffin et al., 1991). This observation does not negate their value as biogeographic indicators, but it does complicate their interpretation. Shallow-water limestones encased in deep-marine tectonized mélange matrix may have arrived in the trench environment as olistoliths from the overriding plate, or as tectonic fragments “bulldozed” from the subducting plate. Such mélange blocks may have moved from one biogeographic province to another, or even been translated through a succession of environments by passively riding on moving plates or by margin-parallel transform faulting. Highly endemic species may develop in reefs fringing isolated island chains before forcibly joining an accretionary complex and mixing with species from a different faunal province. Neoproterozoic Biogeography The identification of Ediacaran cyclomedusoids (LindsleyGriffin et al., 2006) in the Yreka subterrane of northern California (Fig. 1) is immensely important in clarifying the origin and history of the Klamath superterrane. As far as we know, these are the first reported Ediacaran cyclomedusoids from a Cordilleran accreted terrane. The Yreka subterrane cyclomedusoids are typical of Ediacaran biota, the first Metazoan body fossils that appeared suddenly at the end of the Precambrian, and which characterize the Ediacaran Period (e.g., Glaessner, 1971; Runnegar, 1992b; Conway Morris, 1993; McMenamin, 1998; Narbonne, 1998). Biogeography of the Ediacaran biota was examined by Waggoner (1999), who defined three data sets: one consisting of questionable genera, a second data set comprising the 49 genera that are best understood, and a third data set of 88 taxa that he considers to be intermediate between the other two. Waggoner (1999) applied parsimony analysis of endemism (PAE) and phenetic clustering to analyze cladistic similarities and differences between the three data sets. He tested these results with other techniques to select the most prudent conclusion, then plotted his results on five different paleogeographic reconstructions to test which tectonic model fitted the data best. Waggoner (1999) concluded that the best paleogeographic fit for Ediacaran biota is the reconstruction of Rodinia shown in our Figure 6. Three biogeographic groups have been recognized (Waggoner, 1999, 2003; Gehling, 2001): the White Sea assemblage, which includes fossils from Australia, northwest Canada, Siberia, and Baltica (Table 5); the Nama assemblage, which includes Laurentian (Mojave Desert and eastern British Columbia) and South China fossils; and the Avalon assemblage, which includes the Newfoundland and England (Charnwood Forest) fossils (Fig. 6). In Waggoner’s (1999) analysis, the Ediacaran Hills biota of Australia evidenced a strong correlation with biota of Baltica, NW Laurentia, and Siberia. He included the “Windermere fauna” of the Mackenzie Mountains localities (e.g., Hofmann
26
Realm, Region, or Province
TABLE 5. LATE PRECAMBRIAN AND PALEOZOIC BIOGEOGRAPHIC PROVINCES Biota Localities included
Ediacaran Period (600–542 Ma) White Sea assemblage Ediacaran biota Nama assemblage
Ediacaran biota
Ediacaran Hills; Baltica (including Finnmark); Siberia; northwestern Laurentia (Yukon); Yreka subterrane Namibia; Mojave Desert; South America
Avalon assemblage
Ediacaran biota
Charnwood Forest; Newfoundland; Carolina Slate Belt
Middle Cambrian (513–501 Ma) Columban or North Trilobites (isolated from American Province others, minimum internal barriers) Viking or EuropeanTrilobites (limited Mediterranean Province interchange with Altai Mtns. of China; Siberia) Tollchuticook or AsianTrilobites Australian Province Early Ordovician (488–472 Ma) Northern Ordovician Brachiopods, trilobites Region Balto-Scandian Region Southern Ordovician Region
Brachiopods, trilobites Brachiopods, trilobites
Siberian (Tungusian) Province
Brachiopods, trilobites
Middle Ordovician (472–461 Ma) North American Province Brachiopods, trilobites, conodonts Scoto-Appalachian Brachiopods, trilobites, Province conodonts Siberian (Tungusian) Brachiopods, trilobites Province Pacific Province Planktonic graptolites
Atlantic (European) Province Not named
Planktonic graptolites Sphinctozoan sponges
Late Ordovician (461–444 Ma) North American Province Brachiopods, trilobites Mediterranean (Bohemian) Province (locally mixed with Hiberno-Salerian) Hiberno-Salerian Province (locally mixed with Medit.) North Estonian Central Asian Province Not named
Brachiopods, trilobites
Brachiopods, trilobites Brachiopods, trilobites, corals Corals Sphinctozoan sponges
Reference
Waggoner, 1999, 2003; Gehling, 2001; this paper Waggoner, 2003; Gehling, 2001 Waggoner, 1999, 2003; Gehling, 2001
Laurentian miogeocline; northwestern Greenland; South America
Jell, 1974; Palmer, 1979
Western Europe; Turkey; Israel; Atlas Mtns. of Morocco; Appalachian and maritime-Canada terranes
Jell, 1974; Palmer, 1979
Altai Mtns. of China; South China; southeast Asia; Antarctica; Australia; Siberia; western Cordilleran and Alaskan terranes
Jell, 1974; Palmer, 1979; Blodgett et al., 2002
Laurentian miogeocline; Canadian Arctic; Greenland; northwestern Scotland and Ireland; Siberia; China; Kazakhstan Baltica; Norway; Scandinavia Argentina; southern Mexico; Appalachian and maritime Canada terranes; Wales; Scandinavia; Morocco; South China; Altai Mtns. of China; Australia Siberian Platform; Kazakhstan; southwestern Siberia
Jaanusson, 1979
Laurentian craton and miogeocline; Canadian Arctic; Greenland Appalachian terranes; Girvan Scotland; Siberia; Kazakhstan; Nevada (Copenhagen Fm.) Siberian Platform; Kazakhstan; southwestern Siberia Southeastern Australia; New Zealand; Texas (Marathon terrane); Cordilleran terranes; Canadian Arctic; Appalachian terranes; northeastern Siberia; China; Kazakhstan Scandinavia; Europe (Wales, England, Ireland, Belgium, Poland, Bohemia); N Africa Known only from Australia; Canadian Yukon; Alaska terranes (Alexander); and the Yreka subterrane
Jaanusson, 1979 Jaanusson, 1979
Jaanusson, 1979
Jaanusson, 1979 Jaanusson, 1979 Jaanusson, 1979 Jaanusson, 1979
Jaanusson, 1979 Rigby and Potter, 1986; Rigby et al., 1988
Laurentia; Canadian Arctic; Greenland; Altay Mtns.; northeastern Siberia; Siberian Platform Bohemia; Sweden; western Ireland and Scotland; Appalachian and maritime Canada terranes (Magog, Percé)
Jaanusson, 1979
Siljan district; Sweden; northern England; Ireland; southwestern Siberia Northern Estonia; Lithuania; Oslo Region
Jaanusson, 1979
Kazakhstan; Ural Mtns.; China Known only from Australia; Canadian Yukon; Alaskan terranes (Alexander); and the Yreka subterrane
Jaanusson, 1979
Jaanusson, 1979 Jaanusson, 1979 Rigby and Potter, 1986; Rigby et al., 1988 Continued
27
Realm, Region, or Province
TABLE 5. LATE PRECAMBRIAN AND PALEOZOIC BIOGEOGRAPHIC PROVINCES (continued) Biota Localities included
Silurian (444–416 Ma) North Silurian Realm, North Atlantic Region, European Province
Shelly faunas in CO2rich, warm water
Appalachian accreted terranes of Acadia (eastern New England–New Brunswick, Nova Scotia); Europe to western edge of Russian Platform; western Baltica; northernmost Africa (Guinea); Merida terrane of Venezuela; Cordillera Oriental of northeastern Mexico; Marmara Sea area of western Turkey Laurentian craton and miogeocline (Utah, central Nevada, Texas, Chihuahua); Andean accreted terranes Cordilleran accreted terranes; Ural Mtns. terranes; Siberian Platform; Kazakhstan; South China; India; Arabia; Carnic Alps–Karawanken Eastern Australia; China; northern Vietnam; Afghanistan; FSU central Asia
Boucot, 1988; Boucot and Blodgett, 2001
Shelly faunas
Northernmost China–Mongolia; Okhotsk terranes
Boucot and Blodgett, 2001
Shelly faunas
Iranian terranes
Boucot and Blodgett, 2001
Southern two-thirds of South America; Antarctica; South Africa; Ghana; Guinea; Falkland Islands; Argentine Precordillera (brachiopods closer to Eastern Americas Realm)
Boucot, 1988; Benedetto and Sánchez, 1996; Blodgett et al., 1990; Boucot and Blodgett, 2001
Derived from Uralian-Cordilleran Region and European Province of the Silurian
Boucot, 1988; Blodgett et al., 1990; Boucot and Blodgett, 2001
Laurentian miogeocline (western North America) from Nevada to Canadian Arctic Islands
Boucot and Blodgett, 2001; Blodgett et al., 2002
Alaskan accreted terranes; northwestern Canada (Laurentia) Europe west from western edge of Russian Platform; North Africa from Guinea north; Arabia; Acadian accreted terranes of New England, New Brunswick, Nova Scotia New Zealand terranes
Blodgett et al., 1990
North Atlantic Region, North American Province
Shelly faunas
Uralian-Cordilleran Region, undivided
Shelly faunas
Uralian-Cordilleran Region, Sino-Australian Province Uralian-Cordilleran Region, Mongolo-Okhotsk Province Uralian-Cordilleran Region, Iran Province?
Shelly faunas
Early to Middle Devonian (416–391.8 Ma) Malvinokaffric Realm or Shelly faunas, rugose Afro-South American corals (in cool Realm temperate to cold high latitude, low diversity) Early to Middle Devonian (416–385.3 Ma) Old World Realm, Brachiopods, American Province gastropods, rugose corals, trilobites (warmer than Mediterranean Region and Eastern Americas Realm, highest diversity) Old World Realm: Nevadan brachiopods, gastropods, corals, Cordilleran Region (includes Western mixed with Canada Province) Appalachian taxa Cordilleran Region: Gastropods Alaska-Yukon Province Old World Realm: Shelly faunas Rhenish-Bohemian Province Old World Realm: New Zealand Province Old World Realm: Uralian and Mongolo-Okhotsk Province Old World Realm: Tasman Province Old World Realm: North African Province Eastern Americas Realm
Reference
Emsian brachiopods Lochkovian–Givetian shelly faunas Lochkovian–Emsian shelly faunas Rugose corals Lochkovian–Givetian brachiopods, gastropods, rugose corals (warmer than Mediterranean Region, moderate diversity)
Eastern and western Ural Mtns.; Siberia; Tien Shan; southeastern Kazakhstan; Farewell, Alexander, Arctic terranes of Alaska Eastern Australia Northwestern Africa Laurentia, derived from Old World Realm, American Province
Boucot and Blodgett, 2001
Boucot and Blodgett, 2001
Boucot and Blodgett, 2001
Boucot, 1988; Boucot and Blodgett, 2001
Boucot, 1988; Boucot and Blodgett, 2001 Boucot, 1988; Boucot and Blodgett, 2001 Boucot, 1988; Boucot and Blodgett, 2001 Oliver and Pedder, 1979; Boucot, 1988 Boucot, 1988; Blodgett et al., 1990; Boucot and Blodgett, 2001
Continued
28
Realm, Region, or Province
Lindsley-Griffin et al. TABLE 5. LATE PRECAMBRIAN AND PALEOZOIC BIOGEOGRAPHIC PROVINCES (continued) Biota Localities included
Early to Middle Devonian (416–385.3 Ma) (continued) Eastern Americas Realm: Lochkovian–Givetian Appohimchi Subprovince shelly faunas (Appalachian Basin) (subtropical to warm temperate climate, parallel to paleolatitude) Eastern Americas Realm: Pragian–Givetian Colombian Subprovince brachiopods, gastropods (moderate to cool temperate climate) Eastern Americas Realm: Emsian–Eifelian, early Amazon Subprovince Givetian gastropods, mixture of Eastern Americas Realm brachiopods and Mediterranean Realm trilobites (moderate to cool temperate climate) Eastern Americas Realm: Brachiopods, Michigan Basin—Hudson gastropods, Bay Subprovince tetracorals; Givetian rugose corals (tropical to subtropical) Nevadan Subprovince: Early Pragian–early shifted from Old World Emsian brachiopods, Realm in Lochkovian to gastropods, Eastern Americas Realm– tetracorals Appalachian in Pragian– early Emsian, then back to Old World Realm in Eifelian South China Region Emsian-Eifelian shelly faunas Carboniferous (359.2–299.0 Ma) Not named Shelly faunas Permian (Early) (299–270.6 Ma) McCloud Province Brachiopods, corals, fusulinids Central American Shallow-water faunas Province Tethyan Province Shallow-water faunas Eastern Panthalassa Endemic shelly faunas
Reference
Laurentian: Michigan Basin east to Montreal and Chihuahua north to Hudson Bay; includes Appalachian accreted terranes
Johnson and Dasch, 1972; Boucot, 1988; Blodgett et al., 1990; Boucot and Blodgett, 2001
Perija terrane, northern Andes Mtns.
Boucot, 1988; Boucot and Blodgett, 2001
Amazon basin; possibly Parnaiba Basin
Boucot, 1988
Laurentian: midcontinent North America north to the Hudson Bay area (area of mixing between Eastern Americas Realm and Old World Realm for Eifelian gastropods)
Oliver and Pedder, 1979; Koch, 1981; Boucot, 1988; Blodgett et al., 1990
Central Nevada and small area in southern California: many endemic species = Laurentian margin inboard of accreted terranes
Boucot, 1988; Blodgett et al., 1990; Boucot and Blodgett, 2001
South China; northern Vietnam (north of Red River Fault); western Kunlun
Boucot and Blodgett, 2001
Poorly understood; peri-Gondwanan faunas seem to be distinct from some isolated endemic groups
Ross, 1979
Klamath superterrane; northern Sierra; Quesnellia; Stikinia, Wrangellia Laurasian and Gondwanan miogeoclines
Stevens et al., 1990; Belasky et al., 2002 Ross and Ross, 1979
Europe; Ural Mtns. New Zealand; China; North and South American Cordilleran accreted terranes
Ross and Ross, 1979 Ross and Ross, 1979; Belasky et al., 2002
et al., 1985; Narbonne, 1994) in the NW Laurentia group. We include the Yreka subterrane cyclomedusoids in the White Sea assemblage because of their strong similarity to the Finnmark locality of Farmer et al. (1992). In addition, Yreka subterrane fauna exhibit many biogeographic affinities to the Mackenzie Mountains and Canadian Arctic, Siberia, and Baltica throughout the early Paleozoic, suggesting that these biogeographic connections continued at least into the Middle Devonian (Table 3). The recently discovered Ediacaran fossils in the Salient Mountain area of eastern British Columbia (Hofmann and Mountjoy, 2001) are grouped with the Nama assemblage by Waggoner
(2003) and probably represent the Neoproterozoic Laurentian continental margin. Using the paleomagnetic pole determined by Mankinen et al. (2002) for the Ediacaran Trinity ophiolite as a proxy for the future location of both subterranes during the breakup of Rodinia, we can hypothesize that the Trinity-Yreka subterranes formed in one of three possible rift basins (Fig. 6). Although locations between Australia and Antarctica or Australia and Laurentia at ~7°N latitude fit the paleomagnetic data, faunal affinities fit better with a position between northern Laurentia and Siberia (Fig. 6). It may be significant that ophiolites of Ediacaran to Early Cambrian age
Paleogeographic significance of Ediacaran cyclomedusoids occur in the Altai fold belt of Mongolia and within the Kazakhstan block (Talent et al., 1987), indicating that these fragments underwent oceanic rifting at the same time the Trinity ophiolite was forming. Both the Altai and Kazakhstan blocks share many biogeographic affinities with the Yreka subterrane (Table 3; Potter et al., 1990a), and the Yreka and Trinity subterranes may have been part of the same island chain from which parts of Altai and Kazakhstan were derived. For the small Avalon assemblage of Ediacaran biota (Fig. 6), two localities are within terranes accreted along the eastern edge of Laurentia during the Acadian orogeny. According to most paleogeographic reconstructions (cf. Dalziel, 1997; Lawver et al., 2002; Scotese, 2002) these terranes represent oceanic plateaus or volcanic arc assemblages that formed outboard of Baltica within the Rheic Ocean. The Charnwood Forest locality is in Leicestershire, Great Britain (Glaessner, 1979, p. A83), which likely would have occupied the other side of the same ocean basin. The Avalonian localities remained isolated from the future location of the Yreka-Trinity subterranes throughout the late Neoproterozoic, first by the continent of Rodinia and later by the continent of Pannotia (cf. Figs. 6 and 7). The Nama assemblage, comprising Africa, South America, SW Laurentia, and Antarctica, also lay on the opposite side of Rodinia and Pannotia relative to the YrekaTrinity subterranes in the late Proterozoic (Figs. 6 and 7). Cambrian Biogeography The Yreka subterrane has yielded no Cambrian fossils (Table 3). We assume that the subterrane would not have moved far from its hypothetical position in the Ediacaran Period (Fig. 7), and that its geological and paleontological similarities to Alaskan terranes are as valid for the Cambrian Period as for the rest of the Paleozoic. Thus, we propose to use Cambrian faunas of Alaska as proxies for the Cambrian biogeography of the Yreka subterrane (Fig. 8). Middle Cambrian trilobites of the Nixon Fork subterrane of the Farewell terrane, Alaska, are most similar to those from eastern Siberia and the Siberian fold belt (Palmer et al., 1985; Blodgett et al., 2002). The most likely position for the Yreka-Trinity subterranes in the Cambrian would be in the Uralian seaway between Laurentia and Siberia, possibly aligned along developing active-margin island chains providing a migration path with Siberia (Fig. 8). Paleomagnetic data suggest a low latitude (Mankinen et al., 2002). Because Ordovician graptolites of the Yreka subterrane show “Pacific” (Australian-American) affinities (Table 3), some interchange with Australia may have been possible for far-traveled genetic material. Other permissible locations for the Yreka subterrane would be in the belt between 7°N and 7°S extending from Siberia to North China, South China, and Australia (Fig. 8), although we have no Cambrian faunal evidence to support this hypothesis. Ordovician Biogeography Early Ordovician fossils in the Yreka subterrane are known only from the Facey Rock limestone block in the Facey-Duzel mélange (Fig. 2); they are cosmopolitan (Table 3: conodonts).
29
Middle and Late Ordovician fossils are more helpful because they are relatively provincial (Table 4). Because not much paleogeographic change occurred between the Early and Late Ordovician (Lawver et al., 2002), Middle and Late Ordovician biogeographic links are shown on a single paleogeographic reconstruction for 460 Ma (Fig. 9). Species endemic to the Yreka subterrane indicate isolation for at least some species and strongly suggest that both the Yreka and the Trinity subterranes were imperfectly linked to the other tectonic elements. However, affinities with the Laurentian margin also indicate persistent links between the Yreka-Trinity subterranes and Laurentia. Such a mixed biogeographic message with high endemism is typical of oceanic island chains (Newton, 1987). The abundant limestone blocks in the Gregg Ranch Complex suggest an equatorial to temperate latitude, consistent with the paleolatitudes determined by Mankinen et al. (2002). Ordovician fossils from the Yreka subterrane are derived from blocks in mélange, so they may have transferred from one tectonic plate to another during the accretionary process. However, their biogeographic affinities are useful because they demonstrate links to a variety of tectonic elements in the Panthalassic equatorial to temperate ocean (Fig. 9), including Australia, China, Kazakhstan, Baltica, Siberia, and NW Europe (Tables 3 and 5). Laurentian affinities include both accreted terranes (Alaska, California, Maine) and the outer edges of Laurentia proper (Table 3). Genetic exchange would likely have been facilitated by an eastto-west equatorial current connecting these tectonic elements. Because Laurentia straddled the equator (Fig. 9) it likely would have generated gyres similar to the present-day South Equatorial Current and North Pacific Current, facilitating migration along both miogeoclinal margins of Laurentia. For example, the Pacific province for Middle Ordovician planktonic graptolites (Table 5) includes a number of accreted terranes that all shared the same Panthalassic ocean basin, including the White Mountains of Alaska, part of the Farewell terrane, and Girvan, Scotland, then part of the NW Europe terrane (Fig. 9). Middle Ordovician brachiopods, conodonts, and trilobites (Table 3) also show ties to the White Mountains and Girvan, as well as to other wandering terranes such as Kazakhstan, eastern Appalachia, and Baltica. Kazakhstan consists of ensimatic terranes that developed in the early Paleozoic, and then were trapped between Europe and Siberia during a complex series of collisions that lasted ~50 m.y., forming the Uralian deformed belt (Talent et al., 1987). The blocks on either side of the Uralian suture zone were faunally segregated from each other at least through the Early Devonian, and the Uralian belt probably did not finish forming until the late Carboniferous, ca. 320 Ma (Gratsianova et al., 1988; Torsvik and Cocks, 2004). In the Middle Ordovician, eastern Appalachian terranes lay somewhere between NW Europe and southern Laurentia (Fig. 9), and did not accrete to the Laurentian margin until the middle or late Paleozoic (Hatcher et al., 1989; Osberg et al., 1989). The Middle Ordovician Effna Limestone of Virginia (Table 3: Llanvirn trilobites) is part of the Laurentian miogeocline in the Appalachian Valley and Ridge province, which was
30
Lindsley-Griffin et al.
deformed by a series of terrane amalgamations and accretions beginning in the Middle to Late Ordovician and culminating in the Late Mississippian to Permian (e.g., Drake et al., 1989; Hatcher et al., 1989; Osberg et al., 1989). We assume that the “Virginia-Tennessee-Alabama” locality (Table 3: Llandeilo brachiopods) cited by Potter et al. (1990a) is also in the Laurentian miogeocline. A number of Late Ordovician corals are known only from the Yreka subterrane or from both the Yreka subterrane and the Montgomery Limestone of the Shoo Fly complex (Table 3: Ashgill solitary rugose corals, tabulate corals), although other corals are linked to Appalachian and Uralian terranes as well as to Australia, China, Europe, and Siberia (Potter et al., 1990a, 1990b). The Shoo Fly complex is part of the northern Sierra Nevada terrane (Harwood, 1992); the Montgomery Limestone is within the Sierra City mélange, which occupies the structurally highest thrust sheet of the Shoo Fly Complex and is a polygenetic mélange formed in an Ordovician-Silurian accretionary complex (e.g., Hannah and Moores, 1986; Saleeby et al., 1987; Girty et al., 1996; Potter et al., 1990b). The Sierra City mélange also contains a Neoproterozoic (600 ± 10 Ma) plagiogranite block (Saleeby, 1990), which might have been derived from the same continental source as the plagiogranite blocks in the Yreka subterrane’s Skookum Gulch mélange. Middle to Late Ordovician sphinctozoan sponges (Tables 3 and 5) are particularly interesting, because they occur only in Australia, the Canadian Yukon, Alaskan terranes, and the Yreka subterrane (Rigby and Potter, 1986; Rigby et al., 1988; Potter et al., 1990a). Recently identified Late Ordovician sponges from limestone conglomerate clasts in the Lower Devonian Karheen Formation of the Alexander terrane are conspecific with a species from Horseshoe Gulch (Rigby et al., 2005). Late Ordovician shelly faunas of the Nixon Fork subterrane (Farewell terrane of Alaska) are similar to Yreka subterrane faunas, as well as to Siberia and western Laurentia (Blodgett et al., 2002). Figure 9 shows several possible locations for the Yreka-Trinity subterranes in the Ordovician, but the likeliest positions are in the relatively low latitude part of the basin between Laurentia, Siberia, Kazakhstan, Baltica, and NW Europe—the Uralian Seaway of Soja and Antoshkina (1997). Other permissible locations consistent with faunal, geologic, and paleomagnetic evidence would lie between Kazakhstan, North and South China, and Australia (Fig. 9). Silurian Biogeography Silurian faunas (Table 3) are relatively sparse in the Yreka subterrane, with no known fossils of Early Silurian age. Middle and Late Silurian fossils of the Yreka subterrane are all from mélange blocks. They show biogeographic links to many of the same tectonic elements as the Ordovician fossils (Fig. 10 and Table 3). The Yreka subterrane tabulate corals are similar to those of “Altai” (Table 3; Wenlock and Ludlow), the Altai fold belt of Mongolia, northwest China (Potter et al., 1990a). Talent et al. (1987) consider North China to have been adjacent to south-central Siberia in the Silurian, but the reconstruction we
are using in Figure 10 places Kazakhstan between Siberia and North China. Yreka subterrane shelly faunas and corals also show affinities to Siberia and the Uralian fold belt (Table 3) so either interpretation is consistent with our data. “Appalachia” (Table 3: Silurian tabulate corals) is most likely the Appalachian accreted terranes, which were not in North America (Laurentia) at the time, but out in the bordering ocean between NE Laurentia (present coordinates) and NW Europe (Fig. 10). Girvan, Scotland, is in NW Europe. Late Silurian brachiopods from the Gregg Ranch Complex are typical of the “Uralian-Cordilleran region” (Potter et al., 1990a, p. 67). The “Uralian-Cordilleran region” (Tables 3 and 5) includes the Cordilleran accreted terranes, Uralian terranes, Siberian Platform, Kazakhstan, South China, India, Arabia, and the Carnic Alps-Karawanken (Boucot and Blodgett, 2001). Silurian gastropods described as “Laurentian” by Rohr (1980) are actually linked to parts of Alaska and Scotland now recognized as accreted terranes. Middle and Late Silurian tabulate corals from Horseshoe Gulch and the Gregg Ranch Complex are similar to those from Appalachian terranes, Europe, the Uralian fold belt, Siberia, China, and Australia. Thus, in the Silurian the Yreka-Trinity subterranes most likely were located in the Northern Hemisphere, intermediate between the 7° Ediacaran paleopole and the 31° Devonian paleopole of Mankinen et al. (2002), somewhere within the ocean basin surrounded by Siberia-Baltica-Kazakhstan-China (Fig. 10). Although a Southern Hemisphere location between Laurentia and Gondwana (Fig. 10) would be consistent with the paleomagnetic data, such a location is not supported in the absence of faunal affinities to southern Gondwana and Laurentia. Devonian Biogeography Devonian provinces and subprovinces (Table 5) are different for different biota, and some overlap with each other, an observation attributed to “boundary mixing” by Boucot (1988), and easily explained by variations in how different reproductive strategies exploit the available tectonic elements. The Early Devonian is highly provincial, but Middle and Late Devonian became progressively more cosmopolitan (Table 4). Early Devonian biogeography. Early Devonian brachiopods from the Gregg Ranch Complex are “Cordilleran” (Boucot and Potter, 1977), that is, similar to faunas in Nevada and northern Canada and thus consistent with an origin along the western edge of Laurasia (Euramerica of Scotese, 2002). Early Devonian conodonts from the Gregg Ranch Complex either are cosmopolitan or occur in Yreka subterrane and Nevada only. Early Devonian corals from the Gregg Ranch Complex are also Laurasian, although some species are unique to the Yreka subterrane or to the YrekaTrinity subterrane and Nevada (Table 3). The Gazelle Formation of Lindsley-Griffin et al. (1991) is dated only by conodonts of Emsian, Early Devonian age, described as “Cordilleran” (Boucot et al., 1974; Potter et al., 1990a). Thus, Yreka subterrane endemism increased somewhat in the Early Devonian (cf. Silurian and Early Devonian, Table 3) but links to western Laurasia and to Cordilleran accreted terranes persisted, suggesting island chains
Paleogeographic significance of Ediacaran cyclomedusoids or an oceanic plateau lying some distance from the Laurasian continent but able to exchange some genetic material. Middle Devonian biogeography. The Yreka subterrane has yielded no fossils clearly identified as younger than Early Devonian. Geologic evidence (Lindsley-Griffin et al., 2006) suggests that by the Middle Devonian it was buried under lavas of a nascent volcanic arc that forms the base of the Redding subterrane (Fig. 1), developing over the composite Forest Mountain– Yreka–Trinity subterrane after they amalgamated. The early Middle Devonian collision probably uplifted the composite assemblage, producing unconformities over a wide area (Boucot et al., 1974) before post-accretion subsidence allowed deposition of the Copley-Balaklala and Kennett Formations on the YrekaTrinity basement. Deformation of all Redding subterrane strata occurred simultaneously in the Late Jurassic or Early Cretaceous, presumably as a consequence of the accretion of the Klamath superterrane to the North American continental margin. Thus, the Redding subterrane served as a passive recorder of events from mid-Devonian to Late Jurassic time. The Middle Devonian orogeny predates the Antler (Devonian-Mississippian) and Sonoma (Permo-Triassic) orogenies to the east, indicating that active tectonic elements must have continued to block the eastern Klamaths from the North American core at least until the Triassic. Some active tectonic elements may have persisted until the accretion of the Klamath Mountains to North America was complete in the Early Cretaceous (Mankinen et al., 1988). In the Redding subterrane, the Copley Greenstone and Balaklala Rhyolite interfinger with, and grade rapidly upward into, marine shales and graywackes of the Kennett Formation. A placoderm fish plate (Boucot et al., 1974) dates the Balaklala Rhyolite as Middle Devonian, and brachiopods (Boucot and Potter, 1977) date the Kennett Formation as Middle Devonian (Table 3). Biogeographic affinities (Table 3) for these units at the base of the Redding subterrane either are cosmopolitan (placoderm fish plate, conodonts) or suggest limited communication with Laurentia. The conodont Polygnathus kennettensis (Table 3: late Eifelian) is particularly significant because it is known only from the Redding subterrane (Potter et al., 1990a), thus indicating a high degree of faunal isolation for this species. The Middle Devonian tabulate corals (Table 3) are similar to those of Eureka, Nevada, and of Suplee, Oregon (Potter et al., 1990a). We assume the Nevada reference is to the lower-plate miogeoclinal carbonates rather than the upper-plate Roberts Mountain allochthon; thus, it indicates a connection to the Laurentian miogeocline. However, the Suplee and Suplee Butte quadrangles of central Oregon lie within the Blue Mountains, which are within the accretionary belt of the Cordilleran orogen (cf. Mankinen and Irwin, 1990; Burchfiel et al., 1992), well outboard of other accreted terranes that would have separated it from the Middle Devonian Laurentian continental margin (Fig. 11). Faunal affinities to Siberia and the Uralian fold belt that are seen in the Silurian seem to be lacking in the Devonian faunas, suggesting that communication with those elements had been shut off. Thus, the Devonian Forest Mountain–Yreka–Trinity–Redding nucleus of the emerging
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Klamath superterrane was most likely located off the western edge of Laurentia where it had limited communication with carbonate reefs of Nevada as well as with other unaccreted offshore terranes (Fig. 11). A Northern Hemisphere paleolatitude is more likely, considering that other early Paleozoic locations were most likely in the Northern Hemisphere. Permian Biogeography Late Devonian and Carboniferous faunas are cosmopolitan (Tables 4 and 5) and are of little biogeographic use. Ross and Ross (1979) defined two distinct biogeographic provinces for the Permian (Table 5), the Central American Province and the Tethyan Province. Ross and Ross (1979) separated a number of endemic shelly faunas into an eastern Panthalassic group; these now occupy accreted terranes of New Zealand, China, the South American Cordillera, and the North American Cordillera (Table 5). Early Permian faunas of the McCloud Limestone (Redding subterrane) comprise a distinct biogeographic province (Table 5) that extends from the Northern Sierra and Eastern Klamath terranes through Quesnellia and Stikinia to Wrangellia (Stevens et al., 1990; Belasky and Runnegar, 1994; Belasky et al., 2002). Belasky et al. (2002) concluded that during the Permian, these terranes formed a linear belt, probably an island chain, isolated from North America by a 2000–3000-km-wide ocean basin. If this McCloud fauna belt had been tied to North America at its southern end, as suggested by Colpron et al. (2007), the Northern Sierra and Eastern Klamath faunas would be a mixture of Tethyan and McCloud elements with endemism increasing along the belt toward Wrangellia. However, Belasky et al. (2002) concluded that all of the terranes in the McCloud fauna belt were equally distant from the North American craton. The McCloud fauna is highly dissimilar to the Eastern Panthalassa and Tethyan faunas (Belasky et al., 2002); thus, the Klamath superterrane in the Permian was located at some unknown distance from Laurasia (Fig. 12) and was isolated from it by environmental or tectonic elements that blocked migration of shallow-water faunas. A northern hemisphere location is more likely than a southern position, considering earlier locations and assuming the Klamath superterrane was not drifting rapidly. SUMMARY AND CONCLUSIONS The Yreka subterrane and Trinity subterrane record ~180 m.y. of active margin events somewhere in Panthalassa, the Proto-Pacific Ocean. Yreka subterrane features include late Precambrian? to Cambrian? greenschist metamorphism and mélange formation, Late Ordovician blueschist metamorphism and mélange formation, Silurian and Siluro-Devonian trench sedimentation and greenschist metamorphism, and an Early Devonian trench-slope basin resting on coeval mélange. Trinity subterrane events include Ediacaran rifting of an emerging arc to form an ophiolite, Cambrian? amalgamation of disparate oceanic elements, Early Ordovician plutonism, and the development of a supra-subduction zone ophiolite on the older
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composite oceanic basement. Reevaluating Paleozoic biogeographic data for these two subterranes and for similar terranes such as the northern Sierra Nevada (California) and the Alexander and Farewell (Alaska) terranes, we conclude that the Yreka subterrane and Trinity subterrane were located throughout the early Paleozoic in the part of Panthalassa bordered by Australia, NW Laurentia, Siberia, China, Baltica, and the Uralian terranes. Biota suggest considerable isolation, based on the high degree of endemism, as well as some biotic communication with each of these elements, based on persistent links to other elements (Table 3), suggesting an oceanic plateau or island chain well outboard of other tectonic elements. Paleomagnetic data, biogeographic affinities, and regional relationships indicate that the Alexander terrane of Alaska formed well outboard (west in present coordinates) of the Cordilleran margin and close to Siberia and Baltica in the Neoproterozoic through Late Silurian–Early Devonian (e.g., Bazard et al., 1995; Gehrels et al., 1996; Soja and Antoshkina, 1997; Blodgett et al., 2002; Wright and Wyld, 2006). Most detrital zircons in the Alexander terrane were derived by erosion of igneous rocks within the terrane itself, but the Lower Devonian Karheen Formation has yielded detrital zircons with apparent ages of 3.0–1.0 Ga, indicating continental source rocks ranging in age from late Archean to middle Proterozoic (Gehrels et al., 1996). The seven age clusters for the Alexander terrane zircons (Table 1) are distinctively different from those of the Antelope Mountain Quartzite, suggesting that they did not share source areas. Furthermore, the Alexander terrane 1.58–1.48 Ga zircons could not have been derived from Laurentia because no source areas of the appropriate age are known, although they could have been derived from the Scandinavian shield of Baltica (Bazard et al., 1995). The 1.8–1.0 Ga zircons are consistent with a provenance from either the Tasman orogen-Adelaide geosyncline or the Scandinavian shield (Gehrels et al., 1996). On the basis of biogeographic affinities between Alaskan and Uralian faunas, together with paleomagnetic and detrital zircon evidence, Soja and Antoshkina (1997) postulated that the Alexander terrane was located in the Uralian Seaway between Laurentia-Baltica and Siberia during the Silurian and Devonian. Wright and Wyld (2006) have pointed out that the Shoo Fly complex must have formed between a continental source to the west (present orientation) and a magmatic arc to the east, with the Roberts Mountains allochthon separating it from the Laurentian margin, and that its detrital zircons are similar to those of the Alexander terrane. They concluded that the arc and continental fragments that sourced it have been translated away, and it was nowhere near Laurentia in the early Paleozoic. Biogeographic data support this interpretation and place the Yreka subterrane and Trinity subterrane in an island chain or other intraoceanic location near enough to the Alexander terrane for limited faunal migration. Siberia and Kazakhstan are likely sources of the Shoo Fly continental and arc debris. Cambrian through Devonian strata of the Roberts Mountains allochthon were derived from a continental fragment out-
board of Laurentia that has been translated away, according to its sedimentology (Ketner, 1991) and detrital zircons (Wright and Wyld, 2006). These strata contain detrital zircons very similar to the Shoo Fly complex (Wright and Wyld, 2006) and may also have been derived from Siberia and Kazakhstan. Wright and Wyld (2006) interpreted the detrital zircon evidence for the Yreka, Shoo Fly, and Alexander terranes as indicating a relationship to the peri-Gondwanan terranes, including the Avalon, Carolina, and Suwannee terranes of eastern North America as well as terranes of western Africa, Spain, Britain, and France. They consider the Alexander terrane a piece of one of these periGondwanan terranes, and the Sierra City mélange of the Shoo Fly terrane as being derived from it. The other three thrust sheets of the Shoo Fly terrane appear to have different sources of detrital zircons; Wright and Wyld (2006) considered the Amazon craton to be a likely source, with the amalgamation of the Shoo Fly terrane occurring during the Caledonian orogeny (Late Silurian–Early Devonian). They include most strata of the Roberts Mountains allochthon in this same setting but consider that the lower Vinini Formation was more likely derived from the western Laurentian miogeocline. They consider the Yreka and Trinity subterranes to be closely linked to the Alexander terrane and Shoo Fly complex, as well as to the Roberts Mountains allochthon. We agree with Wright and Wyld (2006) that the Yreka and Trinity subterranes were not derived from Laurentia, nor were they close to Laurentia when they originated. We also agree that the geologic links between the Alexander terrane, Shoo Fly complex, and the Yreka-Trinity subterrane are strong. These links are supported by some faunal evidence, although the detrital zircon record for each terrane is different enough to suggest that they were not contiguous. However, the biogeographic evidence does not support connections with the Amazon craton, Avalonia, or West Africa, and only weakly supports ties to parts of Europe other than Baltica. We suggest that the sources of Antelope Mountain continental detritus more likely were elements identified by the biogeographic links discussed above, or were tectonic elements that are not yet recognized or not preserved. The Yreka subterrane and Trinity subterrane amalgamated together when the newly recognized Forest Mountain subterrane collided with them ca. 380 Ma, producing lavas with a probable paleolatitude of 31°N. The superjacent Redding subterrane began to develop over the Forest Mountain subterrane–Yreka subterrane– Trinity subterrane core in mid-Devonian (post-Emsian) time and continued to develop through the Jurassic as younger terranes successively collided with the Forest Mountain–Yreka–Trinity– Redding nucleus. The Yreka subterrane and Trinity subterrane, as part of the Klamath superterrane, did not attach to North America until the Early to Middle Cretaceous, according to paleomagnetic evidence of rotation and the age of overlying deposits. ACKNOWLEDGMENTS We gratefully acknowledge the many years of hard work by the numerous paleontologists cited herein. Without their dedication
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The Geological Society of America Special Paper 442 2008
Silurian-bearing terranes of Alaska Constance M. Soja* Department of Geology, Colgate University, Hamilton, New York, 13346, USA ABSTRACT Silurian rocks in Alaska have been identified in 12 accreted terranes and in the Tatonduk-Nation River area of east-central Alaska, which represents part of autochthonous North America. Most of the terranes are in situ or structurally imbricated portions of the North American (or Siberian) continental margin. An exception is the Alexander terrane of southeastern Alaska, which originated as an offshore island arc. Discontinuously exposed and (or) highly altered sequences have precluded detailed investigations of Silurian rocks in most parts of Alaska, but reconnaissance-level studies reveal that graptolitic shales of turbidite or hemipelagic origin record deepwater or “shale out” conditions west or north of the ancient continental margin of North America. Platform carbonates are also exposed in many areas and are particularly well represented in southwestern (Nixon Fork subterrane of Farewell terrane) and southeastern (Alexander terrane) Alaska, indicating that much of Alaska resided close to the paleoequator in the Silurian. Subtidal stromatolite reefs in southwestern and southeastern Alaska that are similar to those in Salair and the Ural Mountains, Russia, indicate a paleobiogeographic connection between these two parts of Alaska, Siberia, and eastern Baltica via the Uralian Seaway in the Late Silurian. Deposition of vast accumulations of red beds and other siliciclastic rocks beginning in the Late Silurian suggests that parts of Alaska may have been affected by late stages in Caledonian orogenesis and (or) early stages in the Ellesmere orogeny during formation of the Laurussian landmass. Keywords: Silurian, terranes, Alaska, Laurentia, Baltica, Siberia. INTRODUCTION AND GENERAL SETTING
most Silurian rocks must be understood within the context of accreted, tectonostratigraphic terranes. More than 50 terranes and subterranes have been identified in various parts of Alaska, representing pieces of continental and oceanic crust, island arcs, fragments of unknown origin, and composite (amalgamated) terranes that were accreted to North America beginning in the Mesozoic (Coney et al., 1980; Plafker and Berg, 1994a). Although terrane terminology in Alaska remains in a state of
In the northwestern part of North America, Alaska comprises an area comparable to ~40% of the onshore conterminous United States, stretching across more than 50 degrees of longitude with maximum east-west dimensions extending 3600 km from the Alaskan mainland to the western tip of the Aleutian arc (Plafker and Berg, 1994a). Within this vast area, the geology of
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*
Soja, C.M., 2008, Silurian-bearing terranes of Alaska, in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 39–50, doi: 10.1130/2008.442(02). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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flux, as does agreement on paleogeographic reconstructions of terranes, Silurian rocks occur in 12 terranes extending from the Arctic Alaska terrane of the North Slope to the Alexander terrane in the southeastern panhandle (Fig. 1). As shown in Figure 1, Silurian rocks also form part of autochthonous North America in the Tatonduk-Nation Arch Region but are absent, or have yet to be identified, in large portions of Alaska. Debates about terrane origins persist. With the exception of the Alexander terrane, all Silurian-bearing crustal fragments in Alaska appear to be (1) in situ and either largely undeformed or structurally imbricated fault slices of cratonic North America (e.g., Seward terrane, Porcupine terrane); (2) displaced portions of the ancient continental margin of North America or Siberia (e.g., Farewell terrane [including Nixon Fork, Dillinger, and Mystic subterranes], Arctic Alaska terrane, Minchumina terrane, and York terrane), as suggested by the affinities of fossils and by stratigraphic continuity or gradual facies transitions between contiguous areas (Dillon et al., 1987; Dover, 1994; Moore et al., 1994; Blodgett and Boucot, 1999; Blodgett et al., 1999, 2002); or (3) an accreted microcontinent (portions of Arctic Alaska, Seward, York, and Farewell terranes) that was positioned in the early to mid-Paleozoic along the Uralian Seaway but separated by deep ocean from Siberia and northern Laurentia (Dumoulin et al., 2000, 2002). The Alexander terrane represents a Proterozoic (?) to early Paleozoic island arc
that was in proximity to Laurentia-Baltica-Siberia in the Late Silurian–Early Devonian and was accreted to North America probably in the late Mesozoic–early Cenozoic (Coney et al., 1980; Gehrels and Berg, 1994; Soja, 1994, 1996; Bazard et al., 1995; Butler et al., 1997; Soja and Antoshkina, 1997, 1998; Blodgett and Boucot, 1999; Soja et al., 2000; Blodgett et al., 2002; Newton et al., 2002; Antoshkina et al., 2004; Antoshkina and Soja, 2006). Figure 1 in Soja and Krutikov (this volume) illustrates the paleogeographic setting of the Alexander terrane in the Late Silurian. An abundance of carbonate deposits in most Silurian sections of Alaska implies deposition in subtropical-tropical zones. Thick carbonate deposits and associated barrier reefs, calcareous green algae, an absence of evaporites, and Devonian gastropods that are highly ornamented are evidence that parts of Alaska had an equatorial location, as is particularly well demonstrated in southwestern and east-central Alaska (Decker et al., 1994). Although the paleogeographic setting of southeastern Alaska remains in dispute, an extraordinary thickness of limestones (~1000–3000 m) suggests the Alexander terrane was also located close to the equator under conditions favorable for widespread lime precipitation in the Silurian and through much of the Paleozoic. Paleomagnetic evidence indicates that the Alexander terrane resided within 14° of the paleoequator by the Early Devonian (Bazard et al., 1995; Butler et al., 1997).
Figure 1. Map showing distribution of Silurian rocks in Alaska: (1) Arctic Alaska terrane; (2) York terrane; (3) Seward terrane; (4) Porcupine terrane (part of autochthonous North America); (5) White Mountains terrane; (6) Livengood terrane; (7) Nixon Fork terrane; (8) Minchumina terrane; (9) Dillinger terrane; (10) Mystic terrane; (11) Windy terrane; (12) Alexander terrane; and (13) Tatonduk-Nation River area (part of autochthonous North America). Lower to Upper Silurian rocks are exposed in each terrane but are limited to restricted areas. See text for details. Note that Decker et al. (1994) redefined the Mystic, Dillinger, and Nixon Fork terranes as subterranes of the Farewell terrane on the basis of interfingering and gradational facies relationships. Modified from Coney et al. (1980) and Plafker and Berg (1994b).
Silurian-bearing terranes of Alaska BIOSTRATIGRAPHY AND SEDIMENTATION PATTERNS In many parts of Alaska, Silurian rocks are incompletely or discontinuously exposed (Churkin and Brabb, 1965) and may be extremely deformed, metamorphosed, structurally complicated, recrystallized, and (or) unfossiliferous (Dumoulin and Harris, 1987; Moore et al., 1994). Thus, in most instances sedimentary units have not been formally named; many macrofossils cannot be identified taxonomically to levels that are mandated for precise biostratigraphic age control, and generally only reconnaissancelevel studies or paleoenvironmental analyses lacking in detail are available (Moore et al., 1994). Two exceptions, which are discussed below, include upper Silurian (Ludlovian–Pridolian?) rocks of the Nixon Fork terrane (now Nixon Fork subterrane of the Farewell terrane; Blodgett et al., 1999) in southwestern and west-central Alaska (Clough and Blodgett, 1985, 1989; Patton et al., 1994) and a thick sedimentary sequence of late Llandovery–Pridoli? age in the Alexander terrane of southeastern Alaska (Ovenshine and Webster, 1970; Soja, 1990, 1991, 1993, 1994, 1996; Soja et al., 2000, 2003). In addition, a nearly complete suite of Silurian graptolite zones was recognized in the Road River Formation where it is discontinuously exposed in the Tatonduk-Nation River area in east-central Alaska (Churkin and Brabb, 1965). Reconnaissance-level studies and lack of detailed stratigraphic investigations preclude meaningful discussions of the sequence stratigraphy of Silurian rocks exposed in many parts of Alaska. An exception exists in southeastern Alaska (Alexander terrane), where a detailed stratigraphic analysis revealed the existence of cyclically repeated limestones in an Upper Silurian (Ludlovian?) section 47 m thick (Kittredge and Soja, 1993). Seven shallowing-upward cycles (parasequences) were identified ranging in thickness from 3 to 9 m. Coral-stromatoporoid wackestones form the base of each cycle and grade upward into oncoid packstones with silty, lime mudstones at the top of each sequence. The coral-stromatoporoid deposits are characterized by a low-diversity assemblage of dendroid corals, massive stromatoporoids, Atrypoidea brachiopods, and rare occurrences of biostromes associated with “Solenopora” algae, high-spired gastropods, and crinoids. Oncoids typically are 2–6 mm in diameter and form massive, meter-thick units. Interbedded lime mudstones lack shelly biotas but have well-developed vertical burrows and a fine-grained siliciclastic component. The succession of lithofacies within each cycle reflects an increase in energy levels from relatively deeper water environments to relatively shallower ones. The lack of abrasion in the corals and stromatoporoids suggests predominantly quiet-water conditions in shallow subtidal areas affected by periodic turbulence. Symmetrical oncoids and diminished diversity of associated faunas indicate an increase in wave and current activity related to a relative drop in sea level. An increase in the amount of siliciclastic detritus, absence of normal marine shelly benthos, and presence of vertical burrows reflect deposition of silty lime muds in a tidal flat environment
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during ongoing regression and the onset of orogenesis (Kittredge and Soja, 1993). Silurian graptolitic shales that formed as turbidites and as hemipelagic or periplatform oozes in deep-water slope and basinal settings are geographically widespread in Alaska. Examples include Upper Silurian metalimestones and highly deformed argillites that are interbedded with thin graywackes and deeply buried under the Arctic coastal plain (Arctic Alaska terrane) (Fig. 1) as part of the “Franklinian” sequence in northern Alaska (Grantz et al., 1983, 1994; Bird and Molenaar, 1987); graptolitic shales alternating with thin limestones and carbonate breccias (debris flows) of the Deceit Formation (Llandovery?–Pridoli? unit >300 m thick) in the Seward terrane of western Alaska (Ryherd and Paris, 1987) (Fig. 1); dark graptolite shales that alternate with laminated cherts and rare limestones, dolostones, chert arenites, and conglomerates that together exceed 270 m in thickness in the Road River Formation of east-central Alaska (Tatonduk-Nation River area) (Churkin and Brabb, 1965); graptolitic limestones and shales that are 1000 m thick in the Paradise Fork Formation (late Llandovery–Wenlock) of the Farewell terrane (Nixon Fork subterrane) in southwestern and west-central Alaska (Dutro and Patton, 1982) (Fig. 1); Early to “Mid” Silurian graptolitic black shales, limestones, and chert in the Farewell terrane (Dillinger subterrane) of south-central Alaska (Nokleberg et al., 1994) (Fig. 1); and thin-bedded, cherty graptolitic shales of the Descon Formation (Llandovery) in southeastern Alaska’s Alexander terrane (Churkin and Carter, 1970) (Figs. 1 and 2). In general, most of these deposits appear to be “shale out” facies that formed adjacent to shallow-marine carbonate platform sequences on the ancient continental margin of North America or Siberia (Churkin et al., 1984; see also Blodgett and Boucot, 1999; Dumoulin et al., 2000, 2002; Blodgett et al., 2002). These graptolitic rocks indicate that deep-marine conditions prevailed contemporaneously in most parts of the state through much of the Silurian. In addition, Silurian biostratigraphic surface and subsurface studies completed thus far on graptolites, conodonts, and chitinozoans in northern Alaska (Arctic Alaska terrane) and on the Seward Peninsula (York and Seward terranes) suggest a partial correlation in the temporal development of shallow- and deep-water deposits (Carter and Laufeld, 1975; Churkin, 1975; Till et al., 1986; Dumoulin and Harris, 1987, 1988, 1994; Repetski et al., 1987; Ryherd and Paris, 1987). In some areas, graptolitic shales accumulated from the early Llandovery–Ludlow, Pridoli?. Shallow-water, coral- and stromatoporoid-bearing limestones (or marbles) and dolostones or dolomitic mudstones began forming in most areas in the Llandovery and early Wenlock, persisting into the Late Silurian (Ludlow or Pridoli) (Dumoulin and Harris, 1987, 1994; Clough and Blodgett, 1989; Soja, 1990, 1996; Till and Dumoulin, 1994; Soja et al., 2000). Shallowing-upward sequences from basinal to carbonate platform settings are recorded in the York terrane of northwestern Alaska, the Nixon Fork subterrane of the Farewell terrane in southwestern and west-central Alaska, and the Alexander
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Figure 2. Generalized stratigraphy of the Lower Silurian (Llandovery) part of the Descon Formation, based on a composite section exposed on and near Prince of Wales Island (Craig area in southeastern Alaska [Alexander terrane]). Dots indicate actual zones documented by graptolite faunas. Detailed stratigraphic or sedimentologic investigations of this formation, which has a total thickness of at least 3000 m, have not been undertaken. Graptolite data and generalized stratigraphy are from Churkin and Carter (1970) and Churkin et al. (1971). Grain size and texture abbreviations: m.s.—mudstone; w.s.—wackestone; p.s.—packstone; g.s. —grainstone; si.—silt; f.s.—fine sand; c.s.—course sand; g.—gravel; sh.—shale; v.f.s.—very fine sand; m.s.—medium sand; v.c.s.—very course sand.
terrane of southeastern Alaska (Figs. 2 and 3). The shift from deep- to shallow-water conditions appears to have been synchronous in only limited portions of these three regions. One of the oldest widespread carbonates that records in situ, shallowmarine conditions on a subtidal platform is the Tolovana Limestone, which is a highly recrystallized, Lower (LlandoveryWenlock) to Upper Silurian deposit 1220 m thick in east-central Alaska (White Mountains terrane) (Fig. 1). That this formation originated under extremely shallow marine conditions is indicated by the abundance of peloids, ooids, and poorly preserved pentameroids in lime mudstones, wackestones, packstones, and grainstones, which overlie green and maroon mudstones 14.6 m thick at the base of the section (Blodgett et al., 1987; Weber et al., 1988). In most other parts of Alaska, Silurian limestones are generally younger in age (Wenlock-Pridoli) and are characterized by reef-related deposits and by the downslope transport of shallow-water detritus as debris flows and turbidites. For example, the Whirlwind Creek Formation in the Farewell terrane (Nixon Fork subterrane) is 1000–1500 m thick and comprises Upper Silurian (Ludlovian?-Pridolian?) to Middle Devonian (Eifelian) cyclically deposited, shallow-water, peloidal or silty limestones, some of which are characterized by biostromal Favosites and stromatoporoids, and “algal” laminated dolostones or dolomite
breccias (Dutro and Patton, 1982; Patton et al., 1989; R.B. Blodgett, 1996, personal commun.; Blodgett et al., 2000). In the York terrane (Seward Peninsula), ~18–270 m of unnamed, fine-grained or laminated limestones and dolomitic limestones (Wenlock-Ludlow) comprise abundant pentamerid brachiopods, laminar and massive stromatoporoids, and favositid corals preserved in growth position (Sainsbury et al., 1971; Till and Dumoulin, 1994). Other reefal deposits in which microbial organisms predominated characterize the Ludlow-Pridoli? sections in the Farewell and Alexander terranes (Clough and Blodgett, 1989; Soja, 1991, 1994; Soja and Antoshkina, 1997; Soja et al., 2000, 2003) (Fig. 3). In various parts of Alaska, Wenlock-Ludlow limestones typically consist of shallow-water constituents that were redeposited in deep-marine sites as debris flows and turbidites, including those exposed in the Ambler River section of the Arctic Alaska terrane (Dumoulin and Harris, 1988); upper Deceit Formation in the Seward terrane (Ryherd and Paris, 1987); “Lost Creek” unit in the Livengood terrane (Blodgett et al., 1988b); Road River Formation in the Porcupine terrane (Coleman, 1987); Holitna Lowlands of the Farewell terrane (Clough and Blodgett, 1989); and the upper Heceta Formation in the Alexander terrane (Soja, 1993; Soja et al., 2000) (Fig. 3). These deep-water limestones represent the contemporaneous
Figure 3. Generalized stratigraphy of a composite section of the Heceta Formation exposed on and near Heceta and Tuxekan Islands in southeastern Alaska (Alexander terrane). Conodont data are from Ovenshine and Webster (1969, 1970) and Savage (1985). Dots indicate actual zones documented by conodont faunas. Thickness, lithology, texture, and depth curve are modified from Soja (1993). Note that the upper parts of the Heceta Formation may be as young as Pridoli, and lower parts of the Karheen Formation may be as old as Pridoli. Data on red beds (R) in the Karheen Formation (Pridoli?–Early Devonian), which exceeds 1800 m in thickness, are generalized from Ovenshine et al. (1969), Ovenshine (1975), and Eberlein et al. (1983). Grain size and texture abbreviations as in Figure 2.
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development of basinal conditions adjacent to shallow-marine platforms in westernmost, northern, central, and southeastern Alaska (Sainsbury et al., 1971; Dutro and Patton, 1981; Dumoulin and Harris, 1987). CASE STUDIES Southeastern Alaska The best studied, most continuous, and thickest stratigraphic section of Silurian rocks is in Alaska’s southeastern island archipelago (Alexander terrane). Although the rocks exposed there crop out discontinuously along the shorelines of densely forested, mountainous islands, Silurian deposits are generally very well preserved and provide important insights into sedimentary dynamics in an ancient offshore oceanic island (Soja, 1992, 1993, 1996; Soja et al., 2000, 2003). Near Prince of Wales Island in the southern part of the terrane, the Silurian section comprises three formations listed in ascending stratigraphic order: Descon (Lower Ordovician–upper Llandovery deposit 3000 m thick), Heceta (upper Llandovery–Ludlow or Pridoli? deposit >3000 m thick), and Karheen (Ludlow or Pridoli?–Lower Devonian deposit 1800 m thick) (Ovenshine and Webster, 1969, 1970; Churkin and Carter, 1970; Eberlein and Churkin, 1970; Eberlein et al., 1983; Savage, 1985) (Figs. 2 and 3). In the Descon Formation, a variety of volcanic flows, breccias, tuffs, and agglomerates is interbedded with graywackes, mudstones, cherts, shales with minor limestone lenses, and quartzofeldspathic arenites (Churkin and Carter, 1970; Eberlein et al., 1983). These deposits, which accumulated largely in a basinal setting adjacent to volcanically active eruptive centers, form a shallowing-upward sequence gradational with the overlying Heceta Formation, as indicated in a composite stratigraphic section >1000 m thick in the eastern outcrop area (Soja, 1990) (Figs. 2 and 3). Finely laminated, graded lime mudstones and grainstones (turbidites) and a localized, massively bedded reef (stromatoporoid-coral rudstones and Renalcis boundstones) occur at the base of the Heceta Formation (Soja, 1990, 1993) (Fig. 3). These deposits are overlain by shallow-water limestones (peloidal wackestones, packstones, and grainstones) that originated under restricted conditions on an incipient carbonate platform, as suggested by the lack of diverse invertebrate faunas (Soja, 1991, 1993) (Fig. 3). Abrupt transition to massive, recrystallized limestone and to polymictic conglomerates represents a disruption in carbonate platform development in the early- to mid- (?) Ludlow during the earliest pulses of an orogenic event (Klakas orogeny) (Gehrels et al., 1983; Soja and Krutikov, this volume). Subsequent rejuvenation in carbonate platform development is reflected in Ludlovian skeletal limestones that are characterized by a diversity of shallow-marine shelly benthos (e.g., corals, stromatoporoids, Atrypoidea brachiopods, gastropods, Pycinodesma bivalves) and microbial organisms (Riding and
Soja, 1993; Soja, 1993; Soja and Riding, 1993) (Fig. 3). Stromatolite boundstones of Ludlovian-Pridolian? age represent barrier reefs that were constructed at the seaward edge of a shallow-marine platform by a diverse microbial-sponge consortium (Soja, 1991, 1994, 1996; Rigby et al., 1994; Soja et al., 2000, 2003). Sphinctozoan (aphrosalpingid) and stromatoporoid sponges as well as the problematic hydroid Fistulella played a secondary role to the microbes in reef construction; ostracodes, brachiopods, coral, and crinoid fragments were trapped within stromatolitic laminae and contributed skeletal debris to interfingering packstones. Of additional interest are the aphrosalpingid sponges that clearly encrusted Fistulella and other substrates, reflecting complex ecologic relationships in the microbial reefs (Soja et al., 2003). By the end of the Ludlow-Pridoli?, the margin of the shallow-marine platform in the Alexander terrane had experienced collapse during a relative rise in sea level, as indicated by the abrupt transition from shallow-water deposits to deep-marine turbidites, debris flow breccias with reef-derived clasts, and stromatolite slump blocks (Soja, 1993; Soja et al., 2000) (Fig. 3). Evidence for rapid deepening followed by abrupt shallowing during onset of orogenesis in the Alexander arc is derived from upper parts of the Heceta Formation, which become increasingly silty or argillaceous and are gradational with the overlying Karheen Formation (Fig. 3). Conglomerates, cross-bedded and pebbly sandstones, shales, and rare limestones of the Karheen Formation form a thick sequence (1800 m) of terrigenous red beds and shallow-marine deposits. Upper Heceta and Karheen deposits represent a classic flysch-molasse sequence that accumulated as a result of uplift, erosion, and clastic wedge progradation during culminating phases in the Klakas orogeny in the Late Silurian–Early Devonian (Ovenshine, 1975; Eberlein et al., 1983; Gehrels et al., 1983; Soja, 1993, Soja et al., 2000; Soja and Krutikov, this volume). Farther to the north, in the Kuiu Islands area and the Chilkat Mountains, coeval Silurian rocks in southeastern Alaska belong to the Bay of Pillars Formation (mid-Llandovery–early Ludlow), Kuiu limestone (Wenlock-Ludlow), and Point Augusta Formation (Upper? Silurian) (Muffler, 1967; Karl and Griffen, 1992). Suites of similar deposits, including graywackes, mudstones, subordinate limestones, conglomerates, and volcanic units 1500 m thick, occur in the Bay of Pillars and Point Augusta Formations. These rocks originated in deep-sea conditions as turbidites and hemipelagic sediment, and as debris flows and slump blocks transported downslope. Although not yet studied in detail, the Kuiu Limestone is a massive limestone deposit 800 m thick (Muffler, 1967) that appears to have formed under shallow-water conditions similar to those recorded in the Heceta Formation (Soja, 1997, pers. observation). As reported in Soja et al. (2000), the Willoughby Formation is the oldest bedrock unit in Glacier Bay and part of a SilurianTertiary sedimentary, volcanic, and metamorphic sequence that exceeds 6000 m in thickness. Willoughby thickness is estimated to be >1500 m, and the presence of the megalodont bivalve
Silurian-bearing terranes of Alaska Pycinodesma, aphrosalpingid sponges, and microbes Hecetaphyton, Ludlovia, and Sphaerina indicate formation in the Late Silurian (Kirk, 1927a, 1927b; Rossman, 1963; Soja et al., 2000). Although more highly deformed than coeval rocks in the southern part of Alaska’s panhandle, the Glacier Bay sequence where exposed on Drake Island is similar, and in places identical, to other coeval deposits in the Alexander terrane. Bedded limestones in the Willoughby Limestone are predominantly peloidal or mollusk-rich wackestones, packstones, and grainstones that are intercalated with burrowed peloidal mudstones-wackestones, microbial boundstones including stromatolites and thrombolites, and rare skeletal grainstones. These deposits comprise abundant fossils representing low-diversity suites of mollusks, small clusters of massive stromatoporoids, Amphipora, rugose corals, and rare atrypid brachiopods, with a significant proportion of skeletal debris coated by microbial encrusters. Pycinodesma megalodont bivalves and Euomphalopterus (and other) gastropods compose as much as 50% of rock volume and form conspicuous monospecific concentrations on several bedding surfaces (Soja et al., 2000). Evidence for relatively low-energy subtidal conditions in a restricted, shallow-marine lagoon is extensive: abundant micrite and micritized skeletal grains; whole gastropods with microbial encrustations; whole and articulated pelecypods preserved in growth position; amphiporoids with well-preserved delicate branches; lack of intertidal indicators; and general absence of stenohaline (e.g., normal marine) taxa such as brachiopods, crinoids, and bryozoans. Higher-energy events associated with strong waves and currents caused by periodic storms are suggested by the high degree of skeletal fragmentation (micritized grains), grainstone deposits, and symmetrical oncoids in some beds (Soja et al., 2000). The mollusk-rich peloidal wackestones and packstones grade upward into mudstones, wackestones, and gastropod packstones. These are overlain by microbialdominated carbonates with well-developed concentric laminae, abundant Ludlovia, Fistulella, and aphrosalpingid sponges. The microbial-dominated rocks represent an offshore (e.g., platform margin) stromatolite reef complex that is >100 m thick where it is exposed along the west-central shoreline of Drake Island. Stromatolite facies are repeated cyclically three or four times along the shoreline and grade upward into slumped and/ or faulted turbidites and megabreccias. The foreslope debris flows that are interbedded with in situ tidal flat laminites, lagoonal deposits, and stromatolites also form cyclic repetitions, each a few hundred meters thick. These deposits reveal that the platform and its margin experienced a long-term process of periodic spalling under fluctuating environmental conditions. The regrowth of stromatolites and microbial-cement crusts imply that conditions favorable for reef development resumed periodically at the shelf rim. Several 100 m of interbedded turbidites and megabreccias record the beginning stages in the widespread collapse of the platform margin that eventually terminated stromatolite reef growth. Coincidence in timing suggests that culminating stages in the Klakas orogeny induced
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large-scale, catastrophic collapse of the platform, as recorded by similar debris flows and megabreccias in areas to the south and by the progradation of a clastic wedge in the Early Devonian (Soja et al., 2000). Southwestern Alaska Upper Silurian carbonate deposits in the Holitna Lowlands (Sleetmute, Taylor Mountains, McGrath, and Lime Hills quadrangles; Decker et al., 1994) of southwestern Alaska (Farewell terrane) are strikingly similar to those in the Alexander terrane of southeastern Alaska. In particular, both regions show the impressive and apparently coeval development of distinctive stromatolite-sphinctozoan sponge deposits in the Late Silurian (Ludlow-Pridoli?). In southwestern Alaska, an unnamed carbonate unit represents an Upper Silurian–Lower Devonian microbial (not technically “algal”) reef complex ~500 m thick, 1 km wide, and several hundred kilometers long (Clough and Blodgett, 1985, 1989). These deposits interfinger with limestone debris flows and graptolitic shales of basinal origin, which belong to the Dillinger terrane of Coney et al. (1980) and the East Fork Hills subterrane (first defined by Dutro and Patton, 1982) of the Minchumina terrane of Patton et al. (1994) (note: the Dillinger terrane and East Fork Hills subterrane are now subsumed within the Farewell terrane of Decker et al., 1994; see also Blodgett and Boucot, 1999). Stacked, shallowing-upward sequences define the microbial reef complex, beginning with basal thrombolite mud mounds 20 m thick that are succeeded upsection by stromatolite boundstones. These laminated deposits are tens of meters thick and characterized locally by channels and small cavities. Each sequence is generally capped by intertidal-supratidal “cryptalgal” laminated peloidal mudstones and wackestones (Clough and Blodgett, 1985, 1989). Diverse microbial floras (e.g., Solenopora, Epiphyton, Renalcis, and Sphaerocodium) constructed the buildups associated with dasycladaceans, gypidulinid brachiopods, sphinctozoan sponges (aphrosalpingids), problematic hydroids (e.g., Fistulella), and rare corals and stromatoporoids. Future study will help to show in detail the degree of similarity between these stromatolite-related communities and those preserved in southeastern Alaska and Russia. FAUNAL AFFINITIES AND TERRANE PALEOGEOGRAPHY Graptolite sequences in nearby northern Canada (e.g., Road River Formation) are similar to those in Alaska and suggest that northwestern North America was a lateral continuation of the passive continental margin in the northern Yukon, sharing with the Canadian Cordillera a depositional history spanning 200 m.y. (Lane, 1991). Further evidence may be derived from Upper Silurian limestones in the Arctic Alaska terrane that appear to be similar to unnamed Silurian carbonate rocks in the Selwyn Basin and Mackenzie Platform of northern Canada
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(Dillon et al., 1987; Dumoulin and Harris, 1987; Lane, 1991; Moore et al., 1994). Alternatively, Moore (1986), Dover (1994), and more recently Blodgett and Boucot (1999), Blodgett et al. (2002), and Dumoulin et al. (2000, 2002) promote the recognition of pre-Devonian terranes in northern Alaska that originated beyond the depositional margin of North America. Similarities between rocks in Northern Alaska (Arctic Alaska terrane) and various terranes in interior Alaska (Nixon Fork and East Fork subterranes of Farewell terrane; Livengood terrane; and Porcupine terrane) are also indicated by the coeval occurrence of deep-marine graptolitic shales in some areas throughout the Silurian and of carbonates (or metalimestones, metasandstones, and phyllites) with assorted shallow-water megafossils (e.g., corals, gastropods, bryozoa, brachiopods, orthocone cephalopods, crinoids, ostracodes, trilobites) that were transported downslope as debris flows (e.g.., the “Lost Creek unit” in east-central Alaska of Blodgett et al., 1988b) and turbidites in the early Wenlock– Ludlow (Coleman, 1985, 1987; Blodgett et al., 1988b; Dumoulin and Harris, 1988; Grantz et al., 1994). These units, largely unnamed, appear to be equivalent in age to the graptolite shales associated with black cherts and rare limestones in the Road River Formation of east-central Alaska and the Yukon Territory (Churkin and Brabb, 1965; Coleman, 1985; Blodgett et al., 1988b). Coarse-grained, Silurian siliciclastic units become predominant in parts of Alaska during the Late Silurian and appear to correspond to “Old Red Sandstone” facies, which formed in response to late Caledonide events and early stages in the Ellesmere orogeny, as described in Ziegler (1988). For example, Silurian marine deposits grade upward into nonmarine rocks of Early or Middle Devonian age in the Arctic Alaska terrane of northern Alaska (Churkin, 1975; Grantz et al., 1994; Moore et al., 1994) and in the Alexander terrane of southeastern Alaska (Soja and Krutikov, this volume; also see discussion below). An extensive, sub-Mississippian unconformity in the subsurface of some subterranes of the Arctic Alaska terrane suggests that parts of northern Alaska may have been linked with the Ellesmere fold belt of the Canadian Arctic and northern Greenland (Moore et al., 1994). The lack of deformation and (or) metamorphism in Lower?–Middle Devonian rocks in northern Alaska (northeastern Brooks Range and western North Slope) coupled with evidence of strongly deformed Upper Silurian–Lower Devonian rocks of the Franklinian succession in northern Alaska (Arctic continental margin) suggest that orogenesis may have occurred in northern parts of Alaska in the Late Silurian–Early Devonian (Grantz et al., 1994; Moore et al., 1994). Blodgett et al. (1988a) noted that a disconformity exists between Lower Devonian (Emsian) limestones and Upper Ordovician rocks in the northeastern Brooks Range. Parallel orientation of beds above and below the unconformity suggests a lack of significant tectonism from the Late Ordovician–Early Devonian in the Shublik Mountains area of northern Alaska. Although the causes of a postulated pre–Middle Devonian orogenic event in northern Alaska are unknown, Moore et al. (1994) noted that it may have been related to convergence of North
America with Siberia or to North America’s collision with an island arc. Because the Alexander terrane of southeastern Alaska also experienced an orogenic event (Klakas orogeny) during this same time interval (Gehrels et al., 1983), an intriguing possibility exists that the Alexander arc collided with cratonic portions of Laurentia in the Late Silurian–Early Devonian and became welded to the Laurussian landmass during late-stage Caledonide activity (Soja and Krutikov, this volume). Lower Silurian deposits of Greenland may record similar tectonic events: carbonate platform destruction during back-stepping of the shelf margin, megabreccias induced during platform collapse, and progradation of flysch derived from the Caledonide orogen (Sønderholm and Harland, 1989; Higgins et al., 1991; Trettin, 1991; Trettin et al., 1991; Surlyk and Ineson, 1992). By the Late Silurian, proximity of the Alexander terrane to Laurussia and other areas that bordered the Uralian Seaway is confirmed by similar stromatolites that developed as subtidal reefs in southwestern and southeastern Alaska, eastern Baltica, and western Siberia (Salair). Detailed studies of Silurian Communities in southeastern and southwestern Alaska (Alexander and Farewell terranes, respectively) reveal the existence of a distinctive consortium of calci-microbes (not primarily “algae”) and aphrosalpingid (sphinctozoan) sponges in Upper Silurian (Ludlow-Pridoli?) stromatolites (Clough and Blodgett, 1985, 1989; Soja, 1991, 1994; Soja and Riding, 1993; Soja and Antoshkina, 1997, 1998; Soja et al., 2000, 2003; Antoshkina and Soja, 2006). Well-developed cavities in microbial boundstones and preservation of stromatolite clasts in foreslope slump blocks and debris flows indicate that these biotas constructed shallow subtidal stromatolites within Benthic Assemblage 3 (Boucot, 1975) at highenergy platform margins. Downslope slumping of platform-edge debris led to their preservation mixed with other shallow-water taxa in Benthic Assemblage 4 (or 5?). Associated Upper Silurian (Ludlovian?) limestones in southeastern Alaska characterized by in situ, low-diversity assemblages of dendroid corals, massive stromatoporoids, Atrypoidea brachiopods, “Solenopora” algae, high-spired gastropods, and crinoids also corroborate evidence of Benthic Assemblages 2 and 3 in backreef sites. Comparison of microbial taxa (e.g., Ludlovia, Hecetaphyton, and Sphaerina) and conspecific aphrosalpingid sponges reveals that the stromatolite deposits in southeastern Alaska have closest affinities with Ludlow (and Pridoli?) rocks in the northern and southern Urals and in western Siberia (Magkova, 1955; Antoshkina, 1994; Soja and Antoshkina, 1997, 1998; Newton et al., 2002; Soja et al., 2000, 2003; Antoshkina and Soja, 2006). Specifically, co-occurrence of identical microbe, algal, and sponge species in the microbial reefs indicates a high degree of similarity between the Alexander terrane’s Ludlovia and Epiphyton-Sphaerina associations (Soja and Riding, 1993) and the Cyanophyta-Aphrosalpingata and Fistulella-Ikella Communities identified in the Urals by Sapelnikov et al. (1999). Microbe-dominated communities were examined at two sites in the Ilych Reef Formation in the Northern Urals and in the
Silurian-bearing terranes of Alaska Vishnevaya Gora Formation in the Southern Urals, deposits that accumulated as passive margin sequences on the eastern edge of Baltica in the Late Silurian (Newton, 2002; Newton et al., 2002). As in the Alaskan deposits, microbial taxa, such as Ludlovia, Hecetaphyton, Sphaerina, Rothpletzella, Renalcis, and Epiphyton, were the primary constructors of the cavity-riddled, shelfmargin reefs. Preliminary data from coeval deposits in Salair, a Paleozoic island arc welded to present-day western Siberia, extend the range of these communities into present-day western Siberia (Antoshkina et al., 2004; Antoshkina and Soja, 2006). That the Alaskan reefs share Silurian microbial and sponge taxa with coeval stromatolites in the Urals and Salair implies the existence of a seaway that fostered paleobiogeographic connections across many degrees of longitude. In the Late Silurian, the Uralian Seaway bordered three major continents—Laurentia, Siberia, and Baltica. The sharing of paleobiogeographic signatures among disparate terranes and continental areas indicates that the Uralian Seaway functioned as an equatorial-subequatorial marine “corridor” that enabled the transmigration of biotas between northwestern Laurentia, Siberia, and eastern Baltica and between islands or microcontinents situated within it (Soja, 1994; Soja and Antoshkina, 1997, 1998; Soja et al., 2000, 2003; Antoshkina et al., 2004; Antoshkina and Soja, 2006). Origin of the Alexander terrane as an island arc characterized by juvenile oceanic crust uncontaminated by continental material (Samson et al., 1989) contradicts suggestions made by Blodgett and Boucot (1999) and Blodgett et al. (2002) that it is a rifted portion of ancient Siberia later accreted to northwestern North America. Its volcanosedimentary stratigraphic profile and the biogeographic affinities of preserved organisms suggest that the Alexander terrane resided as an offshore oceanic island in the Uralian Seaway from the early Paleozoic until its collision with an unknown continent in the Late Silurian–Early Devonian (Soja et al., 2000). Encroachment and eventual collision of the Alexander terrane with another landmass (Laurussia appears to be the only viable option) near the end of the Silurian is also suggested by detrital zircons. Those extracted from the Karheen Formation represent a shift in sediment input from intra-arc deposits (predominantly of Ordovician age) to continental basement rocks 1–3 b.y. old (Gehrels et al., 1996; Butler et al., 1997). Upper Silurian?–Lower Devonian red beds of the Karheen Formation represent a very thick (1800 m) sequence of syn- and post-orogenic shallow-marine and continental clastic material. These deposits accumulated as molasse in a foreland basin following rapid erosion of adjacent uplifted areas (Ovenshine et al., 1969; Ziegler, 1988), and they appear to be similar to red beds deposited in many parts of Laurussia as a late-stage manifestation of Caledonide orogenic activity (Soja and Antoshkina, 1997; Soja, unpub. data). Detailed provenance studies of detrital grains in the future should help to support or refute two related hypotheses: (1) the Klakas orogeny of southeastern Alaska (Alexander terrane) was a Caledonide event, and (2) the Karheen Formation of southeastern Alaska is a facies of the Old Red Sandstone
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that was dismembered from the Laurussian margin and accreted millions of years later to northwestern Laurentia (Soja et al., 2000; Soja and Krutikov, this volume). The origins of the Farewell terrane are still debated, but there is general agreement that this terrane was part of a continental margin; it is either an in situ peninsular extension of North America or it was displaced and (or) rotated from Canada or Siberia to its present position (Abbott, 1995; Plumley et al., 1981; Churkin et al., 1984; Blodgett and Clough, 1985; Decker et al., 1994; Patton et al., 1994). More recent interpretations favor a Siberian origin for the Farewell and associated terranes (Blodgett and Boucot, 1999; Blodgett et al., 2002). For example, Middle Cambrian trilobites as well as brachiopods of Late Ordovician, Late Silurian, and Early Devonian age with Siberian affinities suggest the Farewell terrane had Laurentian origins (but experienced transoceanic exchange of invertebrate larvae) or was attached to but later rifted from Siberia in the Paleozoic (Babcock et al., 1995; Blodgett et al., 1999, 2002). SUMMARY Abundant carbonate deposits as well as warm-water marine biotas preserved in most Silurian sections of Alaska suggest that the northwestern part of North America formed throughout the Silurian under a subtropical-tropical climate. These conditions were favorable for thick accumulations of microbe-rich limestones and the growth of barrier reefs. In the Alexander terrane, deep- and shallowwater limestones grade upward into shallow-marine and nonmarine red beds (conglomerates, sandstones, siltstones, and mudstones) of Late Silurian?–Early Devonian age (Ovenshine et al., 1969; Eberlein and Churkin, 1970; Churkin, 1975; Ovenshine, 1975; Soja, 1993, 1996; Soja and Krutikov, this volume). These molasse deposits suggest that hot, arid, and equatorial conditions prevailed similar to the continental climate that developed across the Laurussian landmass, or Old Red Sandstone continent, following late Caledonide and early Ellesmerian orogenic events, as described in Ziegler (1988). Much of Alaska lay offshore of the ancient continental margin of North America in the Silurian; thus, few indicators of sea-level change are recorded in the extensive graptolite shales that occur in most areas. Platform limestones studied in the greatest detail in southeastern Alaska (Alexander terrane) reveal that relative sea level fluctuated significantly beginning in the early Wenlock (Soja, 1993, 1996) (Fig. 3). To a first approximation, comparison of seven parasequences of Ludlow (?) age with correlative sections in other parts of Alaska and lack of correspondence with global sea-level curves suggest that tectonic perturbations with secondary eustatic effects were the primary cause of cyclicity in southeastern Alaska (Kittredge and Soja, 1993). Cyclic deposition in peri/subtidal sites was terminated in the Alexander terrane by rapid drowning of the carbonate platform during Late Silurian orogenesis. Related environmental factors that may have favored proliferation and preservation of microbial reefs in the Alexander and Farewell terranes, Siberia, and eastern Baltica in the Late Silurian appear to be tied to late stages in Caledonian-Scandian orogenesis
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and partial closure of the Uralian Seaway (Bazard et al., 1995; Soja et al., 2000). Contributing factors to be examined in future studies include (1) increased terrestrial runoff and elevated nutrient concentrations that fueled microbial-algal “blooms”; (2) relaxed ecological landscapes that were conducive to opportunistic microorganismal growth after reef-building competitors (i.e., metazoans) suffered widespread extinction; and/or (3) elevated seawater alkalinity that promoted a significant calcification event, which enhanced the cementation and preservation of microbial structures during tectonically induced fluctuations in sea level (Soja, pers. observation). ACKNOWLEDGMENTS Supported by the National Science Foundation and Colgate University’s Research Council. Thanks to Robert B. Blodgett, David M. Rohr, Paula Noble, and Julie Dumoulin for carefully reviewing this paper and suggesting improvements. REFERENCES CITED Abbott, G., 1995, Does Middle Cambrian rifting explain the origin of the Nixon Fork terrane?: Geological Society of America Abstracts with Programs, v. 27, no. 5, p. 1. Antoshkina, A.I., 1994, Rify v Paleozoe Pechorskovo Urala: Saint Petersburg, Russia, Nauka, 154 p. Antoshkina, A.I., and Soja, C.M., 2006, Late Silurian reconstruction indicated by migration of reef biota between Alaska, the Urals, and Siberia (Salair): Geologiska Föreningens i Stockholm Förhandlingar, v. 128, p. 75–78. Antoshkina, A.I., Soja, C.M., and Gutak, Y.M., 2004, Paleobiogeographic implications of Silurian reef biotas in the Urals, southeastern Alaska, and Salair: 32nd International Geological Congress, Florence, Italy, Abstracts, Part 2, p. 1082. Babcock, L.E., St. John, J., Jacobson, S.R., Askin, R.A., and Blodgett, R.B., 1995, Neoproterozoic to early Paleozoic geological history of the Nixon Fork subterrane of the Farewell terrane, Alaska: Geological Society of America Abstracts with Programs, v. 27, no. 5, p. 2–3. Bazard, D.R., Butler, R.F., Gehrels, G., and Soja, C.M., 1995, Early Devonian paleomagnetic data from the Lower Devonian Karheen Formation suggest Laurentia-Baltica connection for the Alexander terrane: Geology, v. 23, p. 707–710, doi: 10.1130/0091-7613(1995)023<0707:EDPDFT> 2.3.CO;2. Bird, K.J., and Molenaar, C.M., 1987, Chapter 5. Stratigraphy, in Bird, K.J., and Magoon, L.B., eds., Petroleum geology of the northern part of the Arctic National Wildlife Refuge, northeastern Alaska: U.S. Geological Survey Bulletin 1778, p. 37–248. Blodgett, R.B., and Boucot, A.J., 1999, Late Early Devonian (Late Emsian) eospiriferinid brachiopods from Shellabarger Pass, south-central Alaska, and their biogeographic importance; further evidence for a Siberian origin of the Farewell and Alaskan accreted terranes: Senckenbergiana Lethaea, v. 79, p. 209–221. Blodgett, R.B., and Clough, J.G., 1985, The Nixon Fork terrane—Part of an insitu peninsular extension of the Paleozoic North American continent: Geological Society of America Abstracts with Programs, v. 17, no. 6, p. 342. Blodgett, R.B., Wheeler, K.L., Rohr, D.M., Harris, A.G., and Weber, F.R., 1987, A Late Ordovician age reappraisal for the upper Fossil Creek Volcanics, and possible significance for glacio-eustasy, in Hamilton, T.D., and Galloway, J.P., eds., Geologic studies in Alaska by the U.S. Geological Survey during 1986: U.S. Geological Survey Circular 998, p. 54–58. Blodgett, R.B., Rohr, D.M., Harris, A.G., and Jia-Yu, R., 1988a, A major unconformity between Upper Ordovician and Lower Devonian strata in the Nanook Limestone, Shublik Mountains, northeastern Brooks Range, in Galloway, J.P., and Hamilton, T.D., eds., Geologic studies in Alaska by the U.S. Geological Survey during 1987: U.S. Geological Survey Circular 1016, p. 18–23. Blodgett, R.B., Zhang, N., Ormiston, A.R., and Weber, F.R., 1988b, A Late Si-
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Silurian-bearing terranes of Alaska Dover, J.H., 1994, Geology of part of east-central Alaska, in Plafker, G., and Berg, H. C., eds., The geology of Alaska: Boulder, Colorado, Geological Society of America, Geology of North America, v. G-1, p. 153–204. Dumoulin, J.A., and Harris, A.G., 1987, Lower Paleozoic carbonate rocks of the Baird Mountains quadrangle, western Brooks Range, Alaska, in Tailleur, I., and Weimer, P., eds., Alaskan North Slope geology: Bakersfield, California, Pacific Section, SEPM, and Anchorage, Alaska Geological Society, v. 1, p. 311–336. Dumoulin, J.A., and Harris, A.G., 1988, Off-platform Silurian sequences in the Ambler River quadrangle, in Hamilton, T.D., and Galloway, J.P., eds., Geologic studies in Alaska by the U.S. Geological Survey during 1987: U.S. Geological Survey Circular 1016, p. 35–38. Dumoulin, J.A., and Harris, A.G., 1994, Depositional framework and regional correlation of pre-Carboniferous metacarbonate rocks of the Snowden Mountain area, central Brooks Range, northern Alaska: U.S. Geological Survey Professional Paper 1545, 66 p. Dumoulin, J.A., Harris, A.G., Bradley, D.C., and DeFreitas, T.A., 2000, Facies patterns and conodont biogeography in Arctic Alaska and the Canadian Arctic Islands: Evidence against juxtaposition of these areas during Early Paleozoic time: Polarforschung, v. 68, p. 257–266. Dumoulin, J.A., Harris, A.G., Gagiev, M., Bradley, D.C., and Repetski, J.E., 2002, Lithostratigraphic, conodont, and other faunal links between lower Paleozoic strata in northern and central Alaska and northeastern Russia, in Miller, E.L., Grantz, A., and Klemperer, S.L., eds., Tectonic evolution of the Bering Shelf-Chukchi Sea-Arctic Margin and adjacent landmasses: Geological Society of America Special Paper 360, p. 291–312. Dutro, J.T., Jr., and Patton, W.W., Jr., 1981. Lower Paleozoic platform carbonate sequence in the Medfra quadrangle, west-central Alaska, in Albert, N.R.D., and Hudson, T., eds., The United States Geological Survey in Alaska: Accomplishments during 1979: U.S. Geological Survey Circular 823-B, p. B42–B43. Dutro, J.T., Jr., and Patton, W.W., Jr., 1982, New Paleozoic formations in the northern Kuskokwim Mountains, west-central Alaska: U.S. Geological Survey Bulletin 1529-H, p. H13–H22. Eberlein, G.D., and Churkin, M., Jr., 1970, Paleozoic stratigraphy in the northwest coastal area of Prince of Wales Island, southeastern Alaska: U.S. Geological Survey Bulletin 1284, 67 p. Eberlein, G.D., Churkin, M., Jr., Carter, C., Berg, H.C., and Ovenshine, A.T., 1983, Geology of the Craig quadrangle, Alaska: U.S. Geological Survey Open-File Report 83-91, 28 p. Gehrels, G.E., and Berg, H.C., 1994, Geology of southeastern Alaska, in Plafker, G., and Berg, H.C., eds., The geology of Alaska: Boulder, Colorado, Geological Society of America, Geology of North America, v. G-1, p. 451–467. Gehrels, G.E., Saleeby, J.B., and Berg, H.C., 1983, Preliminary description of the Klakas orogeny in the southern Alexander terrane, southeastern Alaska, in Stevens, C.H., ed., Pre-Jurassic rocks in western North American suspect terranes: Los Angeles, Pacific Section, SEPM, p. 131–141. Gehrels, G.E., Butler, R.F., and Bazard, D.R., 1996, Detrital zircon geochronology of the Alexander terrane, Alaska: Geological Society of America Bulletin, v. 108, p. 722–734. Grantz, A., Tailleur, I.L., and Carter, C., 1983, Tectonic significance of Silurian and Ordovician graptolites, Lisburne Hills, northwest Alaska: Geological Society of America Abstracts with Programs, v. 15, no. 5, p. 274. Grantz, A., May, S.D., and Hart, P.E., 1994, Geology of the Arctic continental margin, in Plafker, G., and Berg, H.C., eds., The geology of Alaska: Boulder, Colorado, Geological Society of America, Geology of North America, v. G-1, p. 17–48. Higgins, A.K., Ineson, J.R., Peel, J.S., Surlyk, F., and Sønderholm, M., 1991, Cambrian to Silurian basin development and sedimentation, north Greenland, in Trettin, H.P., ed., Geology of the Innuitian Orogen and Arctic Platform of Canada and Greenland: Ottawa, Geological Survey of Canada, Geology of Canada, no. 3 (Geological Society of America, Geology of North America, v. E), p. 111–161. Karl, S.M., and Griffen, C.F., 1992, Sedimentology of the Bay of Pillars and Point Augusta Formations, Alexander archipelago, Alaska, in Bradley, D.C., and Dusel-Bacon, C., eds., Geologic studies in Alaska by the U.S. Geological Survey, 1991: U.S. Geological Survey Bulletin 2041, p. 171–185. Kirk, E., 1927a, Pycnodesma, a new molluscan genus from the Silurian of Alaska: Proceedings of the U.S. National Museum, v. 71, p. 1–9.
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Kirk, E., 1927b, Pycinodesma, new name for Pycnodesma Kirk not Schrammen: Journal of the Washington Academy of Sciences, v. 17, p. 543. Kittredge, L.E., and Soja, C.M., 1993, Cyclicity in Silurian island-arc carbonates, Alaska: Geological Society of America Abstracts with Programs, v. 25, no. 2, p. 29. Lane, L.S., 1991, The pre-Mississippian “Neruokpuk Formation,” northeastern Alaska and northwestern Yukon: Review and new regional correlation: Canadian Journal of Earth Sciences, v. 28, p. 1521–1533. Moore, T.E., 1986, Stratigraphic framework and tectonic implications of preMississippian rocks, northern Alaska: Geological Society of America Abstracts with Programs, v. 19, no. 2, p. 159. Moore, T.E., Wallace, W.K., Bird, K.J., Karl, S.M., Mull, C.G., and Dillon, J.T., 1994, Geology of northern Alaska, in Plafker, G., and Berg, H.C., eds., The geology of Alaska: Boulder, Colorado, Geological Society of America, Geology of North America, v. G-1, p. 49–140. Muffler, L.J.P., Jr., 1967, Stratigraphy of the Keku Islets and neighboring parts of Kuiu and Kupreanof Islands, southeastern Alaska: U.S. Geological Survey Bulletin 1241-C, 52 p. Myagkova, E.I., 1955, Kharakteristike klassa Aphrosalpingoida Miagkova, 1955: Doklady Akademii Nauk SSSR, v. 104, p. 478–481. Newton, A., 2002, Paleoenvironmental implications of Silurian sponge-microbe associations, Northern and Southern Ural Mountains, Russia [unpublished honors thesis]: Hamilton, New York, Colgate University, 49 p. Newton, A., Soja, C.M., Antoshkina, A.I., and White, B., 2002, Paleoenvironmental implications of Silurian sponge-microbe associations from the Northern Ural Mountains, Russia: Geological Society of America Abstracts with Programs, v. 34, no. 1, p. 72. Nokleberg, W.J., Plafker, G., and Wilson, F.H., 1994, Geology of south-central Alaska, in Plafker, G., and Berg, H.C., eds., The geology of Alaska: Boulder, Colorado, Geological Society of America, Geology of North America, v. G-1, p. 311–366. Ovenshine, A.T., 1975, Tidal origin of parts of the Karheen Formation (Lower Devonian), southeastern Alaska, in Ginsburg, R.N., ed., Tidal deposits: A casebook of recent examples and fossil counterparts: New York, SpringerVerlag, p. 127–133. Ovenshine, A.T., and Webster, G.D., 1969, Silurian conodonts from southeastern Alaska: Geological Society of America Abstracts with Programs, v. 1, no. 3, p. 51. Ovenshine, A.T., and Webster, G.D., 1970, Age and stratigraphy of the Heceta limestone in northern Sea Otter Sound, southeastern Alaska: U.S. Geological Survey Professional Paper 700-C, p. C170–C174. Ovenshine, A.T., Eberlein, G.D., and Churkin, M., Jr., 1969, Paleotectonic significance of a Silurian-Devonian clastic wedge, southeastern Alaska: Geological Society of America Abstracts with Programs, v. 1, no. 3, p. 50. Patton, W.W., Jr., Box, S.E., Moll-Stalcup, E.J., and Miller, T.P., 1989, Geology of west-central Alaska: U.S. Geological Survey Open-File Report 89-554, 41 p. Patton, W.W., Jr., Box, S.E., Moll-Stalcup, E.J., and Miller, T.P., 1994, Geology of west-central Alaska, in Plafker, G., and Berg, H.C., eds., The geology of Alaska: Boulder, Colorado, Geological Society of America, Geology of North America, v. G-1, p. 241–269. Plafker, G., and Berg, H.C., 1994a, Introduction, in Plafker, G., and Berg, H.C., eds., The geology of Alaska: Boulder, Colorado, Geological Society of America, Geology of North America, v. G-1, p. 1–16. Plafker, G., and Berg, H.C., 1994b, Overview of the geology and tectonic evolution of Alaska, in Plafker, G., and Berg, H.C., eds., The geology of Alaska: Boulder, Colorado, Geological Society of America, Geology of North America, v. G-1, p. 990. Plumley, P.W., Coe, R.S., Patton, W.W., and Moll, E.J., 1981, Paleomagnetic study of the Nixon Fork terrane, west-central Alaska: Geological Society of America Abstracts with Programs, v. 13, no. 7, p. 530. Repetski, J.E., Carter, C., Harris, A.G., and Dutro, J.T., Jr., 1987, Ordovician and Silurian fossils from the Doonerak Anticlinorium, central Brooks Range, Alaska, in Hamilton, T.D., and Galloway, J.P., eds., Geologic studies in Alaska by the U.S. Geological Survey during 1986: U.S. Geological Survey Circular 998, p. 40–41. Riding, R., and Soja, C.M., 1993, Silurian calcareous algae, cyanobacteria, and micro-problematica from the Alexander terrane, Alaska: Journal of Paleontology, v. 63, p. 710–728. Rigby, J.K., Nitecki, M.H., Soja, C.M., and Blodgett, R.B., 1994, Silurian aphrosalpingid sphinctozoans from Alaska and Russia: Acta Palaeontologica Polonica, v. 39, p. 341–391.
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Rossman, D.L., 1963., Geology of the eastern part of the Mount Fairweather quadrangle, Glacier Bay, Alaska: U.S. Geological Survey Bulletin 1121-K, p. K1–K53. Ryherd, T.J., and Paris, C.E., 1987, Ordovician through Silurian carbonate base-of-slope apron sequence, northern Seward Peninsula, Alaska, in Tailleur, I., and Weimer, P., eds., Alaskan North Slope geology: Bakersfield, California, Pacific Section, SEPM, and Anchorage, Alaska Geological Society, v. 1, p. 347–348. Sainsbury, C.L., Dutro, J.T., Jr., and Churkin, M., Jr., 1971, The OrdovicianSilurian boundary in the York Mountains, western Seward Peninsula, Alaska: U.S. Geological Survey Professional Paper 750-C, p. C52–C57. Samson, S.D., McClelland, W.C., Patchett, P.J., Gehrels, G.E., and Anderson, R.G., 1989, Evidence from neodymium isotopes for mantle contributions to Phanerozoic crustal genesis in the Canadian Cordillera: Nature, v. 337, p. 705–709, doi: 10.1038/337705a0. Sapelnikov, V.P., Bogoyavlenskaya, O.V., Mizens, L.I., and Shuysky, V.P., 1999, Silurian and Early Devonian benthic communities of the Ural-Tien Shan Region, in Boucot, A.J., and Lawson, J.D., eds., Paleocommunities: A case study from the Silurian and Lower Devonian: New York, Cambridge University Press, p. 510–544. Savage, N.M., 1985, Silurian (Llandovery-Wenlock) conodonts from the base of the Heceta Limestone, southeastern Alaska: Canadian Journal of Earth Sciences, v. 22, p. 711–727. Soja, C.M., 1990, Island arc carbonates from the Silurian Heceta Formation of southeastern Alaska (Alexander terrane): Journal of Sedimentary Petrology, v. 60, p. 235–249. Soja, C.M., 1991, Origin of Silurian reefs in the Alexander terrane of southeastern Alaska: Palaios, v. 6, p. 111–125, doi: 10.2307/3514877. Soja, C.M., 1992, Potential contributions of ancient oceanic islands to evolutionary theory: The Journal of Geology, v. 100, p. 125–134. Soja, C.M., 1993, Carbonate platform evolution in a Silurian oceanic island: A case study from Alaska’s Alexander terrane: Journal of Sedimentary Petrology, v. 63, p. 1078–1088. Soja, C.M., 1994, Significance of Silurian stromatolite-sphinctozoan reefs: Geology, v. 22, p. 355–358, doi: 10.1130/0091-7613(1994)022<0355: SOSSSR>2.3.CO;2. Soja, C.M., 1996, Island-arc carbonates: characterization and recognition in the ancient geologic record: Earth-Science Reviews, v. 41, p. 31–65, doi: 10.1016/0012-8252(96)00029-3. Soja, C.M., and Antoshkina, A.I., 1997, Coeval development of Silurian stromatolite reefs in Alaska and the Ural Mountains: Implications for paleogeography of the Alexander terrane: Geology, v. 25, p. 539–542, doi: 10. 1130/0091-7613(1997)025<0539:CDOSSR>2.3.CO;2. Soja, C.M., and Antoshkina, A.I., 1998, Coeval development of Silurian stromatolite reefs in Alaska and the Ural Mountains: Implications for paleogeography of the Alexander terrane: Reply: Geology, v. 26, p. 383–384. Soja, C.M., and Krutikov, L., 2008, this volume, Provenance, depositional setting, and tectonic implications of Silurian polymictic conglomerates in Alaska’s Alexander terrane, in Blodgett, R.B., and Stanley, G.D., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy
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Printed in the USA
The Geological Society of America Special Paper 442 2008
Silurian Gastropoda from the Alexander Terrane, southeast Alaska David M. Rohr Department of Earth and Physical Sciences, Sul Ross State University, Alpine, Texas 79832, USA Robert B. Blodgett U.S. Geological Survey–Contractor, 4200 University Drive, Anchorage, Alaska 99508, USA
ABSTRACT Gastropods are described from Ludlow-age strata of the Heceta Limestone on Prince of Wales Island, southeast Alaska. They are part of a diverse megabenthic fauna of the Alexander terrane, an accreted terrane of Siberian or Uralian affinities. Heceta Limestone gastropods with Uralian affinities include Kirkospira glacialis, which closely resembles “Pleurotomaria” lindströmi Oehlert of Chernyshev, 1893, Retispira cf. R. volgulica (Chernyshev, 1893), and Medfracaulus turriformis (Chernyshev, 1893). Medfracaulus and similar morphotypes such as Coelocaulus karlae are unknown from rocks that are unquestionably part of the North American continent (Laurentia) during Late Silurian time. Beraunia is previously known only from the Silurian of Bohemia. Pachystrophia has previously been reported only from western North American terranes (Eastern Klamath, York, and Farewell terranes) and Europe. Bathmopterus Kirk, 1928, is resurrected and is only known from the Silurian of southeast Alaska. Newly described taxa include Hecetastoma gehrelsi n. gen. and n. sp. and Baichtalia tongassensis n. gen. and n. sp. Keywords: Alaska, gastropods, terranes, Ludlow, Prince of Wales Island. INTRODUCTION
Heceta Limestone of southeast Alaska that is highly provincial and indicative of a non–North American origin for the Alexander terrane, one of the primary component terranes of this part of Alaska, in which the Heceta Limestone is a key element of its Silurian stratigraphic succession. The Heceta Limestone, named by Eberlein and Churkin (1970), is a dominantly thick to massive bedded platform limestone unit that contains locally thick lenses of polymictic conglomerate, limestone breccia, sandstone, and argillaceous rocks. The type section was designated on the northeast side of Warm Chuck Inlet on the eastern side of Heceta Island. The formation was recognized to be of variable thickness over short lateral distances (Eberlein and Churkin, 1970, p. 18), with the thickest known
Paleozoic gastropods have proven in the past two decades to have great utility in the study of Paleozoic biogeographic affinities and possible origins of accreted terranes along the western margin of North America (Blodgett, 1992; Blodgett et al., 2002, 2003; Frýda and Blodgett, 2004; Rohr and Blodgett, 2003a; Rohr et al., 2003). This is due to dominance of faunas of these ages by the members of the Archaeogastropoda, a group characterized by direct or short larval life spans, not permitting wide-scale geographic dispersal such as is commonly found among many more “advanced” gastropod groups. In this paper we describe a diverse Ludlovian (middle Late Silurian) gastropod assemblage from the
Rohr, D.M., and Blodgett, R.B., 2008, Silurian Gastropoda from the Alexander Terrane, southeast Alaska, in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 51–61, doi: 10.1130/2008.442(03). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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section situated on western Heceta Island, where it attains a minimum thickness of 10,000 ft. The formation is extensively developed on the western part of Prince of Wales Island and its neighboring outlying islands to the west (i.e., Heceta, Tuxekan, and Kosciusko Islands). A number of lithofacies can be recognized within the unit, including massive algal (or microbial) reef complexes (Soja and Antoshkina, 1997), thinner-bedded forereef deposits, as well as restricted lagoonal (back-reef facies). The unit has been recognized to range in age from Middle to Late Silurian by Eberlein and Churkin (1970) and from late Early (late Llandoverian) to Late Silurian (Ludlovian) by Ovenshine and Webster (1970). The Heceta Limestone contains a rich megafauna, some elements of which have previously been described and/or illustrated elsewhere: brachiopods (Kirk, 1922, 1925, 1926; Kirk and Amsden, 1952; Savage, 1989); sphinctozoan sponges (Rigby et al., 1994); and corals (Oliver, 1964; Oliver et al., 1975). However, the greater part of the megafauna remains undescribed to date. Elements currently undergoing further study include brachiopods (Boucot and Blodgett), scaphopods (Rohr and Blodgett), and corals (A.E.H. Pedder). Calcareous algae are extremely common in the algal (microbial) reef complexes of the Heceta Limestone (Riding and Soja, 1993; Soja and Riding, 1993). The large bivalve Pycinodesma is typical of many back-reef (lagoonal) localities in the Heceta Limestone (Rohr and Blodgett, personal observ.), and although it is typical of age-correlative strata of the Willoughby Limestone and Kennel Creek Limestone (Kirk, 1927a, 1927b) farther to the north within the Alexander terrane, formal illustrations and study of specimens from the Heceta Limestone are lacking. Likewise, gastropods are common and have been described from lagoonal facies of the correlative Willoughby Limestone of the Glacier Bay area to the north (Kirk, 1928; Rohr and Blodgett, 2003a; Rohr et al., 2003). The mollusc-rich silicified fauna described here was recovered by acidization of the limestone from a locality in the Heceta Limestone on Prince of Wales Island. Gastropods and calcareous algae are most abundant, followed by tabulate corals (heliolitids are notably common), minor brachiopods, scaphopods, bivalves, and crinoids. The described specimens are from field locality 2002RB32 (Fig. 1), talus blocks derived from a horizon near the top of the wall face at the northwest corner of a large quarry in the Heceta Limestone situated on the border between sections 20 and 21, T70S, R81E, Craig D-4 quadrangle (55°47.26' N, 133°00.74' W; UTM 8 624531E 6184192N, NAD 83). The limestone is a silty wackestone to packstone and the contained gastropod and brachiopod elements of the fauna indicate an undifferentiated Ludlovian age. This locality plots out in the DSs unit (“rocks of Staney Creek and Tuxekan Passage region”) of Eberlein et al. (1983). This unit was considered by them to represent a transitional facies of the Heceta Limestone and overlying Karheen Formation developed to the east along the western side of Prince of Wales Island. Our field study suggests that both the Heceta Limestone and Karheen Formation can still be recognized in this region, but that the Heceta Limestone is much thinner than to the
Figure 1. Silicified gastropods are from locality 2002RB32, situated in a quarry in the Heceta Limestone, Craig D-4 quadrangle, Prince of Wales Island, southeast Alaska.
west on Heceta Island, and can be divided into two subunits on western Prince of Wales Island: (1) an underlying, dominantly clastic sequence that is correlative with the basal Heceta to the west; and (2) an overlying, dominantly carbonate succession that includes our fossil gastropod locality reported here. Both subunits are not completely exposed, but we would estimate a minimum thickness of several hundred meters for both. PALEOBIOGEOGRAPHIC AFFINITIES OF THE HECETA LIMESTONE GASTROPOD FAUNA AND RELATED ALEXANDER TERRANE FAUNAS The Alexander terrane of southeastern Alaska is a major piece of the accretionary terrane collage that forms much of
Silurian Gastropoda from the Alexander Terrane southeast Alaska. Ordovician, Silurian, and Devonian rocks of Prince of Wales and neighboring islands have been interpreted to be the remnants of an island arc that was rifted from near the Urals or Siberia and later sutured to North America (Blodgett et al., 2002, 2003). As presently recognized, Alexander terrane rocks extend from southeast Alaska, northward into northwestern British Columbia and Yukon Territory, to its western terminus in the Wrangell Mountains of south-central Alaska (Gehrels and Berg, 1994). The paleobiogeographic affinities of some taxonomic groups from the terrane consistently show no similarities with North American faunas from the western Cordillera of the United States or western or Arctic Canada, but rather demonstrate Siberian and/or Uralian affinities. The Asiatic Russian character of the Lower Devonian rugose and tabulate corals from the Alexander terrane was noted by Churkin et al. (1969) and Tchudinova et al. (1974). The Late Silurian brachiopod fauna of the Alexander terrane shows its strongest affinity with that of the Ural Mountains of Russia, as is well demonstrated by the large, distinctive pentamerid genera Brooksina Kirk, 1922, Harpidium Kirk, 1925, and Cymbidium Kirk, 1926, all based on specimens found in the Heceta Limestone in the area of Prince of Wales Island (see Blodgett et al., 2002, for detailed summation of the biogeographic affinities of the Late Silurian brachiopods of the Alexander terrane). Like the Farewell terrane of southwest Alaska, the Heceta Limestone of the Alexander terrane contains extensive buildups of Late Silurian algal-reef-mound complexes, containing an algal flora and associated sphinctozoan sponge complex known also in the Urals and the Farewell terrane (Riding and Soja, 1993; Soja and Riding, 1993; Rigby et al., 1994; Soja and Antoshkina, 1997; Soja et al., 2000). Similar buildups are unknown from nonaccreted rocks of equivalent age in North America. Although few faunal studies exist on Late Silurian gastropods from either Alaska or Russia, our limited database indicates strong faunal ties between Late Silurian gastropod fauna of the Alexander terrane and that of the Farewell terrane of southwestern Alaska (Rohr and Blodgett, 2003a; Rohr et al., 2003) and the Ural Mountains (Chernyshev, 1893). The work of Chernyshev focused on Early Devonian faunas of the Urals; however, the gastropod fauna that is particularly similar to the Late Silurian Alexander and Farewell terrane faunas is that reported from the Taltiya River (at the mouth of the Bobrovka), now considered to be of Ludlovian (middle Late Silurian) age (Melnikov and Khodalevich, 1965, p. 177). Prior publications on Silurian gastropods of Alaska are few and consist of four papers (Kirk, 1928; Rohr and Blodgett, 2003a, 2003b; Rohr et al., 2003) and two abstracts (Blodgett and Rohr, 1990, 1991). The most significant co-occurrence in southeast Alaska and the eastern Urals is Medfracaulus turriformis (Chernyshev, 1893). This distinctive species with its slow rate of expansion and deep umbilicus exhibits an unusual curved axis of coiling in some specimens. The curved spire is also seen in some shells of the same genus from west-central Alaska (Rohr, unpub. data).
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Medfracaulus is characteristic of Upper Silurian strata of westcentral Alaska (Nixon Fork subterrane of the Farewell terrane) and southeastern Alaska (Alexander terrane). Medfracaulus and similar morphotypes such as Coelocaulus karlae are unknown from rocks located undoubtedly on the North American continent (Laurentia) during the Late Silurian. Other Heceta Limestone gastropods with Uralian affinities include Kirkospira glacialis, which appears similar to Pleurotomaria lindströmi Oehlert of Chernyshev, 1893, and a species of Retispira that is similar to Bellerophon volgulicus Chernyshev, 1893. Beraunia is known only from the Silurian of Bohemia. Pachystrophia has previously been reported only from accreted terranes of western North America (Alaska and California) and Europe. In summary, the paleobiogeographic affinities of the Late Silurian gastropod fauna of the Alexander terrane, like those of the brachiopods, sponges and algal-reef-mound complexes, indicate close ties with coeval biotas of the Urals. TAXONOMIC NOTES Family BELLEROPHONTIDAE M’Coy, 1851 Genus RETISPIRA Knight, 1945 Retispira aff. R. volgulica (Chernyshev, 1893) Figures 2.1–2.3 Description. Small (7 mm wide), planispiral, rapidly expanding rounded shell with fine spiral threads crossed by fine growth lines; narrow umbilici, aperture slightly expanded, inductral deposits thin. Illustrated specimen. USNM 528217. Discussion. Retispira is rare at this locality. From the few specimens, it appears very similar in size, shape, and ornament to the Ludlow-age “Bellerophon” volgulicus Chernyshev, 1893, from the Urals. Family EUOMPHALIDAE de Koninck, 1881 Genus PACHYSTROPHIA Perner, 1903 Pachystrophia gotlandica (Lindström, 1884) Figures 2.10–2.13 Euomphalus gotlandicus Lindström, 1884, p. 139, plate 13, figures 19–31. Description. Flat to slightly concave spire; phaneromphalous shells with rounded, vertically elongated whorl; suture impressed, each whorl ~4 times as wide as previous; weak growth lines indicate shallow U-shaped sinus on shoulder of whorl; about half of previous whorl visible in umbilicus. Illustrated specimens. USNM 528220 and 528221. Discussion. Ordovician species of Pachystrophia from Alaska have spiral threads. Pachystrophia has been reported from the Upper Ordovician of the Seward Peninsula (Rohr, 1988), west-central Alaska (Rohr and Blodgett, 1985), and the Eastern Klamath terrane of northern California (Rohr, 1980), but not
Figure 2. All specimens are from locality 2002RB32 in the Heceta Limestone. (1–3) Retispira aff. R. volgulica (Chernyshev, 1893), apertural, umbilical, and dorsal views, × 3, USNM 528217. (4–6) Operculum of Oriostoma sp., exterior, side, and interior views, × 3, USNM 528218. (7–9) Operculum of oriostomatid gastropod, exterior, interior, and side views, × 3, USNM 528219. (10–13) Pachystrophia gotlandica (Lindström, 1884). (10–12) Apertural, top, basal views, × 3, USNM 528220. (13) Top view showing growth lines and sinus at shoulder of whorl, × 2, USNM 528221. (14–16) Beraunia bifrons (Perner, 1903), oblique apertural, oblique abapertural, and top views, × 3, USNM 528223. (17–19) Baichtalia tongassensis n. gen., n. sp., abapertural, oblique side and apertural views, × 2, USNM 528224.
Silurian Gastropoda from the Alexander Terrane cratonic North America. The genus also is known from the Upper Ordovician of Europe. Lindström’s (1884) illustrations of Euomphalus gotlandicus from the Silurian (Wenlockian) of Gotland are very similar in size and shape, but have a slightly wider umbilicus. Lindström (1884, p. 139) noted considerable variation in the form of the species’ whorls. Family ORIOSTOMATIDAE Wenz, 1938 Operculum of oriostomatid gastropod Figures 2.7–2.9 Description. Conical-shaped operculum, circular in cross section, closely spaced concentric growth lines on exterior; interior smooth and concave surrounded by a rounded, raised rim. Illustrated specimen. USNM 528219. Discussion. The operculum is a thicker version of the tabular operculum described below, and is similar to Silurian opercula illustrated by Lindström (1884, pl. 17, figs. 32–35 and 49) and Koken (1925, pl. 33, figs. 15a–15b). None were found in life position, but they may correspond to either of the oriostomatid gastropods from this locality. Genus ORIOSTOMA Munier-Chalmas, 1876 Operculum of Oriostoma sp. Figures 2.4–2.6 Description. Circular, disk-shaped operculum, ~2 mm in thickness; closely spaced, concentric growth lines on the exterior, wider toward edge; interior surface smooth and concave surrounded by raised, rounded rim. Illustrated specimen. USNM 528218. Discussion. The description is based on a single specimen. It is typical of the operculum of Oriostoma (see Lindström, 1884, pl. 17) and may correspond to the species from this locality; however, none were found in place. Similar opercula have been reported from the Lower Silurian (Llandovery) of the Taylor Mountains D-2 quadrangle, and the Lower Devonian (Emsian) Medfra B-4 quadrangle (Rohr and Blodgett, 2003a). The interior of these opercula is convex rather than concave. Genus BERAUNIA Knight, 1937 Beraunia includes the objective synonym Cyclotropis Perner, 1903, and the invalid Rhabdospira Perner, 1903. Beraunia bohemica (Perner, 1903) Figures 3.1–3.6 Rhabdospira bohemica Perner, 1903, plate 76, figures 9–12. Lytospira bohemica Perner, 1903, plate 74, figures 1–4. Cyclotropis bohemica Perner, 1907, p. 194. Description. Discoidal shell up to 3 cm in diameter with all but first two whorls out of contact, whorl
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cross section circular; ornament of up to 25 spiral cords crossed by weaker growth lines; aperture radial with no re-entrants; inside surface of whorl smooth, featureless; last 5 mm of a probable gerontic specimen (Fig. 3.5–3.6) shows no spiral ornament and no shell thickening. Illustrated specimens. USNM 528225 and 528226. Discussion. Perner (1903) reported the species from Late Silurian strata of the Kopanina Formation at Dlouhá Hora, Vyskočilka, and Kosoř in Bohemia, and illustrated the species with a convex concentric operculum. Beraunia bifrons (Perner, 1903) Figures 2.14–2.16 Rhabdospira bifrons Perner, 1903, plate 75, figures 28–31. Cyclotropis bifrons Perner, 1907, p. 196, plate 120, figures 1–3. Description. Discoidal, flat-spired shell with rounded whorls, deep sutures, and a wide umbilicus; ornamentation of at least 15 rounded, spiral threads. Illustrated specimen. USNM 528223. Discussion. The shell appears identical to Perner’s (1903) Bohemian species. Perner (1903) reported the species from Late Silurian age beds of the Kopanina Formation at Dlouhá Hora in Bohemia. Genus HECETASTOMA new genus Diagnosis. Extremely high spired, loosely coiled to uncoiled oriostomatids with strongly developed, widely spaced, spiral cords intersected by welldeveloped growth lines and bearing a prominent basal frill. Comparison. The extremely high spired nature of this loosely coiled to uncoiled shell in combination with its prominent basal frill clearly distinguishes this genus from all other previously described oriostomatid genera. Included species. Hecetastoma gehrelsi n. sp. (described below). Etymology. Combination of the term Heceta (from Heceta Limestone) and the Greek word stoma (mouth). Hecetastoma gehrelsi n. sp. Figures 3.7–3.17 Diagnosis. Loosely coiled to uncoiled; high-spired; circular whorl profile; six strong spiral cords intersected by strongly developed growth lines; frill developed on basal cord of some specimens. Description. Loosely coiled to uncoiled; apical angle ~30°; rounded whorl cross section, narrow but deep umbilicus; six equally spaced, spiral cords, which may
Figure 3. All specimens are from locality 2002RB32 in the Heceta Limestone. (1–6) Beraunia bohemica (Perner, 1903). (1–3) Top, oblique apertural, and basal views, × 2, USNM 528225. (4–6) Last portion of a fragmentary whorl showing (4) circular profile, and (5, 6) closely spaced growth lines in latest stage of growth, × 2, USNM 528226. (7–17) Hecetastoma gehrelsi n. gen., n. sp. (7–10) Apertural, side, top and basal views of a fragmentary specimen, × 2, USNM 528227. (11–14) Apertural, side, top, and basal views of another fragmentary specimen with a welldeveloped basal frill, × 2, USNM 528228. (15, 17) Side and oblique apical views of early whorls, × 3, USNM 528229. (16) Side view of early whorls, × 4, USNM 528230. (18–19) Kirkospira glacialis Rohr and Blodgett, 2003, abapertural and apertural views, × 2, USNM 528231.
Silurian Gastropoda from the Alexander Terrane appear serrated where crossed by strongly developed growth lines or lamella; a more strongly developed flange or frill on basal cord; aperture radial and orthocline; scoop-like re-entrants where each of the spiral cords meets apertural lip; earliest whorls not known. Etymology. In honor of George E. Gehrels, Department of Geosciences, University of Arizona, Tucson, Arizona. Types. Holotype USNM 528229; paratypes USNM 528227, 528228, and 528230. Discussion. Although this shell is much higher spired than typical species of Oriostoma, it has strong similarities to that group and clearly belongs to the family Oriostomatidae. However, the possession of both a high spire and a prominent basal frill indicates that the species belongs to a new genus. The type species of Oriostoma is loosely coiled and has a variable number of revolving costae (Knight, 1941). Several similarly ornamented high-spired shells have been described from the Silurian. These shells, which have all the whorls in contact, include ?Oriostoma clarki Rohr et al., 1981, which has a more rounded whorl with finer revolving cords. Trochonema turritum Lindström, 1884, has about eight cords and a channeled labrum. Tubina and Semitubina are highspired oriostomatids that have uncoiling shells. Peel and Gubanov (1997) described an unrelated Silurian shell and concluded that the uncoiling reflected a change in lifestyle of the organism. Family PLETHOSPIRIDAE Wenz, 1938 Genus KIRKOSPIRA Rohr and Blodgett, 2003 Kirkospira glacialis Rohr and Blodgett, 2003 Figures 3.18 and 3.19 Figures 4.1–4.3 Kirkospira glacialis Rohr and Blodgett, 2003b, p. 118, plates 1 and 2. Description. Moderately high spired (apical angle ~95°), minutely phaneromphalus shells, up to 5.5 cm high, with a vertically elongated, rounded whorl profile; narrow, raised selenizone located above mid-whorl, bounded by spiral threads. Whorl surface curves convexly downward and outward from impressed suture to selenizone and then curves convexly downward and inward to meet umbilicus and slightly thickened columellar lip. Illustrated specimen. USNM 528231 and 528232. Discussion. The genus is previously known only from larger, unsilicified specimens from the Silurian of Willoughby Island, Glacier Bay, Alaska (Rohr and Blodgett 2003b). The species appears to be very similar to “Pleurotomaria” lindströmi Oehlert described by Chernyshev (1893) (reproduced here as Fig. 5) from the Urals.
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Genus BATHMOPTERUS Kirk, 1928 Bathmopterus was established by Kirk in 1928 for shells from Glacier Bay, Alaska. The genus is known only from the Silurian of southeast Alaska, and B. liratus is the only species included. For reasons summarized below, Knight (1941) placed Bathmopterus in synonymy with Euomphalopterus C.F. Roemer, 1876. We here resurrect Bathmopterus as a distinct and valid genus. Bathmopterus liratus Kirk, 1928 Figures 4.4–4.10 Bathmopterus liratus Kirk, 1928, p. 1, plate 1, figures 1–5. Description. Low-spired, rounded, phaneromphalous gastropod with wide skirt-like frill; whorl cross section vertically compressed oval; slightly less than half of each previous convex whorl visible in umbilicus. Thick frill originates at mid-whorl and thins away from shell; frill curves concavely downward and then back upward. Upper suture and top of whorl covered by frill from previous whorl; frill on earlier whorls generally broken off even with edge of succeeding whorl. Growth lines indicate the presence of a U-shaped sinus just outward of and below whorl crest; growth lines below sinus curve strongly forward on frill. Illustrated specimens. USNM 528233 through 528237. Discussion. Kirk (1928, p. 2) concluded that Bathmopterus may readily be distinguished from Euomphalopterus by its wide umbilicus, its deep apertural notch and selenizone, and the very different character of it marginal flange. Knight (1941, p. 50) disagreed about the presence of an apertural notch and selenizone: “I am convinced that Kirk was misled in believing that this specimen retained the outer shell layers showing growth lines on its outer whorl face and upper surface of the frill. In my opinion, the outer shell layers on the areas in question clung closely to the external matrix and were removed with it, exposing the inner probably originally nacreous shell layers, the laminae of which have a very different course from the external lines of growth.” Kirk’s (1928) specimens are unsilicified and were broken free from a massive limestone matrix. Our specimens are silicified, so no shell layers were broken away during preparation. Unless some unusual process of diagenesis did not silicify the outer shell layer, the sinus and selenizone are real external features of the shell. As noted by Knight (1941), many excellent specimens of Euomphalopterus exist and none exhibit a sinus or a selenizone. The presence of these features on Bathmopterus liratus shows that Bathmopterus should be restored as a valid genus.
Figure 4. All specimens are from locality 2002RB32 in the Heceta Limestone. (1–3) Kirkospira glacialis Rohr and Blodgett, 2003, abapertural view, view of nearly complete aperture, and basal view showing the narrow umbilicus, × 2, USNM 528232. (4–10) Bathmopterus liratus Kirk, 1928. (4) Lenticular profile of a single whorl showing frill, × 2, USNM 528233. (5, 6) Top and oblique views of selenizone and frill, × 2, USNM 528234. (7) Top view of shell fragment with selenizone and forward-sweeping growth lines, × 2, USNM 528235. (8) Umbilical view of shell with part of frill preserved, × 2, USNM 528236. (9, 10) Cross sectional and umbilical views, × 2, USNM 528237. (11–12) Coelocaulus karlae Rohr, Blodgett, and Frýda, 2003. (11) Side view of early whorls with curved axis of coiling, × 2, USNM 528238. (12) Side view of later whorls of a large specimen, × 2, USNM 528239.
Silurian Gastropoda from the Alexander Terrane
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Figure 5. “Pleurotomaria” lindströmi Oehlert as illustrated by Chernyshev (1893) from the Urals. (1) Apertural view. (2) Side view. Note close similarity to Kirkospira glacialis Rohr and Blodgett, 2003, illustrated here as Figures 3.18–3.19 and 4.1–4.3.
Genus MEDFRACAULUS Rohr, Blodgett and Frýda, 2003 Medfracaulus turriformis (Chernyshev, 1893) Figures 6.1 and 6.2 Murchisonia turriformis Chernyshev, 1893, p. 36, plate 2, figures 9–11. Medfracaulus cooki Rohr, Blodgett and Frýda, 2003, p. 91, figure 1.1–1.3. Description. High-spired, slowly expanding gastropods with narrow but very deep umbilicus; mid-whorl angulation closer to lower suture line than to upper suture. About 10% of the specimens have curved axis of coiling. Illustrated specimen. USNM 528222. Discussion. The deep, narrow, straight-walled umbilicus, slow rate of expansion, and selenizone below mid-whorl are diagnostic of Medfracaulus. Some specimens from the type locality of the genus in west-central Alaska have a curved axis of coiling, but we dismissed it as tectonic deformation of the shells. The lack of deformation in any other Heceta Limestone shells indicates that the curving axis is a feature of the species. The Heceta Limestone specimens appear identical to Murchisonia turriformis Chernyshev, 1893, from the Urals, and the Russian species is assigned here to Medfracaulus. Chernyshev (1893) originally described the species as “slowly expanding in height and significantly faster in width retaining an open umbilicus” (translation by RBB). Chernyshev (1893) illustrated only one specimen, and that specimen has a curved axis of coiling (re-illustrated herein as Fig. 6.1). His description of the species, however, does not mention the feature. The Russian specimens are from the Taltiya River (at the mouth of the Bobrovka), now considered to be of Ludlovian (middle Late Silurian) age (Melnikov and Khodalevich, 1965, p. 177).
Figure 6. Comparison of Medfracaulus turriformis (Chernyshev, 1893) specimens with curved axis of coiling. (1) From the eastern Urals, Russia, × 1. (2) From the Heceta Limestone, × 2, USNM 528222.
Genus COELOCAULUS Oehlert, 1888 Coelocaulus karlae Rohr, Blodgett, and Frýda, 2003 Figures 4.11 and 4.12 Coelocaulus karlae Rohr, Blodgett, and Frýda, 2003, p. 89, figure 1.4–1.7. Description. High-spired (apical angle 35°), deeply phaneromphalous gastropods with selenizone at mid-whorl; suture impressed, right and left shoulders convex, selenizone slightly raised. Growth lines poorly known; prosocline above selenizone. Base of whorl convex; sharp, circumumbilical angulation; umbilicus deep, flat-walled. Whorl profile D-shaped; the umbilical width is seen to increase at a slower rate than the width of a single whorl; in earlier stages the umbilicus is about equal to the whorl width, and in later stages about half the width. Illustrated specimens. USNM 528238 and 528239. Discussion. Coelocaulus karlae attains a larger size than Medfracaulus turriformis. In addition to its much larger size, this species can be distinguished from C. davidsoni by its lower rate of translation along the axis of coiling and its flatter whorls. It has a much wider umbilicus than Clarke and Ruedemann’s (1903) Silurian occurrences of Coelidium macrospira
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Rohr and Blodgett (Hall, 1852) and C. cf. vitellia (Billings, 1865). With its low whorls and wide umbilicus, the shell of Coelocaulus karlae resembles some Silurian species of Coelocaulus from the Prague Basin. C. pollens is the most similar of the Bohemian species to Coelocaulus karlae (see Perner 1907, pl. 100, figs. 52–54). Both species have the same general shape, relative width of umbilicus, as well as shell size, but Coelocaulus pollens is based on an internal mold, so more detailed comparison of both species is impossible. The relatively wide umbilicus as well as low whorls unite some species of Coelocaulus, including C. karlae, in a well-defined morphological group. Genus BAICHTALIA new genus Diagnosis. High-spired; slowly expanding, rounded whorls slightly out of contact; surface smooth except for selenizone above mid-whorl. Included species. Baichtalia tongassensis n. sp. (described below). Etymology. In honor of James Baichtal, Forest Geologist, Tongass National Forest, Alaska. Baichtalia tongassensis n. sp. Figures 2.17–2.19 Diagnosis. By monotypy, same as the genus. Description. Rounded, high-spired (apical angle of ~35°), slightly uncoiled shells; ~2 cm in diameter; surface smooth except for flat selenizone above midwhorl surrounded by raised edges; whorls circular in cross section; earliest whorls and aperture unknown. Etymology. Species named after Tongass National Forest. Types. Holotype USNM 528224. Discussion. Open coiling and uncoiling in gastropods have been reviewed by Yochelson (1971) and Peel and Gubanov (1997), who noted that this type of shell form is uncommon in the post-Paleozoic. In the Paleozoic open coiling is more common in discoidal than in high-spired shells. Odontomaria linsleyi Rohr and Smith, 1978, from the Lower Devonian of Prince of Wales Island, Arctic Canada, has a similar form, but it is much more openly coiled.
ACKNOWLEDGMENTS We thank James F. Baichtal, U.S. Forest Service, Tongass National Forest, Thome Bay, Alaska, and Susan M. Karl, U.S. Geological Survey, Anchorage, Alaska, for sharing their knowledge of the local geology during our fieldwork on Prince of Wales Island. We are grateful to Doris Heidelberger, Oberursel, Germany, and A.J. Boucot, Dept. of Zoology, Oregon State University, Corvallis, Oregon, for their helpful comments on the manuscript. Fieldwork was supported in part by the U.S.
Forest Service and a grant from the National Geographic Society to Rohr. REFERENCES CITED Billings, E., 1865, Palaeozoic fossils, Volume 1: Ottawa, Geological Survey of Canada, 426 p. Blodgett, R.B., 1992, Taxonomy and paleobiogeographic affinities of an early Middle Devonian (Eifelian) gastropod faunule from the Livengood quadrangle, east-central Alaska: Palaeontographica, Abt. A, v. 221, p. 125–168. Blodgett, R.B., and Rohr, D.M., 1990, Silurian-Devonian gastropod biogeography of Alaska: Geological Society of America Abstracts with Programs, v. 22, no. 7, p. A221. Blodgett, R.B., and Rohr, D.M., 1991, Silurian-Devonian gastropods in Alaska, in Abstracts and proceedings, Annual Meeting of the Western Society of Malacologists on Current Directions in Alaskan Malacology: Western Society of Malacologists Annual Report, v. 23, p. 15–16. Blodgett, R.B., Rohr, D.M., and Boucot, A.J., 2002, Paleozoic links among some Alaskan accreted terranes and Siberia based on megafossils, in Miller, E.L., Grantz, A., and Klemperer, S.L., eds., Tectonic evolution of the Bering Shelf-Chukchi Sea-Arctic Margin and adjacent landmasses: Boulder, Colorado, Geological Society of America Special Paper 360, p. 273–290. Blodgett, R.B., Rohr, D.M., Karl, S.M., and Baichtal, J.F., 2003, Early Middle Devonian (Eifelian) gastropods from the Wadleigh Limestone in the Alexander terrane of southeastern Alaska demonstrate biogeographic affinities with central Alaskan terranes (Farewell and Livengood) and Eurasia, in Galloway, J.P., ed., Studies by the U.S. Geological Survey in Alaska, 2001: U.S. Geological Survey Professional Paper 1678, p. 105–115. Chernyshev, F.N., 1893, Fauna nizhnyago devona vostochnago sklona urala: Trudy Geologicheskago Komiteta, v. 4, no. 3, 221 p. Churkin, M., Jr., Eberlein, G.D., Hueber, F.M., and Mamay, S.H., 1969, Lower Devonian land plants from graptolitic shale in southeastern Alaska: Palaeontology, v. 12, no. 4, p. 559–573. Clarke, J.M., and Ruedemann, R., 1903, Guelph fauna in the State of New York: Memoir of the New York State Museum and Science Service, v. 5, 195 p. de Koninck, L.G., 1881, Faune du calcaire carbonifère de Belgique, Part 3, Gastéropodes. Annales du Musée Royal d’Histoire Naturelle de Belgique, paleontological series, 6:1–170. Eberlein, G.D., and Churkin, M., Jr., 1970, Paleozoic stratigraphy in the northwest coastal area of Prince of Wales Island, southeastern Alaska: U.S. Geological Survey Bulletin 1284, 67 p. Eberlein, G.D., Churkin, M., Jr., Carter, C., Berg, H.C., and Ovenshine, A.T., 1983, Geology of the Craig quadrangle, Alaska: U.S. Geological Survey Open-File Report 83-91, 53 p., scale: 1:250,000, 4 sheets. Frýda, J., and Blodgett, R.B., 2004, New Emsian (Late Early Devonian) gastropods from Limestone Mountain, Medfra B-4 quadrangle, west-central Alaska (Farewell terrane), and their paleobiogeographic affinities and evolutionary significance: Journal of Paleontology, v. 78, p. 111–132, doi: 10.1666/0022-3360(2004)078<0111:NELEDG>2.0.CO;2. Gehrels, G.E., and Berg, H.C., 1994. Geology of southeastern Alaska., in Plafker, G. and Berg, H.C., eds., The geology of Alaska: Boulder, Colorado, Geological Society of America, Geology of North America, v. G-1, p. 451–467. Hall, J., 1852, Descriptions of the organic remains of the lower middle division of the New York system: Paleontology of New York, v. 2, 362 p. Kirk, E., 1922, Brooksina, a new pentameroid genus from the Upper Silurian of southeastern Alaska: Proceedings of the United States National Museum, v. 60, article 19, p. 1–8. Kirk, E., 1925, Harpidium, a new pentameroid brachiopod genus from southeastern Alaska: Proceedings of the United States National Museum, v. 66, article 32, p. 1–7. Kirk, E., 1926, Cymbidium, a new genus of Silurian pentameroid brachiopods from Alaska: Proceedings of the United States National Museum, v. 69, article 23, p. 1–5. Kirk, E., 1927a, Pycnodesma, a new molluscan genus from the Silurian of Alaska: Proceedings of the United States National Museum, v. 71, article 20, p. 1–9. Kirk, E., 1927b, Pycinodesma, a new name for Pycnodesma Kirk not Schrammen: Journal of the Washington Academy of Sciences, v. 17, p. 543.
Silurian Gastropoda from the Alexander Terrane Kirk, E., 1928, Bathmopterus, a new fossil gastropod genus from the Silurian of Alaska: Proceedings of the United States National Museum, v. 74, article 18, p. 1–4. Kirk, E., and Amsden, T.W., 1952, Upper Silurian brachiopods from southeastern Alaska: U.S. Geological Survey Professional Paper 233-C, p. 53–66. Knight, J.B., 1937, Genotype designations and new names for invalid homonyms among Paleozoic gastropod genera: Journal of Paleontology, v. 11, p. 709–714. Knight, J.B., 1941, Paleozoic gastropod genotypes: Geological Society of America Special Paper 32, 510 p. Knight, J.B., 1945, Some new genera of Paleozoic Gastropoda: Journal of Paleontology, v. 19, p. 573–587. Koken, E., 1925, Die Gastropoden des Baltischen Untersilurs (J. Perner, ed.): Mémoires de l’Académie des Sciences de Russie, ser. 8, Classe PhysicoMathématique, v. 37, no. 1, 326 p. Lindström, G., 1884, The Silurian Gastropoda and Pteropoda of Gotland: Kongliga Svenska Vetenskaps-Akademiens Handlingar, 250 p. M’Coy, F., 1851, On some new Silurian Mollusca: Annals and Magazine of Natural History, including Zoology, Botany, and Geology, 2nd ser., p. 45–63. Melnikov, A.S., and Khodalevich, A.N., 1965, Vostochnogo sklon urala— severnyi i srednii ural, in Nikiforova, O.I., and Obut, A.M., eds., Stratigrafiya SSSR v chetyrnatsati tomakh, Siluriiskaya Sistema: Moscow, Izdatelstvo “Nedra,” p. 171–182. Oliver, W.A., Jr., 1964, New occurrences of the rugose coral Rhizophyllum in North America: U.S. Geological Survey Professional Paper 475-D, p. D149–D158. Oliver, W.A., Jr., Merriam, C.W., and Churkin, M., Jr., 1975, Ordovician, Silurian, and Devonian corals of Alaska: U.S. Geological Survey Professional Paper 823-B, p. B13–B44. Ovenshine, A.T., and Webster, G.D., 1970, Age and stratigraphy of the Heceta Limestone in northern Sea Otter Sound, southeastern Alaska, in Geological Survey Research 1970: U.S. Geological Survey Professional Paper 700-C, p. C170–C174. Peel, J.S., and Gubanov, A.P., 1997, Mode of life of an uncoiled Silurian gastropod from Siberia: Bulletin of the Czech Geological Survey, v. 72, p. 339–344. Perner, J., 1903, Gastéropodes, Tome 1, in J. Barrande, ed., Systême silurien du centre de la Bohême 4: Prague, Musée Bohême, 164 . Perner, J., 1907, Gastéropodes, Tome 2, in J. Barrande, ed., Systême silurien du centre de la Bohême 4: Prague, Musée Bohême, 380 p. Riding, R., and Soja, C.M., 1993, Silurian calcareous algae, cyanobacteria, and microproblematica from the Alexander terrane: Journal of Paleontology, v. 67, p. 710–728. Rigby, J.K., Nitecki, M.H., Soja, C.M., and Blodgett, R.B., 1994, Silurian aphrosalpingid sphinctozoans from Alaska and Russia: Acta Palaeontologica Polonica, v. 39, p. 341–391. Roemer, C.F., 1876, Lethaea geognostica oder Beschreibung und Abbildung der für die Gebirgs-Formationen bezeichnendsten Versteinerungen, Theil 1, Lethaea palaeozoica: Stuttgart, E. Schweizerbart’sche Verlagshandlung, Atlas with 62 plates.
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Rohr, D.M., 1980, Ordovician-Devonian Gastropoda from the Klamath Mountains, California: Palaeontographica, Abt. A, v. 171, p. 141–210. Rohr, D.M., 1988, Upper Ordovician gastropods from the Seward Peninsula, Alaska: Journal of Paleontology, v. 62, p. 551–566. Rohr, D.M., and Blodgett, R.B., 1985, Upper Ordovician Gastropoda from west-central Alaska: Journal of Paleontology, v. 59, p. 667–673. Rohr, D.M., and Blodgett, R.B., 2003a, Kirkospira, a new Silurian gastropod from Glacier Bay, Southeastern Alaska, in Galloway, J.P., ed., Studies by the U.S. Geological Survey in Alaska, 2001: U.S. Geological Survey Professional Paper 1678, p. 117–125. Rohr, D.M., and Blodgett, R.B., 2003b, Gastropod opercula from the Silurian and Devonian of Alaska, in Clautice, K.H., and Davis, P.K., eds., Short notes on Alaska Geology, 2001: Alaska Division of Geological and Geophysical Surveys Professional Report 120, p. 73–85. Rohr, D.M., and Smith, R.E., 1978, Lower Devonian Gastropoda from the Canadian Arctic Islands: Canadian Journal of Earth Sciences, v. 15, p. 1228–1241. Rohr, D.M., Boucot, A.J., and Perry, D.G., 1981, Silurian (Wenlockian) gastropods from Baillie-Hamilton Island, Canadian Arctic: Journal of Paleontology, v. 55, p. 331–339. Rohr, D.M., Blodgett, R.B., and Frýda, J., 2003, New Silurian murchisoniid gastropods from Alaska and a review of the genus Coelocaulus, in Clautice, K.H., and Davis, P.K., eds., Short notes on Alaska Geology, 2001: Alaska Division of Geological and Geophysical Surveys Professional Report 120, p. 87–93. Savage, N.M., 1989, The occurrence of the brachiopods Nanukidium and Atrypoidea in the Late Silurian of southeastern Alaska, Alexander terrane: Journal of Paleontology, v. 63, p. 530–533. Soja, C.M., and Riding, R., 1993, Silurian microbial associations from the Alexander terrane, Alaska: Journal of Paleontology, v. 67, p. 728–738. Soja, C.M., and Antoshkina, A.I., 1997, Coeval development of Silurian stromatolite reefs in Alaska and the Urals Mountain: Implications for paleogeography of the Alexander terrane: Geology, v. 25, p. 539–542, doi: 10. 1130/0091-7613(1997)025<0539:CDOSSR>2.3.CO;2. Soja, C.M., White, B., Antoshkina, A., Joyce, S., Mayhew, L., Flynn, B., and Gleason, A., 2000, Development and decline of a Silurian stromatolite reef complex, Glacier Bay National Park, Alaska: Palaios, v. 15, no. 4, p. 273–292. Tchudinova, I.I., Churkin, M., Jr., and Eberlein, G.D., 1974, Devonian syringoporoid corals from southeastern Alaska: Journal of Paleontology, v. 48, p. 124–134. Wenz, W., 1938, Gastropoda, Teil 1: Allgemeiner Teil und Prosobranchia, in Schindewolf, O., ed., Handbuch der Paläozoologie, Volume 6: Berlin, Bornträger, 240 p. Yochelson, E.L., 1971, A new Late Devonian gastropod and its bearing on problems of open coiling and septation, in Dutro, J.T., ed., Paleozoic perspectives: A paleontological tribute to G. Arthur Cooper: Smithsonian Contributions to Paleobiology, v. 3, p. 231–241. MANUSCRIPT ACCEPTED BY THE SOCIETY 14 DECEMBER 2007
Printed in the USA
The Geological Society of America Special Paper 442 2008
Provenance, depositional setting, and tectonic implications of Silurian polymictic conglomerates in Alaska’s Alexander terrane Constance M. Soja Lena Krutikov Department of Geology, Colgate University, Hamilton, New York 13346, USA
ABSTRACT The Heceta Formation of southeastern Alaska (Alexander terrane) comprises a 3000-m-thick limestone-siliciclastic deposit of Early–Late Silurian age. The limestones record the first widespread evidence of carbonate platform development in this ancient island arc. Interbedded polymictic conglomerates represent interruption in platform evolution during onset of the Klakas orogeny, an arc-continent collisional event that occurred in the Late Silurian–Early Devonian. Conglomerates grade upward into finer-grained siliciclastics capped by shallow-marine limestones in sequences that are 200–300 m thick. Clasts range in diameter from 2 to 30 cm, are subangular to well rounded, poorly to moderately sorted, and densely packed in disorganized, poorly stratified beds. Most of the clasts are volcanic (basaltic-andesitic), but limestone clasts predominate in some sections; rare fragments of volcaniclastic, plutonic, and indeterminate rocks also occur. Clast compositions match the lithology of rocks in the underlying Heceta and Descon formations, and sedimentary attributes indicate redeposition of recycled material by debris flows and rivers in a coastal alluvial fan complex. This evidence—together with affinities of marine fossils, paleomagnetic and detrital zircon data, associated Old Red Sandstone-like facies, and coincidence in timing of tectonism—suggests the Klakas orogeny was a Caledonide event that is manifest in Alaska’s Alexander terrane. Keywords: Silurian island arc, conglomerates, coastal alluvial fan, Klakas orogeny. INTRODUCTION
1980; Bazard et al., 1995; Soja et al., 2000; Soja, this volume). Persistent volcanism initiated in the Ordovician led to the construction of an arc edifice. Andesitic and basaltic breccias suggest a volcanic island may have been emergent before volcanism ceased in the mid Early Silurian (Gehrels et al., 1996) and when widespread limestones began to accumulate in a fringing marine
The Alexander terrane in southeastern Alaska represents a Proterozoic(?)–early Paleozoic island arc that was located near Laurentia, Baltica, or Siberia before accretion to western North America in the late Mesozoic–early Cenozoic (Coney et al.,
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[email protected] Soja, C.M., and Krutikov, L., 2008, Provenance, depositional setting, and tectonic implications of Silurian polymictic conglomerates in Alaska’s Alexander terrane, in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 63–75, doi: 10.1130/2008.442(04). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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platform in the late Early Silurian (Soja, 1990, 1993). Prolonged periods of tectonic quiescence persisted after volcanism waned, as indicated by Silurian-Permian limestones that exceed 5 km in thickness (Eberlein et al., 1983; Gehrels and Saleeby, 1987a). A significant punctuation in carbonate deposition was caused by the Klakas orogeny in the Late Silurian–Early Devonian (Gehrels et al., 1983, 1996). As a prelude to culminating orogenic phases, thick accumulations of polymictic conglomerates were interbedded with Silurian fossiliferous limestones. Previous research focused on the limestones to deduce the paleobiogeographic affinities and environmental setting of marine fossils (Soja, 1990, 1991, 1993, 1994, 2006; Soja and Antoshkina, 1997; Soja et al., 2000, 2003; Antoshkina and Soja, 2006). In this study, we investigated the conglomeratic deposits as environmental indicators of the paleogeographic setting and tectonic evolution of this displaced terrane (Miall, 1981; cf. Follo, 1992). In particular, we examined their sedimentology, composition, and stratigraphic relationships to ascertain provenance and depositional conditions associated with the onset of orogenic activity. Our results build on foundational studies (Soja, 1994; Soja and Antoshkina, 1997; Soja et al., 2000; Antoshkina and Soja, 2006) that refute earlier claims that the Alexander terrane was located in the Southern Hemisphere until the Devonian (Gehrels and Saleeby, 1987a, 1987b; Gehrels et al., 1996; Butler et al., 1997). When integrated with paleomagnetic, detrital zircon, faunal, and other sedimentologic indicators, these new data help establish for the first time a paleogeographic-tectonic link to Caledonide orogenesis in the Late Silurian–Early Devonian.
under restricted conditions, as suggested by the lack of diverse invertebrate faunas (Soja, 1990). Abrupt transition to massive, recrystallized limestones and to the polymictic conglomerates under study represents a disruption in carbonate platform development during the earliest pulses of the Klakas orogeny (Gehrels et al., 1983, 1996; Soja, 1993; Soja, this volume). Subsequent rejuvenation in carbonate platform evolution is reflected in Upper Silurian limestones that overlie the conglomerates (Soja, 1993). Lagoonal limestones with low-diversity, shallow-marine benthos grade upward into stromatolitic boundstones, which represent barrier reefs constructed by a diverse suite of microbial-sponge species at the seaward edge of a shallowmarine platform (Rigby et al., 1994; Soja, 1991, 1993; Soja et al., 2000). Similar biotas in the Farewell terrane of southwestern Alaska, the Ural Mountains, and Salair indicate placement of the Alexander terrane along the Uralian Seaway (Soja and Antoshkina, 1997; Soja et al., 2000; Antoshkina and Soja, 2006). This oceanic corridor favored transmigration of marine life between Laurentia, Baltica, and Siberia during the Late Silurian (Fig. 1). By the end of the Late Silurian, the margin of the shallow-marine platform in the Alexander terrane had collapsed, as indicated by the abrupt transition from shallow-water deposits to deep-marine turbidites, debris flow breccias with reef-derived clasts, and
GEOLOGIC SETTING In southeastern Alaska, the Descon Formation of Early Ordovician–Early Silurian age comprises volcanic flows, breccias, tuffs, and agglomerates interbedded with volcaniclastic conglomerates, wackes, mudstones, cherts, shales with minor limestone lenses, and quartzo-feldspathic arenites (Eberlein et al., 1983). These deposits are 3000 m thick; lithologic textures and compositional data suggest that they accumulated adjacent to active eruptive centers during convergent plate margin activity and arc-type magmatism (Eberlein et al., 1983; Gehrels and Saleeby, 1987a; Soja, this volume). Descon facies form shallowing-upward sequences gradational with the overlying Heceta Formation, a deposit of Early–Late Silurian age that exceeds 3000 m in thickness (Eberlein et al., 1983; Soja, 2006). An onshore-to-offshore suite of shallow- and deep-marine limestones is assigned to the Heceta Formation. It records the evolution of a carbonate platform during waning volcanism followed by subsidence, erosion, and marine transgression (Soja, 1990, 1993). At the base of the formation, finely laminated, graded lime mudstones are interbedded with skeletal grainstones (turbidites) and localized, massive rudstones and boundstones (microbialstromatoporoid-coral reefs). These deposits are overlain by shallow-water limestones (peloidal wackestones, packstones, and grainstones) that originated on an incipient carbonate platform
Figure 1. Paleogeographic map showing Alexander terrane (AT) and its postulated placement with respect to ancient Siberian and Laurussian landmasses (stippled areas) in the Late Silurian–Early Devonian. Paleomagnetic data indicate that the AT was located within 14° of the paleoequator in the Early Devonian (Bazard et al., 1995). Stars represent present-day sites where similar microbial-sponge communities occur at the AT, Farewell terrane (SW Alaska), western slope of the Urals (eastern Baltica), and Salair. Modified from Soja et al., 2000.
Silurian polymictic conglomerates in Alaska’s Alexander terrane stromatolite olistostromes (Soja, 1993; Soja et al., 2000; Soja, this volume). Evidence for rapid deepening followed by abrupt shallowing is derived from upper parts of the Heceta Formation, which become more argillaceous where they grade upward into the overlying Karheen Formation. Conglomerates, cross-bedded and pebbly sandstones, shales, and rare limestones in the Karheen Formation form a thick sequence (1800 m) of terrigenous red beds and shallow-marine deposits. These accumulated as a result of uplift, erosion, and clastic wedge progradation during culmination of the Klakas orogeny (Ovenshine, 1975; Eberlein et al., 1983; Gehrels et al., 1983, 1996; Soja, 1993, 2006). Paleomagnetic data indicate that the Alexander terrane was within 14° of the paleoequator by the Early Devonian (Bazard et al., 1995).
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MATERIALS, METHODS, AND STRATIGRAPHY Conglomerates in the southern part of the Alexander terrane were examined at six sites near Sea Otter Sound on Heceta, Orr, Tuxekan, and northwestern Prince of Wales Islands (Fig. 2). These coarse-grained deposits range in thickness from 200 to 300+ m in the east-southeast but may be as thick as 1000 m in the west-northwest (Eberlein et al., 1983; Soja, 1993). Although precise ages of the conglomerates are unknown, conodonts preserved in limestones interbedded with the conglomerates and also limestone clasts containing aphrosalpingid sponges indicate they are early to mid- (?) Ludlow in age (Ovenshine and Webster, 1970; Soja, 1993; Rigby et al., 1994). Petrographic analysis
Figure 2. Map showing the geology and location of sample sites 1–6 near Sea Otter Sound in southeastern Alaska (inset). Legend below map corresponds to pie diagrams, which denote clast % by lithology at these sites: 1—Sarkar Point (n = 52); 2—unnamed island west of Orr Island (n = 20); 3—southeast Tuxekan Island (n = 9); 4—north shore of Indian Garden Bay, Heceta Island (n = 68); 5—northeast shore of Warm Chuck Inlet, Heceta Island (n = 22); 6—southeast Tuxekan Island (n = 63). W—Wadleigh Formation. Geology modified from Eberlein et al. (1983).
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was used to assess average clast size, roundness and sphericity, sorting, and sedimentary structures. A total of 247 clasts and 61 petrographic samples was examined in thin section to compile compositional data (Fig. 2). Dense forests that cover the islands prevent any detailed investigation of rocks beyond the modern intertidal zone, where lateral exposure is limited to a narrow swath of shoreline. Furthermore, stratigraphic sequences along many beach sections are discontinuous because of erosion, marine embayments, or cover by glacial debris. The localities investigated reveal that the polymictic conglomerates grade upward into finer-grained siliciclastics, which show rare examples of sedimentary structures, such as cross-bedding or ripple lamination in the overlying sandstones, siltstones, and shales. Fossiliferous, shallow-marine limestones cap each sequence; limestone conglomerates, such as those exposed along Tuxekan Passage (Fig. 2, site 6), occur locally at the base of some polymictic units but are generally rare (Soja, 1993).
Figure 3. Generalized stratigraphic section showing upper 23 m of conglomerates exposed at Indian Garden Bay, eastern Heceta Island (Fig. 2, site 4). For legend, see Figure 7. Modified from Krutikov (1996).
One of the most continuous exposures allows the stratigraphic context to be reconstructed for conglomerates that crop out on eastern Heceta Island along Indian Garden Bay (Fig. 2, site 4; Figs. 3 and 4). Samples were collected for petrographic analysis at every 3 m from the uppermost 23 m of a graveldominated sequence, which may be 200–300 m thick where it is incompletely exposed to the south. Total thickness is difficult to assess because of numerous areas along the shoreline that are devoid of outcrop. Discontinuities in the conglomerate sequence may be caused by faults, and strata may be repeated; the absence of recurring key marker beds precludes confirmation of these hypotheses. Compositional data were obtained from thin sections of 20 clasts that were collected randomly at each of two localities and from an additional 38 clasts that were sampled randomly throughout the section (Fig. 2, site 4). Particle size was assessed for 45 clasts measured on site at three localities where small, medium, and large size grains are concentrated (Fig. 5). The units strike NW-SE (orientations vary in minor folds) and dip 54°–80° toward the northeast. As shown in Figures 3 and 4, most of the measured section in Indian Garden Bay comprises massive beds typical of all conglomeratic units exposed near Prince of Wales Island. Clasts are sub- to well-rounded, poorly to moderately sorted, densely concentrated, and polymictic in composition, generally composed of volcanic clasts and <10% limestone fragments (Fig. 6). Beds typically are sheet-like, crudely laminated, grain-supported with tightly packed pebbles or matrix-supported in a pebbly or sandy medium, and lack sedimentary structures except for rare evidence of imbrication (Fig. 4D) and normal (Fig. 4B) or reverse grading (Fig. 4C). For example, the basal 2 m of measured section shown in Figure 3 is a coarse conglomerate with well-rounded, relatively well sorted clasts that fine upward and comprise mostly volcanic rocks and <10% limestone fragments. This is overlain by a 0.2-m-thick fine-grained laminated mudstone that lacks abundant pebbles or clasts. From 2.5 to 3.5 m above the base, a tightly packed pebble horizon has a sandy matrix with small clasts coarsening upward into larger clasts in a reverse-graded conglomeratic bed. This is overlain by a conglomerate with a sandy matrix and large, fractured clasts that show a slight grading upward from the base. An abrupt change at 6.0 m above the base (Fig. 3) is marked by a 1.2-m-thick conglomerate with tightly packed, poorly sorted, well-rounded clasts. An erosional gully creates a discontinuity in the outcrop. Conglomerates with poorly sorted, tightly packed clasts occur from 7.7 to 9.5 m above the base of the section; these fine upward into mudstones. The overlying 2 m of section is composed of thin, alternating conglomerates and sandstones with small pebbly clasts. Sandstone units without pebbles become increasingly more prevalent at 12 m above the base. The section from 12 to 14.5 m above the base (Fig. 3) shows considerable lateral variation where sandy units are interbedded with conglomerates characterized by medium-sized cobbles. In a scour channel at 16.7 m above the base (Fig. 4C), sandstones are interbedded with conglomerates that comprise reverse-graded, medium-sized
Silurian polymictic conglomerates in Alaska’s Alexander terrane
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Figure 4. Conglomeratic section at Indian Garden Bay (Fig. 2, site 4). Locations denote where each sample was collected in meters above the base of the section. (A) Massive conglomerates interbedded with thick sandstones; loc. 7–10. (B) Graded cobbles; loc. 2.5. (C) Reverse-graded clasts concentrated in scour channel; loc. 16.7. (D) Rare evidence of imbricated clasts in massive sandstone; loc. 21. Hammer in A and C is 35 cm long; ruler in B and D is 15 cm.
clasts. Farther north along the shore, discontinuous outcrops of conglomerates fine upward into >200 m of sandstone, siltstone, mudstone, and shallow-marine limestones (Krutikov, 1996). These grade upward into lagoonal limestones, characterized by amphiporid stromatoporoids and Atrypoidea brachiopods, which in turn are overlain by cross-bedded sandstones at the base of the Karheen Formation. Another stratigraphic section exposed on eastern Tuxekan Island (Fig. 2, site 6; Figs. 7–9) deserves special attention because it allows inferences to be made about the provenance and depositional setting of conglomerates that comprise a greater proportion of limestone clasts. These conglomerates extend along the shoreline of Tuxekan Passage for 0.6 km and, in general, strike NE-SW with an average dip of 45–60°NW. Samples for petrographic analysis were collected every 3 m through the lower 60 m of the sequence and sporadically thereafter to the top of the section at 290 m. Numerous gaps in the sequence caused
Figure 5. Size range (in cm) of small, medium, and large pebbles from three localities (n = 45 at each) at Indian Garden Bay, Heceta Island (Fig. 1, site 4).
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Figure 6. Photomicrographs of conglomerate clasts in cross-polarized light. (A) Volcanic clast shows embayed plagioclase with corona rim in groundmass of fine plagioclase microlites; ID5W-30C-96 (Fig. 2, site 4). (B) Vesicular basalt with plagioclase phenocrysts in fine-grained groundmass; ID5W28C-96 (Fig. 2, site 4). (C) Volcaniclastic sandstone comprising small limestone (L) and volcanic rock fragments in calcite cement; SP-43-93 (Fig. 2, site 1). (D) Plutonic clast with plagioclase and mafic minerals forming a coarsegrained intergranular texture; MP111C-96 (Fig. 2, site 2). Scale bars are 1 mm.
by intrusive dikes and erosional gullies 10 m wide (perhaps associated with faults that duplicated stratigraphy) preclude an accurate assessment of true thickness. Thin sections of 79 clasts that were collected randomly at two localities (Fig. 2, site 6) provide compositional evidence. Particle size was assessed for 125 clasts measured on site through a 5 m interval where they are concentrated in five successive horizons (Fig. 9). As shown in Figures 7 and 8, the measured section exposed at site 6 (Fig. 2) along Tuxekan Passage comprises massive beds typical of the other conglomeratic units that crop out near Prince of Wales Island. Sub- to well-rounded clasts are poorly to moderately sorted, densely concentrated, and polymictic, composed at this locality of 50% limestone fragments (Figs. 8A–8C). Beds typically are crudely laminated, and clasts are tightly packed in grain-supported units that exhibit no imbrication (Figs. 8B–8C). Bedded-to-massive limestones that underlie the conglomerates on eastern Tuxekan Island are characterized by >50 m of pink, finely laminated, intertidal stromatolites and stromatolite breccias interbedded with peloidal-skeletal grainstones (Fig. 10A) and packstones containing abundant Pycinodesma bivalves. At 56 m above the base of the section (Fig. 7), the basal contact between the conglomerates and underlying limestones is channelized, characterized in the lower 1 m by repetitions of poorly sorted mixtures of angular to rounded limestone clasts that exhibit normal grading and become finer upsection (Figs. 7 and 8). At 70 m above the base, pebbles and cobbles are less tightly packed and supported in a sandy matrix; meter-thick sandstone intervals
also become more prevalent, forming a monotonous sequence of massive conglomerate and sandstone beds that, in general, lack evidence of cross-bedding, normal or reverse grading, imbrication, and other sedimentary structures. At 118 m above the base (Fig. 7), sandstone units without clasts or with clasts no larger than 6 cm (maximum dimension) become increasingly more prevalent. At 193 m, pebbly sandstones predominate, characterized by subrounded limestone clasts (Figs. 7 and 10E). At 245 m, the pebbly sandstones grade upward into sandstones that generally are pebble- and cobble-free (Fig. 7). Massive sandstones persist to the top of the section, characterized by small scours and cross-beds with rare conglomeratic lenses (Figs. 7 and 8D). Farther north along the shore, discontinuous sandstone outcrops grade upward into siltstones, mudstones, and lagoonal limestones that consist of abundant corals associated with massive and amphiporid stromatoporoids (Fig. 10F). PROVENANCE With the exception of the conglomerates exposed along Tuxekan Passage, most (72% average) of the clasts are volcanic (Fig. 2), as indicated by porphyritic textures, glomerocrysts, embayed quartz, abundant plagioclase laths, other altered volcanic fragments, and rare evidence of vesicular textures. Many of the volcanic clasts are silicified or highly altered; abundant plagioclase and rarer phenocrysts of K-spar, pyroxene, olivine, and quartz are replaced by calcite, chlorite, hematite, chert, or
Silurian polymictic conglomerates in Alaska’s Alexander terrane
Figure 7. Generalized stratigraphic section showing 290 m of conglomerates exposed on eastern Tuxekan Island (Fig. 2, site 6).
pyrite (Figs. 6 and 10D). Most of the phenocrysts not replaced by calcite are plagioclase associated with ~10% quartz and accessory minerals, including pyroxene and amphibole (Fig. 6). A typical example shows a plagioclase phenocryst with a corona rim surrounded by a groundmass of plagioclase microlites (Fig. 6A). The high percentage of plagioclase and the small proportion of feldspathoid, quartz, and accessory minerals indicate that these clasts are primarily volcanics of intermediate composition, comprising 20% felsic, 65% intermediate, and 15% basaltic fragments. Volcaniclastic clasts (15% average) comprise poorly sorted, angular to rounded volcanic rock fragments and rare limestone pebbles in a sandy matrix of feldspar, quartz, and other altered grains (Figs. 2 and 6C). Plutonic clasts are rare (6% average) and typically altered but appear to be granitic to gabbroic in composition with equigranular textures of plagioclase, rare quartz, chlorite-replaced biotite, and altered mafic minerals (Figs. 2 and 6D).
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In contrast, 50% of the clasts in the lower 5 m of the conglomeratic section exposed on eastern Tuxekan Island are limestone (Fig. 2, site 6). These clasts are identical to the massive carbonates that underlie the conglomerates—either stromatolites or packstones-grainstones consisting of finely laminated peloidal and skeletal grains, including micritized fragments of gastropods, coral, and brachiopods (Figs. 6C, 10A–10C, and 10E). Associated igneous clasts are highly altered but comprise remnant flow textures, phenocrysts (some embayed) of plagioclase, as well as unidentifiable glomerocrysts and equigranular microlites embedded in a silicified, fine-grained groundmass indicative of intermediate-mafic volcanic facies (Fig. 10D). Less than 15% of these clasts are pebbly volcaniclastic sandstones comprising approximately equal proportions of skeletal limestone and volcanic fragments. These compositional data indicate that the clasts had two primary sources. The lithology of the limestone clasts is identical to interbedded shallow-marine lagoonal and stromatolite reef facies of the Heceta Formation, primarily those that immediately underlie the conglomerates. These compare well with an additional 50 limestone clasts examined during an earlier study of strata exposed at site 1 (Figs. 2 and 6C). At both sites, limestone clasts are characterized primarily by laminated, graded skeletalpeloidal packstones and grainstones; crinoidal-peloidal wackestones and packstones; and stromatolite or Renalcis boundstones associated with skeletal fragments of corals, gastropods, brachiopods, ostracodes, and an abundance of small, micritized grains. These shallow-water carbonate clasts suggest that active faulting and erosion during progradation of coarse-grained siliciclastics produced localized accumulations of synsedimentary (Heceta) material at the base of some of the conglomerate sequences. Depositional conditions are considered in the next section. The non-limestone clasts, in contrast, probably were eroded from the underlying Descon Formation, which forms the oldest known basement exposed in the northwest area of Prince of Wales Island and represents the early Paleozoic development of a volcanic arc (Eberlein et al., 1983; Gehrels and Saleeby, 1987a). During or subsequent to Heceta deposition, many of the phenocrysts, glomerocrysts, and granular microlites, as noted above, were replaced by calcite and other minerals. Replacement of phenocrysts also took place earlier; Eberlein and Churkin (1970) report that plagioclase and olivine typically show alteration to other minerals in volcanic sections of the Descon Formation exposed nearby. Despite diagenetic alterations, the range of basaltic volcanic facies associated with Descon deposits— volcanic flows, pillow basalts, tuff breccias, and volcanic wackes (Eberlein and Churkin, 1970)—closely matches >90% of the clasts we examined. In addition, porphyritic vesicular basalts, which typify some Descon lavas (Eberlein and Churkin, 1970), are represented by <1% of clasts we investigated (Fig. 6B). A Descon source is also supported by detrital zircons extracted from sandy facies interbedded with the conglomerates. From site 1 (Fig. 2), four grains were dated as 460–465 Ma; at site 6 (Fig. 2), four grains ranged in age from 439 to 465 Ma (G.E. Gehrels,
Figure 8. Conglomeratic section exposed along Tuxekan Passage (Fig. 2, site 6). Locations denote where each sample was collected in meters above the base of the section. (A) Massive conglomerates eroded into pillar comprise abundant limestone clasts; bracketed area is enlarged in B, and arrow points to the same clast shown in B; loc. 56–59. (B) Poorly sorted mixture of limestone and other cobbles in pebbly matrix; loc. 57.5 (note arrow points to the same clast shown in A). (C) Close-up showing well-laminated limestone clasts; loc. 59.5. (D) Conglomerates fine upward into cross-bedded, pebbly sandstones; loc. 268.5. Ruler in B and C is 15 cm; hammer in D is 35 cm long.
Figure 9. Size range (in cm) of small to large pebbles (n = 125) measured in five successive horizons located 56–60 m above the section base, Tuxekan Island (Fig. 2, site 6).
Silurian polymictic conglomerates in Alaska’s Alexander terrane 1997, personal commun.). This provides additional support for an intra-arc, Middle–Upper Ordovician (Descon) source for the conglomeratic debris that interrupted carbonate deposition during the Late Silurian. Contrary to previous studies, Krutikov (1996) notes that none of the Silurian conglomerates exposed near northern Prince of Wales Island is composed primarily of plutonic rock fragments (Gehrels et al., 1996; Butler et al., 1997). Gehrels et al. (1996) describe clasts of diorite, quartz diorite, granite, and syenite in
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Heceta conglomerates on southern Prince of Wales Island. Southward thickening of siliciclastic deposits correlated with thinning and disappearance of Heceta limestones on southern Prince of Wales Island suggest that a significant influx of detrital material had prograded from an uplifted source in the south-southeast and overwhelmed the carbonate platform(s) in the north-northwest (Gehrels and Saleeby, 1987a; Soja, 1993). This study shows, in contrast, that cobbles in the northern outcrop area of the Heceta Formation are primarily andesitic-basaltic or intermediate volcanic
Figure 10. Photomicrographs of samples collected from Tuxekan Passage (Fig. 2, site 6). (A) Peloidal-skeletal grainstones with coated grains and gastropods represent bedded limestones underlying conglomerates; YCS-54.5-96. (B, C) Subangular-subrounded volcanic and limestone clasts in conglomerate’s pebbly matrix; YCS-58-96 and YCS-59-96, respectively. (D) Typical volcanic clast comprising altered glomerocysts in fine-grained groundmass; YCS-40CL-96. (E) Conglomerates fine upward into pebbly sandstones consisting of small subrounded to rounded volcanic and limestone pebbles; YCS-193-93. (F) Sandstones grade upward into silty limestones with brachiopods and coated corals; YCS-290-96. Sample numbers for A–C and E–F indicate collection location in meters above base of section. Scale bars are 2 mm; arrows in E and F indicate way up toward stratigraphic top.
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rocks and also limestones, with only rare examples of plutonic rocks. Thus clasts in the polymictic conglomerates distributed throughout the Prince of Wales Island area were eroded from locally uplifted portions of the volcanic arc–shallow marine platform complex (Krutikov, 1996). During onset of the Klakas orogeny, progressive unroofing of the arc’s intrusive roots must have produced erosive debris derived from Silurian plutons uplifted on southern Prince of Wales Island. This material was deposited locally, however, and was not transported as far north (>150 km) as the sites investigated on central and northern Prince of Wales Island. DEPOSITIONAL SETTING Most of the clasts at the six localities are pebble to cobble size (<15–30 cm in maximum dimension), similar to those measured at Indian Garden Bay and Tuxekan Passage (Figs. 5 and 9). All of the conglomerates sampled form massive, sheet-like units that are disorganized to crudely stratified. They comprise clasts that are subangular to well rounded, have low sphericity (prolate to oblate in shape), are poorly to moderately sorted, are embedded in a pebbly to sandy matrix, are grain or matrix supported, show rare imbrication, and rarely exhibit normal or reverse grading (Krutikov, 1996). In addition, their tight packing, similar composition (predominantly volcanic), and local derivation suggest that the clasts were transported from uplifted sources nearby with a fast rate of sediment supply (Nilsen, 1982; Rust and Koster, 1984). These attributes are generally produced by sediment-charged, gravel-dominated debris flows in alluvial fan complexes (Muir and Rust, 1982; Nilsen, 1982; Rust and Koster, 1984; Kim et al., 1995; Miall, 1996; Stratford and Aitchison, 1996; Major, 1997). Typically these gravity-driven mixtures of highly concentrated sediment and water are deposited rapidly close to the source area, producing homogeneous, massive alluvial “fanglomerates” in tectonically active areas (Miall, 1981; Nilsen, 1982; Rust and Koster, 1984; Major, 1997). This helps explain why plutonic clasts in the Heceta conglomerates are restricted to areas adjacent to the uplifted plutons on southern Prince of Wales Island. An alluvial fan forms closest to the sediment source, proximal to the uplifted rocks where large, angular, poorly sorted rock fragments undergo little reworking during rapid denudation. The general lack of internal stratification and of large- or smallscale sedimentary structures in the conglomerates under study suggests high-energy, rapid deposition as debris flows that initially lacked extensive development of braided rivers and plains (Nilsen, 1982; Rust and Koster, 1984). Angular to subangular, poorly sorted clasts represent proximal, first-cycle debris. The roundness and sphericity of many other clasts suggest reworking in repeat transport and storage cycles from the underlying Descon volcanic breccias and conglomerates (Nilsen, 1982; Jo et al., 1997). These had probably accumulated as volcanic and volcaniclastic debris adjacent to an emergent arc edifice before volcanism ceased. Recycling of Descon material on land and (or)
in shallow-marine sites could have proceeded once widespread limestones began to form an incipient carbonate platform in the mid-Early Silurian (Gehrels et al., 1996). Localized faulting and (or) basal scouring during progradation of coastal alluvial fans over adjacent tidal flats could account for the incorporation of shallow-marine Heceta limestone clasts into some of the conglomeratic units, such as those exposed along Tuxekan Passage (Fig. 2, site 6). At that site, clasts range in size from 1 to 25 cm with an average maximum dimension of 8.9 cm (Fig. 9). Most are subrounded to rounded, subspherical, and poorly to moderately well sorted. This suggests periodic reworking in coastal sites followed by rapid deposition of coarse material and winnowing of finer-grained sediments. Similar processes operated in Paleozoic coastal alluvial fan complexes in South Australia and Arctic Canada, where gravels produced as a result of fault-induced coastal uplift prograded over shallow-marine facies (Daily et al., 1980; Rust, 1981; Muir and Rust, 1982). Island arc collision zones favor coastal alluvial fan development adjacent to relatively narrow and steep marine platforms (Rust and Koster, 1984). Cyclic repetition of the conglomerate-to-mudstone sequences suggests periodic tectonism as the arc approached collision with an unknown entity (Gehrels et al., 1983, 1996; Gehrels and Saleeby, 1987a). Rejuvenation of alluvial fans caused repeated incursions of coarsegrained sediment into small, isolated basins (Miall, 1981; Rust and Koster, 1984; Jo et al., 1997). Reduced sediment supply and (or) a decrease in flow energy allowed sand to settle as matrix in between clasts and to produce fining-upward beds without clasts (Rust and Koster, 1984). As rejuvenative effects diminished during waning tectonic intervals, more thinly bedded sand and silt horizons became predominant higher in each stratigraphic section. Pebble horizons became less common, exhibiting rare imbrication of cobbles (Fig. 4D), scour channels (Fig. 4C), as well as variations in bed thickness and rapid vertical facies changes. Together these features represent aqueous activity in fluvial channels and bars that developed on distal portions of a coastal alluvial fan complex (Turner, 1980; Nilsen, 1982; Rust and Koster, 1984; Miall, 1996). Finer-grained sediments accumulated during cyclic alluvial infilling as source regions were denuded and stream competence decreased rapidly down-fan (Turner, 1980; Rust and Koster, 1984; Jo et al., 1997). Marine transgression led eventually to a resumption of carbonate deposition across a spectrum of shallow- to deepmarine environments. This is indicated by the consistent upward transition from the conglomerates into finer grained siliciclastic and marine limestone facies, similar to fan-delta progradation over shallow-marine carbonates described from the Devonian of Arctic Canada (Muir and Rust, 1982) and Pennsylvanian of Texas (Dutton, 1982). The lagoonal deposits, characterized by low-diversity suites of marine fossils, grade upward into calciturbidites and limestone slumps and subaqueous debris flows deposited along a deep-marine slope (Soja,
Silurian polymictic conglomerates in Alaska’s Alexander terrane 1990, 1993). Upward rapid gradation into conglomerates and red beds of the Karheen Formation records the transition from flysch to molasse deposition during culminating phases of the Klakas orogeny. TECTONIC IMPLICATIONS In parts of Alaska during the Late Silurian, coarse-grained, siliciclastic units became predominant. These voluminous deposits appear to correlate with “Old Red Sandstone” facies, which reflect late Caledonide events and early stages in the Ellesmere orogeny (Turner, 1980; Ziegler, 1988). In the Arctic Alaska terrane of northern Alaska, for example, Silurian marine deposits grade upward into nonmarine rocks of Early or Middle Devonian age, similar to those in the Alexander terrane (Ovenshine et al., 1969; Ovenshine, 1975; Grantz et al., 1994; Moore et al., 1994). Parts of northern Alaska may have been linked with the Ellesmere fold belt of the Canadian Arctic and northern Greenland, as suggested by an extensive, sub-Mississippian unconformity in the subsurface of some subterranes of the Arctic Alaska terrane (Moore et al., 1994). Orogenesis may have occurred in northern parts of Alaska in the Late Silurian–Early Devonian. This is suggested by the strongly deformed Upper Silurian–Lower Devonian rocks of the Franklinian succession in northern Alaska (Arctic continental margin) and by the lack of deformation and (or) metamorphism in Lower?–Middle Devonian rocks in northern Alaska (northeastern Brooks Range and western North Slope) (Grantz et al., 1994; Moore et al., 1994). Cessation of platform carbonate deposition in much of northern Alaska (York Mountains, Brook Range, and Sadlerochit and Shublik Mountains) in the Early or Middle Devonian also suggests syn- or post– Early Devonian orogenesis (or clastic influx) (Dumoulin and Harris, 1994). The timing of this deformational activity has implications for interpreting the conglomerates under study within a larger tectonic context. Moore et al. (1994) noted that although the causes of a postulated pre–Middle Devonian orogenic event in northern and southeastern Alaska are unknown, it may have been related to convergence of North America with Siberia or to North America’s collision with an island arc. The Alexander arc-related terrane also experienced the Klakas orogeny during this time interval. Paleomagnetic data and fossils from Heceta limestones indicate that this event took place in a subequatorial site along the Uralian Seaway (Bazard et al., 1995; Soja and Antoshkina, 1997; Soja et al., 2000; Soja, this volume). Thus there is an intriguing possibility that the Alexander arc collided with cratonic portions of Laurentia in the Late Silurian–Early Devonian and became welded to the Laurussian landmass during late-stage Caledonide activity (Bazard et al., 1995; Butler et al., 1997; Soja and Antoshkina, 1997; Soja et al., 2000) (Fig. 1). Some similarities should also be noted with the Pearya composite terrane. Pearya is characterized by plutonism and
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arc-related volcanism in the Ordovician-Devonian, carbonatesiliciclastic successions of Late Ordovician–Late Silurian age, as well as by accretion to the Franklinian mobile belt in the Late Silurian and subsequent accumulation of Upper Silurian (?) –Lower Devonian terrestrial and shallow-marine clastics (Trettin, 1987; G.E. Gehrels, 2004, personal commun.). These events coincide in time with pre- and syn-Klakas orogenic activity in southeastern Alaska. Dissimilarities include evidence that Pearya experienced Taconic-related activity in the mid-Ordovician and had closest affinities with Svalbard-Laurentia, not with Baltica-Siberia (Trettin, 1987). Deposition of voluminous amounts of Heceta conglomeratic material was associated with a cessation in arc-type magmatism during the earliest manifestations of compressional stresses that produced the prolonged Klakas orogenic event (Gehrels et al., 1983, 1996). Conodonts preserved in limestones interbedded with the conglomerates indicate that carbonate deposition was disrupted episodically by siliciclastic deposition during the early to mid- (?) Ludlow (Ovenshine and Webster, 1970; Soja, 1993). Above these conglomerates, Heceta facies on central-northern Prince of Wales Island are in conformable contact with the overlying Karheen Formation. They form mixed carbonate-siliciclastic successions that become siltier and more argillaceous as they grade upsection into the subaerial reds beds and shallow-marine lithic wackes of the Karheen Formation (Soja, 1990, 1993). Culminating phases in uplift, thrusting, deformation, and metamorphism associated with the Klakas orogeny are evident in Karheen deposits. These represent syn- and post-orogenic nearshore and continental clastics that accumulated as molasse following rapid erosion of adjacent uplifted areas (Ovenshine et al., 1969; Gehrels et al., 1996). Major inundation of and burial by 1500 m of siliciclastic debris reflects the migration of the Karheen clastic wedge from these uplifted areas during final infilling of marine basins in the late Early Silurian–Early Devonian (Gehrels et al., 1983; Soja, 1993; Soja, this volume). The Heceta conglomerates described here and the red beds of the Karheen Formation appear to be very similar to Old Red Sandstone-related facies that accumulated along the rising Caledonide front in many parts of Laurussia in the Late Silurian–Early Devonian. They share greatest similarities with kilometer-thick deposits that formed as cyclic fanglomerates and fluvial sediments in coastal alluvial fans and intermontane sites affected by ongoing tectonism in late-orogenic, rapidly subsiding basins (Turner, 1980; Gloppen and Steel, 1981; Ziegler, 1988; Friend and Williams, 2000). In the Early Devonian, the Alexander terrane was receiving sediment from an unknown continental source, as indicated by Precambrian zircons extracted from the Karheen Formation that predate the oldest rocks exposed in the Alexander terrane (Bazard et al., 1995; Gehrels et al., 1996; Butler et al., 1997). Ascertaining the provenance and depositional setting of Silurian polymictic conglomerates in the Heceta Formation thus provides baseline data for using detrital zircon and geochemical evidence in the future to test two interrelated hypotheses: the Klakas orogeny
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of southeastern Alaska was a Caledonide event and the Karheen Formation is a dismembered facies of the Old Red Sandstone preserved in a terrane accreted to western North America. ACKNOWLEDGMENTS We are grateful for financial support from the National Science Foundation (EAR-9417407) and from Colgate University’s Research Council and Division of Natural Sciences and Mathematics. We thank C. Brogenski, V. Hayssen, C. Preston, R. Quick, and B. White for field support and D. Keller, D. Linsley, W. Peck, and B. White for lab assistance. Improvements to earlier drafts of the manuscript were suggested by G. Gehrels, N. Lindsley-Griffin, and B. White. REFERENCES CITED Antoshkina, A.I., and Soja, C.M., 2006, Late Silurian reconstruction indicated by migration of reef biota between Alaska, the Urals, and Siberia (Salair): Geologiska Föreningens i Stockholm Förhandlingar, v. 128, p. 75–78. Bazard, D.R., Butler, R.F., Gehrels, G., and Soja, C.M., 1995, Early Devonian paleomagnetic data from the Lower Devonian Karheen Formation suggest Laurentia-Baltica connection for the Alexander terrane: Geology, v. 23, p. 707–710, doi: 10.1130/0091-7613(1995)023<0707:EDPDFT> 2.3.CO;2. Butler, R.F., Gehrels, G.E., and Bazard, D.R., 1997, Paleomagnetism of Paleozoic strata of the Alexander terrane, southeastern Alaska: Geological Society of America Bulletin, v. 109, p. 1372–1388, doi: 10.1130/00167606(1997)109<1372:POPSOT>2.3.CO;2. Coney, P.J., Jones, D.L., and Monger, J.W.H., 1980, Cordilleran suspect terranes: Nature, v. 288, p. 329–333, doi: 10.1038/288329a0. Daily, B., Moore, P.S., and Rust, B.R., 1980, Terrestrial-marine transition in the Cambrian rocks of Kangaroo Island, South Australia: Sedimentology, v. 27, p. 379–399, doi: 10.1111/j.1365-3091.1980.tb01189.x. Dumoulin, J.A., and Harris, A.G., 1994, Depositional framework and regional correlation of pre-Carboniferous metacarbonate rocks of the Snowden Mountain area, central Brooks Range, northern Alaska: U.S. Geological Survey Professional Paper 1545, 66 p. Dutton, S.P., 1982, Pennsylvanian fan-delta and carbonate deposition, Mobeetie Field, Texas panhandle: American Association of Petroleum Geologists Bulletin, v. 66, p. 389–407. Eberlein, G.D., and Churkin, M., Jr., 1970, Paleozoic stratigraphy in the northwest coastal area of Prince of Wales Island, southeastern Alaska: U.S. Geological Survey Bulletin 1284, 67 p. Eberlein, G.D., Churkin, M., Jr., Carter, C., Berg, H.C., and Ovenshine, A.T., 1983, Geology of the Craig quadrangle, Alaska: U.S. Geological Survey Open-File Report 83-91, 28 p. Follo, M.F., 1992, Conglomerates as clues to the sedimentary and tectonic evolution of a suspect terrane: Wallowa Mountains, Oregon: Geological Society of America Bulletin, v. 104, p. 1561–1576, doi: 10.1130/0016-7606 (1992)104<1561:CACTTS>2.3.CO;2. Friend, P.F., and Williams, B.P.J., eds., 2000, New perspectives on the Old Red Sandstone: London, Geological Society Special Publication 180, 623 p. Gehrels, G., and Saleeby, J., 1987a, Geologic framework, tectonic evolution, and displacement history of the Alexander terrane: Tectonics, v. 6, p. 151–173. Gehrels, G., and Saleeby, J., 1987b, Geology of southern Prince of Wales Island, southeastern Alaska: Geological Society of America Bulletin, v. 98, p. 123–137, doi: 10.1130/0016-7606(1987)98<123:GOSPOW>2.0.CO;2. Gehrels, G.E., Saleeby, J.B., and Berg, H.C., 1983, Preliminary description of the Klakas orogeny in the southern Alexander terrane, southeastern Alaska, in Stevens, C.H., ed., Pre-Jurassic rocks in western North American suspect terranes: Los Angeles, Pacific Section, SEPM, p. 131–141. Gehrels, G.E., Butler, R.F., and Bazard, D.R., 1996, Detrital zircon geochronology of the Alexander terrane, Alaska: Geological Society of America Bulletin, v. 108, p. 722–734, doi: 10.1130/0016-7606(1996)108<0722: DZGOTA>2.3.CO;2.
Gloppen, T.G., and Steel, R.J., 1981, The deposits, internal structure and geometry in six alluvial fan-fan delta bodies (Devonian-Norway)—A study in the significance of bedding sequence in conglomerates, in Ethridge, F.G., and Flores, R.M., eds., Recent and ancient nonmarine depositional environments: Models for exploration: Tulsa, Oklahoma, SEPM Special Publication 31, p. 49–69. Grantz, A., May, S.D., and Hart, P.E., 1994, Geology of the Arctic continental margin, in Plafker, G., and Berg, H.C., eds., The geology of Alaska: Boulder, Colorado, Geological Society of America, Geology of North America, v. G-1, p. 17–48. Jo, H.R., Rhee, C.W., and Chough, S.K., 1997, Distinctive characteristics of a streamflow-dominated alluvial fan deposit: Sanghori area, Kyongsang Basin (Early Cretaceous), southeastern Korea: Sedimentary Geology, v. 110, p. 51–79, doi: 10.1016/S0037-0738(96)00083-8. Kim, S.B., Chough, S.K., and Chun, S.S., 1995, Bouldery deposits in the lowermost part of the Cretaceous Kyokpori Formation, SW Korea: Cohesionless debris flows and debris falls on a steep-gradient delta slope: Sedimentary Geology, v. 98, p. 97–119, doi: 10.1016/0037-0738(95)00029-8. Krutikov, L., 1996, Origin and provenance of polymictic conglomerates in the Heceta Formation, Alexander terrane, southeastern Alaska [unpublished senior thesis]: Hamilton, New York, Colgate University, 36 p. Major, J.J., 1997, Depositional processes in large-scale debris-flow experiments: The Journal of Geology, v. 105, p. 345–366. Miall, A.D., 1981, Alluvial sedimentary basins: Tectonic setting and basin architecture, in Miall, A.D., ed., Sedimentation and tectonics in alluvial basins: Waterloo, Ontario, Geological Association of Canada Paper Special 23, p. 1–33. Miall, A.D., 1996, The geology of fluvial deposits: New York, Springer-Verlag, 582 p. Moore, T.E., Wallace, W.K., Bird, K.J., Karl, S.M., Mull, C.G., and Dillon, J.T., 1994, Geology of northern Alaska, in Plafker, G., and Berg, H.C., eds., The geology of Alaska: Boulder, Colorado, Geological Society of America, Geology of North America, v. G-1, p. 49–140. Muir, I.D., and Rust, B.R., 1982, Sedimentology of a Lower Devonian coastal alluvial fan complex: The Snowblind Bay Formation of Cornwallis Island, Northwest Territories, Canada: Bulletin of Canadian Petroleum Geology, v. 30, p. 245–263. Nilsen, T.H., 1982, Alluvial fan deposits, in Scholle, P.A., and Spearing, D., eds., Sandstone depositional environments: Tulsa, Oklahoma, American Association of Petroleum Geologists, p. 49–86. Ovenshine, A.T., 1975, Tidal origin of part of the Karheen Formation (Lower Devonian), southeastern Alaska, in Ginsburg, R.N., ed., Tidal deposits: A case book of recent fossil counterparts: New York, Springer-Verlag, p. 141–148. Ovenshine, A.T., and Webster, G.D., 1970, Age and stratigraphy of the Heceta Limestone in northern Sea Otter Sound, southeastern Alaska: U.S. Geological Survey Professional Paper 700-C, p. 170–174. Ovenshine, A.T., Eberlein, G.D., and Churkin, M., Jr., 1969, Paleotectonic significance of a Silurian–Devonian clastic wedge, southeastern Alaska: Geological Society of America Abstracts with Programs, v. 1, p. 50. Rigby, J.K., Nitecki, M., Soja, C.M., and Blodgett, R.B., 1994, Silurian aphrosalpingid sphinctozoans from Alaska and Russia: Acta Palaeontologica Polonica, v. 39, p. 341–391. Rust, B.R., 1981, Alluvial deposits and tectonic style: Devonian and Carboniferous successions in eastern Gaspé, in Miall, A.D., ed., Sedimentation and tectonics in alluvial basins: Waterloo, Ontario, Geological Association of Canada Special Paper 23, p. 49–76. Rust, B.R., and Koster, E.H., 1984, Coarse alluvial deposits, in Walker, R.G., ed., Facies models (second edition).: Toronto, Geoscience Canada, reprint series 1, p. 53–69. Soja, C.M., 1990, Island arc carbonates from the Silurian Heceta Formation of southeastern Alaska (Alexander terrane): Journal of Sedimentary Petrology, v. 60, p. 235–249. Soja, C.M., 1991, Origin of Silurian reefs in the Alexander terrane of southeastern Alaska: Palaios, v. 6, p. 111–126, doi: 10.2307/3514877. Soja, C.M., 1993, Carbonate platform development in a Silurian oceanic island: A case study from Alaska’s Alexander terrane: Journal of Sedimentary Petrology, v. 63, p. 1078–1088. Soja, C.M., 1994, Significance of Silurian stromatolite-sphinctozoan reefs: Geology, v. 22, p. 355–358, doi: 10.1130/0091-7613(1994)022<0355: SOSSSR>2.3.CO;2. Soja, C.M., and Antoshkina, A.I., 1997, Coeval development of Silurian stro-
Silurian polymictic conglomerates in Alaska’s Alexander terrane matolite reefs in Alaska and the Ural Mountains: Implications for paleogeography of the Alexander terrane: Geology, v. 25, p. 539–542, doi: 10. 1130/0091-7613(1997)025<0539:CDOSSR>2.3.CO;2. Soja, C.M., White, B., Antoshkina, A.I., Joyce, S., Mayhew, L., Flynn, B., and Gleason, A., 2000, Development and decline of a Silurian stromatolite reef complex, Glacier Bay National Park, Alaska: Palaios, v. 15, p. 273–292. Soja, C.M., Mitchell, M., Newton, A.J., Vendetti, J., Visaggi, C., Antoshkina, A.I., and White, B., 2003, Paleoecology of sponge-?hydroid associations in Silurian microbial reefs: Palaios, v. 18, p. 225–235, doi: 10.1669/08831351(2003)018<0225:POSHAI>2.0.CO;2. Stratford, J.M.C., and Aitchison, J.C., 1996, Devonian intra-oceanic arc rift sedimentation—Facies development in the Gamilaroi terrane, New Eng-
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land orogen, eastern Australia: Sedimentary Geology, v. 101, p. 173–192, doi: 10.1016/0037-0738(95)00063-1. Trettin, H.P., 1987, Pearya: A composite terrane with Caledonian affinities in northern Ellesmere Island: Canadian Journal of Earth Sciences, v. 24, p. 224–245. Turner, P., 1980, Continental red beds: Developments in Sedimentology, v. 29, 562 p. Ziegler, P.A., 1988, Laurussia—The old red continent, in McMillan, N.J., Embry, A.F., and Glass, D.J., eds., Devonian of the world: Canadian Society of Petroleum Geologists Memoir 14, v. 1, p. 15–48. MANUSCRIPT ACCEPTED BY THE SOCIETY 14 DECEMBER 2007
Printed in the USA
The Geological Society of America Special Paper 442 2008
Devonian brachiopods of southwesternmost Laurentia: Biogeographic affinities and tectonic significance Arthur J. Boucot Department of Zoology, Oregon State University, Corvallis, Oregon 97331-2914, USA Forrest G. Poole U.S. Geological Survey, Box 25046, MS-973, Federal Center, Denver, Colorado 80225-0046, USA Ricardo Amaya-Martínez Departamento de Geología, Universidad de Sonora, Hermosillo, Sonora 83000, México Anita G. Harris 1523 East Hillsboro Boulevard., Apartment 1031, Deerfield Beach, Florida 33441-4307, USA Charles A. Sandberg U.S. Geological Survey, Box 25046, MS-939, Federal Center, Denver, Colorado 80225-0046, USA William R. Page U.S. Geological Survey, Box 25046, MS-980, Federal Center, Denver, Colorado 80225-0046, USA ABSTRACT Three brachiopod faunas discussed herein record different depositional and tectonic settings along the southwestern margin of Laurentia (North America) during Devonian time. Depositional settings include inner continental shelf (Cerros de Los Murciélagos), medial continental shelf (Rancho Placeritos), and offshelf continental rise (Rancho Los Chinos). Ages of Devonian brachiopod faunas include middle Early (Pragian) at Rancho Placeritos in west-central Sonora, late Middle (Givetian) at Cerros de Los Murciélagos in northwestern Sonora, and late Late (Famennian) at Rancho Los Chinos in central Sonora. The brachiopods of these three faunas, as well as the gastropod Orecopia, are easily recognized in outcrop and thus are useful for local and regional correlations. Pragian brachiopods dominated by Acrospirifer and Meristella in the “San Miguel Formation” at Rancho Placeritos represent the widespread Appohimchi Subprovince of eastern and southern Laurentia. Conodonts of the early to middle Pragian sulcatus to kindlei Zones associated with the brachiopods confirm the ages indicated by the brachiopod fauna and provide additional information on the depositional setting of the Devonian strata. Biostratigraphic distribution of the Appohimchi brachiopod fauna indicates continuous Early Devonian shelf deposition along the entire southern margin of Laurentia. The largely emergent southwest-trending Transcontinental arch apparently formed a Boucot, A.J., Poole, F.G., Amaya-Martínez, R., Harris, A.G., Sandberg, C.A., and Page, W.R., 2008, Devonian brachiopods of southwesternmost Laurentia: Biogeographic affinities and tectonic significance, in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 77–97 , doi: 10.1130/2008.442(05). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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Boucot et al. barrier preventing migration and mixing of many genera and species of brachiopods from the southern shelf of Laurentia in northern Mexico to the western shelf (Cordilleran miogeocline) in the western United States. Middle Devonian Stringocephalus brachiopods and Late Devonian Orecopia gastropods in the “Los Murciélagos Formation” in northwest Sonora represent the southwesternmost occurrence of these genera in North America and date the host rocks as Givetian and Frasnian, respectively. Rhynchonelloid brachiopods (Dzieduszyckia sonora) and associated worm tubes in the Los Pozos Formation of the Sonora allochthon in central Sonora are also found in stratiform-barite facies in the upper Upper Devonian (Famennian) part of the Slaven Chert in the Roberts Mountains allochthon (upper plate) of central and western Nevada. Although these brachiopods and worm tubes occur in similar depositional settings along the margin of Laurentia in Mexico, they occur in allochthons that exhibit different tectonic styles and times of emplacement. Thus, the allochthons containing the brachiopods and worm tubes in Sonora and Nevada are parts of separate orogenic belts and have different geographic settings and tectonic histories. Devonian facies belts and faunas in northern Mexico indicate a continuous continental shelf along the entire southern margin of Laurentia. These data, in addition to the continuity of the late Paleozoic Ouachita-Marathon-Sonora orogen across northern Mexico, contradict the early Late Jurassic Mojave-Sonora megashear as a viable hypothesis for large-magnitude offset (600–1100 km) of Proterozoic through Middle Jurassic rocks from California to Sonora. Keywords: Appohimchi Subprovince, Devonian brachiopods and conodonts, Sonora, biogeography.
INTRODUCTION Devonian brachiopods from the southwestern margin of Laurentia (North America) in the State of Sonora, Mexico, are found in four principal depositional settings—cratonic platform, inner and medial continental shelf, and continental rise (Fig. 1). They occur in strata of Early, Middle, and Late Devonian ages (Fig. 2). Of the three brachiopod faunas discussed here, two occupied environments on the continental shelf (Figs. 3 and 4) and the third occupied a unique environment within the continental rise (Fig. 5). The brachiopods at Rancho Placeritos (loc. 1 on Fig. 1) in west-central Sonora and near Rancho Los Chinos (loc. 3 on Fig. 1) in central Sonora were discovered by Poole and coworkers during geologic mapping and stratigraphic studies since 1996. The brachiopods at Cerros de Los Murciélagos (loc. 2 on Fig. 1) in northwestern Sonora were recognized by Boucot in 1974, and their stratigraphic position was established by Poole and AmayaMartínez during geologic mapping in 2000. In the Rancho Placeritos area, middle Early Devonian (Pragian) brachiopods are present in a medial carbonate-shelf setting in the “San Miguel Formation.” At Cerros de Los Murciélagos, late Middle Devonian (Givetian) brachiopods indicate an inner carbonate-shelf depositional environment. About 1 km westsouthwest of Rancho Los Chinos in the Minas de Barita area in central Sonora, late Late Devonian (Famennian) brachiopods and worm tubes in the Los Pozos Formation are present near seafloor
hydrothermal-vent (and possibly methane cold-seep) deposits that probably accumulated in an offshelf continental-rise setting. The Los Pozos Formation is part of the Sonora allochthon of the late Paleozoic Ouachita-Marathon-Sonora orogen along the southern margin of Paleozoic Laurentia (Poole et al., 2005). The Early Devonian brachiopod fauna in west-central Sonora was noted by Poole et al. (2003); the Middle Devonian brachiopod fauna in northwestern Sonora was discovered by Boucot during fieldwork in 1974; and the Late Devonian brachiopod fauna in central Sonora was noted by Noll (1981) and described by Noll et al. (1984). Subsequently, collections were made by Poole from several new localities in central Sonora (Poole and Dutro, 1988; Poole et al., 1991). These three brachiopod faunas provide information about Devonian depositional settings along the southwestern margin of Laurentia. Conodonts occur within some brachiopod-bearing beds and associated strata, and provide supplemental information on age and depositional setting. Stratigraphic sections were measured at Rancho Placeritos and Cerros de Los Murciélagos (Figs. 3 and 4), and near Rancho Los Chinos (Fig. 5). Harris and Sandberg studied conodonts from the Rancho Placeritos and Cerros de Los Murciélagos sections. Fossils are scarce in the section near Rancho Los Chinos, but elsewhere in the area radiolarians and conodonts in the Los Pozos Formation confirm a Famennian age for the brachiopodbearing stratiform barite beds. Investigations of the Pragian brachiopod fauna provide significant information concerning biogeographic affinities and
Devonian brachiopods of southwesternmost Laurentia
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Figure 1. Index map of Sonora showing locations of the three brachiopod collections discussed in this report. Locality 1—“San Miguel Formation”; locality 2—“Los Murciélagos Formation”; locality 3—Los Pozos Formation. Map modified from Poole et al. (2005); note that Baja California and Alta California have been restored ~400 km to their pre-Tertiary position; heavy lines offsetting inferred southern edge of continental Laurentian crust are hypothesized Neoproterozoic transform faults related to the breakup of Rodinia. The Cerro El Tejano cratonic-platform section is located in northeastern Sonora (see González-León, 1986, for details on the fauna and stratigraphy).
depositional and tectonic settings on the southwest margin of Laurentia (Figs. 1, 3, and 6). Quotation marks identify formation names at Cerros de Los Murciélagos and Rancho Placeritos that have not been documented in a formal publication. These names will be formalized in reports and maps in preparation. Formal names used for strata in the Minas de Barita region (Rancho Los Chinos area) were defined in previous reports (see Poole et al., 2005). DEVONIAN BRACHIOPOD FAUNAS STUDIED IN SONORA Early Devonian (Pragian) Brachiopods from Rancho Placeritos The middle Early Devonian (Pragian) silicified-brachiopod and conodont fauna occurs within a disconformity-bounded,
1.5-m-thick limestone unit, 105 m above the base of the Devonian “San Miguel Formation” (>240 m thick) near Rancho Placeritos, 40 km northwest of Hermosillo, Sonora (loc. 1 on Fig. 1, Figs. 2 and 3). During fieldwork in 2001, Poole, Amaya-Martínez, and three students collected 90 kg of limestone (sample 01FP-300F) containing silicified brachiopod specimens. Rensselaeria as well as conodonts indicating an early or middle Pragian age are present in this bed. Discomyorthis is unknown below the Pragian, which eliminates the possibility for a pre-Pragian age. Plicoplasia cooperi is unknown above the Pragian, but this argument is weakened by the possibility that the Sonora material is a species close, but not identical, to P. cooperi. The brachiopod fauna is described in the section on Systematic Paleontology. A stratigraphic column of the lower two-thirds of the “San Miguel Formation” (Fig. 3) shows the position of the Pragian brachiopod and conodont collections. The Pragian brachiopod-bearing unit is disconformably underlain by limestone beds containing
AGE
EPOCH
EARLY CARB.
AGE ~Ma
361
STANDARD CONODONT ZONES sulcata Late praesulcata Middle Early Late expansa
Middle Early Late
postera
Early Late
FAMENNIAN
trachytera
Early Latest
LATE DEVONIAN
marginifera
(89FP–104F) Dzieduszyckia sonora (Rancho Los Chinos area)
Late Early Late
rhomboidea
Early Latest Late
crepida
Middle Early Late
FRASNIAN
376
triangularis Middle Early linguiformis Late rhenana Early semichat jamieae Late hassi Early punctata
? ?
(01FP–14F, -15F, and 06FP-24F) Inferred position of Orecopia sp. (Cerros de Los Murciélagos)
transitans
384
Late
falsiovalis
Early
GIVETIAN
×
11
Late
hermannicristatus
Early
varcus
Middle
Late
Early
(05FP–41F) Stringocephalus sp. (Cerros de Los Murciélagos). No conodonts found in sample
?
hemiansatus
388
ensensis kockelianus
EIFELIAN
MIDDLE DEVONIAN
disparilis
australis costatus
392
partitus
×
6 –10
Conodont collections from "San Miguel Formation" in Rancho Placeritos area (see text and Table 1)
5
(01FP–300F) Pragian brachiopods and conodonts (East section, Rancho Placeritos)
patulus
EMSIAN
inversus-laticostatus nothoperbonus gronbergi (excavatus)
PRAGIAN
410
413
SILURIAN
dehiscens (kitabicus) pireneae kindlei
LOCHKOVIAN
EARLY DEVONIAN
serotinus
sulcatus pesavis delta eurekaensis
418
woschmidti eosteinhornensis
×
×4 × 2–3 ×1
Figure 2. Devonian conodont biochronologic time scale (Sandberg and Ziegler, 1996) showing positions of the three brachiopod collections reported herein, and 11 conodont collections from the Rancho Placeritos section (dates [Ma] revised herein by Sandberg).
Devonian brachiopods of southwesternmost Laurentia
81
Lithologic Symbols limestone 98FP–139F
U ? — ? 11 M
32
Middle-Late Devonian conodonts and styliolinids
conglomeratic sandy silty
00FP–236F &-237F 99FP–216Fa & b 96FP–158F 01FP–300F 00FP–239F
7–10 6 M 5 L
8 1.5 6 1 7 2 5 1.5
early Eifelian conodonts in matrix and clasts of debris flow early Pragian brachiopods and conodonts (brachiopod fauna belongs to Appohimchi Subprovince of Eastern Americas Realm)
argillaceous dolomitic disconformity
Fossil Symbols brachiopods gastropods pelmatozoans styliolinids
8
fish burrows
9 98FP–238F
middle Lochkovian conodonts
4
conodonts
8
COVERED
6 11
97FP–332F
3
3
Numbers 1–8 on left side of column are age-diagnostic conodont samples Numbers on right side of column are unit thicknesses in meters early Lochkovian conodonts
11 97FP–331F
2
00FP–241F
1
Approximate DevonianSilurian boundary
early Lochkovian conodonts
27
early Lochkovian—latest Pridolian conodonts
20 METERS
0 "San Antonio Dolostone" (Silurian)
Figure 3. Generalized section of the lower two-thirds of the Devonian “San Miguel Formation” in the eastern part of the Rancho Placeritos area, west-central Sonora. Section measured in 1997 by Poole, Amaya-Martínez, and Page, and in 1998 by Poole and Amaya-Martínez; conodonts studied by Harris and Sandberg. The eastern measured section contains one debris-flow limestone unit, whereas the western measured section contains at least three separate debris-flow units.
Lochkovian shallow-water conodonts and disconformably overlain by a limestone debris-flow conglomerate containing early Eifelian deeper-water conodonts both in the matrix and in clasts (Figs. 2 and 3; Table 1). Some clasts in the conglomerate contain two-hole crinoid ossicles that support a latest Emsian to early Eifelian conodont age (Poole, unpublished data). Reworked detritus
is present directly above each of the erosional surfaces marking the disconformable contacts bounding the Pragian brachiopodbearing limestone unit. Fossils in the “San Miguel Formation” range in age from early Lochkovian to middle Famennian; the uppermost part of the formation is not exposed (Figs. 2 and 3; Tables 1 and 2).
82
Boucot et al. ALLUVIUM Cemented talus breccia
(01FP-14F and -15F)— early Late Devonian gastropod samples Orecopia interval
06FP–24F 4 Block III
Block II 00FP–33F 3 99FP–20F
Stringocephalid interval (05FP–41F)
2
Middle-Late Devonian conodont in sample OOFP-33F; Stringocephalus sp. at 99FP-20F
Lithologic Symbols talus breccia megabreccia limestone dolostone/dolomitic limestone sandy and silty carbonate/sandstone unconformity
Block I
Fossil Symbols brachiopods gastropods corals stromatoporoids/ algal stromatolites conodonts
00FP–35F
1
Middle-Late Devonian conodonts
30 METERS
FAULT Tertiary megabreccia, blocks (up to 3 m across) contain stromatoporoids, stringocephalids, etc. 0
Figure 4. Generalized composite section of part of the Devonian “Los Murciélagos Formation” at Cerros de Los Murciélagos, northwestern Sonora. Section measured in 2001 by Poole, Amaya-Martínez, Page, and Harris, and in 2003 by Poole and Amaya-Martínez; conodonts studied by Harris.
Middle Devonian (Givetian) Brachiopods from Cerros de Los Murciélagos Calcite-replaced stringocephalids of late Middle Devonian (Givetian) age occur in a 14-m interval in the upper part (Block II on Fig. 4) of the exposed “Los Murciélagos Formation” at Cerros de Los Murciélagos, 22 km west of Caborca, Sonora (loc. 2 on Fig. 1, Figs. 2 and 4). Brunner (1975) measured two stratigraphic
sections and reported numerous fossils in them, including foraminifers and conodonts, but did not report stringocephalids. Conodonts reported in her section A (main hill at Cerros de Los Murciélagos) are not age diagnostic. She considered the entire section to be Late Devonian in age, presumably on the basis of foraminifers. Our composite section includes both Middle and Upper Devonian strata, on the basis of regional lithologic correlation and coral fauna (see below). Although most of the Devonian is present in this area,
Devonian brachiopods of southwesternmost Laurentia
83
Los Chinos Conglomerate (Mississippian) 12 m
Lithologic Symbols conglomeratic barite chert argillite/mudstone siltite/siltstone quartzite/sandstone chert-gravel conglomerate
2 m conglomeratic barite Estimated position of fossiliferous limestone unit exposed 1 km WSW of Rancho Los Chinos
Los Pozos Formation 115 m
dolomitic limestone unconformity
Fossil Symbols brachiopods graptolites radiolarians conodonts ichnofossils
30 METERS
FAULT BRECCIA
0
El Torote Limestone (Upper Ordovician and Lower Silurian?) 1.5 m El Yaqui Chert (Upper Ordovician) > 120 m
Figure 5. Generalized section of the Los Pozos Formation and associated formations north of Rancho Los Chinos, central Sonora. Section measured in 1985 by Poole and R.J. Madrid.
faulting and dolomitization have complicated our reconstruction of the Devonian sequence, and we plan additional stratigraphic and paleontologic studies to determine the age and stratigraphic context of the entire Devonian sequence. During a visit to the Cerros de Los Murciélagos area in 1974, Boucot identified the brachiopod Stringocephalus on the west side of the main hill of Cerros de Los Murciélagos (Block II on Fig. 4). Stromatoporoids and corals are associated with the brachiopods.
W.A. Oliver Jr. (1975, personal commun.) identified large stromatoporoids, auloporoids?, Thamnopora sp., Acanthophyllum sp., cf. Grypophyllum sp., and cf. Neostringophyllum sp. (USGS 95434-SD) in material collected by Boucot at the Stringocephalus locality. According to Oliver, these rugose corals are most likely Middle Devonian in age, which is compatible with the Givetian age of the associated Stringocephalus. Harris processed 4 kg of carbonate rock (sample 00FP-33F), but her 248 g insoluble residue contained
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Boucot et al. MIDDLE PRAGIAN PALEOGEOGRAPHY AND LITHOFACIES CRATON
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Figure 6. Map of present-day United States and northern Mexico showing middle Pragian paleogeography and lithofacies and coeval brachiopod localities. RP—Rancho Placeritos locality; DV—Death Valley; CN—Central Nevada; bold black dots represent brachiopod data points; MSM— approximate trace of hypothetical Jurassic left-slip Mojave-Sonora megashear (dotted line) of Anderson and Silver (2005) and Stewart (2005).
only one conodont, which was not age diagnostic. During fieldwork in 2005, Poole and Amaya-Martínez collected 70 kg of limestone blocks (sample 05FP-41F) containing calcite-replaced stringocephalid specimens; however, the dense lime mudstone to wackestone matrix made it impossible to recover complete specimens suitable for species determination and photography. Stringocephalus valves measured in outcrop are as large as 13 cm in length and 8 cm in width, and most of the embedded shells reveal the prominent median septum. Sandberg and Poole processed 4.3 kg of the carbonate matrix for conodonts, but found none in the insoluble residue. According to X-ray diffractograms, the residue from sample 05FP-41F consists of dolomite and trace amounts of quartz and muscovite (Rhonda Driscoll, 2005, personal commun.), indicating significant secondary dolomitization of the host limestone. Conodonts are typically rare to absent in Stringocephalus-bearing facies worldwide because of the restrictive depositional environment. At Cerros de Los Murciélagos, G.A. Cooper (Cooper and Arellano, 1946) collected the gastropod Orecopia, which is generally considered to be early Late Devonian (Frasnian) in age. Gastropods subsequently collected (samples 01FP-14F and 01FP-15F) from a 20–50-cm-thick biostrome in the same area were identified by R.B.
Blodgett (2004, personal commun.) as Orecopia sp., which he considers to range in age from late Middle Devonian (possibly in Germany) to early Late Devonian. Orecopia is known only from the Frasnian in North America. We found the calcite-replaced Orecopia shells to be abundant in beds as thick as 50 cm, and individuals to be as large as 3 cm in axial height and 3 cm in basal width. From Cooper’s collection, W.A. Oliver Jr. (1975, personal commun.) obtained large stromatoporoids, amphiporoids, Smithiphyllum sp., and Tabulophyllum sp., the age of which is consistent with the age range indicated by Orecopia. The stratigraphic position of the upper “Los Murciélagos Formation” (Block III on Fig. 4), which includes a 0.5–.20-mthick interval of lime mudstone to wackestone containing the gastropod Orecopia sp., has not been firmly established because of complex faulting. However, the age range of Orecopia sp. in the western United States is confined to the early Late Devonian (early Frasnian). The most precise dating of this gastropod is in the type section of the Devils Gate Limestone in central Nevada (CN on Fig. 6). There, R.B. Blodgett (2006, personal commun.) has identified Orecopia sp. in talus at the base of the cliff-forming lower member of this formation. This talus probably is derived
Devonian brachiopods of southwesternmost Laurentia
Locality, field no., and USGS collection no.
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TABLE 1. CONODONT COLLECTIONS THAT CONTAIN Icriodus latericrescens robustus Orr FROM THE “SAN MIGUEL FORMATION” IN WEST-CENTRAL SONORA* Latitude N, Lithology and stratigraphic unit Conodont fauna Conodont zone Longitude W (in order of decreasing taxon and age abundance)
Rancho Placeritos West section (El Batamote 1:50,000 quad) 29°20ƍ39ƍƍ, Lowest debris flow (10 m thick) 96FP-158F 111°16ƍ50ƍƍ of at least three in the 12545-SD formation
CAI
Icriodus latericrescens robustus Orr Polygnathus c. costatus Klapper Po. c. patulus Klapper, Ziegler, and Mashkova Po. serotinus Telford
costatus Zone
5.0 and minor 5.5
99FP-216Fa 12683-SD
29°20ƍ35ƍƍ 111°16ƍƍ42.5ƍƍ
Highest debris flow (9 m thick) in formation; pelmatozoan lime wackestone and packstone clasts
Icriodus latericrescens robustus Po. linguiformis bultyncki Weddige Belodella resima (Philip) Icriodus corniger Wittekindt group
serotinus Zone through costatus Zone (likely costatus Zone, as indicated by the matrix of debris flow)
5.0–5.5
99FP-216Fb 12684-SD
29°20ƍ35ƍƍ 111°16ƍƍ42.5ƍƍ
Mixed-fossil lithoclast lime wackestone to grainstone matrix at same stratigraphic level as 99FP-216Fa
l. latericrescens robustus Po. c. costatus Polygnathus spp. Belodella sp. Panderodus unicostatus (Branson and Mehl) Po. serotinus Pandorinellina expansa Uyeno and Mason
costatus Zone
6.0, lesser 5.5, and 6.5, and minor 7
l. latericrescens robustus Dvorakia sp. Belodella sp. Polygnathus cooperi cooperi Klapper Polygnathus spp.
serotinus Zone through costatus Zone (very late Emsian–early Eifelian)
4.5–5.0
l. latericrescens robustus Po. c. costatus Belodella resima l. culicellus (Bultynck) Panderodus unicostatus Po. linguiformis bultyncki I corniger retrodepressus Bultynck Dvorakia? sp. Po. linguiformis linguiformis Hinde Pandorinellina expansa
costatus Zone
Rancho Placeritos East section (El Batamote 1:50,000 quad) 29°20ƍ38ƍƍ, 00FP-236F Talus surrounding and derived 111°15ƍ35ƍƍ 12723-SD from pelmatozoan-rich cobbles in 8-m-thick debris flow. Medium-dark-gray, bioclastic packstone and grainstone containing small pelmatozoan ossicles 00FP-237F 12724-SD
29°20ƍ38ƍƍ, 111°15ƍ35ƍƍ
Mixed-fossil lithoclast lime wackestone to grainstone matrix at same stratigraphic level as 00FP-236F
4.5
Note: CAI—color alteration index. *Compiled by Harris.
from a conspicuous 1.2-m-thick gastropod biostrome located at the level of conodont sample DVG-21 in a recent revision of the type section (Sandberg et al., 2003, fig. 5). The abundant occurrence of Orecopia sp. at Devils Gate occurs in a ~20-m-thick interval, extending from 7 m below to 13 m above sample DVG-21. More recent conodont redating of this part of the lower member by Sandberg (in Casier et al., 2006) now places the gastropod-bearing interval unequivocally in the late part of the punctata Zone. This dating of Orecopia sp. may be applicable to its Sonoran occurrence (Fig. 2).
Late Devonian (Famennian) Brachiopods from near Rancho Los Chinos Late Late Devonian (Famennian) silicified rhynchonelloid brachiopods (Dzieduszyckia sonora) occur in the middle part of the Los Pozos Formation ~1 km west-southwest of Rancho Los Chinos, 100 km southeast of Hermosillo, Sonora (loc. 3 on Fig. 1; Figs. 2 and 5). The fauna includes silicified worm tubes in a 10-m-thick limestone unit a few meters below a stratiform barite unit. Most brachiopod valves are articulated and some appear to be in growth
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Boucot et al. TABLE 2. NUMBERS OF BRACHIOPOD VALVES PRESENT IN THE PRAGIAN LIMESTONE BED IN THE “SAN MIGUEL FORMATION” NEAR RANCHO PLACERITOS (EAST SECTION). Taxon Ventrals Dorsals Articulated Total no. Percent Anoplia? 4 N.P. N.P. 4 00.2 Discomyorthis 2 5 N.P. 5 00.2 Platyorthis 9 207 1 207 9 Leptocoelia 41 64 17 64 3 Costellirostra N.P. N.P. 1 1 0.05 Acrospirifer 1464 439 7 1464 64 Meristella 439 206 12 439 19 Plicoplasia 32 25 N.P. 32 1 Rensselaeria 52 80 N.P. 80 4 TOTAL 2296 100.4 Note: N.P.—not present.
position. Dark-gray bedded chert units occur directly above and below the fossiliferous limestone. The association of brachiopods and worm tubes in limestone beds has been found in only one other outcrop in the region (5 km northeast of Rancho Los Chinos near the Tunas barite deposit). According to Sandberg (unpublished data), sample 89FP-104F from the Tunas barite locality yielded middle Famennian conodonts of the Late marginifera Zone (Fig. 2). Conodonts of the Early and possibly Late marginifera Zone are also associated with Dzieduszyckia in the upper part of the barite sequence of the Slaven Chert in the northern Shoshone Range in north-central Nevada. Throughout the Minas de Barita and Cerro Cobachi areas, the brachiopods and worm tubes are generally confined to stratiform barite deposited at or near seafloor hydrothermal-vent sites and possibly methane cold seeps. No measured sections in the region include these local beds of fossiliferous limestone; however, we estimate that the limestone units project into the middle part of the Los Pozos Formation measured 2.5 km northeast of Rancho Los Chinos (Fig. 5). In the Minas de Barita and Cerro Cobachi areas, the brachiopod valves and worm tubes are replaced by either barite or quartz. The depositional setting of the brachiopods and worm tubes was previously discussed by Poole and Dutro (1988) and Poole et al. (1991). Abundant radiolarians of a Holoeciscus assemblage in this part of the formation also indicate a Late Devonian age (Benita L. Murchey, 1984, personal commun.). CONODONT DATING OF THE BRACHIOPODBEARING BEDS In 2000, Poole, Harris, and Amaya-Martínez collected two conodont samples (00FP-236F and 00FP-237F) within the lower 3 m of the 8-m-thick debris-flow limestone unit that disconformably overlies the brachiopod-bearing bed in the “San Miguel Formation” in the eastern measured section (Fig. 3). One sample (00FP-236F) was taken from pelmatozoan lime wackestone and packstone clasts and the other sample (00FP-237F) was taken from the mixedfossil lime wackestone to grainstone matrix at the same stratigraphic interval (Fig. 3). Both samples yielded abundant conodonts characteristic of the costatus Zone of early Eifelian age (Fig. 2; Table 1). In 1996, Poole, Amaya-Martínez, and Page collected a conodont sample (96FP-158F) from a 1 m interval within the 10-m-thick
debris-flow limestone unit that is the lowest debris flow in the western section. Harris recovered a relatively abundant fauna of chiefly platform elements of icriodids and polygnathids from the debrisflow limestone (Table 1). The fauna includes, in order of decreasing abundance, Icriodus latericrescens robustus Orr, Polygnathus costatus costatus Klapper, Icriodus sp. fragments (likely fragments of I. lat. robustus), and a small number of Dvorakia sp., Po. costatus patulus Klapper, Po. serotinus Telford, and Belodella sp. The overlapping range of Po. c. costatus and Po. c. patulus restricts the age of the collection to the costatus Zone (early but not earliest Eifelian). Because all the icriodids belong to the I. latericrescens group, the species association probably represents the deeper part of the icriodid-polygnathid biofacies. In addition, the scarcity of ramiform and coniform elements in the collection suggests postmortem hydraulic transport of these lighter conodont elements to even deeper water depositional sites. In 2000, Poole and Harris collected two supplemental conodont samples from within the highest limestone debris-flow unit (9 m thick) in the western measured section, at least 40 m above the base of the debris-flow interval that contains at least three discrete flows. One sample (99FP-216Fa) was taken from pelmatozoan lime wackestone and packstone clasts and the other sample (99FP-216Fb) was taken from the mixed-fossil lime wackestone to grainstone matrix at the same stratigraphic interval (Fig. 3; Table 1). Both samples yielded conodonts. The clasts produced, in order of decreasing abundance, I. lat. robustus, Belodella resima (Philip), Icriodus sp. indet. fragments, Polygnathus linguiformis bultyncki Weddige, and Polygnathus sp. indet. fragments. The matrix sample contained a more diverse fauna, including I. lat. robustus, Po. c. costatus Klapper, Polygnathus sp. indet. (likely fragments of Po. c. costatus), Belodella sp., and a few specimens each of Panderodus unicostatus (Branson and Mehl), Po. serotinus Telford, and Pandorinellina expansa Uyeno and Mason. Conodonts from the clast and matrix samples place both collections in the costatus Zone (Fig. 2; Table 1). Conodont identifications from 11 sample levels in the lower two-thirds of the “San Miguel Formation” are listed below and their stratigraphic positions are shown on Figures 2 and 3. Sample 1 (00FP-241F) contained only a few generically identifiable conodonts (Ozarkodina sp., coniform elements of
Devonian brachiopods of southwesternmost Laurentia Icriodus sp., and two Sb conodont elements of a digyrate apparatus), indicating a latest Pridolian to early Lochkovian age. Sample 2 (97FP-331F) yielded Ozarkodina paucidentata Murphy and Matti, and overlying sample 3 (97FP-332F) contained Icriodus woschmidti Ziegler, O. paucidentata, and O. cf. O. pandora Murphy, Matti, and Walliser. The samples indicate the woschmidti to eurekaensis Zones. Sample 4 (98FP-238F) produced Erika? n. sp. and O. r. remscheidensis Ziegler of probable middle Lochkovian age. Sample 5 (01FP-300F, analyzed by Sandberg), from the brachiopod bed, yielded conodonts of the early to middle Pragian sulcatus to kindlei Zones. Sandberg found that the conodonts agree with the Pragian age indicated by the brachiopods, while suggesting that early to middle Pragian is most consistent with the conodont information. The conodont fauna includes abundant specimens of Icriodus claudiae together with a single specimen of I. steinachensis Morphotype Beta. In Nevada, the joint occurrence of these two taxa characterizes the early to middle Pragian sulcatus to kindlei Zones. Samples 6–10 (96FP-158F; 99FP-216Fa and 99FP-216Fb; 00FP-236F and 00FP-237F) contained Po. c. costatus, Po. c. patulus, and Po. serotinus in 96FP-158F, Po. linguiformis bultyncki Weddige in 99FP-216Fa and 00FP-237F, Po. c. costatus, Po. serotinus, and Pandorinellina expansa Uyeno and Mason in 99FP-216Fb, and abundant Icriodus latericrescens robustus Orr in 00FP-236F. Stratigraphic position and species composition indicate that all five samples are within the costatus Zone (Fig. 2; Table 1). Sample 11 (98FP-139F) yielded Middle to Late Devonian conodonts and styliolinids. Most of the conodonts collected at Rancho Placeritos have a color alteration index (CAI) of 4.5–5.5, indicating short-term, locally higher thermal regimes and/or hydrothermal activity. We interpret the Paleozoic section at Rancho Placeritos to be a large roof pendant in the early Tertiary (Paleocene) batholithic belt. Samples collected from the brachiopod-bearing interval at Cerros de Los Murciélagos yielded no age-diagnostic conodonts. The sparse conodonts collected at Cerro de Los Murciélagos have a color alteration index (CAI) of 2–2.5, indicating that the host rock reached a temperature of at least 100 °C. Conodonts collected from the brachiopod- and worm tubebearing limestone bed near the Tunas barite deposit (sample 89FP-104F) have a CAI of 3, indicating that the host rock reached a temperature of at least 150 °C. BIOGEOGRAPHIC IMPLICATIONS The Rancho Placeritos Pragian-age brachiopod fauna is a strictly Appohimchi Subprovince, Eastern Americas Realm fauna. Its specific, as well as generic, affinities are distinct from those of the Nevada Subprovince of about the same age (Trematospira Zone of the Eureka, Nevada area). The dominant Acrospirifer cf. A. murchisoni is an Appohimchi Subprovince taxon distinct from the Nevada Patriaspirifer kobehana, as is also the case with
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the abundant Meristella cf. M. lata. Platyorthis is widespread in the Appohimchi Subprovince (Boucot and Johnson, 1967) but unknown in the Nevada Subprovince. The Leptocoelia, moderately abundant in Sonora, is unknown in the Nevada Subprovince, where its place is taken by Pacificocoelia and Leptocoelina. Plicoplasia aff. P. cooperi is a form close to the Appohimchi Subprovince P. cooperi that is absent in the Nevada Subprovince. In summary, the overall biogeographic affinities of the “San Miguel Formation” Pragian brachiopods are clearly within the Appohimchi Subprovince, Eastern Americas Realm, although having more distant affinities with the Nevada Subprovince of that Realm despite the latter’s being much closer than the nearest Appohimchi Subprovince occurrences in Oklahoma and northern Chihuahua (Boucot and Johnson, 1967). The Transcontinental arch probably formed a barrier separating the Nevada Subprovince faunas from the Appohimchi Subprovince “San Miguel Formation” brachiopod and conodont faunas. What oceanic processes this barrier affected are unknown, but they may have involved surface current circulation patterns governed in part by the shoreline positions in the region during Pragian time. The paleogeography and lithofacies shown on Figure 6 are based on our present knowledge. The fauna and facies provide evidence for a Pragian lowstand or emergence in Sonora, which is compatible with a worldwide lowering of sea level at this time (see Johnson et al., 1985). Geographical separation of the two shelves may have caused the brachiopod generic assemblages to evolve in reproductive isolation. It is notable that the San Miguel Pragian brachiopod fauna does not contain any Old World Realm endemics, including characteristic North American Cordilleran Region and Australian Tasman Region genera. The presence of Stringocephalus in the “Los Murciélagos Formation” is consistent biogeographically with its widespread Old World Realm occurrences (see Boucot et al., 1966, for the widespread Old World Realm distribution of stringocephalids, and Sun and Boucot, 1999, for some recently noted occurrences). Occurrence of the gastropod Orecopia, known elsewhere only in western North America, also is consistent with the “Los Murciélagos Formation” occurrence. West-Central Sonora and Central Nevada Occurrences of Icriodus latericrescens robustus Orr Compilation of the conodont faunas recovered from upper Lower and Middle Devonian rocks in Sonora revealed that at least one relatively long ranging conodont species Icriodus latericrescens Branson and Mehl, thought to be an Eastern Americas Realm endemic, is also present at several localities in Nevada, according to Sandberg (work in progress). Klapper and Johnson (1980), in their landmark paper on endemism and dispersal of Devonian conodonts, presented 13 tables showing absence or presence of important conodont species in many parts of the world. These include areas in the United States (Great Basin, Midcontinent, Appalachians, and Alaska) and Canada. Subspecies of
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Boucot et al.
I. latericrescens Branson and Mehl were shown to occur from the late Emsian serotinus Zone into the late Givetian Lower dengleri Zone (disparilis Zone of current usage). Icriodus latericrescens robustus Orr, the older subspecies, is known to range from the serotinus Zone into the australis Zone and the younger I. lat. latericrescens has been reported from the Early varcus Zone into the disparilis Zone. Icriodus lat. latericrescens has been documented from several North American, mainly Midcontinent and eastern U.S. areas (Illinois, Iowa, Missouri, New York, Ohio, and Indiana) and from Canada (Ontario), England, Germany, and Spain. Icriodus lat. robustus has been reported previously from New York, Kentucky, and Indiana, and now from several localities in Sonora, Mexico, and in collections made by Sandberg (unpublished data) from central Nevada and southern Utah. Sandberg has identified the conodont I. lat. robustus from the Coils Creek Member of the McColley Canyon Formation and from the overlying Denay Limestone in central Nevada (CN on Fig. 6). These formations are part of the western Laurentian (Cordilleran) continental shelf. The late Early Devonian (Emsian) to early Middle Devonian (Eifelian) conodont Icriodus latericrescens robustus, which co-occurs with the Sonora brachiopods, is ubiquitous around the margins of Laurentia. It has been recorded from New York State, throughout the Midcontinent area (e.g., Missouri, Michigan, Indiana, Illinois), as well as in Texas. It is also widely distributed in central Nevada (Sandberg, unpublished data). The apparent absence of its Givetian descendant Icriodus latericrescens latericrescens from the western United States is puzzling, especially because so many conodont faunas of its age have been recorded. This absence is even more puzzling because the Givetian Taghanic onlap produced a sea-level rise that inundated the margins of Laurentia farther inland than Eifelian seas. Why should the subspecies I. l. robustus have given rise to I. l. latericrescens only east of the Transcontinental arch when its lifespan is a time of cosmopolitanism? Additional study of Givetian rocks in Sonora may provide some answers to this mystery. Table 1 lists occurrences of I. latericescens robustus in westcentral Sonora (loc. 1 on Fig. 1, and RP on Fig. 6). DEPOSITIONAL ENVIRONMENTS The lower two-thirds of the “San Miguel Formation” (Fig. 3) includes two major environments of deposition. On the basis of lithofacies and biofacies, the lower part of the Devonian sequence was deposited in shallow-subtidal and lagoonal settings. The upper part was deposited in a deeper-water setting, beginning with the debris flow that disconformably overlies the Pragian brachiopod bed and continuing upward through the dacryoconarid-bearing beds (Fig. 3). The sequence below the debris-flow bed contains invertebrate fossils characteristic of shallow-water environments. The Pragian brachiopod fauna from the “San Miguel Formation” is dominated by Acrospirifer and abundant Meristella, lesser numbers of Rensselaeria, Leptocoelia, Platyorthis, and
rare Discomyorthis, Costellirostra, and questionable Anoplia. The generic association is not quite the same as previously described communities of this age (Boucot, 1982), including the somewhat similar Beachia, Mutationella, Plicoplasia cooperi, and Hipparionyx Communities. Since 1982, Boucot has decided that the Mutationella Community is best assigned to inner Benthic Assemblage (BA) 3 rather than BA2. Benthic Assemblage 3 corresponds to an approximately mid-shelf position. The largely disarticulated condition of most of the brachiopods suggests some current activity, but not strongly turbulent conditions. An overall BA3 assignment fits best with what is known elsewhere in the Appohimchi Subprovince brachiopod Communities, but this Acrospirifer-dominated Community is distinct from previously described communities, and its defining feature is moderate diversity with dominant Acrospirifer and Meristella. It could be termed the Acrospirifer-Meristella Community. The total number of brachiopods present in the Pragian limestone bed in the “San Miguel Formation” (greatest number of dorsal or ventral valves/genus) (Table 2) makes the dominance relations clear. The varying numbers of ventral versus dorsal valves for different genera is, in large part, an indication of their differing tendency to disarticulate under locally turbulent conditions. Leptocoelia has a high resistance to disarticulation, as is also true elsewhere in the Appohimchi Subprovince, followed by Meristella and then Acrospirifer; the numbers of the other genera are too small to have any significance in this regard. The Devonian section at Cerros de Los Murciélagos (Fig. 4) consists of a lower part of lagoonal to intertidal quartz sandy and silty dolostone (Block I), and an upper part of intertidal to shallow-subtidal limestone and quartz sandy limestone that is locally dolomitized (Blocks II and III). A transitional unit between the lower and upper parts is composed of intertidal rhythmic cycles of quartz sandy dolostone and dolomitized limestone (upper part of Block I). The lower part is devoid of megafossils but contains reworked sparse conodonts that range in age from Middle to Late Devonian. The upper part contains megafossils, including stromatoporoids, brachiopods, corals, gastropods, ostracods, and crinoids. Stringocephalus sp., a large brachiopod characterized by a prominent median septum, occurs in Block II in lime mudstone to wackestone that we interpret as a shallow-subtidal environment. According to Harris, sample 00FP-33F, near the top of the Stringocephalus interval, yielded one juvenile I element of a short-platform Icriodus of indeterminate species, indicating a probable age of Middle to Late Devonian. Sandberg found no conodonts in the matrix containing stringocephalids (05FP-41F). Because Stringocephalus occurs in the upper part of the measured section (Block II on Fig. 4), most of the exposed section may be Middle Devonian in age. The position of Orecopia in our reconstructed section (Block III on Fig. 4) is above Stringocephalus and considered early Frasnian in age on the basis of correlation with Orecopia beds in central Nevada. The limestone sequence of Block III consists of rhythmic cycles (generally 0.5–2 m thick) of shallow-subtidal to intertidal mixed-fossil lime wackestone to mudstone with intercalated quartz sandstone. Orecopia samples 01FP-14F, 01FP-15F, and 06FP-24F are from
Devonian brachiopods of southwesternmost Laurentia Block III (Fig. 4). Hence, the exposed “Los Murciélagos Formation” ranges in age from Middle to Late Devonian (Figs. 2 and 4). The well-sorted, disarticulated condition of the stringocephalids at the “Los Murciélagos Formation” locality is consistent with the bulk of the global occurrences of this relatively rough water, Benthic Assemblage 3 group of brachiopods (inner-shelf position). As is known elsewhere, the Sonora occurrence features a high-dominance, very low diversity stringocephalid fauna (see Wang Yu et al., 1987, p. 53, for a brief discussion of the Stringocephalus Community). The Famennian brachiopod Dzieduszyckia sonora fauna (illustrated in Noll et al., 1984) and worm tubes in the Los Pozos Formation unit west of Rancho Los Chinos are believed to occur in the middle part of the formation (Fig. 5). Nereites-facies ichnofossils present in adjacent beds in the Los Pozos Formation indicate bathyal-zone depths of 200–2000 m. The detrital barite beds were deposited in multiple submarine-fan systems and near seafloor vents associated with the fan deposits (Poole et al., 1991). The Los Pozos Formation consists primarily of very thin layers of interbedded argillite, siltite, sandstone, chert, barite, and sparse limestone deposited in an offshelf deep-marine environment (continental rise) along the southern margin of Laurentia.
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endemic during the Early, but not earliest, Devonian. There is an evolutionary transition between the small to medium-size Dalejina and the large Discomyorthis. Johnson (1970) emphasized the relatively large size of the flabellate muscle field, a character evident in the Mexican specimens. The Sonora Early Devonian fauna yielded a few, poorly preserved specimens, mostly posterior portions, of Discomyorthis. Their chief value is biostratigraphic, as Discomyorthis is abundant in Pragian- and Emsian-age beds, with no evidence for its presence in the Lochkovian. Note that the Treatise (Harper, 2000, p. 818) attribution of a Pragian age, and only from Nevada, needs to be supplemented by its widespread occurrence in the Pragian and Emsian of eastern North America, including such Emsian units as the Bois Blanc and Tomhegan Formations with D. alsus (see Boucot and Johnson, 1968). Boucot and Amsden (1958) provided a listing of Dalejina (=Rhipidomelloides) species that included Orthis musculosa Hall, 1857, now assignable to Discomyorthis. Material. Five dorsal valves, two ventral valves.
SYSTEMATIC PALEONTOLOGY All brachiopod specimens figured in Plates 1–4 are reposited and registered at Estación Regional del Noroeste (ERNO), Instituto de Geología, Universidad Nacional Autónoma de México (U.N.A.M.), Apartado Postal 1039, 83000 Hermosillo, Sonora, México. Catalog numbers accompanying the ERNO abbreviation were offered by Carlos M. González-León of the Instituto de Geología in Hermosillo. Superfamily CHONETOIDEA Bronn, 1862 Family ANOPLIIDAE Muir-Wood, 1962 Subfamily ANOPLIINAE Muir-Wood, 1962 Genus ANOPLIA Hall & Clarke, 1892 Anoplia? sp. (Plate 2, Figures 12–14) Discussion. Four poorly preserved, very small ventral valves of a strongly convex, probably smooth shell with a straight hinge line appears to belong to Anoplia of the A. nucleata type but are too poorly preserved to be identified with certainty. Material. Four ventral valves. Superfamily DALMANELLOIDEA Schuchert, 1913 Family RHIPIDOMELLIDAE Schuchert, 1913 Subfamily RHIPIDOMELLINAE Schuchert, 1913 Genus DISCOMYORTHIS Johnson, 1970 Discomyorthis sp. (Plate 2, Figures 4–7) Discussion. Discomyorthis is basically a very large Dalejina that appears to be an Appohimchi
Family PLATYORTHIDAE Harper, Boucot, & Walmsley, 1969 Genus PLATYORTHIS Schuchert & Cooper, 1931 Platyorthis sp. (Plate 1, Figures 1–7) Discussion. A moderate number of partly silicified specimens, with most having only the dorsal cardinalia preserved, were recovered; these retain the diagnostic character of Platyorthis. Only a few specimens are well enough preserved to make clear the typical plano-convex, circular form of the genus. However, the typical cardinalia, ventral muscle field, and costellae are present on a number of specimens. Material. 207 dorsal valves, 9 ventral valves. Superfamily RHYNCHOTREMATOIDEA Family LEPTOCOELIIDAE Boucot & Gill, 1956 Genus LEPTOCOELIA Hall, 1857 Leptocoelia sp. (Plate 3, Figures 7–12) Discussion. A reasonable number of moderately well preserved specimens of Leptocoelia were recovered. They are smaller than typical L. flabellites adult specimens, but all possess the internal and external morphology of the genus. Contrary to the revised Treatise account (Savage, 2002), hinge teeth in Leptocoelia are not crenulate. The dorsal valve tends to be flat or very gently convex, as contrasted with the moderately convex ventral valve, and the shell outline is subcircular. The available
Devonian brachiopods of southwesternmost Laurentia material is coarsely silicified, which precludes specific determination. Material. 17 articulated specimens, 64 dorsal valves, 41 ventral valves. Family EATONIIDAE Schmidt, 1965 Genus COSTELLIROSTRA Cooper, 1942 Costellirostra cf. C. peculiaris (Conrad, 1841) (Plate 4, Figures 1–4) Discussion. The single, articulated specimen of Costellirostra has the typical dorsibiconvex form of the type species, as well as its costellate ornament. Internally, the ventral muscle field conforms to the definition of the genus (Savage, 2002). Material. One articulated specimen. Superfamily DELTHYRIDOIDEA Phillips, 1841 Family DELTHYRIDIDAE Phillips, 1841 Subfamily ACROSPIRIFERINAE Termier & Termier, 1949 Discussion. Contrary to Carter et al.’s (1994) placement of Patriaspirifer in the Hysterolitinae, together with many other non-hysterolitinid genera, Boucot (1975, p. 364–365) pointed out that the cardinalia of typical Hysterolites and Multispirifer are very distinct from those present in the acrospiriferids. Genus ACROSPIRIFER Helmbrecht & Wedekind, 1923 Acrospirifer cf. A. murchisoni (De Castelnau, 1843) (Plate 3, Figures 1–6) Discussion. Johnson (1995) reviewed the reasons, based on fine ornamentation, for removing “Spirifer” murchisoni from Acrospirifer. However, the use of fine ornamentation for differentiating genera and even suprageneric categories within the delthyrids is a problem as it is so bound up with the vagaries of preservation. For example, Boucot (1973, p. 46, pl. 17, fig. 9) made the point with Acrospirifer atlanticus, the fine ornamentation of which varies from smooth to fimbriate and in between on a single specimen! Johnson maintained that the type species of Acrospirifer, A. primaevus, is capillate, on the basis of Vandercammen’s (1963) study of Belgian
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material assigned to that species (Steininger’s, 1853, brief discussion of the type species did not consider fine ornamentation). Until a thorough study of this Transatlantic problem is undertaken, there is a good possibility that Patriaspirifer is a junior synonym of Acrospirifer. However, the distinctive external morphology of P. kobehana, Johnson’s type species of Patriaspirifer, makes it clearly distinct from “Spirifer” murchisoni, including the Sonora material considered here. Material. 439 dorsal valves, 1464 ventral valves, 7 articulated specimens. Superfamily ATHYRIDOIDEA Davidson, 1881 Family MERISTELLIDAE Waagen, 1883 Genus MERISTELLA Hall, 1859 Meristella cf. M. lata (Hall, 1859) (Plate 1, Figures 8–14) Discussion. Second in abundance only to Acrospirifer, a large number of reasonably well preserved Meristella were recovered. Their characteristics in all regards are very similar to those of M. lata, but without carrying out a careful biometric study of this species from a number of localities, it would be premature to be certain of the specific identification. For example, the Sonora material compares very closely with that of Oriskany Sandstone age from northern Maine (Boucot, 1973). The conspicuous dorsal median septum supports a prominent cardinal plate bearing a distinct medial depression, whereas the ventral muscle field is distinctly flabellate, and the shell externally is subequally biconvex with an almost circular outline; such characters are also present in typical M. lata. Material. 206 dorsal valves, 439 ventral valves, 12 articulated specimens. Superfamily AMBOCOELIODEA, George, 1931 Family AMBOCOELIIDAE George, 1931 Genus PLICOPLASIA Boucot, 1959 Plicoplasia aff. P. cooperi Boucot, 1959 (Plate 2, Figures 8–11) Discussion. Some reasonably well
PLATE 1. Figures 1–7 Platyorthis sp. (×2). (1) Poorly silicified posterior of ventral valve. Note the costellate ornamentation; ERNO-8475. (2) Posterior portion of ventral valve interior. Note the cordate muscle field and dental lamellae; ERNO-8475. (3) Side view of exterior of articulated specimen. Note the plano-convex morphology; ERNO-8476. (4) Dorsal view of exterior. Note the relatively flat, costellate valve; ERNO-8476. (5) Ventral view of exterior. Note the gentle curvature of the costellate valve; ERNO-8476. (6) Partially silicified dorsal valve exterior. Note the costellate ornamentation; ERNO-8477. (7) Dorsal valve interior. Note the typical platyorthid cardinal process and brachiophores; ERNO-8477. Figures 8–14 Meristella cf. M. lata (8) Dorsal interior (×2). Note the hinge plate with a well-developed median trough supported by a lengthy median septum; ERNO-8478. (9) Dorsal valve exterior (×2). Note the smooth exterior; ERNO-8478. (10) Ventral valve exterior (×2). Note the stout hinge teeth and the flabellate muscle field; ERNO-8479. (11) Ventral valve interior (×2). Note the stout hinge teeth and the flabellate muscle field; ERNO-8479. (12) Exterior side view of articulated specimen (×2). Note the incurved ventral beak; ERNO-8480. (13) Exterior dorsal view (×3); ERNO-8480. (14) Exterior posterior view (×3); ERNO-8480.
Devonian brachiopods of southwesternmost Laurentia preserved, disarticulated specimens of Plicoplasia aff. P. cooperi were recovered from the Sonora fauna. Despite the relatively coarse silicification, these specimens clearly are very close to P. cooperi, as illustrated by Boucot (1959). Plicoplasia cooperi is also known in beds of Oriskany age from Gaspé and Saint Helens Island, Montreal (Boucot et al., 1986; in both Oriskany Sandstone and Glenerie Limestone lithologies). The Mexican specimens may be a distinct species as they have only one distinct plication lateral to the large plications bounding the dorsal sulcus, with only a weak suggestion of a second, as contrasted with the two present in P. cooperi. From the limited available material it is unclear whether the cardinal process is medially cleft as is the case with P. cooperi. The ventral interarea of P. cooperi is much higher than the very narrow one present in the Sonora material, and the anterior margins are unpreserved owing to the poor silicification. Material. 25 dorsal valves, 32 ventral valves. Superfamily STRINGOCEPHALOIDEA King, 1850 Family CENTRONELLIDAE Waagen, 1882 Subfamily RENSSELAERIINAE Raymond, 1923 Genus RENSSELAERIA Hall, 1859 Rensselaeria aff. R. ovoides (Eaton, 1832) (Plate 2, Figures 1–3) Discussion. Disarticulated dorsal and ventral valves of a relatively small Rensselaeria are one of the moderately abundant elements of the fauna. The posterior portions of the shells are reasonably well preserved, as contrasted with the total absence of the anteriormost portions. None of the exteriors, probably owing to relatively coarse silicification, show any evidence of the relatively fine costellae present anteriorly on typical Rensselaeria ovoides. The umbonal region appears smooth, as is characteristic of the genus, but again the relatively coarse
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silicification would obscure any costellate ornament should it have been present. The beak region of the ventral valve is not as strongly incurved as in younger, Oriskany-age examples of the genus. The convexities of the posterior portions of both valves are convex, with the ventral valve apparently being somewhat more convex than the dorsal, as is the case with Rensselaeria ovoides. Internally, however, the typically short, virtually obsolete dental lamellae are prominent, combined with the deeply impressed, elongate muscle field. The umbonal cavities are filled with secondary material that bounds the deeply impressed muscle field. The dorsal cardinalia are typical for Rensselaeria, with their contronellid form featuring a cardinal plate formed from medially fused hinge plates, supported by well-developed crural plates. Because of the absence of complete specimens, a new species is not erected. Boucot and Johnson (1967, p. 39) laid out the evolutionary relations of the NanothyrisRensselaeria-Etymothyris-Amphigenia lineage. The relatively small size of the Sonora Rensselaeria fits very well with the Sonora specimens and represents an evolutionary stage earlier than the large adult size R. ovoides, but is larger in turn than the largest Nanothyris, as well as having a more deeply impressed ventral muscle field. Critical in Boucot and Johnson’s (1967, pl. 2, figs. 11–13) thinking was the morphologically and stratigraphically intermediate position of what they designated as Prorensselaeria nylanderi (1967, pl. 2, figs. 11–13) from Aroostook County, Maine, now known to be a large Nanothyris (Boucot in Boucot and Wilson, 1994, p. 1003) unrelated to the mutationellinid Prorensselearia. The small Sonora Rensselaeria is more advanced than the large Nanothyris in the more deeply impressed nature of its ventral muscle field, i.e., intermediate between the large Nanothyris from northern Maine
PLATE 2. Figures 1–3 Rensselaeria aff. R. ovoides (×2) (1) Posterior portion of ventral interior. Note the deeply impressed muscle field and the well-developed dental lamellae; ERNO-8481. (2) Posterior portion of ventral exterior. Note the relatively coarse silicification and the apparently smooth exterior; ERNO-8481. (3) Posterior portion of dorsal interior. Note the cardinal plate supported by stout crural plates; ERNO-8482. Figures 4–7 Discomyorthis sp. (4) Posterior portion of dorsal interior (×4). Note the typical rhipidomellid cardinalia; cardinal process, brachiophores and posterior, broad raised area dividing the adductor field; ERNO-8483. (5) Posterior portion of dorsal interior (×4). Note the typical rhipidomellid cardinalia and peripheral traces of the costellae; ERNO-8484. (6) Exterior portion of dorsal valve (×4). Note the costellate ornamentation; ERNO-8484. (7) Fragmentary ventral interior (×2). Note the large, flabellate muscle field; ERNO-8485. Figures 8–11 Plicoplasia aff. P. cooperi (×6). (8) Exterior of ventral valve. Note the presence of one costa lateral to the fold plus a partly developed second lateral costa; ERNO-8486. (9) Interior of ventral valve. Note the very narrow interarea and the stout hinge teeth; ERNO-8486. (10) Exterior of dorsal valve. Note the well-developed costa in the sulcus, the well-developed costa lateral to the sulcus, and the weak beginnings of a second costa; ERNO-8487. (11) Interior of dorsal valve. Note the relatively narrow interarea, short brachiophores medial to the dental sockets and apparently undivided cardinalia; ERNO-8487. Figures 12–14 Anoplia? sp. (×6). (12) Dorsal view of articulated specimen; ERNO-8488. (13) Ventral exterior. Note the smooth exterior; ERNO-8489. (14) Poorly preserved ventral interior. Note the typical outline, straight hinge line, and posterior curvature of the valve; ERNO-8489.
PLATE 3. Figures 1–6 Acrospirifer cf. A. murchisoni (×2). (1) Posterior view of articulated specimen, ventral valve above; ERNO-8490. (2) Side view of exterior, ventral valve above. Note the slightly incurved ventral valve; ERNO-8490. (3) Ventral valve exterior. Note the relatively U-shaped interspaces and costal cross sections; ERNO-8490. (4) Posterior portion of dorsal exterior. Note the relatively broad fold; ERNO-8491. (5) Posterior portion of interior of dorsal valve. Note the position of the dental sockets, and the low myophragm; ERNO-8491. (6) Interior of ventral valve. Note the stout hinge teeth; ERNO-8492. Figures 7–12 Leptocoelia sp. (×6). (7) Posterior view of exterior, ventral valve above. Note the plano-convex valve curvatures; ERNO-8493. (8) Anterior view of exterior, ventral valve above; ERNO-8493. (9) Dorsal interior. Note the cardinal process, cardinal plate, and brachiophores; ERNO-8494. (10) Ventral view of exterior. Note the strong costae; ERNO-8493. (11) Dorsal view of exterior. Note the strong costae; ERNO-8493. (12) Ventral interior. Note the strong hinge teeth; ERNO-8495.
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PLATE 4. Figures 1–4 Costellirostra cf. C. peculiaris. (1) Posterior view, ventral valve above (×2). Note the strongly convexo-planar curvature; ERNO-8496. (2) Exterior of ventral valve (×2). Note the valve outline and the broad, weak, anterior sulcus; ERNO-8496. (3) Interior of ventral valve (×2). Note the elevated rim surrounding the deeply impressed muscle field; ERNO-8496. (4) Exterior of ventral valve (×6). Note the presence of costellae, despite the presence of coarse silicification; ERNO-8496.
and Rensselaeria ovoides. It is worth noting that the Lochkovian Becraft Limestone of Schoharie Valley, New York, where Conrad (1842) presumably obtained his “A.” aequiradiata, which Cloud (1942) assigned to Nanothyris, is a large form with welldeveloped costellae, and a more slender form than the Aroostook County shell. The interior and dorsal exterior of the Schoharie Valley shell have never been described or illustrated, but one would predict that it would have cardinalia and ventral musculature similar to those of the Aroostook County specimens. Rickard (1975) indicated that in Schoharie Valley, the Oriskany Sandstone rests disconformably on Becraft Limestone or the lowest Alsen Formation. All of this information is consistent with the Sonora specimens belonging to a pre–Oriskany age Rensselaeria. Material. 80 dorsal valves, 52 ventral valves.
TECTONIC SIGNIFICANCE Biogeographic distribution of the Appohimchi brachiopod fauna at Rancho Placeritos indicates continuous Early Devonian shelf deposition along the entire southern margin of Laurentia (Fig. 6). Many coeval brachiopods in central Nevada (Nevada Subprovince) are significantly different from those in central Sonora. The southwest-trending Transcontinental arch probably formed a barrier preventing migration and mixing of many genera and species of brachiopods from the southern shelf of Laurentia in northern Mexico to the western shelf (Cordilleran miogeocline) in the western United States. In addition to the arch in New Mexico and Arizona (Fig. 6), environmental factors, such as sedimentation, ocean currents, latitude, and climate, probably contributed to confining some taxa to their respective shelves. The Late Permian Sonora allochthon (Fig. 1) juxtaposes different Devonian facies. Allochthonous Paleozoic deep-water
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continental-rise and ocean-basin rocks, including the Devonian Los Pozos Formation (Fig. 5), were thrust 50–200 km onto coeval Paleozoic shallow-water carbonate-shelf rocks and overlying Permian foredeep flysch of the southern margin of Laurentia (Poole et al., 2005). Paleozoic stratigraphic, faunal, and structural characteristics of rocks of the southern shelf and superjacent late Paleozoic Sonora allochthon are unique to Sonora, and are significantly different from coeval rocks and their superjacent allochthons of the western shelf in Nevada and California (Poole et al., 2005). These biogeographic and tectonic findings contradict the hypothesis of large-magnitude (600–1100 km) left-lateral offset of Cordilleran carbonate-shelf rocks from the Death Valley region (DV on Fig. 6) of California to Sonora along a northwest-striking Jurassic Mojave-Sonora megashear (Fig. 6), as advocated by Anderson and Silver (2005), Stewart (2005), and other workers (see papers in Anderson et al., 2005, for other interpretations). ACKNOWLEDGMENTS AND RESPONSIBILITY The U.S. Geological Survey supported fieldwork of the American geologists and the University of Sonora supported fieldwork of Amaya-Martínez. Boucot’s fieldwork in 1974 was supported by the U.S. National Science Foundation. The brachiopods at Rancho Placeritos were collected in 2000 by Poole, Amaya-Martínez, and Page (00FP-239F) and in 2001 by Poole, Amaya-Martínez, and University of Sonora students Daniel Arturo Amaya-Zepeda, Francisco Montiel-Salas, and José Alfredo Vindiola (01FP-300F). The brachiopods and worm tubes near Rancho Los Chinos were collected in 2003 by Poole (03FP-397F and 03FP-401F) and in 2004 by Poole and barite mine visitors (04FP-190F and 04FP-296F). Boucot identified the brachiopods. Harris identified most of the conodonts in the Rancho Placeritos and Cerros de Los Murciélagos sections, and Sandberg identified the conodonts in the brachiopodbearing beds at Rancho Placeritos and Tunas barite deposit near Rancho Los Chinos. Boucot is responsible for discussing the brachiopods, interpreting their biogeographic significance and paleoecology, whereas Poole and others are responsible for interpreting their depositional and tectonic setting. Boucot thanks Robert B. Blodgett for photographing the Pragian brachiopods and preparing the plates. We thank James E. (Jed) Day, Daniel R. Shawe, Peter E. Isaacson, and Robert B. Blodgett for their helpful reviews of the manuscript, and Barbara J. Ramsey and Norma J. Maes for computer preparation of the figures. REFERENCES CITED Anderson, T.H., and Silver, L.T., 2005, The Mojave-Sonora megashear—Field and analytical studies leading to the conception and evolution of the hypothesis, in Anderson, T.H., Nourse, J.A., McKee, J.W., and Steiner, M.B., eds., The Mojave-Sonora megashear hypothesis: Development, assessment, and alternatives: Geological Society of America Special Paper 393, p. 1–50, doi: 10.1130/2008.393(01).
Anderson, T.H., Nourse, J.A., McKee, J.W., and Steiner, M.B., editors, The Mojave-Sonora megashear hypothesis: Development, assessment, and alternatives: Geological Society of America Special Paper 393, 712 p. and CD-ROM. Boucot, A.J., 1959, Early Devonian Ambocoeliinae (Brachiopoda): Journal of Paleontology, v. 33, no. 1, p. 16–24. Boucot, A.J., 1973, Early Paleozoic brachiopods of the Moose River Synclinorium, Maine: U.S. Geological Survey Professional Paper 784, 81 p. Boucot, A.J., 1975, Evolution and extinction rate controls: Amsterdam, Elsevier, 427 p. Boucot, A.J., 1982, Ecostratigraphic framework for the Lower Devonian of the North American Appohimchi Subprovince: Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, v. 163, p. 81–121. Boucot, A.J., and Amsden, T.W., 1958, Part IV, New genera of brachiopods, in Stratigraphy and paleontology of the Hunton Group in the Arbuckle Mountain region: Oklahoma Geological Survey Bulletin 78, p. 159–174. Boucot, A.J., and Johnson, J.G., 1967, Paleogeography and correlation of Appalachian Province Lower Devonian sedimentary rocks: Tulsa Geological Society Digest, v. 35, p. 35–87. Boucot, A.J., and Johnson, J.G., 1968, Brachiopods of the Bois Blanc Formation: U.S. Geological Survey Professional Paper 584-B, 27 p. Boucot, A.J., and Wilson, R.A., 1994, Origin and early radiation of terebratuloid brachiopods: Thoughts provoked by Prorensselaeria and Nanothyris: Journal of Paleontology, v. 68, no. 5, p. 1002–1025. Boucot, A.J., Johnson, J.G., and Struve, W., 1966, Stringocephalus, ontogeny and distribution: Journal of Paleontology, v. 40, no. 6, p. 1349–1364. Boucot, A.J., Brett, C.E., Oliver, W.A., Jr., and Blodgett, R.B., 1986, Devonian faunas of the Sainte-Hélène Island breccia, Montréal, Quebec, Canada: Canadian Journal of Earth Sciences, v. 23, no. 12, p. 2047–2056. Brunner, P., 1975, Estudio estratigráfico del Devónico en el área de El Bísani, Caborca, Sonora: Revista del Instituto Mexicano del Petróleo, v. 7, no. 1, p. 16–45. Carter, J.L., Johnson, J.G., Gourvennec, R., and Hou, H.-f., 1994, A revised classification of the spiriferid brachiopods: Annals of the Carnegie Museum, v. 64, no. 4, p. 327–374. Casier, J.-G., Berra, I., Olempska, E., Sandberg, C., and Préat, A., 2006, Ostracods and facies of Early and Middle Frasnian at Devils Gate in Nevada: Relationship to the Alamo Event: Acta Palaeontologica Polonica, v. 51, no. 4, p. 813–828. Cloud, P.E., Jr., 1942, Terebratuloid brachiopods of the Silurian and Devonian: Geological Society of America Special Paper 38, 182 p. Conrad, T.A., 1841, Fifth annual report on the paleontology of the state of New York: New York Geological Survey Annual Report, v. 5, p. 25–57. Conrad, T.A., 1842, Observations on the Silurian and Devonian systems of the United States, with descriptions of new organic remains: Journal of the Academy of Natural Sciences of Philadelphia, v. 8, part 2, p. 228–280. Cooper, G.A., and Arrellano, A.R.V., 1946, Stratigraphy near Caborca, northwest Sonora, Mexico: American Association of Petroleum Geologists Bulletin, v. 40, no. 4, p. 606–619. De Castelnau, F., 1843, Essai sur le système silurien de l’Amérique septentrionale: Paris, C.P. Bertrand, Académie Royale des Sciences, Institut de France, 56 p. and 27 pls. (p. 36–45 and pls. 12–15 describe brachiopods). Eaton, A., 1832, Geological text-book, for aiding the study of North American geology; being a systematic arrangement of facts, collected by the author and his pupils (second edition): Albany, New York, Websters and Skinners, 134 p. (p. 45–46 describe brachiopods). González-León, C.M., 1986, Estratigrafía del Paleozoico de la Sierra del Tule, noreste de Sonora: Universidad Nacional Autónoma de México, Instituto de Geología, Revista, v. 6, no. 2, p. 117–135. Hall, J., 1859, Palaeontology of New York, v. 3, part 1 (text) and part 2 (plates), containing descriptions and figures of the organic remains of the Lower Helderberg group and Oriskany sandstone: Natural History of New York: Albany, Geological Survey of New York, Charles van Benthuysen (printer), part 1, p. 1–532 (text), and part 2, p. 102–478 (120 pls.). Text printed in 1859 and plates printed in 1861. Harper, D.A.T., 2000, Suborder Dalmanellidina, p. 782–844, in Williams, A., et al., Treatise on Invertebrate Paleontology, Part H, Brachiopoda, Revised, Volume 3, Linguliformea, Craniiformea, and Rhynchonelliformea (part): Geological Society of America (and University of Kansas Press), p. 424–919. Johnson, J.G., 1970, Great Basin Lower Devonian Brachiopoda: Geological Society of America Memoir 121, 421 p.
Devonian brachiopods of southwesternmost Laurentia Johnson, J.G., 1995, Patriaspirifer, a new genus of Lower Devonian spiriferid brachiopods: Journal of Paleontology, v. 69, no. 1, p. 198. Johnson, J.G., Klapper, G., and Sandberg, C.A., 1985, Devonian eustatic fluctuations in Euramerica: Geological Society of America Bulletin, v. 96, no. 5, p. 567–587, doi: 10.1130/0016-7606(1985)96<567:DEFIE>2.0.CO;2. Klapper, G., and Johnson, J.G., 1980, Endemism and dispersal of Devonian conodonts: Journal of Paleontology, v. 54, no. 2, p. 400–455. Noll, J.H., Jr., 1981, Geology of the Picacho Colorado area, northern Sierra de Cobachi, central Sonora, Mexico [M.S. thesis]: Flagstaff, Northern Arizona University, 169 p. Noll, J.H., Jr., Dutro, J.T., Jr., and Beus, S.S., 1984, A new species of the Late Devonian (Famennian) brachiopod Dzieduszyckia from Sonora, Mexico: Journal of Paleontology, v. 58, no. 6, p. 1412–1421. Poole, F.G., Boucot, A.J., Amaya-Martínez, R., Sandberg, C.A., Harris, A.G., and Page, W.R., 2003, Early Devonian brachiopods in west-central Sonora indicate depositional continuity along southern margin of Laurentia: Geological Society of America Abstracts with Programs, v. 35, no. 4, p. 13–14. Poole, F.G., and Dutro, J.T., Jr., 1988, Devonian fossils in seafloor hydrothermalvent barites of Nevada (USA) and Sonora (Mexico) [extended abstract]: Symposium on Barite and Barite Deposits, Kutná Hora, Czechoslovakia: Prague, Geological Survey of Czechoslovakia, Abstracts, p. 51–53. Poole, F.G., Madrid, R.J., and Oliva-Becerril, J.F., 1991, Geological setting and origin of stratiform barite in central Sonora, Mexico, in Raines, G.L., Lisle, R.E., Schafer, R.W., and Wilkinson, W.H., eds., Geology and ore deposits of the Great Basin: Proceedings, Geological Society of Nevada Symposium, Reno/Sparks, Nevada, April 1990, v. 1, p. 517–522. Poole, F.G., Perry, W.J., Jr., Madrid, R.J., and Amaya-Martínez, R., 2005, Tectonic synthesis of the Ouachita-Marathon-Sonora orogenic margin of southern Laurentia: Stratigraphic and structural implications for timing of deformational events and plate-tectonic model, in Anderson, T.H., Nourse, J.A., McKee, J.W., and Steiner, M.B., eds., The Mojave-Sonora megashear hypothesis: Development, assessment, and alternatives: Geological Society of America Special Paper 393, p. 543–596.
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Rickard, L.V., 1975, Correlation of the Silurian and Devonian rocks in New York State: New York State Museum and Science Service, Map and Chart Series no. 24, 16 p., 4 pls. Sandberg, C.A., and Ziegler, W., 1996, Devonian conodont biochronology in geologic time calibration: Senckenbergiana Lethaea, v. 76, no. 1/2, p. 259–265. Sandberg, C.A., Morrow, J.R., Poole, F.G., and Ziegler, W., 2003, Middle Devonian to Early Carboniferous event stratigraphy of Devils Gate and Northern Antelope Range sections, Nevada, U.S.A.: Courier Forschungsinstitut Senckenberg, v. 242, p. 187–207. Savage, N.M., 2002, Superfamily Rhynchotrematoida, p. 1047–1091, in Williams, A., et al., Treatise on Invertebrate Paleontology, Part H, Brachiopoda, Revised, Volume 4, Rhynchonelliformea (part): Geological Society of America (and University of Kansas Press), p. 921–1688. Steininger, J., 1853, Geognostische Beschreibung der Eifel: Trier, Germany, Druck und Verlag der Fr. Lintz’schen Buchhandlung, 144 p., 10 pls. Stewart, J.H., 2005, Evidence for Mojave-Sonora megashear—Systematic leftlateral offset of Neoproterozoic to Lower Jurassic strata and facies, western United States and northwestern Mexico, in Anderson, T.H., Nourse, J.A., McKee, J.W., and Steiner, M.B., eds., The Mojave-Sonora megashear hypothesis: Development, assessment, and alternatives: Geological Society of America Special Paper 393, p. 209–231. Sun, Y.L., and Boucot, A.J., 1999, Ontogeny of Stringocephalus gubiensis and origin of Stringocephalus: Journal of Paleontology, v. 73, no. 5, p. 860–871. Vandercammen, A., 1963, Spiriferidae du Dévonien de la Belgique: Institut Royal des Sciences Naturelles de Belgique, Mémoire 150, 181 p. Wang, Yu., Boucot, A.J., Rong Jia-yu, and Yang Xue-chang, 1987, Community paleoecology as a geologic tool: The Chinese Ashgillian-Eifelian (latest Ordovician through early Middle Devonian) as an example: Geological Society of America Special Paper 211, 100 p.
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The Geological Society of America Special Paper 442 2008
Devonian brachiopods from Northeastern Washington: Evidence for a non-allochthonous terrane and Late Devonian biogeographic update Peter E. Isaacson* Department of Geological Sciences, University of Idaho, Moscow, Idaho 83844-3022, USA
ABSTRACT There are several isolated outcrops (outliers?) of Devonian rocks in the Pacific Northwest (Washington and Oregon), USA. A locality in northeastern Washington, Limestone Hill, is considered in detail, and other small outcrops in northwest Washington and central Oregon are discussed. Limestone Hill is a Paleozoic outlier. The locality has Ordovician and Silurian (Llandovery and Wenlock) strata, Lower Devonian (Loch stone conglomerate, and Upper Devonian (Frasnian) carbonates with fossils. It has long been known that the area has many allochthons, and it has been assumed that Limestone Hill represents lithologies deposited much farther west. More recent data suggest that Limestone Hill is parautochthonous. Several brachiopod taxa, previously unknown from the Frasnian portion of Limestone Hill, have been found recently and are described herein. The brachiopod faunule consists of Emanuella sp., “Allanella” engelmanni, Cyrtina sp., Thomasaria sp., and an athyridid. These brachiopods appear to be like coeval faunas in Idaho, Montana, Utah, and Nevada, although more species assignments must be made. Frasnian brachiopods are in serious need of updates, as Famennian miospore and acritarch data suggest significant basin restriction and reduced seaway connectivity, with at least ephemerally extensive land areas with ubiquitous land plant taxa. Keywords: Late Devonian, biogeography, Washington State, cosmopolitanism.
INTRODUCTION
Located in the tectonically complex Kootenay Arc (Yates, 1970), Limestone Hill is surrounded by Jurassic igneous rocks and gives every appearance of an outlier. There are no coeval Devonian rocks for 200 km in all directions. Elsewhere in the immediate region are Ordovician graptolitic shales and rather extensive Cambrian carbonates. The latter have received much attention in
A small yet very important fossil locality in northeastern Washington state has yielded a variety of conodont, coral, and brachiopod faunules. They occur in an areally small carbonate outcrop outlier, known locally as Limestone Hill (Fig. 1).
*
[email protected] Isaacson, P.E., 2008, Devonian brachiopods from Northeastern Washington: Evidence for a non-allochthonous terrane and Late Devonian biogeographic update, in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 99–106, doi: 10.1130/2008.442(06). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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Figure 1. Location of Limestone Hill, Pend Orielle County, northeastern Washington.
the recent past. Devonian corals are known from Limestone Hill, but this paper presents the first silicified brachiopods recovered from the carbonate succession. There has been speculation about where the North American craton/margin terminates and possible outboard or allochthonous terranes form a distinct geologic boundary between “Laurentia craton” systems and possible suspect terranes across northern Washington State. Owing to the limited exposure and lateral continuity of carbonates at Limestone Hill, as well as several thrust faults, it is tempting to assign these rocks to an accreted terrane. GEOLOGY OF LIMESTONE HILL Park and Cannon (1943) and Enbysk (1956) reported on a carbonate outlier in Pend Oreille County, Washington (Fig. 1). Dings and Whitebread (1965) gave a comprehensive account of the Devonian units, stating that “overlying the [Lower and Middle Ordovician] Ledbetter Slate ... is a thick sequence of black slates and argillites which contains a heterogeneous assemblage of limestone, conglomerate, sandstone, and quartzite lenses.” They further stated that the bulk of these “Silurian and Devonian” units are “black carbonaceous slate and argillite.” Sorauf (1972) reported that there are few actual carbonate banks on Limestone Hill, though some massive carbonate outcrops suggested small bioherms in the succession. Colonial corals and stromatoporoids were probably occupants. Greenman et al. (1977) provided the most comprehensive reassessment of Limestone Hill, showing its stratigraphic units (Fig. 2). The five units at Limestone Hill begin with Silurian graptolites in Ledbetter lithologies (Unit A) as shown, followed by a poorly exposed granule conglomerate (Unit B). Unit C has limestone conglomerate with “Middle Devonian matrix,” and Unit D, massive carbonates, records the varcus (Givetian) conodont Zone. Unit E contains the brachiopod faunule described herein, and it represents the upper Frasnian gigas Zone. A fault appears to have removed lower Frasnian units. Greenman et al. (1977) suggested that the conglomerates represent a “Silurian or
Devonian disturbance” of unknown affinities. The Antler orogeny was considered, but its age is seemingly too young. Structural Geology and Basement Affinities Northeast Washington lies within an intensely folded and faulted region (the Kootenay Arc) that was also affected by Cretaceous (Kaniksu) plutons. There are several faults in the areas near Limestone Hill, and it is unusual to find sedimentary units that have not been deformed. Dings and Whitebread (1965) described several complex thrust faults (the Metaline, Washington Rock, Lee, and Argillite thrusts) that on first inspection appear to emplace Cambrian carbonates over relatively incompetent Ledbetter lithologies. Drilling, however, showed that the thrusts are complexly imbricated and all sedimentary units in the area are allochthonous. Although Dings and Whitebread gave no estimates of amount of displacement afforded by these faults, it is quite clear that original deposition of the Paleozoic units was farther to the west. This has stimulated speculation that Limestone Hill may well be part of a suspect terrane. Given the hypothesis of Ross et al. (1992) that source area(s) for the Mesoproterozoic part of the Belt Series was from the west, and that Australia may have been part of this source, speculation on the western margin of Laurentia in Paleozoic time may become complicated. The usual model for Laurentia is a passive margin developing after Proterozoic rifting (Stewart and Suczek, 1977). Establishing a position for Limestone Hill within the passive-margin setting depends upon which part of the hill’s succession one uses. Ordovician Ledbetter lithologies suggest “eugeoclinal” deposition similar to western assemblage (eugeoclinal) rocks of Idaho (Saturday Mountain Formation) and Nevada (Vinini and Valmy Formations, with their graptolitic shales). Armstrong et al. (1977) suggested that the 0.706 strontium isotope ratio line should be west of northeastern Washington, at the Okanogan Valley. They worried, however, that the craton between there and Limestone Hill could possibly be an accreted
Devonian brachiopods from Northeastern Washington
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Figure 2. Stratigraphy of Limestone Hill (Greenman et al., 1977). Metaline Limestone is Cambrian; Ledbetter Formation shales are Ordovician; Unit A, a mix of mudstone and limestone, is Silurian. Units B and C are limestone conglomerates, presumably Early and Middle Devonian in age. Unit D, varcus Zone (Givetian, Middle Devonian), is massive limestone. Unit E, gigas Zone (Frasnian, Late Devonian), contains the brachiopod faunule described herein.
terrane, thereby suggesting that Limestone Hill could be similarly accreted. Whitehouse et al. (1992) studied Pb and rare earth elements (Sm, Nd) within Mesozoic granitic rocks in the same region as Limestone Hill. They concluded that low Pb values within the plutons demonstrated that the Belt Supergroup was a possible source for later Proterozoic and Paleozoic sediments, and that the values support an old metamorphic crust as a source. Values suggested a Late Proterozoic Wyoming Subprovince correlation, and that the basement of western Laurentia was well established (in NE Washington) before the Paleozoic. However, they did not dismiss the possibility of a Cretaceous accretion event, or that the area was part of the Cordillera terrane. Other “Outliers” Shown in Figure 3 are other localities in Washington and Oregon, whose lateral extents are unknown or nonexistent. The Suplee, Oregon, locality (Johnson and Klapper, 1978) has the oldest sedimentary rocks in that state (with a small possibility of the Klamath Mountains’ Paleozoic belt encroaching on SW
Oregon). Very little has been said about the nature of the carbonate outcrops near Suplee, although it appears that the two outcrops have very divergent ages. There is a northern outcrop in Crook County, which contains a very small brachiopod and coral faunule (see below). The southern outcrop, in Harney County and once considered to be Devonian, appears to be Triassic(?) according to Johnson and Klapper (1978). The northern locality is not coeval with rocks known from Limestone Hill: corals indicate a Givetian age (Sorauf, 1972), and brachiopods indicate a “Middle Devonian” age (Johnson and Klapper, 1978). Danner (1977) reported on small limestone and marble outcrops and quarries in western Washington and adjacent Canada. He described Devonian occurrences within the Eastern and Western Paleozoic belts of NW Washington, and the San Juan Islands, respectively. The Eastern Paleozoic “non-Tethyan” outcrops in NW Washington (excluding the San Juan Islands) are assigned to the Sumas Mountain Subgroup (Devonian) of the Chilliwack Group (Devonian, Early Pennsylvanian, and Permian). Danner (1977, p. 487–488) described several lithologies for the poorly exposed and structurally deformed Devonian rocks, including
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“argillaceous micrites, to layered skeletal reefs and reefs breccias. …[and] shale, greywacke, argillite, minor amounts of chert and a thick section of volcanic rocks.” In the Western Paleozoic “Tethyan” outcrops (principally the San Juan islands) are Devonian rocks (“of most likely Late Devonian age”) with sandstones, turbidites, graywackes, conglomerates, limestones, and rare bedded cherts. Paleontology of the two belts is given below. PALEONTOLOGY Limestone Hill Girty, in Park and Cannon (1943, p. 22) identified three coral taxa from Limestone Hill: Favosites sp., Favosites limitaria, Alveolites sp.?, and a zaphrentid coral. Enbysk (1954, 1956) recovered several coral taxa from the Middle Devonian portion of Limestone Hill, including Aulacophyllum princeps, Favosites limitaris, Hexagonaria cf. H. alternata, Coenites palmata, Syringopora sp., zaphrentid corals, and stromatoporoids. Her work was the first to offer a detailed list of taxa from the locality. Following her work, Sorauf (1972) recovered “Middle
Devonian” corals, presumably from Unit D. Sorauf concluded that his corals indicated an Eifelian age, though a Givetian age was possible. He further described three new coral taxa, Peneckiella metalinae, Phillipsastrea enbyskae, and Synaptophyllum occidentalis. Sorauf (1972, p. 428) further stated, “The varied fauna of northeastern Washington contains species similar to elements of faunas from Australia, Europe, and eastern North America” (i.e., the Onondaga Limestone). A Cordilleran faunal affinity was not discussed. Greenman et al. (1977, p. 477), described conodont collections from Limestone Hill. They recovered Emsian, mid–late Eifelian, Givetian, Frasnian (“probably early”), and Frasnian (gigas Zone) conodonts. The gigas Zone has the following conodont taxa: Polygnathus decorosus, Polygnathus webbi, Polygnathus normaliz, Ancyrodella lobata, Ancyrodella triangularis, Palmatolepis delicatula delicatula, Palmatolepis foliacea, Palmatolepis gigas, Polygnathus normalis, Polygnathus unicornis, and Polygnathus brevis. Frasnian Brachiopods The brachiopod faunule at Limestone Hill consists of Emanuella sp., “Allanella” engelmanni (Meek), Cyrtina sp., Thomasaria sp., and an athyridid (Fig. 4). “Allanella” engelmanni (Meek), derived from Spirifer engelmanni (Meek) by Beus (1965), is ubiquitous in the western United States. It is found in Idaho (Beus,1965), Montana (Laird, 1947), Utah (Kindle, 1908; Williams, 1948), and Nevada (Nolan et al., 1956). In Nevada, “A.” engelmanni occurs in the Frasnian part of the Devils Gate Limestone, within the “Spirifer” argentarius Zone (Nolan et al., 1956). Emanuella is a cosmopolitan genus (e.g., Veevers, 1959). Without a significant number of species identifications, the biogeographic affinities of the brachiopods is tentative. They are, however, suggestive of coeval faunas in Idaho, Montana, and Utah (Upper Jefferson and Bierdneau Formations), and Nevada (Upper Denay and Devils Gate Formations). Their generally small sizes and thin shells suggest a deeper-water setting on muddy substrate, whereas shallower assemblages (Johnson and Trojan, 1982) have a larger representation of alate spiriferids, including Tecnocyrtina, which also appears to occur in the Frasnian part of the Jefferson Formation in Idaho (Isaacson et al., 1989). Paleontology of Other “Outliers”
Figure 3. Other Devonian “outliers.” Suplee, Oregon (Middle Devonian) is described by Johnson and Klapper (1978). Whatcom and Skagit counties have “non-Tethyan” Upper Devonian rocks (Danner, 1977), whereas the San Juan Islands (Maple Bay and Orcas Island) have “Tethyan” affinity Upper Devonian rocks.
For the Suplee, Oregon (Fig. 1) locality, Johnson and Klapper (1978) found a very low diversity Middle Devonian fauna, including the brachiopods Warrenella sp., Spinatrypa sp., Desatrypa? sp., Schizophoria sp., and Gypidula? sp. Conodonts were also recovered. This assemblage clearly is not coeval with the Limestone Hill faunule described herein. Danner’s (1977) Eastern, “non-Tethyan” belt in NW Washington (Whatcom County) has the brachiopod Spinatrypa sp., stromatoporoids (including Amphipora), algae, corals, and protozoa. Lecompte and McLaren (in Danner, 1977, p. 488) suggested a Givetian or more probably Frasnian age for this fauna, with a
Figure 4. All specimens ×2. (1–7) Emanuella sp. (8–12) “Allanella” engelmanni. (13–20) Indet. athyridids. (21) Cyrtina sp. (22–25) Thomasaria sp.
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coral, Scolipora (Plagiopora), being a species Lecompte did not know and that was “different” from coeval species in Belgium. The implication is this fauna is allochthonous. The Western Paleozoic “Tethyan” belt in the San Juan Islands (Washington) has stromatoporoids (Amphipora), gastropods, and brachiopods, including Stringocephalus sp. (a Givetian genus), Productella? sp., and Pseudogruenwaldtia sp. The latter taxon, described by Paul Copper (in Danner, 1977, p. 494), is “known only from the Frasnian,” and is “not comparable to any described European or North American forms,” thereby demonstrating that Devonian rocks in these islands are allochthonous. Pseudogruenwaldtia has been confused with Atrypa reticularis, a common North American taxon. It therefore appears that at least two realms of brachiopods are found in Washington state. The first, containing Limestone Hill and possibly the NW Washington localities, appears to be North American. The San Juan Islands appear to contain brachiopods of another, non–North American affinity. Collections from all these localities, however, are rather lean and conclusions made from them remain tentative. LATE DEVONIAN (FAMENNIAN) BIOGEOGRAPHY Remarkably little has been written on latest Devonian (Famennian) paleobiogeography. This is partly due to a paucity not only of shelly fossils of this age, but also of any investigative literature. Vavrdová and Isaacson (1999) demonstrated a curious paleo-phytogeographic provincialism in marine acritarchs, relative to the ubiquitous land plant spore Retispora lepidophyta. The distribution patterns of fossil marine microplankton cysts (acritarchs), fossil dinoflagellate cysts (dinocysts), and fossil phycomata of green prasinophycean algae represent an important tool for paleogeographic and paleoclimatic reconstructions. Many paleogeographic reconstructions are based on paleomagnetic data, and the potential of the non-megafaunal fossil record is still underestimated. In combination with other sedimentary data, such as development and occurrences of organic buildups, karst, glacial deposits, and temperature-sensitive calcareous algae (Roux, 1991), fossil marine microplankton can be utilized for determination of global paleoclimatological variations and differentiation. This appears to be applicable to the Late Devonian. Although the biological affinities of acid-resistant organicwalled microfossils attributed to the group Acritarcha are not yet well established (Servais et al., 1997), it appears that the majority of acritarchs represent abandoned cysts of unicellular marine green algae (Colbath and Grenfell, 1995; Vavrdová, 1996). The chemical composition of the polymeric and resistant wall, their overall morphology, and a consistent presence of regular vesicle opening (aperture) are the most compelling arguments for their identification with resting stages of phytoplanktonic algae. Factors controlling the distribution of marine fossil microplankton are probably identical to those influencing the distribution pattern of phytoplankton in modern oceans.
Spatial distribution of recent oceanic microplankton depends on the mean annual temperatures of oceanic water masses, influenced by major currents and position of continental blocks. Major modern microplankton communities clearly show latitudinal control. Eight main planktonic bioprovinces in the Pacific Ocean (Bromwell, 1977) form narrow, well-defined belts that parallel latitudes. The subarctic and subantarctic planktonic communities are confined to 50° and 60° latitude in each hemisphere. The relation between acritarch biofacies and paleolatitudinal position was originally proposed by Cramer (1971) and Cramer and Díez (1972) for Silurian rocks. Cramer’s latitudinal model has been challenged (Tappan, 1980; Fortey and Mellish, 1992) and modified and complemented by many subsequent investigations (Colbath, 1990; Le Herissé and Gourvernnec, 1995). Provincialism in acritarch distributions has been demonstrated in the Ordovician (Li Jun, 1987; Tongiorgi et al., 1995) and suggested in the Early Cambrian (Fatka and Vavrdová, 1998). Early Ordovician acritarch bioprovinces appear to reflect cold peri-Gondwanan, temperate Baltic, and warm AustraloLaurentian Provinces (Tongiorgi et al., 1995). The bioprovince ranges from SE Portugal across Saharan Africa (western Libya, Algeria, Morocco) to the Gulf of Guinea, Brazil, Bolivia, Peru, and eastern North America. Marginal regions not included within the bioprovince are characterized by the occurrence of U. deflandrei (Moreau-Benoit) Jardine et al. (sensu lato) without U. saharicum. These regions include eastern Libya, Argentina, Michigan, Tennessee, Belgium, and Iran. At present, data are insufficient to decide whether the absence of U. saharicum is caused by paleoclimatological or by other factors. U. saharicum and other representatives of the genus were not reported from the low-latitude, warm to tropical “belt” ranging from western Canada (Alberta, Great Slave Lake Region, Saskatchewan) to Poland, Siberia, and Australia (Nautiyal, 1977; Playford, 1993). Isaacson et al. (1999) have offered evidence for at least a Famennian glaciation, and it provided sea-level fluctuations that made seaway connections ephemeral. Land plants, on the other hand, had abundant space for propagation during sealevel drawdowns. CONCLUSIONS There are several outliers of Devonian age in Washington and Oregon, USA. Some have suggested that these outliers (especially those in NW Washington and the U.S. San Juan Islands) represent allochthonous terranes of unknown affinities. The Suplee, Oregon locality is Middle Devonian. New information provided by brachiopods from Limestone Hill, NE Washington, suggests that the locality was parautochthonous. This locality’s Middle Devonian coral fauna has suggested a non-Laurentian affinity, although its Frasnian (Late Devonian) brachiopod faunule is very suggestive of coeval faunas in Nevada. Newer
Devonian brachiopods from Northeastern Washington information from crustal studies further suggests that Limestone Hill is on Proterozoic Laurentian crust, with the continental margin (as indicated by the Sr isotope line) located some distance to the west of the locality. This study recommends updates on Frasnian brachiopod faunas, as well as on their biogeographic affinities. This recommendation is supported by data from the Famennian (post-Frasnian) time. Selected marine acritarchs (e.g., Umbellaspaeridium saharicum) show provincialism, centered around western Gondwana and the eastern United States. Coeval miospores of Retispora lepidophyta, conversely are ubiquitous. This corresponds to extreme sea-level fluctuations caused by Late Devonian glaciation, which may have also occurred through at least part of Frasnian time. ACKNOWLEDGMENTS Linda Jo Ellis, a student in paleontology in 1978, discovered the silicified brachiopod fauna on a paleontology field trip. Subsequent visits to Limestone Hill have not discovered any more brachiopods. J.G. Johnson kindly identified the brachiopods. Milada Vavrdová has provided significant information on Late Devonian phytogeography, based on her work in North Africa, the eastern United States, and western Gondwana. D. Geist provided information on crustal studies of NE Washington. A. Boucot and R. Blodgett reviewed the manuscript and provided information for its improvement. REFERENCES CITED Armstrong, R.L., Taubeneck, W.H., and Hales, P.O., 1977, Rb-Sr and K-Ar geochronology of Mesozoic granitic rocks and their Sr isotopic composition: Geological Society of America Bulletin, v. 88, p. 397–411, doi: 10. 1130/0016-7606(1977)88<397:RAKGOM>2.0.CO;2. Beus, S.S., 1965, Devonian faunule from the Jefferson Formation, central Blue Spring Hills, Utah-Idaho: Journal of Paleontology, v. 39, no. 1, p. 21–30. Bromwell, M., ed., 1977, Atlas of the oceans: New York, Crescent Books, p. 1–209. Colbath, G.K., 1990, Palaeobiogeography of Middle Palaeozoic organic-walled phytoplankton, in McKerrow, W.S. and Scotese, C.R., eds., Palaeozoic palaeogeography and biogeography: Geological Society [London] Memoir 12, p. 207–213. Colbath, G.K., and Grenfell, H.R., 1995, Review of biological affinities of Paleozoic acid-resistant, organic-walled eukaryotic algal microfossils (including acritarchs): Review of Palaeobotany and Palynology, v. 86, p. 287–314, doi: 10.1016/0034-6667(94)00148-D. Cramer, F.H., 1971, A palynostratigraphic model of Pangaea during Silurian time—Colloque Ordovicien-Silurien, Brest: Mémoires du Bureau de Recherches Géologiques et Minières, v. 73, p. 229–235. Cramer, F.H., and Díez, M.C.R., 1972, North American Silurian palynofacies and their spatial arrangement: Acritarchs: Palaeontographica, Abt. B, v. 138, p. 107–180. Danner, W.R., 1977, Paleozoic rocks of northwest Washington and adjacent parts of British Columbia, in Stewart, J.H., Stevens, C.H., and Fritsche, A.E., eds, Paleozoic paleogeography of the western United States: Los Angeles, Pacific Section, SEPM, Pacific Coast Paleogeography Symposium 1, p. 481–502. Dings, M.G., and Whitebread, D.H., 1965, Geology and ore deposits of the Metaline Zinc-Lead District, Pend Oreille County, Washington: U.S. Geological Survey Professional Paper 489, 109 p. Enbysk, B., 1954, Geology of Limestone Hill, Pend Orielle County, Washington [M.S. thesis]: Pullman, Washington State University, 120 p. Enbysk, B., 1956, Additions to the Devonian and Carboniferous faunas of
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and palaeogeographic significance: Review of Palaeobotany and Palynology, v. 86, p. 13–48, doi: 10.1016/0034-6667(94)00097-4. Vavrdová, M., 1996, Excystment mechanism and the affinities of acritarchs: Acta Universitatis Carolinae Geologica, v. 3–4, p. 361–383. Vavrdová, M., and Isaacson, P.E., 1999, Late Famennian phytogeographic provincialism: Evidence for a limited separation of Gondwana and Laurentia, in Feist, R., Talent, J.A., and Daurer, A., eds., North Gondwana: MidPaleozoic terranes, stratigraphy and biota: Jahrbuch der Geologischen Bundesantstalt (Wien), v. 54, p. 453–463. Veevers, J.J., 1959, The type species of Productella, Emanuella, Crurithyris, and Ambocoelia (Brachiopoda): Journal of Paleontology, v. 33, no. 5, p. 902–908.
Whitehouse, M.J., Stacey, J.S., and Miller, F.K., 1992, Age and nature of the basement in northeastern Washington and northern Idaho: Isotopic evidence from Mesozoic and Cenozoic granitoids: The Journal of Geology, v. 100, p. 691–701. Williams, J.S., 1948, Geology of the Paleozoic rocks, Logan Quadrangle, Utah: Geological Society of America Bulletin, v. 59, no. 11, p. 1121–1163. Yates, R.G., 1970, Geologic background of the Metaline and Northport mining districts, Washington, in Weissenborn, A.E., ed., Lead-zinc deposits in the Kootenay Arc, northeastern Washington and adjacent British Columbia: Washington Division of Mines and Geology Bulletin 61, p. 17–39. MANUSCRIPT ACCEPTED BY THE SOCIETY 14 DECEMBER 2007
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The Geological Society of America Special Paper 442 2008
Paleobiogeographic affinities of Emsian (late Early Devonian) gastropods from Farewell terrane (west-central Alaska) Jiří Frýda Czech Geological Survey, Klárov 3, 118 21 Praha 1, Czech Republic Robert B. Blodgett* U.S. Geological Survey–Contractor, 4200 University Drive, Anchorage, Alaska 99508, USA
ABSTRACT The vast majority of Emsian gastropods from Limestone Mountain, Medfra B-4 quadrangle, west-central Alaska (Farewell terrane) belong to species with lecithotrophic larval strategy. The present data show that there is no significant difference in the paleobiogeographic distribution of Emsian gastropod genera with lecithotrophic and planktotrophic larval strategies. Numerical analysis of the faunal affinities of the Emsian gastropod fauna from the Farewell terrane reveals that this terrane has much stronger faunal connections to regions like Variscan Europe, eastern Australia, and the Alexander terrane of southeast Alaska than to cratonic North America (Laurentia). The Canadian Arctic Islands is the only region of cratonic North America (Laurentia) that shows significant faunal affinities to the Emsian gastropod faunas of the Farewell terrane. The analysis also indicates a close faunal link between the Farewell and Alexander terranes. Published paleontological and geological data suggest that the Farewell and Alexander terranes represents tectonic entities that have been rifted away from the Siberia, Baltica, or the paleo-Pacific margin of Australia. The results of the present numerical analysis are not in conflict with any of these possibilities. However, the principle of spatial continuity of the wandering path prefers Siberia as the most probable “parental” paleocontinent for the derivation of both the Farewell and Alexander terranes. Keywords: Gastropoda, Devonian, Farewell terrane, Alaska, paleobiogeography. INTRODUCTION A locality discovered by R.B. Blodgett (his field number 83RB9) on the south flank of Limestone Mountain (SW¼, NE¼, sec. 26, T26S, R23E, Medfra B-4 quadrangle, west-central Alaska, 63°16′01″N, 154°32′44″W) provided a highly diverse
gastropod fauna; it is the most diverse Emsian gastropod fauna known from North America. Analyses of its taxonomic composition (Blodgett et al. 1988; Blodgett and Rohr 1989; Frýda and Blodgett 1998, 2001a, 2004) allow close evaluation of the larval strategies of these gastropods as well as their paleobiogeographic significance. Paleobiogeographic studies have proven to be of
*corresponding author:
[email protected] Frýda, J., and Blodgett, R.B., 2008, Paleobiogeographic affinities of Emsian (late Early Devonian) gastropods from Farewell terrane (west-central Alaska), in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 107–120, doi: 10.1130/2008.442(07). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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utmost utility in reconstructing the paleopositions of many of Alaska’s accreted terranes, providing much more consistently reliable data than paleomagnetism; in many examples from Paleozoic strata from Alaska studied with paleomagnetism, the paleopositions have been drastically overprinted (often several times). The primary aim of the present study is to analyze the paleobiogeographic significance of this highly diverse gastropod fauna from the Emsian strata on the south flank of Limestone Mountain (west-central Alaska, Farewell terrane), as well as to discuss the results of the analysis in the context of previously published models on the paleogeographic position of the Farewell terrane. GEOLOGICAL SETTING The described gastropod taxa are from a highly diverse, gastropod-dominated fauna recovered from a 1.5-m-thick, silicified, richly fossiliferous rubble zone in an unnamed Devonian limestone unit. The locality (field number 83RB9) occurs on the south flank of Limestone Mountain, SW¼, NE¼, sec. 6, T26S, R23E, Medfra B-4 quadrangle, west-central Alaska, 63°16′01″N, 154°32′44″W (see Fig. 1 for location). This silicified horizon occurs 13.7–15.2 m above the base of a carbonate sequence that is at least 100 m thick. About 40 species of gastropods, most of which are relatively small sized, dominate the fauna. Other molluscs are present but less common, including bivalves, scaphopods, orthoconic cephalopods, tentaculitids, and small, indeterminate juvenile ammonoids. Other accessory faunal elements include brachiopods, solitary rugose and tabulate corals, lamellar stromatoporoids, and ostracodes. The overall aspect of the fauna
A
is indicative of an undifferentiated Emsian (late Early Devonian) age. On the basis of regional stratigraphic correlations and faunal composition, the host beds appear to be slightly younger than the earliest Emsian (Polygnathus dehiscens Zone) Soda Creek Limestone (Blodgett et al., 2000), the type section of which is situated to the northeast in the Medfra B-3 quadrangle. The dominance of the fauna by a highly diverse assemblage of gastropods, together with the abundance of the tubular stromatoporoid genus Amphipora in underlying and overlying strata, indicates a relatively shallow-water (shallow part of the photic zone), inner carbonate platform environment. The locality is situated within the Nixon Fork terrane of Patton (1978), which was recognized by Decker et al. (1994) to be genetically related to the adjacent Mystic and Dillinger terranes. All of these terranes were reduced in rank to subterranes of a larger terrane, termed by Decker et al. (1994) as the Farewell terrane (see Fig. 1A for location). DISPERSAL POTENTIAL OF GASTROPODS Gastropods, like all Mollusca, have a biphasic life cycle (i.e., larval and post-metamorphosis stages), and this feature is shared with closely related phyla (e.g., Kamptozoa, Sipunculida, Polychaeta). Like other molluscan groups, the embryonic development is characterized by spiral cleavage, which differs slightly in the main gastropod groups (e.g., Biggelaar and Haszprunar, 1996). The subsequent larval stage is called the trochophore larva, and a similar larval type is developed in all molluscan groups. The trochophore larvae may be free swimming, as in the ancient gastropod groups (Patellogastropoda and Archaeogastropoda), or may
B
Figure 1. (A) Location of the fossil locality (LM) in the Farewell terrane of west-central Alaska. (B) Exact location of the fossil locality (field number 83RB9) on the south flank of Limestone Mountain, SW¼, NE¼, sec. 6, T26S, R23E, Medfra B-4 quadrangle, west-central Alaska. Numbered sections are one mile (1.61 km) square.
Paleobiogeographic affinities of Emsian gastropods occur in egg capsules, as in more advanced gastropods. The last larval stage is termed the veliger, which typically bears two ciliate paddles (velum), sometimes subdivided into several lobes. If free-swimming gastropod larvae use planktic organisms for their nutrition, their development is termed planktotrophic. Marine gastropods with such development have small eggs, but numbering over half a million. Planktotrophic larvae may stay planktic for several months and thus can be carried for long distances by oceanic currents. Thus, even currents of only 0.5 km/h can transport these larvae during three months over 1000 km before they settle. On the other hand, it has been suggested that <1% of planktotrophic larvae survive to metamorphosis (Thorson, 1950, 1966). The gastropods, however, developed another ontogenetic strategy in which their larvae were not dependent on an external food source, but on the yolk of their eggs. Gastropods with such a lecithotrophic (non-planktotrophic) development typically produce fewer eggs, which are relatively large. Most of the freeswimming lecithotrophic larvae remain in the plankton for a few minutes, hours, or days (Scheltema, 1978). Such ontogenetic strategy considerably decreases their dispersal potential. Good preservation of gastropod protoconchs, such as in case of the Emsian fauna from Limestone Mountain, can be used to infer
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modes of development. The size of the embryonic shell (protoconch I) is a key parameter for distinguishing planktotrophic and lecithotrophic larval modes. Small embryonic shells (~<100 μm) are typical for larvae with planktotrophic development (Shuto, 1974; Bandel, 1982; Scheltema, 1977, 1978, 1979; Jablonski and Lutz, 1983). The morphology of the protoconch II can also be used for distinguishing planktotrophic and lecithotrophic modes. For example, the presence of a distinct sinusigerous lip on the larval aperture is good evidence for planktotrophy. EMSIAN GASTROPODS FROM LIMESTONE MOUNTAIN AND THEIR LARVAL STRATEGIES The Emsian locality (83RB9) on the south flank of Limestone Mountain (Fig. 1) provided a highly diverse gastropod fauna. It is the most diversified Emsian gastropod fauna known from North America. Analyses of its taxonomic composition (Blodgett et al., 1988; Blodgett and Rohr, 1989; Blodgett and Johnson, 1992; Frýda and Blodgett, 1998, 2001a, 2004) offers the possibility of closely evaluating its paleobiogeographic significance. The gastropod fauna from Limestone Mountain includes members of 24 gastropod genera (Table 1) belonging to three of the five mod-
TABLE 1. LIST OF GASTROPOD GENERA AND SUBGENERA FROM LIMESTONE MOUNTAIN SHOWING THEIR DISPERSAL POTENTIAL AND OCCURRENCES OUTSIDE THE FAREWELL TERRANE Genus Dispersal Occurrence outside the Source potential Farewell terrane Bellerophon (Bellerophon) ? Yes 1 Tropidodiscus
?
Yes
1
Balbinipleura
Low
Yes
6
Quadricarina (Quadricarina)
Low
Yes
6
Mourlonia
Low
Yes
1
Arctozone
Low
Yes
6
Murchisonia (Hormotoma)
Low
Yes
1
Farewellia
Low
Yes
6
Alaskacirrus
Low
Yes
1, 3, 6
Alaskiella
Low
No
1, 3, 6
Chlupacispira
Low
Yes
2
Decorospira
Low
Yes
6
Euomphalus?
Low
Yes
1
Serpulospira
Low
Yes
1
Straparollus (Eleutherospira)
Low
Yes
3
Euomphalopterus
?
Yes
1
Oriostoma
?
Yes
1
Beraunia
?
Yes
1
Naticopsis (Naticopsis)
?
Yes
1
Medfrazyga
Low
No
6
Palaeozygopleura (Rhenozyga)
Low
Yes
6
Stylonema
Low
Yes
1
Nanochilina
High
Yes
6
Kuskokwimia
High
No
5
Note: Sources: 1—Blodgett et al., 1988; 2—Blodgett and Rohr, 1989; 3—Blodgett and Johnson, 1992; 4—Frýda and Blodgett, 1998; 5—Frýda and Blodgett, 2001; 6—Frýda and Blodgett, 2004.
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ern gastropod orders (Archaeogastropoda, Caenogastropoda, and Heterobranchia), as well as to two extinct groups (Bellerophontoidea and Euomphaloidea). Ten genera belong to the Archaeogastropoda (Figs. 2A–2E, 2H, 2J, and 2L), a high-level taxon with a simple shell ontogeny. The general morphology and size of their protoconchs, which are always formed by only an embryonic shell, has not changed during the last 400 m.y. (Frýda, 1999a). Because the Devonian archaeogastropods have protoconchs consisting of the same type of large embryonic shell as in living archaeogastropods, we can assume that they also had the same type of early ontogeny and similar dispersal potential as modern archaeogastropods. Three genera (Medfrazyga, Palaeozygopleura (Rhenozyga) and Stylonema) belong to the Loxonematoidea Koken, 1889. This superfamily represents a very common group of Paleozoic gastropods whose higher taxonomic position has recently been discussed by several authors (Bandel, 1991; Nützel, 1997; Frýda and Bandel, 1997; Frýda, 1999a, 1999b; Frýda et al., 2008). The different protoconch morphologies found in middle and late Paleozoic loxonematoideans indicate that this superfamily (sensu Wenz, 1938, and Knight et al., 1960) is a polyphyletic group uniting members of the Caenogastropoda, Heterobranchia, and Archaeogastropoda (Bandel, 1991; Nützel, 1997; Frýda and Bandel, 1997; Frýda, 1999a, 2000). Frýda and Bandel (1997) reported large protoconchs (with the diameter of the first whorl ranging from 0.3 to 0.5 mm) in some Early Devonian members of the families Loxonematidae (Katoptychia and Stylonema) and Palaeozygopleuridae (Palaeozygopleura). These protoconchs have less than one whorl and were interpreted as a non-planktotrophic protoconch I consisting of only an embryonic shell. Nützel (1997) suggested that the non-planktotrophic nature of the Palaeozygopleuridae might be the result of their living in a deeper-water environment. However, the problem with this explanation is that these families were found in different limestone facies deposited at variable depths including the shallow-water reef environment (Frýda, 1999b). Fortunately, the excellent preservation of palaeozygopleurid gastropods from the Farewell terrane enables a study of their protoconchs. The diameter of the first whorl in Palaeozygopleura (Rhenozyga) reifenstuhli Frýda and Blodgett, 2004, is slightly less than 0.3 mm (Fig. 9.7 in Frýda and Blodgett 2004; Fig. 2F and 2G). The size of the embryonic shell in Medfrazyga clauticae Frýda and Blodgett, 2004, is even larger, having a diameter ~0.35 mm (Figs. 2M and 2N). Thus, both palaeozygopleurid species most probably had lecithotrophic larval development. The genera Farewellia (Fig. 2O) and Murchisonia (Hormotoma) are members of the superfamily Murchisonioidea Koken, 1896, which has been considered to belong to the Archaeogastropoda (e.g., Knight et al., 1960) or to the Caenogastropoda (Ponder and Warén, 1988). Recent discovery of well-preserved protoconchs in Early and Middle Devonian murchisonoideans, including the type genus Murchisonia, provides evidence that these gastropods belong to the Archaeogastropoda (see Frýda and Rohr, 2004, for references). Even though well-preserved
protoconchs have not yet been found in any species of Murchisonia (Hormotoma) and Farewellia, it is very probable that they are similar to those in all other Devonian murchisonoideans. Thus, a lecithotrophic larval strategy in species of Farewellia and Murchisonia (Hormotoma) seems likely. The unusual change of shell coiling (from dextral to sinistral) in the early shells of Alaskacirrus bandeli Frýda and Blodgett, 1998, and Alaskiella medfraensis Frýda and Blodgett, 1998 (Figs. 2C, 2D, 2J, and 2L), places both genera into the extinct superfamily Porcellioidea. Archaeogastropod affinities of this gastropod group (archaeogastropod-type protoconch as well as the presence of nacre) are well documented (Frýda, 1993; Yoo, 1994; Bandel, 1993; Frýda and Blodgett, 1998, 2001a; Kiel and Frýda, 2004). Frýda and Blodgett (1998) documented the large size of the first whorl in both Alaskacirrus and Alaskiella. Thus, Alaskacirrus and Alaskiella also had a lecithotrophic larval strategy. The genus Nanochilina was placed by Frýda (1998) into the Subulitidae because of its teleoconch features. Recent discoveries of the protoconch morphology in some middle and late Paleozoic “subulitoideans” (Frýda and Bandel, 1997; Frýda and Manda, 1997; Nützel et al., 2000; Frýda, 2001; Bandel, 2002), as well as that in Nanochilina gubanovi Frýda and Blodgett, 2004, support this placement. Emsian Nanochilina gubanovi (Fig. 2I) has orthostrophic early whorls with an elevated spire and a diameter of the first preserved whorl at ~0.1 mm. In these features Nanochilina resembles some members of the Soleniscidae (Nützel et al., 2000; Frýda, 2001; Bandel, 2002). The small size of the first whorl suggests that Nanochilina gubanovi had planktotrophic development. The genus Kuskokwimia Frýda and Blodgett, 2001, based on Kuskokwimia moorei Frýda and Blodgett, 2001 (Fig. 2K), is the only gastropod from Limestone Mountain belonging to the order Heterobranchia. The occurrence of this gastropod in Emsian (upper Lower Devonian) strata represents the oldest known record of the Heterobranchia (Frýda and Blodgett, 2001a). The small size of the first preserved protoconch whorl (with a diameter ~0.1 mm; Frýda and Blodgett, 2001a, Fig. 2K) testifies to its planktotrophic larval strategy. Data on the early shell development in bellerophontoidean molluscs are still rather confused. These molluscs have been interpreted as untorted, exogastrically oriented monoplacophorans, or torted, endogastrically oriented archaeogastropods, or a polyphyletic combination of both. Limited knowledge of the soft-part morphology of these molluscs is based mainly on interpretations of such shell characters as their muscle scar pattern, the presence of a dorsal slit, and the shape of their aperture. Recent data on their protoconchs suggest that the bellerophontiform molluscs unite several groups with differing early shell ontogenies. A sharp peak in the size-frequency distribution for species of Bellerophon and Kokenospira, each with about three whorls, has been interpreted as an effect of an increase in mortality during their change from a larval into a benthic mode of life (Dzik, 1978; Frýda, 1999a). Frýda (1999a) showed that the diameter of the embryonic shell in Silurian Bellerophon scaber
Figure 2. Emsian (late Early Devonian) gastropods from the south flank of Limestone Mountain, Medfra B-4 quadrangle, west-central Alaska. (A) Decorospira lepaini Frýda and Blodgett, 2004, ×14. (B) Balbinipleura krawczynskii Frýda and Blodgett, 2004, ×21. (C, L) Alaskiella medfraensis Frýda and Blodgett, 1998, ×18. (D, J) Alaskacirrus bandeli Frýda and Blodgett, 1998; D ×36, J ×21. (E) Arctozone cooki Frýda and Blodgett, 2004, ×17. (F, G) Palaeozygopleura (Rhenozyga) reifenstuhli Frýda and Blodgett, 2004, ×21. (H) Quadricarina (Quadricarina?) noklebergi Frýda and Blodgett, 2004, ×17. (I) Nanochilina gubanovi Frýda and Blodgett, 2004, ×22. (K) Kuskokwimia moorei Frýda and Blodgett, 2001, ×12. (M, N) Medfrazyga clauticae Frýda and Blodgett, 2004, ×18. (O) Farewellia heidelbergerae Frýda and Blodgett, 2004, ×11.
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is only ~0.05 mm. Thus, the small size of the embryonic shells in Bellerophon indicates that it does not belong to the subclass Archaeogastropoda (Frýda, 1999a; Frýda and Rohr, 2004). In addition, such a small size testifies to its planktotrophic larval strategy (Frýda, 1999a). On the other hand, some Ordovician sinuate bellerophontiform molluscs (e.g., Sinuitopsis and Modestospira) have relatively large (up to 0.4 mm), smooth, symmetrical protoconchs that are formed by only the first half of the whorl (Dzik, 1981). Their lecithotrophic larval strategy is very likely. Thus, bellerophontiform molluscs unite groups with both larval strategies. Two genera of bellerophontiform molluscs, Bellerophon (Bellerophon) and Tropidodiscus, are present among the fossils from Limestone Mountain. Both genera belong to widespread taxa, a fact that might suggest planktotrophic larval strategies. However, in gastropod groups where both planktotrophy and lecithotrophy are developed (Caenogastropoda, Heterobranchia, and Neritimorpha), these strategies often occur among genera within the same family or even among species of the same genus. Thus, without data on protoconch morphologies of the species of Bellerophon (Bellerophon) and Tropidodiscus from the Limestone Mountain, we cannot determine their larval strategies. Three of the 24 Limestone Mountain genera (Euomphalus?, Serpulospira, and Straparollus (Eleutherospira)) belong to the superfamily Euomphaloidea de Koninck, 1881, which is considered to belong to the Archaeogastropoda (Yochelson 1956; Knight et al., 1960). Yoo (1994), and later Bandel and Frýda (1998), found an unusual protoconch morphology in some Devonian and Carboniferous euomphaloidean genera that distinguishes them from members of extant gastropod groups (Patellogastropoda, Archaeogastropoda, Neritimorpha, Caenogastropoda, and Heterobranchia). For this reason, Bandel and Frýda (1998) considered them to represent an independent extinct gastropod group. Nützel (2002) confirmed earlier observations (Yoo, 1994; Bandel and Frýda 1998) on the nature of the protoconch and shape of the boundary between protoconch and teleoconch and noted some affinities among Euomphaloidea, Docoglossa, Cocculiniformia, and Neomphalidae. The protoconch of euomphaloidean gastropods (Euomphalomorpha), formed by a capacious, short tubular embryonic shell (protoconch I; see Yoo, 1994; Bandel and Frýda, 1998; Nützel, 2002; and Frýda et al. 2006), testifies to their lecithotrophic larval strategy. The order-level position of the gastropod genus Euomphalopterus is uncertain (Linsley et al., 1978) as is its protoconch morphology. Thus, no data are available for determining its larval strategy. The same is true for species of Oriostoma, Beraunia, and Naticopsis (Naticopsis). The genus Naticopsis is the type genus for the subfamily Naticopsinae Waagen, 1880, which most probably belongs to the Neritimorpha. Even though the early phylogeny of neritimorph gastropods is still unclear (Bandel and Frýda, 1999; Bandel and Heidelberger, 2001; Frýda and Rohr, 2004), all Paleozoic groups involved in different evolutionary models of the Neritimorpha (including the extinct Cyrtoneritimorpha; see Frýda and Heidelberger, 2003 for discus-
sion and references) had the ability to develop true larval shells (protoconch II). Thus, both planktotrophy and lecithotrophy are possible in species of Naticopsis (Naticopsis) from Limestone Mountain, though lack of data on its protoconch morphology prevents determination of its larval strategy. ANALYSIS OF PALEOBIOGEOGRAPHIC AFFINITIES To evaluate the paleobiogeographic affinities of the gastropod fauna from the Farewell terrane, we summarized data on the distribution of Emsian gastropods from various published and unpublished sources. The Emsian gastropod faunas from many areas of the world are still not described, nor have they not been restudied since their description in the nineteenth century. For North America we have used a data set published by Blodgett et al. (1988) that was complemented by later publications (Blodgett and Rohr, 1989; Blodgett and Johnson, 1992; Frýda and Blodgett, 1998, 2001a, 2004) and unpublished data. In addition to the data from Limestone Mountain (Farewell terrane), we included seven other areas of North America with Emsian gastropod faunas. Among faunas belonging to cratonic North America the following areas were included: east-central Alaska, British Columbia, Nevada, Hudson Bay Lowlands, and the Canadian Arctic Islands. Data on the Emsian gastropod fauna from east-central Alaska, corresponding to locality 1 of Blodgett et al. (1988), come from the triangular area bounded roughly on the northwest by the Porcupine River and on the southwest by the Yukon River. This portion of Alaska has not been tectonically accreted (Coney et al. 1980; Blodgett 1998; Blodgett and Boucot 1999). The examined collections were made over the course of many separate field seasons by R.B. Blodgett from outcrops of the Ogilvie Formation exposed on both side of the Alaska-Yukon border. Data on Emsian gastropod faunas from British Columbia, corresponding to locality 8 of Blodgett et al. (1988), come from examination of a single gastropod collection recovered by D.G. Perry, A.J. Boucot, and H. Gabrielse from the Mount Lloyd George area (see Perry et al., 1981 for additional locality details and the accompanying brachiopod fauna). Data on the Emsian gastropod faunas from Nevada, corresponding to localities 12, 13, and 15 of Blodgett et al. (1988), come from examination of the illustrated gastropod fauna presented by Walcott (1884) and Merriam (1940, 1973), as well as of collections made by A.J. Boucot. Data on Emsian gastropod faunas from the Hudson Bay Lowlands, corresponding to locality 11 of Blodgett et al. (1988), come from examination of the scant gastropod material present in the collections of the Geological Survey of Canada. Data on Emsian gastropod faunas from the Canadian Arctic Islands, corresponding to localities 5 and 6 of Blodgett et al. (1988), come from examination of illustrated gastropods shown in Smith (1984) from southwestern Ellesmere Island (locality 5) and from direct examination of gastropod collections of the Geological Survey of Canada and of Eric C. Prosh from the Disappointment Bay Formation of Lowther Island (locality 6).
Paleobiogeographic affinities of Emsian gastropods The rich and highly diverse Emsian gastropod fauna from the Disappointment Bay Formation on Lowther Island has been newly revised and is being prepared for publication. Data on the Emsian gastropod faunas from the Alexander terrane, corresponding to locality 2 of Blodgett et al. (1988), come from Kasaan Island and are based on examination of collections made separately by Connie Soja and N.M. Savage, as well as the original collections from this horizon listed by E.M. Kindle (1907). Unfortunately our knowledge of Emsian gastropod fauna from areas outside of North America is limited to a few areas. In some large regions, such as different Asian plates (including Siberia), only a few reports exist, and the exact stratigraphical positions of reported gastropod faunas remain uncertain. For these reasons we included in our analysis data only from two regions outside of North America—Variscan Europe and eastern Australia. The data on the Emsian gastropods of both areas come from publications (see for references in Blodgett et al., 1988, 1990, and Frýda and Bandel, 1997) and from an unpublished data set (JF) that is being prepared for publication with a complete taxonomic revision of these faunas. The present study of paleobiogeographic affinities of the gastropod fauna of the Farewell terrane is based on a numerical analysis of generic level data and not on the geographic distribution of Emsian gastropod species. We chose this approach largely because species-level taxonomy of a large part of the Emsian gastropod faunas has not been newly revised. However, because the geographic range of a genus is generally greater than that of any species belonging to the genus, the disadvantage of using generic-level rather than species-level data for paleobiogeographic analysis is the much lower resolution of faunal affinities. Another feature limiting the size of the data set is the fact that only taxa that occur in at least two analyzed areas bear paleobiogeographic signals useful for an evaluation of the paleobiogeographic relationships. Data on endemic taxa reveal other paleobiogeographic information such as the level of isolation of certain area, a feature that may be caused by different biotic and abiotic factors. For these reasons, for our numerical analysis (Table 2) we used only data on gastropod genera that are shared between at least two analyzed regions. Many parameters have been defined for measuring the distance between two samples containing taxon occurrences. We selected Jaccard and Dice (Sorensen) association similarity indices because we are using presence/absence data for the analysis (Table 3). We used cluster analysis for finding hierarchical groupings of the analyzed regions. The latter multivariate statistical method helps us to recognize analyzed areas hierarchically grouped together so as to reflect their paleobiogeographic similarity. From a large number of existing algorithms we chose the unweighted pair-group average (UPGMA) algorithm, in which clusters are joined on the basis of the average distance (similarity) between all members in the two groups (areas). Two Jaccard and Dice association similarity indices were used for distance measure.
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RESULTS The vast majority of the Emsian gastropods from Limestone Mountain (Farewell terrane) belong to species with a lecithotrophic larval strategy (all archaeogastropods including Pleurotomarioidea, Microdomatoidea, Porcellioidea, and Murchisonioidea; Loxonematoidea and Euomphalomorpha). Planktotrophic larval strategy was documented only in two species belonging to the genera Nanochilina (Caenogastropoda) and Kuskokwimia (Heterobranchia). Larval strategies in two bellerophontoidean species belonging to the genera Bellerophon and Tropidodiscus and four species belonging to the genera Euomphalopterus, Oriostoma, Beraunia, and Naticopsis are uncertain (Table 1). Tables 1 and 2 show that there is no significant difference in distribution between taxa with low dispersal potential (lecithotrophic strategy) and those with high dispersal potential (planktotrophic strategy). The size of a given taxon’s geographic range has been influenced not only by its dispersal potential but also by many other biotic and abiotic factors, of which the taxon’s evolutionary history is one of the most important. It should be noted that the great majority of gastropods from the Limestone Mountain fauna belong to genera that have been reported from older strata of different areas of the world (see Frýda and Blodgett, 2004). In summary, even gastropods with a lecithotrophic larval strategy had enough geological time to cover a large geographic range. The Jaccard and Dice association similarity indices (Table 3) reveal that the Emsian gastropod fauna of the Farewell terrane has its closest faunal affinities to Emsian faunas of Eastern Australia and Variscan Europe (Dice similarity coefficients [DSC] equal to 0.56 and 0.50, respectively), and slightly lower faunal affinities with coeval faunas of the Canadian Arctic Islands (DSC = 0.32) and with those of the Alexander terrane (DSC = 0.29). In contrast, similarity indices reveal very low faunal affinity to the Emsian gastropod faunas of the east-central Alaska, British Columbia, Nevada, and Hudson Bay Lowlands (Table 3). The cluster analysis resulted in dendrograms of the same branching topology for both the Jaccard and Dice indices. The only differences exist in the values of similarity and these differences have no significant influence on the evaluation of the paleobiogeographic affinities of the studied fauna or on the grouping of studied areas. Figure 3 shows a dendrogram based on clustering of the Dice distance measure. This measure put more weight on joint occurrences than on the Jaccard coefficient. The branching topology of the dendrogram shows two fairly well separated clusters (Fig. 3). One cluster unites the Emsian gastropod faunas from east-central Alaska, British Columbia, Nevada, and the Hudson Bay Lowlands, all areas that belong to cratonic North America. The second cluster unites the Emsian gastropod fauna from the Farewell terrane with those of the Alexander terrane, Variscan Europe, Eastern Australia, and the Canadian Arctic Islands, the last region being the only one of this cluster belonging to cratonic North America. Thus the Canadian Arctic Islands are the only region of cratonic North America (Laurentia) that has significant faunal affinities to the Emsian gastropod faunas of the Farewell terrane.
East-central Alaska
Alexander terrane
Farewell terrane
British Columbia
Canadian Arctic Islands
Hudson Bay Lowlands
Nevada
Variscan Europe
Eastern Australia
TABLE 2. OCCURRENCES (PRESENCE/ABSENCE DATA) OF THE EMSIAN GASTROPOD GENERA USED IN PRESENT NUMERICAL ANALYSIS OF THEIR FAUNAL AFFINITIES
Alaskacirrus
0
0
1
0
0
0
0
1
0
Angyomphalus (Eoangyomphalus)
0
0
0
0
1
0
0
1
0
Arctozone
0
0
1
0
0
0
0
0
1
Australonema
0
0
0
0
0
0
0
1
1
Balbinipleura
0
0
1
0
0
0
0
1
0
Bellerophon (Bellerophon)
0
1
1
0
1
0
1
1
1
Boiotremus
1
1
0
0
0
0
0
1
0
Blodgettinotus
0
0
0
0
0
0
0
1
1
Brokenriveria
0
0
0
0
1
0
0
0
1
Coelocaulus
0
1
0
0
0
0
0
1
0
Coelocyclus
0
0
0
0
0
0
0
1
1
Cyclonema
0
0
0
1
1
0
0
0
0
Decorospira
0
0
1
0
0
0
0
1
0
Devonicornu
0
0
0
0
1
0
0
1
0
Eurekaphon
0
0
0
0
0
0
1
0
1
Farewellia
0
0
1
0
0
0
0
0
1
Fusispira
0
1
0
0
1
0
0
1
1
Gyronema
0
0
0
0
1
0
0
1
1
Holopea
0
1
0
0
0
0
0
1
1
Chlupacispira
0
0
1
0
0
0
0
0
1
Kodymites
0
0
0
0
0
0
0
1
1
Loxonema
0
0
0
0
1
0
1
1
1
Mourlonia
0
0
1
0
1
0
0
1
0
Murchisonia (Hormotoma)
0
0
1
0
0
0
0
1
0
Murchisonia (Murchisonia)
0
1
1
0
1
0
1
1
1
Nanochilina
0
0
1
0
1
0
0
1
1
Naticopsis (Naticopsis)
0
0
1
0
1
0
0
0
1
Oehlertia
0
1
0
0
1
0
0
1
1
Oriostoma
0
1
1
0
0
0
0
1
0
Palaeozygopleura (Rhenozyga)
0
0
1
0
0
0
0
0
1
Paleuphemites
0
0
0
0
1
0
0
1
0
Platyceras (Orthonychia)
1
0
0
1
1
0
1
1
1
Platyceras (Platyceras)
1
0
0
1
1
1
1
1
1
Platyceras (Platyostoma)
1
0
0
0
1
0
1
0
0
Platyceras (Prosigaretus)
0
0
0
0
1
0
0
0
0
Serpulospira
0
0
1
0
0
1
0
0
1
Stenoloron
0
0
0
0
1
0
0
1
0
Straparollus
0
0
1
0
0
0
1
1
1
Stylonema
0
0
1
0
0
0
0
1
1
Tropidodiscus
0
1
1
0
1
0
0
1
1
Tubina
0
1
0
0
0
0
0
1
0
Genus
Paleobiogeographic affinities of Emsian gastropods
115
0.45
1 Eastern Australia Note: Jaccard indices are given in the upper triangle of the matrix (above and to the right of the diagonal), and Dice indices are given in the lower.
0.62
0.27
0.15
0.34
Variscan Europe
0.14
0.56
0.53
0.14
0.42
1
0.19
0.06
0.32
1 0.20
0.12 0.60 0.50
0.50
0.22
0.18
Nevada
Hudson Bay Lowlands
0.50
0.23
0.43
0.36
0.08
0.08 0.03
0.06 0.22
0.11
0.25
0.09
1
1
0.40
0.26
0.00
0.10
0.00
0.33
Canadian Arctic Islands
British Columbia
0.57
0.00
0.39
0.36 0.43 0.27 0.15 0.33
Farewell terrane
0.25
0.32
1
0.05
0.21 0.33
0.33 0.13 0.05 0.00 0.19 0.29
Alexander terrane
0.00
1
0.33
0.13 0.00 0.00 0.20 1
East-central Alaska
0.14
0.17
Eastern Variscan Nevada
Europe 0.10 Lowlands 0.20
Hudson Bay British
Columbia 0.40
Canadian
Arctic Islands 0.14 terrane 0.00 terrane 0.08 Alaska 1
Farewell Alexander East-central
TABLE 3. JACCARD-DICE SIMILARITY COEFFICIENTS
Australia 0.07
DISCUSSION Reconstruction of the paleogeographic position of any paleocontinent far in the geologic past is difficult. It is even more difficult to determine the paleogeographic position of small crustal segments (microplates or terranes) because they generally provide a smaller amount of biogeographically significant information than larger paleocontinents. In addition, small crustal segments are often completely involved in orogenic processes that might totally destroy fossil paleogeographic indicators during regional metamorphism. Paleogeographic maps have been drawn with much confidence for post-Paleozoic time, because the positions of continental plates can be determined from the magnetic reversal patterns present in ocean-floor deposits. The construction of paleogeographic maps for Paleozoic time, however, is much more difficult, because no extant oceanic crust of this age is present in modern ocean basins, and the paleomagnetic measurements from continental rocks of this age are usually not reliable, owing to thermal or chemical overprinting of the magnetic signatures. For this reason, many differing global reconstructions have been proposed for the Paleozoic. To increase the chance of correctly determining the paleogeographic position of a terrane it is necessary to analyze the affinities of that terrane to other crustal blocks at different time intervals. The wandering path of the terrane should be inferred from a series of temporally different paleopositions. Only when they define a “smooth” wandering path, and when none of the inferred paleopositions contradicts the others, does the “objectivity” of the paleogeographic reconstruction appear to be high. Thus, information on pre-Emsian as well as post-Emsian faunal affinities of the Farewell terrane together with the principle of spatial continuity can be used to verify the results of the present analysis. Pre-Emsian Paleogeographic Position of Farewell Terrane The evaluation of the faunal affinities of the Middle Cambrian trilobite faunas from the Nixon Fork subterrane (Farewell terrane) reveals their Siberian aspect, showing closest similarity to the eastern part of the Siberian Platform (Palmer et al., 1985, and Blodgett et al., 2002; but see also Babcock and Blodgett, 1992, Babcock et al., 1993, and St. John and Babcock, 1997). Upper Ordovician (Ashgillian) strata of the Lone Mountain area (Nixon Fork subterrane) contain rich, diverse brachiopod, gastropod, and coral faunas (Rohr and Blodgett, 1985; Blodgett et al., 1992; Potter and Blodgett, 1992; Rohr et al., 2003). These faunas also suggest a link to Ashgillian faunas from Siberia (Blodgett et al. 2002) as well as to those of the Seward Peninsula (York terrane; Blodgett et al., 1992; Rohr et al., 2003; but see also Frýda et al., 2001). Older Ordovician faunas known from the Nixon Fork subterrane contain taxa of both North American and Siberian affinities (Rohr and Blodgett, 1988; Rohr et al., 1992; Rohr and Gubanov, 1997; Blodgett et al., 2002). Upper Silurian and lower Lower Devonian rocks of the Farewell terrane differ from coeval rocks of cratonic North America (Laurentia) by the presence of
Frýda and Blodgett
0.6
i
Accreted terranes
a
Alexander terrane
t
Farewell terrane
n
Eastern Australia
e
Variscan Europe
0.7
r
Canadian Arctic Islands
SIMILARITY
0.8
u
Nevada
0.9
British Columbia
1
a
East-central Alaska
L
Hudson Bay Lowlands
116
0.5 0.4 0.3 0.2 0.1
Figure 3. Results of a cluster analysis (UPGMA) based on the estimated faunal similarities between Emsian gastropod faunas. Binary similarities estimated using the Dice coefficient.
extensive buildups of algal reef mound complexes (Clough and Blodgett, 1985, 1988, 1992; Blodgett and Gilbert 1992). Similar algal buildups of equivalent age have been reported from the Alexander terrane as well as from the Urals (Soja et al., 2003 and Antoshkina, 1999, and references therein). Thus, the available data suggest strong links between the Farewell terrane and Siberia for pre-Emsian time. Emsian Paleogeographic Position of Farewell Terrane Rich and diverse brachiopod and coral faunas from the Soda Creek Limestone of the Nixon Fork subterrane have a strong affinity to various parts of Siberia (Kolyma, Taimyr, and Kuznetsk Basin) and Arctic Russia (Novaya Zemlya), but lack links to cratonic North America (Laurentia) (Blodgett et al., 1995, 2000, 2002). In addition, late Emsian-age brachiopods of the Shellabarger Pass area of the Mystic subterrane (Farewell terrane) show faunal affinities to the Urals and Kolyma (Blodgett and Brease, 1997; Blodgett, 1998; Blodgett and Boucot, 1999; Garcia-Alcalde and Blodgett, 2001; Blodgett et al., 2002). Like the pre-Emsian faunas, those of Emsian age from the Farewell terrane also suggest strong links to Siberia. Post-Emsian Paleogeographic Position of Farewell Terrane The diverse Eifelian (early Middle Devonian) fauna of the Cheeneetnuk Limestone (Blodgett and Gilbert 1983) of the Nixon
Fork subterrane provides the biogeographically most important Middle Devonian fauna of the Farewell terrane. This fauna has not yet been completely evaluated taxonomically owing to its richness. It includes brachiopods, rugose and tabulate corals, gastropods, bivalves, rostroconchs, nautiloids, goniatites, tentaculitids, trilobites, ostracodes, and sponges (see Blodgett et al., 2002, for references). The goniatite fauna of the Cheeneetnuk Limestone suggests a link to Eurasia (House and Blodgett, 1982). The gastropod fauna of the Cheeneetnuk Limestone has only been partially evaluated (Blodgett and Rohr, 1989; Blodgett, 1992, 1993; Blodgett and Cook, 2002; Blodgett and Johnson, 1992); nevertheless, it differs at the species level from coeval faunas of cratonic North America. At the generic level it shows some links to the Northwest Territories, Nevada, and Australia. The algal flora of the Cheeneetnuk Limestone (Poncet and Blodgett, 1987) differs from that of cratonic North America but indicates a link to the Russian Platform and Variscan Europe (Germany). Mamet and Plafker (1982) noted close faunal relationships between the Late Devonian fauna of the Mystic subterrane (Farewell terrane) and coeval faunas of the Russian Platform and the Urals. Hahn et al. (1985) and Hahn and Hahn (1985) described Pennsylvanian-age trilobites from Farewell terrane (derived from the Nixon Fork and Mystic subterranes) and showed their affinities with Eurasian, rather than North American, taxa. The Permian flora of the Farewell terrane (Mamay and Reed, 1984) indicates links to the southwestern part of the United States as well as to Siberia. Similarly, Sunderlin (2001 and this volume) recollected
Paleobiogeographic affinities of Emsian gastropods the aforementioned Permian flora and noted it to be of a mixed phytogeographic affinity with temperate Angaran and Euramerican floristic Realms. Evolution of Farewell and Alexander Terrane Assemblage The Farewell and Alexander terranes are the only parts of present-day North America with Late Silurian and early Early Devonian buildups formed by of algal reef mounds, which also indicate a Uralian-Siberian link (Clough and Blodgett 1985, 1988, 1992; Blodgett and Gilbert 1992; Soja et al. 2003; and Antoshkina 1999). The Emsian-age Kasaan Island gastropod fauna of the Alexander terrane is very similar to the Emsian gastropod fauna of the Farewell fauna (Blodgett et al. 1988, 2002; Table 3). On the basis of the paleobiogeographic affinities of various early and middle Paleozoic faunas, the Farewell and Alexander terranes are most closely allied with Siberia (notably Kolyma, Taimyr, Kuznetsk Basin), as well as with the Urals. It seems most likely that these terranes represent tectonic elements that have been rifted away from the Siberian continent (Blodgett and Brease, 1997; Blodgett, 1998; Blodgett and Boucot, 1999; Blodgett et al., 2002; Boucot and Blodgett, 2001; GarciaAlcalde and Blodgett, 2001). On the other hand, Bazard et al. (1995) analyzed paleomagnetic data from 29 sites within the Lower Devonian Karheen Formation and suggested that Eastern Australia and the Scandinavian part of Baltica represent possible candidates for derivation of the Alexander terrane. In addition, both paleocontinents are possible sources of Precambrian detrital zircon grains found in the Karheen Formation. However, Bazard et al. (1995) noted that paleontological data from the underlying Heceta Formation are only consistent with a position near northern Laurentia-Baltica and, therefore, they suggested for the Alexander terrane an Early Devonian position near Baltica. Similarly, Gehrels et al. (1996), on the basis of U-Pb analyses of 101 detrital zircon grains from the Paleozoic and Triassic clastic strata of the Alexander terrane, indicated that most detritus in the terrane was derived from intraterrane igneous rocks. They noted that it appears unlikely that the Alexander terrane was in proximity to the Cordilleran margin during Early Devonian time because these Precambrian zircons are not the same age as grains that were accumulating along the western margin of North America. Gehrels et al. (1996) therefore suggested the paleo-Pacific margin of Australia and the Scandinavian portion of Baltica as a possible source for detrital zircon grains from the Paleozoic and Triassic clastic strata of the Alexander terrane. Johnston (2001) published a different geotectonic model, pointing out that Alaska and Yukon Territory are divisible into a series of geological belts, including a northern E-W-trending Arctic Alaska belt, a central SW-trending Ruby belt, a southern E-trending Dillinger belt, and the SE-trending Yukon-Tanana belt. According to Johnston, these belts previously formed part of a linear ribbon continent (named Saybia), which originally extended from eastern Siberia south and was likely 8000 km long.
117
Belasky et al. (2002) suggest that Wrangellia-Alexander, Stikinia, and the Eastern Klamath terranes were at approximately the same distance (~2000–3000 km) from their latitudinally equivalent segments of the North American craton during the Early Permian. On the other hand, data on the biogeographic affinities of the early Mesozoic (Triassic) gastropod faunas of the Alaskan terranes (Chulitna, Farewell, Alexander, and Wrangellia) suggest that the Chulitna-Farewell-Alexander terrane assemblage was separated during Late Triassic time by a reproductively significant distance from the Wrangellia-Wallowa terrane couplet in various tropical areas within the Panthalassa Ocean (Blodgett et al., 2001; Blodgett and Frýda, 2001; Frýda and Blodgett, 2001b). These facts suggest that the wandering path of the Farewell and Alexander terranes from their early Paleozoic positions to their present positions was probably very complex and is still poorly known. CONCLUSIONS The vast majority of Emsian gastropods from Limestone Mountain (Farewell terrane) belong to species with a lecithotrophic larval strategy. The present data show that there is no significant difference in the paleogeographic distribution of Emsian gastropod genera having low (lecithotrophic strategy) and high (planktotrophic strategy) dispersal potential. Numerical analysis of faunal affinities of the Emsian gastropod fauna from the Farewell terrane reveals that this terrane had much stronger faunal connections to regions such as Variscan Europe, Eastern Australia, and the Alexander terrane than to cratonic North America (Laurentia). The Canadian Arctic Islands are the only region of cratonic North America (Laurentia) that has significant faunal affinities to the Emsian gastropod faunas of the Farewell terrane. The present analysis also indicates a close faunal link between the Emsian gastropods faunas of the Farewell terrane with those of the Alexander terrane (Kasaan Island). Previously published models indicate close links of the Farewell and Alexander terranes since the Silurian. Paleontological and geological data suggest that the Farewell terrane represents a continental margin sequence that has been rifted away from the Siberian continent or from the Scandinavian portion of Baltica. Results of our analysis of faunal affinities of the Emsian gastropod fauna from the Farewell terrane are not in conflict with either of these possibilities. Lack of data from Siberia does not allow for a test of the faunal relationships between Siberia and the Farewell and Alexander terranes based on Emsian gastropod faunas. Application of the principle of spatial continuity of the wandering path prefers Siberia over two other possible paleocontinents, Baltica and Australia, as the most probable “parental” paleocontinent for a derivation of the Farewell and Alexander terrane assemblage. ACKNOWLEDGMENTS This work was supported by the Alexander von HumboldtStiftung, grants from the Czech Academy of Science
118
Frýda and Blodgett
(KJB307020602), the Grant Agency of the Czech Republic (205/08/0062), the Czech-American Cooperation Programme (Kontakt ME08011), and by the Committee for Exploration and Research of the National Geographic Society (project no. 6739-00) to Jiří Frýda. We thank the Alaska Division of Geological and Geophysical Surveys for providing helicopter support to Blodgett, which allowed for the collection of the Limestone Mountain fauna. REFERENCES CITED Antoshkina, A.I., 1999, Origin and evolution of lower Paleozoic reefs in the Pechora Urals, Russia: Bulletin of Canadian Petroleum Geology, v. 47, p. 85–103. Babcock, L.E., and Blodgett, R.B., 1992, Biogeographic and paleogeographic significance of Middle Cambrian trilobites of Siberian aspect from southwestern Alaska: Geological Society of America Abstracts with Programs, v. 24, no. 5, p. 4. Babcock, L.E., Blodgett, R.B., and St. John, J., 1993, Proterozoic and Cambrian stratigraphy and paleontology of the Nixon Fork terrane, southwestern Alaska: Proceedings of the First Circum-Pacific and Circum-Atlantic Terrane Conference, Guanajuato, Mexico, 9–14 November 1993, p. 5–7. Bandel, K., 1982, Morphologie und Bildung der fruhontogenetischen Gehause bei conchiferen Mollusken: Facies, v. 7, p. 1–198, doi: 10.1007/ BF02537225. Bandel, K., 1991, Úber triassische ‘‘Loxonematoidea’’ und ihre Beziehungen zu rezenten und palaeozoischen Schnecken: Paläontologische Zeitschrift, v. 65, no. 3–4, p. 239–268. Bandel, K., 1993, Evolutionary history of sinistral archaeogastropods with and without slit (Cirroidea, Vetigastropoda): Freiberger Forschungshefte, v. C450, p. 41–81. Bandel, K., 2002, Reevaluation and classification of Carboniferous and Permian Gastropoda belonging to the Caenogastropoda and their relation: Mitteilungen aus dem Geologisch-Paläontologischen Institut der Universität Hamburg, v. 86, p. 81–188. Bandel, K., and Frýda, J., 1998, Position of Euomphalidae in the system of the Gastropoda: Senckenbergiana Lethaea, v. 78, no. 1–2, p. 103–131. Bandel, K., and Frýda, J., 1999, Notes on the evolution and higher classification of the subclass Neritimorpha (Gastropoda) with the description of some new taxa: Geologica et Palaeontologica, v. 33, p. 219–235. Bandel, K., and Heidelberger, D., 2001, The new family Nerrhenidae (Neritimorpha, Gastropoda) from the Givetian of Germany: Neues Jahrbuch für Geologie und Paläontologie, Monatshefte, v. 12, p. 705–718. Bazard, D.R., Butler, R.F., Gehrels, G., and Soja, C.M., 1995, Early Devonian paleomagnetic data from the Lower Devonian Karheen Formation suggest Laurentia-Baltica connection for the Alexander terrane: Geology, v. 23, p. 707–710, doi: 10.1130/0091-7613(1995)023<0707:EDPDFT> 2.3.CO;2. Belasky, P., Stevens, C.H., and Hanger, R.A., 2002, Early Permian location of western North American terranes based on brachiopod, fusulinid, and coral biogeography: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 179, p. 245–266, doi: 10.1016/S0031-0182(01)00437-0. Biggelaar, J.A.M., and Haszprunar, G., 1996, Cleavage patterns and mesentoblast formation in the Gastropoda: An evolutionary perspective: Evolution (International Journal of Organic Evolution), v. 50, p. 1520–1540. Blodgett, R.B., 1992, Taxonomy and paleobiogeographic affinities of an early Middle Devonian (Eifelian) gastropod faunule from the Livengood quadrangle, east-central Alaska: Palaeontographica, Abt. A, v. 221, p. 125–168. Blodgett, R.B., 1993, Dutrochus, a new microdomatid (Gastropoda) genus from the Middle Devonian (Eifelian) of west-central Alaska: Journal of Paleontology, v. 67, p. 194–197. Blodgett, R.B., 1998, Emsian (late Early Devonian) fossils indicate a Siberian origin for the Farewell terrane, in Clough, J.G., and Larson, F., eds., Short notes on Alaska geology, 1997: Alaska Division of Geological and Geophysical Surveys Professional Report 118, p. 53–61. Blodgett, R.B., and Boucot, A.J., 1999, Late Early Devonian (Late Emsian) eospiriferinid brachiopods from Shellabarger Pass, Talkeetna C-6 Quadrangle, south-central Alaska and their biogeographic importance; further
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The Geological Society of America Special Paper 442 2008
Significance of detrital zircons in Upper Devonian ocean-basin strata of the Sonora allochthon and Lower Permian synorogenic strata of the Mina México foredeep, central Sonora, México Forrest G. Poole U.S. Geological Survey, Box 25046, Federal Center, Denver, Colorado 80225 George E. Gehrels Department of Geosciences, University of Arizona, Tucson, Arizona 85721 John H. Stewart U.S. Geological Survey, 345 Middlefield Road, Menlo Park, California 94025 ABSTRACT U-Pb isotopic dating of detrital zircons from a conglomeratic barite sandstone in the Sonora allochthon and a calciclastic sandstone in the Mina México foredeep of the Minas de Barita area reveals two main age groups in the Upper Devonian part of the Los Pozos Formation, 1.73–1.65 Ga and 1.44–1.42 Ga; and three main age groups in the Lower Permian part of the Mina México Formation, 1.93–1.91 Ga, 1.45–1.42 Ga, and 1.1–1.0 Ga. Small numbers of zircons with ages of 2.72–2.65 Ga, 1.30–1.24 Ga, ca. 2.46 Ga, ca. 1.83 Ga, and ca. 0.53 Ga are also present in the Los Pozos sandstone. Detrital zircons ranging in age from 1.73 to 1.65 Ga are considered to have been derived from the Yavapai, Mojave, and Mazatzal Provinces and their transition zones of the southwestern United States and northwestern Mexico. The 1.45–1.30 Ga detrital zircons were probably derived from scattered granite bodies within the Mojave and Mazatzal basement rocks in the southwestern United States and northwestern Mexico, and possibly from the Southern and Eastern Granite-Rhyolite Provinces of the southern United States. The 1.24–1.0 Ga detrital zircons are believed to have been derived from the Grenville (Llano) Province to the east and northeast or from Grenvilleage intrusions or anatectites to the north. Several detrital zircon ages ranging from 2.72 to 1.91 Ga were probably derived originally from the Archean Wyoming Province and Early Paleoproterozoic rocks of the Lake Superior region. These older detrital zircons most likely have been recycled one or more times into the Paleozoic sandstones of central Sonora. The 0.53 Ga zircon is believed to have been derived from a Lower Cambrian granitoid or metamorphic rock northeast of central Sonora, possibly in New Mexico and Colorado, or Oklahoma. Detrital zircon geochronology suggests that most of the detritus in both samples was derived from Laurentia to the north, whereas some detritus in the Permian synorogenic foredeep sequence was derived from the evolving accretionary wedge to the south. Poole, F.G., Gehrels, G.E., and Stewart, J.H., 2008, Significance of detrital zircons in Upper Devonian ocean-basin strata of the Sonora allochthon and Lower Permian synorogenic strata of the Mina México foredeep, central Sonora, México, in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 121–131, doi: 10.1130/2008.442(08). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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Poole et al. Compositional and sedimentological differences between the continental-rise Los Pozos conglomeratic barite sandstone and the foredeep Mina México calciclastic sandstone imply different depositional and tectonic settings. Keywords: Southwest Laurentia, Paleozoic stratigraphy and tectonics, sediment provenance, geochronology, Sonora allochthon, Permian foredeep.
INTRODUCTION
Gondwana (Poole et al., 2005). Information on the source of detritus in allochthon and foredeep strata is important in reconstructing Paleozoic paleogeography and paleotectonics of southern Laurentia prior to and during northward emplacement of the allochthon in Permian time. The purpose of this study was to analyze detrital zircons from an ocean-basin conglomeratic barite sandstone in the Sonora allochthon and a calciclastic sandstone in the Mina México foredeep to determine their provenance for comparison with the provenance of detrital zircons from two previously studied
Ages of detrital zircons in Paleozoic sandstones of northwestern Mexico provide information for locating source terranes, for establishing the amount of mixing from multiple sources, and for determining sedimentary dispersal patterns. The Sonora allochthon is a major tectonic feature in northern Mexico (Fig. 1), consisting of Paleozoic offshelf rocks that were thrust onto carbonate-shelf and foredeep-basin rocks of southern Laurentia during the late Paleozoic collision with northwest CALIFORNIA MSM of Anderson & Silver (2005)
h arc
Yuma al ent ntin zoic) o c o ns ale Tra eP ddl (mi
MSM of Stewart (2005)
NEW MEXICO ARIZONA
Tucson
UN
ITE ME D ST XIC ATE S O
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Caborca
Continental shelf and cratonic-platform hingeline (middle Paleozoic)
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Approximate leading edge of SSonora onora allochthon
?
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o si lo mosillo mo Herm Hermosillo
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Inferred edge of continental crust
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Figure 1. Stratigraphic and tectonic setting of samples in central Sonora. Sample localities discussed in text: (1) Los Pozos Formation and (2) Mina México Formation in Minas de Barita area, (3) Mina México Formation in Sierra Santa Teresa, and (4) unnamed unit at Cerro El Pollo in Sierra Agua Verde. Map modified from Poole et al. (2005); note that Baja California and Alta California have been restored ~400 km to their pre-Tertiary position; heavy lines offsetting inferred southern edge of Laurentian crust are hypothesized Neoproterozoic transform faults. Traces of hypothetical Jurassic left-slip MojaveSonora megashear (MSM) of Anderson and Silver (2005) and Stewart (2005) are plotted on map.
Detrital zircons in Paleozoic strata of central Sonora
Age (Ma)
MINAS DE BARITA Mina México foredeep Sonora allochthon & Laurentia shelf
SYSTEM/ SERIES
2
PERMIAN
Mina México Fm > 1000
No Data
L
Rancho Nuevo Formation > 1000
U
Los Chinos Cgl 0–15
MISSISSIPPIAN L
?
Mina México Fm > 500 Unnamed limestone 120
No Data
? Sierra Martínez Group (undivided limestone) >2000
Unnamed limestone > 688
Unnamed limestone > 1500
?
Unnamed limestone 783
Cerro Tasajo Ls 0–75
1 U
3
?
300
U PENNSYLVANIAN M L
350
Laurentia shelf
No Data
?
M
SIERRA AGUA VERDE
SIERRA SANTA TERESA Mina México foredeep & Laurentia shelf
No Data
U
123
Los Pozos Fm 50–150
4 DEVONIAN
Unnamed limestone and sandstone 425
M
400
L No Data
U SILURIAN 450
ORDOVICIAN
L U
El Torote Ls 0–20 El Yaqui Chert 5–150
M
El Mezquite Shale 100–150
L
El Quemado Shale > 300
U 500
CAMBRIAN
M L
Peña Blanca Quartzite 30
?
Pozo Nuevo Ls >1200
Unnamed limestone 898
?
?
Unnamed limestone 137
No Data
Undivided limestone & dolostone
?
No Data
?
Unnamed limestone > 700
?
Proveedora Quartzite > 120
Figure 2. Generalized stratigraphic chart of preorogenic, synorogenic, and carbonate-shelf rocks in the Minas de Barita area; synorogenic and carbonate-shelf rocks in Sierra Santa Teresa; and carbonate-shelf rocks at Cerro El Pollo in Sierra Agua Verde. Stipple indicates flysch facies of Rancho Nuevo and Mina México Formations. Sampled units indicated by numbered circles 1–4. Time scale modified after Gradstein et al. (2005). Unit thicknesses in meters.
samples (Gehrels and Stewart, 1998) of Paleozoic foredeep and carbonate-shelf sandstones in central Sonora (Fig. 2). The Los Pozos detrital zircons are the first Sonora allochthon rocks to be dated by the single crystal analytical method (ID-TIMS). We present analytical data and discuss the possible source areas for the detrital zircons, which have an important bearing on the Paleozoic erosional, depositional, and tectonic history of the southern margin of Laurentia in northern Mexico and the southwestern United States. DEPOSITIONAL AND TECTONIC SETTING OF SAMPLES Paleozoic rocks within the Sonora allochthon in central Sonora comprise Lower Ordovician through Pennsylvanian and possibly lowermost Permian preorogenic and synorogenic strata (Figs. 1–3). Total thickness of the sequence is unknown because of structural complications, but a composite of partial stratigraphic sections indicates that it probably exceeds 2000 m (Fig. 2). The Sonora allochthon is the westernmost segment of
the Ouachita-Marathon-Sonora orogenic system, which borders the southern edge of the Paleozoic North American (Laurentian) craton (Poole and Madrid, 1988; Poole and Perry, 1997; Poole et al., 2005). It formed in late Paleozoic time by oblique collision of South America (Gondwanan continental-margin arc) and the subducting southern margin of the Laurentian craton (Poole et al., 2005). The Sonora allochthon consists of preorogenic (pre–Upper Mississippian) continental-rise deposits 600–800 m thick, and synorogenic (post–Lower Mississippian) flysch deposits at least 1000 m thick. The best exposed Paleozoic ocean-basin (eugeoclinal) sequence in central Sonora is in the Sonora allochthon in the vicinity of Minas de Barita ~100 km east of Hermosillo (Figs. 1–3). The Los Pozos Formation (50–150 m thick) is within the Sonora allochthon, which was thrust northwestward 50–200 km above Laurentian continental-shelf and foredeep rocks. The Lower and Middle Permian Mina México Formation (>1000 m thick) represents the Mina México foredeep fill that accumulated in front of the evolving Sonora allochthon during Permian time. The foredeep formed on the outer part of the Laurentian continental shelf.
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Poole et al.
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EXPLANATION
UNMAPPED C′
28°52.5'
Contact Thrust fault
Early Tertiary (Laramide) plutons
Normal fault
Synorogenic foredeep flysch (Permian)
UNMAPPED
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Quaternary-Tertiary gravels
1
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3 KILOMETERS
Strike-slip fault Anticline
Synorogenic turbidites and deepwater flysch (Permian-Mississippian)
Syncline Overturned anticline
Preorogenic deep-water black shale, turbidite, and chert (Mississippian-Ordovician)
Overturned syncline
Carbonate-shelf rocks (Permian-Cambrian)
Sample localities 1 and 2
109°57.5'
Figure 3. Simplified geologic map of the Minas de Barita area showing the location of sandstone samples 1 and 2. Map from Poole et al. (2005).
A sediment transport direction to the southwest is indicated by sedimentary features in turbidites and submarine-fan systems of the Upper Devonian part of the Los Pozos Formation and Upper Mississippian–Lower Pennsylvanian part of the Rancho Nuevo Formation. SAMPLE DESCRIPTIONS AND ZIRCON POPULATIONS Our two new samples are from the Minas de Barita area in central Sonora (localities 1 and 2, Figs. 1–3). Sample one is from an upper Upper Devonian (Famennian) conglomeratic barite sandstone unit within the Sonora allochthon. It was collected from a 3-m interval ~56 m above the base of the Los Pozos Formation, which measures 115 m in total thickness at this locality (sample 1, Figs. 2 and 3). The second sample is from an upper Lower Permian (Leonardian) calciclastic sandstone in the Lower
and Middle Permian Mina México Formation (sample 2, Figs. 2 and 3). Geochronologic data, generated by isotope dilution— thermal ionization mass spectrometry (ID-TIMS), are listed in Table 1 and shown on Figures 4–6. Two sandstone samples previously analyzed are from sedimentary units in (1) the lower part of the foredeep-basin sequence in Sierra Santa Teresa (upper Lower Permian, Leonardian), and (2) the carbonate-shelf sequence in Sierra Agua Verde (lower Upper Devonian, Frasnian) (localities 3 and 4, respectively; Figs. 1 and 2). Detrital zircon data from these autochthonous to parautochthonous foredeep-basin and carbonate-shelf rocks were published previously by Gehrels and Stewart (1998), but are summarized here for comparison with the suite of zircons from the two sandstones featured herein (Figs. 2 and 6). Zircons were extracted from two ~20 kg samples of sandstones from the Sonora allochthon and Mina México foredeep,
Detrital zircons in Paleozoic strata of central Sonora
125
TABLE 1. U/Pb ZIRCON DATA OF SANDSTONES IN THE SONORA ALLOCHTHON AND MINA MÉXICO FOREDEEP Apparent ages (Ma) 206 207 207 206 206 U Grain Grain wt. Pbc Pb* Pb* Pb* Projected age Pbm Pbc 204 208 238 235 206 type (Ma) (Pg) (pg) (ppm) Pb Pb U U Pb* Mina México Formation (N28°55.33', W109°56.75') CR 6 6 43 447 3.8 931 ± 34 951 ± 39 998 ± 34 1002 ± 36 CE 12 4.6 76 2165 6.2 1047 ± 7 1057 ± 10 1077 ± 11 1079 ± 12 LE 15 11 185 2773 3.2 1079 ± 6 1081 ± 8 1085 ± 8 1085 ± 9 ME 10 9 305 3866 7.7 1086 ± 7 1085 ± 8 1085 ± 8 1085 ± 9 ME 22 5 203 10340 5.7 1084 ± 6 1085 ± 7 1086 ± 7 1086 ± 7 CE 9 6 34 652 6.7 1085 ± 29 1086 ± 34 1086 ± 33 1086 ± 34 LE 12 7 217 4196 6.0 1023 ± 7 1043 ± 8 1084 ± 6 1088 ± 7 LE 9 7 196 2890 5.4 1041 ± 8 1055 ± 9 1086 ± 10 1089 ± 10 CE 14 8 78 1690 6.1 1108 ± 11 1107 ± 13 1105 ± 13 1105 ± 14 LE 10 6 105 2690 6.5 1346 ± 12 1373 ± 14 1415 ± 9 1418 ± 9 CE 14 6 58 2085 7.2 1417 ± 14 1421 ± 16 1426 ± 10 1426 ± 10 LE 21 13 92 2330 4.5 1416 ± 10 1424 ± 12 1436 ± 9 1437 ± 10 LE 17 9 101 2970 5.0 1440 ± 10 1441 ± 11 1442 ± 8 1442 ± 9 LR 20 5 177 11705 6.1 1439 ± 9 1440 ± 10 1441 ± 6 1441 ± 6 LE 14 6 153 5070 5.3 1378 ± 10 1403 ± 11 1441 ± 7 1444 ± 7 ME 7 6 329 5990 6.8 1384 ± 9 1408 ± 10 1445 ± 7 1448 ± 7 CR 8 6 39 1160 4.2 1908 ± 28 1910 ± 31 1912 ± 12 1912 ± 12 LR 15 6 128 7540 6.1 1917 ± 10 1917 ± 12 1917 ± 6 1917 ± 6 CR 9 7 59 1688 4.9 1917 ± 19 1917 ± 21 1918 ± 9 1918 ± 9 MR 14 5 284 17400 4.0 1918 ± 10 1921 ± 11 1925 ± 6 1925 ± 6 Los Pozos Formation (N28°55.94', W109°55.58') LR 4 7 939 1207 6.1 511 ± 3 514 ± 7 525 ± 26 527 ± 30 CE 4 9 397 466 3.4 1224 ± 8 1230 ± 21 1239 ± 28 1240 ± 29 CE 4 6 205 669 6.8 1293 ± 7 1296 ± 19 1300 ± 24 1300 ± 25 LR 3 8 987 1687 3.7 1379 ± 9 1393 ± 13 1415 ± 11 1417 ± 12 CE 4 6 178 541 4.4 1407 ± 7 1412 ± 16 1420 ± 18 1421 ± 19 LR 1 6 105 1003 6.0 1410 ± 8 1418 ± 12 1431 ± 15 1432 ± 16 CR 4 7 473 1346 5.0 1323 ± 9 1365 ± 14 1432 ± 14 1437 ± 14 CR 4 7 800 2190 5.4 1391 ± 9 1411 ± 12 1441 ± 10 1442 ± 10 CR 6 7 214 1732 7.9 1596 ± 9 1620 ± 14 1652 ± 10 1654 ± 10 MR 1 7 214 1539 6.8 1441 ± 7 1528 ± 12 1651 ± 11 1658 ± 12 LE 6 8 782 3000 5.1 1601 ± 9 1632 ± 11 1672 ± 7 1675 ± 8 MR 8 8 157 8290 4.5 1569 ± 7 1613 ± 9 1672 ± 6 1675 ± 6 CR 4 7 395 1232 6.8 1638 ± 11 1656 ± 16 1679 ± 13 1680 ± 12 LR 2 6 807 1553 4.6 1604 ± 9 1641 ± 14 1689 ± 11 1692 ± 10 LE 4 7 625 2366 3.4 1669 ± 11 1679 ± 11 1692 ± 8 1693 ± 8 LR 1 5 126 2188 8.9 1687 ± 10 1694 ± 13 1703 ± 9 1704 ± 9 CR 5 7 116 467 6.2 1690 ± 10 1706 ± 23 1726 ± 20 1727 ± 20 CR 4 5 745 1170 5.5 1822 ± 9 1826 ± 16 1832 ± 12 1832 ± 12 MR 3 9 137 2249 6.2 2310 ± 11 2390 ± 15 2459 ± 6 2462 ± 6 LR 1 7 232 5257 4.0 2562 ± 14 2610 ± 17 2646 ± 5 2648 ± 5 CR 2 7 264 802 4.7 2632 ± 17 2644 ± 23 2654 ± 9 2654 ± 9 LR 1 7 682 991 5.2 2642 ± 13 2684 ± 19 2715 ± 8 2716 ± 8 Note: Grain type: C = colorless, L = light pink, M = medium pink, E = euhedral, R = rounded. All grains abraded to 206 204 ~75% of original diameter with air abrador. Pbm/ Pb—m is measured ratio, uncorrected for blank, spike, or 206 208 fractionation; Pbc/ Pb—c is corrected for blank, spike, and fractionation. Concentrations have an uncertainty of up to 235 –10 238 –10 238 235 25% due to uncertainty of weight of grain. Constants used: O = 9.8485 × 10 ; O = 1.55125 × 10 , U/ U = 137.88. All uncertainties are at the 95% confidence interval. Pbc (pg) is total common Pb in the analysis, in picograms. Pb blank = 5 pg; U blank was consistently <1 pg. Projected ages are upper intercepts projected from 80 ± 40 Ma. *Radiogenic Pb.
each of which was collected from a narrow stratigraphic interval at a single locality. The samples were processed utilizing standard mineral separation techniques (Gehrels and Stewart, 1998). Zircons in the two samples were heterogeneous in color, ranging from colorless to medium pink (Table 1). Morphologies ranged from euhedral, with well-preserved edges and tips on crystals, to well-rounded and spherical grains. Euhedral grains are predominant in the Mina México sample, whereas most grains in the Los Pozos sample are moderately to well rounded (Table 1).
Zircon grains from the larger size fractions (>125 μm) were divided into populations based on color and morphology, and representatives of each were abraded in an air abrasion apparatus (Krogh, 1982) to ~75% of their original diameter. Twentytwo grains from each sample were analyzed by conventional ID-TIMS. The data were reduced by using programs of Ludwig (1991a, 1991b), with the parameters listed in Table 1. Most results are concordant to slightly discordant (Figs. 4 and 5), and are accordingly interpreted to yield robust crystallization ages.
2.75–2.45 Ga events
2.07 Ga event
1.85 Ga event 1.92 Ga event
1.8–1.6 Ga events
1.45–1.34 Ga event
1.25 Ga event
2400
640 Ma event
Los Pozos (Devonian)
1.0–1.1 Ga Grenville
Poole et al.
527 Ma event
126
Pb*/238U
2000
206
1600
Mina México (n=20)
Permian foredeep strata
1200 Permian foredeep strata
800
Santa Teresa (n=21) (unit 7)
Los Pozos (n=22)
Devonian preorogenic strata 207
Pb*/235U
Figure 4. U-Pb concordia diagram of single detrital zircon grains from the middle part of the Los Pozos Formation in the Minas de Barita area. Results are plotted as squares rather than error ellipses because ellipses are too small to be seen at the scale of this figure. Shaded squares represent results that are concordant to slightly discordant.
Devonian carbonate-shelf strata
0.6
0.8
Cambrian Neoproterozoic
1.0
1.2
Cerro El Pollo (n=29) 1.4
Mesoproterozoic
1.6
1.8
2.0
2.2
Paleoproterozoic
2.4
2.6 Archean
Detrital zircon age (Ga)
Figure 6. Age spectra for 92 detrital zircon grains from preorogenic ocean-basin, synorogenic foredeep, and carbonate-shelf rocks in central Sonora. Sample name and number of grains contained in each curve are shown on the right.
Mina México (Permian)
1800
206Pb*/238U
1600
1400
1200
1000
207Pb*/235U
Figure 5. U-Pb concordia diagram of single detrital zircon grains from the lower part of the Mina México Formation in the Minas de Barita area. Results are plotted as squares rather than error ellipses because ellipses are too small to be seen at the scale of this figure. Shaded squares represent results that are concordant to slightly discordant.
This high degree of concordance also suggests that most zircons analyzed have experienced only one phase of growth. Imaging of the zircons by cathodoluminescence (CL) was accordingly not necessary. The interpreted age for concordant grains is based on the 207Pb/206Pb ages, whereas interpreted ages for discordant grains are based on projection to an upper intercept age from 80 ± 40 Ma. This lower intercept is a reasonable minimum age of metamorphism and isotopic disturbance in the region (Gehrels and Stewart, 1998). The results of the analyses are plotted on concordia diagrams in Figures 4 and 5, and are plotted as age probability plots in Figure 6. Zircon grains from the Los Pozos Formation came from several layers of olive- to light olive-gray, conglomeratic, very coarse to fine-grained quartz and barite sandstone (litharenite) in a 3 m interval of the section located on the southeast limb of the Los Chinos syncline (Fig. 3). The sandstones are interbedded with chert and argillite, and the sample contains abundant barite granules and pebbles as large as 2 cm in length. In thin section, ~20% of the rock consists of well-rounded and moderately spherical monocrystalline quartz grains, which appear to be igneous quartz; 5%–11% angular to subrounded chert grains; sparse tourmaline; and rare feldspar and zircon grains. About 50% of the
Detrital zircons in Paleozoic strata of central Sonora rock is crystalline barite, and ~20% is secondary clay, hematite, and limonite. Zircon grains from the Mina México Formation came from an olive- to brownish-gray, graded, very coarse to mediumgrained calciclastic sandstone with some fossil-fragmental and pebbly layers. The sandstone is located ~50 m north of a northwest striking, dextral-slip oblique fault (Fig. 3). The sandstone has abundant very coarse and coarse angular chert and subordinate coarse to medium subangular to rounded quartz grains set in the lime grainstone matrix, which occurs within an olive-gray turbidite sequence of pyritic fine and very fine grained quartz sandstone and siltstone with sparse calciclastic (lime grainstone) interbeds. In thin section, ~25% of the rock consists of subangular to subrounded limestone grains; ~10% well-rounded and highly spherical monocrystalline quartz grains; ~5% angular monocrystalline quartz grains; ~14% subangular shale fragments; ~10% subrounded chert grains; ~5% sandstone grains; ~1% oolite grains; and rare grains of feldspar, tourmaline, and zircon. About 6% of the rock is calcite cement, and 4%–10% is limonite. Sparse fragments of plant and invertebrate remains occur in some layers. Euhedral pyrite is disseminated throughout the rock. AGES OF DETRITAL ZIRCONS Preorogenic Offshelf Sample Los Pozos Formation Twenty-two grains were analyzed from a barite sandstone unit in the Upper Devonian part of the Los Pozos Formation, as previously described. The main clusters of interpreted ages are 1.73–1.65 Ga (n = 9) and 1.44–1.42 Ga (n = 5). In addition, there are three grains at 2.72–2.65 Ga, two grains at 1.30–1.24 Ga, and three additional grains at ca. 2.46, ca. 1.83, and ca. 0.53 Ga (Table 1 and Fig. 6). A previously analyzed quartz sandstone sample (85FP-497) from the Los Pozos Formation, which was collected near our new sample site, yielded multipopulation zircons with an average Pb/ Pb age of ca. 1675 Ma (Poole et al., 1991).
127
et al. (2005) assigned unit 7 to the Mina México Formation, which is transitional with the subjacent Paleozoic carbonate-shelf sequence of Laurentia. The sample analyzed is from a feldspathic sandstone-quartzite with an average zircon grain size of 120 μm. According to Gehrels and Stewart (1998), all of the zircons in this sample are <100 μm, and all analyzed grains were 63–100 μm in sieve size. Most grains in this size range are colorless, some are light to dark pink in color, and a few grains are brownish tan. The brownish-tan grains are euhedral and show no sign of rounding, whereas all of the other grains are moderately rounded and spherical. Most of the zircon grains in the sample yield slightly to moderately discordant ages (Gehrels and Stewart, 1998, fig. 8). The main clusters of interpreted ages are 1.11–1.08 Ga (n = 4), ca. 1.43 Ga (n = 4), and 1.77–1.63 Ga (n = 6); there are two grains at ca. 2.72 and ca. 2.64 Ga, and one grain at 0.63 Ga. Carbonate-Shelf Sample Quartzite of Cerro El Pollo Twenty-nine zircon grains were analyzed from a lower Upper Devonian carbonate-shelf quartz sandstone-quartzite within subunit 24 of Stewart et al. (1999) at Cerro El Pollo in Sierra Agua Verde (Gehrels and Stewart, 1998). The sample is a calcite-cemented quartz sandstone-quartzite with moderately rounded and sorted quartz grains averaging 500 μm in diameter (Stewart et al., 1999). According to Gehrels and Stewart (1998), all of the zircons recovered from this sample were <125 μm, and the analyzed grains were all 80–125 μm in sieve size. Of these, most are light pink in color with subordinate dark-pink and colorless grains. Morphologies range from elongate euhedral crystals to moderately well rounded and spherical grains. Most of the grains in this sample yield slightly to moderately discordant ages, and there are a few highly discordant ages that apparently lie along discordia lines with zircons yielding lessdiscordant ages (Gehrels and Stewart, 1998, fig. 7). According to Gehrels and Stewart (1998), the main age groups are ca. 1.43 Ga (n = 10) and 1.78–1.62 Ga (n = 14), and there are two additional grains at ca. 2.48 and ca. 2.07 Ga.
Foredeep Flysch Sample Mina México Formation Twenty grains were analyzed from a calciclastic sandstone unit in the lower part (Leonardian) of the Permian Mina México Formation, as previously described. The main clusters of interpreted ages are 1.1–1.0 Ga (n = 9), 1.45–1.42 Ga (n = 7), and 1.93–1.91 Ga (n = 4) (Table 1 and Fig. 6). Foredeep Flysch Sample Mina México Formation (unit 7 of Stewart et al., 1997) Twenty-one zircon grains were analyzed from an upper Lower Permian turbiditic sandstone in unit 7 in the northern part of Sierra Santa Teresa described by Stewart et al. (1997). Poole
SOURCES OF DETRITAL ZIRCONS AND PALEOGEOGRAPHIC IMPLICATIONS Most of the detrital zircon ages from the sandstone samples in the Sonora allochthon and Mina México foredeep of central Sonora suggest derivation from igneous and metamorphic rocks that are widespread in the southwestern part of the Laurentian craton (Fig. 7). Zircons from the Yavapai, Mazatzal, and Mojave crustal provinces and their transition zones (1.8–1.6 Ga) of the southwestern United States and northwestern Mexico are recognized in the samples from the Sierra Santa Teresa unit 7 of Stewart et al. (1997), Los Pozos, and Cerro El Pollo, but not in our Mina México sample (Fig. 6). The 0.53 Ga zircon is believed to have been derived from a Lower Cambrian granitoid
Poole et al. 105°
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Mesoproterozoic
PRECAMBRIAN CRUSTAL PROVINCES PACIFIC OCEAN
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Grenville (Llano) Province 1.3–1.0 Ga
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Mazatzal Province
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Transition (Mojave-Mazatzal) zone 1.8–1.6 Ga Yavapai Province 1.8–1.7 Ga
Transition (Mojave-Yavapai) zone 1.8–1.6 Ga Mojave Province (includes Caborca crust) 1.8–1.6 Ga
Figure 7. Distribution and ages of crystalline basement rocks in southwestern Laurentia following Neoproterozoic–early Paleozoic breakup of Rodinia. Compiled by Poole et al. (2005) primarily from data of Anderson, 1989; Hoffman, 1989; Wooden and DeWitt, 1991; Anderson et al., 1993; Karlstrom and Bowring, 1993; Van Schmus and Bickford et al., 1993; Sims et al., 1993; Smith et al., 1999; Barnes et al., 1999; Karlstrom et al., 1999, 2002; Barth et al., 2000; Iriondo, 2001; Ross and Villeneuve, 2003; Iriondo et al., 2004; and Farmer et al., 2005. AOT—AlabamaOklahoma transform fault; TT—Texas transform fault, CST—Chihuahua-Sonora transform fault; BCT—Baja California transform fault.
or metamorphic rock northeast of central Sonora, possibly in New Mexico and Colorado, or Oklahoma. Detrital zircon geochronology suggests that most of the detritus in both samples was derived from Laurentia to the north, whereas some detritus in the Permian synorogenic foredeep sequence was derived from the evolving accretionary wedge to the south. Most of the grains in the Los Pozos sample were probably derived from both nearby and distant sources to the north, including the Yavapai, Mojave, and Mazatzal Provinces and their transition zones (1.8–1.6 Ga) and scattered granite bodies (1.44–1.24 Ga) intruding the older basement. Magmatism ranging in age from 1630 to 1220 Ma is known in southwestern New Mexico (McLemore et al., 2000; Rämö et al., 2003). Alternatively, the 1.44–1.24 Ga grains may be from the Southern and Eastern Granite-Rhyolite Provinces and Grenville (Llano) Province. The other grains include a mixture
of Wyoming Province (2.72–2.46 Ga), Trans-Hudson and Penokean orogens (ca. 1.83 Ga), Grenville (Llano) Province (1.30–1.24 Ga), and Lower Cambrian metamorphic rocks and granitoids (0.53 Ga) to the north and east of central Sonora (Table 1 and Fig. 6). See Figure 7 for distribution and inferred ages of crystalline basement rocks of southern Laurentia. The younger grains (1.1–1.0 Ga) in the Mina México calciclastic sandstone were probably derived from nearby granitic sources in the Grand Canyon region of Arizona (Timmons et al., 2001), and from the Grenville (Llano) Province to the east of central Sonora and Grenville-age intrusions in northern Sonora and Colorado; the intermediate-age grains (1.45–1.42 Ga) were probably derived from granitic bodies intruding older basement rocks in the southwestern United States and northwestern Mexico. The older grains (1.93–1.91 Ga) may be from
Detrital zircons in Paleozoic strata of central Sonora Paleoproterozoic basement rocks of the Lake Superior region (fold belts associated with the Trans-Hudson and Penokean orogens) or recycled from older Paleozoic strata exposed north and east of central Sonora. No zircons of Wyoming Province age (>2.0 Ga) are present in our Mina México sample. Gehrels and Stewart (1998) concluded that most of the grains in the Cerro El Pollo sandstone were probably derived from nearby basement rocks of the southwestern United States. The older zircons, however, must have been recycled from underlying Ordovician and older strata or perhaps shed from >2.0 Ga rocks of the Wyoming Province to the north. One zircon of apparent Grenville (Llano) (1.3–1.0 Ga) age is present in the sample (Gehrels and Stewart, 1998, table 1 therein). As with the upper Upper Devonian Los Pozos sample, most of the zircon grains in the upper Lower Permian Mina México samples were probably derived from Precambrian rocks that are widespread in Arizona, New Mexico, Colorado, and northern Sonora and Chihuahua. Another possible source for some of the Permian foredeep-basin detritus was the evolving late Paleozoic accretionary wedge (Sonora allochthon) to the south (Fig. 1). The provenances of most zircon grains in our upper Upper Devonian Los Pozos Formation sample may be the Mazatzal (1.7–1.6 Ga), Yavapai (1.8–1.7 Ga), and Mojave (1.8–1.6 Ga) Provinces, and scattered granite bodies north of central Sonora. Other possible provenance areas on the Laurentian craton include the Wyoming Province, Grenville (Llano) Province, and Grenville-age intrusions (Hoffman, 1989; Van Schmus and Bickford et al., 1993; Timmons et al., 2001). Numerous Lower Cambrian metamorphic rocks and granitoids occur to the north and east of central Sonora. Other possible sources for the 527 Ma zircon may be Cambrian plutonic and volcanic rocks in New Mexico and Colorado (Loring and Armstrong, 1980; Armbrustmacher, 1984; Evans and Clemons, 1988; Matheney et al., 1990; McLemore et al., 1999; McMillan et al., 2000; McMillan and McLemore, 2004) and Oklahoma (Ham et al., 1964; Gehrels and Dickinson, 1995), or recycled Cambrian detritus during Devonian erosion—for example, the Cambrian Bolsa Quartzite (which contains a 525 Ma zircon grain)—may have been exposed on the craton to the north (Gross et al., 2000; Stewart et al., 2001). Another possible source area is in Colorado where Cambrian alkaline intrusive complexes in the Wet Mountains intrude Proterozoic metamorphic and plutonic rocks (Armbrustmacher, 1984). The probable provenances of most zircon grains in our sample of the upper Lower Permian part of the Mina México Formation are the Grenville (Llano) Province to the east and Grenvilleage intrusions to the north, granitoid bodies in the southwestern United States and northwestern Mexico, and the Southern and Eastern Granite-Rhyolite Provinces to the northeast (Anderson, 1989; Van Schmus and Bickford et al., 1993; also plate 1 compiled by J.C. Reed Jr., 1993; Stewart et al., 2001, fig. 6; Timmons et al., 2001; Rämö et al., 2003; Poole et al., 2005, fig. 2). Zircons with ages of 1.93–1.91 Ga may be from Paleoproterozoic rocks to the north in the Trans-Hudson and Penokean orogens (Sims et al., 1993; Barnes et al., 1999; Ross and Villeneuve, 2003),
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or they may be recycled from eroded Proterozoic and Paleozoic sandstones in the southwestern United States. All four samples contain ca. 1.43 Ga zircons, suggesting a Mesoproterozoic tectonic and intrusive event within the older Paleoproterozoic Mojave Province (1.8–1.6 Ga) basement in Sonora. Some basement rocks in the Caborca crust have yielded zircons of ca. 1.45 Ga in age (Anderson and Silver, 1981). In addition, zircons of this age may have been sourced in the Southern and Eastern Granite-Rhyolite Provinces (1.45–1.34 Ga) to the northeast (Figs. 6 and 7). Granites of similar age (ca. 1.4 Ga) occur throughout New Mexico (Rämö et al., 2003; V.T. McLemore, 2006, personal commun.). The age spectra of detrital zircons from ocean-basin (eugeoclinal) and carbonate-shelf (miogeoclinal) strata are similar for samples from rocks of similar age, but the sample pair of one age is different from the sample pair of another age. Specifically, similar spectra for the preorogenic upper Upper Devonian eugeoclinal (continental rise) Los Pozos sample and lower Upper Devonian miogeoclinal (carbonate shelf) Cerro El Pollo sample differ from spectra for the synorogenic upper Lower Permian foredeep Mina México sample in the Minas de Barita area and synorogenic upper Lower Permian foredeep Mina México unit 7 sample in Sierra Santa Teresa (Fig. 6). This implies that the Los Pozos and Cerro El Pollo samples had a similar source terrane that was different from that of the Mina México samples in the Minas de Barita and Sierra Santa Teresa areas. The compositional and sedimentological differences reflect their different geologic settings. Clearly, the provenance of these sampled strata varied through time, but the change in provenance was the same through time in ocean-basin, foredeep-basin, and carbonate-shelf strata. CONCLUSIONS AND DISCUSSION Although multiple sources of zircons are indicated by our study of ocean-basin and foredeep sandstone samples, the major sources were crustal provinces and granitoid bodies exposed in the Laurentian craton to the north and east of central Sonora (Figs. 6 and 7). We conclude that (1) zircons older than 2.0 Ga were derived originally from the Wyoming Province in the northern United States; (2) zircons ranging in age from 1.93 to 1.85 Ga may have been derived originally from the Trans-Hudson and Penokean orogens of the Lake Superior region; (3) zircons ranging in age from 1.73 to 1.65 Ga probably came from the Mazatzal and Mojave Provinces and their transition zones in the southwestern United States and northwestern Mexico; (4) zircons ranging in age from 1.45 to 1.42 Ga most likely were derived from widespread granitoid plutons in the southwestern United States and northwestern Mexico, and possibly the Southern and Eastern Granite-Rhyolite Provinces in the southern United States; and (5) zircons ranging from 1.3 to 1.0 Ga probably were derived from the Grenville (Llano) Province east of central Sonora and scattered coeval intrusive bodies (Stewart et al., 2001, fig. 6; Poole et al., 2005, fig. 2) north and east of central Sonora. The zircon from the Upper Devonian part of the Los Pozos Formation, dated as ca. 527 Ma, most likely was derived from Lower Cambrian
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metamorphic rocks and granitoids exposed in the craton north and northeast of central Sonora, or was a recycled grain from erosion of the Cambrian Bolsa Quartzite or other pre-Devonian sandstones in northern Mexico. The lower Upper Devonian carbonate-shelf and upper Lower Permian foredeep-basin sandstones were derived mainly from 1.45–1.40 Ga granitoids that are widespread in the southwestern United States and northwestern Mexico, 1.77–1.62 Ga rocks of the Mazatzal Province, and 1.12–1.08 Ga rocks of the Grenville (Llano) Province east of central Sonora and granitoids in Sonora and adjacent United States. Rocks of the Mojave crustal province (1.8–1.6 Ga) are recognized in northwestern Mexico, but zircons of that age are absent in our Mina México sample. Also, zircons of Late Archean age (>2.4 Ga) are absent from our Mina México sample but are present in the other three samples (Fig. 6). The suite of zircons from the Los Pozos Formation and Cerro El Pollo unit differ from those in the Mina México Formation of the Minas de Barita and Sierra Santa Teresa areas, indicating different geological settings. Also, compositional and sedimentological differences between the continental-rise Los Pozos conglomeratic barite sandstone and the foredeep Mina México calciclastic sandstone imply different depositional and tectonic settings. Sedimentary features in turbidites and submarine-fan systems of the Upper Devonian part of the Los Pozos Formation and Upper Mississippian–Lower Pennsylvanian part of the Rancho Nuevo Formation indicate a southwesterly transport direction. These features are evidence for a south to southwest regional gradient along the margin of the Laurentian craton and continental shelf during much of the Paleozoic. Most of the zircons in Paleozoic sandstones in northern Mexico are believed to have been derived initially from basement rocks of the Laurentian craton. ACKNOWLEDGMENTS AND RESPONSIBILITY The U.S. Geological Survey supported fieldwork of Poole and Stewart, and the U.S. National Science Foundation (EAR-9416933) supported field and laboratory work of Gehrels. Stewart and Gehrels collected the two samples from the Sonora allochthon and Mina México foredeep strata in the Minas de Barita area, Gehrels processed the samples and compiled the analytical data, and Poole evaluated the local geologic setting and regional significance of the samples. We thank Matthew S. Spurlin for assistance in processing the two samples from the Minas de Barita area. Paula L. Hansley of Petrographic Consultants International provided modal analyses of thin sections. Manuscript reviews by John N. Aleinikoff, Karl V. Evans, Virginia T. McLemore, and Wayne R. Premo improved the paper. We are grateful to Barbara J. Ramsey and Norma J. Maes for computer drafting of Figures 1–3 and 7. REFERENCES CITED Anderson, J.L., 1989, Proterozoic anorogenic granites of the southwestern United States, in Jenney, J.P., and Reynolds, S.J., eds., Geologic evolution of Arizona: Tucson, Arizona, Arizona Geological Society Digest 17, p. 211–238.
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The Geological Society of America Special Paper 442 2008
The flora, fauna, and sediments of the Mount Dall Conglomerate (Farewell Terrane, Alaska, USA) David Sunderlin* Department of Geology and Environmental Geosciences, Lafayette College, Easton, Pennsylvania 18042, USA ABSTRACT New collections of floral and faunal remains were recovered from late Paleozoic sediments of the Mount Dall conglomerate in the Alaska Range of south-central Alaska. This isolated unit’s type section is ~1500 m thick and comprises thick to very thick conglomerate beds with interbedded sandstones and siltstones in a series of fining-upward intervals each tens of meters thick. The unit is interpreted to be a coastal braidplain deposit of Early Permian age in the upper Farewell terrane (Mystic subterrane sequence). Genus-level taxonomic composition of paleobotanical collections from lenticular mudstones to siltstones is discussed with regard to taphonomy and the interpreted lowland paleoenvironment of deposition. Poorly to moderately preserved megafossil compressions and impressions of the foliage genera Pecopteris, Zamiopteris, Rufloria, Angaropteridium, Cyclopteris, and Cordaites are consistent through several hundred meters of section and suggest a locally dense floral community. Horizons with sideritic rhizoliths indicate the presence of immature soils. The co-occurrence of these foliar and reproductive organs in the Mount Dall conglomerate suggests a mixed phytogeographic affinity to both the temperate Angaran Floristic Province of northern Pangea and the Euramerican Province of lower paleolatitudes. The brachiopod genera ?Stenoscisma and ?Schuchertella also were recovered and indicate a coastal depositional setting. These new biogeographic data complement exclusively marine zoogeographic data from the Farewell terrane’s older strata and may be used to test hypotheses regarding the paleogeography of this displaced continental fragment. The paleofloral data support the placement of this terrane within a midlatitude climate belt during the Early Permian. Keywords: Farewell terrane, Mount Dall conglomerate, Mystic subterrane, Permian, Alaska, Zamiopteris. INTRODUCTION The regional geology of most of Alaska has been recognized as a complex belt of discreet blocks each with internally continuous stratigraphy. These blocks, or “tectonostratigraphic
terranes,” are bounded by fundamental discontinuities in stratigraphy that are not genetically related and may be due to either faulting or stratigraphic overlay (Coney et al., 1980; Jones et al., 1982). Determining the depositional, tectonic, and paleogeographic origins of these terranes with respect to paleolatitude, one
*
[email protected] Sunderlin, D., 2008, The flora, fauna, and sediments of the Mount Dall Conglomerate (Farewell Terrane, Alaska, USA), in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 133–150, doi: 10.1130/2008.442(09). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
133
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Sunderlin
another, and nearby cratons has been the subject of much study in efforts to reconstruct the geologic history of Alaska, the northern Pacific, and the Arctic Basin as well as the entirety of the Cordillera of North America (Jones et al., 1982; Plafker and Berg, 1994; Johnston, 2001). Studies of the biogeographic affinities of fossil biotas have both generated novel hypotheses of a particular terrane’s geologic history and provided independent tests of hypotheses initially derived from geological methods. The Farewell terrane is a geologically integral and geographically central component in the Alaskan terrane system but its origin, development, and paleogeographic history remain incompletely understood. Field mapping and stratigraphic work of Reed and Nelson (1980) provided a major portion of the geologic groundwork for the establishment of an early terrane model of Jones et al. (1982). A thorough treatment of the Farewell terrane’s definition was most recently synthesized by Decker et al. (1994), who built on this work and stratigraphically linked the unit’s component subterranes. Simplified lithological successions from the most complete sections of the Farewell terrane as well as subterrane sequence nomenclature, as it will be used in this paper, are presented in Figure 1. Blodgett et al. (2002) and Dumoulin et al. (2002) have recently summarized the growing body of paleontological and biogeographic work on the Farewell terrane throughout the Paleozoic and this, too, is summarized in Table 1. Considerable progress in answering questions about the Farewell terrane’s history has been made recently in southern Alaska from the synthesis of these and other data sources (Blodgett et al., 2002; Dumoulin et al., 2002; Bradley et al., 2003). Here, I provide new data from the Early Permian Mount Dall conglomerate of the Farewell terrane that may prove important in furthering an understanding of the tectonostratigraphic development of south and central Alaska. These data include the most detailed description of sediments of the Mount Dall conglomerate (Mystic subterrane sequence) yet published and an interpretation of the paleoenvironment of their deposition. I also build on the work of Mamay and Reed (1984) by introducing the most WM
complete collection of the paleofloral compression-impression assemblage of this unit yet obtained and the first examination of this flora in a sedimentological and taphonomic context. The floral remains from the Mount Dall conglomerate provide the only existing terrestrial biogeographic datum for use in regional reconstruction of the geologic history of the Farewell terrane. Additionally, the Mount Dall flora is the only paleofloral datum of this age known within a 1500-km radius (present geography) and, if originally deposited close to its present geographic coordinates (see Bradley et al., 2003), provides insight into broader phytogeographic patterns in the temperate latitudes of northern Pangea in the Early Permian. These sedimentological and biogeographic data together may help refine the regional geologic history of south and central Alaska as well as assist in understanding global floral provinciality in the late Paleozoic. GEOLOGIC SETTING The “Farewell composite terrane” of south-central Alaska, USA (herein referred to simply as the Farewell terrane, and of which the Mount Dall conglomerate is a component) preserves a geological record from the Precambrian through Early Cretaceous with most of its stratigraphic thickness composed of sediments deposited in the marine realm (Decker et al., 1994). These highly metamorphosed to unmetamorphosed sediments and volcanics are discussed by Decker et al. (1994); the most complete Paleozoic sections are summarized in stratigraphic context in Figure 1. The Farewell terrane is delimited geologically by fault or stratigraphic overlay contacts with the neighboring Ruby, Innoko, Yukon-Tanana, and Kahiltna terranes. The Ruby terrane to the north is a continental margin sequence of Paleozoic age (with metamorphism to blueschist facies). The Innoko terrane is to the west and is a Devonian– Early Cretaceous ocean basin and arc sequence. The heterogeneous Yukon-Tanana terrane is located to the east of the Farewell terrane and lacks evidence of any tectonic interaction with the Farewell terrane in the late Paleozoic, and it is likely that the proximity of the two was achieved later on (Bradley et al., 2003). The Kahiltna ter-
HB
Triassic AR
Dall Basin sediments
Mystic subterrane sequence KM
Devonian
Nixon Fork subterrane
Dillinger subterrane
Carbonate Platform
Basinal Facies
Cambrian
Continental Basement
Unknown Basement
Figure 1. Simplified lithological and component relationships within the Farewell terrane. Sections modified from Decker et al. (1994). KM—Kuskokwim Mountains (northern), WM—White Mountain, HB— Holitna Basin, AR—Alaska Range (southwestern). Terrane component scheme modified from Bradley et al. (2003).
Flora, fauna, and sediments of Mount Dall Conglomerate TABLE 1. BROAD BIOGEOGRAPHIC AFFINITIES OF FOSSIL BIOTAS COLLECTED FROM PALEOZOIC FAREWELL TERRANE SEDIMENTS Biotic group References
Period
Biogeographic affinity Mixed
Permian
Early
Plants (MY) Brachiopods (MY)
Carboniferous
Late
Trilobites (MY)
Early
N.D.
Late
Foraminifers (MY)
Mamet and Plafker (1982)
Middle
Brachiopods (NF) Corals (NF) Gastropods (NF) Cephalopods (NF) Bivalves (NF) Rostroconchs (NF) Trilobites (NF) Ostracodes (NF) Crinoids (NF) Sponges (NF) Algae (NF) Brachiopods (NF) Brachiopods (MY) Conodonts (MY)
House and Blodgett (1982); Rigby and Blodgett (1983); Poncet and Blodgett (1987); Baxter and Blodgett (1994); Blodgett and Johnson (1994); Johnson and Blodgett (1993); Blodgett and Rohr (1989); Blodgett (1992, 1993); Blodgett et al. (2002)
North America/ Uncertain
Blodgett and Brease (1997); Blodgett et al. (1995, 2000); Blodgett (1998); Savage and Blodgett (1995); Blodgett et al. (2002), Frýda and Blodgett (2004)
Siberia
Late
Aphrosalpingids (NF; DL) Brachiopods (NF) Algae (NF)
Blodgett et al. (1984); Soja and Antoshkina (1997); Clough and Blodgett (1985, 1988, 1992); Blodgett and Clough (1985); Blodgett and Gilbert (1992); Soja (1994)
Siberia/ "Old World"
Early
N.D.
Late
Brachiopods (NF) Gastropods (NF) Corals (NF)
Middle
Marine (NF)
Early
Conodonts (NF)
Late
N.D.
Middle
Trilobites (NF)
Early
N.D.
Devonian
Early
Silurian
Ordovician
Cambrian
This study
135
Hahn et al. (1985); Hahn and Hahn (1985)
N.A.
Siberia
N.A.
N.A.
Siberia
N.A.
Rohr and Blodgett (1985); Blodgett et al. (1992); Potter et al. (1988); Potter and Blodgett (1992)
Mixed
Rohr et al. (1992); Rohr and Gubanov (1997); Rohr and Yochelson (1999); Measures et al. (1992)
Mixed
Dumoulin et al. (2002)
Mixed
N.A.
N.A.
Babcock et al. (1995); Palmer et al. (1985); Babcock and Blodgett (1992); Babcock et al. (1993); St. John and Babcock (1997); Kingsbury and Babcock (1998) N.A.
Siberia/ Greenland N.A.
Note: MY—Mystic Subterrane; DL—Dillinger Subterrane; NF—Nixon Fork Subterrane; N.D.—no data.
rane is a Jurassic–Middle Cretaceous flysch belt that is thought to cover the suture between the Wrangellia terrane and the Farewell terrane. Collision of Farewell and Wrangellia was likely to have taken place by the Cretaceous providing the provenance for the sedimentation in the Kahiltna terrane. Nonmarine and marine predominantly siliciclastic facies of the Late Cretaceous–earliest Tertiary Kuskokwim Group unconformably overlie much of the Farewell terrane, structurally “stitching” the unit and limiting Farewell outcrop area (Fig. 2). The Dillinger (environmental equivalent to the Minchumina), Nixon Fork, and Mystic subterranes are the main stratigraphic components of the Farewell terrane, with described
stratigraphic sequences from outcrops in the Alaska Range, Holitna Basin, White Mountain area, and the northern Kuskokwim Mountains (Decker et al., 1994) (Figs. 1 and 2). The Dillinger subterrane is characterized mainly by basinal carbonates with interbedded mudstones and shales of Cambrian to Devonian age. These rocks are correlative with the Nixon Fork subterrane facies of shallower platform carbonates, and a slope transitional connection between the subterranes is clearly evident in the Ordovician and Silurian sections of the Holitna Basin and White Mountain area (Churkin et al., 1984; Decker et al., 1994). Both of these subterrane sequences unconformably
136
Sunderlin
156 W
153 W
150 W Fairbanks
IN
Alaska
MN NF
64 N
nt ain ou
Ra ng e
wi
MY
Al as ka
-
ok
od tar Idi
DL
Ku sk
63 N
o Nix
M
lt Fau k or nF
m
RB
s
IN
WR
White Mountain
KU 62 N
Figure 2. Map of the components of the Farewell terrane. Dark gray—Dillinger (DL) and Minchumina (MN); light gray—Nixon Fork (NF); textured—Mystic (MY). Other units: CG—Chugach terrane; IN—Innoko terrane; KU—Kuskokwim Group; KH—Kahiltna terrane; PE—Peninsular terrane; RB—Ruby terrane; WR—Wrangellia terrane. Boxed area within MY enlarged in Figure 3. (Modified from Wilson et al., 1998.)
t
ali Den
Fa
Holitna Basin
aul ell F w e r
PE KH
DL KU
50 km
overlie Precambrian metasediments and metavolcanics and are themselves unconformably overlain by the so-called Mystic subterrane sequence of the Late Devonian–Early Jurassic (Fig. 1). The Mystic sequence is laterally heterogeneous on a regional scale, being composed of cherts, argillites, pillow basalts, platform carbonates and, at this study’s field area at Mount Dall in the Alaska Range, ~1500 m of thick to very thick cobble to boulder conglomerate beds with semicontinuous gravel to mudstone interbeds. MOUNT DALL CONGLOMERATE The Mount Dall conglomerate crops out in rugged, alpine terrain south of the Denali-Farewell fault in the western region of Denali National Park and Preserve, Alaska, USA (Talkeetna C-5 quadrangle) (Figs. 3 and 4A) (Reed and Nelson, 1980). It is exposed in a broad, ~N60°E trending (5–8° plunging) syncline with occasional intrusive dikes and local contact metamorphism of the country rock. For this study, alpine conditions restrict helicopter, ski-plane, and foot access to only ~600 m of the estimated total ~1500 m thickness of the Mount Dall conglomerate. Access begins at least 500 m from the presumed base of the unit. The field area is located at 63.6°N, 152.2°W, 6 km N75°E of Mount Dall proper, at elevations between 1200 m and 1500 m, west of Dall Glacier and between its ultimate and
Anchorage
CG
penultimate unnamed tributary glaciers draining from the west (Figs. 3, 4A, 4B). Major facies associations within the studied section include (1) conglomerate beds (Fig. 4C and 4D) that are horizontally variable in thickness (~1–10 m) and (2) interbedded mudstone to sandstone beds (~0.5–3 m) (Fig. 4E and 4F). Crosscutting felsic intrusive rocks of presumably Late Cretaceous or Tertiary age (“TKi” of Reed and Nelson, 1980) are infrequent and do not disturb stratigraphic continuity in the field area. The Mount Dall conglomerate unconformably overlies a thick, Pennsylvanian (?) turbidite-flyschoid facies succession that is isoclinally folded (Reed and Nelson, 1980; Bradley et al., 2003). The stratigraphic log of the measured section is shown in Figure 5 and, more completely, in Appendix 2 of Sunderlin (2005). Facies Association I: Conglomerates The dark-brown to gray, erosionally resistant conglomerate units of the Mount Dall conglomerate are laterally traceable across the 6 km2 field area and into inaccessible terrain to the west toward the southern face of Mount Dall itself (Fig. 4B). Data on cobble composition and size were gathered from meter-long vertical transects across accessible faces at many of the 3- to >10-m-thick conglomerate beds in the field area. These data were inserted into
Flora, fauna, and sediments of Mount Dall Conglomerate
137
152 16' W
152 08' W
62 36'N B C
Pd?
A
Pd Pd
Mt. Dall
x
Pzus
Pzus
62 34'N 1 km
Pd?
Figure 3. Mount Dall field area. Area enlarged from box in Figure 2. Light gray—Mount Dall Conglomerate summer outcrop (Pd); dark gray— Kahiltna terrane sediments or lower Mystic subterrane sediments (Pzus); texture—glacial moraine; white—alpine snow-glacial ice. Sections in Figure 5 labeled A–C. Dashed line indicates axis of broad plunging synclinal structure. Modified from Reed and Nelson, 1980.
simple calculations of volumetric contributions of each lithology to the conglomerate bed (Fig. 6). Clasts range in diameter from 1 to 20 cm in the sampled 1 m transects, with maximum diameters of ~40–50 cm observed on some conglomerate outcrops. Clasts are composed of radiolarian chert, limestone, intraformational sandstone-conglomerate, and rare igneous pebbles (Sunderlin, 2002), all of which are derived from underlying strata of the Farewell terrane (Bradley et al., 2003). The pebble-boulder conglomerate units are clast-supported with a well-indurated, medium-coarse light-gray silt to fine sand matrix of the same lithological composition as the gravel-cobble fraction (Fig. 4D). The volumetric contribution of carbonate clasts increases relative to chert clasts through the measured section with noticeable but statistically weak correlation (Fig. 6), suggesting a shift in the provenance of the sediments through the accumulation of the unit. Black and red chert clasts are subrounded and most likely derived from basinal deposits in the Mystic or Dillinger subterrane sequences. The limestone clasts are well rounded and have been faunally and lithologically linked to limestone units elsewhere (Shellabarger Pass) within the Farewell terrane by faunal association of Middle Devonian (?) megafossils (Reed and Nelson, 1980; Blodgett and Boucot, 1999). Bradley et al. (2003) processed three limestone clasts that yielded conodonts with Late Carboniferous (Morrowan) to earliest Permian age ranges and attributed the provenance of at least some limestone cobbles to marine strata of this age elsewhere within the Mystic subterrane (possibly the Holitna Basin or the Cheeneetnuk River–White Mountain area). Clast
imbrications, where discernible, indicate a flow direction toward the paleoeast, which is constant through the measured section (Sunderlin, 2002; Bradley et al., 2003). Facies Association II: Gravelly Sandstones to Mudstones The conglomerates described above are interbedded with gravelly (pebble) sandstones to subordinate clayey mudstones that are erosionally less resistant, laterally discontinuous, normally graded, and variable in bed thickness from 1 to 5 m. These strata show internal scour-and-fill structures, slump structures, and small-scale cross-bedding, and are truncated by scour surfaces. These scour surfaces result in discrete sediment sequences, where the thinnest such units grade from pebbles through coarse sandstones (Fig. 4C). The thicker (>3 m) sections of this facies group fine upward into siltstones and mudstones that are laterally continuous up to 15 m. Occasional plant impressions, siderite-replaced comminuted plant debris, and sideritic rooted horizons (Fig. 4E) are preserved within the finest-grained layers, along with shallow-flow ripples and cross-laminations. Highly organic-rich or leaf mat beds (Fig. 4F) are rare through the measured section but, where preserved, are richly fossiliferous layers of siltstone to mudstones topping a normally graded gravelly sandstone to mudstone sequence. The typical sedimentary succession within the subconglomerate interbeds is laterally discontinuous on a 10-m scale with cross-bedding and channel scours 5–15 m wide.
B C
D
F
E Figure 4. Field photographs at Mount Dall. (A) Photograph of Mount Dall field area. View toward WSW. (B) Photograph of Mount Dall Conglomerate strata looking W toward south slope of Mount Dall. Conglomerate units appear as erosionally resistant benches. (C) Gravels with cross bedding (scale, 50 cm). Dark clasts are predominantly cherts, light clasts are carbonates or intraformational sandstones. (D) Photograph of conglomerate unit (hammer for scale, ~35 cm). (E) Sideritic rhizomorph horizon with individual root trace clusters indicated with arrows. Laminar bedding in mudstones is perturbed by rooting horizons (scale, 5 cm). (F) Cordaitalean leaf mat (ice axe pick for scale, maximally 2 cm wide).
m
10
A
B
9
m
330
m
9
300
C
270 300
7
8
8
250
Mud
Sand Gravel Congl.
7
260 250
6
200
150
Interval
12
200
5
150
4
11
100
100
3
50
10
50
Mud
Fern foliage
Sand Gravel Congl.
2
0
Seed plant foliage
Mud
Sand Gravel Congl.
1
Marine fossils 0
Figure 5. Composite stratigraphic log of the Mount Dall conglomerate. Sections A–C correspond to measured outcrop labeled in Figure 3. Log begins at least 500 m from base of the unit. Section C stratigraphically overlies sections A and B. Intervals 1–12 labeled on section log axis.
140
Sunderlin 1 0.9
R2 = 0.4565
0.8 0.7 0.6
Figure 6. Scatterplot and linear regressions of proportional volumes of conglomerate cobbles in conglomerate beds. Samples taken over 1 m transects by interval in Figure 5. Closed circles— cherts; open squares—carbonates.
0.5 0.4 0.3 0.2
R2 = 0.5064
0.1 0 0
1
2
3
4
5
6
7
8
9
10
11
12
Interval DEPOSITIONAL ENVIRONMENT Both the coarse-grained nature of the section overall and the lateral discontinuity of the finer-grained sedimentary packages suggest a high-energy depositional environment. The chert and limestone boulder to pebble clasts are subrounded to well-rounded and it is likely that they were derived from local topography of uplifted sediments of older marine units in the Farewell terrane located to the paleo-west. The spatial arrangement and frequency of erosional scours, the rare development of immature rhizomorphic paleosols, and the overall thickness of the Mount Dall conglomerate all indicate the frequent restructuring of a fluvial, braided-stream system in which sediments were deposited, perhaps with climatically or tectonically driven events of increased conglomerate deposition demarking the mesoscale stratigraphic intervals (Fig. 5) (Sunderlin, 2001). Siderite deposition associated with refractive organic material suggests the presence of constant soil moisture at moderate pH (Baas Becking et al., 1960) and fresh-water or, at most, brackish salinities (Berner and Cochran, 1998). FLORAL REMAINS S. Nelson, R. Detterman, and B. Reed made a small collection of Mount Dall conglomerate plant fossils (n ~20) while field mapping in the Alaska Range in 1976. The collection and sediments were originally assigned a Pennsylvanian age by Reed and Nelson (1980), but this was changed to Early Permian after a description of this collection presented by Mamay and Reed (1984). Three species of the fern foliage form genus Pecopteris (P. unita Brongniart, P. arborescens Schlotheim, P. hemitelioides Brongniart), leaf impressions of a cyclopteroid form of unknown generic affinity, and the cordaitalean leaf genus Zamiopteris were
identified from the original collection. The last, known elsewhere only from Permian deposits, provided limits for the age of the Mount Dall flora and the sediments in which it was preserved. These specimens are fragmentary impressions in lenticular siltstone horizons and are among the collections of the National Museum of Natural History, Smithsonian Institution, in Washington, D.C., USA (USNM 312727–312734). Genus Pecopteris The three species of the genus Pecopteris identified by Mamay and Reed (1984) are known elsewhere from the equatorial regions of Carboniferous-Permian Pangea now within present-day Europe and southern North America. Pecopteris unita Brongniart is well known for its strong pinnule venation visible even under moderate preservational conditions (Haubold, 1985). P. arborescens (alt. arborenscens) Schlotheim is a compact form with stout pinnules that originate at right angles to the secondary rachis. P. hemitelioides (alt. hemiteloides) Brongniart is a species with larger pinnules that are observed to recurve distally toward the pinna apex and have blunt terminations of the ultimate pinnule rachises. These Pecopteris species occur in many Late Paleozoic lowland depositional basins including strata in Germany (Haubold, 1985), New Brunswick (Bell, 1962), Spain (Wagner, 1971), and the Dunkard Group in southern Pennsylvania–northern West Virginia (Gillespie et al., 1975). All three of these forms are stratigraphically long-ranging and cross the Carboniferous-Permian boundary (Mamay and Reed, 1984). In the Mount Dall conglomerate, remains of this genus are not found with associated reproductive structures and therefore cannot be assigned more specific natural nomenclature.
Flora, fauna, and sediments of Mount Dall Conglomerate
141
Cyclopteroid-like Foliage
Angaropteridium sp. (Fig. 7C)
Five specimens referable to the cyclopteroid foliage form are discussed by Mamay and Reed (1984) and two are figured (their figs. 62C and 62D). These specimens have many veins entering the pinnule and have moderately restricted bases but the apical parts of the pinnules were not preserved. These specimens are ~1.2 cm wide at the base, are entire margined where visible, and have open and strong divergent venation with rare but apparent vein dichotomies.
Description Foliar remains assignable to the Carboniferous–Permian genus Angaropteridium Zalessky occur in moderate abundance in some plant beds in the Mount Dall conglomerate. These fossils occur as isolated pinnules and this fact limits their taxonomic assignment. Specimens show arcuate, dichotomizing cardiopteroid venation, but instead of the foliar venation deriving from a single entry vein (Cardiopteridium), venation originates from a petiole where a wide vein or multiple (5–10) veins enter the pinnule. Pinnules have markedly auriculate bases and entire margins with an overall ovoid shape. Specimens measure 8–12 mm wide and 10–16 mm long on average. The insertion of multiple veins into the pinnule and its overall shape provide basis for the assignment of these specimens to the genus Angaropteridium. The form genus Cardioneura Zalessky is a possible alternative assignment for some specimens, as this genus is defined by lack of a midrib and the restricted basal morphology similar to that of neuropterids. Cardioneura, however, possesses stronger venation in the midleaf and near the pinnule petiole than these specimens exhibit. Assignment of these specimens to the genus Neuropteris is not possible either, judging from the basal venation patterns and the lack of a midleaf vein cluster from which veins arc toward the margin. These specimens are preserved as both impressions and compressions. Intervals: 7–8 Specimens: AK0057A, AK0101E, AK0103B, AK0108D, AK0109B, AK0112B, AK0116A, AK0117C, AK0118A.
Genus Zamiopteris Linear-lanceolate symmetrical leaves with acute apices, entire margins, and open venation from the Mount Dall conglomerate are assigned to the genus Zamiopteris Schmalhausen by Mamay and Reed (1984) (their figs. 62G and 62I). These leaves lack a midrib but rather possess a clear central region of preferential vein dichotomy from which venation arcs toward intersection with the leaf margin at acute angles. Figured specimens in Mamay and Reed (1984) show vein densities of 16–18 veins/cm and are maximally 20 mm wide. Although the pecopterids and cyclopteroids provided little age control (late Paleozoic–early Mesozoic) and limited resolution into the phytogeographic affinity of the flora, the presence of Zamiopteris provided the first biostratigraphic constraint on the deposition of the Mount Dall conglomerate as being Permian in age (Mamay and Reed, 1984). Because Zamiopteris is known only from Permian Siberia, Kazakhstan, and other associated northern Pangean landmasses (with the later exception of a single occurrence reported from Venezuela [Ricardi et al., 1998]), its occurrence here also suggested that the Farewell terrane had a phytogeographic affinity to the Angaran Province sensu lato in the Late Paleozoic (Mamay and Reed, 1984). That the same genus provides both an age and the biogeographic affinity of the flora is not ideal and may lead to misinterpretations of depositional timing and regional phytogeographic variation. However, the Siberian aspect of at least the genus Zamiopteris does provide the youngest biogeographic datum for the Paleozoic of the Farewell terrane and this both refutes and corroborates previous conclusions based on older marine biotas (Table 1). NEW MACROFLORAL COLLECTION All genera originally reported by Mamay and Reed (1984) were again obtained from the Mount Dall conglomerate in the summer field seasons of 2000 and 2001. Additionally, putative sphenophytes were found in plant hash deposits in medium to course sands but are not distinguishable beyond the class level of taxonomy. As before, no lycopsids were found from the sediments of the Mount Dall conglomerate. The following genera were identified from new collections of the Mount Dall flora and are to be deposited among the collections of the Denali National Park and Preserve Museum, Alaska, USA (accession #564).
Cordaites sp. (Fig. 7E) Description Parallel-veined and parallel-margined specimens lacking dorsal furrows and exhibiting subordinate vein fibers (“false veins” [Meyen, 1987]) between the main veins fit within the description of the genus Cordaites. Specimens of this genus do not exhibit intersection of veins with the lateral margins of the leaf. These specimens are preserved in coaly dark shales as leaf mats suggesting a deciduous habit. Their typical width is 18–24 mm with ~25 veins/cm. All specimens of this genus that were recovered are fragmentary, and thus a measurement of the leaf length is impossible. However, a minimum length estimate of many specimens is ~15 cm. Intervals: 4–5, 7–8 Specimens: AK0001A, AK0001C, AK0011A, AK0011B, AK0015A, AK0018A, AK0021C, AK0029B-D, AK0046A-B, AK0062A, AK0064B, AK0095B, AK0096B, AK0099B, AK0101B-C, AK0110A, AK0111A, AK0115B. Cyclopteris sp. (Fig. 7A) Description Similar to Angaropteridium, Cyclopteris Brongniart pinnules in the Mount Dall conglomerate measure ~16–18 mm long
A
B D
C
E
F
Figure 7. Foliar plant remains from the Mount Dall flora. (A) Cyclopteris sp. (AK0119). (B) Pecopteris cf. hemitelioides (AK0020). (C) Angaropteridium sp. (AK0116). (D) Pecopteris cf. arborenscens (AK0086). (E) Cordaites sp. (AK0062). (F) Rufloria sp. (AK0105). View of stomatiferous (“dorsal”) furrows in relief. Scale for all images: 1 cm.
Flora, fauna, and sediments of Mount Dall Conglomerate and ~10–18 mm wide, with entire margins and dichotomizing open venation that lacks a midleaf vein cluster. Their bases are restricted, often show evidence of being apetiolate, and are occasionally weakly chordate to auriculate. These specimens could be aphlebiae of Angaropteridium (alt. Abacanidium) fronds, more specifically assignable to Cardioneura, or perhaps Zamiopteris leaves at early foliar growth stages. Intervals: 5–8 Specimens: AK0020B, AK0029E, AK0032A, AK0032C, AK0034A, AK0061B-C, AK0091B, AK0109E-F, AK0119A-D, AK0120A-B, AK0122A, AK00123C. Pecopteris cf. arborescens (Fig. 7D) (alt. P. arborenscens) Description This species of Pecopteris is characterized by small lengthto-width ratios of the pinnules (1.8 versus >2.5 in P. hemitelioides), close spacing of the pinnules along the secondary rachis, acute termination of the ultimate rachises, and pinnule insertion into the rachis at right angles. P. arborescens was identified in the collection of Mamay and Reed (1984), and similar specimens occur in these new collections. The details of pinnule venation are not visible and therefore assignment to this species is uncertain (see Zodrow, 1990). Interval: 7, 9 (8?) Specimens: AK0035A, AK0086A, AK0087A, AK0093A. Pecopteris cf. hemitelioides (Fig. 7B) (alt. P. hemiteloides) Description Pecopteris hemitelioides was identified by Mamay and Reed (1984) and recovered again in the field seasons of 2000 and 2001. Pinnules are crenulate, inserted on the secondary rachis at high angles, and show simple singular dichotomies in the pinnule venation in the best-preserved specimens. Some specimens may belong to the better known species P. unita with its strong pinnule venation, but that this venation is not distinct on these specimens inhibits this assignment. Interval: 4–7, 10 Specimens: AK0013A, AK0016A, AK0017A, AK0020A, AK0032B, AK0033A, AK0046C, AK0050C, AK0052A, AK0058B, AK0059A, AK0069A, AK0070A, AK0073A, AK0075A, AK0080A, AK0088A, AK0090A, AK0094A, AK0097A.
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Interval: 3–5, 8 Specimens: AK0006B, AK0009A, AK0021B, AK0041A, AK0044A, AK0067A, AK0100A-B, AK0105A, AK0114B, AK0111C? Trigonocarpus sp. (Fig. 8A) Description A three-dimensionally preserved partial seed that is 22 mm long and 9 mm wide was recovered from sandy sideritic sediments of the Mount Dall conglomerate and can be assigned to the long-ranging and cosmopolitan genus Trigonocarpus. This specimen is one-third of the original seed split along its dehiscence (angle of dehiscence faces is ~100°–140°). It possesses an acute apex, a short stalk attachment, weak longitudinal ribs, and internal pits. Specimens are preserve in three dimensions by siderite replacement. Interval: 9 Specimens: AK0025, AK0037A? Zamiopteris sp. (Fig. 8B) Description Many specimens with a central midleaf vein concentration and coarse venation gently arching to an entire margin occur in the Mount Dall conglomerate. The leaves are lanceolate in overall shape and are up to 50 mm wide and 140 mm long in fragmentary specimens. No complete leaves are preserved of this type but overall lengths may have exceeded 300 mm. These specimens have an average venation density of ~15 veins/cm and are assignable to the genus Zamiopteris, consistent with material previously described by Mamay and Reed (1984) as well as similar material assigned to this genus (Meyen, 1982; LePage et al., 2003). Leaves often occur in dense monotypic leaf mats and, like Cordaites and Rufloria, may be deciduous. Interval: 3–5, 7–8 Specimens: AK0001E, AK0006C, AK0021A, AK0040A, AK0042A, AK0043A, AK0045A, AK0047A, AK0048A, AK0049A-B, AK0050A-B, AK0051A-B, AK0057B, AK0059B, AK0066A, AK0071A, AK0099A, AK0108B, AK0109A, AK0109C, AK0110B, AK0110D, AK0111D, AK0112A, AK0115A, AK0118B, AK0121A, AK0123A TAPHONOMY AND PRESERVATION OF PLANT REMAINS
Rufloria sp. (Fig. 7F) Description The presence of dorsal furrows in specimens with parallel venation provides the basis for their assignment to Rufloria (Meyen, 1982, 1987). These leaves range from 15 to 30 mm wide and up to 120 mm long though no full leaf impressions were obtained. The leaves are entire margined, have a vein density of 14–26 veins/cm, and occur in leaf mats suggesting a deciduous life habit.
Foliar, axial, and root remains in the Mount Dall flora are most frequently preserved as impressions in siltstones and organic-rich mudstone beds. Occasional compression preservation is noted in cordaitalean (Cordaites, Rufloria, Zamiopteris) foliage, whereas siderite-replacement is common among the more refractory elements such as woody tissues and seeds, especially in Interval 9. Concentrated assemblages of parallel- and open dichotomous-veined leaves of seed plants occur in thin,
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monotypic “leaf mat” lenses in local low-energy fluvial settings, where mud accumulated in what are presumed to be small-scale point bars or standing-water deposits adjacent to, but separate from, the main flow in braided streams (Fig. 4F). Fern foliage is not found in concentration but rather as dispersed frond fragments in slightly coarser siltstones more proximal to the flow and in locally higher energy deposits. In a few instances the preservation of cordaitalean foliage shows small-scale leaf folding both parallel to and across the venation pattern. At first glance, this
Figure 8. Foliar, seed, and faunal remains from the Mount Dall conglomerate. (A) Trigonocarpus sp. (AK0025) (scale, 0.5 cm). (B) Zamiopteris sp. (AK0049) (scale, 1 cm). (C) ?Stenoscisma (Rhynchonellida) (AK0007) (scale, 1 cm). (D) ?Schuchertella (Strophomenida) (AK0036). Scale for all images: 1 cm.
suggests that at least some leaves were constructed of a rather thin lamina in life much like leaves that are for seasonal use only (deciduous). I cannot rule out sediment compaction as a cause for this observation, however. Seeds are concentrated with small stem and wood fragments in coarser-grained sandstones, where their resistant tissues are often replaced by siderite and preserved three dimensionally. There is no discernible parallel orientation of the long axes of any fossil material in any subconglomerate facies. Occasional root traces are apparent and disturb fine
Flora, fauna, and sediments of Mount Dall Conglomerate laminations within siltstones and mudstones. These root traces average <0.25 cm in diameter, are not associated by attachment with foliar or axial remains and, where observed, are present across the full lateral extent of siltstone-mudstone lenses (Fig. 4E). The taphonomy of the fossil materials, the depositional environment, and the local stratigraphic relationships of bedding within the Mount Dall conglomerate all show local variation in depositional setting typical of that within a braided-stream system. These variations in both the horizontal and vertical dimensions reveal the spatial heterogeneity of ecological and taphonomic environments within this unit. FAUNAL REMAINS Syndepositional marine faunal remains are infrequent and poorly preserved in the Mount Dall conglomerate. The few observed and collected specimens of brachiopods and gastropods were concentrated in the middle of the measured section within Interval 9 and in medium to coarse sandstones. Their occurrence does not appear to be the result of redeposition to judge by the condition of the shells and the fact that there is no lithological difference between the surrounding matrix and the fossil infill. Conodonts are also known from the section but are secondarily deposited via the limestone cobbles in the massive conglomerate beds (Bradley et al., 2003). These brachiopods and poorly preserved putative gastropod fossils together constitute the entire fauna of possible syndepositional age known from the sediments of the Mount Dall conglomerate. ?Stenoscisma sp. (Rhynchonellida) (Fig. 8C) Description A single articulated rhynchonellid brachiopod was recovered measuring 10 mm wide by 9 mm high. The specimen is preserved as an internal and external mold, shows strong ribbing, and is moderately sulcate. On the basis of its morphological attributes, assignment to Stenoscisma is only tentative (Jin Yugan, 2001, personal commun.). Although the genus ranges through the Tethyan and Boreal Provinces in the Permian (Stehli, 1971), this specimen most resembles the Early Permian species S. hueconianum Girty and S. venustum Girty found in North America (Grant, 1965). However, strong angular costae on the pedical valve resemble those of the Early Permian (Sakmarian–Artinskian) genus Septacamera Stepanov, and this is another possible assignment. Septacamera ranges from Svalbard and the Urals west into Arctic North America (Grant, 1971) but is usually twice the size of this specimen. Interval: 9 Specimen: AK0007. ?Schuchertella sp. (Strophomenida) (Fig. 8D) Description A poorly preserved articulated strophomenid brachiopod measuring ~22 mm wide was recovered from the same sandy sediments in Interval 9 that contained ?Stenoscisma sp. The spec-
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imen is partially preserved as a sideritic crust of the exterior. On the basis of its outline in impression, its weakly sulcate and strongly ribbed shell, and the characteristics of its hinge, it can be assigned tentatively to the cosmopolitan and long-ranging (Silurian–Permian) genus Schuchertella. Other clearly strophomenid brachiopods with similar shell ribbing occur within this interval but are too poorly preserved to determine to the genus level. Interval: 9 Specimen: AK0036. DISCUSSION Age of Deposition No radiometric ages are yet available for the deposition of any of the late Paleozoic rocks of the Mystic sequence including the Mount Dall conglomerate. Unfortunately, samples of coarseand fine-grained material analyzed for palynomorph remains yielded nothing, apparently because of overbaking (J. Utting, 2002, personal commun.). An Early Permian age is assigned here to the Mount Dall conglomerate on the basis of three lines of biostratigraphic evidence. The first, which admittedly is based on a somewhat circular argument in the context of the phytogeographic discussion and conclusions below, is the presence of late Paleozoic macrofloral forms (including the, as far as we know, exclusively Permian Zamiopteris foliage). Second is the occurrence of the rhynchonellid brachiopod genus Stenoscisma-form that most closely resembles the Early Permian S. hueconianum. The third line of biostratigraphic evidence, which provides a lower age constraint on the unit, is based on Late Pennsylvanian–Early Permian conodonts extracted from carbonate cobbles (Bradley et al., 2003). These macro- and microfossil remains constrain the oldest possible age of the Mount Dall conglomerate and its flora to be most likely earliest Permian. An upper age constraint within the Permian is not firmly set by these biostratigraphic data. However, it is presumed to be no later than the Middle Permian (Roadian– Capitanian), owing to the floral assemblage components (specifically the Pecopteris species identified by Mamay and Reed [1984]) and the fact that the most-similar rhynchonellid species within the genus Stenoscisma are Early Permian in age (S. hueconianum and S. venustum). Paleoenvironmental Interpretation The Mount Dall conglomerate was likely deposited in a setting that combined an alluvial fan and a fluvial braidplain environment where ruderal (disturbance-tolerant, sensu Grime (1977) vegetation existed on ephemeral, lens-shaped islands in a highenergy fluvial plain (Fig. 9). The fact that there is at least one marine incursion, as evidenced by the occurrence of primarily preserved marine fossils, suggests that the unit was associated with an alluvial fan delta in a coastal environment. The floral community growing within the depositional regime is ecologically or successionally immature and consists of ferns and
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Mystic subterrane sequence & Dillinger subterrane units
Mystic Sequence Flysch
Figure 9. A landscape reconstruction of the depositional environment of the Mount Dall conglomerate. Braidplain vegetation heterogeneity from distal open-canopy seed-plant cover to ruderal fern-dominated vegetation on small islands within the braided stream. (Cordaitalean symbols from Willis and McElwain, 2002.)
Angara
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pteridosperms. Similarly, the vegetation appears to be structurally simple, lacking a long-lived canopy and having no multilayered stratification and probably no large trees at all. As a result, the soil horizons are either nonexistent or poorly developed, consisting of only rare laterally discontinuous horizons that penetrate into a shallow water table. Lateral to a continually restructuring main flow was an area of more stable substrate conducive to, and perhaps the result of, a more locally dense, open canopy vegetation where a flora of cordaites, pteridosperms, and ferns persisted in situ through many generations. Although not directly observed in the Mount Dall sediments, mature broad-leaved cordaitaleans may have thrived nearby and shed their leaves annually to produce the observed leaf mats that are often preserved in the low-energy regions of the stream system in abandoned channel fill (as in Darby et al., 1990). What remains of unclear significance are the periods of deposition of sediments with increased grain size, presumably
Figure 10. Orthographic projection above 45°N on the Sakmarian reconstruction of Ziegler et al. (1997) and Rowley et al. (2003). Star symbol indicates projected placement of the Farewell terrane in the Early Permian and arrows show phytogeographic links both to the Angaran Province of Siberia and peri-Siberian landmasses and to the Euramerican Province of tropical mainland Pangea.
Flora, fauna, and sediments of Mount Dall Conglomerate deposited under high-energy conditions as evidenced by the horizontally traceable conglomerate beds. I propose that these beds are the result of coarse-grained sheetflood events and gravity-driven debris flows of large (order 10 km2) horizontal extent. The finergrained mesoscale (~5 m vertically, ~10–15 m horizontally) fluvial sedimentary structures between these conglomeratic units can be explained by the proposed depositional model. The dynamical origin of the regional cobble-boulder incursions most likely reflects times of increased upstream erosion, perhaps due to decadal to millennial climate fluctuations or to tectonic changes in base-level (as in Dorsey and Umhoefer, 2000) and related gravity-flow deposits.
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The entire plant assemblage contains forms typical of cratonic North America as well as some previously thought to be endemic to Siberia and the Angaran Floristic Province. This combination suggests a mixed phytogeographic affinity (Fig. 10). Unfortunately, a variety of lithologies from the Mount Dall conglomerate failed to yield suitable palynomorphs for either biostratigraphic data or biogeographic information (J. Utting, 2002, personal commun.), but new collections of samples from elsewhere within the field area or in other understudied localities of the Mystic subterrane sequence may provide these data for their comparison with syntheses of the microfloral biogeography of the Permian temperate regions (Utting and Piasecki, 1995).
Phytogeographic Importance CONCLUSIONS The floral remains at Mount Dall are the only Permian plant megafossils that have been found within a ~1500 km radius. The closest coeval assemblages occur in the southwest contiguous United States, the Verkhoyansk Region of eastern Russia, and a pair of localities on Axel Heiberg Island, Canada, and North Greenland (LePage and Pfefferkorn, 1999; LePage et al., 2003; Wagner et al., 1982; Wagner et al., 2002). If this flora was deposited in a zone of plate convergence offshore of the western Cordilleran margin of North America as new geological evidence suggests (Bradley et al., 2003), then its occurrence in isolation from other known floral zones is of great significance for global phytogeographic reconstructions and analyses. With regard to the biogeographic affinities of individual taxa appearing in the Mount Dall conglomerate, Zamiopteris is known almost exclusively from the Permian of the Siberian and peri-Siberian region (the Angaran Phytogeographic Province, which includes Kazakhstan, the Russian Platform, Mongolia, etc. [Meyen, 1987; Ziegler, 1990]) and isolated reports from the Korean Peninsula and Venezuela (Ricardi et al., 1998). The cordaitalean genus Rufloria is an Angaran and Sub-Angaran (southern Angaraland) leaf form (Meyen, 1982) that has, as of yet, only one reported occurrence outside of the Angaran Floristic Province and its immediately neighboring paleogeographic blocks (in the Kungurian of the Sverdrup Basin, Arctic Canada [LePage et al., 2003]). Cyclopteroid forms assignable to the genus Angaropteridium (sensu Meyen 1982, 1987) have not been reported outside of the Angaran region. The remaining genera (Cordaites, Pecopteris, Trigonocarpus, Cyclopteris) in the Mount Dall flora are more cosmopolitan, with paleolatitudinal ranges that extend from the temperate regions of Angaraland southward into the Euramerican and Cathaysian Phytogeographic Provinces of the Early Permian and, in some cases, Gondwana at southern temperate paleolatitudes. For the latter wide-ranging genera as well as those that are endemic to the Angaran-Subangaran Phytogeographic Provinces, species-level provincialities have been proposed (e.g., Meyen, 1982). Unfortunately, the preservation of the material from Mount Dall restricts phytogeographic discussions to the generic level. Perhaps more sampling from the intervals and beds of best preservation would allow for confident species-level identifications, which could be even more phytogeographically informative.
Two facies associations in the Mount Dall conglomerate are described on the basis of fieldwork in the southwestern Alaska Range in Denali National Park and Preserve. Interbedded massive cobble conglomerates and subordinate mudstone-gravel facies were deposited in a fluvial braided stream succession on the coast of a developing foreland basin. Discreet intervals of normal gradation within these faces are noted from the measured section. Provenance analysis of conglomerate clasts suggests that they were derived from lower stratigraphic units in the Farewell terrane and specifically from older sedimentation elsewhere in the Mystic and Dillinger subterrane sequences. Point counts through the measured section indicate a change in the lithological composition of the dominant conglomerate beds and suggest a gradual change in clast provenance upsection. Fossil plants preserved as impression-compression remains with subordinate siderite replacement are Early Permian in age and reflect mixed phytogeographic similarity to both the Angaran (north temperate Pangea) and Euramerican (tropical and subtropical Pangea) Floristic Provinces. Additional age constraints come from the occurrence of a likely Early Permian rhynchonellid brachiopod preserved within the unit and latest Pennsylvanian– earliest Permian conodonts occurring in the carbonate cobbles of the massive conglomerate beds. These foreland basin sediments represent the sedimentological manifestation of the Browns Fork orogeny and the coeval Dall Basin (Bradley et al., 2003). ACKNOWLEDGMENTS Funding from the Geological Society of America John T. Dillon Alaska Research Award, the Society for Sedimentary Geology, the Paleontological Society, the U.S. Geologic Survey, and the Department of Geophysical Sciences at the University of Chicago made this work possible. Logistic support was provided by Talkeetna Air Taxi and USGS-Anchorage. I would like to thank T.A. Rothfus, D.C. Bradley, and J.A. Dumoulin for field assistance and D.C. Bradley, A.M. Ziegler, P.M. Rees, J.C. McElwain, W.A. DiMichele, D.B. Rowley, J. Yugan, C.K. Boyce, S.M. Kidwell, and H. Pfefferkorn for discussions and laboratory assistance.
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mental changes—Quaternary, Carboniferous–Permian, and Proterozoic: New York, Oxford University Press, p. 111–146. Zodrow, E.L., 1990, Revision and emendation of Pecopteris arborescens group, Permo-Carboniferous: Palaeontographica, Abt. B, v. 217, p. 1–49. MANUSCRIPT ACCEPTED BY THE SOCIETY 14 DECEMBER 2007
Printed in the USA
The Geological Society of America Special Paper 442 2008
Late Triassic silicified shallow-water corals and other marine fossils from Wrangellia and the Alexander terrane, Alaska, and Vancouver Island, British Columbia Andrew H. Caruthers* George D. Stanley Jr.* The University of Montana, 32 Campus Drive, Missoula, Montana 59812, USA ABSTRACT Study of Late Triassic biofacies and associated paleoecology reveals new silicified shallow-water corals and other fossils from new and previously known localities within the Alexander terrane (Keku Strait and Gravina Island, southeast Alaska) and Wrangellia (Wrangell Mountains, Alaska, and Vancouver Island, British Columbia). Twenty-five species of coral are identified from eight localities within the Alexander terrane and 34 species are identified from four localities in Wrangellia. Distributions of silicified shallow-water marine fossils contribute to Late Triassic (Norian–Rhaetian) paleoecology, biotic diversity, and terrane paleogeography. Depositional environments establish the conditions in which these organisms lived as well as provide evidence for lithological correlation between tectonically separate fragments. This study also confirms the presence of biostrome reef buildups in the southern Alexander terrane (Gravina Island), indicating warm, clear, and nutrient-free water with lots of sunlight; this differs from the central Alexander terrane (Keku Strait) and northern Wrangellia (Wrangell Mountains), where corals grow as individual colonies, not in a structured, reef-like buildup, and are accompanied by filter- and detritus-feeding organisms indicating warm, cloudy and nutrient-rich water in a back-reef environment. Paleobiogeographic results from silicified Upper Triassic corals show faunal similarity between Gravina Island and Keku Strait (Alexander terrane) and no similarity between northern and southern Wrangellia. Likewise, no similarity was found between the Alexander terrane and either northern or southern Wrangellia. Keywords: Triassic, Alexander terrane, Wrangellia, scleractinian corals, paleobiogeography.
*E-mails: Caruthers:
[email protected]; Stanley:
[email protected]. Caruthers, A.H., and Stanley, G.D., Jr., 2008, Late Triassic silicified shallow-water corals and other marine fossils from Wrangellia and the Alexander terrane, Alaska, and Vancouver Island, British Columbia, in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 151–179, doi: 10.1130/2008.442(10). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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INTRODUCTION Alexander Terrane The Alexander terrane is located throughout much of southeast Alaska and parts of western British Columbia, Canada; Wrangellia, on the other hand, is recognized as occurring in south-central Alaska, Vancouver Island, and other scattered islands along coastal British Columbia, Canada (Berg et al., 1972; Jones et al., 1972; Gehrels and Saleeby, 1987; Jones et al., 1977). They are two of many displaced terranes that constitute the Cordilleran Region of western North America (Fig. 1). Prior to accretion, Wrangellia and the Alexander terrane are considered to have been island-arc systems located in the Panthalassan Ocean (Coney et al., 1980). However, stratigraphic successions within each terrane suggest the two terranes had considerably different tectonic histories throughout the
Phanerozoic (Berg et al., 1972; Jones et al., 1972). The Alexander terrane contains volcanic and sedimentary stratigraphy that is Cambrian through Triassic age, whereas Wrangellia contains only Carboniferous through Jurassic stratigraphy. Both terranes, however, have comparable Upper Triassic reef-like assemblages surrounding them (Gehrels and Saleeby, 1987; Gardner et al., 1988; Stanley, 1993; Soja, 1996). The Alexander terrane is considered to be a displaced continental fragment owing to its foundation of Late Proterozoic and Early Paleozoic continental crust (Gehrels, 1990; Wilson, 1968). Furthermore, the Alexander terrane is unique because it contains sedimentary rocks from every Phanerozoic system, making it ideal for study and interpretation. The Alexander terrane was a separate entity throughout much of its Paleozoic history (Wilson, 1968; Monger and Ross, 1971; Jones et al., 1972; Monger et al., 1972; Hillhouse and Gromme, 1980). It contains similar-age rocks with lithologic, paleomagnetic, and paleobiogeographic properties distinct from adjacent Wrangellia (Jones et al., 1977). Wrangellia
Figure 1. Generalized map of western North America showing approximate positions of Wrangellia and the Alexander terrane (after Jones et al., 1972; Jones et al., 1977; from Katvala and Stanley, this volume).
In contrast to the Alexander terrane, Wrangellia is composed principally of fragmented blocks separated from each other by nearly 200 km distance along the western coast of North America (Jones et al., 1977). The type section, or northern block, is located in the southern Wrangell Mountains of south-central Alaska, whereas the southern block is best documented on Vancouver Island in western British Columbia, Canada (Fig. 1). Two smaller sub-blocks (Fig. 2), located on the Queen Charlotte Islands, British Columbia, and on Chichagof Island, Alaska, have also been tied to this terrane (Jones et al., 1977) but were not included in this study. All four blocks of this terrane are grouped together by internally consistent Triassic lithological sequences (MacKevett, 1976; Jones et al., 1977) (Fig. 2). The lithological sequence of Triassic rock in northern Wrangellia begins with a basal unit of bedded chert (Ladinian) overlain by the Nikolai Greenstone, a thick volcanic unit intermixed with pillowed aa and pahoehoe basalt. This unit is disconformably overlain by the Carnian to Norian Chitistone and Nizina Formations, together comprising a carbonate sequence of rocks 1400 m thick (MacKevett, 1976; Armstrong and MacKevett, 1982). Triassic rocks of the southern Wrangellian block (Vancouver Island) have stratigraphic sequences and lithologies similar to those of northern Wrangellia (Jones et al., 1977). Muller et al. (1974) and Yorath et al. (1999) noted that the Triassic sequence of the southern block begins with 200 m of black silicified shale and siltstone intruded by numerous Ladinian sills. This unit is overlain by the Karmutsen Formation, ~6600 m of volcanic rock containing (1) pillow lava, (2) pillow breccia and aquagene tuff, and (3) basalt flows with minor pillow lava and sedimentary layers. The volcanic unit is overlain by Carnian to early Rhaetian carbonate sequences of the Quatsino and Parson Bay Formations (Yorath et al., 1999).
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Figure 2. Correlative stratigraphic columns showing Middle–Upper Triassic rock sequences for localities within the recognized borders of Wrangellia (modified from Jones et al., 1977).
Triassic Fossil Investigations Prior to plate tectonic reconstructions, Smith (1927) recognized Upper Triassic shallow-water marine invertebrate fossils from a variety of localities spanning the western coast of North America and Canada. Smith concluded that fossils from rock now regarded as Wrangellia and the Alexander terrane were comparable in age. They, along with rocks from other terranes, were thought to represent fringing reefs spanning the entire western coast of North America. Later Tozer (1967) and Stanley (1979) clarified the situation by putting Smith’s Triassic coral and spongiomorph reef occurrences within the emerging concepts of terrane theory, thus explaining high-latitudinal anomalies of their occurrence. Subsequent paleontologic, lithologic, structural, and paleomagnetic work
has separated these Triassic limestones into many different terranes, including Wrangellia and the Alexander terrane, by interpreting their tectonic and stratigraphic histories as well as post-Triassic accretion onto the North American craton. Gardner et al. (1988) interpreted a Pennsylvanian pluton cutting across the boundary between Wrangellia and the Alexander terrane. This interpretation indicates amalgamation of the two terranes in the Carboniferous. Other criteria needed for terrane amalgamation consist of having inherently similar geology, paleomagnetism, and faunal composition including diversity and ecology (paleobiogeography). Paleobiogeographic results generated in this study support findings by Blodgett and Frýda (2001, p. 240–241) indicating no similarity in marine fauna between these two terranes in post-Pennsylvanian strata, despite the fact
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that both terranes contain Mesozoic and Paleozoic fossils such as bivalves, corals, sponges, spongiomorphs, brachiopods, gastropods, conodonts, and fusulinids. Having a wealth of invertebrate fossils offers independent ways to test the amalgamation of Wrangellia and the Alexander terrane in post-Carboniferous rocks. Marine organisms offer excellent potential for biogeographic reconstruction analysis (Belasky and Runnegar, 1993). Paleobiogeographic Studies Previous paleobiogeographic study of rugose corals suggests that during the Permian, Wrangellia was situated 5000 km west of the craton (Belasky and Runnegar, 1994). However, a more in-depth study using rugose corals, brachiopods, and fusulinids infers that during the Permian, Wrangellia and the Alexander terrane were situated ~2000–3000 km west of the North American craton (Belasky et al., 2002). Newton (1983) conducted a paleobiogeographic study of Late Triassic bivalves from Wrangellia and the Alexander terrane. Through analysis of similarity coefficients, she found Wrangellia to occupy an eastern Pacific location and to be statistically similar to Late Triassic (early Norian) faunas from Hells Canyon and cratonal North America. In comparison, the Alexander terrane was found to occupy a more southern position during Late Triassic time having a loose similarity with Triassic (Norian) bivalves from Peru (no statistical analysis conducted, just one bivalve found to be similar). Aberhan (1999) also conducted a paleobiogeographic study using the Early Jurassic bivalve Weyla. He determined that Wrangellia has been displaced northward by at least several hundred to possibly >1000 km since Early Jurassic time. Katvala and Henderson (2002) studied Pennsylvanian– Permian sequence stratigraphy of conodonts from Vancouver Island (southern Wrangellia) and found fauna that indicates this part of Wrangellia was located in cooler waters, suggesting a more temperate climate ~25°N paleolatitude during this time period. Further paleobiogeographic study of Late Triassic corals by Yarnell (2000) suggested a high degree of similarity between Wrangellia and the Alexander terrane. Preliminary paleobiogeographic results of Late Triassic corals given by Yarnell (2000) concerning Wrangellia and the Alexander terrane were generated from older literature (i.e., Smith, 1927). Many of the taxonomic identifications made by Smith (1927) need revision, and new taxonomic schemes have been proposed (Cuif, 1965, 1967, 1975a, 1975b, 1976; Beauvais, 1980; Roniewicz, 1989). Furthermore, previous studies of Late Triassic corals from the Alexander terrane (Montanaro Gallitelli et al., 1979; Stanley, 1979) only included a small subpopulation of corals from the western coast of Gravina Island and did not include any material from the Keku Strait area, a region now much better studied (Katvala and Stanley, this volume). This paper studies Late Triassic silicified corals from localities within Wrangellia and the Alexander terrane. Alexander terrane corals were collected from both Gravina Island and Keku Strait in order to provide a larger sample size with updated taxonomic identifications, thus increasing statistical accuracy.
Paleobiogeographic results in this paper were generated from previously collected and new Late Triassic coral populations from many localities within Wrangellia and the Alexander terrane; these corals are systematically described in Caruthers and Stanley (2008). Paleomagnetic Studies Paleomagnetism has been used quantitatively to reveal changing paleogeography of terranes through time. Paleolatitude results for the Alexander terrane yield 14°N ± 4° in the Early Devonian, 25°N–30°N in the Early Permian, and 10°N–23°N in the Late Triassic (Hillhouse and Gromme, 1980; Butler et al., 1997). To contrast, Panuska (1984) revealed a Southern hemisphere location for the Alexander terrane from Paleozoic through Late Jurassic time. Investigations by Hillhouse (1977), as well as Irving and Yole (1972), indicate low-lying paleolatitudes (within 15° of the paleoequator) for both northern Wrangellia (Wrangell Mountains) and southern Wrangellia (Vancouver Island) during Triassic time. Stone (1981) supported these findings by concluding that northern Wrangellia contained a paleolatitude close to and probably north of the equator during Triassic time. Panuska (1984) suggested that Wrangellia began to move into the Southern Hemisphere during Late Triassic or Early Jurassic time, and concluded an amalgamation with the Alexander terrane at mid to low southern paleolatitudes during Late Jurassic time. Together these data suggest Wrangellia has moved several thousand kilometers northward relative to North America (also shifting) since Triassic time, whereas the Alexander terrane did not move as far northward during post-Triassic time relative to Wrangellia which contains low (equatorial) paleolatitudes (Hillhouse and Gromme, 1980; Panuska, 1984). Paleomagnetic results for Wrangellia also correspond to Jurassic paleobiogeography presented by Aberhan (1999), which suggests northward movement out of the tropics during post-Jurassic time. Geologically, these results were affirmed by Jones et al. (1977) with the recognition of distinctly different Triassic sequences between Wrangellia and the Alexander terrane, thus directly challenging the idea that these two terranes were amalgamated during Carboniferous time. However, these results conflict with the preliminary data of Yarnell (2000), which suggested statistically similar Late Triassic coral assemblages from Wrangellia and the Alexander terrane. Reefs We define the term “reef” as an organic buildup of corals, sponges, and spongiomorphs in the studied Upper Triassic units, similar to the definition of Jackson (1997). Stanley (2001) reviewed the original definition, emphasizing a fourfold concept of the term—from a stratigraphic reef, to “mud mounds,” degraded reef, and ecologic reef. The reef concept is strongly rooted in biological characteristics, being centered on complex biological and ecological interaction, topographic relief, and
Late Triassic silicified shallow-water corals and other marine fossils wave resistance within a shallow-water setting. The term “stratigraphic reef” is a descriptive term pertaining merely to a thickened mass of carbonate differing appreciably from surrounding rocks (Stanley, 2001). Carbonate reefs have been recognized and their characteristics plotted throughout the Phanerozoic (Kiessling and Flügel, 2002); however, the quintessential framework of these reefs is highly variable, incorporating a multitude of organisms including cyanobacteria, stromatolites, archaeocyathids, crinoids and blastoids, stromatoporoids, receptaculitids, tabulate, rugose, and scleractinian corals, rudistid bivalves, brachiopods, coralline algae, bryozoans, sponges and spongiomorphs. The framework concept has been challenged by Hubbard et al. (2001) who remarked on the taphonomy of Holocene reefs, the framework of which is frequently transformed into rubble by physical and biological processes. Reefs (modern and ancient) will flourish best under conditions of optimal nutrients, sunlight, water temperature, and depth, and degrade under conditions of higher environmental stress, including increased sedimentation or volcanism (Stanley, 2001; Zonneveld et al., 2002; Flügel, 2002). Therefore, the presence of a reef indicates a multitude of environmental (biological and sedimentological) conditions. Equally important in reef ecosystems is photosymbiosis, especially when established between zooxanthellae and their coral hosts (Wood, 1999). During the Late Triassic (Norian–Rhaetian), the paleotropics extended from >30°N to >35°S of the paleoequater, and scleractinian corals, spongiomorphs, and calcified sponges were among the major reef builders in the marine realm (Flügel, 2002). This produced a wide, warmer, and more arid tropical to subtropical band across Late Triassic Earth, thus promoting the development of large-scale framework reefs especially in the former Tethys Region. Some Norian examples left deposits over 2000 m thick (Flügel, 2002). These thick Norian reefs grew along rift-shoulders of the Tethys seaway across what is now central Europe and Asia (Flügel, 2002). Small-scale reefs and buildups sporadically occurred surrounding the continental fragments and volcanic island arcs situated within the eastern Tethys seaway and Panthalassa Ocean (Stanley, 2001; Zonneveld et al., 2002; Flügel, 2002); particularly within certain terranes of the North American Cordillera (Quesnellia, Stikinia, and the Wallowa terrane), where Tethyan-type reef development has been recorded. This paper describes a new small-scale reef buildup from the Alexander terrane, with little evidence of reefs from Wrangellia. Zonneveld et al. (2002) suggested that higher rates of sedimentation and increased volcanism within the North American Cordilleran terranes were most likely the cause for an absence of reef ecosystems. Zonneveld et al. (2002) as well as Sanders and Baron-Szabo (2005) reported that in the Mesozoic, turbid-water bioconstructions, dominated by scleractinian corals, accumulate in conditions of high nutrients and siliciclastic input. This paper discusses corals and the situation with reefs in both the Alexander terrane and Wrangellia. Criteria for reef classification include an in situ framework of
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organisms displaying definite structure within a laterally confined carbonate rock sequence. Fossil Preservation Quality of preservation in recovered fossils and processing techniques limit accurate identification and associated paleobiogeographic designations. Modern schemes of Late Triassic higher-level (family and genus) coral taxonomy and the resulting revisions of Norian to Rhaetian taxa proposed by Cuif (1965–1976), Beauvais (1980), and Roniewicz (1989) group corals into higher taxa by differences in the microstructural arrangement of their aragonitic skeletons. Applying this scheme is difficult because coral faunas of this study have undergone some degree of replacement by silica, calcite, or other minerals. Once corals have undergone mineralogical replacement, their original delicate microstructure is obliterated, thus rendering microstructural classification impossible. In order to cope with this difficulty, we identified North American Late Triassic shallow-water corals largely on the basis of traditional morphological comparisons originally made by Frech (1890) and Volz (1896), which Roniewicz (1989) has summarized and updated. Furthermore, our processing techniques used dissolution of silicified corals by acetic acid treatment (as opposed to previous coral studies in the Alexander terrane by Smith [1927] and Stanley [1979] where samples were collected as crack-out material). Etching samples in dilute acetic acid greatly enhances morphological detail and number of specimens recovered, thus enabling more accurate statistical results. Therefore, statistical accuracy of past fossil data from the Alexander terrane was likely skewed by many factors and is in need of revision. Replacement by Silica Variable silicification can increase or decrease the quantity and quality of preserved fossils recovered from etched limestone blocks, consequently destroying morphology and subtle structure vital for proper species identification (Caruthers, 2005; Caruthers and Stanley, 2008). Within silicified corals, over-silicification can (1) destroy detailed corallite walls; (2) lump together thinly bladed septa, destroying their delicate arrangement; (3) produce an entirely diagenetic pseudocolumella mimicking essential characteristics of the species; and (4) obliterate minute ornamentation along septal surfaces (Fig. 3A). Under-silicification of a coral colony usually results in an extremely thin and brittle skeleton having only recognizable colony and corallite shape without diagnostic features such as septa, septal arrangement, dissepiments, or columella (Fig. 3B). When etching limestone blocks that contain under-silicified fauna, care should be taken by using an acetone-based hardener such as Alvar, Butvar, or Vinac. Alvar was used in this study. Replacement by Calcite Non-silicified corals show distinct taxonomic advantages and disadvantages for study. The majority of non-silicified
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Figure 4. Map of southeast Alaska showing geographic locations of the Keku Strait area and Gravina Island (modified from Katvala and Stanley, this volume). Figure 3. (A) Over-silicification in the holotype specimen Kuhnastraea cowichanensis (Clapp and Shimer); note accretion of silica between corallite walls creating illusion of a thick, shared wall between adjacent corallites. (B) Under-silicification of the type specimen Meandrostylis vancouverensis (Clapp and Shimer) from Lake Cowichan (MI 8302), showing basic corallite shape without sepal ornamentation and minimal septal preservation.
corals have undergone replacement by calcite; however, occasionally the original aragonite is still present. Thus etching in acid would obliterate the entire colony. In this type of preservation most three-dimensional colony shapes cannot be determined or need to be inferred from thin sections. Nonetheless, this type of preservation may be advantageous when looking at thin sections of corallites. Thin and polished sections may reveal (1) more detailed morphological differences; (2) minute microstructure including trabecular linkages and septal ornamentation (spines); (3) vesicles along the epitheca; and (4) thickness of septa. The distinct advantages of calcite preservation may result in positive upper-level taxonomic identification using the modern scheme set forth by Cuif (1965–1976), Beauvais (1980), and Roniewicz (1989).
TERRANE LOCALITIES AND FOSSILS Alexander Terrane Keku Strait General History and Age. The Alexander Terrane encompasses many stratigraphic units throughout most of southeast Alaska. Keku Strait and Gravina Island (Fig. 4) lie within the recognized borders of the Alexander terrane (Jones et al., 1972; Berg, 1973; and Muffler, 1967) and were chosen as study areas because of an abundance of silicified, shallow-water marine fossils within Upper Triassic strata. The Keku Strait area contains the best preserved and most complete Triassic section within the Alexander terrane (Muffler, 1967). Katvala and Stanley (this volume) extensively studied and described stratigraphic units in the Keku Strait area, paying close attention to the biostratigraphy and lithostratigraphy of the Paleozoic and Triassic units. Therefore, specific descriptions of the geology and fossils used to determine relative ages of Triassic units in this area have been excluded from this paper.
Late Triassic silicified shallow-water corals and other marine fossils The late Carnian to early Norian Cornwallis Limestone and a limestone interval within the early Norian to Rhaetian Hound Island Volcanics are two units within the larger Triassic Hyd Group (Muffler, 1967; Katvala and Stanley, this volume). These units were chosen for study because of their abundance of silicified shallow-water marine fossils, especially corals. Fossils etched from limestone blocks included corals, sponges, spongiomorphs, the globular hydrozoan Heterastridium, brachiopods, crinoid ossicles, echinoid fragments, branching algae, plant material (wood), stromatolites, conodonts, fish remains, and occasional bone fragments from marine reptiles, as well as a wealth of mollusks including gastropods, large oysters, other sessile and freeliving bivalves, nautiloids, and ammonoids (Fig. 5). Localities. Many localities have been established within the Cornwallis Limestone by Katvala and Stanley (this volume), from which four main localities are discussed (Table 1): (1) Flounder Cove, Montana Invertebrate number (MI) 0099 (= U.S. Geological Survey [USGS] Mesozoic locality M1911); (2) Southwest of Kousk Island MI 0074 (= USGS Mesozoic locality M2136); (3) Big Spruce Island MI 0056 (= USGS Mesozoic locality M2135); and (4) Cornwallis Peninsula East MI 0070 (= USGS Mesozoic locality M1906). Only one relevant locality was established within the thin fossiliferous grainstone unit of the Hound Island Volcanics from the Gil Harbor mudflat MI 0087 (= USGS Mesozoic locality M1912). These sites occur along the Cornwallis Peninsula on Kuiu Island as well as on adjacent islands in the Keku Strait area (Fig. 6). Recovered Corals. The Flounder Cove locality proved to be not only diverse ecologically, but also supported one of the most highly diversified coral faunas within the Alexander terrane (Appendix). Corals belonging to the genera Crassistella Roniewicz, Distichomeandra Cuif, Distichophyllia Cuif, Gablonzeria Cuif, Kompsasteria Roniewicz, Kuhnastraea Cuif, Margarosmilia Volz, Meandrostylis Frech, Pamiroseris Melnikova, Paracuifia Melnikova, Retiophyllia Cuif, and Stylophyllum Reuss were all identified from this locality (Appendix). A locality along the shores of a small island southwest of Kousk Island contained large quantities of corals of low diversity. Multiple specimens from the genera Crassistella Roniewicz, Distichophyllia Cuif, and Gablonzeria Cuif were recovered from etched limestone blocks. These genera are commonly found worldwide in Upper Triassic strata, possibly indicating teleplanic larval dispersion as well as the presence of a highly adaptive species. The colonial cerioid coral Crassistella Roniewicz is highly variable in growth form, seemingly type-specific to certain localities. Corals from this genus are observed growing as hemispherical mounds, encrusting shells, or conglomerate clasts, even growing as large plates suspended in the water column. An encrusting growth form is evident from this locality. Silicification at the small island southwest of Kousk Island proved to be rather poor and most colonies are under-silicified as mentioned earlier. Most are preserved only as colony surfaces, in many instances with the obliteration of internal structure. This is the primary reason encrusting species such as Crassistella juvavica Roniewicz were
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preserved at this locality. Specimens of Distichophyllia norica (Frech) and Gablonzeria profunda (Reuss) were preserved with little significant detail, owing in part to coarse-grained silicification, abundant “beekites,” and modern intertidal algal encrustation. A modern alga on the surfaces of silicified coral is not only difficult to remove, it also seems to destroy much of the delicate detail preserved between adjacent corallites. Specifically corallite walls along colony surfaces that have had algae removed often seem smooth and worn, resembling over-silicification because individual corallite walls cannot be distinguished. Big Spruce Island contained silicification almost identical to the small island southwest of Kousk Island; however, this locality had a higher diversity of coral including species from genera Astraeomorpha Reuss, Crassistella Roniewicz, Distichomeandra Cuif, Kuhnastraea Cuif, and Pamiroseris Melnikova. Large, flat coral colonies growing along the paleo-ocean floor and car-sized spongiomorph colonies (Erik Katvala, 2006, personal commun.) were observed using other spongiomorphs, bivalves and other corals as attachment substrate, thus demonstrating ability for vertical growth or tiering structure within the water column and the possibility for reef-like development (UMIP 302895, 302534). Only two genera of corals (Distichophyllia Cuif and Kuhnastraea Cuif) were recovered from the fourth locality along the east side of Cornwallis Peninsula. Although this locality recorded the lowest overall diversity in corals from Keku Strait, it contained many varieties of sphinctozoid sponges, colonies of the calcified sponge Stromatomorpha californica Smith up to 30 cm, brachiopods, bivalves, and echinoderm spines. Corals were found growing in solitary states within bedded limestone; therefore, no observable structure or reef-like growth was recorded for this particular locality. Erik Katvala (2004, personal commun.) noted many uncollected corals from this locality within the tidal zone, which probably reflects the lower diversity of coral fauna from this locality. The upper Norian to lower Rhaetian Gil Harbor locality contained lower overall coral diversity in comparison to the lower Norian Flounder Cove locality. Several biofacies are present, judging from the extreme variation in the gastropods, corals, and bivalves as well as minor occurrences of the spherical hydrozoan Heterastridium and echinoderms (including plates and spines) within limestone beds from this locality (Katvala and Stanley, this volume) (Fig. 7). Initial gastropod identification by Robert B. Blodgett (2005, personal commun.) included a new genus within the subfamily Astraeinae, Tectus? n. sp. aff. T. interruptum Cox, Planospirina sp., Chulitnacula alaskana (Smith), Andangularia wilsoni, Neritopsis (Wallowiella) n. sp., Cryptaulax aff. C. tilarniocnesis Haas, Chartroniella pacifica (Jaworski), Omphaloptycha jaworski Haas, Ptychostoma sp., Toxoconcha aff. T. gracilis (Haas), and Protorcula? sp. Bivalves are currently being identified by Chris McRoberts. Corals are represented by five genera: Astraeomorpha Reuss, Crassistella Roniewicz, Distichophyllia Cuif, Gablonzeria Cuif, and Meandrostylis (Frech). Coral growth is represented by small mound-shaped and platelike colonies not associated with each other, as well as solitary
Late Triassic silicified shallow-water corals and other marine fossils
Locality
TABLE 1. LOCALITIES WITH RESPECT TO TERRANE/CRATON Sub-Area Terrane/Craton Formation
–
AT
Southeast Alaska
Alexander
–
Early Norian
Montana Invertebrate Number (MI) –
Flounder Cove
FC
Keku Strait
Alexander
Cornwallis Limestone
Late Norian
MI 0099
Gil Harbor Southwest of Kousk Big Spruce
Symbol
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Age
GH
Keku Strait
Alexander
Hound Island Volcanics
Early Norian
MI 0087
SWK
Keku Strait
Alexander
Cornwallis Limestone
Early Norian
MI 0074 MI 0056
BS
Keku Strait
Alexander
Cornwallis Limestone
Early Norian
CPE
Keku Strait
Alexander
Cornwallis Limestone
Early Norian
MI 0070
NB
Gravina Island
Alexander
Nehenta Formation
Early Norian
MI 9930-32
Phocena Bay
PB
Gravina Island
Alexander
Nehenta Formation
Early Norian
MI 9933
Nelson Cove
NC
Gravina Island
Alexander
Nehenta Formation
Early Norian
MI 9935-39
NWR
Wrangell Mountains
Northern Wrangellia
–
Norian
–
GB
Wrangell Mountains
Northern Wrangellia
Chitistone Formation
Early Norian
MI 9934
SWR
Vancouver Island
Southern Wrangellia
–
Rhaetian
–
Cornwallis Peninsula East Nehenta Bay
– Green Butte – Lake Cowichan
LC
Vancouver Island
Southern Wrangellia
Parson Bay Formation
Rhaetian
MI 8302
Pender Point
PP
Vancouver Island
Southern Wrangellia
Parson Bay Formation
Rhaetian
MI 9947 MI 9913
Tahsis Inlet
TS
Vancouver Island
Southern Wrangellia
? Parson Bay Formation
Rhaetian
–
WA
Wallowa Mountains
Wallowa
–
Norian
–
Hells Canyon
HC
Wallowa Mountains
Wallowa
Martin Bridge Limestone
Early Norian
MI 8417
–
PU
Andean Highlands
Peru
–
Norian?
–
PU
Andean Highlands
Peru
Chambara Formation
Norian?
MI 8811
Shalipayco
Note: Symbols used throughout text: AT—Alexander terrane; FC—Flounder Cove; GH—Gil Harbor; SWK—southwest of Kousk; BS—Big Spruce; CPE—Cornwallis Peninsula East; NB—Nehenta Bay; PB—Phocena Bay; NC—Nelson Cove; NWR—Northern Wrangellia; GB—Green Butte; SWR—Southern Wrangellia; LC—Lake Cowichan; PP—Pender Point; TS—Tahsis Inlet; WA—Wallowa Terrane; HC—Hells Canyon; PU—Peru.
corals individually found within mainly coarse-grained limestone matrix (thin section reveals a more conglomeratic matrix; Katvala and Stanley, this volume) and therefore not exhibiting the kind of structured community that would indicate a reef. Specimens of the genus Crassistella Roniewicz exhibit both hemispherical mound shapes and encrusting plate-like growth. The encrusting variety of Crassistella Roniewicz is illustrated in Figure 8. In comparison to all other Alexander terrane localities, the preserved fauna from the Gil Harbor mudflat exhibit the best examples of silicification. Fine-grained silicification helped to preserve minute details such as septal ornamentation including spines in corals, small ribs and growth bands in bivalves, and detailed suture patterns in ammonoids. Pyritization and mudcasting of worm tubes and other macrofauna also helped to
Figure 5. Partially etched limestone block from MI 0099 showing (A) various silicified fossils as they appear following etching in acetic acid (scale bar, 10 cm); (B) diversity of silicified fauna with jumbled mixed nature of fossils (scale bar, 10 cm); (C) partially etched ammonite and gastropod; (D) sphinctozoid sponge Parauvanella sp. identified by Baba Senowbari-Daryan (2005, personal commun.), with annelid worm tubes protruding through matrix; and (E) rejuvenated solitary coral, Distichophyllia norica Frech emerging, note immaculately wellpreserved ornaments along septal surfaces.
preserve fossils from this locality, but the primary mode of preservation is silicification. Gravina Island General History, Age, and Lithology. Like the Upper Triassic units from the Keku Strait area, Gravina Island was chosen for study because of its known abundance of silicified, shallowwater marine fossils originally described by Smith (1927) and later mapped by Berg (1973). Stanley (1979) contributed to the tectonic implications of the then “suspect” Alexander terrane by looking at the reef-building potential and geographic distribution of carbonate buildups from two localities along the west coast of Gravina Island. Montanaro Gallitelli et al. (1979) conducted the last major paleontological study of Late Triassic corals from Gravina Island by updating the original taxonomy of Smith (1927). The resulting taxonomic list was subsequently used by Yarnell (2000) for paleobiogeographic analysis. The Upper Triassic stratigraphy from Gravina Island includes only one stratigraphic unit, the Nehenta Formation of Berg (1973). Past work has indicated a Late Triassic age based on the presence of the flat clam genus Halobia (not formally speciated by Smith, 1915; Chapin, 1918; Martin, 1926; and Berg, 1973); however, an early Norian age has now been confirmed based on conodonts identified by Erik Katvala (2005, personal commun.) (Fig. 9).
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Figure 6. Large scale maps of the Cornwallis Peninsula (A) and Eastern Kuiu Island (B) areas. Dots represent Triassic localities visited by Erik Katvala (in 2002 and 2003), and locality numbers indicate Montana Invertebrate (MI) locality numbers used in this contribution. Trk—Keku Volcanics, Trc—Cornwallis Limestone, and Trv—Hound Island Volcanics (modified from Katvala and Stanley, this volume).
The Nehenta Formation has been subdivided into three principal members: (1) a calcareous member that encompasses carbonaceous limestone and siltstone in its lower part and calcareous conglomerate, grit, and sandstone in its upper part; (2) a coarse conglomerate that intertongues with the lower part of the calcareous member; and (3) a basaltic volcanic member occurring in the lower part of the formation (Berg, 1973). Although these three members have been identified, structural complexities have made stratigraphic order among the assigned members difficult to determine. Facies variation occurring along the western coast of Gravina Island makes it difficult to distinguish between members one and two. Following the generalized geological map published by Berg (1973), the Nehenta Formation is observed cropping out along the intertidal zone on both sides of Gravina Island (Fig. 10). A section
of the Nehenta Formation was measured on the east side of the island at Bostwick Inlet (A–A′ on Fig. 10) where beds strike almost perpendicular to the beach. The section along this beach (Fig. 11) was the only place on the island where the Nehenta Formation could be measured. Biofacies. Several biofacies were noted within the measured section at Bostwick Inlet as well as in exposures along the western side of Gravina Island. Biofacies include (1) coquina biofacies containing densely packed, shallow-water bivalves and a few gastropods within limestone interbeds cropping out at 361–364 m in the measured section at Bostwick Inlet (Figs. 11 and 12); (2) Halobia-bearing black calcareous shale with interbedded siltstone biofacies within all three members of the Nehenta Formation; (3) Heterastridium conglomerate biofacies rich with the spherical hydrozoan Heterastridium
Late Triassic silicified shallow-water corals and other marine fossils
Figure 7. Partially etched limestone block from the Gil Harbor mud flat (MI 0087) showing (A) rich diversity of silicified bivalves (pectinacian located top of image) somewhat fragmented by original processes, gastropods (new genus from subfamily Astraeinae above), and annelid worm tubes with random orientation (scale bar, 1 cm); (B) silicified and fragmented bivalves and spherical hydrozoan Heterastridium (scale bar, 1 cm).
Figure 8. Silicified coral colony of the genus Crassistella Roniewicz preserved encrusting a bivalve (scale bar is 1 cm).
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Figure 9. Late Triassic biochronology presenting relevant conodont, ammonoid and bivalve biozones. Genus abbreviations: M.—Metapolygnathus, E.—Epigondolella, Mi.—Misikella, H.—Halobia. Ammonoid zones after Tozer (1967, 1984, 1994) and Orchard and Tozer (1997). Approximate bivalve ranges after McRoberts (1993, 1997), Silberling et al. (1997), and Chris McRoberts (2003, personal commun.). Conodont zones after Orchard (1991) and Orchard and Tozer (1997). (Figure adapted from Katvala and Stanley, this volume.)
cropping out at 250–265 m within the measured section at Bostwick Inlet (Fig. 11A) and north of Nelson Cove on the west coast, which may indicate a stratigraphic continuity between both sides of the island; and (4) coral- and spongiomorph-rich biofacies abundant along the west coast of the island. This biofacies contains several tectonically convoluted depositional environments described in the following sections. Recovered Fossils. Early Norian macrofossils were recovered from both sides of Gravina Island, southeast Alaska. Along the east coast at Bostwick Inlet, recovered fossils include bivalves, gastropods, Heterastridium, algal balls, and scattered, often splintered, bone fragments. Recovered microfossils include conodonts, a few vertebrate bone fragments, and fish scales.
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Figure 10. General geological map of Gravina Island adapted from Berg (1973) showing the Nehenta Formation (outlined in white). Montana Invertebrate (MI) locality numbers and field localities are given, as well as the section measured along the exposure at Bostwick Inlet (A–A′).
The majority of macrofossils were not silicified and instead were replaced by secondary calcite; however, some specimens were preserved with low-level silicification and were recovered by acid dissolution. The west side of Gravina Island had many genera and species of coral, sponges, spongiomorphs, Heterastridium, brachiopods, bivalves, gastropods, echinoderms, crinoid ossicles, a large nautiloid, and aulacocerid cephalopods, as well as a few bone fragments and conodonts. Preservation can be described as low-grade, poor, or under-silicified for many of the fossils. Within the outcrop fossils seem well silicified; however, after acid dissolution, the minute details are often obliterated, even with the use of a hardener such as Alvar. Localities. Variation in facies was recorded along the west side of the island where limestone beds strike in a NE–SW direction, having a general trend approximately parallel to the shoreline. Structural deformation is more intense to the south, along the shores of Nehenta and Phocena Bays (Fig. 13A) and less intense along the coastline north of Nelson Cove (Fig. 13B). Fossiliferous localities are concentrated in three large areas along the intertidal zone of the western coast: (1) to the south, Nehenta Bay contains two described localities; (2) one locality is located farther north at Phocena Bay; and (3) the area north of Nelson Cove provides three described fossiliferous localities. The rocky tidal flats of Nehenta Bay (MI 9930–9932) contain large olistolith limestone reef blocks (up to 7.5 m long), resting in a black, folded, and highly fractured carbonaceous
and sometimes calcareous shale (Fig. 14A). Olistolith boulders, weather resistant compared to the enclosing sedimentary layers, yield abundant silicified fauna including corals (Distichophyllia Cuif, Crassistella Roniewicz, Pamiroseris Melnikova, Cuifia Melnikova, Gablonzeria Cuif, and Paracuifia Melnikova), sponges, spongiomorphs, gastropods, and brachiopods. Corals and sponges are dominant within these boulders, often preserved in situ consistent with the reef-like structure inside the blocks (Fig. 14B) prior to sliding into deeper water. This facies in interpreted as a forereef zone within the Alexander terrane. Adjacent beds contain large colonies of encrusting colonial corals (Crassistella juvavica Roniewicz and Astraeomorpha crassisepta Reuss) as well as the calcified sponge Stromatomorpha californica Smith. Corals can also be observed within imbricated limestone clasts in a bedded fossiliferous limestone unit within the volcaniclastic, argillite of locality MI 9930 (Fig. 15). This bedded fossiliferous facies interfingers with the Halobia-bearing black calcareous shale with interbedded siltstone biofacies (described above). Farther north at Phocena Bay, the Nehenta Formation crops out along the intertidal zone of the northern, northeastern, and western shorelines. Only one recrystallized, poorly preserved, fossiliferous limestone bed was recognized along the northern shore. Abundant halobiid bivalves were collected from the black calcareous shale within interbedded siltstone (biofacies 2) cropping out along the tide line of the bay. Along the western shoreline, locality MI 9933 contained abundant, large, complete, but
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Figure 11. Measured section of the Nehenta Formation at Bostwick Inlet, Gravina Island (A–A′ on Fig. 10). (A) Base of section up to 340 m, showing limestone interbeds with Heterastridium-rich conglomerate interval at 255 m and locations of conodont samples taken. (B) Continuation of section up to 680 m showing bivalve coquinas, with scattered marine reptile bones at ~364 m as well as intrusive sills with accompanying metamorphosed shale within the upper part of the section.
poorly preserved, coral colonies from a sandy, medium-grained limestone unit. All coral colonies seemed to belong to Meandrostylis (Frech); however, only the corallite shape is preserved and no septa can be recognized from any of the colonies which exhibit large plate-like growth without encrustation. The Nelson Cove area contains an unfossiliferous, massively bedded, sandy dolomite unit on the southwest facing arm, stratigraphically overlying black calcareous shale. North of Nelson
Cove, the limestone beds thicken and include at least four different facies exhibiting small-scale structural deformity. The first facies is a fossiliferous limestone with poorly preserved corals, sponges, and spongiomorph bioclast fragments (locality MI 9936). Coral and spongiomorph colonies are the most abundant organisms at this locality and exhibit either encrusting or plate-like growth. Colonies are not typically in life position but are visibly broken and often upside-down, growing on top of each other. This particular outcrop
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Figure 12. Photograph of non-silicified bivalves (trigoniid left and pteriid right) (Chris McRoberts, 2005, personal commun.) from coquina beds at 364 m in measured section at Bostwick Inlet, Gravina Island (Fig. 11); pteriid bivalve is ~2 cm wide.
Figure 14. Nehenta Bay (MI 9931) showing (A) olistolith boulder measuring 7.5 m long and (B) multiserial branching columnar coral colony (arrows) preserved in life position within olistolith boulder (scale in centimeters).
Figure 13. Field pictures showing (A) intense structural folding of the Nehenta Formation along the tide flat of Nehenta Bay; dotted line traces fold, rock hammer in foreground for scale; (B) smaller-scale folding of Nehenta Formation north of Nelson Cove, Gravina Island.
is interpreted as having a broken-up reef-type ecology or floatstone appearance, possibly in a backreef setting. Facies number two is a calcareous tuff and tuffaceous limestone interbedded within a rounded basalt pebble, cobble, and boulder limestone unit (Fig. 16). This dominantly volcanic conglomerate facies flanks both sides of the volcanic member established by Berg (1973) and interfingers with the fossiliferous facies described above. Facies number two compares well with the description given by Berg (1973), especially with respect to the basalt pillow flows, pillow breccia, and subordinate calcareous tuff cropping out of the intertidal zone ~0.25 km north of Nelson Cove. This facies is interpreted as a volcanic episode entering a shallow backreef setting. One mile north of Nelson Cove, the Nehenta Formation consists of fine- to medium-grained limestone that is massively bedded (4–5 m thick) and locally may be thrust on top of black calcareous shales, constituting facies number three. The massive limestone beds contain sporadically concentrated areas of poorly silicified and fragmented colonial corals mostly of the genus Crassistella
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Figure 16. (A) Volcanic conglomerate facies of the Nehenta Formation with abundant basalt pebbles and cobbles exposed along the surface. (B) Surface view of this facies with rounded vesicular basalt cobbles trapped in the matrix (hammer in center for scale). Figure 15. Limestone rip-up clasts surrounded by argillite and shale in Nehenta Bay, Gravina Island showing (A) slight imbrication of limestone clasts and (B) a silicified coral colony of the genus Crassistella Roniewicz (scale bar in centimeters).
Roniewicz, as well as few calcareous sponges. Samples were not collected because of the poor preservation and massive and inaccessible nature of the outcrop. The massively bedded, fine-grained, unfossiliferous limestone of facies number three grades conformably into facies number four, a structurally intact fossiliferous limestone reef or biostrome (locality MI 9935) averaging 8–10 m thick and 50–58 m long (Fig. 17). Corals are the dominant taxa. They form sinuous, encrusting colonies several meters long with large plates, up to 0.5 m long (Fig. 17B). Thick multiserial columns are interspersed within coralline plates and even mound-shaped phaceloid- and dendroid-growing colonies. Corals are frequently intergrown with several calcified sponges and spongiomorphs, most notably Stromatomorpha californica Smith and Spongiomorpha ramosa Frech. These framework organisms typically grow on top of one another, forming reef framework within this deposit, exhibiting a reef crest facies. The
coral fauna at this locality is the most diverse fauna from Gravina Island. Identified coral genera include Kuhnastraea Cuif, Chondrocoenia Roniewicz, Astraeomorpha Reuss, Crassistella Roniewicz, and Retiophyllia Cuif. The top of the deposit is characterized by a small transgressive sequence of unfossiliferous limestone and shale interbeds grading upward into black calcareous shale. This unit of the Nehenta Formation cannot be mistaken for an olistolith boulder, primarily because of its bedded nature and conformable contacts stratigraphically above and below the interval. The presence of a well-developed framework reef, which encompasses forereef, backreef, and reef crest biofacies, within the Nehenta Limestone suggests two conditions that enabled smallscale reef growth: (1) low rates of sedimentation and a decreased intensity of volcanism within this part of the Alexander terrane; and (2) shallow, warm water surrounding the terrane with low nutrient levels and sufficient sunlight for possible zooxanthellate coral growth (Hallock, 2001). North of locality MI 9935 and 1.7 miles north of Nelson Cove, massive fine-grained limestone beds become thinner and less pronounced. Fossiliferous, medium- to coarse-grained calcareous
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Figure 17. Bedded biostrome reef in Nehenta Formation north of Nelson Cove, Gravina Island showing (A) conformable contact with massive limestone below with in situ structured growth of laminar corals and calcified sponges (fallen tree for scale); (B) surface view of a silicified coral colony with laminar, plate-like growth; colony in center is ~20 cm long.
sandstone and limestone interbeds crop out within black calcareous shale (MI 9938). Interbeds range in thickness from 5 to 10 cm closer to tide line to 1–2 m closer to high-tide mark, eventually grading into black calcareous shale. No cross-beds, ripple-marks, or transport indicators were seen in the sandstones. Silicified fossils include bivalves, brachiopods, echinoid spines, gastropods, corals (from genus Paracuifia Melnikova), and sponges (locality MI 9938). Wrangellia Localities from this terrane are separated into two regions of Canada and Alaska. These are discussed in the following sections.
Northern Wrangellia, South-Central Alaska Another paleontologically important Upper Triassic site is located in the Wrangell Mountains, south-central Alaska, in what is known today as northern Wrangellia. Originally, MacKevett (1965, 1970, 1974, and 1976), Armstrong et al. (1969), and Armstrong and MacKevett (1982) extensively mapped and made lithologic, stratigraphic, and structural comparisons throughout the Wrangell Mountains. They indicated the presence of a silicified fossiliferous packstone bed at Green Butte, which they designated USGS Mesozoic locality M1708. Jones et al. (1977) subsequently provided a stratigraphic correlation linking the Wrangell Mountains, Alaska (type section), with Vancouver Island, British Columbia; Queen Charlotte Islands, British Columbia; Chichagof Island, Alaska; and the Wallowa Mountains, Oregon, suggesting these five landmasses, separated by a distance of up to 200 km, were part of a much larger terrane named Wrangellia (Fig. 2). Newton (1983) revisited M1708 and initiated a systematic study of the bivalves, applying her results for paleobiogeographic analysis. She further interpreted paleoenvironmental conditions as well as paleoecologic interactions between bivalves and gastropod faunas. Montanaro Gallitelli et al. (1979) did not visit this site but studied the USGS material collected by N.J. Silberling; thus providing initial systematic description of the corals. However, the resulting study was based on a small sample and subsequent work was needed. General Stratigraphy and Age. The Upper Triassic carbonate succession in northern Wrangellia at Green Butte is represented by the Chitistone and Nizina Limestones, which overlie a thick upper Triassic sequence of volcanic rocks assigned to the Nikolai Greenstone (Fig. 2). The carbonate succession at Green Butte was chosen for study because of the well-preserved, shallow-water silicified fossils concentrated in shell beds within the upper part of the Chitistone Limestone. USGS Mesozoic locality M1708 was first made known by Silberling in Armstrong et al. (1969), where he identified the ammonoids Tropites cf. T. welleri Smith and Arcestes, and the bivalve Halobia cf. H. superba Mojsisovics, which occurred in a unit 152 m above the base of the section. These fossils indicated a late Carnian to early Norian age for the Chitistone Limestone. Lithofacies at Green Butte. The Triassic limestone at Green Butte is 1067 m thick (701 m of Chitistone Limestone and 366 m of Nizina Limestone) as measured above the smooth unaltered surface of the underlying Nikolai Greenstone (Armstrong et al., 1969). The limestone is consistently bedded and shows little evidence of any reef development (Fig. 18A). The basal 2 m of the Chitistone Limestone were described by Armstrong and MacKevett (1982) as interbedded, black calcareous shale and paleyellow, weathered, thinly bedded, argillaceous lime mudstone. A pale-orange to pale-gray medium-bedded dolomite makes up the bulk of the unit and is exposed up to 87 m above the underlying Nikolai Greenstone. Within the Green Butte section, two concentrated fossiliferous deposits were mentioned by Armstrong and MacKevett (1982), but were not described in enough detail to assign a specific bed or location within the section. A
Late Triassic silicified shallow-water corals and other marine fossils
Figure 18. (A) Wrangell Mountain locality MI 9934 (USGS Mesozoic locality M1708) along the east face of Green Butte at conformable contact between Chitistone and Nizina Formations. Arrows indicate camp and localities where sparse and scattered aulacocerid cephalopods and Halobia were found. (B) Locality MI 9934 situated along a steeply inclined talus slope.
peloid-algae-molluscan packstone is present between 87 and 90 m above the Nikolai Greenstone as well as a massive, gray, algaemolluscan-echinoderm wackestone and packstone at 103.5 m. The latter horizon is the well-known USGS Mesozoic locality M1708 designated herein as University of Montana locality MI 9934 from the newly collected material. Silicified Bed at the Green Butte Locality. This locality is observable cropping out along the eastern edge of a northeast facing talus slope within the large, northeast drainage of Green Butte striking NW–SE at 307° and dipping 32° to the NE, in the McCarthy B-5 Quadrangle (Fig. 18). This deposit is extremely important for its highly diversified shallow-water marine fauna.
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An early Norian age is indicated by conodonts recovered by Erik Katvala (2005, personal commun.). The fossils are of exceptional quality and preserved by silicification, as noted by Newton (1983) and Montanaro Gallitelli et al. (1979). Locality MI 9934 lies close to, if not directly on, the conformable contact between the thick- to medium-bedded Chitistone Limestone and the thinly bedded Nizina Limestone. However, recent conodont work by Erik Katvala (2005, personal commun.) reveals a small disconformity within the Chitistone Limestone, separating the underlying Late Carnian, thick to medium-bedded limestone and the Early Norian bioclastic, grain-supported packstone of MI 9934. Lithologically, the outcrop consists of a very well silicified, bioclastic, tabular, grain-supported coquina or packstone (Fig. 19A) within a bedded, fine- to medium-grained limestone matrix, 2.12 m thick and 6.6 m in outcrop length. This grain-supported coquina contains cobble-sized limestone clasts (2–15 cm in diameter) that are scattered throughout the deposit. No apparent clast imbrication, transport indicators, or preferential orientation of bioclasts are present in the outcrop. However, carbonate clasts were soft and unlithified when they were redeposited, as evidenced by intrusions of fossil lenses into limestone clasts (Figs. 19B and 19C). The limited lateral extent of the deposit as well as the wavy or irregular basal “scoured” contact (Fig. 19C) argues for a slopechannel deposit. Furthermore, the grain-supported nature, random and jumbled consistency of bioclasts (lack of preferential orientation), and large trapped clasts suggest deposition on the proximal edge of a debris flow (i.e., submarine fan) or the “A” horizon of a turbidite sequence (Howell and Normark, 1998). This slurry-like mass became activated and moved downhill, possibly from local tectonics or gravity on a slope, causing a submarine slide prior to lithification. A species of Halobia was collected and identified as Halobia cf. H. austriaca (Mojsisovics) by Chris McRoberts (2005, personal commun.) (Fig. 20) from fine-grained limestone directly below the disconformity. This taxon indicates an earliest Norian (Kerri ammonoid biozone) age for the sediments directly below MI 9934 (Fig. 9). Recovered Fossils. Well-silicified shallow-water marine organisms, including bivalves, gastropods, ammonoids, echinoderms, calcified sponges, spongiomorphs, algae, corals, and lots of bone fragments (mostly in conodont residue), were recovered from etched limestone blocks retrieved from MI 9934. Gastropods identified by Robert B. Blodgett (2005, personal commun.) included Amphiscapha sp., Wortheniella spp., Pleurotomaria subcancellata d’Orbigny, Zygites sp., Temnotropis magnus n. sp., three new species of a new discohelicid genus, Naticopsis sp., a new genus (and new species) aff. Hyperacanthus, Trypanocochlea n. sp., an undetermined zygopleurid, a loxonematid with numerous spiral cords, Angularia n. sp. aff. A. subpleurotomaria, Spinidelphinulopsis whaleni Blodgett, Frýda, and Stanley, Paradelphinulopsis valleuri Blodgett, Frýda, and Stanley, Neritopsis (Wallowiella) vallieri Frýda, Blodgett and Stanley, Nuetzelopsis tozeri Frýda, Blodgett and Stanley, Eucycloscala sp., Chartroniella n. sp. cf. pacifica (Jaworski), Ompahloptycha n. sp.,
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Figure 20. Halobiid bivalve Halobia cf. H. austriaca (Mojsisovics) found in butterfly position, below scour of MI 9934 (Fig. 19), indicating earliest Norian (Kerri ammonite biozone) age (scale in centimeters).
Figure 19. Locality MI 9934 depicting (A) slightly tabular, bioclastic, grain-supported nature; (B) intrusion of a silicified high-spired gastropod into limestone clast (scale bar in centimeters); and (C) evident scour marking base of the outcrop (dashed black line) with intrusions of fossils into a carbonate clast (arrow) (scale in centimeters).
and Protorcula sp. Coral taxa from 12 genera were identified: Ampakabastraea Alloiteau, Astraeomorpha Reuss, Chondrocoenia Roniewicz, Crassistella Roniewicz, Distichomeandra Cuif, Distichophyllia Cuif, Gablonzeria Cuif, Kuhnastraea Cuif, Margarosmilia Volz, Pamiroseris Melnikova, Paracuifia Melnikova, and Retiophyllia Cuif. Within these 12 genera, ten species had previously been described by Montanaro Gallitelli et al. (1979), and thirteen are new to the locality (Appendix). Montanaro Gallitelli et al. (1979) also identified species from the genera Thamnasteriomorpha Melnikova and Guembelastraea Cuif, which were not identified in this study. Astraeomorpha Reuss, Crassistella Roniewicz, Kuhnastraea Cuif, Gablonzeria Cuif, and Distichophyllia Cuif are the most abundant corals within this locality. Colonies commonly exhibit mound-shaped growth and are sporadically intergrown with other corals, spongiomorphs, and algae. Such relationships indicate a low level of ecological interaction or competition for growth space. Although the genus Retiophyllia Cuif occurs frequently, specimens are highly fragmented and likely came from larger phaceloid-dendroid colonies. By comparison with taxa from the Alps, corals from this deposit are largely recognized as being Norian; however, the genera Distichomeandra Cuif and Retiophyllia Cuif contain species that are recognized in Carnian deposits in the Tethys (Caruthers and Stanley, 2008; Table 2). The fauna at Green Butte presents the best silicification and the greatest diversity of any Late Triassic locality from the North American Cordillera. Upon etching, fossils are often welded together from silicification and need to be gently separated with
Late Triassic silicified shallow-water corals and other marine fossils TABLE 2 (A) RCSI ANALYSIS MATRIX SHOWING SPECIES LEVEL FAUNAL SIMILARITY IN CORALS FROM THE ALEXANDER TERRANE, WRANGELLIA, WALLOWA TERRANE, AND LOCALITIES FROM THE PUCARÁ GROUP, PERU AT NWR SWR WA PU AT
1
NWR
0.195
0.11725
0.0825
0.2625
1
0.195
0.575
0.325
SWR
1
WA
0.2975
0.9675
1
0.1275
PU
1
(B) SIMPLIFIED RCSI MATRIX FROM DATA COMPILED ABOVE, WITH MOST SIGNIFICANT VALUES INDICATED BY SYMBOLS AT NWR SWR WA PU AT NWR SWR WA PU
1
§
†
1
§
§
1
* 1
† 1
Note: AT—Alexander terrane; NWR—Northern Wrangellia; SWR—Southern Wrangellia; WA—Wallowa terrane; PU—Peru; n = 438. Scale adapted from Yarnell (2000). *Similarity 96%–100% (greatly significant). † Dissimilarity 0–13% (greatly significant). § Dissimilarity 14%–20% (significant).
a dental pick or other tool. Preservation of fine details such as delicate ornamentation or spines along the lateral surfaces of septa are often observed within some corals, as well as in the shell ornamentation and growth lines of certain gastropods and bivalves. Even echinoid plates and crinoid stems retain their minutely detailed structure. Southern Wrangellia, Vancouver Island General Stratigraphy and Age. Jones et al. (1977) documented the strata on Vancouver Island as a southern fragment of Wrangellia and correlative to Triassic rocks from the Wrangell Mountains (Fig. 2). Triassic rocks of Vancouver Island are composed of three formations that have been studied by many geologists (e.g., Fyles, 1955; Muller et al., 1974; Massey and Friday, 1987; Yorath et al., 1999). Stratigraphically, from oldest to youngest, these are the Karmutsen, Quatsino, and Parson Bay Formations, which together make up the larger Vancouver Group. The underlying volcanic Karmutsen Formation is loosely correlated with the Nikolai Greenstone of the Wrangell Mountains in south-central Alaska, differing with respect to percentage of pillow basalts and aquagene tuffs (Jones et al., 1977). As in the Wrangell Mountains, the overlying Triassic units from Vancouver Island are composed of carbonate rocks. The Quatsino Formation is a light-gray, massive to thickly bedded or blocky to flaggy dark-gray limestone, interbedded with shale and recognized as a shallow-water platform carbonate (Jeletzky, 1970; Muller et al., 1974; Carlisle and Susuki, 1974; Yorath et al., 1999). This formation correlates stratigraphically with the upper Carnian/lower Norian Chitistone Limestone in south-central
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Alaska (Jones et al., 1977; Yorath et al., 1999). From exposures occurring at many sites on Vancouver Island, Carlisle and Susuki (1974, p. 258–263) recognized an upper Carnian conformable contact between the Quatsino Formation and the overlying Parson Bay Formation. The entire extent of the Parson Bay Formation and its associated stratigraphic relationship to the “Sutton Limestone” has been subject to different interpretations. Muller (1977), Jones et al. (1977), and Massey and Friday (1987) placed this unit as a member within the Parson Bay Formation; however, Clapp (1912) and Yorath et al. (1999) recognized it as an individual formation overlying the Parson Bay Formation, largely by its distinctive character and widespread distribution (Yorath et al., 1999). This paper follows the stratigraphy set forth by Muller (1977), Jones et al. (1977), and Massey and Friday (1987) designating the Sutton Limestone as a member within the Parson Bay Formation. The Parson Bay Formation. Bancroft (1913) established the Parson Bay Group for exposed outcrops of shale, limestone, and minor volcanic rocks located along Parson Bay on Harbledown Island, north of Vancouver Island. However, upon recognition of both Jurassic and Triassic fossils within the group, Crickmay (1928) subdivided the Parson Bay Group into two formations: (1) the Parson Bay Formation including Triassic rocks and (2) the Harbledown Formation including Jurassic rocks. Yorath et al. (1999) stratigraphically correlated the Parson Bay Formation of Vancouver Island with the upper Norian Nizina Formation of south-central Alaska. General Lithology and Locality. The Parson Bay Formation has a highly variable lithology, containing thinly bedded shale, limestone, argillite, sandstone, and minor volcanic rocks, as well as many benthic and pelagic shallow-water marine fossils (largely within the Sutton member). Fossils include corals, bivalves, gastropods, and ammonoids (Clapp and Shimer, 1911; Fyles, 1955; Jones et al., 1977; Massey and Friday, 1987; Stanley, 1979, 1989). These fossils have been interpreted by Carlisle and Susuki (1974) as late Norian on the basis of the presence of the flat clam Monotis subcircularis. The Parson Bay Formation has been observed cropping out at various localities throughout Vancouver Island. For this work, three localities were selected for study: (1) the southern shore of Lake Cowichan, (2) Tahsis Inlet on the west coast, and (3) Quatsino Sound to the north (Fig. 21). At Tahsis Inlet, fossiliferous localities were identified and sampled (unpublished data) by M. Orchard and D. Erwin as well as G.D. Stanley and K. Paisley. Field notes from Stanley and Paisley indicate that samples were collected from four sites. Site 1 is a steeply dipping bedded limestone with interbedded volcanics, calcareous mudstone, and argillite, cropping out along a reclaimed logging road near Lutes Creek in an unnamed valley located west of Tahsis Inlet near the town of Tahsis (Fig. 22A). Here limestone is leached with molds of solitary and colonial corals, trigoniid bivalves, gastropods, brachiopods, echinoids, and crinoid ossicles. Site 2, ~1 km from site 1, is located in a small ditch within the valley. Abundant fossil molds of stromatoporoid-like organisms as well as trace fossils and possible molds of crab claws were collected. Site 3
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Figure 21. Satellite image of Vancouver Island showing the location of the three study areas used in this paper (modified from www.googlemaps.com).
Figure 22. (A) Coral horizon of Parsons Bay Formation at Tahsis Inlet, Vancouver Island (MI 9913). (B) Fossil mold of the meandroid coral Campesteria proloxia Caruthers and Stanley from MI 9913 (rock hammer for scale).
(MI 9913) is stratigraphically below site 2 and contains molds of phaceloid and meandroid corals (Fig. 22B) as well as thick trigoniid bivalves within bedded limestone. Site 4 is located in the Quatsino Formation along the eastern shoreline of Tahsis Inlet. It contains molds of corals from the genera Anthostylis (Frech), Campesteria Caruthers and Stanley, Distichophyllia Cuif, and Distichomeandra Cuif, as well as sponges, bivalves, brachiopods, and the calcified sponge Stromatomorpha californica Smith. Sutton Member. Clapp first recognized the Sutton Member as a formation that included all of the intercalated limestones within the Vancouver Group of southern Vancouver Island (Clapp and Shimer, 1911). Silicified shallow-water marine fossils collected from the type section at Lake Cowichan were mistakenly identified as Early Jurassic by Clapp and Shimer (1911); however, subsequent work by Fyles (1955) and Shimer (1926) showed them to be Late Triassic in age. Likewise, detailed mapping and stratigraphic revision have placed this limestone as a member within the Parson Bay Formation (Muller, 1977; Massey and Friday, 1987). The bulk of the Sutton Member occurs along the southwest flank of the Cowichan Anticlinorium, cropping out at three main localities: (1) the type section, identified by Clapp and Shimer (1911), along the south shore of Lake Cowichan (MI 8302), Vancouver Island, ~4.8 km west of Sutton Creek (Fig. 23); (2) a section 60 m thick at Redbed Creek, identified by Yorath et al. (1999) as well as the adjacent creek to the north; and (3) a locality near Sproat Lake also identified by Yorath et al. (1999). The Sutton Member lies unconformably on top of the Karmutsen Formation along the Cowichan Anticlinorium and contains abundant coarse-grained cobbles and breccia of Karmutsen Formation origin that are trapped within its limestone matrix (Yorath et al. 1999). Fyles (1955) noted that the Sutton Member at Lake Cowichan grades laterally into basalt flows and volcaniclastic
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Figure 23. Satellite image of Lake Cowichan, southern Vancouver Island showing location of the Sutton Member outlined in white. (Image from www. googlemaps.com; geology adapted from Massey and Friday, 1987.)
Figure 24. Medium bedded, fine-grained limestone beds of MI 8302 (dipping 45°, striking NE–SW) along the southern shores of Lake Cowichan, Vancouver Island (sledge hammer, 78 cm, centerleft for scale).
rock, which Stanley (1989) asserted had been correlated mistakenly with the Karmutsen Formation along subtle fault contacts. Biofacies. Clapp and Shimer (1911) as well as Stanley (1979, 1989) noted the existence of at least three distinct zones or biofacies within the thin- to medium-bedded, fine-grained gray fossiliferous and flaggy limestone striking approximately parallel to the beach of Lake Cowichan (MI 8302; Fig. 24). These
include (1) a coarse, grain-supported, diverse bivalve packstone in a fine-grained limestone matrix that includes bivalves, gastropods, echinoderms, cephalopods, and worm tubes (Fig. 25A); (2) an in situ branching coral zone of moderate diversity represented by the genus Retiophyllia Cuif (Fig. 25B); and (3) an encrusting flat coral zone dominated by the genera Anthostylis Roniewicz, Campesteria Caruthers and Stanley, Astraeomorpha Reuss,
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Figure 25. Biofacies of MI 8302 showing (A) packstone biofacies with silicified bivalves, gastropods, echinoids, and corals; (B) in situ branching coral biofacies evident by the phaceloid coral Retiophyllia clathrata Emmrich (sledge hammer head bottom-right for scale); (C) encrusting coral biofacies showing a fragmented colony genus Meandrostylis (Frech) (scale bar in centimeters).
Kuhnastraea Cuif, Crassistella Roniewicz, and Gablonzeria Cuif (Fig. 25C). Stanley (1979, 1989) reported on the gross paleoecology and biofacies at Lake Cowichan and compiled a species and faunal list for recovered fossils in a measured section 71.7 m thick. Collected limestone blocks from this site were used for acid etching, and well-preserved material was retrieved. Fine silicification has preserved significant amounts of detail and diversity within shallow-water marine fauna at Lake Cowichan. In silicification and texture of biofacies the Lake Cowichan locality resembles that of Green Butte (MI 9934), where minute details are preserved; cementation of fossils within the bivalve-packstone biofacies (1) as described above is also similar. Some earlier collected material was etched in HCl, but subsequent collected material was processed with acetic acid, which yielded better results (Fig. 26).
Unpublished data from Stanley and Fois-Erickson reveal another locality within the Sutton Member at Quatsino Sound in northern Vancouver Island (MI 9947). Lithologically, locality MI 9947 is a fossiliferous dark-gray, medium- to fine-grained relatively pure limestone with chert stringers. The locality crops out near a rockslide at Pender Point in Quatsino Sound, northern Vancouver Island. Retiophyllia Cuif was identified in the field, and the genera Astraeomorpha Reuss and Gablonzeria Cuif were identified from etched samples. Silicification of corals from Pender Point is not as complete (under-silicified) as in samples from Lake Cowichan (MI 8302). Within this coral population, most of the external features including epitheca and colony shape are preserved by fine-grained silicification; however, silicification does not penetrate the entire colony. This is especially evident in samples of the genus Retiophyllia Cuif, making specimens extremely fragile and easily fragmented after etching.
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Figure 26. Chondrocoenia schafhaeutli (Winkler) from Lake Cowichan, MI 8302 showing (A) specimen processed in HCl with corallites possessing little or no structural detail, and (B) specimen etched in acetic acid, retaining much corallite detail throughout the colony (scale bars = 1 cm).
FINDINGS AND CONCLUSIONS Biofacies and Lithofacies Comparisons Several biofacies and lithofacies are recorded for the first time from the Upper Triassic stratigraphy in the Keku Strait area and on Gravina Island, southeast Alaska (Alexander terrane). Biostrome reef deposits are recognized in situ north of Nelson Cove and as olistolith boulders resting in deeper water calcareous shales in Nehenta Bay (Gravina Island), as well as Big Spruce Island and Cornwallis Peninsula East (Keku Strait). Clear nutrient-poor water is needed to support a coral-rich biostrome reef development of this magnitude. Furthermore, large colony size, intergrown nature of corals with other organisms, absence of significant filter- and detritusfeeding organisms, and structured growth of coral colonies argue for zooxanthellate symbiosis. During early Norian time, therefore, these localities within the Alexander terrane were areas of deposition along the fore-reef zone of small fringing reefs where warm, clear-water currents provided a relatively nutrient-free environment with low volcanism and sedimentation rates enabling unaltered structured reef growth. In contrast, several other localities within the Alexander terrane (Keku Strait) and northern Wrangellia (Wrangell Mountains) contain coral colony growth that do not show structure or frame-
work, are smaller in size, and grow separately rather than being intergrown with various sponges and spongiomorphs. Fossiliferous deposits from this area also contain many types of filter- and detritus-feeding organisms such as bivalves, gastropods, annelid worm tubes, algae encrusted corals, oncolites, crinoids, and echinoderms, which indicate nutrient-rich, cloudy water typical of unrestricted slope/ramp environments. Southern Wrangellia (Lake Cowichan) contains three distinct biofacies (Stanley, 1989) indicating variable environments of deposition during Late Triassic time. Coral-rich biofacies including in situ branching corals and flat encrusting corals indicate patch-reef buildups occurring below the intertidal zone, whereas the highly fragmented bivalves of biofacies three represent deposition along the shallow-water beach zone. Biofacies did not contain many detritus and filter feeding organisms such as echinoderms, gastropods, crinoids, algal mats, or oncolites, which would indicate cloudy nutrient-rich water. However, the presence of wood within the in situ branching coral biofacies signifies a close proximity to land, and the fine-grained limestone matrix indicates relatively quiet water. Reefs in North American Terranes Triassic coral reefs in North America have been discussed by Smith (1915, 1927). He recognized a strong equatorial
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affinity of taxa, presenting the possibility that an extensive Carnian–Norian reef system developed along the western coast of North America during Late Triassic time. In these publications, Smith attempted to make sense out of the broad latitudinal distribution of coral assemblages that he envisioned as reefs fringing the western coast of North America. Smith, however, lacked a clear understanding of plate tectonics, and so his interpretations, invoking only broader climatic extensions of the tropical belt to account for this distribution of tropical shallow-water marine fauna, were stretched beyond credibility. Stanley (1979, 1996) consolidated these early ideas of Triassic reefs into a framework of plate tectonics and the concept of terranes. As discussed earlier in this contribution, application of the terms “reefs” in general and “coral reefs” in particular has inherent problems (Stanley, 2001). Most of Smith’s (1915, 1927) early examples of “coral reefs” are simply assemblages of corals and associated shallow-water marine organisms that lived in former island-arc settings. Some notable examples, such as those in Quesnellia, the Wallowa terrane, and Stikinia, contain massive, coral-dominated carbonate deposits that are similar to those occurring in the former Tethys Region. The latitudinal distribution of these occurrences confirms tropical and subtropical climate as well as postdepositional tectonic movement of terranes into higher latitudes. Such reef deposits have not been found in either Wrangellia or the Alexander terrane (except for small-scale biostromes as indicated in this contribution). Flügel (2002) presented world paleogeographic maps incorporating terranes of the eastern Panthalassan Ocean. These maps illustrated the importance of terrane tectonics in the interpretation of paleogeography. A rise in sea level in the Norian made conditions optimal for the development of tropical reefs, thus precipitating their expansion worldwide (Flügel, 2002). Coral Discussion and Paleobiogeographic Analysis A total of 458 silicified corals were collected and described from 12 localities within northern and southern Wrangellia and the Alexander terrane (Table 1). Twenty-five species were identified from eight localities (five localities from the Keku Strait area and three from Gravina Island) within the Alexander terrane. Of these, 19 species had not previously been recognized from the Alexander terrane (Appendix). Thirty-four species of coral were identified from four principal localities within Wrangellia (one locality from northern Wrangellia in the Wrangell Mountains and three from southern Wrangellia on Vancouver Island). Twenty of these species were previously unrecognized (Appendix). Seven new species and one new genus, Campesteria, were identified from localities of Wrangellia and the Alexander terrane (Caru-
thers and Stanley, 2008), three from genus Paracuifia Melnikova, two from genus Retiophyllia Cuif, and one from the genus Gablonzeria Cuif. Paleobiogeographic results were obtained using the Similarity Index of Raup and Crick (1979). Statistical analysis was based on Late Triassic coral faunas described by Caruthers and Stanley (2008), Stanley and Whalen (1989), and Stanley (1994). Similarity was first assessed by locality (not shown) in order to accurately group localities respective of terrane (a preliminary analysis of the dataset is in Caruthers, 2005, fig. 27, table 3). Results indicate (1) no statistical similarity between the Alexander terrane with either the Wallowa terrane, northern Wrangellia, or southern Wrangellia; (2) no similarity between northern and southern Wrangellia; and (3) a high degree of statistical similarity between southern Wrangellia and the Pucará Group, Peru (Table 2A and 2B). Despite having numerous (eight) specieslevel taxonomic coral links between northern Wrangellia and the Wallowa terrane (Caruthers and Stanley, 2008; Table 2), our analysis could not confirm a statistical similarity between these two terranes. However, Newton (1983) and R. Blodgett (2008, personal commun.) maintain this connection based on bivalve and gastropod analyses. Furthermore, it should be noted that many of the coral species (five) linking northern Wrangellia and the Wallowa terrane are considered to be cosmopolitan and are commonly found throughout the Tethys Region and the North American Cordillera. Statistical results agree with previous stratigraphic and lithologic data suggesting separate paleogeographic histories for the Alexander terrane and Wrangellia during the Late Triassic (Jones et al. 1977), as well as with paleobiogeographic analysis of Newton (1983), Aberhan (1999), and Blodgett and Frýda (2001) and with paleomagnetic data from Irving and Yole (1972), Hillhouse (1977), Hillhouse and Gromme (1980), and Butler et al. (1997). Paleobiogeographic coral data from northern and southern Wrangellia reveal no statistical similarity despite having a very broadly similar stratigraphy and lithology (Jones et al., 1977) and therefore do not support an amalgamation in Triassic time. Statistical comparisons of assemblages of different ages may be responsible for dissimilar coral data from northern and southern Wrangellia. The coral fauna from the Wrangell Mountains (northern Wrangellia) has been found to be early Norian, whereas the associated faunas from Vancouver Island (southern Wrangellia) are documented as Rhaetian. This time gap of ~16 m.y. (Gradstein and Ogg, 2004) may have been long enough for evolution to alter Late Triassic coral taxonomy within southern Wrangellia, thus creating two statistically separate coral assemblages.
Conodont biozone Epigondolella quadrata, Metapolygnathus primitius Epigondolella quadrata, E. triangularis Epigondolella bidentata? Epigondolella triangularis triangularis Metapolygnathus primitius, M. sp. cf. M. polygnathiformis Metapolygnathus sp. aff. M. primitius Epigondolella tozeri, E. englandi, E. sp. aff. E. mosheri Metapolygnathus primitius Epigondolella sp. cf. E. quadrata Corals ?Ampakabastraea cf. A. nodosa Cuif Anthostylis cf. A. acanthophora (Frech) Astraeomorpha confusa (Winkler) A. crassisepta Reuss Campesteria prolixia Caruthers and Stanley Chondrocoenia schafhaeutli (Winkler) Ch. cf. C. paradoxa (Melnikova) Crassistella juvavica (Frech) C. parvula (Melnikova) Cuifia cf. C. marmorea (Frech) Distichomeandra austriaca (Frech) D. minor (Frech) D. sp. A D. sp. B Distichophyllia norica (Frech)
Relative Age
O
O O
X O
O
Early Norian
O
X
X
O
O
X
O
Late Late Norian– Carnian– Rhaetian Early Norian
Keku Strait FC GH SWK (M1910/1911) (M1912) (M2136) MI 0099 MI 0087 MI 0074
X
X
X
O
BS (M2135) MI 0056
O
O
O
X
O
Early Norian
X
O
O
O O X
O O
X
X
O
Early-Middle Early Norian Norian
Gravina Island Northern NB PB NC GB (M1708) MI 9930- MI 9933 MI 9935-9938 MI 9934 9932
Appendix
CPE MI 0070
Alexander Terrane
X
X O
X
X
Late Norian– Rhaetian
X
X
O
Late Norian– Rhaetian
Wrangellia Terrane Southern LC PP MI 8302 MI 9947
X
O
O
X
Late Norian– Rhaetian
TS MI 9913
2 10 1 44 Continued
1 42 1 1 4
3
39 7
2
14
3
Specimen Count
Appendix 175
Corals (continued) Gablonzeria profunda (Reuss) G. major (Frech) G. grandiosa Caruthers and Stanley Kompsasteria cf. K. oligocystis (Frech) Kuhnastraea cowichanensis (Clapp and Shimer) K. decussata (Reuss) K. incrassata (Frech) Margarosmilia cf. M. charlyana (Frech) M. cf. M. richthofeni Volz M. confluens (Münster) Meandrostylis grandiseptus Stanley and Whalen M. vancouverensis (Clapp and Shimer) Pamiroseris meriani (Stoppani) P. borealis (Smith) Paracuifia jennieae Caruthers and Stanley P. anomala Caruthers and Stanley P. smithi Caruthers and Stanley Parastraeomorpha cf. P. similis Roniewicz Recticostastraea wallowaensis Stanley and Whalen Retiophyllia alfurica (Wilckens) R. caespitosa (Reuss) R. clathrata (Emmrich) R. dawsoni (Clapp and Shimer) R. dendriformis Caruthers and Stanley R. cf. R. frechi Roniewicz R. cf. R. norica (Frech)
Relative Age
O O
O
O
O
O
O O O
X
O
O
O
O
50 6 1
O
O
O
O X X O
O
O
X
O
O
X X O
X
4
O
3 6 Continued
10 35 25 17 15
2
1
1
11 9 4
28
1 5 2
16 1 4
O
O
O
X
X
X X
Late Norian– Rhaetian
X
X
X X
Late Norian– Rhaetian
12
X
X
Late Norian– Rhaetian
X
O
X
Early-Middle Early Norian Norian
Specimen Count
X
X
Early Norian
TS MI 9913
1
X
Late Late Norian– Carnian– Rhaetian Early Norian
Wrangellia Terrane Southern LC PP MI 8302 MI 9947
O
X
Early Norian
Keku Strait FC GH SWK (M1910/1911) (M1912) (M2136) MI 0099 MI 0087 MI 0074
Appendix (continued) Alexander Terrane Gravina Island Northern BS CPE NB PB NC GB (M1708) (M2135) MI 0070 MI 9930- MI 9933 MI 9935-9938 MI 9934 MI 0056 9932
176 Appendix
Corals (continued) R. obtusa Caruthers and O 1 Stanley R. oppeli (Reuss) X 1 R. parviseptum (Squires) X 6 R. cf. R. robusta Roniewicz O 2 R. tenuicosta (Reuss) O 3 Stylophyllum cf. S. pygmaeum O 1 Frech Total coral species according to 13 n/4 o 2 n/3 o 1 n/2 o 2 n/5 o 1 n/1 o 3 n/2 o 1n 4 n/2 o 13 n/10 o 6 n/9 o 1 n/5 o 2 n/2 o 458 terrane new/old Note: O—not formerly known from terrane; X—found and previously known from terrane; FC—Flounder Cove; GH—Gil Harbor; SWK—Southwest of Kousk; BS—Big Spruce; CPE—Cornwallis Peninsula East; NB—Nehenta Bay; PB—Phocena Bay; NC—Nelson Cove; GB—Green Butte; LC—Lake Cowichan; PP—Pender Point; TS—Tahsis Inlet.
Late Norian– Rhaetian Late Norian– Rhaetian Late Norian– Rhaetian Early-Middle Early Norian Norian Early Norian Late Late Norian– Carnian– Rhaetian Early Norian Early Norian Relative Age
Keku Strait FC GH SWK (M1910/1911) (M1912) (M2136) MI 0099 MI 0087 MI 0074
Appendix (continued) Alexander Terrane Gravina Island Northern BS CPE NB PB NC GB (M1708) (M2135) MI 0070 MI 9930- MI 9933 MI 9935-9938 MI 9934 MI 0056 9932
Wrangellia Terrane Southern LC PP MI 8302 MI 9947
TS MI 9913
Specimen Count
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177
ACKNOWLEDGMENTS We would like to acknowledge a grant from the National Science Foundation (EAR-9624501). We wish to thank Jim Baichtal and the U.S. Forest Service for logistical and field support at Keku Strait and Gravina Island, as well as Erik Katvala for reviewing our manuscript, his detailed fieldwork, sharing his knowledge on the stratigraphy with us, conodont identifications, and his collections of Keku Strait fossils. Further gratitude goes to Devi Sharp and Danny Rosenkrans (National Park Service staff members at Wrangell-St. Elias National Park and Preserve) for arranging transportation to the Green Butte locality. Robert Blodgett is thanked for reviewing an earlier version of our manuscript, for his field collaboration, and for sharing with us his knowledge of Alaskan geology. REFERENCES CITED Aberhan, M., 1999, Terrane history of the Canadian Cordillera: Estimating amounts of latitudinal displacement and rotation of Wrangellia and Stikinia: Geological Magazine, v. 136, no. 5, p. 481–492, doi: 10.1017/ S001675689900299X. Armstrong, A.K., MacKevett., E. M. Jr., and Silberling, N. J., 1969, The Chitistone and Nizina Limestones of part of the southern Wrangell Mountains, Alaska— A preliminary report stressing carbonate petrography and depositional environments: U.S. Geological Survey Professional Paper 650-D, p. D49–D62. Armstrong, A.K., and MacKevett, E.M., Jr., 1982, Stratigraphy and diagenetic history of the lower part of the Triassic Chitistone Limestone, Alaska: U.S. Geological Survey Professional Paper 1212-A, p. 1–26. Bancroft, J.A., 1913, Geology of the coast and islands between the Strait of Georgia and Queen Charlotte Sound, British Columbia: Geological Survey of Canada Memoir 23, 146 p. Beauvais, L., 1980, Sur la Taxonomie des Madréporaires mesozoiques: Acta Palaeontologica, v. 25, p. 345–360. Belasky, P., and Runnegar, B., 1993, Biogeographic constraints for tectonic reconstructions of the Pacific region: Geology, v. 21, p. 979–982, doi: 10.1 130/0091-7613(1993)021<0979:BCFTRO>2.3.CO;2. Belasky, P., and Runnegar, B., 1994, Permian longitudes of Wrangellia, Stikinia, and Eastern Klamath terranes based on coral biogeography: Geology, v. 22, p. 1095–1098, doi: 10.1130/0091-7613(1994)022<1095:PLOWSA> 2.3.CO;2. Belasky, P., Stevens, C.H., and Hanger, R.A., 2002, Early Permian locations of western North American terranes based on brachiopod, fusulinid, and coral biogeography: Palaeogeography: Palaeoclimatology, v. 179, p. 245–266, doi: 10.1016/S0031-0182(01)00437-0. Berg, H.C., 1973, Geology of Gravina Island, Alaska: U.S. Geological Survey Bulletin 1373, 41 p. Berg, H.C., Jones, D.L., and Richter, D.H., 1972, Gravina-Nutzotin Belt— Tectonic significance of an Upper Mesozoic sedimentary and volcanic sequence in southern and southeastern Alaska: U.S. Geological Survey Professional Paper 800-D, p. 1–24. Blodgett, R.B., and Frýda, J., 2001, On the occurrence of Spinidelphinulopsis whaleni [Late Triassic (early Norian) Gastropoda] in the Cornwallis Limestone, Kuiu Island, southeastern Alaska (Alexander terrane) and its paleobiogeographic significance: Bulletin of the Czech Geological Survey, v. 76, no. 4, p. 267–274. Butler, R.F., Gehrels, G.E., and Bazard, D.R., 1997, Paleomagnetism of Paleozoic strata of the Alexander terrane, southeastern Alaska: Geological Society of America Bulletin, v. 109, no. 10, p. 1372–1388, doi: 10.1130/ 0016-7606(1997)109<1372:POPSOT>2.3.CO;2. Carlisle, D., and Susuki, T., 1974, Emergent basalt and submergent carbonateclastic sequences including the Upper Triassic Dilleri and Welleri Zones on Vancouver Island: Canadian Journal of Earth Sciences, v. 11, no. 2, p. 254–279. Caruthers, A.H., 2005, Upper Triassic Carbonates and Scleractinian corals from Wrangellia and the Alexander terrane (Alaska and Vancouver
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of Energy, Mines and Resources, Earth Physics Branch, Publication 42, p. 87–95. Jackson, J.A., ed., 1997, Glossary of geology (fourth edition): Alexandria, Virginia, American Geological Institute, 769 p. Jeletzky, J.A., 1970, Some salient features of early Mesozoic history of insular tectonic belt, western British Columbia: Geological Survey of Canada Paper, 26 p. Jones, D.L., Irwin, W.P., and Ovenshine, A.T., 1972, Southeastern Alaska— A displaced continental fragment?: U.S. Geological Survey Professional Paper 800-B, p. 211–217. Jones, D.L., Silberling, N.J., and Hillhouse, J., 1977, Wrangellia—A displaced terrane in northwestern North America: Canadian Journal of Earth Sciences, v. 14, p. 2565–2577. Katvala, E.C., and Henderson, C.M., 2002, Conodont sequence biostratigraphy and paleobiogeography of the Pennsylvanian-Permian Mount Mark and Fourth Lake formations, southern Vancouver Island: Canadian Society of Petroleum Geologists Memoir 19, p. 461–478. Katvala, E.C., and Stanley, G.D., Jr, 2008, this volume, Conodont biostratigraphy and facies correlations in a Late Triassic Island Arc, Keku Strait, Southeast Alaska, in Blodgett, R.B., and Stanley, G.D., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, doi: 10.1130/2008.442(11). Kiessling, W., and Flügel, E., 2002, Paleoreefs—A database on Phanerozoic reefs, in Kiessling, W., Flügel, E., and Golonka, J., eds., Phanerozoic reef patterns: Tulsa, Oklahoma, SEPM Special Publication 72, p. 77–92. MacKevett, E.M., Jr., 1965, Preliminary geologic map of the McCarthy C-6 quadrangle, Alaska: U.S. Geological Survey Miscellaneous Geologic Investigations Map 1-444, scale 1:63,360. MacKevett, E.M., Jr., 1970, Geologic map of the McCarthy C-5 quadrangle, Alaska: U.S. Geological Survey Geologic Quadrangle Map GQ-899, scale 1:63,360. MacKevett, E.M., Jr., 1974, Geologic map of the McCarthy B-5 quadrangle, Alaska: U.S. Geological Survey Geologic Quadrangle Map GQ-1146, scale 1:63,360. MacKevett, E.M., Jr., 1976, Geologic map of the McCarthy quadrangle, Alaska: U.S. Geological Survey Miscellaneous Field Studies Map MF-733A, scale 1:250,000. Martin, G.C., 1926, The Mesozoic stratigraphy of Alaska: U.S. Geological Survey Bulletin 776, 493 p. Massey, N.W.D., and Friday, S.J., 1987, Geology of the Cowichan Lake area, Vancouver Island (92 C/16): British Columbia Ministry of Energy, Mines and Petroleum Resources, Paper 1987-1, p. 223–229. McRoberts, C.A., 1993, Systematics and biostratigraphy of halobiid bivalves from the Martin Bridge Formation (Upper Triassic), northeast Oregon: Journal of Paleontology, v. 67, p. 198–210. McRoberts, C.A., 1997, Late Triassic North American halobiid bivalves: Diversity trends and circum-Pacific correlations, in Dickens, M.J., Zunyi, Y., Hongfu, Y., Lucas, S.G., and Acharyya, S.K., eds., Late Paleozoic and Early Mesozoic circum-Pacific events and their global correlation: World and Regional Geology, v. 10, p. 198–208. Monger, J.W.H., and Ross, C.A., 1971, Distribution of fusulinaceans in the Western Canadian Cordillera: Canadian Journal of Earth Sciences, v. 8, no. 2, p. 259–278. Monger, J.W.H., Souther, J.G., and Gabrielse, H., 1972, Evolution of the Canadian Cordillera: A plate-tectonic model: American Journal of Science, v. 272, p. 577–602. Montanaro-Gallitelli, E., Russo, A., and Ferrari, P., 1979, Upper Triassic coelenterates of western North America: Bollettino della Società Paleontologica Italiana, v. 18, no. 1, p. 133–156. Muffler, L.J.P., 1967, Stratigraphy of the Keku Islets and neighboring parts of Kuiu and Kupreanof Island Southeastern Alaska: U.S. Geological Survey Bulletin 1241-C, 52 p. Muller, J.E., 1977, Evolution of the Pacific Margin, Vancouver Island, and adjacent regions: Canadian Journal of Earth Sciences, v. 14, no. 9, p. 2062–2085. Muller, J.E., Northcote, K.E., and Carlisle, D., 1974, Geology and mineral deposits of Alert-Cape Scott map-area, Vancouver Island, British Columbia: Geological Survey of Canada Paper 74-8, 77 p. Newton, C.R., 1983, Paleozoogeographic Affinities of Norian bivalves from the Wrangellian, Peninsular, and Alexander Terranes, Western North America, in Stevens, C.H., ed., Pre-Jurassic rocks in western North American suspect terranes: Los Angeles, Pacific Section, SEPM, p. 37–48.
Late Triassic silicified shallow-water corals and other marine fossils Orchard, M.J., 1991, Upper Triassic conodont biochronology and new index species from the Canadian Cordillera, in Orchard, M.J., and McCracken, A.D., eds., Ordovician to Triassic conodont paleontology of the Canadian Cordillera: Geological Survey of Canada Bulletin 417, p. 299–335. Orchard, M.J., and Tozer, E.T., 1997, Triassic conodont biochronology, its calibration with the ammonoid standard, and a biostratigraphic summary for the Western Canada Sedimentary Basin: Bulletin of Canadian Petroleum Geology, v. 45, no. 4, p. 675–692. Panuska, B.C., 1984, Paleomagnetism of the Wrangellia and Alexander terranes and the tectonic history of southern Alaska [Ph.D. thesis]: Fairbanks, University of Alaska, 197 p. Raup, D.M., and Crick, R.E., 1979, Measurement of faunal similarity in paleontology: Journal of Paleontology, v. 53, p. 1213–1227. Roniewicz, E., 1989, Triassic scleractinian corals of the Zlambach beds, northern Calcareous Alps, Austria: Vienna Denkschrift Österreiche Akademie Wissenschaften, v. 126, p. 1–152. Sanders, D., and Baron-Szabo, R.C., 2005, Scleractinian assemblages under sediment input: Their characteristics and relation to the nutrient input concept: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 216, p. 139–181, doi: 10.1016/j.palaeo.2004.10.008. Shimer, H.W., 1926, A Triassic coral reef fauna in British Columbia: Geological Survey of Canada, Museum Bulletin, v. 42, p. 85–89. Silberling, N.J., Grant-Mackie, J.A., and Nichols, K.M., 1997, The Late Triassic Bivalve Monotis in accreted terranes of Alaska: U.S. Geological Survey Bulletin 2151, 21 p. Smith, J.P., 1927, Upper Triassic marine invertebrate faunas of North America: U.S. Geological Survey Professional Paper 141, 262 p. Smith, P.S., 1915, Notes on the geology of Gravina Island, Alaska: U.S. Geological Survey Professional Paper 95, p. 100–104. Soja, C.M., 1996, Island-arc carbonates: Characterization and recognition in the ancient geologic record: Earth-Science Reviews, v. 41, p. 31–65, doi: 10.1016/0012-8252(96)00029-3. Stanley, G.D., Jr., 1979, Paleoecology, structure and distribution of Triassic coral buildups in western North America: University of Kansas Paleontological Contributions, v. 65, 68 p. Stanley, G.D., Jr., 1989, An Upper Triassic reefal limestone, southern Vancouver Island, British Columbia: Canadian Society of Petroleum Geologists Memoir 13, p. 766–775. Stanley, G.D., Jr., 1993, Volcanic island reefs from circum-Pacific terranes: Canadian Society of Petroleum Geologists Abstracts with Programs, Calgary, Alberta, p. 298.
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Stanley, G.D., Jr., 1994, Upper Triassic corals from Peru: Palaeontographica, Abt. A, v. 233(1–6), p. 75–98. Stanley, G.D., Jr., 1996, Paleobiology and biology of corals: Introduction, in Stanley, G.D., Jr., ed., Paleobiology and biology of corals: Pittsburgh, Pennsylvania, Paleontological Society, p. 3–7. Stanley, G.D., Jr., 2001, Introduction to reef ecosystems and their evolution, in Stanley, G.D., Jr., ed., The history and sedimentology of ancient reef systems: New York, Kluwer Academic/Plenum Publishers, Topics in Geobiology, v. 17, p. 1–39. Stanley, G.D., Jr., and Whalen, M.T., 1989, Triassic corals and spongiomorphs from Hells Canyon, Wallowa Terrane, Oregon: Journal of Paleontology, v. 63, no. 6, p. 800–819. Stone, D.B., 1981, Triassic paleomagnetic data and paleolatitudes for Wrangellia, Alaska: Short notes on Alaskan Geology: Alaska Division of Geological and Geophysical Surveys Geologic Report, v. 73, p. 55–62. Tozer, E.T., 1967, A standard for Triassic time: Geological Survey of Canada Bulletin 156, 103 p. Tozer, E.T., 1984, The Trias and its ammonoids: The evolution of a time scale: Geological Survey of Canada Miscellaneous Report 35, 171 p. Tozer, E.T., 1994, Canadian Triassic ammonoid faunas: Geological Survey of Canada Bulletin 467, 663 p. Volz, W., 1896, Die Korallenfauna der Trias. II. Die Korallen der Schichten von St-Cassian in Süd-Tyrol: Palaeontographica, v. 43, p. 1–124. Wilson, J.T., 1968, Static or mobile earth; the current scientific revolution: Proceedings of the American Philosophical Society, v. 112, no. 5, p. 309–320. Wood, R., 1999, Reef evolution: Oxford, UK, Oxford University Press, 414 p. Yarnell, J.M., 2000, Paleontology of two North American Triassic reef faunas: Implications for terrane paleogeography [M.S. thesis]: Missoula, University of Montana, 141 p. Yorath, C.J., Sutherland Brown, A., and Massey, N.W.D., 1999, Lithoprobe, Southern Vancouver Island, British Columbia: Geology: Geological Survey of Canada Bulletin 498, p. 45–58. Zonneveld, J.P., Gingras, M.K., Orchard, M.J., Stanley, G.D., Jr., Blakney, B.J., and Henderson, C.M., 2002, Triassic reefs of the Canadian Rocky Mountain front ranges: Recovery of hard-bottom communities in the aftermath of the Permian-Triassic extinction: 16th International Sedimentological Congress, Johannesburg, South Africa, 8–12 July 2002, Abstract Volume, p. 425.
MANUSCRIPT ACCEPTED BY THE SOCIETY 14 DECEMBER 2007
Printed in the USA
The Geological Society of America Special Paper 442 2008
Conodont biostratigraphy and facies correlations in a Late Triassic island arc, Keku Strait, southeast Alaska Erik C. Katvala University of Calgary, Dept. of Geoscience, 2500 University Drive NW, Calgary, Alberta T2N 4M1, Canada George D. Stanley Jr. The University of Montana, Dept. of Geosciences, 32 Campus Drive #1296, Missoula, Montana 59812 USA ABSTRACT Upper Triassic rocks in the Keku Strait area of southeast Alaska record a variety of facies in an intra-arc setting. The Hyd Group consists of the Burnt Island Conglomerate, Keku sedimentary strata, Cornwallis Limestone, Hamilton Island Limestone, and the Hound Island Volcanics. The Burnt Island Conglomerate represents initial infill of the basin and underlies the Hamilton Island Limestone, which is coeval with the Cornwallis Limestone and Keku sedimentary strata. Volcanic and sedimentary rocks of the Hound Island Volcanics overlie the entire area. An improved biostratigraphic framework indicates deposition from early Carnian through late Norian time. Conodonts originating in the late Carnian include Metapolygnathus polygnathiformis, Metapolygnathus carpathicus, Metapolygnathus nodosus, Metapolygnathus sp. cf. M. reversus, Metapolygnathus sp. aff. M. zoae, Metapolygnathus sp. aff. M. nodosus, and Metapolygnathus primitius. Early Norian conodonts include Epigondolella quadrata, Epigondolella triangularis, Epigondolella sp. aff. E. triangularis, and the longer-ranging Neogondolella sp. and Misikella longidentata. Middle Norian conodonts include Epigondolella spiculata, Epigondolella transitia, Epigondolella matthewi, Epigondolella postera, and Neogondolella steinbergensis. Late Norian conodonts include Epigondolella bidentata, Epigondolella englandi, Epigondolella sp. aff. E. mosheri, and Epigondolella tozeri. This study resulted in three major accomplishments. Reworked Paleozoic conodonts in Upper Triassic rocks, combined with geologic evidence, suggest major pre– Late Triassic uplift due to compressional tectonics. Late Carnian and early Norian ages support the correlation between the Keku sedimentary strata, shallow-marine limestone of the Cornwallis Limestone, and deeper-water limestone of the Hamilton Island Limestone. Precise conodont biostratigraphy establishes the base of the Hound Island Volcanics as late early Norian, within the Epigondolella triangularis Zone. Keywords: Alexander terrane, conodont biostratigraphy, Late Triassic, southeast Alaska.
Katvala, E.C., and Stanley, G.D., Jr., 2008, Conodont biostratigraphy and facies correlations in a Late Triassic island arc, Keku Strait, Southeast Alaska, in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 181–226, doi: 10.1130/2008.442(11). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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INTRODUCTION Alaska and the North American Cordillera are composed predominantly of allochthonous tectonostratigraphic terranes that accreted to the continental margin during the Mesozoic and Cenozoic (Coney et al., 1980; Jones et al., 1983). While a unique internal stratigraphy defines tectonostratigraphic terranes, tectonic processes shape them geographically, structurally, and depositionally (Coney et al., 1980; Saleeby, 1983). The tectonic fragment known as the Alexander terrane encompasses most of southeast Alaska, as well as parts of western British Columbia, southwestern Yukon, and eastern Alaska (Fig. 1) (Berg et al., 1972; Jones et al., 1972; Gehrels and Saleeby, 1987). The Alexander terrane is a displaced continental fragment that was a separate tectonic entity throughout much of the Phanerozoic
Figure 1. Generalized map of western North America showing position of Alexander and Wrangellia terranes (modified from Jones et al., 1972, 1977).
(Wilson, 1968; Monger and Ross, 1971; Jones et al., 1972; Monger et al., 1972). Its distinctive foundation of late Proterozoic (Gehrels, 1990) and early Paleozoic continental crust sets it apart from neighboring terranes (Berg et al., 1972; Jones et al., 1972). Unlike other terranes, rock of every Phanerozoic period occurs in the Alexander terrane, providing a long geologic record for interpretation (Gehrels and Saleeby, 1987). In the center of southeast Alaska, Keku Strait lies between Kuiu and Kupreanof islands (Fig. 2). Tidal activity exposes bedrock along most shorelines, and shorelines are abundant on the many smaller islands in the strait. Outcrop along shorelines is readily accessible by boat, and newer roads permit access to the less-exposed inland outcrop on both Kuiu and Kupreanof islands. The area includes rocks ranging in age from Late Silurian through Tertiary and encompasses the most complete stratigraphic section in southeast Alaska (Muffler, 1967). Wright and Wright (1908) were the first to comment on fossils and stratigraphy in the Keku Strait area. Atwood (1912) followed up their report and noted some of the Permian and Triassic fossils from the Hamilton Bay region south of Kake. Martin (1916) also made a few notes on the Triassic fossils and stratigraphy of Kupreanof Island. Smith (1927) followed this with a formal description of some of the Keku Strait fossils. Buddington
Figure 2. Map of southeast Alaska with major and referenced islands. Keku Strait map area is outlined (after Muffler, 1967; Berg, 1973).
Conodont biostratigraphy and facies correlations in a Late Triassic island arc, Keku Strait and Chapin (1929) included the first detailed description of the Keku Strait area in their compilation of the geology of southeast Alaska. Many years later, Muffler (1967) published an even more detailed description and map focusing solely on the geology in the Keku Strait area. Furthermore, this paper also provided the first Triassic biostratigraphy of the region (Silberling in Muffler, 1967). These studies primarily used shoreline geology, as road access to the interior of the larger islands was then extremely limited and the creeks are typically too vegetated to provide exposed strata. Upper Triassic units crop out throughout the north end of the Keku Strait, recording deposition over a variety of different environments. These deposits are part of a larger northwest-trending belt in southeast Alaska (Gehrels and Berg, 1984) that overlies both the Admiralty and Craig subterranes of the Alexander terrane (Berg et al., 1978; Van Nieuwenhuyse, 1984; Gehrels et al., 1987; Gehrels and Berg, 1994). These units unconformably overlie Paleozoic rock and are in turn overlain by Jurassic, Cretaceous, or younger rock (Muffler, 1967; Berg et al., 1972; Rubin and Saleeby, 1991). In the Late Triassic, volcanic rock, volcaniclastic and lithoclastic sedimentary rock, and limestone record proximal and distal marine environments, and possibly terrestrial environments as well. The lithologies, sedimentary structures, and preserved fossil biota are consistent with deposition in an island-arc setting (see Soja, 1996). Abundant fossils, especially in limestone, provide a strong basis for biostratigraphic correlation. In the field area, deposition began in the Carnian and continued through the late Norian. Muffler (1967) assigned the Upper Triassic rocks to the Hyd Group and the Keku Volcanics. New age data suggest that the felsic igneous rock of the Keku Volcanics is Cretaceous in age (Mortenson, 2004, personal commun.). However, sedimentary beds previously assigned to the Keku Volcanics occur in succession with other Triassic units and are therefore still a part of the Upper Triassic Hyd Group. Briefly, the previously described units within the Hyd Group are as follows (Fig. 3): • Sedimentary rock of the Keku Volcanics: Bedded lithoclastic sandstone and conglomerate derived from underlying units; neptunian dikes; • Burnt Island Conglomerate: Poorly bedded to massive basal conglomerate with angular and rounded clasts derived from underlying Paleozoic units; • Cornwallis Limestone: Notably oolitic limestone that is commonly fossiliferous and has beds with variable amounts of sand to cobble sized clasts; • Hamilton Island Limestone: Very thinly bedded aphanitic limestone with subordinate argillaceous laminae and sandy intervals; • Hound Island Volcanics: Mainly basaltic pillow lava, basaltic pillow breccia, massive basalt, andesitic volcanic breccia, and hyaloclastic tuff with subordinate tuffaceous polymict conglomerate, limestone, and sandstone. Additional studies in the Keku Strait area have focused on individual portions of the geology, most commonly parts of the
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Figure 3. Simplified stratigraphic column for the Upper Triassic Hyd Group in the Keku Strait area.
Triassic stratigraphy. Brew and Muffler (1965) and Muffler et al. (1969) studied undevitrified volcanic glass in the Hound Island Volcanics. Several publications documented Triassic marine invertebrates from Keku Strait for taxonomic and paleogeographic studies (ammonoids and bivalves: Silberling and Tozer, 1968; corals: Montanaro-Gallitelli et al., 1979; bivalves: Newton, 1983; brachiopods: Hoover, 1991; bivalves: Silberling et al., 1997; bivalves: McRoberts and Blodgett, 2000; gastropods: Blodgett and Frýda, 2001; gastropods: Frýda and Blodgett, 2001). Hillhouse and Grommé (1980), and Haeussler et al. (1992) examined paleomagnetic data in the Hound Island Volcanics. Karl et al. (1999) mapped the area to the east of Keku Strait, including many adjacent Triassic outcrops. Finally, abundant mineral deposits in the Keku Strait area related to the Keku Volcanics have been studied by a number of authors (Berg, 1981; Taylor et al., 1995; McDonald et al., 1998; Bittenbender et al., 2000; Still et al., 2002). Muffler (1967) and Orchard et al. (2001) emphasized the importance of age data in the terranes. Without a temporal scale, understanding structural and tectonic histories is extremely difficult. In terranes, this problem is compounded by incoherent regional stratigraphy, abrupt facies changes, and structural telescoping of dissimilar stratigraphic profiles (Orchard et al., 2001). The major goal of this study was to refine the Triassic biostratigraphic data of the region, primarily by using conodonts. This work was designed to improve stratigraphic correlation within the area, refine Triassic biostratigraphy in the Alexander terrane, and build a stronger foundation for future biostratigraphic and lithostratigraphic research of Triassic age rocks in the terranes. Because the units in the field area primarily crop out along shorelines, upper and lower contacts are often not present, making
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biostratigraphic data essential to stratigraphic interpretation. Furthermore, this study is associated with ongoing research on Late Triassic invertebrate fossils and their evolution, survival from mass extinction, and biogeographic potential in tectonostratigraphic terranes. An improved biostratigraphic framework will assist paleontologic determinations with Late Triassic invertebrates. Over a total of six weeks during the summers of 2001, 2002, and 2003, we visited Triassic rocks throughout the Keku Strait area (Figs. 4 and 5; Table 1). Past studies (Buddington and Chapin, 1929; Muffler, 1967) guided the selection of sites for paleontological sampling, though we also located new sites (Table 1). Conodont samples were collected from measured sections or more frequently as individual site samples. Lab techniques for conodont recovery included acetic acid dissolution and heavy-liquid separation. Macrofossils were collected whenever encountered. Many silicified macrofossils were recovered by etching blocks mainly in acetic acid, but sometimes in
hydrochloric acid. Biostratigraphically significant macrofossils, including those published in past studies (Muffler, 1967), are included in this paper to enhance or supplement the conodont biostratigraphy. Although the focus of the work was collection of paleontological samples, we examined each geologic unit throughout the study area and collected samples from every major lithology encountered. LITHOSTRATIGRAPHY AND PALEONTOLOGY Loney (1964) defined the Hyd Formation to the north of the Keku Strait area on Admiralty Island. With the better exposures in the Keku Strait area, Muffler (1967) raised the Hyd Formation to group status and subdivided it into four formations. These include the Burnt Island Conglomerate, Cornwallis Limestone, Hamilton Island Limestone, and Hound Island Volcanics. These formations loosely correlate with Loney’s (1964) basal breccia,
Figure 4. Map of Keku Strait with generalized distributions of Paleozoic and Triassic rock, localities visited, and areas of reference for Triassic rocks (modified from Muffler, 1967).
Figure 5 (on this and following pages). Larger-scale maps of areas represented in Figure 4. Dots represent visited Triassic localities, and Loc numbers are University of Montana Museum of Paleontology locality numbers. Corresponding USGS numbers are given when appropriate. Trb—Burnt Island Conglomerate; Trk—Keku Volcanics; Trc—Cornwallis Limestone; Trh—Hamilton Island Limestone; Trv—Hound Island Volcanics. Outcrop extent is unchanged from Muffler (1967).
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Figure 5 (on this and following pages).
limestone, argillite, and volcanic members. Muffler (1967) also defined the Keku Volcanics, a unit mainly comprised of felsic volcanics that was inferred to underlie and partially interfinger with the Hyd Group. However, recently acquired age data from U-Pb zircon ratios measured from several samples of the Keku Volcanics indicate a Cretaceous age (Mortenson, 2004, personal commun.). Thus, the felsic igneous rock of the Keku Volcanics may have a Cretaceous intrusional origin. Sedimentary deposits formerly included in the Keku Volcanics occur in succession with the Cornwallis Limestone. Accordingly, in the descriptions below we remove specific sedimentary units from the Keku Volcanics and associate them with either the newly described Keku sedimentary strata or the Cornwallis Limestone in the Hyd Group.
The focus of this study was on sedimentary, particularly carbonate, units. Basic descriptions of Triassic units are included below to familiarize the reader. The descriptions presented follow Muffler (1967) and supplement the original definitions with new information and interpretations. Eroded Paleozoic rocks provided the detrital source for many Triassic units. The Paleozoic Cannery and Pybus formations are recognizable as clasts, and basic descriptions of these units are included below to provide familiarity. Fossils discovered during the course of this work are curated in the University of Montana Museum of Paleontology, and Table 2 summarizes the major groups of Triassic fossils discovered in the Keku Strait area. Biostratigraphically significant
Figure 5 (on this and following pages).
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Figure 5. (continued).
fossils, including conodonts, ammonoids, and halobiid and monotiid bivalves, are discussed thoroughly in the section on biostratigraphy. Paleozoic Units Loney (1964) named the Cannery Formation for Permian exposures on southeastern Admiralty Island. This formation crops out throughout northeastern Kupreanof Island (Fig. 4), and currently comprises units of Devonian to Carboniferous (Buddington and Chapin, 1929; Jones et al., 1981), possibly Permian (Dutro in Muffler, 1967), age. In the Keku Strait area, the main rock type is thin-bedded, tuffaceous volcanic siltstone to sandstone with locally occurring chert, limestone, and pillow flows (Muffler, 1967). The siltstone and sandstone characteristically weather blue-green or reddish-brown in color and are intensely fractured (Muffler, 1967). Loney (1964) defined the Pybus Dolomite for Permian exposures on southeastern Admiralty Island. Muffler (1967) renamed it the Pybus Formation and included the Permian outcrops of distinctively white limestone, dolomite, and chert in the Keku Strait area and on Admiralty Island. Silicified crinoids, bryozoans, and brachiopods are common. In the Keku Strait area, the Pybus Formation crops out near Hamilton Bay, inland of Cape Bendel, throughout the Keku Islets, and along the west side of Cornwallis Peninsula (Fig. 4). Triassic Units Burnt Island Conglomerate Occurrence: The type locality is the small islands (reefs) below mean high tide between Burnt Island and Grave Island south of Kake (Fig. 5F) (Muffler, 1967). The Burnt Island
Conglomerate also occurs west of Kupreanof Island from Kake down into Hamilton Bay (Fig. 4), in the Cape Bendel region, and on a few of the northern Keku Islets (Figs. 5C, 5E, 5F, and 5G). Around Cape Bendel, Muffler (1967) noted the Burnt Island Conglomerate on the southern end of the Triassic outcrop, though this formation also occurs on the northeast end (Buddington and Chapin, 1929), and further inland between the Paleozoic outcrop and the Hound Island Volcanics (Fig. 5E). Description: This unit is mainly a poorly bedded to massive, grain-supported conglomerate (Fig. 6). The conglomerate is dominated by rounded and angular clasts derived from the underlying Paleozoic units (Muffler, 1967). Buddington and Chapin (1929) likened this unit to both conglomerate and breccia on the basis of the variable rounding of clasts. At the type locality, most grains are pebbles (Muffler, 1967), but sand to boulder-sized clasts also occur. More resistant rock types, such as chert and limestone clasts from the Pybus Formation, typically form larger pieces than less resistant lithologies of the Cannery Formation; these larger clasts are commonly somewhat rounded. The largest clasts observed were on the island east of Burnt Island and in the Cape Bendel region (Fig. 7), though Atwood (1912) observed boulders up to about a meter in diameter from the south end of Hamilton Island. Overall, the rounding and size range of clasts suggest a variable degree of recycling before deposition. Muffler (1967) also described a 1.5-m-thick, fine-grained limestone bed in the lower part of this unit (Fig. 3) just north of little Hamilton Island at U.S. Geological Survey (USGS) Mesozoic locality M1982 (Fig. 5G). This limestone bed does not match the rest of the Burnt Island Conglomerate lithologically, though it is the oldest Triassic deposit in the field area and helps to constrain the age of the unit. Originally, Muffler (1967)
TABLE 1. LOCALITY TABLE WITH UNIVERSITY OF MONTANA MUSEUM OF PALEONTOLOGY LOCALITY NUMBERS Locality no.
USGS no.
Formation
Locality name
N Latitude
W Longitude
Age
MI 0040
–
Cannery
Watershed Rocks—sample 1
56°58.812'
133°55.591'
Permian?
MI 0041
–
Cannery
Watershed Rocks—sample 2
56°58.991'
133°53.455'
Permian?
MI 0042
–
Cannery
Watershed Rocks—sample 3
56°59.547'
133°52.395'
Permian?
MI 0043
–
Cannery
Watershed Rocks—sample 4
57°00.678'
133°51.515'
Permian?
MI 0044
–
Cannery
Watershed Rocks—sample 5
57°01.796'
133°51.384'
Permian?
Keku Volcanics
56°55.069'
134°11.039'
Cretaceous? Late Triassic
MI 0045 MI 0046
–
Hound Island Volcanics
Cornwallis Peninsula East, Keku/Cornwallis Transition—site 1 Black and Tan
56°49.372'
133°57.859'
MI 0047
–
Burnt Island Conglomerate
Island East of Burnt Island
56°56.898'
133°55.959'
Late Triassic
MI 0050
–
Saginaw Bay
Trickling Cave
56°53.172'
134°03.480'
Carboniferous
MI 0051
–
Saginaw Bay
Samtron Monitor Island
56°55.399'
134°08.546'
Carboniferous
MI 0052
–
Muffler's crinoidal limestone
56°54.908'
134°09.886'
Carboniferous?
MI 0053
–
Cannery
Cornwallis Peninsula East, east of the Elephant's Head Hamilton Island Northeast, Cannery Formation White Rock Road Pit
56°56.572'
133°55.104'
Permian?
56°54.534'
133°44.010'
Permian
Neptunian Dike and M1918
56°53.347'
134°04.380'
Late Triassic
Big Spruce Island
56°55.400'
134°09.217'
Late Triassic
Hamilton Island Northeast, Burnt Island Conglomerate Portage Pass, Burnt Island Conglomerate—site 3 Portage Pass, Burnt Island Conglomerate site 4
56°56.671'
133°55.435'
Late Triassic
56°55.031'
133°51.963'
Late Triassic
56°54.875'
133°51.545'
Late Triassic
MI 0054
–
Pybus
MI 0055
1918
MI 0056
2135
MI 0057
–
Keku Volcanics/Cornwallis Limestone Keku Volcanics/Cornwallis Limestone Burnt Island Conglomerate
MI 0058
–
Burnt Island Conglomerate
MI 0059
–
MI 0060
Hamilton Island Northeast NE Shore
56°56.630'
133°55.437'
Late Triassic
MI 0061
1882– 1884 1924
Burnt Island Conglomerate/unnamed shallow water limestone Hamilton Island Limestone Hamilton Island Limestone
Payne Island North, M1924
56°56.571'
134°08.675'
Late Triassic
MI 0062
1904
Payne Island Southwest
56°56.339'
134°06.832'
Late Triassic
MI 0063
–
Burnt Island Conglomerate/Hamilton Island Limestone Hamilton Island Limestone
56°56.207'
134°06.809'
Late Triassic
MI 0064
1903
Hamilton Island Limestone
Small Island south of Payne Island Southwest Payne Island Southwest, M1903
56°56.336'
134°07.454'
Late Triassic
MI 0065
2126?
Hamilton Island Limestone
Top Cathedral Falls
56°54.103'
133°43.281'
Late Triassic
MI 0066
1932
Hamilton Island Limestone
Portage Pass
56°55.218'
133°52.491'
Late Triassic
MI 0067
1889
Hamilton Island Limestone
Hamilton Island Southeast
56°54.559'
133°51.118'
Late Triassic
MI 0068
–
Cornwallis Limestone
Floating Unknown
56°52.838'
134°02.406'
Late Triassic
MI 0069
–
Cornwallis Limestone
Cornwallis Peninsula East—June 22
56°56.215'
134°15.180'
Late Triassic
MI 0070
1906
Cornwallis Limestone
Cornwallis Peninsula East, M1906
56°56.099'
134°14.534'
Late Triassic
MI 0071
–
Cornwallis Limestone
Kuiu Island East
56°52.922'
134°01.397'
Late Triassic
MI 0072
–
Cornwallis Limestone
Kuiu Island East-A
56°52.726'
134°01.336'
Late Triassic
MI 0073
–
Cornwallis Limestone
K(?) and K(not)
56°52.770'
134°02.203'
Late Triassic
MI 0074
2136
Cornwallis Limestone
Southwest of Kousk Island
56°53.690'
134°00.508'
Late Triassic
MI 0075
–
Hound Island Volcanics
Hamilton Island West—site 3
56°55.653'
133°54.135'
Late Triassic
MI 0076
–
Hound Island Volcanics
Hamilton Island West—site 4
56°55.885'
133°54.636'
Late Triassic
MI 0077
–
Hound Island Volcanics
Hamilton Island West—site 5
56°55.920'
133°54.644'
Late Triassic
MI 0078
Hound Island Volcanics
Hamilton Island West—site 6
56°56.260'
133°55.125'
Late Triassic
MI 0079
1886– 1887 1899
Hound Island Volcanics
Hound Island West, M1899
56°52.798'
133°56.805'
Late Triassic
MI 0080
1923
Hound Island Volcanics
Hound Island West, M1923
56°52.323'
133°56.447'
Late Triassic Continued
190
Katvala and Stanley TABLE 1. LOCALITY TABLE WITH UNIVERSITY OF MONTANA MUSEUM OF PALEONTOLOGY LOCALITY NUMBERS (continued)
Locality no.
USGS no.
Formation
Locality name
N Latitude
W Longitude
MI 0081 MI 0082 MI 0083
Age
1921
Hound Island Volcanics
Hound Island West, M1921
56°52.017'
133°56.075'
Late Triassic
–
Hound Island Volcanics
Hound Island North
56°53.132'
133°56.848'
Late Triassic
1913
unnamed shallow water limestone Hamilton Island Limestone
Cape Bendel Day 2
57°04.176'
134°00.905'
Late Triassic
Cape Bendel Day 2-A
57°04.155'
134°00.907'
Late Triassic
Hound Island Volcanics
Hamilton Island Southwest
56°54.696'
133°52.330'
Late Triassic
Hound Island Volcanics
Hound Island East
56°52.635'
133°56.011'
Late Triassic
MI 0084
–
MI 0085
MI 0087
1890/ 1921 1900– 1901 1912
Hound Island Volcanics
Gil Harbor
56°50.050'
134°00.717'
Late Triassic
MI 0088
–
Cannery
Cape Bendel Day 1—site 1
57°02.147'
134°00.697'
Permian?
MI 0089
–
Pybus
Cape Bendel Day 1—site 2
57°02.296'
134°00.652'
Permian
MI 0090
–
Burnt Island Conglomerate
Cape Bendel Day 1—site 3
57°02.448'
134°00.729'
Late Triassic
MI 0091
–
Cape Bendel Day 1—site 4
57°02.599'
134°00.805'
MI 0092
–
Cannery?/ Hamilton Island Limestone Hound Island Volcanics
Cape Bendel Day 1—site 5
57°02.893'
134°00.878'
Permian?/Late Triassic Late Triassic
MI 0093
–
Cannery?
Cape Bendel Day 1—site 6
57°03.763'
133°59.112'
Permian?
MI 0094
–
Hound Island Volcanics
Cape Bendel Day 1—site 9
57°02.682'
134°01.154'
Late Triassic
MI 0095
–
Cornwallis Limestone
56°55.040'
134°11.011'
Late Triassic
MI 0096
–
Cornwallis Limestone
56°55.058'
134°11.038'
Late Triassic
MI 0097
–
56°55.744'
134°05.260'
Permian/Late Triassic
MI 0098
1928
56°55.021'
133°52.468'
Late Triassic
MI 0099
1910– 1911 –
Pybus/Burnt Island Conglomerate/Hamilton Island Limestone Hamilton Island Limestone/ Hound Island Volcanics Keku Volcanics/ Cornwallis Limestone Burnt Island Conglomerate?
Cornwallis Peninsula East, Keku/Cornwallis Transition—site 3 Cornwallis Peninsula East, Keku/Cornwallis Transition—site 2 Squawking Crow
56°51.643'
134°00.661'
Late Triassic
56°55.223'
133°38.259'
Late Triassic?
MI 0086
MI 0100
Hamilton Island Southwest, Hamilton/Hound Transition Flounder Cove
MI 0101
–
Cannery?
Metamorphic 2002: Metaconglomerate rock pit Metamorphic 2002: phyllite A
56°52.940'
133°37.379'
Permian?
MI 0102
–
Keku Volcanics?
Metamorphic 2002: Rhyolite
56°55.756'
133°32.875'
Cretaceous?
MI 0103
–
Cannery?
Metamorphic 2002: phyllite B
56°55.665'
133°29.122'
Permian?
also included limestone beds basal to the Hamilton Island Limestone because of the presence of conglomerate beds among them. However, these limestone beds show closer ties to the depositional environment of the Hamilton Island Limestone and have been reassigned to that formation. In most places, the Burnt Island Conglomerate overlies Paleozoic rock by an erosional unconformity. Either the lowest limestone bed of the Hamilton Island Limestone or volcanic rock of the Hound Island Volcanics overlies this formation. Paleontology: Triassic fossils are extremely rare in this unit. The early Carnian ammonoid Coroceras and the bivalve Halobia sp. cf. H. rugosa were reported from the limestone bed at USGS Mesozoic locality M1892 (Silberling in Muffler, 1967) (Figs. 3 and 5G). This constrains the age of the Burnt Island Conglomerate to early and late Carnian, as upper Carnian rocks of the Hamilton Island Limestone overlie the unit on Hamilton Island. Clasts eroded from the Pybus Formation also contain Permianage crinoids, bryozoans, and brachiopods.
Interpretation: This basal conglomerate represents rapid infill above the pre–Late Triassic unconformity in the region. Triassic basal conglomerates in the Cornwallis Limestone and the Keku sedimentary strata probably correlate laterally, as they also overlie the regional unconformity. However, they are generally thin and of a different character. Thus, they are described with their respective units. Keku Sedimentary Strata Occurrence: Muffler (1967) defined the Keku Volcanics as the felsic igneous rock on Cornwallis Peninsula now thought to be Cretaceous in age. Hence, Triassic sedimentary units that are associated with these volcanics on Cornwallis Peninsula do not have a type locality or name. Extensive lithoclastic conglomerate and sandstone crop out on the island informally referred to as Big Spruce Island (Fig. 5A). Lithoclastic beds also occur inland on Cornwallis Peninsula (Blodgett and Caruthers, 2003, personal commun.) and in scattered locations on the eastern shore of Corn-
Conodont biostratigraphy and facies correlations in a Late Triassic island arc, Keku Strait
Location name Neptunian Dike and M1918 Big Spruce Island
TABLE 2. TRIASSIC MACROFOSSIL GROUPS RECOVERED FROM THE KEKU STRAIT AREA Loc. no. Fossil groups Fm.
191
Age
55
algae?, aulacocerids, ammonoids, gastropods
C
late Carnian or early Norian
56
C
60
ammonoids, brachiopods, corals, crinoids, echinoids, gastropods, oysters, sponges, spongiomorphs, Stromatomorpha ammonoids, aulacocerids, halobiid bivalves, nautiloids, wood
HL
late Carnian? to early Norian late Carnian
Hamilton Island Northeast Payne Island North, M1924 Small Island south of Payne Island SW Payne Island Southwest, M1903 Top Cathedral Falls
61
ammonoids, halobiid bivalves
HL
late Carnian
63
ammonoids, halobiid bivalves
HL
–
64
halobiid bivalves
HL
late Carnian or early Norian
65
halobiid bivalves
HL
late Carnian or early Norian
Portage Pass
66
ammonoids, halobiid bivalves, nautiloids, wood
HL
late Carnian to early Norian
Hamilton Island Southeast Cornwallis Peninsula East Cornwallis Peninsula East, M1906 Kuiu Island East-A
67
ammonoids, halobiid bivalves
HL
late Carnian
69
brachiopods, corals, crinoids, Stromatomorpha, wood
C
early Norian
70
C
72
branching algae, "shelly" bivalves, brachiopods, corals, crinoids, gastropods, oysters, sponges, Stromatomorpha, wood ammonoids, brachiopods, "shelly" bivalves, corals, crinoids, wood
C
late Carnian or early Norian to early Norian late Carnian or early Norian
K(?) and K(not)
73
leaf and plant matter, wood
C
late Carnian or early Norian
Southwest of Kousk Island
74
C
early Norian
Hamilton Island West— Site 5 Hamilton Island West— Site 6 Hound Island West, M1899 Hound Island West, M1923 Hound Island West, M1921 Hound Island North
77
branching algae, “shelly" bivalves, bone?, brachiopods, corals, crinoids, echinoids, gastropods, oysters, sponges, spongiomorphs, Stromatomorpha brachiopods
HV
–
78
halobiid bivalves
HV
–
79
aulacocerids, halobiid bivalves, trace fossils
HV
early Norian
80
halobiid bivalves
HV
early Norian
81
halobiid bivalves, trace fossils
HV
early Norian
82
trace fossils
HV
early Norian
Cape Bendel, day 2
83
ammonoids, halobiid bivalves, brachiopods, crinoids
un
late Carnian
Cape Bendel, day 2-A
84
halobiid bivalves
HL
late Carnian
Hamilton Island Southwest Hound Island East
85
halobiid bivalves
HV
early Norian
86
ammonoids, halobiid bivalves, ichthyosaur bone
HV
Gil Harbor
87
HV
Squawking Crow
97
ammonoids, aulacocerids, halobiid and monotiid bivalves, "shelly" bivalves, brachiopods, corals, crinoids, echinoids, gastropods, Heterastridium, oysters, sponges, trace fossils, wood ammonoids, halobiid bivalves, brachiopods
late Carnian? or early Norian? and middle Norian late Norian
HL
late Carnian to early Norian
Flounder Cove
99
ammonoids, halobiid bivalves, “shelly" bivalves, bone, brachiopods, C early Norian corals, crinoids, echinoids, gastropods, nautiloids, oysters, sponges, spongiomorphs, Stromatomorpha, trace fossils, wood Note: Loc. no.—University of Montana Museum of Paleontology locality. Formations (Fm.): C—Cornwallis Limestone; HL—Hamilton Island Limestone; HV—Hound Island Volcanics; un—unnamed shallow-water limestone.
wallis Peninsula (Figs. 5A and 5B). The lateral extent of these beds is uncertain. A single site on Cornwallis Peninsula (Fig. 5B, site 55) has neptunian dikes that were included in the Keku Volcanics by Muffler (1967).
Description: Lithoclastic sandstone and granule to cobble conglomerate of the Keku sedimentary strata are mainly composed of clasts eroded from underlying units. The clasts are mostly chert and limestone, and limestone cobbles occur near the
192
Katvala and Stanley
Figure 6. Cut slab of pebble conglomerate in the Burnt Island Conglomerate from the island east of Burnt Island (Fig. 5F, site 47). Scale is in centimeters.
Figure 7. Large Pybus Formation clast in the Burnt Island Conglomerate on the small island east of Burnt Island (Fig. 5F, site 47). Person and hammers (33 cm length) for scale.
base of the unit (Muffler, 1967). Exposures are not common on the shores of Keku Strait, but a thick exposure of these strata on Big Spruce Island (Fig. 5A, site 56; Fig. 8) contains limestone and chert fragments eroded from Paleozoic units. These exposures consist of bedded sandstone and pebble conglomerate with intermittent cobbles containing abundant sedimentary scours (Fig. 9). A large, angular boulder of crinoidal Carboniferous limestone lies within these units (Fig. 10).
Neptunian dikes occur below the limestone beds at USGS Mesozoic locality M1918 (Fig. 5B, site 55; Fig. 11). These dikes are sedimentary fillings in cracks formed in underlying rock, and contain abundant fossils in a siliceous mud matrix. The base of the unit appears to be a widespread chert- and limestone-clast conglomerate overlying the erosional unconformity (Muffler, 1967). In places, the unit interfingers with or is conformably overlain by the Cornwallis Limestone. Although it has not been observed, it is likely that volcanic rock of the widely distributed Hound Island Volcanics also overlies these beds in places. This would account for some of the mafic volcanic rock reported by Muffler (1967) in the interior of Cornwallis Peninsula. Paleontology: The only deposits from the Keku sedimentary strata currently known to contain Triassic fossils are the neptunian dikes from eastern Cornwallis Peninsula (Fig. 5B, site 55). These beds contain abundant gastropods, straight-shelled cephalopods (aulacocerids), and ammonoids in a siliceous mud matrix (Fig. 12). Removal of these fossils is difficult because of silicification and mineralization, though use of concentrated hydrochloric acid can yield molds. Beds of Cornwallis Limestone from USGS Mesozoic locality M1918 that overlie the neptunian dikes (Fig. 5B, site 55) were originally reported as early Carnian on the basis of the ammonoid fauna (Silberling in Muffler, 1967). However, Silberling (2002, personal commun.) informed us that improvements in ammonoid biostratigraphy (Tozer, 1994) actually place the ammonoids Thisbites and Styrites close to the Carnian-Norian boundary. Although Triassic fossils are not known in the lithoclastic beds from the Keku sedimentary strata, lateral tongues of the Cornwallis Limestone provide additional biostratigraphic control. Conodonts from that limestone confirm this age. Cornwallis Limestone overlying the lithoclastic beds on Big Spruce Island (Fig. 5A) contains early Norian conodonts. Overall, these date the upper Keku sedimentary strata as late Carnian and early Norian. On Big Spruce Island, lithoclasts of Carboniferous limestone contain abundant crinoids and bryozoans, and Permian brachiopod bioclasts are reworked into a pebble conglomerate. Interpretation: Overall, the Keku sedimentary strata represent nearshore and/or terrestrial environments. The neptunian dikes contain minimally transported fossils packed together in grain to grain contact, and surrounded and partially infilled by siliceous mud matrix. These packstone beds probably represent a higher-energy environment without significant sediment input, such as in the tidal zone. The neptunian dikes may represent the local base of the Cornwallis Limestone and if they are correlative in age and originally made of carbonate, they should be included within the Cornwallis Limestone. The abundant fluvial sedimentary structures in the lithoclastic beds indicate terrestrial and/or marine fluvial deposition. The basal conglomerate is thin and displays more bedding than the Burnt Island Conglomerate. Muffler (1967) originally included interbeds of oolitic limestone in the Keku Volcanics on the basis of the interpretation that the Keku Volcanics stratigraphically underlies the Cornwallis Limestone. Removal of the Keku
Figure 8. Representative stratigraphic section across the south side of Big Spruce Island with fossil determined ages. USGS Mesozoic locality M2135 (Fig. 5A, site 56) occurs in the limestone.
Figure 9. Sedimentary scour-and-fill structures in the lithoclastic succession of Keku sedimentary strata (Fig. 8) on southwestern Big Spruce Island (Fig. 5A). Hammer (33 cm length) for scale.
194
Katvala and Stanley
Figure 10. Large Carboniferous limestone clast in sandstone beds of the Keku sedimentary strata on southwestern Big Spruce Island (Fig. 5A). Hammer (33 cm length) for scale.
Figure 12. Cut slab of a neptunian dike with fossils (UMIP 303376), including many aulacocerids (Fig. 5B, site 55). Note the light-colored, geopetal, crystalline infill in the fossils. Scale is in centimeters.
Figure 11. Neptunian dikes (outlined) of the Keku sedimentary strata that underlie Cornwallis Limestone of USGS Mesozoic locality M1918 on eastern Kuiu Island (Fig. 5B, site 55). Meter stick (10 cm subdivisions) and hammer (33 cm length) for scale.
Volcanics from the Hyd Group and interpretation of the Keku sedimentary strata as lateral facies of the Cornwallis Limestone (Fig. 3) make it prudent to include the limestone outcrops entirely within the Cornwallis Limestone. These lithoclastic units represent the non-calcareous equivalent of the lithoclastic limestone beds in the Cornwallis Limestone. Cornwallis Limestone Occurrence: The type locality of the Cornwallis Limestone is the 2–3-km stretch of outcrop on the northeasternmost shore of Cornwallis Peninsula on Kuiu Island (Muffler, 1967). Overall,
Figure 13. Lithoclastic limestone with pebbly and sandy layers from the Cornwallis Limestone on the eastern side of Kuiu Island (Fig. 5B, site 68). Hammer (33 cm length) for scale.
the unit occurs on the Cornwallis Peninsula and on some of the adjacent Keku Islets (Figs. 5A, 5B, and 5D). Description: Muffler (1967) defined the Cornwallis Limestone as a medium- to very thick- bedded, characteristically oolitic limestone. This paper expands the definition to include
Conodont biostratigraphy and facies correlations in a Late Triassic island arc, Keku Strait
Figure 14. Photograph of a thin section of Cornwallis Limestone from eastern Kuiu Island (Fig. 5B, site 72) in cross-polarized light. This sandy carbonate contains ooids (O), and quartz (Q), chert (Ch), detrital calcite (C), and carbonate (Ca) grains. Picture width is 7.94 mm.
Figure 15. Sandy to pebbly lithoclastic limestone with cross-beds overlain by horizontal bedding. The horizontal beds contain a rounded boulder of Pybus Limestone. Float block of Cornwallis Limestone from eastern Kuiu Island (Fig. 5B, site 72). Meter stick (10 cm subdivisions) for scale.
both more massive, fossiliferous limestone and well-bedded, less fossiliferous, lithoclastic limestone, both characteristically oolitic (Fig. 13). Variations in fossil type and abundance are apparent within the fossiliferous limestone and probably represent variation in biofacies or depositional environment. Muffler (1967) also reported wispy, locally occurring interbeds of darker, aphanitic limestone as a minor part of the unit, though these are interpreted as diagenetic in origin. The lithoclastic limestone generally has thinner, betterdefined bedding, and contains a variety of clasts from eroded Paleozoic units. Ooids are also common in these lithoclastic units
195
Figure 16. Flounder Cove succession of Cornwallis Limestone with coarser-grained limestone in cliff unconformably overlying finergrained, bedded limestone (Fig. 5B, site 99). Large, loose boulder in center-left of picture is ~1 m high.
(Fig. 14), so it is easy to recognize them as Cornwallis Limestone. Pybus Formation clasts are the most common, though Devonian limestone, Carboniferous limestone, and unidentified chert clasts have been distinguished. The lithoclastic beds display sorting and occur as sandy, pebbly, and cobbly limestone beds. These lithoclastic units are more common near the base of the unit (Muffler, 1967), though in some places they dominate the section. In the area of the prominent point west of Hound Island on eastern Kuiu Island (Fig. 5B, sites 68 and 71–73), the limestone is particularly lithoclastic. Shallow-water fossils are rare to absent here, and much of the rock is calcareous oolitic sandstone (Fig. 14). Clastic grains comprise a variety of mineral and rock types, including quartz, feldspar, biotite, zircon, limestone fragments, volcanic rock fragments, and metamorphic rock fragments. A loose block of sandy to pebbly limestone displays large-scale cross-beds that underlie horizontal layers containing a round boulder of Pybus Formation limestone (Fig. 15). To the south of this area, at the Flounder Cove locality (Fig. 5B, site 99), the limestone beds are finer-grained, and are adjacent to rhyolite of the Keku Volcanics. Upsection, these beds are unconformably overlain by coarser-grained limestone (Figs. 16 and 17) before being capped by basalt of the Hound Island Volcanics. The finer-grained beds do not contain as many shallow-water fossils, but extremely fossiliferous, coarsergrained beds are present. These concentrated fossil beds are up to several decimeters thick, scour into the underlying carbonate mud, have randomly oriented fossils, and have upward-fining layers of carbonate grains and bioclasts. These fossil beds probably represent concentrations of reworked material during higher-energy erosional events. An oncoidal bed of undetermined thickness (Fig. 18) characterizes the base of the section at the Flounder Cove locality.
196
Katvala and Stanley
Figure 17. Representative stratigraphic section at the Flounder Cove locality with fossil determined ages. This section includes USGS Mesozoic Localities M1910 and M1911.
Figure 18. Cut slab of oncoidal limestone with chert lithoclasts (UMIP 303377) from Cornwallis Limestone at the base of the Flounder Cove succession (Fig. 5B, site 99). Scale is in centimeters.
The Cornwallis Limestone directly overlies either the Paleozoic units via an erosional unconformity, or the Keku sedimentary strata. The Cornwallis Limestone also intertongues into the
Keku sedimentary strata. Volcanic rocks of the Hound Island Volcanics overlie the Cornwallis Limestone. Paleontology: The Cornwallis Limestone is typified by abundant shallow-water fossils, including corals, sponges, spongiomorphs, gastropods, large oysters, brachiopods, Stromatomorpha, echinoid fragments, nautiloids, branching algae (Fig. 19), and algal laminations. The unit also contains ammonoids, aulacocerids, other bivalves, and uncommon vertebrate and plate material; boring and soft-sediment trace fossils occur at the Flounder Cove locality. Many of these fossils, including the corals, spongiomorphs, and gastropods, are under study by George Stanley and Andrew Caruthers of the University of Montana and by Robert Blodgett of Anchorage, Alaska. Halobiid bivalves occur throughout the Flounder Cove locality and are under study by Chris McRoberts of the State University of New York (SUNY) at Cortland. One large, toppled spongiomorph colony from Big Spruce Island (Fig. 5A, site 56) was at least 1 m wide by 2 m high in life. Also occurring in the unit is Stromatomorpha californica, a large laminar fossil common to the terranes of western North America. Smith (1927) classified this fossil as a hydrozoan, but it is probably a stromatoporoid sponge. In the area on eastern Kuiu Island area with sandier lithoclastic carbonates, carbonized plant remains are actually common (Fig. 5B, sites 71–73). Most of the plant fossils display no diagnostic features, but
Conodont biostratigraphy and facies correlations in a Late Triassic island arc, Keku Strait
197
Figure 19. Representative macrofossils from the Cornwallis Limestone and one from the Hound Island Volcanics. (1) Colonial coral, Kompsasteria oligocystis; UMIP 228229, Loc 0099. (2) Colonial coral Crassistella juvavica; UMIP 228216, Loc 0099. (3) Branching alga; UMIP 302961, Loc 0074. (4) Partial test of cidarid echinoid; UMIP 228300, Loc 0056. (5) Two brachiopod specimens, Spondylospira sp.; UMIP 303089, Loc 0099. (6) Nautiloid; UMIP 228281, Loc 0099. (7) Gastropod, Spinidelphinulopsis whaleni; UMIP 228247, Loc 0099. (8) Two hydrozoan specimens, Heterastridium conglobatum; UMIP 302043 (top), 228365 (bottom), Loc 0087. (9) Sphinctozoid sponge Amblysiphonella sp.; UMIP 302765, Loc 0056. (10) Indeterminate oyster; field photo (not collected), Loc 0056. (11) Stromatoporoid?, Stromatomorpha californica; UMIP 302763, Loc 0056. (12) Branching hydrozoan, Spongiomorpha ramosa; UMIP 302764, Loc 0056. Scale bars represent 1 cm.
some are preserved well enough to see detail of leaves and branches. Well-preserved plant fossils do not occur in the purer carbonates; however, carbonized wood fragments are common there. Conodonts, a scleractinian coral, and a brachiopod confirm a Late Triassic age for these plant fossils. Acid-etched carbonate blocks from the Flounder Cove locality (Fig. 5B, site 99) yielded small bones of uncertain taxonomic affinity. A long rib-like bone was also found in calcareous sandstone on the eastern side of Kuiu Island (Fig. 5B, site 72). Microfossil resi-
dues from the Cornwallis Limestone include conodonts, bony fish teeth, bony fish scales, fish bones, shark teeth, shark dermal denticles, bivalves, gastropods, ammonoids, ostracodes, foraminifers, holothurian sclerites, sponge spicules, and echinoid spines (Fig. 20). Conodonts reveal most of the Cornwallis Limestone to be of earliest Norian to late early Norian age, though outcrops at the base of the Flounder Cove locality are of latest Carnian age, passing into early Norian upsection (Fig. 17).
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Figure 20. Representative microfossils from the Cornwallis Limestone. Magnification is at ×65 unless otherwise indicated. Photos 1–5: Foraminifers. (1) UMIP 303230, Loc 0099; (2) UMIP 303192, Loc 0056; (3) UMIP 303231, Loc 0099; (4) UMIP 303247, Loc 0099; (5) UMIP 303239, Loc 0099. Photos 6–10: Radiolarians. (6) UMIP 303228, Loc 0099; (7) ×130, UMIP 303244, Loc 0099; (8) ×130, UMIP 303229; Loc 0099; (9) ×30, UMIP 303252, Loc 0099; (10) UMIP 303234, Loc 0099. (11) Framboid; UMIP 303208, Loc 0070. (12) Ostracode; ×43, UMIP 303242, Loc 0099. (13) Holothurian sclerite; UMIP 303232, Loc 0099. (14) Sponge spicule; UMIP 303187, Loc 0056. (15) Echinoid spine; UMIP 303237, Loc 0099. Photos 16–26: Various teeth. (16) UMIP 303217, Loc 0074; (17) UMIP 303213, Loc 0070; (18) UMIP 303198, Loc 0056; (19) UMIP 303215, Loc 0070; (20) UMIP 303219, Loc 0074; (21) UMIP 303203, Loc 0069; (22) UMIP 303201, Loc 0069; (23) UMIP 303202, Loc 0069; (24) ×26, UMIP 303207, Loc 0070; (25) ×43, UMIP 303222, Loc 0070; (26) UMIP 303204, Loc 0069. Photos 27–33: fish scales and dermal denticles. (27) ×26, UMIP 303211, Loc 0070; (28) UMIP 303205, Loc 0069; (29) UMIP 303199, Loc 0069; (30) UMIP 303194, Loc 0056; (31) UMIP 303200, Loc 0069; (32) UMIP 303218, Loc 0074; (33) ×43, UMIP 303206, Loc 0070.
Interpretation: The abundant shallow-water fossils (Silberling in Muffler, 1967), mildly reworked plant fossils, ooids, and oncoids strongly indicate a shallow-marine environment for
the Cornwallis Limestone. Lithoclastic and volcaniclastic clasts reveal sediment input from terrestrial erosion. Sedimentary structures such as variously cross-laminated beds and erosional sur-
Conodont biostratigraphy and facies correlations in a Late Triassic island arc, Keku Strait
199
However, the relationship between Triassic units on either side of the strait is unclear, as the amount of post-Triassic tectonic compression or lateral offset is unknown. The detached relationship of these unnamed limestone units to the defined Cornwallis Limestone prompts tentative classification of these outcrops as a separate unit. These rocks represent shallow-marine deposition separate from the rocks on Cornwallis Peninsula, but they may be assigned to the Cornwallis Limestone after further study.
Figure 21. Beds of lithoclastic, cross-bedded, unnamed shallow-water limestone from Cape Bendel (Fig. 5E, site 83). Hammer (33 cm length) for scale.
faces in these lithoclastic units indicate higher current velocity and probable proximity to shore. Higher clastic input may have excluded many marine organisms from some deposits. Interbeds of finer-grained limestone may represent facies transitions into lagoonal or deeper-water environments. The Flounder Cove locality may exemplify this as a lagoonal or inner slope deposit. A combination of boring and soft-sediment trace fossils from the Flounder Cove locality indicates intermittent development of hard grounds in the carbonate. An erosional unconformity within the Flounder Cove section resulted in coarser-grained, mildly lithoclastic, massive limestone overlying the muddier beds. Overall, this unit represents shallow-marine environments with good carbonate production. It is distal to the Keku sedimentary strata, and coeval with the deeper-water Hamilton Island Limestone. Unnamed Shallow-Water Limestone Lithoclastic limestone similar to that of the Cornwallis Limestone occurs in two places on the east side of Keku Strait. In the Cape Bendel region (Fig. 5E, site 83), lithoclastic limestone underlies units (Fig. 21) dated by using conodonts as late Carnian. This limestone contains chert pebbles of the Pybus Formation as well as other unidentified clasts. Microfossil residues from Cape Bendel include bony fish teeth, bony fish scales, and shark dermal denticles. The limestone deposits in this area (Fig. 5E, sites 83 and 84) are entirely fault-bounded, so their association with the Hound Island Volcanics is suspect. We remove these beds from the Hound Island Volcanics because volcaniclastic material is absent. In Portage Pass (Fig. 5F, site 59), lithoclastic limestone with Cannery Formation and Pybus Formation clasts contains a suspected Late Triassic oyster. This limestone was previously mapped as Burnt Island Conglomerate (Muffler, 1967), to which it is similar. Both of these lithoclastic limestone outcrops occur on the opposite side of the strait from the Cornwallis Limestone.
Hamilton Island Limestone Occurrence: The type locality of the Hamilton Island Limestone is the northern tip of Hamilton Island (Muffler, 1967). Elsewhere, it crops out west of Kupreanof Island in Hamilton Bay and on Hamilton Island, in the Keku Islets, and in the Cape Bendel region (Figs. 5C–5G). Only two isolated outcrops represent the Hamilton Island Limestone in the Cape Bendel region (Fig. 5E, sites 84 and 91). Description: Very thin-bedded, fine-grained limestone with black argillaceous laminae is typical of the Hamilton Island Limestone (Muffler, 1967). Debris flows and thin to medium beds of calcareous sandstone occur locally. The debris flows are beds with matrix-supported and poorly sorted sand- to cobble-sized clasts; the larger clasts show no preferred orientation. These debris flows contain clasts of Hamilton Island Limestone, the Cannery and Pybus formations, and reworked shallow-water Late Triassic fossils. These debris flows were originally included by Muffler (1967) in the Burnt Island Conglomerate. However, because they are interbeds within the rocks ascribed to the Hamilton Island Limestone and they contain reworked clasts of the Hamilton Island Limestone, they are included within this formation. Strata at the northern end of Hamilton Island (Fig. 5F, site 60; Fig. 22) and elsewhere are interpreted as turbidites as they are well-bedded, often fine upward, contain fossil concentrations, and display low-angle cross-bedding. Soft-sediment deformation interpreted as slump folding (Fig. 23) and deposition of most of the succession in a single conodont zone (Fig. 22; Table 3) suggest rapid deposition. Debris flows containing Paleozoic clasts and Triassic limestone clasts occur in the lower part of the succession and are superficially similar to the Burnt Island Conglomerate (Fig. 5F, site 57; Figs. 22 and 24). Upsection, these conglomeratic beds contain fewer Paleozoic clasts and more Triassic clasts. Many of the Triassic clasts appear to be locally reworked from the Hamilton Island Limestone. These conglomeratic deposits become finer-grained upsection, and calcareous sandstone represents them upsection. The lower contact of this unit is the first limestone bed above the Burnt Island Conglomerate. Although it was not observed, the lower contact may onlap the erosional unconformity truncating Paleozoic units. For example, on the northeast side of the central part of Hamilton Island and on the northeast side of Hamilton Bay, the Hamilton Island Limestone appears to rest disconformably on the Pybus Formation (Muffler, 1967). However, exposure of the contact is not adequate to preclude the presence of a thin unit of
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Katvala and Stanley
Figure 22. Representative stratigraphic section at the northeast shore of Hamilton Island with fossil-determined ages. This section includes USGS Mesozoic localities M1882–84.
Burnt Island Conglomerate (Muffler, 1967). Basalt and volcaniclastics of the Hound Island Volcanics conformably and sometimes transitionally overlie the Hamilton Island Limestone (Muffler, 1967). We speculate that the Hamilton Island Limestone is a deeper-water facies of the Cornwallis Limestone, but lateral transitions confirming such a relationship are absent in the field area.
Paleontology: The bivalve Halobia characteristically dominates beds of the Hamilton Island Limestone (Fig. 25). Bivalves from this unit are under study by Chris McRoberts of SUNY at Cortland. Ammonoid, halobiid, and conodont faunas indicate that this unit is predominantly late Carnian, though the uppermost beds are of early Norian age. Fossiliferous clasts in the
TABLE 3. BIOSTRATIGRAPHICALLY SIGNIFICANT FOSSILS RECOVERED FROM THE KEKU STRAIT AREA Sample Biozone Type Fossil names
Fm
Late Carnian 60
HMNE-C1
"upper" nodosus
C
HMNE-C3
upper nodosus
C
60
HMNE-C6
upper nodosus
C
Metapolygnathus nodosus, M. polygnathiformis, M. carpathicus Metapolygnathus nodosus, M. polygnathiformis, M. carpathicus, M. sp. aff. M. zoae Metapolygnathus nodosus, M. polygnathiformis
HL
60
60
HMNE 34.4
–
A
Tropitid
HL
60
HMNE-C8
Dilleri
A
Shastitas sp., Hannoceras sp., Shastitas? sp.
HL
60
HMNE-C8
Dilleri-Welleri
A
Paratropites?
HL
60
HMNE-C9F2
Welleri
A
Discotropites? sp.
HL
60
HMNE-C10
upper nodosus
C
Metapolygnathus polygnathiformis
HL
60
HMNE-C13
upper nodosus
C
Metapolygnathus carpathicus
HL
60
HMNE-C15
upper nodosus
C
Metapolygnathus nodosus, M. sp. aff. M. zoae
HL
60
HMNE-C16
upper nodosus
C
Metapolygnathus nodosus, M. carpathicus
HL
60
HMNE-C17
upper nodosus
C
Metapolygnathus nodosus, M. polygnathiformis, M. carpathicus
HL
60
HMNE 69.1
Welleri
A
Hannoceras sp.
HL
61
M1924-F
–
A
Tropitid
HL
61
M1924-F
–
B
62
PISW-C2L2
middle nodosus
C
62
PISW-C3L3
nodosus
C
Halobia sp. cf. H. superba, H. sp. cf. H. ornatissima Metapolygnathus nodosus, M. polygnathiformis, M. sp. cf. M. reversus Metapolygnathus polygnathiformis?, M. sp., Neogondolella? sp.
62
PISW-C7
nodosus
C
Metapolygnathus nodosus
HL
66
PP-C2F2
nodosus
C
Metapolygnathus polygnathiformis, M. carpathicus
HL
66
M1932 (photo)
–
B
Halobia ornatissima (silicified)
HL
67
HMSE-C1
uppermost nodosus
C
Metapolygnathus sp. aff. M. nodosus
HL
67
HMSE-C2
uppermost nodosus
C
Metapolygnathus sp. aff. M. nodosus
HL
67
HMSE A5.6-5.7
–
B
Halobia sp. cf. H. superba
HL
67
HMSE A10.75m
Welleri
A
Discotropites? sp.
HL un
Loc. no.
HL HL
HL HL HL
83
CB2-C2
uppermost nodosus
C
Metapolygnathus sp. aff. M. nodosus, M. polygnathiformis
83
CB2-C3
uppermost nodosus
C
Metapolygnathus sp. aff. M. nodosus, M. polygnathiformis
un
84
CB2-C1A
lower nodosus
C
HL
84
CB2-C2A
lower nodosus
C
Metapolygnathus sp. cf. M. nodosus Metapolygnathus polygnathiformis, M. sp. cf. M. reversus, M. reversus?
97
SC-C4
upper nodosus
C
97
SC-F4
–
A,B
99
FL-C2
nodosus
99
FL-C3
nodosus-primitius
99
FC2-C1
99
FC2-C2
Late Carnian or early Norian 55 SM-C1
HL HL
Metapolygnathus nodosus, M. polygnathiformis Tropitids, Halobia ornatissima?
HL
C
Metapolygnathus sp. cf. M. nodosus
C
C
Metapolygnathus sp. aff. M. primitius
C
nodosus-primitius
C
C
nodosus-primitius
C
Metapolygnathus sp. aff. M. primitius Metapolygnathus nodosus, M. sp. aff. M. zoae, M. sp. aff. M. primitius
C
C
nodosus-primitius
C
Metapolygnathus sp. cf. M. primitius
56
BS-C2
–
C
Metapolygnathid or Epigondolellid
C
64
PISW-F3
–
B
Halobia radiata, H. sp. cf. H. austriaca
HL
65
TCF-F
–
B
CP-C2
nodosus-primitius
C
Halobia radiata Metapolygnathus sp. cf. M. primitius, Neogondolella sp.
HL
70 70
CPE-F2
primitius
C
Metapolygnathus primitius, M. sp. aff. M. primitius
C
70
CPE-C3
primitius
C
Metapolygnathus primitius
C
74
SWK 1
primitius
C
Metapolygnathus primitius
C
95
CPE-C4
primitius?
C
Metapolygnathus primitius?
C
96
CPE-C5
primitius
C
Metapolygnathus primitius
C
C Continued
202
Loc. no.
Katvala and Stanley TABLE 3. BIOSTRATIGRAPHICALLY SIGNIFICANT FOSSILS RECOVERED FROM THE KEKU STRAIT AREA (continued ) Sample Biozone Type Fossil Names
Early Norian 56
Fm
BS-C5
quadrata
C
Epigondolella sp. cf. E. quadrata
C
66
PP-C1F1
–
B
Halobia beyrichi, H. sp. cf. H. cordillerana, H. sp. cf. H. lineata
HL
66
PP-C2F2
–
B
Halobia sp. cf. H. beyrichi, H. cordillerana
HL
68
FU-C1
quadrata?
C
Epigondolella sp. cf. E. quadrata, Metapolygnathus sp.
C
69
CPE-C2
triangularis
C
C
70
CP-C1
primitius-quadrata
C
72
KUW-F3A
Kerri
A
Epigondolella triangularis uniformis, E. quadrata, Metapolygnathus primitius Epigondolella sp. cf. E. quadrata, Metapolygnathus sp. aff. M. primitius Guembelites clavatus
73
K(not)-C
–
C
Epigondolella sp.
C
74
SWK 2
primitius-quadrata
C
Epigondolella sp. cf. E. quadrata, Metapolygnathus primitius
C
C C
74
SWK 2
primitius-quadrata
C
Epigondolella sp. cf. E. quadrata, Metapolygnathus primitius
79
HIW-C2
triangularis
C
HV
79
HIW-F1
–
B
Epigondolella triangularis uniformis, E. triangularis triangularis, E. quadrata Halobia sp. cf. H. beyrichi, H. sp. cf. H. lineata
C
80
HIW-F2
–
B
Halobia beyrichi?
HV
81
HIW-C5
triangularis
C
Epigondolella sp. cf. E. triangularis
HV
82
HIN-C2
triangularis?
C
HV
85
HMSW-C1F1
–
B
Epigondolella triangularis?, E? sp., Metapolygnathus? sp., Neogondolella? sp. Halobia sp. cf. H. beyrichi, H. sp. cf. H. fallax
97
SC-F1
–
B
Halobia cordillerana, H. lineata
HL
99
FL-C6
primitius
C
Metapolygnathus primitius
C
99
FL-C11
primitius
C
Metapolygnathus sp. aff. M. primitius
C
99
FL-C12
primitius
C
Metapolygnathus sp. aff. M. primitius
C
99
FL-C16
triangularis
C
C
99
FL-C17
triangularis
C
99
FL
Kerri
A
Epigondolella triangularis uniformis, E. triangularis triangularis, E. quadrata, rounded conodont fragments Epigondolella triangularis uniformis, E. triangularis triangularis, E. quadrata, Misikella longidentata Stikinoceras kerri, Greisbachites? sp.
99
FC2-F3
Kerri
A,B
Stikinoceras kerri, Halobia beyrichi
C
99
FC2-F7
Kerri
B
Halobia cordillerana
C
99
FC2-F6
Kerri
A
Stikinoceras kerri
C
Middle Norian 46
BT-C2
spiculata
C
Epigondolella triangularis triangularis, E. spiculata
HV
86
HIE-C2
–
C
Neogondolella sp. cf. N. steinbergensis, N. sp.
HV
86
HIE-C3
postera
C
Epigondolella sp. cf. E. postera, Neogondolella sp.
HV
86
HL-C1
–
C
Neogondolella sp.
HV
86
HL-C2
–
C
Epigondolella? sp.
HV
86
HL-C6
spiculata
C
Epigondolella spiculata, Neogondolella sp.
HV
86
HIE-F1
–
B
Halobia fallax
HV
86
HIE-F4
–
B
Halobia fallax
HV
87
GH-C1
spiculata
C
HV
87
GH-C2
spiculata
C
Epigondolella spiculata, E. transitia, E. matthewi, E. sp. aff. E. triangularis, E. triangularis triangularis Epigondolella spiculata
HV
HV
C C
HV Continued
debris-flow conglomerates within the northern Hamilton Island succession are of late Carnian age. These indicate the erosion of upper Carnian limestone in the late Carnian. In the Hamilton Island Limestone on the southern shore of Hamilton Island, shallow-water corals rarely occur in debris-flow deposits, indicating a proximal source for the debris flows. Also, carbonized wood fragments occur locally in the Hamilton Island Limestone, and a
silicified piece of wood was found on the northern end of Hamilton Island (Fig. 26), further indicating reworking from proximal environments. A locality in Portage Pass preserved well-silicified Halobia (Fig. 27); silicification of this thin bivalve is rare. Microfossil residues in the Hamilton Island Limestone include conodonts, bony fish teeth, shark dermal denticles, foraminifers, radiolarians, sponge spicules, a pyritized scolecodont tooth,
Conodont biostratigraphy and facies correlations in a Late Triassic island arc, Keku Strait
Loc. no.
TABLE 3. BIOSTRATIGRAPHICALLY SIGNIFICANT FOSSILS RECOVERED FROM THE KEKU STRAIT AREA (continued ) Sample Biozone Type Fossil Names
203
Fm
Late Norian 87
GH-C6
bidentata
C
Epigondolella tozeri, E. sp. aff. E. mosheri, E. bidentata
HV
87
=GH-C6
bidentata
C
Epigondolella tozeri, E. englandi, E. sp. aff. E. mosheri, E. bidentata
HV
87
GBC
Cordilleranus
H
Heterastridium conglobatum
HV
87
GHBC-1
Cordilleranus
H
Heterastridium conglobatum
HV
87
GHBC-1
Columbianus III
A
Hellerites sp. or Parajuvavites sp.
HV HL
Reworked Paleozoic Conodonts 62 PISW-C1L1
Early Permian
C
Mesogondolella sp.
68
FU-C1
Devonian
C
Polygnathus linguiformis, Polygnathus. sp.
C
68
FU-C2
middle Devonian
C
C
97
SC-C2
Early Permian
C
Belodella triangularis, Panderodus sp., Palmatolepis sp. (juvenile), Polygnathus sp., broken specimens Mesogeondolella sp., Sweetognathus sp.
HL
Note: Loc. no.—University of Montana Museum of Paleontology locality. Biozone is the applicable ammonoid or conodont biozone as presented in Fig. 45. Type is the fossil used to determine the biozone, with A—ammonoid, B—bivalve, C—conodont, and H—Heterastridium. Formations (Fm.): C—Cornwallis Limestone; HL—Hamilton Island Limestone; HV—Hound Island Volcanics; un—unnamed shallow-water limestone.
Figure 23. Slump fold from the Hamilton Island Limestone on northeast Hamilton Island (Fig. 5F, site 60). The white “S”, a piece of plastic beach litter, is ~10 cm tall.
and various tubular fossils (Fig. 28). Finally, clasts of the Pybus Formation in the debris-flow conglomerates on northern Hamilton Island have Permian crinoids, bryozoans, and brachiopods. Interpretation: The dominance of Halobia, the presence of ammonoids, and the lack of other kinds of macrofossils are suggestive of a deep-water environment (Silberling in Muffler, 1967). The turbiditic nature and uniform age of the beds overlying the Burnt Island Conglomerate suggest that rapid infill of the basin continued after deposition of the Burnt Island Conglomerate and probably slowed into the Norian. Upper Carnian rock fitting this description crops out in the Cape Bendel region and was formerly included in the Hound Island Volcanics (Fig. 5E, site 84). Similar, undated beds occur inland of Cape Bendel near outcrops of Burnt Island Conglomerate (Fig. 5E, site 91). The absence of volcaniclastic material in these beds,
Figure 24. Debris flow conglomerate with Permian and Triassic clasts from the Hamilton Island Limestone on northeast Hamilton Island (Fig. 5F, site 60). Hammer (33 cm length) for scale.
and their fault-bounded nature, suggests they are separate from the Hound Island Volcanics and equivalent to the Hamilton Island Limestone. Hound Island Volcanics Occurrence: The type locality of the Hound Island Volcanics is the shores of Hound Island (Muffler, 1967). The unit also occurs on Kuiu Island, the Keku Islets, around Hamilton Bay, on Hamilton Island, in the Cape Bendel region, and on Turnabout Island (Figs. 5B–5G). Description: The Hound Island Volcanics consists mainly of basaltic pillow lava, basaltic pillow breccia, massive basalt, andesitic volcanic breccia, and hyaloclastic tuff with subordinate tuffaceous polymict conglomerate, limestone, and volcaniclastic
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Figure 25. Bedding surface covered by Halobia (UMIP 303065) from the Hamilton Island Limestone at the Squawking Crow locality (Fig. 5C, site 97). Scale is in centimeters.
Figure 27. Silicified Halobia ornatissima (uncollected) from the Hamilton Island Limestone in Portage Pass (Fig. 5F, site 66). Scale is in centimeters.
Figure 26. Silicified wood fragment (UMIP 303378) from the Hamilton Island Limestone on northeast Hamilton Island (Fig. 5F, site 60). Scale is in centimeters.
sandstone (Muffler, 1967). Nomenclature for many of the volcanic deposits described by Muffler (1967) followed Carlisle (1963). Undevitrified, basaltic volcanic glass occurs locally in hyaloclastic tuff, preserved by rapid cooling underwater (Brew and Muffler, 1965; Muffler et al., 1969). The andesitic breccia and polymict conglomerate commonly form thick beds that are conspicuous in the field. The matrix of the breccia is commonly calcareous (Muffler, 1967); breccia fragments are mostly Triassic limestone and volcanic rock, but sometimes are derived from Paleozoic units. The breccia beds resemble huge debris flows, as they are matrix-supported, poorly sorted units that have a wide range of grain sizes (Fig. 29). The largest clasts show no preferred orientation. Pumice fragments occur rarely in the Hound Island Volcanics (Muffler, 1967), and Buddington and Chapin (1929) reported volcanic bombs in places. Additionally, a pahoehoe basalt flow is in the Cape Bendel region (Fig. 30). The limestone is fine-grained and generally occurs as interbeds
or sometimes as a sediment drape over basalt pillows (Fig. 31). These interbeds are often lenticular (Fig. 32) or occur as stringers of limestone (Muffler, 1967). Their lithology and appearance are very similar to those of the Hamilton Island Limestone, but their association with volcanic rock leads to larger amounts of finegrained siliciclastic material. A thicker succession of this limestone crops out on the eastern shore of Hound Island (Fig. 5D, site 86; Fig. 33). Limestone beds lower in the succession overlie an andesitic debris flow (Fig. 33). Within this succession are numerous lag deposits containing sulfides, volcaniclastic debris, and abundant bones. Another limestone succession, though structurally isolated, crops out in the Gil Harbor mudflat (Fig. 5B, site 87). The older beds in this area are fine-grained limestone typical of the Hound Island Volcanics, whereas the younger beds are thicker, fossiliferous packstone with abundant shallow-water fossils. Additional shallow-water limestone crops out along the coast south of Gil Harbor and in Kadake Bay (Fig. 5B). The base of the Hound Island Volcanics is marked by the first occurrence of major basaltic and andesitic volcanism or associated volcaniclastic deposits. The pre–Upper Jurassic unconformity overlies this unit. Paleontology: Within the variety of facies in the Hound Island Volcanics, many different fossils types occur. Buddington and Chapin (1929) reported carbonized plant fossils in sandstones of this unit on Turnabout Island, north of Cape Bendel. Trace fossils occur in some of the volcaniclastic sandstone (Fig. 34). The presence of a starfish resting trace (Fig. 34) indicates a marine environment for these volcaniclastic beds. Some of the vertebrate fossils in the limestone succession on the eastern shore of Hound Island are ichthyosaur bone. Neogondolellid conodonts are particularly abundant in the eastern Hound Island succession
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Figure 28. Representative microfossils from the Hamilton Island Limestone. Magnification is at ×65 unless otherwise indicated. Photos 1–2: Foraminifers. (1) UMIP 303265, Loc 0060; (2) UMIP 303276, Loc 0060. (3) Sponge spicule. UMIP 303272, Loc 0060. Photos 4–6: Radiolarians at ×130. (4) UMIP 303268, Loc 0060; (5) UMIP 303269, Loc 0060; (6) UMIP 303273; Loc 0060. (7) Scolecodont. UMIP 303279, Loc 0060. Photos 8, 13–14: Various teeth. (8) ×43, UMIP 303261, Loc 0067; (13) UMIP 303260, Loc 0067; (14) UMIP 303263, Loc 0067. Photos 9–12: Fish scales and dermal denticles. (9) UMIP 303280, Loc 0060; (10) UMIP 303281, Loc 0060; (11) ×43, UMIP 303262, Loc 0067; (12) UMIP 303274, Loc 0060.
Figure 29. Debris flow breccia from the Hound Island Volcanics on western Hamilton Island (Fig. 5F, north of site 77). Meter stick (10 cm subdivisions) for scale.
as well. Throughout the area, halobiid bivalves dominate the finegrained limestone beds, as in the Hamilton Island Limestone. Monotis also occurs in a locality in southeastern Hamilton Bay (Fig. 5G, site M1898), and in deposits at the Gil Harbor mudflat (Muffler, 1967) (Fig. 5B, site 87). These younger deposits in Gil Harbor contain abundant silicified fossils including corals, many
Figure 30. Pahoehoe from the Hound Island Volcanics in the Cape Bendel region. Hammer (33 cm length) for scale.
different bivalves, gastropods, ammonoids, aulacocerids, Heterastridium (Fig. 19), echinoid spines, and trace-fossil tubes (Fig. 35), as well as carbonized wood fragments. Each bed from these deposits contains a distinct combination of fossil taxa,
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Figure 31. Limestone (outlined) draping basalt pillows from the Hound Island Volcanics. Outcrop is in Hamilton Bay, along the shoreline northeast of Little Hamilton Island (Fig. 5G). Hammer (33 cm length) for scale.
Figure 33. Succession of gently dipping, fine-grained, Halobia-rich limestone overlying andesitic conglomerate (foreground). The outcrop is in the Hound Island Volcanics on the east coast of Hound Island (Fig. 5D, site 86). Each bucket is ~40 cm tall.
Figure 32. Pinching out of lensoid limestone bed (behind hammer) in Hound Island Volcanics from the western coast of Hound Island (Fig. 5D, site 81). The succession includes, from bottom to top, basalt, volcaniclastics, limestone, volcaniclastics, and basalt. Hammer (33 cm length) for scale.
probably representing different biofacies. Gastropods from Gil Harbor (Fig. 5B, site 87) include the genus Chulitnacula, which is endemic to the Chulitna, Farewell, and Alexander terranes (Frýda and Blodgett, 2001). Two predominant trace fossils from the Gil Harbor carbonates are sinuous burrows that wind around the macrofossils in three dimensions and variably oriented, tubular borings through the shells in the deposits. A few shallow-water limestone samples from south of Gil Harbor have echinoid spines (Fig. 5B, site 46). Overall, shallow-water microfossil residues yielded conodonts, bony fish teeth, bony fish scales, fish bones, shark dermal denticles, bivalves, gastropods, ostracodes, foraminifers, and various tubular fossils (Fig. 36). Deep-water microfossil residues yielded conodonts, bony fish teeth, bony fish scales, fish bone, bivalves, foraminifers, radiolarians, and various tubular fossils (Fig. 36). Biostratigraphically significant fossils encompass ages from late early Norian through late Norian. The
Figure 34. Trace fossils in volcanic sandstone from the Hound Island Volcanics on western Hound Island (Fig. 5D, site 81). White arrow points to starfish resting trace. Hammer (33 cm length) for scale.
Monotis occurrences represent the only two late Norian localities, and conodonts from Gil Harbor confirm a late Norian age. Interpretation: This unit represents basaltic and andesitic volcanism throughout the area. Fossils associated with basaltic rock and volcaniclastics indicate that volcanism began in the late early Norian. This coincides with the youngest age obtained in carbonate units of the Cornwallis Limestone. Carbonates from the Cornwallis Limestone are not associated with basaltic volcanism but contain late early Norian fossils. Volcanism continued
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Figure 35. Silicified fossils, mostly shallow-water bivalves, from the Hound Island Volcanics in the Gil Harbor mudflat (Fig. 5B, site 87). Scale is in centimeters.
into at least the middle Norian and possibly into the late Norian. Pumice fragments, pahoehoe, volcanic bombs, and erosional debris flows along with pillow lavas suggest volcanism in both subaerial and subaqueous environments. The scarcity of late Norian fossils may suggest post-Triassic erosion. Carbonate units in the Hound Island Volcanics preserve both shallow- and deep-water depositional environments. Packstone beds from Gil Harbor contain abundant silicified shallow-water fossils. These fossils are reworked, but not enough reworking occurred before burial to break up the many large and fragile fossils that manifest nearly complete preservation. Additionally, the representation of many separate biofacies indicates a lack of homogenization through reworking and transport. Trace fossils show no vertical variation within each bed, indicating fairly uniform conditions. Overall, these features in the Gil Harbor packstones and the lack of adjacent beds with finer-grained sediments or turbiditic or storm-related sedimentary structures indicate deposition in a shallow-water environment above normal wave base. On eastern Hound Island, the abundance of neogondolellid conodonts and the fine-grained, well-bedded rocks suggest a deeper-water environment (Mercantel, 1973; Behnken, 1975; Babcock, 1976; Carey, 1984; Carr et al., 1984; Carter and Orchard, 2000). BIOSTRATIGRAPHY Ammonoids, specific bivalves, and conodonts are the most important index fossils in the Upper Triassic, and combining data from these taxonomic groups provides power-
ful biostratigraphic resolution (Fig. 37). Using exposures in northeast British Columbia, Tozer (1967, 1984, 1994) pieced together a complete ammonoid zonation for the Upper Triassic of western Canada. The “flat clams” Halobia and Monotis occur worldwide and are important in Triassic biostratigraphy (McRoberts, 1997; Silberling et al., 1997). Halobia ranged from the Carnian into the middle Norian, where it coexisted with Monotis (Eomonotis) briefly before being succeeded by Monotis (McRoberts, 1997; Silberling et al., 1997). Monotis existed during the late Norian, but went extinct before the end of the Triassic (Silberling et al., 1997). The spherical hydrozoan Heterastridium conglobatum also occurs in the Keku Strait and adds to age determinations. It is abundant in deposits of Cordilleranus Zone age (Fig. 37) in the North American Cordillera (Silberling and Tozer, 1968; Tozer, 1994). Heterastridium may range down into the Columbianus Zone (Fig. 37) of the middle Norian (Tatzreiter, 1975; Krystyn and Wiedmann, 1986; Stanley et al., 1994) and possibly into the Rhaetian (González-León et al., 1996). Krystyn (2003, personal commun.) and Stanley (e.g., Stanley et al., 1994) postulated that within Heterastridium populations, individuals display an increase in their maximum diameter through time, possibly permitting determination of age. Although ammonoids, Halobia, Monotis, and to some extent Heterastridium are valuable and provide good biostratigraphic data, macrofossils can be difficult to locate in tectonostratigraphic terranes. This is due to the original, restricted lateral extent of facies, structural complications, the taphonomic bias against larger fossils, and in part to environmentally restricted taxa. Conodonts occur in a greater variety of facies, are less susceptible
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Figure 36. Representative microfossils from the Hound Island Volcanics. Magnification is at ×65 unless otherwise indicated. Shallow-water carbonate facies. Photos 1–3: Foraminifers. (1) UMIP 303287, Loc 0086; (2) UMIP 303294, Loc 0086; (3) UMIP 303290, Loc 0082. Photos 4–5: Radiolarians at ×130. (4) UMIP 303291, Loc 0082; (5) UMIP 303292, Loc 0082. (6) Sponge spicule. UMIP 303299, Loc 0081. (7) Echinoid spine. ×43, UMIP 303298, Loc 0086. Photos 8–9: Fish dermal denticle and scale. (8) ×43, UMIP 303296, Loc 0086; (9) UMIP 303284, Loc 0086. Photos 10–14: Various teeth. (10) UMIP 303285, Loc 0086; (11) ×43, UMIP 303297, Loc 0086; (12) ×43, UMIP 303288, Loc 0082; (13) ×43, UMIP 303295, Loc 0086; (14) UMIP 303286, Loc 0086. Deep-water carbonate facies. Photos 15–20: Foraminifers. (15) UMIP 303305, Loc 0087; (16) UMIP 303304, Loc 0087; (17) UMIP 303308, Loc 0087; (18) UMIP 303307, Loc 0087; (19) UMIP 303321, Loc 0087; (20) UMIP 303309, Loc 0087. 21: Ostracode. ×43, UMIP 303320, Loc 0087. 22: sponge spicule; UMIP 303311, Loc 0087. Photos 23–27, 34–35: various teeth; (23) ×26, UMIP 303317, Loc 0087; (24) UMIP 303312, Loc 0087; (25) ×43, UMIP 303324, Loc 0087; (26) UMIP 303302, Loc 0046; (27) UMIP 303318, Loc 0087; (34) UMIP 303316, Loc 0087; (35) ×43, UMIP 303315, Loc 0087. 28: Jawbone?. ×26, UMIP 303322, Loc 0087. Photos 29–33: Fish scales and dermal denticles. (29) UMIP 303300, Loc 0046; (30) ×26, UMIP 303323, Loc 0087; (31) UMIP 303314, Loc 0087; (32) ×26, UMIP 303303, Loc 0046; (33) ×43, UMIP 303313, Loc 0087.
to destructive diagenetic processes, and commonly occur in the smallest outcrop or individual clast. Detailed biostratigraphic work with conodonts is still lacking in most terrane localities. A complete conodont biostratigraphy for the Upper Triassic is still
being developed (Kozur, 1980; Krystyn, 1980; Orchard, 1991a, 1991b; Buryi, 1997; Orchard and Tozer, 1997). The zonations presented by Orchard are based on a “highly resolved and intercalibrated conodont-ammonoid zonation” (Carter and Orchard,
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Figure 37. Late Triassic biochronology displaying conodont zones, ammonoid zones, and selected bivalve ranges. Heterastridium conglobatum occurs in the Cordilleranus ammonoid Zone with Monotis subcircularis. Generic abbreviations: M.—Metapolygnathus; E.—Epigondolella; Mi.—Misikella; H.—Halobia. Ammonoid zones after Tozer (1967, 1984, 1994) and Orchard and Tozer (1997). Approximate bivalve ranges after McRoberts (1993, 1997), Silberling et al. (1997), and McRoberts (2003, personal commun.). Conodont zones after Orchard (1991b) and Orchard and Tozer (1997).
2000) from extensive work done in northeast British Columbia (Orchard, 1991b; Orchard and Tozer, 1997; Carter and Orchard, 2000). This zonation is applicable to Triassic rocks in a number of Cordilleran terranes, including the Queen Charlotte Islands of Wrangellia (Orchard, 1991a; Carter and Orchard, 2000). It is also applicable to the biostratigraphic succession in the Keku Strait area. Conodonts were the focus of this study. In addition, L. Krystyn and N.J. Silberling examined and/or identified associated ammonoids and C.A. McRoberts identified halobiid bivalves. Following Orchard (1991b) and Tozer (1994), standard ammonoid zones begin with an upper case letter and are not in italics, whereas conodont zones are in italics as their formal species names. In the sections below, we discuss only biostratigraphically significant fossils, which are listed in Table 3. Reworked Paleozoic fossils occur locally in the Late Triassic deposits of the Keku Strait area. Carboniferous and Permian macrofossils in reworked clasts are common in the Burnt Island Conglomerate, lithoclastic beds of the Keku sedimentary strata, and the Cornwallis Limestone. Devonian clasts also occur in the Cornwallis Limestone. Conodonts of Devonian and Early Permian age were identified in reworked Paleozoic clasts in Upper Triassic rock (Fig. 38; Table 3).
Tozer (1967) picked the base of the Kerri Zone as the Carnian-Norian boundary (Fig. 37), but the precise position of this boundary remains undefined. Previously, the conodont Neogondolella navicula was used as an indicator of the Carnian-Norian boundary. This taxon has since been shown to be facies controlled, and occurrences of the genus in the upper Carnian appear similar to those from the lower Norian (Carter and Orchard, 2000; Orchard, 2003, personal commun.). Furthermore, Metapolygnathus primitius originates in the Macrolobatus Zone and overlaps into the Kerri Zone, so no currently known conodont origination coincides with the base of the Kerri Zone (Orchard, 1983, 1991a, 1991b; Orchard and Tozer, 1997). Radiolarian distributions also do not coincide well with the base of the Kerri Zone (Carter and Orchard, 2000). We use the traditional definition in this paper for convenience (Fig. 37), but it is important to note that the Subcommission on Triassic Stratigraphy is presently striving to define the Carnian-Norian precisely. Middle Triassic Only one fossil indicating a Triassic age earlier than Late Triassic has been reported in the Alexander terrane. The uppermost Anisian (Middle Triassic) conodont Neogondolella acuta
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Figure 38. Late Devonian and Early Permian conodonts reworked into Hyd Group units. All figures are scanning electron micrographs at ×82 magnification except Photo 12. Illustrated specimens are upper surface views of Pa elements unless otherwise indicated. Photos 1–3: Belodella triangularis Stauffer. (1) FU-C2, UMIP 302712; (2) FU-C2, UMIP 302714; (3) FU-C2, UMIP 302718. Photos 4–5: Panderodus sp. (4) FU-C2, UMIP 302717; (5) FU-C2, UMIP 302713. (6) Polygnathus linguiformis (Hinde 1879); FU-C1, UMIP 302724. Photos 7–8: Polygnathus sp. (7) FU-C1, UMIP 302723; (8) FU-C2, UMIP 302719. (9) Pb element; FU-C2, UMIP 302715. Photos 10–11: Juvenile Palmatolepis sp. (10) FU-C2, UMIP 302720; (11) FU-C2, UMIP 302721. (12) Pb element; ×48, FU-C2, UMIP 302716. Photos 13–16: Mesogondolella sp. (13) SC-C2, UMIP 302728; (14) SC-C2, UMIP 302726; (15) SC-C2, UMIP 302727; (16) SC-C2, UMIP 302729. (17) Sweetognathus? sp.; SC-C2, UMIP 302725.
was found on Big Saltery Island, just east of the Keku Strait area, southeast of Kupreanof Island (Wardlaw in Karl et al., 1999). However, because this fossil was in a limestone debris flow, its occurrence may indicate deposition and subsequent erosion of pre–Late Triassic age rock, accounting for the lack of Middle Triassic rock in the field area. Early Carnian Fossils indicating an early Carnian age also are rare. Silberling (in Muffler, 1967) reported the bivalve Halobia sp. cf. H. rugosa and the ammonoid Coroceras sp. cf. C. suessi from a limestone bed at the base of the Burnt Island Conglomerate at USGS Mesozoic locality M1892 (Fig. 5G). Silberling (in Berg, 1981) identified Halobia rugosa (Fig. 37) in fossiliferous metasedimentary strata ~25 km southwest of Petersburg, east of the map area. Wardlaw (1982) described primitive forms of
the early Carnian conodont Metapolygnathus polygnathiformis (Fig. 37) from metamorphosed carbonate east of the map area at USGS Mesozoic locality 32771. The presence of early Carnian fossils in basal beds of the Burnt Island Conglomerate, and early Carnian fossils in units nearby, suggests that early Carnian limestone may have been eroded after deposition. This could account for the lack of preservation of lower Carnian and older units in the field area. Late Carnian Biostratigraphically significant fossils of late Carnian age occur in the Burnt Island Conglomerate, at various localities in the Hamilton Island Limestone, in the Cornwallis Limestone, and in the unnamed shallow-water limestone. Late Carnian conodonts recovered include Metapolygnathus polygnathiformis, M. carpathicus, M. nodosus, M. reversus, M. sp. aff. M. zoae, and
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Figure 39. Late Carnian conodonts from the Hyd Group. All figures are scanning electron micrographs at ×82 magnification. Illustrated specimens are upper surface views of Pa elements unless otherwise indicated. Photos 1–8: Metapolygnathus polygnathiformis (Budurov and Stefanov, 1965). (1) PP-C2F2, UMIP 302675; (2) lower view, PP-C2F2, UMIP 302676; (3) lateral view, PP-C2F2, UMIP 302677; (4) PISW-C2L2, UMIP 302666; (5) lateral view, PISW-C2L2, UMIP 302660; (6) PISW-C2L2, UMIP 302665; (7) CB2-C3, UMIP 302711; (8) HMNE-C6, UMIP 302688. (9) Metapolygnathus sp. cf. M. reversus (Mosher, 1973); PISW-C2L2, UMIP 302671. Photos 10–11: Juvenile Metapolygnathus nodosus (Hayashi, 1968). (10) HMNE-C16, UMIP 302699; (11) HMNE-C15, UMIP 302697. Photos 12–15: Metapolygnathus carpathicus (Mock, 1979). (12) HMNE-C3, UMIP 302686; (13) lower view, PP-C2F2, UMIP 302680; (14) lateral view, PP-C2F2, UMIP 302678; (15) PP-C2F2, UMIP 302679. Photos 16–22, 28–29: Metapolygnathus nodosus (Hayashi, 1968). (16) HMNE-C6, UMIP 302690; (17) HMNE-C6, UMIP 302691; (18) HMNE-C1, UMIP 302683; (19) lateral view, HMNE-C1, UMIP 302685; (20) lower view, HMNE-C1, UMIP 302684; (21) HMNE-C1, UMIP 302682; (22) lateral view, PISW-C2L2, UMIP 302667; (28) HMNE-C15, UMIP 302693; (29) HMSE-C2, UMIP 302708. Photos 23–25: juvenile Metapolygnathus sp.. (23) HMNE-C6, UMIP 302689; (24) HMNE-C15, UMIP 302696; (25) HMNE-C15, UMIP 302695. Photos 26–27: Metapolygnathus sp. aff. M. zoae Orchard, 1991b. (26) HMNE-C3, UMIP 302687; (27) HMNE-C1, UMIP 302681. Photos 30–33: Metapolygnathus sp. aff. M. nodosus (Hayashi, 1968). (30) CB2-C2, UMIP 302709; (31) lower view, HMSE-C2, UMIP 302706; (32) HMSE-C1, UMIP 302703; (33) lateral view, HMSE-C2, UMIP 302707.
M. primitius (Figs. 39 and 40; Table 3). Other than M. primitius, all of these occur in the nodosus Zone sensu Orchard (1991b) (Fig. 37). Metapolygnathus nodosus ranges into the early Norian, but when abundant, and not co-occurring with Norian fossils, it is interpreted as late Carnian in age. Metapolygnathus primitius ranges into the early Norian further than M. nodosus, and is inter-
preted as either late Carnian or early Norian in age. Populations of M. primitius are covered in the next section. Furthermore, in the genus Metapolygnathus, the basal pit begins shifting toward the anterior end of the conodont in the late Carnian (Orchard, 1991b). This progression can indicate an approximate biostratigraphic position within the late Carnian, particularly in the latest
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Figure 40. Late Carnian through early Norian conodonts from the Hyd Group. All photos are scanning electron micrographs at ×82 magnification. Illustrated specimens are upper surface views of Pa elements unless otherwise indicated. Photos 1–5: Metapolygnathus sp. aff. M. primitius (Mosher, 1970). (1) FL-C3, UMIP 302652; (2) FC2-C2, UMIP 302655; (3) lateral view, FC2-C2, UMIP 302659; (4) FC2-C2, UMIP 302656; (5) FC2-C2, UMIP 302658. Photos 6–17: Metapolygnathus primitius (Mosher, 1970). (6) lateral view, FL-C6, UMIP 302615; (7) lower view, CPE-F2, UMIP 302648; (8) FL-C6, UMIP 302617; (9) CPE-C2, UMIP 302613; (10) CPE-C5, UMIP 302638; (11) CPE-C3, UMIP 302641; (12) lateral view, CPE-F2, UMIP 302649; (13) early Norian, FL-C6, UMIP 302614; (14) CPE-F2, UMIP 302642; (15) subadult, CPE-C2, UMIP 302604; (16) CPE-F2, UMIP 302645; (17) lower view, CPE-C2, UMIP 302612. (18) Neogondolella sp.; CP-C2, UMIP 302650. Photos 19–23: Epigondolella quadrata Orchard, 1991b. (19) SWK-2, UMIP 302599; (20) FL-C17, UMIP 302631; (21) HIW-C2, UMIP 302633; (22) lateral view, CPE-C2, UMIP 302607; (23) lower view, FL-C17, UMIP 302625.
Carnian. Metapolygnathus primitius has a medially located pit, and specimens that look like M. primitius, but have a posterior pit, are referred to M. sp. aff. M. primitius. This conodont ranges into the early Norian, and is treated the same as M. nodosus biostratigraphically. Silberling (in Muffler, 1967) reported a number of late Carnian ammonoids and halobiid bivalves; both fossil groups were recovered in this study.
Cornwallis Limestone samples from the base of the section at the Flounder Cove locality (Fig. 5B, site 99; Fig. 17) yielded Metapolygnathus nodosus, M. sp. cf. M. nodosus, M. sp. aff. M. zoae, and M. sp. aff. M. primitius. In conjunction with Halobia ornatissima (McRoberts, 2003, personal commun.), these samples are late Carnian in age (Fig. 37). The overlying beds contain early Norian fossils and are discussed below.
Conodont biostratigraphy and facies correlations in a Late Triassic island arc, Keku Strait On the northeast shore of Hamilton Island (Fig. 5F, site 60; Fig. 22), in a thick section of Hamilton Island Limestone, the conodonts Metapolygnathus nodosus, M. polygnathiformis, M. carpathicus, and M. sp. aff. M. zoae, halobiid bivalves (McRoberts, 2003, personal commun.), and tropitid ammonoids, including Discotropites? sp., Shastites sp., and Hannoceras sp. (Krystyn, 2003, personal commun.), all contribute to a late Carnian age. Anterior migration of the basal pit in some of the specimens of M. nodosus indicates the upper portion of the nodosus Zone (Orchard, 1991b), and the ammonoids indicate the Dilleri and Welleri ammonoid zones (Krystyn, 2003, personal commun.). All of the Dilleri Zone ammonoids occur in a debris-flow conglomerate (Fig. 23), but younger conodonts of the upper portion of the nodosus Zone (Welleri Zone) occur in the host rocks, suggesting intraformational reworking. In the Hamilton Island Limestone on the southeast shore of Hamilton Island (Fig. 5F, site 67), a variant of Metapolygnathus nodosus, called here M. sp. aff. M. nodosus, is believed to be from the youngest part of the nodosus Zone (Orchard, 2003, personal commun.). It co-occurs with the ammonoid Discotropites? sp. (Krystyn, 2003, personal commun.) and Halobia sp. cf. H. superba (McRoberts, 2003, personal commun.), all of which indicate a late Carnian age (Fig. 37). In the Cape Bendel region (Fig. 5E), two different rock types are faulted into proximity. The first of these (Fig. 5E, site 84), the Hamilton Island Limestone, contains Metapolygnathus sp. cf. M. nododus, M. reversus?, M. sp. cf. M. reversus, and M. polygnathiformis. These conodonts are less developed than later specimens of Metapolygnathus and are indicative of the lower portion of the nodosus Zone (Fig. 37). The other rock type (Fig. 5E, site 83) belongs to the currently unnamed shallowwater limestone, and contains M. polygnathiformis and M. sp. aff. M. nodosus. As at the southeast shore of Hamilton Island, this variation of M. nodosus is believed to be from the youngest portion of the nodosus Zone (Orchard, 2003, personal commun.). Several other sites in the Hamilton Island Limestone also produced late Carnian fossils, though these sites did not have well-exposed sections like those on Hamilton Island. In the Hamilton Island Limestone on southwestern Payne Island (Fig. 5C, site 62), the conodonts Metapolygnathus nodosus, M. polygnathiformis, and M. sp. cf. M. reversus indicate the middle portion of the nodosus Zone (Fig. 37). A sample from the Hamilton Island Limestone at the Squawking Crow locality (Fig. 5C, site 97) yielded M. nodosus and M. polygnathiformis. Along with a few tropitid ammonoids, and Halobia ornatissima? (McRoberts, 2003, personal commun.), these conodonts indicate the upper portion of the nodosus Zone (Fig. 37). In the Hamilton Island Limestone on a small island west of Payne Island (Fig. 5C, site 61), the bivalves H. sp. cf. H. superba and H. sp. cf. H. ornatissima (McRoberts, 2003, personal commun.) and a tropitid ammonoid indicate a late Carnian age. Finally, an isolated sample from the Hamilton Island Limestone in Portage Pass (Fig. 5F, site 66) produced M. polygnathiformis and M. carpathicus, indicating the nodosus Zone. This same area yielded silicified
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H. ornatissima (Fig. 5F, site 66, and Fig. 27), which is also late Carnian (Fig. 37). Late Carnian or Early Norian In the absence of a distinct conodont boundary at Tozer’s (1967) Carnian-Norian boundary, conodonts alone indicate either latest Carnian or earliest Norian age. In conjunction with shortranging ammonoids and halobiid bivalves, samples may be specifically assigned to either the late Carnian or early Norian. Not every site or sample can be determined so precisely, however, so this section documents those samples determined to be either late Carnian or early Norian. Samples with Metapolygnathus primitius, or M. primitius and other metapolygnathids (Fig. 40; Table 3), conform to the primitius Zone of late Carnian or early Norian age (Fig. 37). The bivalves Halobia radiata, H. austriaca, and H. superba (McRoberts, 2003, personal commun.) also overlap the boundary (Fig. 37). The Cornwallis Limestone and the Hamilton Island Limestone have produced fossils of this age. A sample from Cornwallis Limestone on a small island southwest of Kousk Island (Fig. 5D, site 74) and two samples from Cornwallis Limestone adjacent to the Keku Volcanics on Cornwallis Peninsula (Fig. 5A, sites 95 and 96) yielded Metapolygnathus primitius. A sample from Cornwallis Limestone overlying the neptunian dikes on the eastern side of Kuiu Island (Fig. 5B, site 55) yielded M. sp. aff. M. primitius. The ammonoids Thisbites and Styrites from these beds (Silberling in Muffler, 1967) indicate proximity to the Carnian-Norian boundary (Silberling, 2002, personal commun.; Tozer, 1994) and place these beds in the primitius Zone (Fig. 37). Several samples from Cornwallis Limestone on Cornwallis Peninsula (Fig. 5A, site 70) contain M. primitius, M. sp. aff. M. primitius, and Neogondolella sp. These are all of late Carnian or early Norian age in the nodosus and/or primitius conodont zones (Fig. 37). The lower portion of the Cornwallis Limestone on Big Spruce Island (Fig. 5A, site 56, and Fig. 8) yielded an indeterminate metapolygnathid or epigondolellid of late Carnian or early Norian age respectively. Fewer samples in the Hamilton Island Limestone show Carnian-Norian boundary affinities. Hamilton Island Limestone from the top of Cathedral Falls (Fig. 5G, site 65) had Halobia radiata (McRoberts, 2003, personal commun.) and the same unit on southwest Payne Island (Fig. 5C, site 64) had H. radiata and H. sp. cf. H. austriaca (McRoberts, 2003, personal commun.). Both of these are of late Carnian or early Norian age (Fig. 37). Early Norian Early Norian fossils occur in the Cornwallis Limestone, the Hamilton Island Limestone, and in the Hound Island Volcanics. The conodonts Metapolygnathus primitius, Epigondolella quadrata, E. triangularis uniformis, E. triangularis triangularis, E. sp. aff. E. triangularis, and Misikella longidentata all occur in the Cornwallis Limestone. However, only E. quadrata, E. triangularis uniformis, E. triangularis triangularis, and E. sp. aff. E. triangularis
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occur in the Hound Island Volcanics during this time interval (Figs. 40–42; Table 3). Epigondolella triangularis progressively develops more ornate forms during the early Norian (Orchard, 1991b), allowing further refinement of age within the triangularis Zone (Fig. 37). Additionally, a sample can be interpreted as early Norian when M. primitius and E. quadrata are abundant and co-occur with E. triangularis. Older, less ornate forms of E. triangularis are typical of the Cornwallis Limestone, and the younger, more ornate forms are common in the Hound Island Volcanics. Halobiid bivalves are also useful in determining an early Norian age (McRoberts, 1997), and are the only definitive early Norian fossils recovered from the Hamilton Island Limestone. Epigondolella triangularis ranges into the middle Norian (Fig. 37) in the field area as supported by the co-occurrence of E. triangularis with middle Norian faunas as discussed in the next section. Faunas that solely contain E. triangularis are assumed to be early Norian. Many samples in the Cornwallis Limestone indicate an early Norian age. On one of the small islands southwest of Kousk Island (Fig. 5D, site 74), Metapolygnathus primitius and Epigondolella sp. cf. E. quadrata occur in the Cornwallis Limestone. This indicates an age of early Norian in the primitius or quadrata zones (Fig. 37). Samples from the Cornwallis Limestone on northern Cornwallis Peninsula (Fig. 5A, site 70) contain M. sp. aff. M. primitius, and E. sp. cf. E. quadrata, indicating an early Norian age in the quadrata Zone (Fig. 37). Another sample from Cornwallis Peninsula with no corresponding USGS locality yielded M. primitius, E. quadrata, and E. triangularis uniformis (Fig. 5A, site 69). These are of the lower part of the triangularis Zone (Fig. 37). The upper portion of the Cornwallis Limestone on Big Spruce Island (Fig. 5A, site 56; Fig. 8) contains E. sp. cf. E. quadrata of early Norian age. Most of the Cornwallis Limestone at the Flounder Cove succession (Fig. 5B, site 99; Fig. 17) is of early Norian age. Samples in the fine-grained limestones above the lower few meters contain Metapolygnathus primitius, M. sp. aff. M. primitius, the bivalves Halobia beyrichi (Halobia cf. H. alaskensis of Silberling in Muffler, 1967) and H. cordillerana (McRoberts, 2003, personal commun.) and the ammonoids Stikinoceras kerri (Mojsisovicsites sp. of Silberling in Muffler, 1967) and Griesbachites? sp. (Krystyn, 2003, personal commun.). These are from the Kerri Zone and correspond to the Norian portion of the primitius Zone (Fig. 37). In the coarse-grained limestone capping the section, the conodonts Epigondolella quadrata, E. triangularis uniformis, E. triangularis triangularis, and Misikella longidentata indicate the triangularis Zone in the early Norian (Fig. 37). In the area of the prominent point west of Hound Island on eastern Kuiu Island, the ammonoid Guembelites clavatus of the Kerri Zone (Fig. 37) was found in calcareous sandstone of the Cornwallis Limestone (Fig. 5B, site 72). West of here (Fig. 5B, site 73), an epigondolellid was found in calcareous sandstone with abundant plant fossils, and even farther west (Fig. 5B, site 68), Epigondolella sp. cf. E. quadrata and a metapolygnathid were found in lithoclastic sediments with reworked Devonian
conodonts (Fig. 38). These indicate an overall early Norian age for this area. In the Hamilton Island Limestone in Portage Pass (Fig. 5F, site 66), the bivalves Halobia beyrichi, H. cordillerana, and H. sp. cf. H. lineata (McRoberts, 2003, personal commun.) together indicate an early Norian age (Fig. 37). Also in the Hamilton Island Limestone, H. cordillerana and H. lineata (McRoberts, 2003, personal commun.) occur at the Squawking Crow locality, indicating an early to middle Norian age (Fig. 5C, site 97). These are the only definitive early Norian ages in the Hamilton Island Limestone. In the Hound Island Volcanics on the southwest side of Hamilton Island (Fig. 5F, site 85) the bivalves Halobia sp. cf. H. beyrichi and H. sp. cf. H. fallax (McRoberts, 2003, personal commun.) indicate a probable early Norian age (Fig. 37). On the west side of Hound Island, USGS Mesozoic locality M1899 (Fig. 5D, site 79) produced Epigondolella quadrata, E. triangularis uniformis, E. triangularis triangularis, and the bivalves H. sp. cf. H. beyrichi and H. sp. cf. H. lineata (McRoberts, 2003, personal commun.), and locality M1923 (Fig. 5D, site 80) had H. beyrichi? (McRoberts, 2003, personal commun.). Locality M1921 on the west side of Hound Island (Fig. 5D, site 81) produced specimens of E. sp. cf. E. triangularis, and a sample on the western side of northern Hound Island (Fig. 5D, site 82) yielded E. triangularis? of early Norian age. All of these indicate a late early Norian age for the western side of Hound Island (Fig. 37). Middle Norian Biostratigraphically significant fossils of middle Norian age occur only in the Hound Island Volcanics. They include the conodonts Epigondolella spiculata, E. matthewi, E. sp. cf. E. postera, E. transitia, Neogondolella sp. cf. N. steinbergensis (Fig. 42; Table 3), the bivalve Halobia fallax, and the conodonts E. triangularis and E. sp. aff. E. triangularis, which continue into the middle Norian (Fig. 37). Epigondolella triangularis and E. transitia have been reported previously only in the early Norian (Orchard, 1991b), and their later ranges and respective ages of extinction are undocumented. Epigondolella triangularis always occurs with E. spiculata in the field area, supporting the presence of middle Norian E. triangularis. Epigondolella sp. aff E. triangularis and E. transitia may have a similar range, though further documentation of range for these species requires populations richer in both species over a larger stratigraphic interval. In the North American Cordillera, the conodont Epigondolella multidentata (Fig. 37) normally signifies the base of the middle Norian (Orchard, 1991b). In Europe, other species are used to identify the base of the middle Norian (e.g., Kozur in Channell et al., 2003). Because neither E. multidentata nor the European species occur in the field area, the base of the middle Norian was not located. In the lower, finer-grained beds in the Gil Harbor mudflat (Fig. 5B, site 87), samples yielded Epigondolella triangularis triangularis, E. sp. aff. E. triangularis, E. spiculata, E. transitia,
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Figure 41. Early through middle Norian conodonts from the Hyd Group. All photos are scanning electron micrographs at ×82 magnification. Illustrated specimens are upper surface views of Pa elements unless otherwise indicated. Photos 1–12: Epigondolella triangularis uniformis Orchard, 1991b. (1) FL-C16, UMIP 302620; (2) FL-C17, UMIP 302622; (3) lower view, FL-C17, UMIP 302628; (4) CPE-C2, UMIP 302609; (5) lower view, CPE-C2, UMIP 302608; (6) FL-C17, UMIP 302623; (7) HIW-C2, UMIP 302634; (8) FL-C16, UMIP 302619; (9) FL-C16, UMIP 302621; (10) CPE-C2, UMIP 302606; (11) FL-C17, UMIP 302627; (12) lateral view, FL-C17, UMIP 302630. Photos 13–18: Epigondolella triangularis triangularis (Budurov, 1972). (13) GH-C1, UMIP 302572; (14) HIW-C2, UMIP 302635; (15) FL-C17, UMIP 302629; (16) lateral view, GH-C1, UMIP 302579; (17) BT-C2, UMIP 302586; (18) GH-C1, UMIP 302581. Photos 19–20: Epigondolella transitia Orchard, 1991b. (19) GH-C1, UMIP 302573; (20) lower view, GH-C1, UMIP 302574. Photos 21–23: Epigondolella sp. aff. E. triangularis (Hayashi, 1968). (21) GH-C1, UMIP 302571; (22) GH-C1, UMIP 302575; (23) lateral view, GH-C1, UMIP 302583. (24) Misikella longidentata Kozur and Mock, 1974; FL-C17, UMIP 302636. (25) Juvenile Epigondolella sp. Mosher, 1968; triangularis Zone, FL-C17, UMIP 302632.
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Figure 42. Middle through late Norian conodonts from the Hyd Group. All figures are scanning electron micrographs at ×82 magnification. Illustrated specimens are upper surface views of Pa elements unless otherwise indicated. Photos 1–7: Epigondolella spiculata Orchard, 1991b. (1) GH-C1, UMIP 302580; (2) lower view, GH-C1, UMIP 302576; (3) BT-C2, UMIP 302587; (4) GH-C2, UMIP 302570; (5) GH-C1, UMIP 302578; (6) HL-C6, UMIP 302588; (7) lateral view, GH-C1, UMIP 302582. (8) Epigondolella matthewi Orchard, 1991b; GH-C1, UMIP 302577. (9) Epigondolella sp. cf. E. postera Kozur and Mostler, 1971; HIE-C3, UMIP 302596. (10) Subadult Epigondolella sp.; spiculata Zone, GH-C1, UMIP 302585. (11) Neogondolella sp. cf. N. steinbergensis (Mosher 1968); HIE-C2, UMIP 302592. Photos 12–14: Neogondolella sp. (12) HL-C6, UMIP 302589; (13) HIE-C3, UMIP 302595; (14) lower view, HL-C1, UMIP 302591. (15) Juvenile Neogondolella sp.; middle Norian, lateral view, HIE-C2 UMIP 302593. (16) Juvenile late Norian Epigondolella sp.; GH-C6, UMIP 302559. Photos 17–22: Epigondolella englandi Orchard, 1991b. (17) GH6, UMIP 302563; (18) lateral view, GH-C6, UMIP 302558; (19) GH6, UMIP 302564; (20) GH-C6, UMIP 302560; (21) lateral view, GH6, UMIP 302568; (22) lower view, GH6, UMIP 302569. Photos 23–25: Epigondolella bidentata Mosher, 1968. (23) Lower view, GH-C6, UMIP 302555; (24) GH-C6, UMIP 302556; (25) GH-C6, UMIP 302557. Photos 26–31: Epigondolella tozeri Orchard, 1991b. (26) GH-C6, UMIP 302554; (27) GH-C6, UMIP 302553; (28) GH-C6, UMIP 302552; (29) GH6, UMIP 302565; (30) lower view, GH6, UMIP 302566; (31) lateral view, GH6, UMIP 302567. Photos 32–34: Epigondolella sp. aff. E. mosheri Kozur and Mostler, 1971. (32) GH-C6, UMIP 302551; (33) lateral view, GH-C6, UMIP 302561; (34) GH-C6, UMIP 302562.
and E. matthewi. These are indicative of the spiculata Zone (Fig. 37). Between Gil Harbor and Kadake Bay along the coast (Fig. 5B, site 46), an individual sample yielded E. triangularis triangularis and E. spiculata of the middle Norian spiculata
Zone (Fig. 37). Finally, samples from the east side of Hound Island (Fig. 5D, site 86) had E. sp. cf. E. postera, E. spiculata, Neogondolella sp. cf. N. steinbergensis, other Neogondolella sp. and the bivalve Halobia fallax (Silberling et al., 1997; McRoberts,
Conodont biostratigraphy and facies correlations in a Late Triassic island arc, Keku Strait 2003, personal commun.). Monotis (Eomonotis) ?pinensis was also reported (Muffler, 1967; Silberling et al., 1997). These indicate the postera Zone higher in the middle Norian (Fig. 37). Late Norian Only two localities in the Hound Island Volcanics yielded late Norian fossils, and this study investigated only the Gil Harbor site. In the Gil Harbor mudflat, conodonts recovered from the silicified fossil beds (Fig. 5B, site 87) include Epigondolella bidentata, E. tozeri, E. englandi, and E. sp. aff. E. mosheri (Fig. 42; Table 3). Epigondolella tozeri originates in the middle Norian, and E. bidentata and E. englandi originate in the late Norian. This association suggests the bidentata Zone in the late Norian (Fig. 37). Monotis subcircularis and the hydrozoan Heterastridium conglobatum occur in these beds and at USGS Mesozoic locality M1898 in Hamilton Bay (Fig. 5G) (Muffler, 1967). Monotis subcircularis occurs only in the late Norian Cordilleranus Zone (Fig. 37), agreeing with the conodont data. The abundant late Norian fossils indicate that the conodont E. sp. aff. E. mosheri is a late Norian predecessor of the Rhaetian E. mosheri. Large individuals of Heterastridium in the Gil Harbor mudflat (Fig. 5B, site 87) have a size range of 2–3 cm. According to the theory that maximum diameters of Heterastridium can assist age deterination (Stanley et al., 1994) these sizes indicate an upper middle Norian age (Stanley, unpublished data). Conodonts and bivalves, however, clearly place these beds as latest Norian in age, contradicting this theory. DISCUSSION Reworking of Index Fossils There is the possibility in any biostratigraphic work that important index fossils have been reworked from lower stratigraphic levels thus obscuring or incorrectly indicating the age of a unit. Of the index fossils used in this study, both ammonoids and conodonts occur as reworked fossils. However, every known rock unit in the field area containing reworked index fossils is conglomeratic. Paleozoic index fossils in Triassic rock were recovered only from conglomerates known to contain Paleozoic clasts. Devonian conodonts occur in the sandy and conglomeratic lithoclastic limestone from the east side of Kuiu Island (Fig. 5B, site 68; Table 3). Complete specimens were extracted from individual limestone clasts, and broken and/or rounded or abraded specimens were extracted from whole-rock samples representing conodonts in sand- to pebble-sized clasts or matrix. These Devonian conodonts represent reworking of Givetian- and Frasnianaged units. Broken Permian conodonts have also been extracted from the matrix of conglomeratic Hamilton Island Limestone in the Keku Islets (Fig. 5C, sites 62 and 97), and whole Permian conodonts have been extracted from limestone clasts in the Burnt Island Conglomerate (Katvala, unpublished data). Besides
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conodonts, Paleozoic brachiopods are known to occur in conglomerate clasts in the Keku sedimentary strata, the Burnt Island Conglomerate, and the Hamilton Island Limestone. Some Late Triassic index fossils are reworked as or within intraclasts within conglomerate, and are the same age as index fossils in the rock surrounding the conglomerate. Late Carnian ammonoids and conodonts occur within intraclasts in the debris flows of the lower Hamilton Island Limestone on northeastern Hamilton Island (Fig. 5F, site 60). Additionally, rounded conodonts that appear to be Late Triassic occur in intraclastic limestone at the Flounder Cove locality (Fig. 5B, site 99). One Late Triassic site at Gil Harbor (Fig. 5G, site 87) had obvious reworking of Late Triassic index fossils; this caused two apparent ages to exist in the same bed. In the youngest described bed (Fig. 35; Table 3), ammonoids of the middle Norian Columbianus Zone (Fig. 37) occur with Monotis subcircularis of the Cordilleranus Zone. Additionally, Epigondolella tozeri, previously ascribed to the middle Norian (Orchard, 1991b), occurs with E. englandi, E. bidentata, and E. sp. aff. E. mosheri of the late Norian. This limestone bed contains reworked limestone intraclasts, which may be the source of the age discrepancy. If there is reworking of middle Norian fossils into the late Norian, this may support the size-to-age correlation of Heterastridium conglobatum (Stanley et al., 1994), though it would be impossible to tell. To contrast, without previous documentation of the ranges for the ammonoids and conodonts, it is possible that middle Norian species range into the late Norian. The ammonoids are often well preserved and complete, and no rounded or otherwise abraded conodonts were found. Finally, the bed may be condensed and represent deposition over a long period of time. Regardless, the final deposit is late Norian in age. In every case, beds with known, reworked index fossils in the field area have conglomeratic lithologies. Furthermore, there are abundant sources of possible reworked fossils with which to contaminate biostratigraphic data. Even mildly reworked conodonts found in the matrix of intraclastic limestone generally show signs of rounding or abrasion. Overall, this indicates that the determinations of age in the Late Triassic of Keku Strait are reliable for further interpretations. Pre–Late Triassic Uplift Throughout the Alexander terrane, uplift and erosion during and/or prior to Late Triassic deposition is recorded by the following: (1) a regional unconformity with incision of Paleozoic units, (2) a lack of reported Middle Permian to Middle Triassic rocks, and (3) the occurrence of rocks of a variety of ages beneath the unconformity (Gehrels et al., 1987). The clasts of Devonian through Early Permian age (Fig. 38; Table 3) found in the Cornwallis Limestone, Hamilton Island Limestone, and lithoclastic beds of the Keku sedimentary strata strongly support pre–Late Triassic uplift in the Keku Strait area. Furthermore, this uplift happened, at least partly, after deposition in the Early Permian and was sufficiently disruptive to expose much of the Paleozoic succession.
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This uplift resulted in a complicated paleotopography during the Late Triassic. Paleozoic outcrops of Silurian, Devonian, Carboniferous, and Permian age occur in close proximity on the west side of Keku Strait. The overlying Triassic units are less complicated structurally, and locally surround Paleozoic units. These “islands” of Paleozoic rock are likely sources for the abundant and locally large Paleozoic clasts deposited in Upper Triassic units. The large, angular block of Carboniferous limestone in Upper Triassic sandstone on Big Spruce Island (Fig. 10) supports the existence of a steep, proximal source area. A similar situation is present on Kupreanof Island, as the Devonian to Permian Cannery Formation is more intensely folded than the overlying Triassic units. Muffler (1967) postulated that Permian thrusting could have juxtaposed dissimilar Permian facies prior to deposition of the Pybus Formation. In fact, Devonian and Mississippian clasts of the Cannery Formation on Kupreanof Island were reported in Permian units underlying the Pybus Formation on Kuiu Island (Berg et al., 1978; Jones et al., 1981). Karl et al. (1999) noted that the Paleozoic rocks on eastern Kupreanof Island were of higher metamorphic grades than the Triassic rocks, supporting a pre-Mesozoic metamorphic event. On Kupreanof Island, numerous thrust faults from at least two generations of thrusting (Karl et al., 1999) support a complex compressional history for the Alexander terrane. Overall, this thrusting could account for much of the pre–Late Triassic uplift, though some uplift after deposition of the Pybus Formation is necessary to account for its shedding of debris into shallow-marine or terrestrial Upper Triassic units. Furthermore, the presence of Devonian clasts within Late Triassic sediments (Fig. 5B, site 68, and Fig. 38) suggests a long history of erosion, supporting earlier uplift. This uplift has been attributed to latest Permian or Triassic rifting in the past (Gehrels and Saleeby, 1987; Gehrels et al., 1987). This interpretation depends partly on the presence of felsic to mafic bimodal volcanism and a lack of deformation associated with uplift. However, the Burnt Island Conglomerate represents the earliest Triassic deposition in the field area and does not contain any late Permian or Triassic volcanic rocks, although it does contain older volcaniclastic rock. Additionally, new age data place the felsic igneous rock of the Keku Volcanics in the Cretaceous (Mortenson, 2004, personal commun.) instead of the Triassic, thereby withdrawing the Keku Volcanics from the Hyd Group. Furthermore, there is abundant evidence of pre–Late Triassic deformation as noted above. The removal of the Keku Volcanics from the Hyd Group, the lack of volcanic rock in both the Burnt Island Conglomerate and Hamilton Island Limestone, and evidence for compressional tectonics in the Permian all seem to preclude rifting as the major source of pre–Late Triassic uplift. Hyd Group Deposition Although a Middle Triassic age was reported nearby (Wardlaw in Karl et al., 1999), the earliest age of deposition indicated in the Hyd Group is early Carnian. The Burnt Island Conglomer-
ate of early Carnian to late Carnian age represents initial, rapid infill of the basin. This probably occurred at the end of a period of uplift and erosion (Gehrels et al., 1987), in conjunction with subsidence and/or sea-level rise. Widespread carbonate deposition began in the late Carnian as represented in the Hamilton Island Limestone and the Cornwallis Limestone. Turbiditic successions and debris flows in the Hamilton Island Limestone record continued and relatively rapid infilling in deeper portions of the basin (Fig. 22). Fossils from the Hamilton Island Limestone indicate deposition through most or all of the late Carnian. These thick, deeper-water upper Carnian deposits without significant facies change suggest relative sealevel rise throughout this interval. Fewer upper Carnian deposits are exposed in the Cornwallis Limestone. These deposits directly overlie either the pre–Late Triassic erosional unconformity or the more proximal facies of the Keku sedimentary strata. This also supports relative sea-level rise during the late Carnian. Although thicker Carnian deposits of the Cornwallis Limestone were not located on the surface, they may be preserved east of Kuiu Island in the waters of Keku Strait or in the subsurface. The unnamed shallow-water limestone on the east side of the strait is upper Carnian and may represent an upper Carnian equivalent of the Cornwallis Limestone on the opposite side of the basin. Placing the felsic igneous rock of the Keku Volcanics in the Cretaceous leaves the formerly associated Keku sedimentary strata in the Hyd Group. The lithoclastic beds and neptunian dikes represent nearshore deposition and they underlie or laterally grade into the Cornwallis Limestone. Lithoclastic limestone beds of the Cornwallis Limestone are transitional between the purely lithoclastic beds and the massive carbonate of the Cornwallis Limestone. Their stratigraphic relationship with the Cornwallis Limestone and the fossils adjacent to the neptunian dikes indicate that these nearshore deposits are mainly late Carnian, but they could be early Carnian to early Norian in age. Widespread carbonate deposition continued into the early Norian, where more extensive shallow-water carbonates developed, as represented by most of the Cornwallis Limestone. Few early Norian deposits are documented in the Hamilton Island Limestone. The only areas where the upper contact of the Hamilton Island Limestone is exposed are on Hamilton Island and in Hamilton Bay (Fig. 5F and 5G). Unfortunately, on northern Hamilton Island, this contact is intruded by Tertiary gabbro, on southern Hamilton Island this contact is intruded and deformed, and in Hamilton Bay this same contact is cut by several faults. This may account for the scarcity of early Norian deposits in the Hamilton Island Limestone. Although no transitional facies have been found in the field area, the shallow-water Cornwallis Limestone and the deeper-water Hamilton Island Limestone are probably lateral facies of each other (Muffler, 1967). Shallow-water corals and plant material are present as reworked bioclasts in the Hamilton Island Limestone, and the same late Carnian ages occur in both units. A west-dipping thrust fault (previously mapped
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Figure 43. (A) Close-up of the thrust fault in the Keku Islets. Dark-colored, folded beds of the Hamilton Island Limestone underlie light-colored, massive Permian limestone of the Pybus Formation (Fig. 5C, site 61). White arrows indicate the fault boundary. Scale is in centimeters. (B) Map of Payne Island and surrounding islets in the northern Keku Islets. Southwest-dipping thrust fault just southwest of Payne Island indicated on map.
as a high-angle fault; Muffler, 1967) occurs just west of Payne Island (Fig. 5C, site 61) and brings the Permian Pybus Formation over top of the ductile Hamilton Island Limestone (Fig. 43). It separates the Keku sedimentary strata and all of the Cornwallis Limestone from outcrops of Burnt Island Conglomerate and Hamilton Island Limestone, and could account for their current proximity. Relatively continuous carbonate deposition through the Carnian-Norian boundary in the field area has the potential for refining the definition of this boundary. Ammonoids, bivalves, conodonts, and radiolarians are preserved in fossil assemblages of the Hyd Group, providing abundant fossil age control. To date, preservation of this interval has been found only in the Cornwallis Limestone. If more Norian fossils can be recovered from the Hamilton Island Limestone, there is potential for examination of this boundary in both shallow- and deep-water sediments. Orchard (1991a, 1991b) suggested the nodosus, communisti, and primitius conodont zones as other potential levels for formal definition of the Carnian-Norian boundary. The abundance of upper Carnian limestone in the Hamilton Island Limestone provides potential for investigating these intervals. Although this study does not have enough data to securely place a boundary, the section at the Flounder Cove locality (Fig. 5B, site 99; Fig. 17) exhibits the most potential. The youngest conodont ages in the Cornwallis Limestone are indicative of the triangularis Zone. The oldest conodonts in the Hound Island Volcanics are more advanced morphologically and are indicative of the upper part of the triangularis Zone. Assuming that the onset of extensive basaltic and andesitic volcanism was relatively synchronous throughout the area, this implies an age of onset within the triangularis Zone of late early Norian age. This is supported by the lack of Triassic volcanic material in Carnian and earliest Norian deposits in the Burnt Island Con-
glomerate, Keku sedimentary strata, the Cornwallis Limestone, and the Hamilton Island Limestone. Deeper-water limestone of the Hound Island Volcanics overlies shallower-water facies of the Cornwallis Limestone on Eastern Kuiu Island (Fig. 5B) and the unnamed shallow-water limestone in the Cape Bendel region (Fig. 5E). This suggests continued relative sea-level rise, possibly from tectonic subsidence linked to the onset of volcanism. Deposition of the Hound Island Volcanics continued into the middle and late Norian, and the youngest deposits are limestones of late Norian age. Owing to the scarcity of youngest Triassic deposits, it is not currently possible to determine the level of volcanism in the late Norian, and volcanism may have diminished or ceased in the middle Norian. The scarcity of late Norian deposits is presumably due to the pre– Upper Jurassic unconformity. Figure 44 depicts chronostratigraphic cross sections across the strait. The northern cross section begins in Cornwallis Peninsula, passes through the northern Keku Islets, and ends in the Cape Bendel region (Figs. 4 and 44). The southern cross section begins on eastern Kuiu Island, passes through Hound Island, and ends in the Hamilton Island region (Figs. 4 and 44). In both cross sections, the break between the Cornwallis Limestone and the Hamilton Island Limestone occurs roughly the same distance across the strait. On the northern cross section, this point matches the position of the post-Triassic thrust fault (Fig. 5C, site 61; Fig. 43). Furthermore, it becomes clear on the cross sections that the Cornwallis Limestone, Hamilton Island Limestone, and Hound Island Volcanics have similar age ranges throughout their respective geographic distributions. Figure 45 portrays a block diagram representing depositional extent and facies correlations for the Burnt Island Conglomerate, Keku sedimentary strata, the Cornwallis Limestone, and the Hamilton Island Limestone in the Carnian to early Norian. This
Figure 44. Generalized chronostratigraphic cross-sections across northern (top) and southern (bottom) Keku Strait. Fossil symbols for ammonoids, bivalves, and conodonts indicate biostratigraphic age control with possible ranges indicated by arrows. Specific information on biostratigraphic fossils is in Table 3.
Conodont biostratigraphy and facies correlations in a Late Triassic island arc, Keku Strait
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Figure 45. Block diagrams depicting facies correlations and depositional extent for (A) the Burnt Island Conglomerate, Keku sedimentary strata, Cornwallis Limestone, and Hamilton Island Limestone during Carnian to early Norian time, and (B) the overlying Hound Island Volcanics during early to late Norian time. Island on left side of diagram represents a local topographic high. Not to scale.
diagram displays deposition of clastic rock in nearshore and terrestrial environments, shallow-water and deep-water limestone farther out in the basin respectively, and the initial basin infill underlying them all in the center of the basin. The island on the left side of the diagram represents a topographic high of Paleozoic limestone in Triassic time. A question mark denotes the possible presence of clastic Triassic rock on the right side of the diagram. A break in the block diagram between the Cornwallis Limestone and the Hamilton Island Limestone represents the post-Triassic thrust fault shown in Figure 43.
CONCLUSIONS This study was successful in providing a revised biostratigraphic framework for the Late Triassic rocks in the Keku Strait area. Conodonts, bivalves, and ammonoids revealed ages from late Carnian through late Norian. Overall, these ages generally confirm the dates reported by Muffler (1967), but allowed greater chronological precision over the entire area. Greater precision permits the precise dating of the onset of volcanism in the Hound Island Volcanics as late early Norian. Furthermore, the biostrati-
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Katvala and Stanley
graphic data allow greatly improved correlation between the geologic units. This includes the placement of the Keku sedimentary strata as a facies of the Cornwallis Limestone and the recognition of correlatable age ranges within the Cornwallis and Hamilton Island limestones. Conodont data from reworked clasts also allows reinterpretation of the pre–Late Triassic tectonic history. At least two phases of uplift exposed the Paleozoic succession before Late Triassic time. This uplift was probably due to thrusting as opposed to rifting, and the latest phase was after deposition of the Permian Pybus Formation. Identification of possible Carnian-Norian boundary successions will contribute to future work on this elusive interval. Paleontological samplings increase the size of Triassic macrofossil collections from Keku Strait by several times and represent the most fossiliferous Triassic strata in the Alexander terrane. These fossils will assist in future paleogeographic studies and in facies determination within the Late Triassic carbonate succession. New age data indicate a Cretaceous age for the felsic intrusions of the Keku Volcanics (Mortenson, 2004, personal commun.). However, the presence of Triassic fossils and stratigraphic relationships with other Triassic units indicate that several sedimentary deposits assigned to the Keku Volcanics by Muffler (1967) are of Triassic age. This revision requires removal of the Cretaceous Keku Volcanics from the Hyd Group. The remaining Keku sedimentary strata have thus been placed in the Hyd Group as an informal unit. Because the geologic history and outcrop extent of the Keku sedimentary strata are poorly understood, it is inappropriate to give a formal designation at this time. However, given their close association with the Cornwallis Limestone, similarity to lithoclastic limestone included within the Cornwallis Limestone, and geographic restriction to Cornwallis Peninsula, the Keku sedimentary strata should be associated with the Cornwallis Limestone. A redefined Cornwallis Limestone would also be able to encompass the outcrops of unnamed shallow-water limestone on the east side of Keku Strait. Future work should focus on the rocks of Cornwallis Peninsula. Overall, the Keku Strait area, with its more extensive exposures and relatively low metamorphism, proves to have the best known exposures of Late Triassic rock in the Alexander terrane. The Burnt Island Conglomerate represents initial infill of the basin, and the overlying Keku sedimentary strata, Cornwallis Limestone, and Hamilton Island Limestone represent a proximal to distal facies succession during the late Carnian and early Norian. An absence of volcanism in older deposits is succeeded by extensive basaltic and andesitic volcanism of the ubiquitous Hound Island Volcanics in the late early Norian. Relative sea-level rise was prevalent throughout deposition of the Hyd Group. The wide variety of volcanic and sedimentary facies displaying rapid facies changes, peri-platform carbonates in the forms of debris flows, slumps, and turbidites, and the presence of terrane-endemic gastropods (Blodgett and Frýda, 2001; Frýda and Blodgett, 2001) support deposition around an island arc (see Soja, 1996). Finally, new ties between different stratigraphic columns on either side of the strait provide a linkage between rocks that might appear unre-
lated elsewhere in the Alexander Terrane. Thus, the Keku Strait area forms the best standard for comparison when examining the Late Triassic succession in the Alexander terrane. TAXONOMIC NOTES Brief descriptions of select Late Triassic conodonts recovered and subsequently applied to the Late Triassic biostratigraphy are included below. Ones without descriptions follow previous work (Budurov and Stefanov, 1965; Hayashi, 1968; Mosher, 1970, 1973; Kozur and Mostler, 1971; Budurov, 1972; Kozur and Mock, 1974; Mock, 1979; Kozur, 1980; Orchard, 1983, 1991a, 1991b). Each species description references the illustrating figure. All conodonts are stored at the University of Montana–Museum of Paleontology. Epigondolella matthewi Orchard, 1991b Figure 42, Photo 8 Epigondolella matthewi has a mostly unornamented, round posterior margin and is characterized by a relatively broad, biconvex platform with two to four large denticles on each anterior platform margin (Orchard, 1991b). The blade has relatively few, large denticles that pass into a carina composed of several discrete nodes that usually do not reach the posterior end of the platform (Orchard, 1991b). Although the figured specimen has one large denticle on one margin and three on the other, its shape is indicative of E. matthewi and is herein considered a variant. Epigondolella sp. aff. E. mosheri Kozur and Mostler, 1971 Figure 42, Photos 32–34 These specimens from Keku Strait, Alaska are larger, more elongate, and have more denticles on the lateral margins than Epigondolella bidentata. They have five carinal nodes posterior of a prominent denticle pair and are not as elongate or as narrow posteriorly as E. mosheri. Orchard (1991b) noted how small growth stages of E. mosheri are very similar to E. bidentata, and how transitional forms in the Cordilleranus Zone are larger, have five carinal nodes, and have a broader posterior platform. Thus, these specimens may be variations on the form of E. mosheri, or transitional forms between E. bidentata and true E. mosheri. Epigondolella sp. cf. E. postera Kozur and Mostler, 1971 Figure 42, Photo 9 This specimen has the anterior platform denticulation common to Epigondolella postera, but the carina is straight with little deflection. Furthermore, it lacks the posterior margin and denticulation found in E. spiculata. It is a juvenile and may be either E. postera, E. spiculata, or transitional between them. Because this specimen most resembles E. postera, it
Conodont biostratigraphy and facies correlations in a Late Triassic island arc, Keku Strait is tentatively called E. sp. cf. E. postera. Also note the Epigondolella specimen represented in Figure 42, Photo 10, which also appears to be a young E. postera, E. spiculata, or similar form. Epigondolella sp. aff. E. triangularis Figure 41, Photos 21–23 These specimens from Keku Strait, Alaska, resemble Epigondolella spatulata with their generally shorter platform and blade, transversely elongate nodes on the anterior platform, and noticeable expansion of the platform posterior of the aforementioned nodes. They differ from E. triangularis in that the unexpanded anterior platforms are relatively longer, the posterior expansion is not as great, and they have lessprominent posterior ornamentation. Specimens of E. sp. aff. E. triangularis occur with E. triangularis in the early Norian and both E. triangularis and E. spiculata in the middle Norian. Epigondolella spiculata Orchard, 1991b Figure 42, Photos 1–7 When compared with the holotype of Epigondolella spiculata, the Keku Strait specimens are not as asymmetric, but they do not have the posterior platform expansion typical of E. triangularis triangularis. As in E. spiculata, the blade is shorter than in E. triangularis and the anterior nodes are large, discrete, and sharp. The carina extends to the posterior and a prominent carinal node on the posterior platform projects posteriorly like E. spiculata. These specimens are strongly denticulate posteriorly and the denticles project outward from the platform margin. The main differences from E. spiculata are on the lower surface and correspond somewhat to E. triangularis. In profile, the lower surface is not flat or convex, but is slightly concave anterior to the posteriorly upturned margin. This concavity is not as pronounced as in other Norian epigondolellids, and the posteriorly projecting denticles make the upturned posterior margin appear similar to E. spiculata. The asymmetric keel is not straight or obliquely truncated as in typical E. spiculata, but is instead weakly bifurcated as in rare occurrences documented by Orchard (1991b). Unlike in E. triangularis, this bifurcation does not extend to the pit. Finally, the loop surrounding the basal pit is more pronounced, as in E. triangularis. Epigondolella tozeri Orchard, 1991b Figure 42, Photos 26–31 Epigondolella tozeri is strongly denticulate with two to four high anterior platform denticles and strong nodes on the tapered to subparallel margins of the
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pointed to narrowly truncated posterior platform (Orchard, 1991b). It has discrete, low carinal nodes that extend to the posterior end and the lower surface has a basal pit beneath the anterior third of the platform and a posteriorly broad basal scar (Orchard, 1991b). When compared with the co-occurring E. sp. aff. E. mosheri of the Keku Strait samples, this species has a more robust platform, more posterior denticulation, the broad basal scar, and a much less tapered posterior platform margin. The specimen in Figure 42, Photo 26 appears to be transitional in form between E. tozeri and E. sp. aff. E. mosheri. This species has not been previously reported from the late Norian, though it may be reworked from middle Norian strata as discussed above. Epigondolella triangularis (Budurov, 1972) sensu lato Figure 41, Photos 1–18 Orchard (1991b) documented a trend toward increased platform ornamentation, increased posterolateral expansion, and decreased relative blade length in early Norian populations of Epigondolella triangularis. This trend is present within specimens found in this study, from the earlier-occurring E. triangularis uniformis Orchard 1991b into E. triangularis triangularis Orchard 1991b. Although Orchard (1991b) did not report middle Norian populations of E. triangularis triangularis, they do occur in Keku Strait, because E. triangularis triangularis always occurred in samples that had the middle Norian E. spiculata. These middle Norian specimens still display the increased ornamentation, increased posterolateral expansion, and decreased relative blade length that are typical of younger forms of E. triangularis. Metapolygnathus sp. aff. M. nodosus Figure 39, Photos 30–33 This species of Metapolygnathus is similar to M. nodosus (Hayashi, 1968) but has a more linguiform shape, a flatter platform that in profile is still slightly raised anteriorly, a higher posterior carina, and a posteriorly located basal pit that is slightly shifted toward the anterior when compared with other Metapolygnathus species. Orchard (2003, personal commun.) believes this species may be an advanced form of M. nodosus distinctive of the latest Carnian. Metapolygnathus ex. gr. M. polygnathiformis (Budurov and Stefanov, 1965) Figure 39, Photos 1–8 For the purpose of this study, Metapolygnathus polygnathiformis represents a group of meta polygnathids with subsymmetrical, subquadrate
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Katvala and Stanley platforms with unornamented margins of uniform height (Orchard, 1991b). This definition includes forms that have been called Metapolygnathus polygnathiformis and M. noah Hayashi 1968, as well as larger forms of M. nodosus with fused nodes that appear smooth and are sometimes hard to distinguish. Figure 39, 6 may represent a smooth form of M. nodosus or a transitional form between nodose and non-nodose species of Metapolygnathus. In this study, every sample with large, non-nodose Metapolygnathus also had smaller, definitive specimens of M. polygnathiformis, allowing positive identification of the conodont population. Although distinctions between early Carnian and late Carnian forms may exist, they are not pursued in this study, as M. polygnathiformis nearly always co-occurred with M. carpathicus and/or M. nodosus, indicating the late Carnian. Metapolygnathus sp. aff. M. primitius (Mosher, 1970) Figure 40, Photos 1–5 Specimens referred to as Metapolygnathus sp. aff. M. primitius (Fig. 40, 1–5) display a similar platform shape and ornamentation as M. primitius (sensu Orchard, 1991b), but the basal pit is not shifted very far to the anterior of the posterior truncation of the keel. In typical M. primitius, the basal pit has shifted to the anterior when compared with M. nodosus, and this trend continues in early Norian species of Epigondolella. Metapolygnathus sp. aff. M. primitius may represent a separate lineage from the M. primitius that leads to Epigondolella quadrata. Metapolygnathus sp. cf. M. reversus (Mosher, 1973) Figure 39, Photo 9 Unlike most metapolygnathids, Metaploygnathus reversus has no free blade. The single specimen referred to M. sp. cf. M. reversus is similar in platform shape and ornamentation to M. reversus, but the exact characters of the anterior blade are not known because it is not preserved. However, most of the specimen is present, and either the free blade is absent as in M. reversus, or, if preserved, it would represent a very small portion of the total platform length. Metapolygnathus sp. aff. M. zoae Orchard, 1991b Figure 39, Photos 26–27 This relatively elongate species of Metapolygnathus is characterized by about four large, well-defined, but low, circular nodes on each anterior platform margin (Orchard, 1991b). When compared with M. nodosus, this species has larger, more prominent, broader, and, in upper view, more rounded anterior platform
nodes (Orchard, 1991b). The specimens of M. sp. aff. M. zoae differ from M. nodosus by having circular nodes that are not fully differentiated as in M. zoae, but are prominent and may be fused together in larger specimens. ACKNOWLEDGMENTS The Robert and Leigh M. Besancon Fellowship and the Bertha Morton Scholarship of the University of Montana, National Science Foundation grants EAR-9624501 and EAR-0229795, and the Subcommission on Triassic Stratigraphy and IGCP project 467 provided financial support for this research. The people and organizations of Kake, Alaska, and Mark Nay of the U.S. Forest Service in Corvallis, Oregon, provided additional logistic support. Jim Mortenson of the Pacific Centre for Isotopic and Geochemical Research at the University of British Columbia provided unexpected isotopic age dates under cooperation with the Geological Survey of Canada CCGK project on Triassic time. We thank Robert Blodgett, Norm Silberling, Marc Hendrix, Don Winston, Michael Hofmann, and Steve Sheriff for helpful geologic discussions and Mike Orchard, Chris McRoberts, Robert Blodgett, Heinz Kozur, Leo Krystyn, and Norm Silberling for generous paleontologic discussions and fossil identification. Sid Ash, Mike Sandy, and John Utting all provided additional interest in fossils from the area. Mike Orchard, Norm Silberling, Marc Hendrix, and Eric Edlund provided constructive reviews of the manuscript. Roy Pescador of the University of Montana Electron Microscopy Facility provided electron microscopy services and resources with support by grant RR-16455-01 from the National Center for Research Resources (Biomedical Research Infrastructure Network program), National Institutes of Health. Andrew Caruthers, Craig Dugas, Adam Bender, and Tim Wheeler provided field and laboratory assistance. REFERENCES CITED Atwood, W.W., 1912, Some Triassic fossils from southeastern Alaska: The Journal of Geology, v. 20, no. 7, p. 653–655. Babcock, L.C., 1976, Conodont paleoecology of the Lamar Limestone (Permian), Delaware Basin, West Texas, in Barnes, C.R., ed., Conodont paleoecology: Geological Association of Canada Special Paper 15, p. 279–294. Behnken, F.H., 1975, Leonardian and Guadalupian (Permian) conodont biostratigraphy in western and southwestern United States: Journal of Paleontology, v. 49, no. 2, p. 284–315. Berg, H.C., 1973, Geology of Gravina Island, Alaska: U.S. Geological Survey Bulletin 1373, 44 p. Berg, H.C., 1981, Upper Triassic volcanogenic massive-sulfide metallogenic province identified in southeastern Alaska, in Albert, N.R., and Hudson, T., eds., The United States Geological Survey in Alaska: accomplishments during 1979: U.S. Geological Survey Circular 0823-B, p. 104–108. Berg, H.C., Jones, D.L., and Coney, P.J., 1978, Map showing pre-Cenozoic tectonostratigraphic terranes of southeastern Alaska and adjacent areas: U.S. Geological Survey Open-File Report 78-1085, 2 sheets. Berg, H.C., Jones, D.L., and Richter, D.H., 1972, Gravina-Nutzotin Belt— Tectonic significance of an upper Mesozoic sedimentary and volcanic sequence in southern and southeastern Alaska, in Geological Survey Research 1972: U.S. Geological Survey Professional Paper 800-D, p. 1–24.
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The Geological Society of America Special Paper 442 2008
Stratigraphy of the Triassic Martin Bridge Formation, Wallowa terrane: Stratigraphy and depositional setting George D. Stanley* Jr. Department of Geosciences, The University of Montana, Missoula, Montana 59812, USA Christopher A. McRoberts* Department of Geology, State University of New York at Cortland, Cortland, New York 13045, USA Michael T. Whalen* Department of Geology and Geophysics, University of Alaska Fairbanks, Fairbanks, Alaska, 99775-5780, USA ABSTRACT The Upper Triassic (Carnian–Norian) Martin Bridge Formation of northeastern Oregon, southeastern Washington, and western Idaho is characterized by rapidly shifting depositional processes within a tropical volcanic island arc setting. Martin Bridge sequences in the Hells Canyon and northern Wallowa Mountains document shallow-water peritidal evaporitic sediments that are succeeded by deeper and predominantly subtidal deposits. This indicates drowning of the carbonate platform and a transition to deeper-water turbiditic sedimentation before a late Triassic transition into the overlying mid-Norian to Jurassic Hurwal Formation. At the type locality in the southern Wallowa Mountains, dysaerobic shales, carbonate debris sheets, and turbiditic sediments indicate distal slope and basinal environments while other facies at other sites in the Wallowa Mountains and Hells Canyon areas indicate reef and shallow-water platform settings. In this paper we formally recognize the name Martin Bridge Formation and reinstate the type locality in the southern Wallowa Mountains as the principal unit stratotype. An additional reference section is given at Hurricane Creek in the northern Wallowa Mountains. The Martin Bridge is formally divided into four members: the Eagle Creek and Summit Point Members are introduced and formally proposed herein and the BC Creek and Scotch Creek Members also are elevated to formal status. A partial reconstruction of the Wallowa terrane during deposition of the Martin Bridge Formation suggests a north-south (or northeast-southwest) trending platform margin facing a forearc basin situated to the east (or southeast). The lithofacies and paleontological characteristics of the Martin Bridge can be put into the framework of a depositional and a tectonic model to help better explain many of the stratigraphic
*E-mails: Stanley:
[email protected]; McRoberts:
[email protected]; Whalen:
[email protected]. Stanley, G.D., Jr., McRoberts, C.A., and Whalen, M.T., 2008, Stratigraphy of the Triassic Martin Bridge Formation, Wallowa terrane: Stratigraphy and depositional setting, in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 227–250, doi: 10.1130/2008.442(12). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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Stanley et al. and paleontologic problems previously encountered. We believe that the Wallowa terrane provides one of the best and most complete examples yet known for shallowwater carbonate depositional patterns in an oceanic island arc setting. Keywords: Triassic, Wallowa, Oregon, stratigraphy, paleontology.
INTRODUCTION The Martin Bridge Formation is the only conspicuous Triassic limestone unit exposed in northeastern Oregon and adjacent Idaho. It is part of a thicker late Paleozoic to early Mesozoic volcanic and sedimentary succession called the Wallowa terrane (Silberling and Jones, 1984). The Wallowa terrane is one volcanic island arc and related sedimentary rock assemblage separated from another island arc, the Olds Ferry terrane (Brooks and Vallier, 1978), by the intervening Baker terrane, which represents a subduction mélange complex (Mullen, 1985). Together these tectonostratigraphic terranes constitute part of the Blue Mountains Region of northeast Oregon, southeast Washington, and western Idaho (Fig. 1). In late Mesozoic time, these terranes amalgamated and by Cretaceous time were accreted to the North American craton. They share few stratigraphic relationships with coeval Permian and Triassic rocks on the craton (Dickinson and Thayer, 1978; Vallier and Brooks, 1986) but have been compared with Wrangellia and other island arc terranes in the North American Cordillera (Jones et al., 1977; Mortimer, 1986). This tectonic model and the related rock types are quite different from those developed for cratonal sequences and are more comparable to modern-day western Pacific island arcs and associated oceanic crust (Brooks and Vallier, 1978; Follo, 1992).
Figure 1. Generalized map showing three principal terranes in the Blue Mountains Region. Cross-hatch—island arc; stipple—subduction mélange.
Figure 2. Index map showing the outcrop pattern of the Martin Bridge Formation and the various regions discussed in the text (modified from Brooks and Vallier, 1978). 1—southern Wallowa Mountains; 2—northern Wallowa Mountains; 3—Hells Canyon and Seven Devils Mountains.
The Wallowa terrane is now known to represent an Early Permian to Late Triassic volcanic island arc mantled by a cover of Mesozoic sediment. Initiation of a tropical carbonate platform, represented by the Martin Bridge Formation, began in Late Triassic (late Carnian) time. By the early Norian, a steep-sided, shallow-water platform with protected lagoons and a shelf margin of carbonate sand shoals and coral-sponge-algal patch reefs developed. The platform-to-basin transition is marked downslope by an abrupt facies change to coarse-grained gravity-flow breccia, conglomerate and other slope deposits, which laterally grade into deep-water starved-basin facies (Follo, 1992). This transition from platform to slope and basinal facies is one of the best documented examples of a platform-basin transition in any Cordilleran terrane. Emplacement of the Wallowa batholith and subsidiary satellites during Late Jurassic and Early Cretaceous time (Armstrong et al., 1977) altered the country rocks and obscured many stratigraphic relationships. Early Mesozoic tectonic events deformed
Stratigraphy of the Triassic Martin Bridge Formation, Wallowa terrane pre-batholith rocks (Nolf, 1966; Follo, 1986; Mirkin, 1986) and Tertiary uplifts resulted in erosion of much of the early Mesozoic rocks. Finally, the Neogene eruption of Columbia River basalt covered and deeply buried much of the Wallowa terrane and restricted exposures to three main regions: (1) the southern Wallowa Mountains, near Halfway, Oregon, (2) the northern Wallowa Mountains, near Enterprise, Oregon, and (3) portions of Hells Canyon and Seven Devils Mountains on the OregonIdaho border where the Snake River cuts through the overlying Columbia River basalt (Fig. 2). Other areas where the Wallowa terrane is exposed include metamorphosed rocks west of Riggins and exposures at Pittsburg Landing and adjacent areas, near the Washington, Oregon, Idaho borders (Fig. 2). The Martin Bridge Formation is a well-known and distinctive unit in the succession of the Wallowa terrane and it remains one of the best studied and dated early Mesozoic shallow-water sequences in North America. Ranging from the Carnian through Norian stages of the Upper Triassic, intervals within the Martin Bridge display a relatively complete succession containing ammonoids, conodonts, and flat clams belonging to the genus Halobia. The carbonate rocks have been compared with other Carnian–Norian carbonate sequences such as those found in Wrangellia (Jones et. al., 1977). This paper focuses on the Martin Bridge and synthesizes biostratigraphic and sedimentologic data to improve correlations and better interpret its history and paleogeography. The principal areas to be discussed are presented in Figure 3. Our purpose is to reconcile some of the stratigraphic problems, formalize some of the units, and offer a concept of a Martin Bridge stratotype that is both intellectually acceptable in the context of depositional processes and operational in the sense of a practical stratigraphy.
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calcareous algae, sponges, spongiomorphs, corals, and bivalves (Newton, 1986; Stanley, 1986; Stanley and Senowbari-Daryan, 1986; Newton et al., 1987; Senowbari-Daryan and Stanley, 1988; Stanley and Whalen, 1989), together with the thick carbonate rocks of the Wallowa terrane, provides evidence of a tropical setting. Evidence of carbonate deposition and tropical marine fauna is corroborated by paleomagnetic results indicating Triassic paleolatitudes of 18°–24° (±4°) north or south of the equator (Hillhouse et al., 1982; May and Butler, 1986). Although Newton (1983) and Malmquist (1991) favored a paleoposition in the Southern Hemisphere, Stanley and Vallier (1992) supported a location in the Northern Hemisphere on the basis of Permian paleolatitudes derived from paleomagnetic investigations of Harbert et al. (1988). The Late Triassic paleolongitudinal positions of the tropical Wallowa terrane relative to North America are even more equivocal. Studies of early Norian silicified faunas from Hells Canyon yield conflicting results. Shallow-water bivalves, for example, indicate that the Wallowa terrane was in the eastern Pacific and close enough to the craton to allow exchange with cratonal faunas (Newton, 1987, 1988; Malmquist, 1991), but sponges, scleractinian corals, and spongiomorphs revealed some endemic elements with no links to the craton and strong Tethyan connections, suggesting that a substantial body of ocean lay between the Wallowa terrane and the North American craton (Senowbari-Daryan and Stanley, 1988; Stanley and Whalen, 1989; Stanley and Yancey, 1990; Stanley and Vallier, 1992). Follo (1992, p. 1572) believed that an eastern Pacific site for Wallowa would make it
TECTONOSTRATIGRAPHIC SETTING The Martin Bridge Formation is part of the ~8-km-thick Lower Permian to Upper Jurassic sequence of volcanic and sedimentary rocks of the island arc referred to as the Wallowa terrane. It is best exposed in northeastern Oregon, western Idaho, and southeastern Washington (Fig. 2). Although the outcrops of the Martin Bridge are small in extent and scattered, it is important to remember that the Wallowa terrane was not a single volcanic edifice but a series of volcanic island groups within an island arc. Furthermore it is not an isolated terrane but is complexly associated with four other tectonostratigraphic terranes of the Blue Mountains Province (Fig. 1). These terranes amalgamated during the Jurassic, prior to Cretaceous accretion to the North American craton (Brooks and Vallier, 1978; Silberling and Jones, 1984). The Martin Bridge Formation contains diverse carbonate and argillaceous rock types that represent patch reefs, platform shoals, restricted peritidal basins, lagoons, slope deposits, and basinal rocks. All of these were deposited within an island arc setting following the abrupt cessation of volcanism (Stanley, 1986; Stanley and Senowbari-Daryan, 1986; Whalen, 1988; Follo, 1994). The diverse shallow-water invertebrate fauna of
Figure 3. Locality map showing outcrop pattern (stippled) of Martin Bridge Formation and metamorphosed equivalents (modified from Follo, 1986). Localities discussed in text: 1a and 1b—principal stratotype at Eagle and Paddy Creeks; 2—Summit Point Member stratotype; 3—Torchlight Gulch; 4—East Eagle Creek near Bradley Mine; 5—reference section at Hurricane Creek; 6—BC Creek stratotype at Chief Joseph Mountain; 7—Scotch Creek Member stratotype; 8—Black Marble Quarry; 9—Spring Creek; 10—Kinney Creek .
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geometrically difficult to account for the large amounts of northward displacement and extensive oblique convergence indicated for the Wallowa terrane between Triassic time and accretion to North America at the end of the Early Cretaceous. On the other hand, May and Butler (1986) suggest there has been little or no northward displacement of Wallowa in relation to the craton since the Late Triassic, but during this time, North America had been moving northward from Pangea. Compared with other terranes in Canada and Alaska, the Wallowa terrane is a relatively small tectonic fragment and it has been suggested that it might correlate with larger terranes farther north. To what larger parent might the orphaned Wallowa terrane belong? The Wallowa terrane was originally interpreted as a southern extension of Wrangellia on the basis of paleolatitude and similarity of the general stratigraphic successions (Jones et al., 1977). Subsequently, a range of volcanic rock types and paleontological evidence based on silicified Late Triassic corals has been used to recognize a distinction between the Wallowa terrane and Wrangellia (Sarewitz, 1983; Mortimer, 1986; Stanley, 1986; Whalen, 1988). Studies of some silicified gastropods from the Wallowa terrane, on the other hand (Blodgett et al., 2001; Frýda et al., 2003; Nützel et al., 2003), show some endemic taxa and paleogeographic linkage with Alaskan Wrangellia as well as parautochthonous rocks in Peru but indicate less similarity with most other terranes of North America. By Middle Jurassic (Bajocian) time, coral and bivalve faunas from the Wallowa terrane show stronger links to the Western Interior Embayment of the craton (Stanley and Beauvais, 1990). In conjunction with paleobotanical findings (Ash, 1991a, 1991b), these data suggest that by the Middle Jurassic the Wallowa terrane was closer to North America and in its northward journey, moving out of the tropics into temperate paleolatitudes (Stanley and Beauvais, 1990; White et al., 1992). CORRELATIONAL AND NOMENCLATURAL PROBLEMS Since the first description of the Martin Bridge Formation, conceptual and nomenclatural misunderstandings have plagued stratigraphers. Without designating a stratigraphic name, Smith (1912) described a partial stratigraphic sequence of Upper Triassic limestone, limy shales, and associated invertebrate fossils near the confluence of Paddy and Eagle Creeks at Martin’s Bridge in the southern Wallowa Mountains. Smith (1912) discussed these fossils in the context of reef development and recognized the presence of corals, the flat bivalve Halobia, and ammonoids similar to taxa from central Europe. Later Smith (1927) described a more detailed stratigraphic section containing Carnian to Norian fossils. The term “Martin Bridge” first appeared in an International Geological Congress guidebook by Chaney (1932) and was used not only for limestone but also for a variety of rock types including shale, basalt, andesite, and tuff. At nearly the same time, Gilluly et al. (1933, p. 12, citing work in preparation by C.P. Ross), without defining a type section, designated
the Martin Bridge in the Wallowa Mountains as “1000 to 3000 feet of limestone, limy shale, and interbedded volcanic rocks of Upper Triassic age.” Subsequently, Ross (1938, p. 32) used the name Martin Bridge Formation for rocks in the southern Wallowa Mountains. He described characteristic rock types and well-preserved and abundant fossils, and designated a type section at Smith’s (1912, 1927) locality near Martin’s Bridge (a bridge once existing near the confluence of Eagle and Paddy Creeks). Because Smith (1912) did not designate a name, Ross (1938) must be regarded as the original author for the Martin Bridge Formation. With little reference to the section in the southern Wallowa Mountains, Smith and Allen (1941, p. 10) defined the Martin Bridge on the basis of stratigraphic sections at Hurricane Creek and the Upper Imnaha River drainage in the northern Wallowa Mountains as 200–2000 ft of “grey to black, crystalline limestone, which toward the top is both intercalated with and grades into the argillaceous Hurwal Formation.” Their definition and scope of the Martin Bridge clearly establishes it as a prominent rock type composed of limestone and marble that overlies the informally named “Lower Sedimentary Series.” The latter is a structurally and petrographically complex volcanic and sedimentary unit regarded as part of the Clover Creek Formation by Nolf (1966) and subsequent workers. According to Smith and Allen (1941), the Martin Bridge is identified as the first prominent limestone unit above the “Lower Sedimentary Series.” As Nolf (1966, p. 56) pointed out, a strict application of this definition of the Martin Bridge would be impossible to apply in the southern Wallowa Mountains, where several prominent limestone beds occur in both the Martin Bridge and the overlying Hurwal Formations. Other inconsistencies have occurred in the nomenclature and concept of this formation. Hamilton (1963) introduced the name “Martin Bridge Limestone” for metamorphosed limestone exposed in the Riggins region of western Idaho. Following Hamilton’s lead, Vallier (1977) also used the name “Martin Bridge Limestone” to describe thick limestone and dolomite exposed in Hells Canyon at a stratigraphic section measured just south of Kinney Creek near Hells Canyon Dam. It seems clear that Hamilton (1963) attributed the term “limestone” to the Martin Bridge only in an informal sense and his usage followed Smith and Allen’s (1941) concept of Martin Bridge solely as a carbonate unit. This, as well as other cited examples, is inconsistent with the rule of priority (Article 7c, North American Stratigraphic Code) because the original stratotype in the southern Wallowa Mountains (Ross, 1938) was disregarded. References to the unit “Martin Bridge Limestone” in the northern and southern Wallowa Mountains and Hells Canyon (e.g., Newton, 1986; Stanley, 1986; Vallier and Brooks, 1986; Whalen, 1988; Follo, 1992) perpetuated the problem. Despite a reexamination of the originally designated stratotype of the Martin Bridge Formation on Eagle Creek (McRoberts, 1993), misuse under the name “Martin Bridge Limestone” continued (Follo, 1994; White, 1994; White and Vallier, 1994). The proliferation of informal, and often conflicting, stratigraphic names poses
Stratigraphy of the Triassic Martin Bridge Formation, Wallowa terrane further problems. Many of the subunits designated within the Martin Bridge (Nolf, 1966; Follo, 1992, 1994) either are informal or, in the case of Nolf (1966), are proposed in a thesis and therefore not sanctioned by the North American Stratigraphic Code (North American Commission on Stratigraphic Nomenclature, 2005, Article 4). These inconsistencies result from improper use of stratigraphic terminology, but they also reflect the complex nature of depositional patterns and processes in an island arc setting where rapid vertical and lateral facies changes occurred. STRATIGRAPHY AND REDEFINITION OF THE MARTIN BRIDGE FORMATION The Martin Bridge Formation conformably overlies Upper Triassic (Carnian) volcanic and volcaniclastic strata (Fig. 3) informally designated the “Lower Sedimentary Series” (Smith and Allen, 1941), the Clover Creek Formation in the northern Wallowa Mountains (Nolf, 1966), or the Doyle Creek Formation (Seven Devils Group) of Hells Canyon (Vallier, 1977). Where observed in both the northern and southern Wallowa Mountains (Follo, 1994), the contact is gradational. The Martin Bridge is conformably overlain by and/or may grade laterally into the Hurwal Formation (Smith and Allen, 1941; Vallier, 1977; Follo, 1992). Formational boundaries between the Hurwal, Martin Bridge, and underlying volcanic and volcaniclastic rocks are not exposed at the Martin Bridge stratotype. They are known from other localities in the southern and northern Wallowa Mountains and Hells Canyon, however. The Martin Bridge is discussed below at its better occurrences in three principal regions: (1) southern Wallowa Mountains, (2) northern Wallowa Mountains, and (3) Hells Canyon (Fig. 3). Erection of a Composite Stratotype A composite stratotype (North American Commission on Stratigraphic Nomenclature, 1983, p. 853) is necessary for the Martin Bridge Formation because the present reference section does not adequately represent the diversity of Martin Bridge rock types present throughout the Wallowa terrane. Follo (1994, p. 7) discussed the stratigraphic nomenclature and presented some informal lithofacies subdivisions of the Martin Bridge, but he emphasized that formal designation was outside the scope of his work. In redefining the Martin Bridge Formation, we propose retention of both the original name—Martin Bridge Formation— and the original type section as one of the principal reference sections. Furthermore, we formally propose four members defined by reference stratotypes within the northern and southern Wallowa Mountains (Fig. 3). Generalized columnar sections are presented in Figure 4, which is keyed to the map in Figure 3. Southern Wallowa Mountains Excellent exposures of the Martin Bridge Formation can be found in the southern Wallowa Mountains (Table 1) but many outcrops, like the one at Summit Point (Fig. 3, site 2), are more or less
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isolated and preserved as windows in the Columbia River Basalt. Significant structural deformation in the southern Wallowa Mountains makes exact stratigraphic relationships uncertain. Locally, the Martin Bridge is deformed by numerous thrust faults of unknown displacement, and by innumerable high-angle normal faults, and in certain areas is tightly to isoclinally folded (Vallier, 1977; Mirkin, 1986; McRoberts, 1990; Follo, 1994). Structurally repeated or omitted strata are common, as illustrated by the carbonate “ridge bed,” which is found to be repeated several times in the primary stratotype at Eagle Creek (McRoberts, 1990, 1993). Structural deformations in the southern Wallowa Mountains, following Late Triassic–Early Jurassic amalgamation of the island arcs, appear to be related to their accretion to the continent and to emplacement of the Wallowa Batholith (Late Jurassic–Early Cretaceous) and later during the initial phases of Basin and Range extension (Mirkin, 1986). We designate the principal reference section of the greater composite stratotype of the Martin Bridge Formation, in the southern Wallowa Mountains at Eagle Creek near Ross’s original site. This reference section also serves as the stratotype for the proposed Eagle Creek Member. A second reference section is designated at Summit Point in the southern Wallowa Mountains. Eagle Creek Member The Eagle Creek Member is here designated to include alternating calcareous shale, calcareous mudstone, and well-bedded limestone as well as bioclastic and lithoclastic rudstones exposed in the Wallowa-Whitman National Forest, along the Eagle Creek drainage (Fig. 3). The lowermost 100 m of the Eagle Creek Member is exposed along Paddy Creek (Fig. 3; Table 1) and the remainder of the member, comprising ~125 m of the original type section (Figs. 4 and 5A; Table 1) of the Martin Bridge Formation, is exposed along Eagle Creek (Smith, 1927; Ross, 1938; McRoberts, 1990, 1993). It showed that post-Triassic structural deformation, including at least three low-angle thrust faults of unknown displacement and nearly 20 high-angle faults of limited (<5 m) displacement, disrupted the sequence at the stratotype. In spite of the structural complexities, occurrences of abundant biostratigraphic fossils allowed the stratigraphic reconstruction of the principal stratotype by rearrangement of five structural blocks into an upper Carnian to lower Norian sequence (McRoberts and Stanley, 1991; McRoberts, 1993). Sedimentologic studies of the unit stratotype by Follo (1992, 1994) and stratigraphic investigations by McRoberts (1990, 1993) indicated that the Eagle Creek Member is primarily composed of finely laminated, slightly calcareous, organic-rich shale, dark mudstone, and impure limestone (Eagle Creek facies A of Follo, 1994). Common invertebrates are ammonoids and halobiids as well as rarer shallow-water bivalves (McRoberts, 1992, 1993; Tamura and McRoberts, 1993). The shales of the Eagle Creek Member are interpreted as representing normal background sedimentation on a carbonate slope environment (McRoberts, 1990; Follo, 1994), rather than shallow-water backreef lagoonal muds as suggested by Prostka (1962). This deeperwater interpretation is substantiated by the fining-upwards in
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Figure 4. General lithostratigraphic and chronostratigraphic correlations of the Martin Bridge and other units within the Wallowa terrane. Abbreviations: hu—Hurwal Formation; eg—Excelsior Gulch conglomerate (Follo, 1992); sp—Summit Point Reef (Stanley and SenowbariDaryan, 1986); ec—Eagle Creek locality and type section (McRoberts, 1993); lss—“Lower Sedimentary Series; cc—Clover Creek greenstone; dlu—Deadman Lake unit; sc—Scotch Creek facies; bc—BC Creek facies (Nolf, 1966); dc—Doyle Creek Formation (Nolf, 1966; Vallier, 1977); wsc—Wild Sheep Creek Formation (Vallier, 1977). Stratigraphic sections: 1a and 1b—the principal stratotype; 2—Summit Point; 3—Hurricane Creek; 4—BC Creek.
TABLE 1. GEOGRAPHIC AND STRATIGRAPHIC DETAILS FOR LOCALITIES No. Locality name Member Geographic reference 1a Eagle Creek Eagle Creek Confluence of Eagle and Paddy Creeks 1b Paddy Creek Eagle Creek Along Forest Road 360, east side of Paddy Creek 2 Summit Point Summit Point Southwest from the top of Summit Point 3 Torchlight Gulch Summit Point Along Forest Road 075, south side of Torchlight Gulch 4 East Eagle Creek Scotch Creek Near the Bradley Mine 5 Hurricane Creek Scotch Creek West side of Hurricane Creek 6 BC Creek BC Creek North Fork of Creek Between 6000ƍ and 6400ƍ on the south fork of Scotch Creek 7 Scotch Creek Scotch Creek Along Murray Creek at 6800ƍ 8 Black Marble Quarry unknown 9 Spring Creek Scotch Creek West side of the Snake River 10 Kinney Creek Scotch Creek North side of Kinney Creek *S—stratotype section; R—reference section; F—fossil locality.
millimeter- to centimeter-thick, carbonate-rich and carbonatepoor couplets (Fig. 5B). These most likely represent the distal fringes of carbonate turbidites. Well-bedded limestone also alternates with calcareous shale in the Eagle Creek Member (Eagle Creek facies B of Follo, 1994). Limestone beds are characterized by coarse-grained, normally graded grainstone and packstone including bioclasts of corals, spongiomorphs, and other shallowwater fossils. The normally graded carbonates of the Eagle Creek
Type* S S S R R R S S F F R
Member are consistent with allochthonous deposition as gravity or turbiditic flows from shallow-water carbonates to the north. Sporadic occurrences of beds, up to 2 m thick, of lumachelle limestone (bioclastic rudstone) form a minor constituent of the Eagle Creek Member (McRoberts, 1993). These monospecific Halobia lumachelle beds are bounded both above and below by organic-rich mudstones and black (dysoxic) shales. Such thick shell beds may represent extremely condensed intervals or
Stratigraphy of the Triassic Martin Bridge Formation, Wallowa terrane protracted population blooms of halobiid bivalves in response to a relaxation of oxygen deficient conditions. Regardless of their genesis, these beds may serve as event horizons locally correlatable in the southern Wallowa Mountains. A distinctive 10-m-thick lithoclastic and bioclastic conglomerate unit, exposed on prominent ridges above Eagle Creek (McRoberts, 1990; Follo, 1994), forms a conspicuous lithofacies of the Eagle Creek Member. A majority of the bioclasts within this conglomerate bed are corals and spongiomorphs, which may have led Smith (1912, 1927) to misinterpret this unit as a coral reef. The rounded nature of the conglomerate clasts within this bed implies a history of erosion and reworking on a shallow shelf prior to transportation into the basin as debris flows. Although Follo (1992, 1994) interpreted this conglomerate unit (his Eagle Creek facies C) to represent numerous gravity-flow debris sheets, field mapping, biostratigraphic relationships, and overall similarity of these debris beds at the Martin Bridge type locality show that they probably represent a single bed that has been repeated by several high-angle thrust faults (McRoberts, 1990). Unfortunately the upper and lower contacts of the Eagle Creek Member are not exposed at the unit stratotype. At many places in the southern Wallowa Mountains the lower contact of the Eagle Creek Member of the Martin Bridge and the Lower Sedimentary Series is unconformable (Ross, 1938), but it may be conformable or slightly diachronous along Paddy Creek (Prostka, 1962; Mirkin, 1986; Follo, 1994) where minor thrust faults occur between competent limestone and incompetent shale (Follo, 1994). Unlike Prostka (1962), who defined the lower contact of the Martin Bridge as the first massive limestone overlying the calcareous argillite of the Lower Sedimentary Series, we regard the boundary as intercalated and gradational and define the transition between the two units as the first calcareous interval, either calcareous mudstone or pure limestone, above the non-calcareous or slightly calcareous green and gray fissile argillite of the Lower Sedimentary Series. The upper contact of the Eagle Creek Member with the Hurwal Formation has not been identified at the stratotype in the southern Wallowa Mountains. Ross (1938) mentioned the presence of rocks overlying the Martin Bridge but never identified them with any particular unit. Prostka (1962) assigned poorly exposed rocks of the southern Wallowa Mountains to the Hurwal but relationships to the Martin Bridge were ambiguous. The uncertain nature of the contact in the southern Wallowa Mountains is highly problematical because there is little lithologic similarity between the dark argillites referred to the Hurwal and the uppermost calcareous argillites occurring at the Martin Bridge stratotype. Smith (1912, 1927), who first studied the faunal sequence of the Martin Bridge type locality, recognized Halobia oregonensis, H. salinarum, H. dilatata, and H. halorica as well as some Tethyan coral species. In the absence of ammonoids from the Martin Bridge Formation, Smith placed the Carnian-Norian boundary within 121 m of barren shale and limestone between the last occurrence of H. oregonensis and an overlying “Coral Zone” of lower Norian age. Subsequently, Stanley (1979, 1986)
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Figure 5. Field photos of principal stratotype of Martin Bridge Formation. (A) Outcrop at the confluence of Eagle and Paddy Creeks; note dashed line represents approximate Carnian-Norian stage boundary based on halobiid bivalves. (B) Carbonate-rich laminated distal turbidite showing soft-sediment structures.
noted a lower Norian coral assemblage. Orr (1986) described the ichthyosaur Shastasaurus, and Gruber (in Kristan-Tollmann and Tollmann, 1983) reported Carnian halobiids (H. rugosa and H. radiata radiata) from structurally complex outcrops near the type section. Work by McRoberts and Stanley (1991) and McRoberts (1993) reveals a rather complete sequence across the CarnianNorian boundary with diverse assemblages of halobiid bivalves and ammonoids. In ascending order, the mudstones and bedded carbonates of the lower part of the Martin Bridge contain the bivalve Halobia superba superba and the ammonoids Discotropites sp., Arietoceltites sp., and possibly Anatropites sp. Together
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Figure 6. Summit Point locality in the southern Wallowa Mountains. (A) Base of section showing exposure of massive limestone of the Summit Point Member, which can be traced up the slope of Summit Point. (B) Coral framestone, Retiophyllia. Surface view of a colony in situ within the massive limestone of the Summit Point Member.
these fossils are indicative of a late Carnian age (Dilleri–Macrolobatus Zones). About 2.7 m above the highest occurrence of H. superba superba, H. beyrichi is found in several horizons, indicating a lower (Kerri Zone) Norian age and thus indicating the proximity of the Carnian-Norian boundary. Abundant shell beds containing the middle Norian bivalve H. halorica occur in the
Summit Point Member A second reference section of the Martin Bridge Formation is located 8 km northeast of the Eagle Creek type locality on the hillside of Summit Point in the Wallowa-Whitman National Forest (Figs. 3 and 4; Table 1). The section, described in more detail below, is designated the Summit Point Member and it has been published by Stanley and Senowbari-Daryan (1986). The Summit Point Member is exposed as a small outlier at a second site at Torchlight Gulch, (Fig. 3, site 3; Table 1). However, this site is much smaller and presents no contacts with any adjacent units. Extensive limestone exposed on the hillside along the southwest slopes of Summit Point (Fig. 6A) display in depositional contact both the Summit Point Member and the overlying Scotch Creek Member (Fig. 7). The best exposures occur from Twin Bridge Creek up the slope. The section begins in massive limestone of the Summit Point Member at the base of a thrust fault. The Summit Point Member is overlain by bedded bioclastic carbonate rocks of the Scotch Creek Member (Fig. 7) that continue northeast up toward the Summit Point lookout tower. The top of the section presents an unconformity with the Columbia River Basalt. The Hurwal Formation is not present at this site because of pre– Columbia River Basalt erosion, which was extensive and varied over the extent of the Wallowa terrane. At this Summit Point site, the Summit Point Member is 33–35 m thick. The overlying Scotch Creek Member (Fig. 7) is difficult to measure because of repeated sections due to thrust faults, but it could reach as much as 150–200 m in thickness. At this locality, the Summit Point Member is quite distinct from the exposures on the section along Eagle Creek because it consists entirely of massive to thick bedded, light-colored to medium-gray limestone. The limestone dips ~45° to the southwest. The stratigraphy is complicated by small thrust faults that offset and duplicate this section of carbonate rocks, causing repetition of the Summit Point Member at both higher and lower elevation along the slope. A large thrust fault, more or less parallel to bedding, is clearly present at the base of the section along a northwest-trending forest road near Twin Bridges Creek. A stratigraphic section of the Summit Point and Scotch Creek Members in Figure 7 is described in more detail below. Framework-building and binding fossils characterize the Summit Point Member (unit 1). These include the phaceloid-dendroid coral Retiophyllia (Fig. 6B), a solitary to pseudocolonial coral Distichophyllia, the thalamid sponge Salzburgia, spongiomorphs, tabulozoans, and masses of the red alga Solenopora. The light-gray limestone contains local breccia and mostly massive limestone. Cavities within the limestone may be lined with biogenic cyanobacterial (porostromate algae) crusts. The base of the Summit Point stratotype, where the Summit Point Member is well exposed, begins along the forest road on
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Figure 7. Columnar section of the limestone at Summit Point showing the various members of the Martin Bridge at this site. The section begins above a thrust fault at the base of the section and is overlain by the Columbia River Basalt. The exact thickness of unit 6 is not known and is most likely much thicker than indicated. CRB—Columbia River Basalt.
the southwest side of Summit Point at a topographic map elevation of 5550 ft and continues up the slope toward the fire lookout tower to a topographic map elevation of 6580 ft. At the lower level of the forest road, the massive limestone begins above a low-angle thrust fault. Below the thrust fault lies tectonically sheared calcareous shale characterized by Halobia and some bedded limestone containing bivalves and gastropods. The Summit Point Member (unit 1, Fig. 7) consists of 33–35 m of massive, unbedded light-gray limestone chiefly dominated by in situ solitary and colonial scleractinian corals, chambered sponges, and red algae. Where well exposed, the limestone presents the fabric of in situ reef framework. It begins on the forest road and continues eastward up the slope of the hill to Summit Point (Fig. 6A). This massive limestone is succeeded by medium-bedded units of coarse-grained bioclastic limestone (units 2–5) of the Scotch Creek Member. These are characterized by solitary and colonial corals, gastropods, crinoid ossicles and echinoid plates, and bivalves. Some beds are rich in oncolites. Unit 4 (Fig. 7) is a marker bed composed of large oncolites, 2–5 cm in diameter, the nuclei of which are coral fragments, crinoid ossicles, and lowspired gastropods. Bedded carbonate lithologies within the upper part of the member are characterized by breccia, dissolution cavi-
ties, and small fissures infilled with fine-grained, yellow-colored, often laminated sediment. Small in situ patch reefs 1–2 m high and 5–6 m long occur within this interval. Channels between the patches are indicated by skeletal debris, breccia, and conglomerate. The top of the section (unit 6) merges into fine- to mediumgrained, dark-gray, recrystallized bioclastic limestone that weathers almost white in color. These rocks attain dips of up to 60° west. Because of the pervasive recrystallization that increases upward in the section, the only discernable fossils are vague indications of bioclasts, gastropods, and bivalves. The carbonate sequence at Summit Point contrasts with that at Eagle Creek in displaying massive and mostly unbedded, pure carbonate rock with a predominance of framestone and bindstone. It contains an abundance of reef-related fossils characteristic also of the Dachstein Reef Limestone of the Northern Calcareous Alps in central Europe (Stanley and Senowbari-Daryan, 1986). Conglomerate and breccia textures also are present in the overlying Scotch Creek Member, most of which contains bioclastic beds and some low-angle crossbeds. A rich and diverse fauna of corals, sponges, tabulozoans, bivalves, gastropods, crinoid ossicles, spongiomorphs, red algae, benthic foraminifers, and cidaroid echinoids characterizes
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the Scotch Creek Member at Summit Point. Many of these are preserved as bioclasts (Fig. 8). Among these are large thickshelled megalodontid bivalves (Fig. 9). This is the first report of these reef-adapted bivalves in the Wallowa terrane, but they are very common to Upper Triassic reef complexes of the Tethys (Flügel, 2002) and also are known from carbonate rocks from Upper Triassic (Norian) reef complexes in the Stikine terrane (Reid, 1988). The Summit Point Member is interpreted to be part of a reef complex very similar in biota, size, and composition to those from the thicker Dachstein Reef Limestone of the Northern Calcareous Alps of Austria and southern Germany (Zankl, 1969; Stanley and Senowbari-Daryan, 1986). Norian Dachstein reefs in the Alps were interpreted as having formed on a broad ramp rather than being shelf-edge buildups, which characterize some later reef complexes of the Rhaetian (Stanton and Flügel, 1989). Age-diagnostic megafossils are absent from the Summit
Figure 9. Two different examples of megalodontid bivalve facies (A and B) in massive unbedded limestone. These thick-shelled, gregarious bivalves are photographed in outcrop in the Summit Point Member, Summit Point, southern Wallowa Mountains. One cent coin for scale.
Figure 8. Field photo of unsorted bioclasts consisting mostly of colonial corals (large flat clasts), bivalve shells, sponges, crinoid ossicles and echinoid spines, and other shallow-water biota in the Scotch Creek Member, Summit Point, southern Wallowa Mountains. Scale in centimeters.
Point Member but the megalodont bivalves, corals, sponges, red algae, and foraminifers are well known from most NorianRhaetian reef complexes in the Northern Calcareous Alps. Coral, sponge, and spongiomorph taxa occurring in conglomerate beds of the Eagle Creek Member appear to be similar, if not identical, to taxa present in the Summit Point Member.
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Conodonts retrieved from this sequence at Summit Point currently are under study with Shunxin Zhang and M.J. Orchard, and preliminary findings indicate late Carnian to early middle Norian ages. Locally, the contact between the Summit Point Member and underlying shale and limestone is a thrust fault. The strata below the thrust fault are reminiscent of the Eagle Creek Member but because of the structural complications, depositional relationships are unclear. A geologic map produced by Follo (1986) suggests a possibility that the Summit Point Member and overlying Scotch Creek Member might be an integral block thrust over the Eagle Creek Member, but structural evidence in support of this idea would require more extensive fieldwork. The faunal composition of the Summit Point Member is similar to fossils found in conglomerate clasts of Eagle Creek Member at the stratotype, and they are unlike the fauna at the Black Marble Quarry in the northern Wallowa Mountains (see below). Northern Wallowa Mountains The northern Wallowa Mountains, located near the town of Enterprise, Oregon, display impressive and rugged exposures of early Mesozoic sedimentary and volcanigenic rocks subjected to the forceful intrusion of granitic rocks related to the Wallowa Batholith (Fig. 10A). The impressive exposures in the northern Wallowa Mountains are situated some 50 km northeast of the Martin Bridge stratotype on Eagle Creek (Fig. 3; Table 1). This part of the Wallowa Mountains displays some of the thickest exposures of the Martin Bridge Formation. A thickness of ~350 m was estimated for the Martin Bridge at the west end of Hurricane Creek as measured from the base where it contacts the underlying Triassic volcaniclastic and volcanic rock below to the contact with the overlying Hurwal Formation. However, folding and ductal flow, resulting from the intense local metamorphism and batholith implacement that characterizes this region, has produced locally apparent thicknesses of over 1000 m (Follo, 1994). Follo (1994) estimated a more realistic thickness for the Martin Bridge of 350–450 m from exposures west and southwest of Enterprise. The strata cannot be physically traced between these two regions because of the intrusion of the Wallowa Batholith and extrusion of the extensive Columbia River Basalt, which cover much of the exposures of early Mesozoic rocks. The stratigraphy and fossils of this region were discussed by Smith and Allen (1941), Stanley (1979), and Follo (1994), but the most detailed treatment was given in a doctoral thesis by Nolf (1966) who informally described three members of the Martin Bridge—the Hurricane Creek, BC Creek, and Scotch Creek. We here elevate two of Nolf’s units, the BC Creek and Scotch Creek Members, to formal status as members within the Martin Bridge Formation. We here designate the section measured by Nolf along Hurricane Creek (Fig. 11) as a reference section for the composite stratotype of the Martin Bridge Formation.
Figure 10. (A) Exposure of the Martin Bridge and associated rocks in the northern Wallowa Mountains. Photo is of the succession near Joseph, Oregon, looking from Mount Howard toward BC Basin and BC Creek (center) and Chief Joseph Mountain (highest point to the right of view). Near the base of the section and halfway up the slope, dark-colored Triassic volcanic and volcaniclastic rocks are succeeded upward by lighter-colored limestone of the Martin Bridge Formation, BC Creek Member (stratotype). Above the tree line on the right, light-colored rock belongs to the Scotch Creek Member. The top ridge (darker) is Hurwal Divide. To the left of BC Basin, light-colored granitic rocks of the Wallowa Batholith are intruded by horizontal feeder dikes. (B) Typical lithology of Scotch Creek Member from south side of Hurricane Creek. Hand specimen contains poorly sorted bivalves, gastropods, and echinoderm fragments. Scale in millimeters.
BC Creek Member This member was informally named and described by Nolf (1966) for 150 m of fine-grained, well-bedded limestone directly overlying volcanic rocks of the Clover Creek Greenstone of the Seven Devils Group. At the stratotype along BC Creek in BC Basin, it is sublithographic and pale yellow-brown, light gray or light brownish gray, or distinctly pink colored, owing to iron
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Figure 11. Detailed section of Martin Bridge Limestone and adjacent units in the Northern Wallowa Mountains as measured by Bruce Nolf along the west side of Hurricane Creek. Index locator map inset at bottom right. Modified from Nolf (1966); here his Hurricane Creek Member is included as a metamorphosed part of the BC Creek Member, the lowest unit of the Martin Bridge Formation.
oxide. Some beds of the BC Creek are finely cross-bedded. Rare fossils consist of echinoid spines and plates and bivalve shells. This unit was redescribed by Follo (1994) as his BC Creek unit. We designate the composite section Nolf (1966) measured and described at the head of the north fork BC Creek between the 8800 and 9450 foot topographic elevation lines (Fig. 3; Table 1) as the stratotype for the BC Creek Member of the Martin
Bridge Formation. This member is exposed on Chief Joseph Mountain also. The BC Creek Member consists of thin-bedded (5.0–25.0 cm) laminated and cross-bedded limestone. Laminae are silty or dolomitic with irregular bedding planes and stylolites. Rip-up clasts and nodular limestone are also common in beds 50 cm to 2 m thick. Replacement chert and “chicken-wire” textures are common.
Stratigraphy of the Triassic Martin Bridge Formation, Wallowa terrane Follo (1994) also noted salt casts, karstic solution collapse structures, and evidence of nodular anhydrite. This member is represented also at the base of the thick limestone section exposed in Hells Canyon near Spring Creek (Whalen, 1988). The Hurricane Creek Member is interpreted to be laterally equivalent to the BC Creek Member (Nolf, 1966). According to Nolf, no fossils and few sedimentary features are known from the marble-like limestone of the Hurricane Creek Member. Nolf’s reef interpretation for the Hurricane Creek Member was based solely on the massive nature of this member, its thickening, and its apparent lateral transition into the bedded BC Creek Member. Field examination of stratigraphic relationships between the Hurricane Creek Member and the BC Creek Member at Nolf’s original sections prompted our reinterpretation. Both Nolf’s (1966) original view of a reef and Follo’s subsequent (1992, 1994) interpretation of the Hurricane Creek as a high-energy barrier are, in our opinion, erroneous. Field relationships suggest to us that the marble lithology of Nolf’s Hurricane Creek Member resulted primarily from contact metamorphism from the intrusion of the Wallowa Batholith. Thus the apparent thickening and lateral transition of the massive Hurricane Creek into the bedded BC Creek Member is mostly a function of decreasing metamorphism away from the batholith. We therefore do not advocate use of the Hurricane Creek Member as a formal stratigraphic member. On the basis of this apparent stratigraphic relationship, sparse fossil content, and evaporitic textures including nodular anhydrites, Nolf interpreted the BC Creek Member as a backreef, evaporitic lagoon facies. These data accord well with the presence of dolomitic textures, cryptalgal laminites, and solution collapse breccias as noted by Follo (1994). Limestone and dolomite exposed at the base of the Martin Bridge section in Hells Canyon contain similar features (Whalen, 1988) and also belong to this member (see below). Scotch Creek Member The Scotch Creek Member was first established by Nolf (1966) for bedded calcarenites exposed at an elevation between 6000 and 6400 ft on the south fork of Scotch Creek. It was described also on the west side of Hurricane Creek. The 150–250-m-thick Scotch Creek Member overlies the BC Creek Member. We retain the original section of Nolf (1966) along Scotch Creek as the stratotype. The unit is well exposed also along the west side of Hurricane Creek where, in its upper part, it is gradational with overlying shales and argillites of the Hurwal Formation. We designate the exposures along the west side of Hurricane Creek as a reference section of the Scotch Creek Member. This member also is present in the southern Wallowa Mountains at Summit Point and in Hells Canyon (Fig. 3; Table 1) where it is exposed at Kinney Creek (Vallier, 1977). The Scotch Creek Member at its stratotype is mostly coarsegrained, poorly sorted bioclastic limestone but also includes abundant silt-size and finer siliciclastic material (Follo, 1994). Texturally, the limestone of this member is mostly grainstone to packstone, but coarser-grained carbonate conglomerate and
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breccia also are present. Beds range from 1.0 cm to 1.0 m or more in thickness. The bioclastic beds are poorly sorted, rich in fossil debris (Stanley, 1982, Fig. 10A) and are normally to inversely graded with channel scouring and slump features commonly present. Some of these beds resemble calcareous tempestites similar to those reported in the Martin Bridge Formation in Hells Canyon (Newton, 1986; Whalen, 1988) but others appear to be slope-related gravity-flow deposits. Shallow-water fossils reported by Nolf (1966) include abundant but fragmentary and abraded bivalves Septocardia and Parallelodon, gastropods Naticopsis, Cirras, and Eucyclis, as well as fragments and debris of corals, spongiomorphs, tabulozoans, echinoid spines, and ammonoids (Fig. 10B). The fauna and bioclastic textures in Scotch Creek carbonate rocks in the northern Wallowa Mountains appear similar to counterparts previously discussed from the Summit Point site. The Scotch Creek Member is interpreted to represent downslope carbonate deposition in slightly deeper water. This member records the initial phases in drowning of the shallow carbonate platform. Slope instability is indicated by the frequent occurrence of debris and gravity-flows associated with turbidite sedimentation. As in the southern Wallowa Mountains, the transition from carbonate platform to basin is marked by a progressive decrease in limestone and increase in mudstone and shale. Nolf (1966) and Follo (1994) correctly interpreted the Scotch Creek to represent drowning of the carbonate platform. Voluminous carbonate debris, including large clasts of corals and other reef organisms, were transported downslope along a carbonate ramp. Whether the slope was characterized by a reef rim is unclear. However, the only evidence for an in situ reef occurs in the southern Wallowa Mountains at Summit Point and at Torchlight Gulch. The voluminous coral and sponge debris characterizing the Scotch Creek in the northern Wallowa Mountains most likely was derived from patch reefs similar in lithologies to those discussed from Summit Point. However, with the notable exception of the Black Marble Quarry (discussed below), no such reef facies have been discovered in the northern Wallowa Mountains. The age of the Scotch Creek is based on the occurrence in this member of ammonoids identified by N.J. Silberling. These include Juvavites sp. and Mojsisovicites kerri (Nolf, 1966) and they indicate the lower Norian Kerri Zone. The Scotch Creek Member exposed on the west side of Hurricane Creek yielded the ammonoids Arcestes sp. indet. and Tropiceltites columbianus, again establishing an early Norian age (Nolf, 1966). The carbonate textures and fossils all correspond closely to some of the limestone exposed at Hells Canyon (see below). Martin Bridge at Hells Canyon The Martin Bridge has been exposed by dissection deep into the Columbia River Basalt in the gorge of Hells Canyon (Fig. 3). It is well exposed and has been mapped on both sides of Hells Canyon by Vallier (1977). The exposure is structurally complex, with tight to isoclinal folds on the west side of the
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canyon that open into a broad synform to the east (Fig. 12) where stratigraphic relationships are best deciphered. The depositional history of the Martin Bridge at this locality has been interpreted by Whalen (1988). Approximately 330 m of bedded limestone is present in the Hells Canyon section. In Hells Canyon, Martin Bridge sediments were deposited atop the volcaniclastic Doyle Creek Formation of the Seven Devils Group (Fig. 12). The Doyle Creek is equivalent to the Clover Creek unit in other parts of the Wallowa terrane. The basal contact of the Martin Bridge is not well exposed in Hells Canyon but intercalation of thin beds of volcaniclastic siltstone in the lower 25 m (Fig. 13) indicates a gradational contact with the underlying Doyle Creek Formation (Whalen, 1988). Similar intercalated relationships were observed near the Martin Bridge unit stratotype in the southern Wallowa Mountains (McRoberts, 1993; Follo,
Figure 12. Field photo on the west side of Hells Canyon showing the dark volcanic rocks of the Clover Creek Greenstone of the Seven Devils Group below and the bedded limestone of the Martin Bridge Formation above.
Figure 13. Field photo from the west side of Hells Canyon showing the contact of the Martin Bridge Formation with the Clover Creek Greenstone (dashed line). The bedded carbonate rocks begin with the BC Creek Member. Unlike Figure 12 which shows some structural dislocation, the contact between the two units appears conformable.
1994). The onset of pure marine carbonate deposition indicates cessation of volcanic activity and a relatively abrupt shutdown of terrigenous sediment supply (Whalen, 1988). Detailed facies analysis of the Martin Bridge Limestone (Whalen, 1988) reveals a suite of shallow-water platform carbonates deposited as an overall-deepening upward sequence. The basal Martin Bridge facies (60 m thick) includes laminated and fenestral dolostones with gypsum casts and algal laminated intervals interbedded with mudstones and peloid wackestones. Supratidal to shallow intertidal conditions are indicated by algal laminated dolomudstones, gypsum casts, and fenestral textures. Interbedded mudstones and peloid wackestones were deposited in relatively quiet water peritidal settings. We assign and correlate this lower portion of the Martin Bridge in Hells Canyon to the BC Creek Member. Peritidal facies are overlain by bioclast, intraclast, peloid wackestone and packstone, bioclast or ooid grainstone, and spongiomorph bafflestone (~150 m thick). The normal marine fauna and coarse-grained packstone and grainstone indicate deposition in shallow to deep subtidal, moderate- to high-energy environments. The uppermost portion of the Martin Bridge (120 m thick) in Hells Canyon is a series of 5–20-m-thick fining-upward subtidal cycles, which probably indicate minor relative sea-level variations from above to below fair-weather wave base (Whalen, 1988). This upper member is well exposed on the west side of Hells Canyon at Spring Creek and contains beds of bioclastic limestone remarkably similar in thickness and faunal composition to those from the Scotch Creek stratotype in the northern Wallowa Mountains. Because silification is better at Spring Creek, tempestite beds here have yielded a wealth of fossils consisting of bivalves, corals, spongiomorphs, and sponges (Newton et al., 1987; Senowbari-Daryan and Stanley, 1988; Stanley and Whalen, 1989). More recently, gastropods have been described (Blodgett et al., 2001; Frýda et al., 2003; Nützel et al., 2003). The upper portion of the limestone at Hells Canyon is similar to the Scotch Creek Member but does not contain any appreciable siliciclastic material or any indication of slope deposition. It thus represents platform interior facies correlative with the Scotch Creek Member in the Wallowa Mountains and we designate the exposure along Kinney Creek (Fig. 3; Table 1), following Vallier (1967), as a reference section. This section well illustrates the level-bottom facies model for this member. The upper contact of the Martin Bridge in Hells Canyon is either a modern erosional surface or an older unconformity overlain by Tertiary Columbia River Basalt (Vallier, 1977; Whalen, 1988). Biostratigraphically useful fossils from the Hells Canyon Region are rather scarce and restricted to only a few occurrences of ammonoids and undescribed halobiid bivalves. Recently conodonts have been retrieved from beds adjacent to the silicified Spring Creek site. They represent the late Carnian–early Norian Primitius and Communisti Zones (Nützel et al., 2003), thus reinforcing the ages derived from megafossils. Stanley (1986) and Newton et al. (1987) described silicified bivalve molluscs and coral-spongiomorph assemblages at Spring Creek. This is a site
Stratigraphy of the Triassic Martin Bridge Formation, Wallowa terrane first described by Vallier (1967) and designated U.S. Geological Survey (USGS) locality M2672. It is assigned to the early Norian Kerri Zone on the basis of ammonoids, including the diagnostic Tropiceltites cf. T. columbianus (identified by N.J. Silberling and cited in Vallier, 1977, p. 49). An early Norian age is consistent with the occurrence of the late Carnian to early Norian bivalve Halobia austriaca found in float nearly 20 m below M2672 by McRoberts (unpublished data). The bivalves reported by Newton et al. (1987) are temporally wide-ranging, and include both Carnian and Norian taxa widely correlative with those of the Tethys. The corals, however, do not include Carnian species but rather relate to those occurring in Norian–Rhaetian rocks of the Alps. The sponge Amblysiphonella cf. A. steinmanni, reported from Hells Canyon, is known from the Rhaetian Zlambach Formation as are most of the coral taxa. Other “Martin Bridge” Occurrences Several isolated outcrops of limestone have been attributed to the Martin Bridge Formation but neither dating nor correlations are well established. Three problematical localities are discussed below. Black Marble Quarry Another possible occurrence of the Martin Bridge Formation is in a quarry in the eastern slopes of the northern Wallowa Mountains, northwest of Enterprise, Oregon (Fig. 3, loc. 8). This locality is referred to as the Black Marble Quarry (Smith and Allen, 1941), formerly mined by the Black Marble Lime Company. Cut into the hillside of a wooded area between Sheep Ridge and Ruby Peak, it presents a fresh exposure of dark limestone (Fig. 14A). As described by Nolf (1966) and later Stanley (1979), the limestone is thick bedded and dark gray to black in color. It appears to be carbonaceous although subsequent analysis of total organic content yielded carbon percentages ranging from 0.86 to 1.18 and not unlike values for average light-colored limestone of the Martin Bridge in Hells Canyon (0.52%–1.52%). The limestone of the Black Marble Quarry is rich in reefal fossils such as molluscs, corals, and spongiomorphs as well as large alatoform bivalves (Fig. 14B). The quarry is located near the upper mapped extent of the Martin Bridge (Fig. 15) but unfortunately it is covered by the Tertiary Columbia River Basalt so that relationships with adjacent units are not visible. However, the quarry is situated slightly higher than the mapped exposure of the Martin Bridge, which could perhaps place it in the overlying Hurwal. The isolation makes the quarry stratigraphically anomalous and difficult to correlate with adjacent units. Interestingly, the fauna does not resemble any reported from reef limestone in the southern Wallowa Mountains. Stanley (1979) illustrated some of the reefal fossils from the Black Marble Quarry and reported foraminifers, chambered sponges, colonial and solitary corals, spongiomorphs, crinoid ossicles, and echinoids. Large branching spongiomorphs (probable hydrozoan colonies) are distinctive and abundant (Fig. 16).
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Thin sections revealed predominantly fine-grained micritic to peloid textures (mudstone, packstone) as well as abundant in situ coral and spongiomorph framestone. The fauna indicates a shallow-water, low-energy back-reef environment characterized by small thickets of branching corals, spongiomorphs, and chambered sponges. Smith and Allen (1941) assigned the Black Marble Quarry to the Martin Bridge Formation and reported ammonoids Juvavites sp. and Sagenites herbicho Mojsisovics (identified by S.W. Muller) from and near the quarry. Many years of subsequent work failed to locate such ammonites and it appears that the reported specimens were collected as float in areas surrounding the quarry rather than from the quarry itself. Although conodonts or diagnostic ammonoids have not been recovered, the fauna of the quarry clearly suggests a Norian age. Nolf (1966) presented two hypotheses to explain the isolated Black Marble Quarry: (1) as a unit of the Martin Bridge moved structurally by normal fault to a higher stratigraphic position, or (2) as an exotic slide block, displaced into the Hurwal basin much later in the Norian. Stanley (1979) pointed out that the faunal composition and lithology were atypical of the Martin Bridge but resembled some dark limestone already known from parts of the overlying Hurwal, notably as clasts within the Deadman Lake unit (Fig. 11), an observation also made by Nolf (1966) who favored the idea that the quarry might be a large olistholith or slide block. Unpublished data on involutinid foraminifers recovered from limestone in the quarry revealed a diverse assemblage including Aulotortus communis Kristan and A. tumidus Kristan-Tollmann. These taxa are characteristic of middle to late Norian in their occurrences in Alpine reef sequences. Giant alatoform bivalves, up to 1 m across (Fig. 14B) (Yancey and Stanley, 1999), occur in the quarry. These authors designated the bivalves as Family Wallowaconchidae, and they also have been reported from sites and terranes very distant from the Wallowa Mountains (Fig. 17). Identical or closely related species of wallowaconchids also occur in Norian limestone in Sonora, Mexico and in the Norian rocks of the Chulitna terrane and Stikinia (Yancey and Stanley, 1999) and similar bivalves also have been reported from Norian rocks of the Tethys (Yancey et al., 2005). Lewiston, Idaho Upper Triassic carbonate rocks have long been known from a limestone quarry on the Lapwai Indian Reservation located on the east side of Mission Creek near Lewiston, Idaho (Cooper, 1942; Haas, 1953). Silicified fossils were collected as limestone blocks by Norman D. Newell. Squires (1956) described a Norian silicified coral fauna and Hoover (1991) described brachiopods. However, no authors have assigned the limestone to the Martin Bridge. The Lapwai quarry lies in a remote position some 50 km northeast of known exposures of the Martin Bridge. The abandoned quarry is exposed on a hillside completely surrounded by Tertiary Columbia River Basalt. Thus its stratigraphic relationship with any overlying and underlying units cannot be established.
Figure 14. The Black Marble Quarry. (A) Photo of a portion of the Black Marble Quarry in the northern Wallowa Mountains. Note the darkcolored limestone is bedded and more or less horizontal. (B) Large alatoform bivalves of the family Wallowaconchidae in life position as exposed in the quarry wall. These bivalves (Yancey and Stanley, 1999) have wing-like septate extensions from the main shell. Two individuals are shown with the wing-like extensions overlapping. The commissure between the valves was vertical.
Figure 15. Generalized geologic map of the Martin Bridge Formation exposed in the northern Wallowa Mountains. The Black Marble Quarry is indicated. Upthrown (U) and downthrown (D) blocks are shown. Modified from Nolf (1966).
Stratigraphy of the Triassic Martin Bridge Formation, Wallowa terrane The quarry consists of thick- to medium-bedded recrystallized limestone. Investigations by Stanley (1979, 1986) and Whalen (1985) showed at least 150 m of thick-bedded recrystallized limestone, which yielded a diverse silicified fauna of sponges, corals, “tabulozoan” brachiopods, bivalves and gastropods, and three new echinoid taxa. Although correlation with the Martin Bridge at Hells Canyon and in the northern Wallowa Mountains would seem logical, the silicified corals show greater resemblance to a silicified fauna in medium- to thick-bedded limestone from the Sutton Formation dated as Rhaetian (Crickmayi Zone) on Vancouver Island (Caruthers and Stanley, 2008), This is part of the greater Wrangellia succession. A rich gastropod fauna from the Lapwai quarry, reported by Haas (1953, p. 310), has been described by Nützel and Erwin (2004) who reported over 60 species and similarity with some Wallowa gastropods. Nützel and Erwin (2004) illustrated from the Lapwai quarry the ammonoid Gnomohalorites cordilleranus, indicative of early late Norian (Sevatian) time, thus confirming the late Norian age suggested by Stanley (1979). Such dates, however, would suggest more appropriate correlations with the Hurwal Formation rather than the Martin Bridge (Fig. 4). Carbonate facies at the Lapwai quarry include mudstones, bioturbated peloid and bioclast wackestone, peloid-intraclast packstone, bioclast grainstone, peloid grainstone, and coralspongiomorph bafflestone (Whalen, 1985). Coral-spongiomorph bafflestone units occur as lenses or thickets 1–5 m in diameter and 1–2 m thick. Mudstone and wackestone facies represent deposition under relatively deep subtidal conditions below fair-weather wave base, whereas packstones and grainstones were probably deposited above fair-weather wave base (Whalen, 1985). Broken and rotated coral colonies (Squires, 1956) are evidence of strong wave or current activity. Although pervasive reef structure is not indicated, the fossil associations and microfacies do suggest that a patch reef-type environment with numerous thickets of corals and spongiomorphs must have existed. Halobia has not been reported from the Lapwai quarry and many invertebrate species do not appear similar to those of the Martin Bridge Formation (Stanley, 1986). At present we can assign this limestone to neither the Martin Bridge nor the Hurwal. Perhaps the best approach is to refer to the limestone of the Lapwai quarry as “the Mission Creek Limestone” (Nützel and Erwin, 2004). Furthermore, we cannot exclude the possibility that the Lapwai quarry might represent a fragment of another terrane, separate from the Wallowa terrane . Other Occurrences? Two other localities of potential Martin Bridge rock types occur in the Seven Devils Mountains. They include the metamorphosed marble, phyllite, schist, and related rocks of the Riggins area (Hamilton, 1963) and a locality near the mouth of the Grande Ronde River (Vallier, 1977, p. 48). The Riggins locality, near the town of Lucile, Idaho (10–11 mi north of Rig-
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gins), is near the suture zone of the Wallowa terrane (Fig. 2). Unfortunately it is too intensely sheared and metamorphosed to be compared profitably with other sections discussed above. The Grande Ronde locality has not been studied in any detail,
Figure 16. Light-colored branching spongiomorphs (recrystallized) and some smaller coral fragments stand out from the darker matrix in a fallen block in the Black Marble Quarry. Scale in centimeters.
Figure 17. Generalized terrane map showing sites in four terranes of western North America that have yielded Norian wallowaconchid bivalves. Shaded area is the North American craton, and line of teeth marks indicates position of fold and thrust belt.
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but according to Follo (1994) it is intensely metamorphosed and structurally complicated and thus not useful for stratigraphic study. Another site at the mouth of Cottonwood Creek (Goldstrand, 1994, p. 1439) near the Triassic-Jurassic unconformity may contain the lower part of the Martin Bridge. Martin Bridge–Type Limestone Occurrences in the Hurwal Formation The post–Martin Bridge history of the Wallowa terrane is revealed in conglomerate and breccia, which occur in the overlying deeper-water Hurwal Formation. Nolf (1966) cited huge chaotic breccia deposits of the Deadman Lake unit (Fig. 11) that contain limestone blocks, some of which reach up to 300 m across. The Black Marble quarry may lie within one of these clasts. Reef corals, spongiomorphs, and other fossils of Norian age contained within some limestone clasts of this unit are similar to those of the Martin Bridge. Follo (1992) believed the deposit was an olistostromal megabreccia derived by downslope slump and/or debris flow from a source area located somewhere north of the outcrop area. The limestone clasts within the Deadman Lake unit most likely came from an extensive, but as yet unknown, carbonate complex in existence during post–early Norian time after the Martin Bridge of the northern Wallowa Mountains had already been drowned (Nolf, 1966). We believe, however, that like the Deadman Lake unit, reef blocks within this Hurwal unit could also have been derived from another Upper Triassic carbonate platform complex within the Wallowa terrane, lying to the north or the northwest. The limestone at Mission Creek, Idaho, could be part of that carbonate platform. A land source is indicated in the rounded nature of the clasts within the Excelsior Gulch, which implies “fluvial transport or wave reworking in a coastal environment prior to redeposition in deeper water” (Follo, 1992, p. 1568). The Excelsior Gulch Conglomerate (Fig. 4) is another distinctive unit in the Hurwal Formation of the southern Wallowa Mountains containing limestone clasts (Follo, 1992). Dated as middle Norian or younger, it contains a much more abundant and diverse reef fauna with lithofacies and fossils nearly identical to those of the Summit Point Member, but it also yields chert clasts like those from the Baker terrane and unknown from the Wallowa terrane. The Excelsior Gulch conglomerate also produced rare clasts containing abundant dasycladacean algae, which are unknown from the Martin Bridge (Flügel et al., 1989). Follo (1992) recognized the clasts and noted the differences between the Excelsior Gulch and the Deadman Lake unit. He suggested the nearby Baker terrane as the source of those clasts. However, it seems equally plausible, owing to the similarity of the reefal fossils in the Excelsior Gulch clasts to those from the Summit Point Member, that the abundant reef limestone clasts could also have been derived from uplifted middle to late Norian carbonate reef limestone within the Wallowa terrane.
DEPOSITIONAL HISTORY OF THE WALLOWA TERRANE The Martin Bridge Formation records the relatively abrupt onset of carbonate sedimentation following cessation of volcanic activity and a decrease in output of volcaniclastic sediment. Vallier (1995) interpreted this as marking a change in plate vector with the consequent firing up of volcanism in the Olds Ferry terrane (Huntington Island Arc). This terrane, adjacent to the Wallowa terrane, is characterized by volcanic and mixed carbonate and volcaniclastic rocks and linked by provenance to the Wallowa terrane (see LaMaskin, this volume). If so, the Wallowa terrane may have become forearc to the Olds Ferry terrane. Eventually the Wallowa terrane amalgamated with other terranes in the Blue Mountains and by Cretaceous time had collided with the North American craton. Relative to other Triassic sequences in the North American Cordillera, limestone deposits of the Martin Bridge are not as thick as those found in other terranes, notably in the Alaskan portion of Wrangellia, but they nevertheless record a whole suite of typical shallow- to deeper-water carbonate facies. These include reef development, downslope debris, and deeper-water, black-shale environments. With the aid of diagnostic fossils, various members and lithologies within the Martin Bridge Formation may be correlated and their depositional history reconstructed. The often abrupt lateral facies changes from one rock type to another are actually predicted in models of carbonate sequences deposited on subsiding volcanic oceanic islands (Soja, 1996), and patterns differ greatly from those from cratonal settings. Paleogeography and Tectonic History Paleomagnetic findings confirm a tropical oceanic setting for the Wallowa terrane at ~18° north or south of the paleoequator in Middle and Late Triassic time (Hillhouse et al., 1982). Early Permian data dictate a northern hemisphere position (24° ± 12°) and possible subsequent southward transport relative to North America (Harbert et al., 1988). The present geographic relationships and paleopole data on the Wallowa terrane relative to the North American craton confirm that the Wallowa terrane has not experienced any major relative latitudinal shifts since the Triassic, although both craton and Wallowa terrane, like those of many other Cordilleran terranes, moved northward. Exposures of shallow- and deeper-water Martin Bridge facies and members in the northern and southern Wallowa Mountains respectively (Follo, 1992), indicate a northeast-southwest platform to basin transition. However, rotation has affected the Wallowa terrane since deposition of the early Mesozoic rocks. The Wallowa terrane is assumed to have experienced ~65° of clockwise rotation following Mesozoic accretion (Wilson and Cox, 1980). Restoration of this tectonic rotation would produce a roughly west to east platform-to-basin transition between early and middle Norian time as depicted in cartoon form (Fig. 18). Follo (1992) interpreted slope and basinal Martin Bridge facies
Stratigraphy of the Triassic Martin Bridge Formation, Wallowa terrane
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Figure 18. Geotectonic-lithofacies “cartoon” for the Wallowa terrane during (A) early Norian and (B) late Norian time (after White et al., 1992). Subduction zone and polarity of subduction indicated by teeth marks. Letters indicate areas discussed in text. Note: Map area has been restored to its Triassic position (unrotated 65° counter clockwise) following Hillhouse et al. (1982), and for simplification, the Wallowa terrane is treated as one island. Question marks indicate uncertainties. By Middle Norian time the terrane is closer to the subduction zone and the reef and lagoon sediments have vanished, being replaced with deeper-water facies as platform drowning occurs.
as having been deposited along the western margin of a forearc basin with accretionary wedge deposits of the Baker Terrane overlying a west-dipping subduction zone. Depositional History of the Martin Bridge The stratigraphic problems discussed earlier illustrate the nature of facies changes within a volcanic island arc setting and the resulting stratigraphic problems in classification and physical lithostratigraphic correlation. Combined localities of the Martin Bridge reveal a late Carnian to middle Norian carbonate sequence that begins with peritidal sedimentation and progressively deepens upward. Although frequently considered an “island” there is actually little, if any, evidence for subaerial exposure or land surfaces after the initial evaporitic deposition characterizing the base of the Martin Bridge. There is no evidence of soil profiles, caliche, root casts, or terrestrial deposits, and most of the Norian sequence indicates the presence of a single (or a series of) shallowly submerged platform(s) where carbonate sedimentation kept pace. By late Norian time, subsidence finally outpaced sedimentation. This resulted in the drowning of the platform (Whalen, 1988) as indicated by the transition to deeper-water platform and slope deposits (Scotch Creek Member) that both overlie and interfinger laterally with platform and reef deposits of the BC Creek and Summit Point Members (Fig. 4). This drowning sequence and an overall relative rise in sea level are clearly seen in the Martin Bridge at Hells Canyon. Following cessation of volcanism and volcaniclastic deposits, carbonate deposition was initially under supratidal, then peritidal,
and finally shallow subtidal conditions. The peritidal interval (60 m) is about one-third thinner than the sequence of similar facies in the northern Wallowa Mountains (Whalen, 1988; Follo, 1992), indicating a more distal depositional and deeper paleogeographic setting. Peritidal deposits in Hells Canyon were quickly replaced by deeper subtidal facies (bioclast, intraclast, peloid wackestones and packstones, spongiomorph bafflestones) as relative sea level rose. Bioclast and ooid grainstones were deposited in open-shelf shoal environments and they are interbedded with subtidal facies. In the Summit Point site the reef facies of the Summit Point Member is overlain by the Scotch Creek Member (Figs. 4 and 7), which represents a deeper-water slope facies. This change would represent a response to the subsiding nature of the platform margin. The sequence in the northern Wallowa Mountains records the deposition of open shelf subtidal deposits and carbonate sand shoals thought to be laterally equivalent to algal laminated and replaced evaporite-bearing peritidal limestones (Follo, 1986, 1994). The Martin Bridge rocks in the northern Wallowa Mountains thus represent a lateral facies change that can only be seen as a vertically stacked sequence in Hells Canyon. Peritidal and open shelf facies are overlain by silty bioclastic carbonate grainstones with graded beds, scour and fill structures, and slump folds, which are interbedded with calcareous argillites that Follo (1986, 1994) interpreted as slope deposits implying a deepeningupward sequence. The Martin Bridge is overlain by even deeper water facies of the Hurwal Formation, but in the northern Wallowa Mountains lateral lithofacies relationships between the Martin Bridge and the Hurwal are not exposed (Follo, 1992).
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Both in the southern Wallowa Mountains (Smith and Allen, 1941; Nolf, 1966) and near Riggins, Idaho (Lund et al., 1983), the Martin Bridge is overlain and intercalated with rocks that indicate a deeper-water origin. The volcaniclastic and calcareous argillites, graywackes, and conglomerates of the Hurwal Formation (Smith and Allen, 1941; Nolf, 1966) and metamorphosed equivalents in the Lucile Slate (Lund et al., 1983) indicate a platform-to-basin transition as well as drowning of the Martin Bridge carbonate platform (Whalen, 1988). Tropical platform carbonates have the potential to outpace most relative rises in sea level (Schlager, 1981) and several factors may have contributed to platform drowning, including (1) rapid rates of relative sea-level rise, (2) increased volcaniclastic input, (3) northward tectonic transport out of the tropics, and (4) lack of a significant reefal rim. Whalen (1988) mentioned the possibilities for points 1 and 4 playing a role. These characteristics are shown in Figure 19.
Late Triassic eustatic sea level is poorly understood, but most localities imply a Norian lowstand (Vail et al., 1977; Haq et al., 1987; Hallam, 1992). The general deepening-upward trend of the Martin Bridge–Hurwal sequence is thus interpreted to represent thermal subsidence of the island arc (Whalen, 1988). High rates of short-term sea-level rise can cause platform drowning (Schlager, 1981) and may have contributed to the lack of a reefal rim around the Martin Bridge platform (Whalen, 1988). High rates of sea-level rise, however, may not have been necessary to drown the Martin Bridge platform, because of the synergistic effects of arc subsidence, volcaniclastic input, northward transport, and lack of a persistent reef rim. Clastic sediment input and change in environmental conditions and nutrient influx are all accepted causes of carbonate platform drowning. Facies of the Hurwal Formation provide clear evidence of increased clastic input. If a northern hemisphere paleolatitude for the Wallowa
Figure 19. Schematic model of deposition for various members of the Martin Bridge Formation. (A) Shoreline (subaerial); (B) evaporitic and peritidal sediments of the BC Creek Member; (C) shallow water, more normal salinity, lagoonal sediments, and patch reefs of the Scotch Creek and Summit Point Members; (D) downslope debris beds, including olistostromal slide blocks of the slope that occur within the sequence at Eagle Creek as well as at other sites in the Wallowa Mountains; (E) deeper-water basin including the dysaerobic black mud environment of the Eagle Creek Member.
Stratigraphy of the Triassic Martin Bridge Formation, Wallowa terrane terrane is accepted (Harbert et al., 1988), then progressive northward transport during Late Triassic time may have caused environmental stress, and cooling may have contributed to the demise of the platform and drowning, similar to the “Darwin point” concept postulated for the Hawaiian-Emperor seamount chain (Grigg, 1982). CONCLUSIONS We suggest that the Wallowa terrane provides one of the best and most complete examples in the Triassic of North America for shallow-water carbonate depositional patterns in an oceanic island arc setting. The stratigraphic problems in both the nomenclature and classification of the Martin Bridge can be attributed to the complex variety of subenvironments and expected depositional patterns within an island arc setting. Detailed stratigraphic, sedimentologic, and paleontologic studies have helped resolve correlation problems and allow reconstruction of depositional patterns within the island arc setting. A better understanding of the depositional settings allows us to establish a composite stratotype with a number of reference sections and four members to better account for variations in rock types within the stratigraphic framework of the island arc. A depositional model for the Martin Bridge takes into account the development in such a setting, where one would expect rather sudden facies changes from deeper dysaerobic mud to downslope accumulations to shallow
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platform (patch reef and lagoon) facies. The various members of the Martin Bridge can best be understood in the context of such a model (Fig. 20). We propose retention of the name Martin Bridge Formation defined by a composite section. This composite section includes the original type section of the Martin Bridge Formation in the southern Wallowa Mountains as well as other sections (Fig. 4). We include a supplemental reference section from Hurricane Creek in the northern Wallowa Mountains near Enterprise and we formally designate four members of the Martin Bridge Formation from stratotypes in both the northern and southern Wallowa Mountains. These are the Eagle Creek and Summit Point Members in the southern Wallowa Mountains, and the BC Creek and the Scotch Creek Members from northern Wallowa Mountains and Hells Canyon. Units of the Martin Bridge Formation record rapidly shifting facies and depositional patterns within the framework of a tropical volcanic island arc. The proposed stratigraphic nomenclature and stratotypes reflect and better account for the complex depositional patterns in an island arc setting. They record the initial establishment of shallow-water, peritidal, evaporitic conditions after an abrupt cessation of volcanic activity. Facies of the Martin Bridge Formation indicate that the platform initially kept pace with subsidence but was eventually drowned, as indicated by the deepening-upward succession. Shoal and reef margin facies grade laterally and vertically into deeper-water platform and slope deposits, which in turn were transitional into
Figure 20. Generalized “cartoon” model for the Martin Bridge lithofacies and members showing sediment types and biotic compositions. In this model one can appreciate how lithofacies and biotic types can laterally change rapidly in the island arc setting of the Wallowa terrane. Also note that the reef belt (best developed on the windward side) is much narrower relative to the more extensive lagoon.
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basinal dysaerobic black-shale and mudstone facies. The end of the Martin Bridge deposition was characterized by the drowning of the carbonate platform with a transition to deeper-water turbiditic sedimentation before a Late Triassic (middle Norian) transition into the overlying Hurwal Formation. The Hurwal Formation contains some carbonate intervals, but most of the rock types (shales, spicular mudstones, and turbidite deposits) are clearly of deeper-water origin and indicate drowning of the platform. Chaotic carbonate breccias and conglomerate deposits in the Hurwal record yet another phase of younger carbonate platform development later in middle to late Norian time somewhere outside of the present outcrop area. Conglomerate and breccia, referred to as the Deadman Lake and Excelsior Gulch units in the northern and southern Wallowa Mountains respectively (Nolf, 1966; Follo, 1992), show an abundance of corals, sponges, spongiomorphs, and a variety of other reef organisms within limestone clasts ranging from a few centimeters to tens or perhaps as much as 100 m in length. Follo (1992) interpreted limestone clasts of the Deadman Lake unit to have been derived from a northern source whereas limestone and volcanic rock fragments and chert of the Excelsior Gulch unit were derived from a uplifted southern source (Baker terrane?). Although the exact source of the two distinctive chaotic deposits and their relationships to each other are unclear, they are important in suggesting the presence of a subaerial land source and the existence of younger (middle–late Norian?) carbonate platforms from which reef debris was shed downslope into the adjacent basin. ACKNOWLEDGMENTS We wish to thank N.J. Silberling and T.L. Vallier for providing valuable insight toward our understanding of Triassic stratigraphy and paleontology of the Wallowa-Hells Canyon Region and for reviewing early drafts of the manuscript. Thanks to J. Scallan for improving an earlier draft. This research was in part funded by National Science Foundation grants EAR-8916664 and EAR-9624501 to Stanley and grants to McRoberts and Whalen from the University of Montana, which supported fieldwork and comparative studies in the Wallowa terrane. Whalen also acknowledges a grant from Sigma Xi which supported the fieldwork. REFERENCES CITED Armstrong, R.L., Taubeneck, W.H., and Hales, P.O., 1977, Rb-Sr and P-Ar geochronometry of Mesozoic granitic rocks and their Sr isotopic composition, Oregon, Washington, and Idaho: Geological Society of America Bulletin, v. 88, p. 397–411, doi: 10.1130/0016-7606(1977)88<397: RAKGOM>2.0.CO;2. Ash, S.R., 1991a, A new Jurassic flora from the Wallowa terrane in Hells Canyon, Oregon and Idaho: Oregon Geology, v. 53, p. 27–33. Ash, S.R., 1991b, A new Jurassic Phlebopteris (Plantae, Filicales) from the Wallowa terrane in the Snake River Canyon, Oregon and Idaho: Journal of Paleontology, v. 63, p. 800–819. Blodgett, R.B., Frýda, J., and Stanley, G.D., Jr., 2001, Delphinulopsidae, a new neritopsoidean gastropod family from the Upper Triassic (upper Carnian
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Printed in the USA
The Geological Society of America Special Paper 442 2008
Late Triassic (Carnian-Norian) mixed carbonate-volcaniclastic facies of the Olds Ferry terrane, eastern Oregon and western Idaho Todd A. LaMaskin* Department of Geological Sciences, 1272 University of Oregon, Eugene, Oregon 97403-1272, USA ABSTRACT The Late Triassic (Carnian-Norian) time interval is well represented by strata in the Wallowa and Baker terranes and in the Izee-Suplee area of central Oregon. Strata of the Olds Ferry terrane, by contrast, have received little attention beyond field mapping and reconnaissance-level efforts. This paper describes the sedimentology of Carnian-Norian deposits in the Olds Ferry terrane and describes a model for arc-flanking mixed carbonate-volcaniclastic sedimentation. A distinctive thin-bedded limestone unit in the lower Huntington Formation is underlain and overlain by volcanic and shallow intrusive rocks. Ammonite and bivalve age assignments indicate that the limestone unit spans the late Carnian to early Norian; portions of the section may, more specifically, represent the welleri or dilleri ammonite zones. Four lithofacies— Graded Skeletal Packstone and Mudstone, Tuffaceous Peloidal Grainstone, Skeletal Peloidal Packstone, and Lapilli-Tuff Breccia—record episodic sedimentation influenced by earthquakes, storms, and volcanic eruptions on an oxygen-poor carbonate slope apron. The limestone unit represents an echinoderm- and mollusk-dominated heterozoan faunal assemblage characterized by abundant delivery of unlithified carbonate sediment to deeper water. Keywords: Blue Mountains, mixed carbonate-volcaniclastic, Late Triassic fauna, slope apron. INTRODUCTION The Blue Mountains Province of eastern Oregon and western Idaho (Fig. 1) is a compound terrane assemblage of accreted Paleozoic-Mesozoic volcanic arcs, sedimentary basins, subduction mélange complexes, and post-tectonic stitching plutons (Brooks and Vallier, 1978; Dickinson, 1979; Avé Lallemant et al., 1980; Silberling et al., 1984; Burchfiel et al., 1992; White et al., 1992; Avé Lallemant, 1995; Vallier, 1995). Four lithotectonic terranes
trend generally east and northeast toward the Salmon River Belt (Gray and Oldow, 2005): the Wallowa, Baker, Izee, and Olds Ferry terranes (Silberling et al., 1984). Strata of the Izee terrane depositionally overlie older rocks and structures of the Wallowa, Olds Ferry, and Baker terranes (Dickinson and Thayer, 1978; Blome et al., 1986; White et al., 1992; Dorsey and LaMaskin, 2007), and thus represent a stratigraphic overlap assemblage linking the Blue Mountains terranes by Middle Jurassic time (i.e., White et al., 1992; Dorsey and LaMaskin, 2007).
*
[email protected] LaMaskin, T.A., 2008, Late Triassic (Carnian-Norian) mixed carbonate-volcaniclastic facies of the Olds Ferry terrane, eastern Oregon and western Idaho, in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 251–267, doi: 10.1130/2008.442(13). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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EXPLANATION Late Jr-Early K Plutons Olds Ferry terrane
Idaho Batholith Izee "terrane" Wallowa terrane Baker terrane
WASH OR
Salmon River Belt
WA LLO WA
WA OR
?
ID
Sr/86Sr .704 and .706 line
87
Snake River
ER
BAK
Huntington
E IZE
0 km 50
S OLD
RY Fig. 3
FER
OR ID
N
Figure 1. Regional geologic map of Blue Mountains Province showing distribution of terranes. Modified from Dickinson (1979), Mann (1989), and Gray and Oldow (2005).
Detailed analysis of sedimentary strata has been presented for units in the Wallowa terrane (e.g., White, 1972, 1986, 1994; Follo, 1986, 1992, 1994; Newton, 1986; Whalen, 1988; White et al., 1992; Goldstrand, 1994; White and Vallier, 1994) and the Izee terrane (e.g., Dickinson and Vigrass, 1965; Dickinson and
Late
AGE Ma TIHONIAN KIMMERIDGIAN OXFORDIAN
Middle
BATHONIAN BAJOCIAN AALENIAN
168 171 176
Late Middle
Triassic
190
LADINIAN
?
Weatherby Formation
Volcanic Lithic Phyllite, Shale and Siltstone Minor Conglomerate and Silicic Tuff
216 228 237
Crystalline Limestone/Marble, Relict Intraclast Pellet Packstone and Foram. Wackestone Textures
Blue Mtns. Overlap Assemblage
Volcanic-Clast, Stretched-Pebble Conglomerate, Volcanic Lithic Arenite and Phyllitic Siltstone
?
Jet Creek Mbr.
?
Volcanic-Clast Conglomerate, Turbiditic Sandstone, Shale and Basaltic to Rhyolitic Pyroclastic Deposits
?
Volcanic Breccia, Hyaloclastite, Basaltic to Rhyolitic Pyroclastic Deposits and Volcanic Flows; Volcanic-Clast Conglomerate and Volcanic Lithic Arenite
?
?
ANISIAN
Early
TERRANE ASSOCIATION
Unconformity
196 200 U. Huntington Fm. 204
NORIAN CARNIAN
The Late Triassic (Carnian-Norian) time interval is well represented by strata in the Wallowa and Baker terranes and in the Izee-Suplee area of central Oregon. In the Wallowa terrane, the
165
PLIENSBACHIAN
HETTANGIAN RHAETIAN
Upper Triassic Strata of the Blue Mountains Province
EXPLANATION
161
183
SINEMURIAN
GEOLOGIC SETTING
151 156
TOARCIAN
Early
Jurassic
CALLOVIAN
144
FORMATION AND LITHOLOGY
Thayer, 1978; Dickinson, 1979; Dickinson et al., 1979; Smith, 1980; Taylor, 1982; Taylor and Guex, 2002). These important works provide the basis for our current understanding of sedimentary basin dynamics and paleogeography during TriassicJurassic development of the Blue Mountains Province. Strata of the Olds Ferry terrane, by contrast, have received little attention beyond field mapping and reconnaissance-level efforts (Beeson, 1955; Berry, 1956; Spiller, 1958; Wagner et al., 1963; Brooks, 1967; 1979a). Herein, I describe in detail a Late Triassic (Carnian-Norian) limestone unit in the Olds Ferry terrane and propose a depositional model for arc-flanking mixed carbonate-volcaniclastic sedimentation. The results of this study will permit a future assessment of stratigraphic linkages between the Olds Ferry terrane and strata in the Izee-Suplee area of central Oregon and possibly other areas of the Blue Mountains Province during Late Triassic time.
?
Lower Huntington Fm.
Olds Ferry
Carnian-Norian age limestones of the Huntington Formation (this study) Faulted Contact
? ?
Age of Contact Uncertain Informal Formation Division
245 248
Figure 2. Mesozoic chronostratigraphy of the Huntington, Oregon region. Data sources are Brooks (1979a), Imlay (1986), and Tumpane et al. (2008).
Late Triassic mixed carbonate-volcaniclastic facies of the Olds Ferry terrane Carnian-Norian Martin Bridge Formation records cessation of volcanic activity and onset of widespread carbonate deposition (Follo, 1986, 1992, 1994; Whalen, 1988; McRoberts, 1993). This transition has been interpreted to represent the end of significant volcano-plutonic activity and initiation of subsidence of the Wallowa arc (Vallier, 1977; Brooks and Vallier, 1978; Pessagno and Blome, 1986; Follo, 1992, 1994; White, 1994). In the Izee-Suplee Region of central Oregon, Carnian-Norian deposits record a significant transition from a large, regional depocenter (Vester Formation) to multiple smaller, individual fault-bounded depocenters (Aldrich Mountains Group) (Dickinson and Vigrass, 1965; Dickinson, 1979; Blome et al., 1986; LaMaskin et al., 2004). In the Baker terrane, Carnian-Norian limestone lenses, possibly indicating shoaling of formerly deep-water basins, have been identified in multiple locations (Morris and Wardlaw, 1986; Nestell and Orchard, 2000). In terms of regional structure, Carnian-Norian (ca. 223–219 Ma) age mylonites from the Oxbow shear zone are interpreted to record sinistral transpression in the Wallowa terrane (Avé Lallemant et al., 1985) and have been used to infer the presence of a major intra-arc strike-slip fault system that was active in the Late Triassic (Avé Lallemant et al., 1985; Follo, 1986, 1992, 1994; Avé Lallemant, 1995; Vallier, 1995). Previous studies of Carnian-Norian stratigraphic and structural relationships in all terranes of the Blue Mountains Province provide an opportunity to assess potential links between terranes. An understanding of Carnian-Norian age strata in the Olds Ferry terrane is needed to evaluate potential associations between the Olds Ferry terranes and other areas of the Blue Mountains Province. Olds Ferry Terrane and Huntington Formation The Olds Ferry terrane of eastern Oregon and western Idaho represents a volcanic island-arc succession of Middle Triassic to Early Jurassic age (Brooks and Vallier, 1978; Dickinson and Thayer, 1978; Brooks, 1979a, 1979b; Silberling et al., 1984; Walker, 1986; Tumpane, 2008). Rocks of the Olds Ferry terrane (Fig. 2) include mafic to silicic hypabyssal and extrusive volcanics and interbedded coarse to fine volcaniclastic strata of the Huntington Formation (Brooks, 1979a), which have been variably altered to greenschist and prehnite-actinolite metamorphic facies (Goebel, 1990). The lower member of the Huntington Formation (Collins et al., 2000a, 2000b) consists of bedded to massive volcanic flows including thick- to medium-bedded volcanic breccia, pillowed greenstone, andesite porphyry, and minor rhyodacite (Beeson, 1955; Berry, 1956; Kennedy, 1956; Spiller, 1958; Juras, 1973; Brooks and Vallier, 1978; Brooks, 1979a; Charvet et al., 1990; Collins et al., 2000a, 2000b). These volcanic lithologies are intercalated with subordinate epiclastic strata including medium- to thick-bedded conglomerate, thin-bedded black siliciclastic mudstone, and thin-bedded, fossiliferous limestone. The dominance of abundant volcanic flows and hypabyssal intrusions with relatively minor epiclastic strata indicates that the lower Huntington Formation likely represents volcanic activity,
253
shallow intrusions, and associated sedimentation on the proximal flanks of the Olds Ferry arc. The upper member of the Huntington Formation (Collins et al., 2000a, 2000b) consists chiefly of laminated shale and thin- to medium-bedded sandstone turbidites interlayered with thick-bedded cobble to boulder conglomerate and thick-bedded rhyodacite porphyry (Fig. 2). Coarse sandstone and conglomerate beds typically display well-defined channel morphologies that record erosion into underlying sandstone units. The relative abundance of epiclastic strata and rhyodacite suggests that the upper Huntington Formation may represent more distal submarine fan sedimentation concurrent with a pronounced shift to more explosive silicic volcanism (San Fillipo, 2006; Dorsey and LaMaskin, 2007). Strata of the Huntington Formation are overlain by Lower Jurassic through Middle Jurassic stretched-pebble conglomerate, crystalline limestone, and slate-phyllite deposits of the Weatherby Formation (Brooks, 1979a; Imlay, 1986). The contact between the Huntington and Weatherby formations has been interpreted as depositional (Brooks and Vallier, 1978; Brooks, 1979a; Dickinson, 1979) and as a faulted unconformity (Blome et al., 1986). Recent work by Tumpane et al. (2008) indicates that the upper Huntington Formation also includes Lower Jurassic rocks. Additional mapping and geochronology is needed to resolve the nature of the contact between the two formations. OCCURRENCE AND AGE OF HUNTINGTON FORMATION LIMESTONE A unit of thin-bedded limestone in the lower member of the Huntington Formation is discontinuously exposed along the Snake River north and south of the town of Huntington, Oregon (Fig. 3), and on the Idaho side of the Snake River (Beeson, 1955; Berry, 1956; Kennedy, 1956; Spiller, 1958; Juras, 1973; Brooks, 1979a). The limestone unit is situated stratigraphically between hypabyssal intrusives and flows of amygdaloidal andesite to dacite porphyry of the lower and upper Huntington Formation and is directly underlain by a distinctive volcanic- and carbonateclast breccia bed (Fig. 4). Limestone exposures are generally poor (~10%–50% exposure) and confined to erosional hollows between ridges, hampering detailed lateral correlation (Fig. 5). Exposed sections of this limestone unit are ~30 m thick and contain a suite of lithofacies interbedded at the decimeter scale that are lithologically similar in different exposures. Lateral variations in lithofacies (e.g., deepening or shallowing, proximal to distal) are not discernable. Faunal collections from northeast of the town of Huntington, Oregon by Beeson (1955, identifications by S.W. Muller), indicated the presence of the ammonites Arcestes sp., Tropites sp., Discotropites sp., and Sagenites sp., and the flat clam Halobia cf. H. superba and were interpreted to represent a Late Triassic, Carnian age. Limestone exposures on the Idaho side of the Snake River examined by Juras (1973) contain halobiid bivalves and spongiomorphs indicative of a late Carnian age (identifications
254
LaMaskin Amygdaloidal andesite to dacite porphyry, mafic dikes and hypabyssal intrusions
Carnian-Norian age limestone unit
I-84
Limestone- and volcanicclast breccia Amygdaloidal andesite to dacite porphyry, mafic dikes and hypabyssal intrusions ~15 m Fig. 6a
Figure 4. Illustration of stratigraphic position of limestone unit within the lower Huntington Formation. Fig. 6b ak e
N
Huntington 5
WA
H
er
km
Riv
0
Sn
I-84
w
y. 3
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OR ID OR
ID
EXPLANATION Weatherby Fm
Huntington Fm.
Normal fault, ball on hanging wall
Limestone in Jet Creek mbr., Weatherby Fm Jet Creek mbr., Weatherby Fm.
Brownlee Pluton
Interstate/State hwy.
Tertiary-Quaternary volcanic and sedimentary rocks and landslide deposits
Figure 3. Geologic map of the Huntington, Oregon, area. Modified from Juras (1973) and Brooks (1979a).
by N.J. Silberling). Brooks (1979a, identifications by R.W. Imlay) noted that the presence of ammonites and Halobia sp. indicates a late Carnian to late middle Norian age for the Huntington Formation. To refine the age assignment for the limestone interval in the Huntington Formation, I collected samples from two locations (Fig. 6). Location 1 is in Spring Gulch, a tributary to Brownlee reservoir, ~9 km north of Huntington. Location 2 is ~2.8 km northeast of Huntington. Abundant ammonites and bivalves are present in float and thus provide an age assignment for the interval as a whole. Ammonites indicative of the Tropites subullatus zone of Smith (1927) were identified
by D. Taylor (2004, personal commun.) including Clionites sp., Arcestes (Proarcestes) cf. carpenteri, Tropites sp. and Discotropites sp. In addition, the Halobiid bivalves Halobia superba?, Halobia cf. austriaca, Halobia cf. superba, and Halobia ornatissima were identified by C. McRoberts (2004, personal commun.), indicating a late Carnian to early Norian age. The ammonite and bivalve age assignments indicate that the limestone unit spans the late Carnian to early Norian. The presence of Halobia superba and Halobia ornatissima indicates that portions of the section may, more specifically, represent the equivalent of the welleri or dilleri ammonite zones of Silberling and Tozer (1968) (McRoberts, 1993). The unique stratigraphic association of the limestone unit and underlying breccia, and the consistency of biostratigraphic ages reported for limestone exposures throughout the area (Beeson, 1955; Juras, 1973; Brooks, 1979a; this study) suggest that the limestone unit represents a single interval in the lower Huntington Formation that is structurally repeated in the area north of Huntington. This inference remains to be tested in future mapping. LITHOFACIES DESCRIPTIONS Four distinct lithofacies are recognized in the Carnian-Norian limestone unit of the lower Huntington Formation on the basis of bed thickness, grain types, sedimentary structures, and fossil content. Lithofacies identified are (1) Graded Skeletal Packstone and Mudstone Facies, (2) Tuffaceous Peloidal Grainstone Facies, (3) Peloidal Skeletal Packstone Facies, and (4) Lapilli-Tuff Breccia Facies. These rock types are intercalated as thin to medium beds throughout the 30 m limestone interval (Figs. 5 and 7). Detailed descriptions and environmental interpretations are summarized in the following and in Table 1.
Thin- to medium-, swaley bedded peloidal skeletal packstone composed of fine- to coarsesand-sized abraded skeletal debris. Peloids are fine- to medium-sand sized. Present as interbeds within graded skeletal packstone to mudstone facies. Abundant pebble-sized volcanic clasts, oncoids, and peloidal intraclasts. Thin- to medium-bedded tuffaceous skeletal peloidal grainstone. Abundant medium- to coarsesand-sized volcanic clasts. Peloids are fine- to coarse-sand sized. Quartz sand constitutes ~5–7% of framework grains.
Thin- to medium-bedded, normally graded, volcanic-clast breccia. Pebble- to cobble-sized clasts grading upward to crystal and glass-rich lapillistone. May be overlain by cuspate- to blockyshard vitric tuff.
Peloidal Skeletal Packstone
Graded LapilliTuff Breccia
Tuffaceous Peloidal Grainstone
Thin bedded, thin (1–3 mm) to thick (3–10 mm) laminae of skeletal packstone/ wackestone composed of silt- to med-sand sized abraded skeletal debris grading upwards to argillaceous lime mudstone. Quartz silt constitutes >1–3% of wackestone portions.
Normal grading of thin to medium beds, planar bases. Vitric tuff beds show planar lamination.
Normal grading in 3–10cm-thick beds, planar lamination. Scour and fill structures <1.5 m wide and <10 cm thick. Flame structures.
Laterally discontinuous scour-and-fill structures <0.5m wide. Normal grading in 1–7-cm beds. Hummocky cross-stratification and symmetrical wave ripples.
Thin graded laminae, ripple cross-lamination, rare scour-and-fill structures.
Barren
Ex situ bivalve, crinoid, gastropod, foraminifer, and algal skeletal debris.
Skeletal packstone-wackestone: Predominantly ex situ crinoid skeletal sand with minor bivalve, gastropod, and sponge spicule debris. Mudstone: Abundant Halobia sp. and Monotis sp. on beddingplane surfaces, abundant juvenile ammonites, calcispheres (replaced radiolaria?). Predominantly ex situ crinoid skeletal sand with minor bivalve, gastropod, algae and sponge spicule debris. Abraded and fragmental ammonites and Halobia sp.
Weathered: Light brownish gray (2.5Y 6/2). Fresh: Grayish brown (2.5Y 5/2).
Weathered: Light olive brown (2.5Y 5/4). Fresh: Light brownish gray (2.5Y 6/2).
Weathered: Very pale brown (10YE7/3) to gray (5Y 5/1). Fresh: Very dark gray (5Y 3/1).
Weathered: Dark gray (2.5Y 4/1) to very dark gray (5Y 3/1); Beddingplane surfaces pale yellow (2.5Y 7/3) to light yellowish brown (2.5Y 6/3). Fresh: Black (5Y 2.5/1).
TABLE 1. DETAILED DESCRIPTIONS OF HUNTINGTON FORMATION LIMESTONES Bedding and lithology Sedimentary structures Biota Color
Graded Skeletal Packstone to Mudstone
Lithofacies
Neomorphic spar cement, dolomitized basaltic glass (caries texture), calcite, quartz, and white mica replacement of volcanic grains.
Dolomitized basaltic glass, patches of complete neomorphic spar with skeletal and peloidal ghosts, calcite, volcanic grains replaced by quartz and white mica.
Aggrading neomorphic spar, dolomitized basaltic glass, patches of complete neomorphic spar with skeletal and peloidal ghosts, calcite, volcanic grains replaced by quartz and white mica.
Abundant solution seams, stylolites, aggraded neomorphic spar.
Diagenetic features
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LaMaskin
Figure 5. Field photograph illustrating typical outcrop expression of the Huntington Formation limestone unit. Outcrop consists of decimeter-scale interbedding of the four lithofacies described in this study. Scale bar represents ~1 m.
1m
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g rin Sp lch Gu 250 m
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Snake River Rd.
2200
EXPLANATION
Paved road Unpaved road Railroad tracks
Exposed section of CarnianNorian limestones Sewage Lagoons
Figure 6. Detailed maps of limestone exposures within the Huntington Formation. See Figure 3 for locations. Topographic base (in feet) is from U.S. Geological Survey 1:24,000 topographic maps.
Graded Skeletal Packstone and Mudstone Facies Graded Skeletal Packstone and Mudstone Facies (Table 1) is the dominant rock type in the lower Huntington Formation limestone unit (Fig. 8). Individual sharp-based, normally graded depositional couplets (Fig. 8A and 8B) include a basal laminated skeletal packstone/wackestone composed of abraded silt- to medium-sand–sized crinoidal skeletal debris with minor bivalve,
gastropod, brachiopod, and sponge skeletal debris (Fig. 8C) overlain by laminated argillaceous lime mudstone with abundant circular, calcite-filled vugs (Fig. 8D). The laminated argillaceous lime mudstone portion of couplets contains abundant juvenile ammonites and Halobia flat clams (i.e., Halobia limestones: Zankl, 1971; Wilson, 1975; Hallam, 1981; Bellanca et al., 1995; Parrish et al., 2001) (Fig. 8E). Both decimeter- and meter-scale intraformational slump folds are observed in this lithofacies (Fig. 8F). This lithofacies is interpreted to record episodic, distal, nonchannelized deposition from turbidity currents in a low-oxygen, slope to toe-of-slope setting. Turbidity currents were likely triggered by earthquakes, storm-induced disturbance of shallow settings, and/or bottom-return flow of storm currents. The basal portion of each couplet was produced by basinward transport and redeposition of reworked intermediate- and shallow-water skeletal sand, whereas the upper laminations in each couplet represent waning-flow deposition and background suspension settling of siliciclastic silt and carbonate mud. Evidence for deposition from distal turbidity currents includes thin- to thick-graded laminae, parallel and ripple cross-lamination, and rare basal scour structures (Markello and Read, 1981; Aigner, 1985; Galli, 1989; Lee and Kim, 1992; Monaco, 1992; Maurer et al., 2003). The presence of intraformational slump folds is evidence for mass failure and subsequent downslope movement of weakly lithified beds and indicates deposition in a slope to toe-of-slope setting. Deposition under low-oxygen conditions is indicated by a lack of bioturbation (i.e., preservation of millimeter-scale laminations) and the presence of a restricted fauna of ammonoids and Halobia bivalves (cf. Zankl, 1971; Wilson, 1975; Finch and Abbott, 1977; Hallam, 1981; Newton, 1986; McRoberts, 1993, 2001; Bellanca et al., 1995; Parrish et al., 2001; Martini et al.,
Late Triassic mixed carbonate-volcaniclastic facies of the Olds Ferry terrane
EXPLANATION
257
13
Graded Skeletal Packstone and Mudstone
Peloidal Skeletal Packstone Graded Lapilli-Tuff Breccia Graded Bedding/Lamination Hummocky Cross Stratification Symmetrical Wave Ripples
Meters above Base
Tuffaceous Peloidal Grainstone
Figure 7. Partial measured section of Huntington Formation limestone unit ~2.8 km northeast of Huntington, Oregon. Section illustrates interbedding and vertical succession of the four lithofacies identified and described in this study.
12
Ripple Cross-Lamination Planar Laminations
11
W Mud ac Pake Br G c ec ra k /C in on g.
Flame Structures
2004; Schatz, 2005). Circular, calcite-filled vugs in argillaceous lime mudstone portions of depositional couplets are likely calcite-replaced radiolaria; however, diagenetic effects have destroyed test walls and distinctive morphology is not visible. The interlamination of an open-marine fauna that is highly abraded and well rounded with a restricted pelagic fauna found in laminated lime mud is evidence for downslope transport of platform-derived carbonate sediment. Tuffaceous Peloidal Grainstone Facies Tuffaceous Peloidal Grainstone Facies (Table 1) consists of thin- to medium-bedded, erosive-based, graded volcaniclasticgrainstone beds (Fig. 9A). Because this lithofacies consists of >50% carbonate allochems with little to no lime mud, it is classified according to depositional texture as a grainstone; the remainder consists of volcanic lithic fragments including lathwork, microlitic, and vitric grains (Dickinson, 1970; Ingersoll and Cavazza, 1991; Marsaglia, 1992) that are variably calcitized, dolomitized, silicified, and altered to clay minerals. These grains are generally <2 mm and thus represent fine to coarse tuff (Fisher and Schminke, 1984). Following the convention of Schmid (1981), Fisher and Schmincke (1984), and McPhie et al. (1993), and to discriminate that the volcanic portion of these deposits is epiclastic, they are referred to as tuffaceous peloidal grainstones.
Tuffaceous peloidal grainstone is massive to planar laminated with rare cross-lamination. Peloids are fine- to medium-sand sized and are composed of lime mud. Skeletal grains include highly abraded crinoid, bivalve, gastropod, foraminiferal, and algal fragments; coated grains and ooids are conspicuously absent (Fig. 9B). Beds are present as both laterally extensive thin beds and channelized medium beds. Rare flame structures are present, where underlying lime mud was injected upward into overlying tuffaceous peloidal grainstone beds. The presence of unaltered lime mud flame structures suggests that the surrounding mud-free grainstone texture is original and is not the result of neomorphism of an original packstone texture. Volcanic Lithic Grains The presence of volcanic lithic grains in the tuffaceous peloidal grainstones represents mixing of volcanic and carbonate allochemical components and provides information about the nature of the volcanic source terrane. Lathwork volcanic lithic grains, derived from basaltic and basaltic-andesite sources, are recognized as sand-sized plagioclase phenocrysts in an intergranular groundmass. Plagioclase phenocrysts are commonly altered to white mica (i.e., sercicitized) or partly to wholly replaced by calcite. Microlitic volcanic lithic grains (Fig. 9C) are commonly altered to clay minerals and consist of plagioclase microlites in a brown to orange groundmass, commonly displaying trachytic textures. Common clay alteration results in a microcrystalline, low-order
Late Triassic mixed carbonate-volcaniclastic facies of the Olds Ferry terrane birefringence. Vitric volcanic lithic grains comprise a wide variety of textures including both classical and banded perlitic glass (Fig. 9D), wispy scoriaceous glass (Fig. 9E), blocky to cuspate shards, and ovoid, globular grains (Fig. 9F). Vitric grains are light brown to yellow (i.e., palagonite), and are commonly replaced by calcite, dolomite, or clay minerals. In these instances, they are recognizable only by iron stains defining grain outlines and former concentric or vesicular internal fabric. Ovoid globular vitric grains (80–500 μm diameter; Fig. 9F) are a common constituent in the tuffaceous peloidal grainstone facies and impart a distinctive yellow color in both hand sample and thin section. They are predominantly altered to clay minerals and are commonly bordered by a rim of prismatic calcite crystals. Some ovoid globular vitric grains display internal concentric structures that superficially resemble coated grain or ooidal textures; however, the grains are composed wholly of basaltic glass and the concentric fabric is interpreted to represent cracking during cooling of basaltic glass droplets (cf. Bertani and Carozzi, 1985a, 1985b). Basaltic glass commonly develops on the margins and within the matrix of pillow lavas both as spalled shards and as quenched glass droplets (Carlisle, 1963; Dimroth et al., 1978). Bertani and Carozzi (1985a, 1985b) interpreted basaltic glass globules in a micritic matrix as hyaloclastite formed by mixing of extruding lava with unconsolidated sediments. The authors suggested that turbidity currents triggered by the turbulence of the mixing process contributed to the downslope transport of glass and sediment. Ovoid globular vitric grains within the tuffaceous peloidal grainstones of the Huntington Formation likely had their origin as a component of pillow lavas which have been identified in the area (Juras, 1973; Charvet et al., 1990). Grains were subsequently reworked and transported offshore as density-driven turbidity currents. Further evidence for erosional mixing of tuffaceous peloidal grainstone includes the abundance of shallow-water skeletal debris, well-rounded basaltic lithoclasts, and a wide variety of glass and lithic textures indicative of both subaerial (e.g., pumice/scoria) and subaqueous (e.g., non-vesicular, blocky shards) textures (cf. Schmincke and Sumita, 1998).
Figure 8. Photographs and photomicrographs of Graded Skeletal Packstone and Mudstone facies. (A) Field photograph. Note individual couplets of skeletal packstone (dark) and argillaceous mudstone (light). Tip of “Sharpie” pen for scale. (B) Plane-polarized light (PPL) photomicrograph of Graded Skeletal Packstone and Mudstone Facies; arrows denote individual graded couplets. (C) Photomicrograph (PPL) illustrating skeletal packstone/wackestone at the base of a couplet. Note abundant skeletal debris including crinoids, bivalves and gastropods. (D) Photomicrograph (PPL) of argillaceous mudstone at the top of a couplet. Note abundant dissolution seams and aggrading neomorphism of lime mud. Recrystallized circular vugs likely represent radiolarian tests. (E) Photograph of bedding-plane surface illustrating high-density accumulation of Halobia ornatissima bivalve shells. (F) Field photograph of intraformational slump fold. Note concordance of beds above and below slumped interval. Faithful canine for scale.
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Depositional Mechanism This lithofacies is interpreted to represent volcaniclasticcarbonate deposits of turbidity currents in a slope setting adjacent to volcanic edifices. The heterogeneous composition of volcanic lithic grains that are rounded to subrounded, and the presence of admixed shallow-water carbonate skeletal debris indicates that sediment was derived from (1) erosion of subaerial volcanic deposits, (2) remobilization of subaqueous, spalled basaltic glass associated with pillow basalts, and (3) arc-proximal, shallow-water carbonate zones (cf. Schmincke and Sumita, 1998; Thordarson, 2004). Laterally extensive thin beds represent unchannelized turbidites, whereas rare channelized medium beds are interpreted to represent channel scour-and-fill in feeder channel lobes. A storm origin for these beds must be considered because of the interbedded nature of these deposits with skeletal peloidal packstone facies (see below), which is demonstrably storm-wave generated. Graded beds that resemble classical turbidites can be generated by storm-induced combined flows; however, where a storm origin has been established for “turbidite beds,” they characteristically show a spectrum of sedimentary structures related to storm-wave modification (Nelson, 1982; Myrow and Southard, 1996; Myrow et al., 2002; Myrow, 2005). Sedimentary structures of the tuffaceous peloidal grainstone facies described here are not consistent with structures developed under combined-flow conditions. Specifically, a lack of storm-wave–induced sedimentary structure associations, such as symmetrical ripples, hummocky cross-stratification, climbing ripples, starved ripples, thickening and thinning lenticular beds, and lag deposits, rules out a direct storm origin for these beds. Peloidal Skeletal Packstone Facies Peloidal Skeletal Packstone Facies (Table 1) is intercalated with Graded Skeletal Packstone and Mudstone Facies as thin to medium, hummocky beds with erosive bases (Figs. 10A–10C). Beds are lenticular and laterally discontinuous, and are composed of stacked graded peloidal skeletal-lag deposits dominated by crinoid and flatclam debris (Fig. 10D). Sedimentary structures include normal and reverse grading of thin beds, ripple cross-lamination, flat-based, convex-up symmetrical ripples, and hummocky cross-stratification (Fig. 10B). Pebble-sized pumice-scoria clasts that have been replaced by calcite, oncoids, and peloidal intraclasts are common. These beds are interpreted to represent winnowing and sediment redistribution by large storm currents such as those generated by winter hurricanes (Marsaglia and Klein, 1983; Quiquerez et al., 2004). Sediment, including poorly sorted skeletal sand and volcanic lithic debris, peloids, intraclasts, and oncoidal pebbles, originated in a shallow to intermediate setting and was remobilized and transported basinward in storm-generated turbidity or return-flow currents during large storm events. Interbedding of Peloidal Skeletal Packstone Facies with low oxygen deposits of Graded Skeletal Packstone and Mudstone facies may record lowering of relative sea level or may be the result of larger storms that effectively increased the depth of storm wave base (Aigner, 1985; Elrick and Snider, 2002), allowing reworking and
Late Triassic mixed carbonate-volcaniclastic facies of the Olds Ferry terrane winnowing of the depositional surface while also remobilizing sediment from shallow water into the basin. Graded Lapilli-Tuff Breccia Facies Graded Lapilli-Tuff Breccia Facies (Table 1) consists of rare, thin- to medium-bedded, normally graded, planar-based units intercalated with other facies described here (Fig. 11A). This lithofacies is unique in that it consists exclusively of volcanic grains and clasts; it is barren of any carbonate allochemical constituents of other facies in the limestone unit (Figs. 11A and 11B). Also conspicuously absent are lenticular beds, basal scour features, clast-imbrication, or cross-stratification, any of which might indicate deposition as a current-emplaced bed (Maicher et al., 2000; White, 2000). The upper 1–2 cm of some beds consists of crystal-rich lapillistone composed of subhedral plagioclase crystals and basaltic glass in association with volcanic lithic grains (Figs. 11A and 11C). Beds of graded lapilli-tuff breccia are commonly overlain by a thin bed of laminated, well-sorted, cuspate- to blocky-shard vitric tuff containing minor amounts of silt-sized quartz and feldspar (Fig. 11D). Volcanic lithic constituents are predominantly amygdaloidal andesite porphyry and intersertal-lathwork and microlitic clasts and vesicular pumice-scoria clasts that have been wholly replaced by calcite. Rare clasts appear to be composed of polycrystalline quartz and may easily be misinterpreted as chert or sandstone grains; however, close inspection indicates that they are the result of selective silicification following calcite replacement of plagioclase crystals or vesicular pumice clasts. Dolomite-replaced zones, indicated by scalloped replacement fronts (i.e., caries texture), are ubiquitous along the margins of volcanic clasts and in places have wholly or partly replaced clasts (Fig. 11B). Where replacement of volcanic clasts can be confirmed by the presence of ghost feldspar crystals and clast outlines, the replacive dolomite is a dusty, coarse, equigranular microspar with prominent enfacial angles. Where clasts are not wholly replaced, pore spaces are lined with dusty microspar that closely mimics clast outlines (Fig. 11B) and is petrographically identical to the dusty, clast-replacive microspar. Pore spaces are subsequently infilled with clear, coarse, equigranular microspar and where pore spaces are particularly large (e.g., 1–2 mm) by a core of clear monocrystalline quartz. No evidence of allochemical ghosts
Figure 9. Photographs and photomicrographs of Tuffaceous Peloidal Grainstone Facies. (A) Field photograph of stacked, normally graded thin beds; dashed lines indicate sharp, erosive contacts. Pen for scale. (B) Photomicrograph (PPL) illustrating abundant peloids and crinoid skeletal fragments as well as volcanic lithic grains and opaque minerals. (C) Cross-polarized light (XPL) photomicrograph of microlitic vitric volcanic lithic gain. Note well-developed trachytic texture. (D) Photomicrograph (PPL) of perlitic vitric volcanic lithic grain. (E) Photomicrograph (PPL) of wispy scoriaceous vitric volcanic lithic grain. (F) Photomicrograph (PPL) of globular vitric volcanic lithic grain. Note the presence of equant, prismatic spar along gain periphery.
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has been observed in the pore spaces of the Graded Lapilli-Tuff Breccia Facies. These deposits are interpreted as originally clast-supported breccia fabrics in which pore spaces have been enlarged by dolomite replacement of volcanic-clast margins. Dusty, coarse, equigranular microspar that tracks clast outlines represents the former grain boundaries, and clear, coarse, equigranular microspar represents the original pore space. On the basis of the above evidence, this lithofacies is interpreted to represent subaqueous suspension-fall of ash and lapilli following either submarine or subaerial pyroclastic eruptions (Carey and Sigurdsson, 1984; Cashman and Fiske, 1991; Schmincke and Sumita, 1998; Carey, 2000; White, 2000). Overlying laminated vitric tuff beds may represent suspension settling of the fine ash portion of erupted materials or may represent reworking and deposition of fine ash via dilute turbidity currents following eruptive events (White and Busby-Spera, 1987; Cousineau, 1994; White, 2000). DISCUSSION Platform Morphology and Sediment Dynamics Limited exposures of the Upper Triassic limestone unit hamper lateral correlation and assessment of regional lithofacies distribution. Thus, any depositional model invoked to explain the variety of lithofacies present is speculative and must be based on assumptions regarding the nature of the shallow carbonate system. Nonetheless, the spectrum of lithofacies permits generation of a plausible depositional model. The Huntington Formation limestone unit records episodic sedimentation influenced by earthquakes, storms, and volcanic eruptions in a low-oxygen marine basin offshore of volcanically constructed edifices (Fig. 12). Lithofacies described here are interpreted to represent deposition on a carbonate slope apron adjacent to volcanic edifices, possibly in a distally steepened ramp system. Carbonate slope aprons are typically the site of interbedding of broad sheets of carbonate sediment gravity flows and suspension deposits but typically lack systematic vertical sequences characteristic of submarine fan turbidite deposits (Coniglio and Dix, 1992). The predominantly volcanic strata of the lower Huntington Formation, which is intruded by Middle and Upper Triassic plutons (Walker, 1986) have been interpreted to represent the arc massif of an intra-oceanic (Avé Lallemant, 1995; Vallier, 1995) or fringing-arc complex (Dickinson, 1979, 2004). Thus, the stratigraphic relationship of mixed volcaniclastic-carbonate deposits with extrusive volcanics suggests that deposition occurred in a proximal forearc or intra-arc setting. Within the Huntington Formation, the upsection change from extrusive and hypabyssal volcanic rocks to a distinctive volcanic- and carbonate-clast breccia, to overlying carbonate turbidites (Fig. 4) may indicate the presence of an initially sloped basin margin and can be interpreted to indicate a fault-bounded basin such as is commonly
Figure 10. Photographs and photomicrographs of Peloidal Skeletal Packstone Facies. (A) Field photograph illustrating normally graded thin bed (arrow), interbedded within Graded Skeletal Packstone to Mudstone Facies (GSPMF). Arrow also denotes thickening of individual bed into a wave ripple, which is in turn draped and onlapped by overlying deposits. Pen for scale. (B) Field photograph illustrating decimeter-scale hummocky cross-stratification. Pen for scale. (C) Field photograph of multiple stacked graded beds of Peloidal Skeletal Packstone Facies overlying Graded Skeletal Packstone to Mudstone Facies (GSPMF). Note base of first marked bed is erosive into underlying Graded Skeletal Packstone to Mudstone Facies. Truncated unit between beds 3 and 4 is Graded Skeletal Packstone to Mudstone Facies. Arrow 1 points to flat-based symmetrical ripple. Arrow 2 points to internal truncation surface. (D) Full-slide scan representing a portion of single thin bed. Volcanic clasts (VC) represent scoriaceous pebbles that have been extensively replaced by coarse spar.
Late Triassic mixed carbonate-volcaniclastic facies of the Olds Ferry terrane present in the forearc or intra-arc region (e.g., Dickinson, 1995; Smith and Landis, 1995; Carey, 2000). Shallow-water carbonate shoals and intermediate-depth crinoid thickets were located on fault-bounded structural highs adjacent to the volcanic arc (e.g., Dorobek, 2005) and were the source of abundant sediment input to slope and basin environments. The slope apron was the site of episodic deposition of thin storm- and earthquake-generated graded beds. During times of tectonic and storm quiescence, the slope accumulated siliciclastic and lime mud, pelagic ammonite and bivalve shells, and radiolarian tests via suspension settling. Thinner beds and laminae of the Graded Skeletal Packstone and Mudstone Facies represent energetically smaller or more distal (i.e., distal apron channel overbank) turbidity current deposits, whereas
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thin to medium beds of Tuffaceous Peloidal Grainstone Facies represent energetically larger, or more proximal (i.e., apron channel and proximal apron channel overbank) turbidity current deposits. The apron surface was extensively reworked and winnowed by wave currents either as a result of large storms that effectively deepened storm wave base or during times of low relative sea level (cf. Aigner, 1982; Handford, 1986; Jennette and Pryor, 1993; Elrick, 1996; Badenas and Aurell, 2001; Elrick and Snider, 2002). Storms also swept shallower-water settings, mobilizing carbonate and volcaniclastic sediment that was subsequently reworked, winnowed, and deposited on the apron surface. Strata of the Huntington Formation limestone unit represents a predominantly non-tropical, heterozoan-type faunal
Figure 11. Photographs and photomicrographs of Lapilli-Tuff Breccia Facies. (A) Scan of a polished slab illustrating three-part division. (B) Scanned image of a thin section illustrating millimeter-scale pores rimmed with dusty spar and cored by clear spar; arrows indicate pervasive dolomite replacement of volcanic clasts and grains. (C) Photomicrograph (PPL) of crystal-rich lapillistone. L—lithic fragment; Pl—plagioclase crystals; P—palagonitic glass; C—calcite cement. (D) Photomicrograph (PPL) of vitric tuff bed. Note abundant blocky and cuspate shards.
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LaMaskin
EXPLANATION Shallow subtidal carbonate and volcaniclastic sand and pebbles Intermediate subtidal carbonate mud and volcaniclastic silt-pebbles
Rainout of Ash and Lapilli
Graded laminae/beds of carbonate-volcaniclastic sand Carbonate-volcanic clast breccia
Shallow Subtidal Volcaniclastic-carbonate Shoals
Andesite porphery Pillow lava Intermediate Subtidal Crinoid Thickets SL FWWB SWB Deep Subtidal (?)
Carb
on
lope ate S
Apro
n
Figure 12. Conceptual depositional model for mixed volcaniclastic-carbonate deposits of the Huntington Formation.
assemblage composed of echinoderms (specifically crinoids) and mollusks (bivalves and gastropods) characterized by abundant sediment delivery of unlithified carbonate material to deep water (James, 1997; Dorsey and Kidwell, 1999; Pomar, 2001) despite the likelihood of a tropical latitude of the basin during deposition (Hillhouse et al., 1982). The proliferation of heterozoan faunal assemblages may occur under nutrient-rich surface-water conditions such as zones of increased nutrient load from runoff, in upwelling zones, or in areas of cold-water intrusion (Whalen, 1995; James, 1997; Parrish et al., 2001, Pomar et al., 2004). Current paleogeographic data are insufficient to establish which of these mechanisms may have operated to impede tropical carbonate platform growth along the Olds Ferry arc. CONCLUSIONS 1.
2.
A distinctive limestone unit within the Huntington Formation (Olds Ferry terrane) spans the late Carnian to early Norian. The presence of Halobia superba and Halobia ornatissima indicates that portions of the section may represent the equivalent of the welleri or dilleri ammonite zones (McRoberts, 1993). Four lithofacies—Graded Skeletal Packstone and Mudstone Facies, Tuffaceous Peloidal Grainstone Facies, Skeletal Peloidal Packstone Facies, and Lapilli-Tuff Breccia Facies record episodic earthquake, storm, and volcanic eruption-influenced sedimentation in a low oxygen, carbonate slope apron adjacent to volcanic edifices, possibly in a distally steepened ramp system. Shallow-
3.
water carbonate shoals and intermediate-depth crinoid thickets were located on fault-bounded structural highs adjacent to the volcanic arc and were the source of abundant sediment input to slope and basin environments. The limestone unit represents a predominantly heterozoan-type faunal assemblage composed of echinoderms and mollusks with abundant sediment delivery of unlithified carbonate material to deep water despite the likelihood of a tropical latitude for the basin during deposition.
ACKNOWLEDGMENTS Funding was provided by grants from the Weimer fund of SEPM (the Society for Sedimentary Geology), the Geological Society of America (Graduate Student Research Grant no. 7721-04), and the Baldwin Fellowship at the University of Oregon. Field assistance with fossil collection was provided by T. Sieber. The manuscript benefited greatly from early reviews by R. Dorsey and S. Boggs. I appreciate conversations with R. Dorsey, S. Boggs, G. Retallack, H. Wright, and J. Roberge that provided clarity. D. Taylor and C. McRoberts identified ammonite and bivalve collections, respectively, and assisted in their biostratigraphic interpretation. B. and L. Orr assisted with the University of Oregon Condon collection and were miraculously able to retrieve the collections of Beeson (1955) from long-term storage. Critical reviews from GSA reviewers M. Whalen, S. Dorobek, and G. Stanley greatly improved the manuscript and are much appreciated.
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The Geological Society of America Special Paper 442 2008
Early Jurassic bivalves of the Antimonio terrane (Sonora, NW Mexico): Taxonomy, biogeography, and paleogeographic implications Annemarie Scholz* Museum für Naturkunde der Humboldt-Universität zu Berlin, Invalidenstraße 43, 10115 Berlin, Germany Martin Aberhan* Museum für Naturkunde der Humboldt-Universität zu Berlin, Invalidenstraße 43, 10115 Berlin, Germany Carlos M. González-León* Instituto de Geología, ERNO, Universidad Nacional Autónoma de México, Apartado Postal 1039, Hermosillo, Sonora, Mexico 83000 ABSTRACT The Early Jurassic (late Hettangian to early Toarcian) bivalve fauna of the Sierra de Santa Rosa Formation of the Antimonio terrane (Sonora, NW Mexico) is analyzed taxonomically and biogeographically. Fifty taxa are recognized, representing 36 genera and subgenera. Thirty-four of these taxa have not been mentioned from the Jurassic of this region previously. This fauna is of great biogeographical interest, because Early Jurassic bivalves from low paleolatitudes of the tectonically complex western margin of North America are still poorly documented. About half of the described species are also known from other localities along the eastern Pacific margin. The second largest group is composed of widespread taxa, which, in addition to eastern Pacific occurrences, are also reported from other regions, particularly from Europe. The smallest group is endemic taxa that appear to be limited to Sonora during the analyzed time intervals. Geological evidence indicates that the Antimonio terrane was tectonically transported southeastward between the Middle and Late Jurassic from an original position at the southwestern margin of the United States by the MojaveSonora megashear. We calculated similarities of contemporaneous pectinoid bivalve faunas from seven eastern Pacific regions to independently constrain Early Jurassic paleolatitudinal positions of this terrane. Cluster analyses and similarity coefficients tentatively suggest that tectonic displacement of the Antimonio terrane toward lower paleolatitudes may already have started in Early Jurassic (Pliensbachian) time. Keywords: bivalvia, Jurassic, paleobiogeography, taxonomy, northwest Mexico.
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Scholz, A., Aberhan, M., and González-León, C.M., 2008, Early Jurassic bivalves of the Antimonio terrane (Sonora, NW Mexico): Taxonomy, biogeography, and paleogeographic implications, in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 269–312, doi: 10.1130/2008.442(14). For permission to copy, contact editing@ geosociety.org. ©2008 The Geological Society of America. All rights reserved.
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INTRODUCTION The western part of North America, i.e., the North American Cordillera, is geologically highly complex. Since the Paleozoic, more than 200 allochthonous terranes have been accreted to an active continental margin (Coney et al., 1980; Frisch and Loeschke, 1993). The origin of many of these terranes remains unresolved, and their fossil faunas are still poorly documented. Here, we study the Jurassic bivalve fauna from one of those terranes, the Antimonio terrane of Sonora, Mexico. In particular, we (1) describe in detail and illustrate the numerous bivalves from the Early Jurassic Sierra de Santa Rosa Formation; (2) analyze these new data biogeographically by comparing them with contemporaneous faunas from other regions along the eastern margin of the paleo-Pacific; and, on this basis, (3) address the question of postdepositional tectonic displacement of the Antimonio terrane. Apart from short faunal lists (see Damborenea and GonzálezLeón, 1997, for references), previous documentation of Early Jurassic bivalves from Sonora was provided by Jaworski (1929) and, more recently, by Damborenea and González-León (1997). From the material collected by Keller (1928) and housed at the Naturhistorisches Museum Basel, Jaworski (1929) described and partly figured seven species from the Sierra de Santa Rosa. Damborenea and González-León (1997) described and illustrated 15 taxa from the Sierra del Álamo and Sierra de Santa Rosa, 12 of which had not been reported previously from this region. The material studied by Damborenea and González-León (1997) is housed in the collections of the Estación Regional del Noroeste, Instituto de Geología, Universidad Nacional Autónoma de México in Hermosillo (ERNO). The bivalve fauna analyzed in the study at hand comprises 50 taxa, 34 of which are documented from the Lower Jurassic of Mexico for the first time. GEOLOGICAL SETTING, LOCALITIES AND STRATIGRAPHY The Antimonio terrane has been proposed as an allochthonous terrane that rests on the Proterozoic to Permian basement of the Caborca terrane (González-León, 1989; Stanley and González-León, 1995). It is composed of the Permian (Guadalupian) Monos Formation (Cooper and Arellano, 1946) and the Upper Permian to Lower Jurassic El Antimonio Group (GonzálezLeón et al., 2005) that disconformably overlies it. The Caborca terrane along with the Antimonio terrane is considered to have been tectonically transported southeastward between the Middle and Late Jurassic from an original position at the southwestern margin of the United States by the Mojave-Sonora megashear, a fault for which geological evidence suggests a left lateral displacement of 800–1000 km (Fig. 1) (Silver and Anderson, 1974; Anderson et al., 1979; Anderson and Schmidt, 1983; Campa and Coney, 1983; Marzolf, 2003). The El Antimonio Group is interpreted as a sedimentary succession that records shallow to deep marine sedimentation in a fore-arc basin that developed adjacent to the southwestern
Figure 1. Location of outcrops of the Lower Jurassic Sierra de Santa Rosa Formation in northwestern Sonora, Mexico, and position of the Mojave-Sonora megashear.
margin of the United States (González-León et al., 2005). It is also inferred that it was thrust over the Caborca terrane, which was part of the North American craton at some time during the Middle Jurassic, just prior to, or contemporaneous with, displacement along the Mojave-Sonora megashear (Stanley and González-León, 1995; González-León et al., 2005). The Early Jurassic bivalve fauna of Sonora occurs in the Sierra de Santa Rosa Formation of the El Antimonio Group. Outcrops of this group, which encompasses the Upper Permian to Triassic Antimonio Formation, the Upper Triassic Rio Asunción Formation, and the Lower Jurassic Sierra de Santa Rosa Formation (Hardy, 1981; Lucas and Estep, 1999; González-León et al., 2005), are scattered in several ranges of northwestern Sonora. The most complete sections of the El Antimonio Group occur in the Sierra del Álamo and Sierra de Santa Rosa mountains (Fig. 1). Other bivalve-bearing sections of the Sierra de Santa Rosa Formation are incomplete and occur in the Pozos de Serna area and in the Sierra de la Jojoba. Sierra del Álamo Section The El Antimonio Group in the Sierra del Álamo is divided into 14 stratigraphic sequences, which are numbered from base upward and have been further subdivided into several units (Fig. 2). Whereas sequences I to VI represent the Antimonio Formation and sequences VII to IX are part of the Río Asunción Formation, sequences X to XIV compose the lower (late Hettangian to late Sinemurian) part of the Sierra de Santa Rosa Formation at that locality. Sequence IX of Rhaetian age and sequence X of latest Hettangian to earliest Sinemurian age are separated by an erosional unconformity that marks the Triassic-Jurassic boundary in this
Figure 2. Sections of the Sierra de Santa Rosa Formation in northwestern Sonora showing fossiliferous intervals with ammonites and bivalves.
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area (Fig. 2). Sequence X is 60 m thick; its base is erosional and composed of coarse-grained to pebbly sandstone that grades upward into mudstone, siltstone, and interbedded limestone of shallow marine origin. This sequence contains abundant ammonites that have been assigned to the Canadensis and Trigonatum Zones by Taylor et al. (2001). From this sequence, the following bivalves were collected within an area of a few hundred square meters around Universal Transverse Mercator (UTM) position 3399704 351128: Grammatodon sulcatus, Agerchlamys wunschae, Weyla unca, Protocardia truncata, Neocrassina gueuxi, Cardinia concinna, Cardinia? sp. A, and Lucinidae gen. et sp. indet. B. Sequence XI is 280 m thick and overlies sequence X with a covered contact (Fig. 2). The basal 20 m are composed of a clast-supported pebble conglomerate that grades upward into interbedded sandstone and siltstone of fluvial origin. The middle part of this sequence is intruded by a dioritic dike and its upper part consists of massive to finely laminated calcareous mudstone and siltstone with ammonites of early Sinemurian (Leslei Zone) age (Taylor et al., 2001). The bivalve fauna from the upper part of this sequence, collected from an area of a few meters around UTM 3398025 351313, includes Palaeoneilo elliptica, Parallelodon cf. hirsonensis, Grammatodon sulcatus, Bakevellia sp. A, Gervillella araucana, G. leesi, Mytiloperna? sp. A, Plicatula sp. A, Oxytoma cf. inequivalvis, Entolium corneolum, Agerchlamys wunschae, Weyla alata, W. bodenbenderi, W. titan, W. unca, Modiolus hillanus, Osteomya dilata, Frenguelliella poultoni, Prosogyrotrigonia sp. A, Protocardia truncata, and Neocrassina gueuxi. The 115-m-thick sequence XII starts at its base with a 5-m-thick pebble conglomerate that grades into coarse-grained, thick-bedded, cross-stratified sandstone and sandy coquina. The upper part of this sequence is composed of massive mudstone and siltstone with intercalations of thin-bedded bioclastic limestone, fine-grained sandstone, and thin volcanic ash tuff beds. The bioclastic limestone beds contain ammonites of the lower upper Sinemurian Carinatum and Jamesi Zones (Pálfy and González-León, 2000, Taylor et al., 2001) and the bivalves Weyla alata, W. unca, Protocardia luggudensis, and Neocrassina gueuxi. The remainder of the Sierra de Santa Rosa Formation at this locality is formed by sequences XIII and XIV. Sequence XIII is 465 m thick, and in its lower part is composed of thick beds of conglomerate, pebbly to coarse-grained sandstone, and minor interbeds of mudstone, siltstone, and tuff beds. Its middle and upper parts are composed of fine-grained sandstone and thin-bedded to laminated mudstone and siltstone with sedimentary and biogenic structures that indicate sedimentation in basinal settings with turbidite currents. Ammonites of these strata suggest a late Sinemurian age (Pálfy and González-León, 2000). Sequence XIV is 150 m thick and in its lower part is composed of very thick beds of conglomerate, sandstone, and siltstone and in its upper part of medium- to coarse-grained sandstone. No bivalves were found in these two sequences.
Sierra de Santa Rosa Section The Lower Jurassic strata that crop out in the Sierra de Santa Rosa mountains were assigned to the Sierra de Santa Rosa Formation by Hardy (1981). He divided this section of shallow- to deep-marine sedimentary strata into lower, middle, and upper members. The most complete thickness of the lower member is 353 m and was measured along section 3, which starts at UTM 3326285 430123 and ends at UTM 3326318 429239 (for location see also Damborenea and González-León 1997, fig. 3). The base of the Lower Jurassic section in this area is tectonically thrust onto a Precambrian metamorphic basement. The measured section of the lower member is composed of fine- to coarse-grained sandstone, minor conglomerate lenses, and dark-gray siltstone and mudstone in its lower part (Fig. 2). The middle part consists of fine-grained sandstone with interbedded laminated mudstone, and sandy limestone with bivalves and ammonites. Its upper part consists of interbedded fine- to coarse-grained sandstone, mudstone, and bioclastic, locally encrinitic limestone. From this lower member we have identified Weyla alata, W. bodenbenderi, Frenguelliella poultoni, Lucinidae gen. et sp. indet. A, and Lucinidae gen. et sp. indet. B. Ammonites reported by Pálfy and González-León (2000) indicate that the lower member belongs to the uppermost Sinemurian Harbledownense Zone. In section 3, the middle member is 160 m thick and gradationally overlies the lower member from UTM 3326318 429239 to 3326256 429037. It consists of thin- to medium-bedded calcareous mudstone and siltstone and interbedded sandy to silty, bioclastic limestone. This member is also well exposed along section 4 (UTM positions 3327598 435179 to 3327489 434972) in this area (for location see also Damborenea and GonzálezLeón 1997, fig. 3). This member is the most fossiliferous part of the Sierra de Santa Rosa Formation at this locality. Ammonites most likely indicate an early Pliensbachian age for this member (Pálfy and González-León, 2000), and the identified bivalves include Parallelodon cf. hirsonensis, Grammatodon sulcatus, Gervillella araucana, G. leesi, Cercomya peruviana, Plagiostoma schimperi, Plagiostoma sp. A, Pinna cf. folium, Antiquilima cf. nodulosa, Ctenostreon sp. A, Entolium corneolum, Agerchlamys wunschae, Eopecten velatus, Weyla alata, W. titan, W. unca, Modiolus giganteus, M. cf. baylei, Pholadomya fidicula, Ph. idea, Pachymya? sp. A, Goniomya sp. A, Pleuromya uniformis, Ceratomya concentrica, Frenguelliella poultoni, Protocardia striatula, P. luggudensis, Neocrassina gueuxi, Myoconcha neuquena, Lucinidae gen. et sp. indet. A, Lucinidae gen. et sp. indet. B, and Isocyprina ancatruzi. Pozos de Serna Section At the Pozos de Serna locality, the Sierra de Santa Rosa Formation is not complete (Figs. 1 and 2) as it is faulted against Proterozoic strata. The lower part consists of interbedded sandstone and conglomerate, which is ~100 m thick and is followed by
Early Jurassic bivalves of the Antimonio terrane a covered interval ~100 m thick. Above is a 50-m-thick package of thin-bedded, fine-grained, calcareous sandstone that is overlain by 75 m of calcareous shale with ammonites and bivalves and a 30 m interval of reddish-brown, cross-stratified, sandy limestone beds with interbedded siltstone. The 180-m-thick middle part of the formation is dominated by thin-bedded to laminated calcareous shale with rare thin interbeds of dark-colored, micritic limestone where an interval with ammonites is present. The upper part is formed by a 150-m-thick section composed of fine- to medium-grained, thin- to medium-bedded calcareous sandstone with interbedded thick beds of sandy limestone and minor intervals of blue shale. An assemblage of ammonites from the lowermost fossiliferous interval indicates an early Pliensbachian age (Linares et al., 1997; Pálfy and González-León, 2000), and bivalves from the same interval, collected from an area of several hundred square meters around UTM 3336001 385940, include Grammatodon sulcatus, Gervillia (Cultriopsis) sp. A, Pinna cf. folium, Plicatula sp. A, Pseudolimea? sp. A, Agerchlamys wunschae, Weyla alata, W. bodenbenderi, Modiolus giganteus, Ceratomya petricosa, Frenguelliella poultoni, Groeberella sp. A, Protocardia sp. A, Cardinia sp. A, and Neocrassina gueuxi. Sierra de la Jojoba Section The Sierra de la Jojoba section, which is considered to be part of the Sierra de Santa Rosa Formation, is 110 m thick and is tectonically placed between Proterozoic and Jurassic? strata (Figs. 1 and 2). It consists of massive calcareous shale and siltstone and interbedded medium- to thick-bedded sandy limestone and calcareous sandstone. The bioclastic limestone contains bivalves (Weyla alata, W. titan, and Pholadomya cf. voltzi) and ammonites of probably early Toarcian age (Pálfy and GonzálezLeón, 2000).
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The original of this catalog is housed at the GeoBio-Center in Munich, but copies also exist in Würzburg and Berlin. The classification of higher taxa follows the schemes of Amler et al. (2000) and Carter (1990). The synonymy lists contain the first reference of the particular taxon and focus on literature records from North and South America or refer to recently published synonymy lists. If a taxon was hitherto unknown from America, the most important references from regions outside America are given. Concerning open nomenclature, the recommendations of Bengtson (1988) were applied. Taxa represented by at least several well-preserved specimens were measured with a sliding caliper. Measured dimensions, together with certain morphological features of some taxa, are given in Figure 3. For the multivariate analyses of paleobiogeographic affinities, we used SPSS 12.0 for Windows. We compiled a presence-absence data set containing all known pectinoid bivalves of Sonora and six other regions. These regions are northern Chile (from 22° to 31°S present-day latitude), central Chile and Argentina (from 31° to 41°S present-day latitude; regionally known as the Neuquén Basin), and, from western North America, the Brooks-Mackenzie Basin (northern and central Yukon and adjacent Northwest Territories and Alaska), the Western Canada Sedimentary Basin (central and southern Canadian Rocky Mountains and foothills of Alberta and British Columbia), and the allochthonous Canadian terranes of Stikinia and Wrangellia (both British Columbia).
MATERIAL AND METHODS This study is based on more than 500 specimens of Early Jurassic bivalves, which were collected by two of the authors (MA and CMG-L) in 1997 and 1998 in Sonora, Mexico. All figured specimens are deposited at the Estación Regional del Noroeste, Instituto de Geología, Universidad Nacional Autónoma de México in Hermosillo (ERNO). Additional material is housed in the collections of the Museum für Naturkunde, Berlin (MB.M.). The material is preserved mostly as external and internal molds, and only a small number of specimens show shell preservation. These shells are often abraded or incompletely preserved. Most specimens were prepared mechanically and latex casts were taken from external molds. Many specimens were identified on the basis of the relevant literature on bivalves from western North and South America. For the identification of species that had not yet been previously described from these regions, the Jurassic bivalve catalog proved very helpful; it contains photocopies of ~85% of Jurassic bivalves figured in the literature including an indication of age and region.
Figure 3. Measured dimensions and morphological features of selected bivalve taxa: A—Modiolus, B—Parallelodon, C, D—trigoniid bivalves, E—bakevelliid bivalves, F—Weyla.
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We analyzed the data set separately for two time slices, Hettangian–Sinemurian and Pliensbachian. Hettangian and Sinemurian bivalves are analyzed together because, in Sonora, several taxa are from the Canadensis Zone of unit 16 of the Sierra del Álamo section. This zone comprises late Hettangian and early Sinemurian sediments, but the Hettangian-Sinemurian boundary cannot be defined precisely. A hierarchical cluster analysis was used in order to find regions with similar faunal composition. The procedure of unweighted pair-group method using arithmetic average (UPGMA) was used herein with the Jaccard coefficient as the distance measure. In addition, we calculated the Dice coefficient to clarify the relations between Sonora, Argentina, and Chile. SYSTEMATIC PALEONTOLOGY Class Bivalvia Linné, 1758 Subclass Palaeotaxodonta Korobkov, 1954 Order Nuculoida Dall, 1889 Family Malletiidae H. Adams and A. Adams, 1858 Genus Palaeoneilo Hall and Whitfield, 1873 Type species. Nuculites constricta Conrad, 1842, subsequent designation by Hall, 1885. Palaeoneilo elliptica (Goldfuss, 1837) Figs. 4A and 4B. 1836 Nucula striata Lamarck—Roemer: 99, pl. 6, fig. 11. 1837 Nucula elliptica sp. nov.—Goldfuss: 153, pl. 124, figs. 16a–e. 1942 Nucula patagonidica sp. nov.—A. Leanza: 151, pl. 1, figs. 1–4. 1987a Palaeoneilo patagonidica (A. Leanza 1942)—Damborenea: 54, pl. 1, figs. 1–3; text-fig. 6. 1987a Palaeoneilo galatea (d’Orbigny 1850) ?—Damborenea: 56, pl. 1, fig. 4; text-fig. 7. 2000 Palaeoneilo elliptica (Goldfuss, 1837)—Hodges: 28, pl. 2, figs. 1–30; text-figs. 25–33 (see for extensive synonymy list). Material: One internal mold of a left valve (ERNO-8151) and one internal mold of a right valve (ERNO-8152) from the lower Sinemurian, unit 18, Sierra del Álamo. Description: Palaeoneilo elliptica is medium-sized, slightly inequilateral, equivalved, gently inflated, and subelliptical in outline. Its length is about twice its height. The anterior margin is convex. The ventral margin is slightly convex and nearly parallel to the slightly convex dorsal margin. The posterior margin is well rounded. The umbones are slightly prominent, prosogyrate, and are situated between one-third and one-half of the shell length from the anterior end. The taxodont hinge has about ten teeth anteriorly and about 18 teeth posteriorly, which decrease in size from both ends toward the umbo. As is typical of the genus, they are chevron-shaped toward the umbones. The internal mold of the left valve shows tear-shaped muscle scars, the posterior one being larger than the anterior one.
Remarks: This is the first record of the genus from Mexico. The Mexican material strongly resembles Early Jurassic specimens from Argentina that have been described by Damborenea (1987a) as Palaeoneilo galatea (d’Orbigny, 1850) ? and P. patagonidica (A. Leanza, 1942). Both have been included into P. elliptica by Hodges (2000), a view that is followed here. In his detailed examination of P. elliptica, Hodges (2000) analyzed several samples with a large number of specimens from the Lower Lias of southwest Britain, and found a wide range of morphological variability. P. galatea from the Middle Lias of Europe can be distinguished from P. elliptica by having an angulate posterior margin (Hodges, 2000). Subclass Pteriomorphia Beurlen, 1944 Order Arcoida Stolizka, 1871 Family Parallelodontidae Dall, 1898 Subfamily Parallelodontinae Dall, 1898 Genus Parallelodon Meek and Worthen, 1866 Type species. Macrodon rugosus Buckman, 1845. Parallelodon cf. hirsonensis (d’Archiac, 1843) Figs. 4E and 4F. cf. 1843 Cucullaea hirsonensis sp. nov.—d’Archiac: 374, pl. 27, figs. 5, 5a. cf. 1969 Parallelodon hirsonensis (d’Archiac)—Fischer: 78, pl. 9, fig. 3a–b (see for extensive synonymy list). cf. 1994 Parallelodon hirsonensis (d’Archiac 1843)—Aberhan: 12, pl. 1, figs. 15, 17–20; text-fig. 5. Material: One internal mold of a left valve (MB.M.4501) from the lower Sinemurian, unit 18, Sierra del Álamo; two internal molds of left valves and one steinkern (ERNO-8155 to 8156, MB.M.4502) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: Parallelodon cf. hirsonensis is mediumsized, equivalved, inequilateral, elongated, and subrectangular in outline. The shell is moderately inflated. Its length is about
Figure 4. (A–B) Palaeoneilo elliptica (Goldfuss, 1837). (A) internal mold of a left valve, × 4, ERNO-8151. (B) internal mold of a right valve, × 4, ERNO-8152; all lower Sinemurian of Sierra del Álamo. (C–D) Grammatodon (Grammatodon) sulcatus Aberhan, 1994. (C) internal mold of a right valve, × 2, ERNO-8153. (D) Latex cast of a left valve, × 2, ERNO-8154; all lower Sinemurian of Sierra del Álamo. (E–F) Parallelodon cf. hirsonensis (d’Archiac, 1843). (E) internal mold of a left valve, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO-8155. (F) internal mold of a left valve, lower Pliensbachian of Sierra de Santa Rosa, × 1.5, ERNO-8156. (G) Bakevellia sp. A; internal mold of a left valve, lower Sinemurian of Sierra del Álamo, × 1.5, ERNO-8157. (H–J) Gervillella araucana Damborenea, 1987; (H) left valve, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO-8158. (I) left valve, lower Sinemurian of Sierra del Álamo, × 1, ERNO-8159. (J) internal mold of a left valve, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO-8160. (K–L) Gervillella leesi Aberhan and Muster, 1997. (K) articulated specimen, left valve view, lower Sinemurian of Sierra del Álamo, × 1, ERNO-8161. (L) articulated specimen, left valve view, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO 8162.
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twice its height. The hinge line is straight, but not as long as the total length. The posterior margin is convex and meets with the hinge margin at an obtuse angle. The ventral margin is nearly parallel to the dorsal margin, but is slightly concave mesially. The anterior margin is short and convex. The umbo is prosogyrate, wide, and low, and is placed about one-fourth of the total shell length from the anterior end. A weak umbonal ridge runs from the umbo to the postero-ventral corner. The surface of the shell is smooth except for commarginal growth lines. The ligament area has the shape of an elongated triangle. Remarks: This is the first record of the genus from the Lower Jurassic of Mexico. In overall shape and outline, P. cf. hirsonensis has a great similarity to the type material of P. hirsonensis d’Archiac (1843, pl. 27, figs. 5, 5a). As the Mexican material is only moderately well preserved, and important characters such as hinge features are not known, it is only referred to P. hirsonensis with reservation. Nevertheless, with its rectangular outline and the obtuse postero-dorsal angle, it shows typical features of P. hirsonensis as described by Aberhan (1994). Damborenea (1987a, p. 66, pl. 1, figs. 10–12) described Parallelodon sp. from the Pliensbachian of Argentina, which is very similar to P. hirsonensis, but has a wider hinge plate and more numerous teeth than P. hirsonensis described by Aberhan (1994). Genus Grammatodon Meek and Hayden, 1861 Type species. Arca (Cucullaea) inornata Meek and Hayden, 1859, by original designation. Subgenus Grammatodon Grammatodon (Grammatodon) sulcatus Aberhan, 1994 Figs. 4C and 4D. 1994 Grammatodon (Grammatodon) sulcatus sp. nov.— Aberhan: 13, pl. 1, figs. 21–24; text-fig. 7. Material: Internal molds of one right valve and one left valve (MB.M.4503 to 4504) from the upper Hettangian–lower Sinemurian, unit 16, three external and two internal molds of right valves, and six internal and four external molds of left valves (ERNO-8153 to 8154, MB.M.4506 to 4518) from the lower Sinemurian, unit 18, Sierra del Álamo; one internal and one external mold of left valves (MB.M.4521 to 4522) from the lower Pliensbachian, Pozos de Serna; one internal mold of a right valve (MB.M.4523) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: Grammatodon (G.) sulcatus is small and inequilateral and has a subtrapezoidal outline. Its height is about the half of the total length. The umbones are situated at about one-third of the shell length from the anterior, and are broad and slightly prosogyrate. The dorsal margin is straight and extends along one-third of the shell length. The posterior margin is obliquely truncated, forming an obtuse angle with the dorsal margin and a right angle with the convex ventral margin. The anterior margin is straight, meets the ventral margin also at a right angle, and is slightly projecting beyond the anterior end of the hinge line. A relatively sharp, straight to slightly curved
umbonal carina extends from behind the umbo to the posteroventral corner of the shell; the carina becomes less sharp during ontogeny. In right valves, anterior to the carina, there is a shallow sulcus, which is very pronounced near the umbo but becomes less prominent and finally disappears toward the ventral margin. The left valve lacks this sulcus. The whole surface of the shell is covered with smooth commarginal growth lines. The postero-dorsal part of the right valve wears numerous radial riblets, which extend to the carina. One to two of these riblets are slightly more prominent than the others. The interspaces are about half as wide as the riblets. On the flanks and the sulcus of the right valve only commarginal growth lines can be seen. The anterior part of both valves also wears faint, fine radial striae, which are more pronounced than the commarginal growth lines. The preservation of the material does not allow an accurate description of the hinge. Some specimens show several small anterior teeth, which are slightly oblique near the umbo and become more horizontal toward the anterior end. Some other specimens possess two to three posterior elongated pseudolaterals, which are oriented in a horizontal to subhorizontal direction. Other internal structures are unknown. Remarks: This is the first record of the species outside Chile. G. sulcatus was first described by Aberhan (1994) from the upper Sinemurian of northern Chile. The specimens described herein agree well with the material from Chile. Only the radial components of the ornamentation of the shell are more pronounced in the Chilean material. Grammatodon (G.) concinnus (Phillips, 1829), which was documented by Aberhan (1994, p. 12, pl. 2, figs. 1–5) from the Lower Jurassic of Chile, can be distinguished from G. sulcatus by its greater inflation and the lack of a sulcus on the right valve. Also similar is Grammatodon (G.) costulatus (A. Leanza, 1942, p. 152, pl. 1, figs. 5–6; Damborenea, 1987a, p. 69, pl. 2, figs. 14–17; text-fig. 16b) from the Lower Jurassic of Argentina. Unlike G. sulcatus, the whole shell of G. costulatus is covered with regular radial riblets, the interspaces being twice as wide as the ribs. Order Pterioida Newell, 1965 Family Bakevelliidae King, 1850 Genus Bakevellia King, 1848 Type species. Avicula antiqua Münster, 1836, by subsequent designation (King 1850). Bakevellia sp. A Fig. 4G. Material: Two internal molds of left valves (ERNO-8157, MB.M.4524) from the lower Sinemurian, unit 18, Sierra del Álamo. Description: Bakevellia sp. A has a rhombic outline, is well inflated and has an anterior auricle and a posterior wing. The umbo is slightly prosogyrate and projects somewhat above the hinge line. The shell is slightly longer than high. The ratio of the length of the diagonal to the maximum width
Early Jurassic bivalves of the Antimonio terrane of the main body of the shell is about 2:1. The hinge line is half the length of the total shell length. The posterior margin has a sigmoidal shape, being concave at the posterior wing and turning highly convex near the posterior end of the shell. The ventral margin is evenly convex and passes gradually into the anterior margin. The anterior auricle is small, narrow, and rounded, and extends along half of the anterior margin. It is separated from the main body of the shell by a shallow groove. The posterior wing is nearly right-angled and pointed, and is separated from the disc of the shell by a wide sulcus. One elongated posterior tooth, which runs parallel to the hinge margin, is developed in both specimens, a juvenile and an adult. Other internal features, such as anterior teeth or muscle scars, are unknown. The surface of the shell is covered with commarginal growth lines. Remarks: This is the first Lower Jurassic record of the genus from Mexico. Bakevellia sp. A is similar to Bakevellia (B.) waltoni (Lycett, 1863), which is a common species in the Lower Jurassic of North and South America (Aberhan, 1994, p. 16, pl. 2, figs. 10–14; text-fig. 8; 1998a, p. 71, pl. 2, fig. 17, pl. 3, figs. 5–7; Aberhan and Muster, 1997, p. 801, textfig. 2D–G). Similar to Bakevellia sp. A, B. waltoni also has a rhombic shape and one (to two) elongated posterior teeth. As the presence of anterior teeth in Bakevellia sp. A cannot be confirmed with the material at hand, a comparison with the anterior teeth of B. waltoni cannot be performed. In addition, Bakevellia sp. lacks some of the prominent features that are significant of B. waltoni, in particular the presence of an acute anterior auricle and a sharply pointed, elongated posterior wing. Among other species of Bakevellia that are similar to Bakevellia sp. A, Bakevellia (B.) nana Fürsich and Werner (in Muster, 1995) from the Upper Jurassic of Portugal has two to three posterior teeth rather than one. Bakevellia (B.) binneyi Brown also has a small and rounded anterior auricle, but two posterior teeth, and is known from the Upper Permian to the Lower Jurassic of Greenland, Europe, and Asia (see Muster, 1995). Gervillaria hartmanni (Münster in Goldfuss, 1835, p. 122, pl. 115, figs. 7a–f) is less curved, the posterior wing is separated from the disc of the shell by only a slight flexure, not a sulcus, and occasionally wears weak radial ribs on the shell. Genus Gervillella Waagen, 1907 Type species. Perna avicuolides J. Sowerby, 1814 (pl. 147, pl. 66), by subsequent designation (Cox 1940). Gervillella araucana Damborenea, 1987 Figs. 4H–4J. 1987b Gervillella araucana sp. nov.—Damborenea: 133, pl. 1, figs. 6–10; text-fig. 6. 1994 Gervillella araucana Damborenea 1987—Aberhan: 18, pl. 3, figs. 8–10. 1995 Gervillella araucana Damborenea 1987—Muster: 61, pl. 12, figs. 3–4; text-fig. 45.
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1997 Gervillella araucana Damborenea 1987—Aberhan and Muster: 807, text-figs. 4F–G, 5C. 1998a Gervillella araucana Damborenea 1987—Aberhan: 74, pl. 2, figs. 14–15. Material: One left valve and two internal molds of left valves (ERNO-8159, MB.M.4525 to 4526) from the lower Sinemurian, unit 18, Sierra del Álamo; one left valve and four internal molds of left valves (ERNO-8158, ERNO-8160, MB.M.4527 to 4529) from the lower Pliensbachian, middle member, section 3, Sierra de Santa Rosa. Description: Gervillella araucana is medium-sized, twisted, and has an elongated, lanceolate outline. The left valves are very globose and twisted anticlockwise in posterior view. The dorsal margin is straight and half as long as the total shell. The posterior margin is concave at the posterior wing and convex near the posterior end and therefore has a sigmoidal shape. The ventral and the anterior margins are evenly rounded. The height is about half of the shell length. The umbo is prosogyrate, acute, and anteriorly placed but not terminal. A relatively sharp carina leads from the umbo toward the postero-ventral corner of the shell. This carina is not at the greatest inflation of the shell, but lies dorsally of it, and is preceded ventrally by a shallow but distinct furrow. The anterior auricle is lobate, extends along the anterior margin, and is separated from the disc of the shell by a shallow sulcus on the left valve. The posterior wing is narrow, obtuse, and separated from the disc of the shell by a deep sulcus. The surface of the shell is covered with regular growth lines. On the anterior part of the shell, they become imbricate. Internal features are not preserved. Remarks: This is the first record of Gervillella from the Lower Jurassic of Mexico. G. araucana can be distinguished from other species of Gervillella by its highly twisted shell and the presence of a carina on the left valve. Gervillella leesi Aberhan and Muster, 1997 Figs. 4K and 4L, 5A–5C, 6A. 1934 Gervillia sp. nov.—Lees: 41, pl. 4, fig. 1. 1997 Gervillella leesi sp. nov.—Aberhan and Muster: 805, text-figs. 4A–E, 5D. 1998a Gervillella leesi Aberhan and Muster 1997—Aberhan: 74, pl. 3, figs. 8–11; pl. 4, figs. 1; text-fig. 4. Material: One articulated specimen (ERNO-8161) from the lower Sinemurian, unit 18, Sierra del Álamo; one articulated specimen, one external and two internal molds of left valves, and one steinkern (ERNO-8162, MB.M.4530 to 4533) from the lower Pliensbachian, middle member, section 3, and one articulated specimen, two external and three internal molds of left valves, and four steinkerns (ERNO-8163 to 8165, MB.M.4534 to 4540) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: Gervillella leesi is large-sized, thick-shelled, nearly equivalved, with the left valve slightly more inflated than the right valve. The shape is elongated, slightly curved
Figure 5. (A–C) Gervillella leesi Aberhan and Muster, 1997. (A) articulated specimen, right valve view, × 1, ERNO-8163; (B) Steinkern, left valve view, × 1, ERNO-8164; (C) articulated specimen, left valve view, × 1, ERNO-8163; all lower Pliensbachian of Sierra de Santa Rosa.
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upward posteriorly, but not twisted. The anterior auricle is lobate, small relative to the overall size of the specimens, and separated from the main body of the shell by a shallow sulcus. The posterior wing is large and obtuse. A cross section through the body of the shell at about mid-length is rhombic. The umbones are prosogyrate, pointed, and low. The shell carries two carinae. One carina leads from below the umbo to the postero-ventral corner of the shell. The other one originates behind the umbo below the posterior wing and separates the main body of the shell from the posterior wing. Between the two carinae runs a shallow sulcus. In mediumsized specimens, the surface of the shell is covered by numerous regular growth lines that bend where they meet the carinae. Large-sized specimens develop conspicuous, strong knobs and wrinkles that follow the growth lines between the two carinae, but cross them in nearly right angles on the antero-ventral part of the shell and on the posterior part of the posterior wing. The ligament is multivincular. One of the specimens shows five ligamental pits. Anteriorly, the interspaces between the pits are narrower than the pits, but get wider in the posterior half of the ligamental area. Other internal features are unknown. Remarks: This is the first record of Gervillella leesi outside western Canada. It was first described by Aberhan and Muster (1997) from the Hettangian and upper Sinemurian of western Canada. Other Lower Jurassic carinate bakevelliids are Gervillella araucana and Gervillaria pallas (A. Leanza, 1942, p. 155, pl. 4, fig. 1). G. araucana, which is also described herein, is twisted, much more elongated, and wears only one carina. G. pallas, which is known from the Lower Jurassic of Argentina, Chile, and Canada (see Aberhan and Muster, 1997, for synonymy list), also is inequivalved and twisted, and has a pointed posterior wing. Although incomplete, the majority of the Mexican material is represented by very large specimens, reaching an estimated length of nearly 20 cm. A very conspicuous feature of these large specimens is well-developed knobs and wrinkles on the shell exterior. We interpret this feature as a phenomenon typical of adult and gerontic growth stages of G. leesi. This kind of ornamentation is very similar to that observed in fully grown specimens of another bakevelliid, Gervilleioperna (Gervilletia) turgida (A. Leanza, 1942), in which transverse, broad, and irregular rugae, which in some instances become true nodes, develop on the upper and lower carinae of the shell.
Description: Gervillia (Cultriopsis) sp. A is small, arched and only slightly inflated. It has a very slender, elongated, ensiform outline, with its height being about one-third of its length (H = 8 mm, L = 26 mm). The ratio of the length of the posterior wing to the length of the diagonal is 1:5; the ratio of the length of the diagonal to the maximum width of the main body of the shell is 5:1. A faint carina runs from behind the umbo toward the postero-ventral corner. The umbo is pointed, orthogyrate, and protrudes slightly beyond the hinge line. The anterior auricle is very small, pointed, and forward-turned; the posterior wing is relatively small (L = 5 mm), acute-angled in a posterior direction, and separated from the disc of the shell by a shallow sulcus. The internal mold shows the imprint of two posterior lateral teeth that are subparallel to the dorsal margin. Anterior teeth cannot be observed. The ligament and muscle scars are unknown. The shell surface is smooth and wears only commarginal growth lines. Remarks: This is the first documentation of the subgenus Cultriopsis from the Lower Jurassic of Mexico. Species of Cultriopsis have been revised by Muster (1995). The Mexican specimen has closest affinities to Gervillia (Cultriopsis) dundriensis Cox, 1946, and Gervillia (Cultriopsis) northamptonensis Cox, 1946. The former agrees well in size and outline, especially in the ratio between the length of the posterior wing and the diagonal (1:5), but lacks an anterior auricle, the umbo is terminal, and the posterior wing is obtuse. G. (C.) northamptonensis is similar in overall shape, but has no teeth and no carina. Another very similar species is Gervillia olifex Quenstedt (1856: 86, pl. 11, figs. 4–5) from the Lower Jurassic of Dusslingen, southern Germany, which can be distinguished by a larger ratio (~1:3) of the length of the posterior wing to the diagonal, and by a slightly larger width (ratio of diagonal to width 4:1). Damborenea (1987b, p. 130) also mentioned G. olifex and noted a similarity to Gervillia cf. angusta von Münster described by Escobar (1980, p. 41, pl. 2, fig. 1d) from the lower Sinemurian of El Culebreada, Chile, which has an obtuse posterior wing and a ratio of the length of the posterior wing to the diagonal of the shell of ~1:3. Damborenea (1987b, p. 131, pl. 3, figs. 4–6; text-fig. 5) also described Gervillia (Cultriopsis) sp. from the Toarcian of Argentina, which differs by the absence of an anterior auricle and the shape of the posterior wing.
Genus Gervillia Defrance, 1820 Type species. Gervillia solenoidea Defrance, 1824, subsequent monotypy. Subgenus Cultriopsis Cossmann, 1904 Type species. Gervillia (Cultriopsis) falciformis Cossmann, 1904, original designation. Gervillia (Cultriopsis) sp. A Fig. 6B. Material: One internal mold of a left valve (ERNO-8166) from the lower Pliensbachian, Pozos de Serna.
Family Isognomonidae Woodring, 1925 Genus Mytiloperna von Ihering, 1903 Type species. Perna americana Forbes in Darwin, 1846. Mytiloperna? sp. A Fig. 6C. Material: One external mold of a left valve (ERNO-8167) from the lower Sinemurian, unit 18, Sierra del Álamo. Description: Mytiloperna? sp. A is subrectangular, slightly inflated, with a terminal umbo. A conspicuous carina extends from the umbo to the antero-ventral corner of the shell. The
Early Jurassic bivalves of the Antimonio terrane part anterior of this carina is slightly concave; the posterior part is slightly convex. The dorsal margin is straight and longer than half the total shell length. The posterior and the ventral margins are slightly convex. The anterior margin is slightly concave. The posterior wing is undifferentiated. The surface of the shell is covered with weak, regularly spaced growth lines. Internal features are not preserved. Remarks: Internal features such as the extension of the ligamental area and the presence or absence of posterior hinge teeth are needed to distinguish Mytiloperna from the externally very similar bakevelliid genus Aguilerella. In particular, A. kobyi (Loriol) is similar in overall shape to the Mexican specimen (e.g., see Muster 1995, p. 15, pl. 1, figs. 1–4, text-fig. 8). Nevertheless, the outline of the specimen also strongly resembles that of other species of Mytiloperna, for example Mytiloperna mytiliformis (Schlippe) (e.g., Ma et al., 1976, p. 298, pl. 31, figs. 7–9). The type species of Mytiloperna, M. americana (Forbes) from the Lower Jurassic of Chile (Forbes, 1846: 266, figs. 4–6; Philippi, 1899, p. 45, pl. 22, fig. 8; Ihering, 1903, p. 123, fig. 1), is very large and has a mytiliform rather than subrectangular outline. From North America, only one species of Mytiloperna has been described until now, i.e., Mytiloperna charlottensis Aberhan (1998a, p. 75, pl. 4, figs. 7–10; text-fig. 5) from the Pliensbachian of the Queen Charlotte Islands, western Canada. In contrast to the subrectangular outline of the Mexican
Figure 6. (A) Gervillella leesi Aberhan and Muster, 1997; Latex cast of an external mold of a left valve, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO-8165. (B) Gervillia (Cultriopsis) sp. A; internal mold of a left valve, lower Pliensbachian of Pozos de Serna, × 2, ERNO-8166. (C) Mytiloperna? sp. A; latex cast of an external mold of a left valve, lower Sinemurian of Sierra del Álamo, × 2, ERNO-8167. (D–F) Cercomya (Capillimya) peruviana Cox, 1956; composite mold of an articulated specimen; (D) dorsal view; (E) left valve view; (F) right valve view, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO-8168. (G, H) Pinna (Pinna) cf. folium Young and Bird, 1822. Steinkern; (G) right valve view; (H) left valve view, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO-8169. (I, J) Antiquilima (Antiquilima) cf. nodulosa Hayami, 1959; (I) internal mold of a left valve, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO-8170; (J) latex cast of an external mold of a right valve, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO-8171. (K) Ctenostreon sp.; composite mold of a right valve, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO-8172. (L) Pseudolimea? sp. A; composite mold of a right valve, lower Pliensbachian of Pozos de Serna, × 2, ERNO-8173. (M, N) Plagiostoma schimperi (Branco, 1879); (M) latex cast of an external mold of a left valve, × 1, ERNO-8174; (N) internal mold of a left valve, × 1, ERNO-8175; all lower Pliensbachian of Sierra de Santa Rosa. (O, P) Plagiostoma sp. A. (O) composite mold of a right valve, × 1, ERNO-8176; (P) internal mold of a left valve, × 1, ERNO-8177; all lower Pliensbachian of Pozos de Serna. (Q) Plicatula (Plicatula) sp. A; latex cast of an internal mold of a left valve, lower Sinemurian of Sierra del Álamo, × 2, ERNO-8178. (R) Oxytoma (Oxytoma) cf. inequivalvis (J. Sowerby, 1819); latex cast of an external mold of a left valve, lower Sinemurian of Sierra del Álamo, × 2, ERNO-8179.
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specimen, however, this species is subtrigonal in shape. If our generic identification can be confirmed when additional material becomes available, this would be the first record of the genus from Mexico. Family Pinnidae Leach, 1819 Genus Pinna Linné, 1758 Type species. Pinna rudis Linné, 1758, subsequent designation by Children, 1823. Subgenus Pinna Pinna (Pinna) cf. folium Young and Bird, 1822 Figs. 6G and 6H. cf. 1822 Pinna folium sp. nov.—Young and Bird: 243, pl. 10, fig. 6. 1987a Pinna (Pinna) cf. folium Young and Bird 1822—Damborenea: 95, pl. 4, figs. 6, 11a–b, 12a–b, 13–14; text-fig. 24 (see for extensive synonymy list). 1994 Pinna (Pinna) cf. folium Young and Bird 1822—Aberhan: 22, pl. 7, figs. 3–4. 1997 Pinna sp.—Damborenea and González-León: 183, fig. 5.1. 1998a Pinna (Pinna) cf. folium Young and Bird 1822—Aberhan: 80, pl. 5, figs. 10–12; pl. 6, fig. 1. Material: One fragmentary valve (MB.M.4544) from the lower Pliensbachian, middle member, section 3, one external mold and five partially fragmented steinkerns (ERNO-8169, MB.M.4545 to 4549) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa; one fragmented steinkern with remains of shell (MB.M.4551) from the lower Pliensbachian, Pozos de Serna. Description: Pinna (P.) cf. folium is medium-sized and has a cuneiform outline. The apical angle is very acute (about 30°). The dorsal margin is straight to slightly convex, the ventral margin is slightly concave. The posterior margin is straight and slightly rounded, and the posterior end has a wide gape. The medium carina is strong. In some well-preserved specimens, a faint furrow is sitting on the medium carina. The shell is ornamented by about eight radial ribs, which cover the area dorsal to the medium carina. The shell also is covered with smooth growth lines and a large number of commarginal folds. These folds intersect with the radial ribs and form a bulging reticulate pattern. The cross section of the shell is rhombic with rounded edges. Remarks: P. cf. folium is a common element of the Lower Jurassic of North and South America with occurrences reported from Argentina, Chile, British Columbia, and Alberta (Damborenea, 1987a; Aberhan, 1994, 1998a). The material studied by Damborenea and González-León (1997) from the same locality in Mexico (section 4, Sierra de Santa Rosa) and described as “Pinna sp.” resembles the specimens described herein and is included in P. cf. folium. We regard the somewhat larger apical angle of about 40° as part of the intraspecific variation of the species. An extensive comparison with similar species of Pinna is given by Damborenea (1987a).
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Order Limoida Rafinesque, 1815 Family Limidae Rafinesque, 1815 Genus Antiquilima Cox, 1943 Type species. Lima antiquata J. Sowerby, 1818. Subgenus Antiquilima Antiquilima (Antiquilima) cf. nodulosa Terquem, 1855 Figs. 6I and 6J. cf. 1855 Lima nodulosa sp. nov.—Terquem: 322, pl. 22, fig. 3. 1929 Lima nodulosa Terqu.—Jaworski: 6, pl. 1, fig. 6. 1942 Lima succincta Schloth.—Leanza: 178, pl. 10, fig. 7. 1984 Antiquilima sp.—Manceñido and Damborenea: 425, pl. 1, fig. 18. 1959 Lima aff. compressa Terquem—Frebold: 10, pl. 4, fig. 3. 1994 Antiquilima (Antiquilima) cf. nagatoensis Hayami 1959—Aberhan: 22, pl. 8, figs. 1–3. 1998a Antiquilima (Antiquilima) cf. nagatoensis Hayami 1959—Aberhan: 81, pl. 7, figs. 9–10. Material: One external mold of a left valve (MB.M.4555) from the lower Pliensbachian, middle member, section 3, and one internal and four external molds of right valves, and one internal mold of a left valve (ERNO-8170 to 8171, MB.M.4553 to 4554, MB.M.4556 to 4557) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: Antiquilima (A.) cf. nodulosa is mediumsized, equivalved, inequilateral, and has an obliquely ovate, rather compressed, and only slightly inflated shell. It shows the typical ornamentation of sinuous radial ribs of two different orders of strength and commarginal growth rugae. The radial ribs carry imbricate lamellae on their crests, which may develop into spines in larger specimens and are conspicuously stronger on the primary than on the secondary ribs. The interspaces between the primary ribs are wider than the ribs and bear no or one to three intercalated secondary riblets, which are of different width and do not reach the same strength as the primary riblets. Remarks: This taxon is also known from the Sinemurian of northern Chile (Aberhan, 1994), the Lower Jurassic of Argentina (e.g., Lima succincta Schlotheim in A. Leanza, 1942) and the Sinemurian to Pliensbachian of western Canada (Aberhan, 1998a). In previous studies, Aberhan (1994, 1998a) tentatively assigned the material from South America and western Canada to A. nagatoensis Hayami, 1959, a species known from the Sinemurian of Japan. Assignment was with reservation because of differences in the length-height ratio. Furthermore, all the above mentioned material from America, including that described here as well as Lima nodulosa Terqu. of Jaworski (1929) from the same locality in Mexico, is characterized by the presence of imbricate lamellae on the crests of the ribs. This feature is less clearly developed in the holotype of A. nagatoensis, in which primarily the higher-order riblets appear to be jagged. For this reason, the American specimens are here referred to Antiquilima nodulosa (Terquem, 1855, p. 322, pl. 22, fig. 3 a–c, Dumortier, 1864, p. 57, pl. 8, figs. 6–8) from the Hettangian of France, with which they share the distinct lamellae on the crests of the radial ribs. Because of a
lower length-height ratio in the European material, assignment to this species is with reservation. A. cf. nodulosa can be distinguished from the closely related Antiquilima (A.) succincta (v. Schlotheim, 1813) by the presence of imbricate lamellae on the crests of the ribs (see also Aberhan, 1994, 1998a). Genus Ctenostreon Eichwald, 1862 Type species. Ostracites pectiniformis Schlotheim, 1820. Ctenostreon sp. A Fig. 6K. Material: One fragmented composite mold of a right valve (ERNO-8172) from the lower Pliensbachian, middle member, section 3, Sierra de Santa Rosa. Description: The single available specimen of Ctenostreon sp. A suggests a subcircular outline of the valve. The fragment carries ten strong, sinuous, rounded radial ribs. The interspaces vary in width but generally have the same width as the ribs. The whole surface of the shell is covered with fine commarginal, imbricate lamellae that form conspicuous knobs on the ribs. Remarks: The genus Ctenostreon was hitherto unknown from the Lower Jurassic of Mexico. Aberhan (1994, p. 23, pl. 8, figs. 4–7; 1998a, p. 84, pl. 7, figs. 1–5) described Ctenostreon cf. rugosum from the Lower Jurassic of Chile and western Canada, which is very similar in size and ornamentation, and considered the better known species Ctenostreon pectiniforme (Schlotheim) and C. proboscideum (J. Sowerby) as likely junior synonyms of C. rugosum. As the Mexican specimen is represented only by a fragment, an identification at species level is not possible at the moment. Genus Pseudolimea Arkell in Douglas and Arkell, 1932 Type species. Plagiostoma duplicata J. de C. Sowerby, 1827. Pseudolimea? sp. A Fig. 6L. Material: One composite mold of a right valve (ERNO-8173) from the lower Pliensbachian, Pozos de Serna. Description: Pseudolimea? sp. A is small, trigonal in outline and slightly inflated. It carries about 16 radial ribs, which have about the same width as their interspaces. Its anterior and posterior ends bear only weak riblets. Other structures such as internal characters cannot be observed. Remarks: Pseudolimea has not been reported from the Lower Jurassic of Mexico before. In the single, poorly preserved specimen at hand, the presence of secondary riblets cannot be verified. Therefore, the specimen is referred to Pseudolimea with reservation. Pseudolimea hettangiensis (Terquem, 1855), which has been described from the Hettangian of Chile (Aberhan, 1994, p. 27, pl. 11, fig. 6), has a slightly higher number of radial ribs (17 to 18), but additional material would be needed for a more detailed comparison. Genus Plagiostoma J. Sowerby, 1814 Type species. Plagiostoma giganteum J. Sowerby, 1814.
Early Jurassic bivalves of the Antimonio terrane Plagiostoma schimperi (Branco, 1879) Figs. 6M and 6N. 1879 Lima schimperi sp. nov.—Branco: 111, pl. 6, fig. 4. 1900 Lima (Plagiostoma) schimperi Branco—Greppin: 130, pl. 15, fig. 7; pl. 16, figs. 2, 5. 1929 Plagiostoma cf. exaltata Terqu.—Jaworski: 5, pl. 1, figs. 5a–b. 1936 Plagiostoma schimperi Branco—Dechaseaux: 22, text-fig. 6. 1943 Lima (Plagiostoma) schimperi Branco—Cox: 160, pl. 11, figs. 15–18, pl. 12, figs. 19–22. 1986 Plagiostoma schimperi (Branco)—Jaitly: 43, pl. 1, fig. 6. 1995 Plagiostoma schimperi (Branco 1879)—Jaitly et al.: 180, pl. 12, fig. 3. 1997 Plagiostoma cf. P. punctatum J. Sowerby, 1814—Damborenea and González-León: 188, fig. 5.2. Material: One external mold of a right valve, one external mold of a left valve, and two almost complete left valves (ERNO-8174 to 8175, BM.M.4558 to 4559) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: Plagiostoma schimperi is medium-sized, equivalved, inequilateral, and has an obliquely ovate, subtrigonal shell, with the ventral margin being well rounded and slightly asymmetrical. The posterior auricle is relatively large. Because preservation is incomplete, the size and outline of the anterior auricle are unknown. The surface of the shell is covered by numerous (50–70) faint radial riblets, which are separated by narrow, linear to undulating, fine, punctate grooves. Toward the ventral margin, the riblets become slightly stronger. They are of unequal width and have flat to slightly rounded tops. The riblets are crossed by irregularly spaced commarginal grooves. Very fine growth lines are visible on the posterior auricle, which also is ornamented with very fine radial threads. Remarks: P. schimperi is very common in the Middle Jurassic of England, France, Switzerland, and Germany (Cox 1943), but so far has not been described from North or South America. Damborenea and González-León (1997) figured one specimen from section 3 and referred to it as Plagiostoma cf. punctatum J. Sowerby 1815. This species has also been reported from the Lower Jurassic of Chile, but differs from P. schimperi by a very high number (100–140) of ribs (Aberhan 1994). Therefore, we include the specimen described by Damborenea and González-León (1997) in P. schimperi. Also from the same locality, Jaworski (1929) identified Plagiostoma cf. exaltatum Terquem. Whereas we include Jaworski’s record in P. schimperi, the European P. exaltatum has a larger size and more numerous ribs (~100) than the Mexican specimens. Plagiostoma sp. A Figs. 6O and 6P. Material: Two internal molds of left valves and one composite mold of a right valve (ERNO-8176 to 8177, MB.M.4563) from the lower Pliensbachian, Pozos de Serna.
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Description: Plagiostoma sp. A is medium-sized, equivalved, and has a conspicuously slender, obliquely ovate shape. It is distinctly higher than long. The anterior auricle is relatively large and obtuse. The ventral margin is well rounded and slightly asymmetrical. The posterior margin is relatively long and straight. The outline of the posterior auricle is unknown. The composite mold carries numerous faint, thin radial riblets, which are broad and flat-topped. The interspaces are narrower than the riblets. Remarks: Plagiostoma sp. A is very conspicuous because of its slender outline, which is much higher than long, a feature that separates this taxon from P. schimperi, which exhibits a more rounded shape and an almost equal length-height ratio. Plagiostoma bilibini Milova (1969, p. 177, pl. 1, fig. 11, pl. 2, fig. 3) from the Pliensbachian of Siberia has a similar outline and ornamentation, but is conspicuously larger than Plagiostoma sp. A. Lima ridigula Phillips, described by Lycett (1863, p. 42, pl. 33, fig. 7a) from the Bathonian of England, has a similar ornamentation, but differs in having a more pronounced anterior auricle. Plagiostoma northamptonensis Cox (1943, p. 165, pl. 16, figs. 35–37) is only slightly higher than long, and has a different ornamentation consisting of sinuous riblets of variable width. Order Ostreoida Férussac, 1822 Family Plicatulidae Watson, 1930 Genus Plicatula Lamarck, 1801 Type species. Spondylus plicatus Linné, 1758; subsequent designation by Schmidt, 1818. Subgenus Plicatula Plicatula (Plicatula) sp. A Fig. 6Q. Material: One internal mold of a left valve (ERNO-8178) from the lower Sinemurian, unit 18, Sierra del Álamo; one internal mold of a right valve (MB.M.4564) from the lower Pliensbachian, Pozos de Serna. Description: Plicatula (Plicatula) sp. A is small sized, higher than long, and only slightly inflated. In one of the internal molds, about four radial plicae can be seen at the ventral margin. Because the material consists only of internal molds, other external features are unknown. Both molds show a triangular resilifer, which is followed anteriorly and posteriorly by narrow ridges, sockets, and finally by prominent crura. The divergent crura are straight, narrow, and slightly crenulated. The adductor scar of the valves is relatively small, sub-orbicular in shape, and situated in the posterior part of the shell. Remarks: This is the first record of Plicatula from the Lower Jurassic of Mexico. Following the thorough comparison of the taxa Plicatula and Harpax by Damborenea (2002), the here described specimens can be placed in the subgenus Plicatula. The main differences between Plicatula and Harpax are the hinge region and the reversed relation of convexity of the valves (see Damborenea, 2002, for details). Plicatula
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(P.) armata Goldfuss, described by Aberhan (1994) from the upper Sinemurian of Chile, differs from Plicatula (P.) sp. A by having more radial plicae (9 to 25) and a more centrally placed adductor muscle scar. Order Pectinoida Rafinesque, 1815 Family Oxytomidae Ichikawa, 1958 Genus Oxytoma Meek, 1864 Type species. Avicula münsteri Bronn, 1830. Subgenus Oxytoma Oxytoma (Oxytoma) cf. inequivalvis (J. Sowerby, 1819) Fig. 6R. cf. 1819 Avicula inequivalvis sp. nov.—J. Sowerby: 78, pl. 244, figs. 2–3. cf. 1987b Oxytoma (Oxytoma) inequivalvis (J. Sowerby 1819)—Damborenea: 160, pl. 6, figs. 9–12; text-fig. 15 (see for synonymy list). cf. 1994 Oxytoma (Oxytoma) inequivalvis (J. Sowerby 1819)—Aberhan: 35, pl. 17, figs. 1–5. cf. 1998a Oxytoma (Oxytoma) inequivalvis (J. Sowerby 1819)—Aberhan: 95, pl. 9, figs. 8–14 (see for synonymy list of North American records). Material: One external mold of a left valve (ERNO-8179) from the lower Sinemurian, unit 18, Sierra del Álamo. Description: The fragmented Mexican specimen assigned to Oxytoma (O.) cf. inequivalvis is relatively small, weakly inflated, and ornamented with thin and straight primary radial ribs. In some of the interspaces second order riblets are developed. Remarks: O. (O.) inequivalvis is a common species with a cosmopolitan distribution, but has not been documented from the Lower Jurassic of Mexico before. It has been described in detail by Duff (1978) and Damborenea (1987b). Ornamentation of the Mexican specimen is very similar to that of O. inequivalvis, but due to poor preservation it is assigned to the latter taxon with reservation. Family Entoliidae von Teppner, 1922 Genus Entolium Meek, 1865 Type species. Entolium demissus Phillips, 1829. Subgenus Entolium Entolium (Entolium) corneolum (Young and Bird, 1828) Figs. 7A, 7B, and 8. 1828 Pecten corneolus sp. nov.—Young and Bird: 234, pl. 9, fig. 5. 1984 Entolium (Entolium) corneolum (Young and Bird 1828)—Johnson: 45, pl. 1, figs. 24–26, ?20, ?22, ?27 (see for extensive synonymy list). 1994 Entolium (Entolium) corneolum (Young and Bird 1828)—Aberhan: 37, pl. 17, figs. 15–19 (see for synonymy list of South American records). 1997 Entolium (Entolium) disciforme (Schübler in Zieten, 1833)—Damborenea and González-León: 190, figs. 5.3–5.5. 1998a Entolium (Entolium) corneolum (Young and Bird
1828)—Aberhan: 105, pl. 11, figs. 8, 12–14, 16, 17 (see for synonymy list of North American records). Material: One right valve (ERNO-8180) from the lower Sinemurian, unit 18, Sierra del Álamo; one articulated specimen, six right valves, three left valves, and three single valves (MB.M.4565 to 4577) from the lower Pliensbachian, middle member, section 3, and six right valves, two left valves, and two fragmented valves (ERNO-8181, MB.M.4579 to 4587) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. The specimens are preserved as molds and show only remains of shell preservation. Description: Entolium (E.) corneolum is medium-sized, sub-equilateral, and nearly equivalved. The disc is suborbicular and compressed. The surface of the shell is covered with regularly spaced, commarginal growth lines, and wears no other ornamentation. The auricles of the right valve extend beyond the hinge line dorsally, whereas the left valve auricles meet at an angle of about 180°. Remarks: E. corneolum differs from E. lunare (Roemer, 1839) by the lack of a byssal notch in early ontogenetic stages and from E. orbiculare (J. Sowerby, 1817) by the lack of commarginal grooves on the right valve (Johnson, 1984). The specimens from the Lower Jurassic of Mexico are indistinguishable from European representatives of this species. From Mexico, Damborenea and González-León (1997) described Entolium (E.) disciforme (Schübler in Zieten, 1833). This species was included in E. corneolum by Johnson (1984), and this view is followed here. Affinities to species of Entolium from the Jurassic of western North America were discussed by Aberhan (1998a).
Figure 7. (A, B) Entolium (Entolium) corneolum (Young and Bird, 1828); (A) latex cast of an external mold of a left valve, lower Sinemurian of Sierra del Álamo, × 1, ERNO-8180; (B) right valve, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO-8181. (C, D) Agerchlamys wunschae (Marwick, 1953); (C) latex cast of an external mold of a left valve, lower Sinemurian of Sierra del Álamo, × 1, ERNO-8182; (D) composite mold of a right valve, upper Hettangian– lower Sinemurian of Sierra del Álamo, × 2, ERNO-8183. (E) Eopecten velatus (Goldfuss, 1833); latex cast of an external mold of a left valve, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO-8184. (F–J) Weyla (Weyla) alata (von Buch, 1838); (F–H) articulated specimen; (F) right valve view; (G) dorsal view; (H) left valve view, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO-8185; (I) right valve, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO-8186; (J) external mold of a left valve, lower Sinemurian of Sierra del Álamo, × 1, ERNO-8187. (K) Weyla (Weyla) bodenbenderi (Behrendsen, 1891); latex cast of an external mold of a right valve, upper Sinemurian of Sierra de Santa Rosa, × 1, ERNO-8188. (L) Weyla (Weyla) titan (Möricke, 1894). External mold of a left valve, lower Toarcian of Sierra de La Jojoba, × 1, ERNO-8189. (M–O) Weyla (Weyla) alata (von Buch, 1838); (M) latex cast of an external mold of a right valve, × 1.5, ERNO-8190; (N) latex cast of an external mold of a left valve, × 1, ERNO-8191; (O) latex cast of an external mold of a left valve, × 1, ERNO-8192; all lower Sinemurian of Sierra del Álamo.
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Figure 8. Scatter diagram showing the length-to-height ratio of Entolium (Entolium) corneolum (Young and Bird, 1828).
Family Pectinidae Rafinesque, 1815 Genus Agerchlamys Damborenea, 1993 Type species. Chlamys (Camptochlamys) wunschae Marwick, 1953a. Agerchlamys wunschae (Marwick, 1953) Figs. 7C and 7D. 1942 Pecten (Chlamys) textorius Schloth. var. torulosa Quenst.—A. Leanza: 173, pl. 7, fig. 4. 1953 Chlamys (Camptochlamys) wunschae sp. nov.— Marwick: 98, pl. 10, figs. 23–24. 1993 Agerchlamys wunschae (Marwick)—Damborenea, figs. 4a–j. 1994 Agerchlamys wunschae (Marwick 1953)—Aberhan: 38, pl. 18, figs. 1–6. 1998a Agerchlamys wunschae (Marwick 1953)—Aberhan: 108, pl. 12, figs. 1–5. 2002 Agerchlamys wunschae (Marwick 1953)—Damborenea: 68, pl. 7, figs. 1–10; text-figs. 8l, 8q, 32, 45c (see for extensive synonymy list). Material: One right valve (ERNO-8183) from the upper Hettangian–lower Sinemurian, unit 16, one right valve and one left valve (ERNO-8182, MB.M.4590) from the lower Sinemurian, unit 18, Sierra del Álamo; one right valve (MB.M.4594) from the lower Pliensbachian, Pozos de Serna; one fragmented left valve (MB.M.4597) from the lower Pliensbachian, middle member, section 3, Sierra de Santa Rosa. Most specimens are preserved as molds, only few show shell preservation. Remarks: This is the first record of the genus from Mexico. A. wunschae has been extensively described recently by Damborenea (2002) and no further description is needed here. Similar to the Chilean specimens described by Aberhan (1994), the Mexican representatives lack the fine divaricate striae reported by Damborenea (1993, 2002) in specimens from the Pliensbachian of Argentina and New Zealand. Taylor and Guex (2002, p. 14, pl. 1, figs. 1–4, 8–11; pl. 4, figs. 1–3) described Agerchlamys boellingi, a new species
Genus Eopecten Douvillé, 1897 Type species. Hinnites tuberculatus Goldfuss, 1835 (errore pro Spondylus tuberculosus Goldfuss, 1835), original designation by Douvillé, 1897. Eopecten velatus (Goldfuss, 1833) Fig. 7E. 1833 Pecten velatus sp. nov.—Goldfuss: 45, pl. 90, fig. 2. 1984 Eopecten velatus (Goldfuss 1833)—Johnson: 150, pl. 5, figs. 4, 5, 7, 8; text-figs. 137–141 (see for extensive synonymy list). 1994 Eopecten velatus (Goldfuss 1833)—Aberhan: 41, pl. 21, figs. 2, 6–7 (see for synonymy of South American records). 2002 Eopecten cf. velatus (Goldfuss 1833)—Damborenea: 54, pl. 6, figs. 1–3. Material: Four partly fragmented left valves (ERNO-8184, MB.M.4608 to 4610) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. The specimens are preserved as internal and external molds, rarely with remains of shell attached. Description: Only left valves of Eopecten velatus were recovered. They are medium-sized, inflated, prosocline, and have a sub-orbicular outline. The umbo is placed at midlength. The anterior wing passes gradually into the disc of the shell, without any demarcation. The disc is covered with about 25 primary radial ribs, between which smaller second- and third-order riblets are intercalated. The ribs and riblets are narrow and have smooth tops. In one specimen, toward the ventral margin some second order riblets reach the same thickness as the primary ribs. The anterior wing is ornamented by numerous ribs of equal strength, which are not as prominent as the primary ribs on the disc. Remarks: Affinities to other species of Eopecten were established by Johnson (1984), Aberhan (1994) and Damborenea (2002) and no further comments are necessary here. These are the first finds of the genus from the Lower Jurassic of Mexico. Family Neitheidae Sobetzky, 1960 Genus Weyla J. Böhm, 1920 Type species. Pecten alatus von Buch, 1838. Subgenus Weyla Weyla (Weyla) alata (von Buch, 1838) Figs. 7F–7J, 7M–7O, and 9. 1838 Pecten alatus sp. nov.—v. Buch: 55. 1839 Pecten alatus sp. nov.—v. Buch: 3, pl. 1, figs. 1–3. 1929 Neithea mexicana sp. nov.—Jaworski: 2, pl. 1, figs. 1–3. 1987b Weyla mexicana (Jaworski) ?—Damborenea: 189, pl. 11, fig. 3; text-fig. 22d. 1987b Weyla (Weyla) alata alata (von Buch
Early Jurassic bivalves of the Antimonio terrane
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1838)—Damborenea: 170, pl. 9, fig. 1; pl. 10, figs. 1a–b, 2–3, 4a–c, 5; text-figs. 18, 19-1, 22a (see for extensive synonymy list). 1987b Weyla (Weyla) alata angustecostata (R. Philippi 1899)—Damborenea: 177, pl. 6, fig. 17b; pl, 7, figs. 1a–b, 2; pl. 8, figs. 1, 2a–b, 3a–b; pl. 9, figs. 2a–b, 3; text-figs. 2, 19-2, 21. 1997 Weyla—Lucas and Estep: 45, figs. 1A–B (only). 1994 Weyla (Weyla) alata (von Buch 1838)—Aberhan: 43, pl. 22, figs. 1–3; pl. 23, figs. 1–2. 1997 Weyla (Weyla) alata (von Buch, 1838)—Damborenea and González-León: 190, figs. 5.6–5.9. 1998a Weyla (Weyla) alata (von Buch 1838)—Aberhan: 119, pl. 15, figs. 3, 5–8; pl. 16, figs. 3, 12 (see for synonymy of North American records). Material: Fifteen right valves, 14 left valves, and one articulated specimen (ERNO-8187, ERNO-8190 to 8192, MB.M.4611 to 4636) from the lower Sinemurian, unit 18, and one left valve (MB.M.4637) from the upper Sinemurian, unit 19, Sierra del Álamo; two fragmented right valves and one left valve (MB.M.4638 to 4640) from the upper Sinemurian, lower member, section 3, Sierra de Santa Rosa; four right valves and three partly fragmented left valves (MB.M.4690 to 4696) from the lower Pliensbachian, Pozos de Serna; four right valves, three left valves, and one articulated specimen (MB.M.4641 to 4648) from the lower Pliensbachian, middle member, section 3, and 21 right valves, 16 left valves, and six articulated specimens (ERNO-8185 to 8186, MB.M.4649 to 4689) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa; one fragmented right valve and two fragmented left valves (MB.M.4697 to 4699) from the lower Toarcian, Sierra de La Jojoba. Preservation is as internal mold, as external mold, or with shell material. Remarks: For a full description of W. (W.) alata refer to Damborenea (1987b) and Aberhan (1994). Recently, Mexican specimens of W. alata were described by Damborenea and González-León (1997). These authors concluded that Weyla mexicana (Jaworski, 1929) is a junior synonym of W. alata. Having studied Jaworski’s type material of W. mexicana and additional material from the type locality, we follow this view here. Aberhan (1998a) regarded W. mexicana as a synonym of Weyla meeki Damborenea (1987b), a problematic species first described as Pecten acutiplicatus by Meek (1864) from the Lower Jurassic of California. According to present knowledge, we cannot exclude W. meeki as another junior synonym of W. alata. Clearly, more well-preserved, articulated specimens of W. meeki are needed from the type locality to settle this issue. In the studied sections, W. alata is by far the most common species of Weyla. In comparison with Weyla (W.) bodenbenderi (see below), its ribs are always simple and do not exhibit intercalations of secondary ribs or splitting of primary ribs. Weyla (Lywea) unca (see below) has triangular ribs on both valves and is biconvex in all ontogenetic stages.
Figure 9. Scatter diagram showing the length-to-height ratio of Weyla (Weyla) alata (von Buch, 1838).
Weyla (Weyla) bodenbenderi (Behrendsen, 1891) Fig. 7K. 1891 Pecten Bodenbenderi sp. nov.—Behrendsen: 391, pl. 22, fig. 3. 1987b Weyla (Weyla) bodenbenderi (Behrendsen 1891)—Damborenea: 178, pl. 7, fig. 3; pl. 10, fig. 6; pl. 11, fig. 1–2; pl. 12, figs. 1–3 ; text-fig. 22b–c (see for extensive synonymy list). 1994 Weyla (Weyla) bodenbenderi (Behrendsen 1891)—Aberhan: 44, pl. 23, figs. 3–4 (see for synonymy list of South American records). 1998a Weyla (Weyla) bodenbenderi (Behrendsen 1891)—Aberhan: 120, pl. 14, fig. 13, pl. 15, fig. 2, pl. 16, fig. 5, pl. 19, fig. 2 (see for synonymy list of North American records). Material: One fragmented left valve (MB.M.4701) from the lower Sinemurian, unit 18, Sierra del Álamo; one external mold of a fragmented right valve (ERNO-8188) from the upper Sinemurian, lower member, section 3, and one fragmented left valve (MB.M.4700) from the lower Pliensbachian, middle member, section 3, Sierra de Santa Rosa; one fragmented right valve (MB.M.4702) from the lower Pliensbachian, Pozos de Serna. Description: Weyla (W.) bodenbenderi has a mediumsized, pectiniform, inequivalved shell. The convex right valves carry strong radial ribs, which are broad and rounded and have slightly flattened tops with steep flanks. The interspaces between the ribs are smaller than the ribs or of the same size. In some specimens, fine radial striae are present in these interspaces. In one specimen, the ribs split up before they
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reach the ventral margin. The flat to slightly convex left valve carries 12 rounded primary ribs, which increase in number by splitting of secondary ribs. Remarks: These are the first records of the species from Mexico. A much more complete description of W. bodenbenderi was given by Damborenea (1987b) and Aberhan (1994) based on Lower Jurassic specimens from Argentina and Chile, respectively. In Mexico, it occurs together with W. alata at the same localities in the Sierra de Santa Rosa, Sierra del Álamo, and Pozos de Serna, although in much smaller numbers. Weyla (Weyla) titan (Möricke, 1894) Fig. 7L. 1894 Vola alata (v. Buch) var. titan—Möricke: 41. 1899 Pecten titan Mör.—Philippi: 26, pl. 15. 1994 Weyla (Weyla) titan (Möricke 1894)—Aberhan: 45, pl. 24, fig. 3; pl. 25, fig. 1. Material: One fragmented external mold of a left valve (MB.M.4703) from the lower Sinemurian, unit 18, Sierra del Álamo; one fragmented right valve and one fragmentary articulated specimen (MB.M.4705 to 4706) from the lower Pliensbachian, middle member, section 3, Sierra de Santa Rosa; one fragmented external mold of a left valve (ERNO-8189) from the lower Toarcian, Sierra de La Jojoba. Remarks: W. titan was described in some detail by Aberhan (1994), and this is the first record of the species from Mexico. Damborenea and González-León (1997) suggested that W. titan could be just a “gerontic” variant of W. alata and that the two taxa be separated at subspecies level at most. The poorly preserved Mexican specimens at hand do not add to this discussion. However, a comparison of right valves of both taxa, based on Early Jurassic specimens from northern Chile, reveals that, at identical stages of growth, the radial ribs of W. titan are more than twice as broad than those of W. alata. Similarly, the interspaces between ribs are considerably narrower than the ribs in W. titan, whereas the interspaces are of about the same width as or wider than the ribs in W. alata. In our view, this warrants distinction of both taxa at the species level. Subgenus Lywea Damborenea, 1987 Type species. Pecten uncus Philippi, 1899. Weyla (Lywea) unca (Philippi, 1899) Figs. 10A and 10B. 1899 Pecten uncus Ph. sp. nov.—Philippi: 30, pl. 17, fig. 3a–c. 1942 Pecten cf. uncus Phil.—A. Leanza: 170, pl. 9, figs. 1–3, 6. 1987b Weyla (Lywea) unca (R. Philippi 1899)—Damborenea: 187, pl. 12, fig. 4; pl. 13, figs. 1–4, 5a–b, 6a–b, 7a–c, 8, 9a–d, 10a–d, 11a–b; text-fig. 25 (see for extensive synonymy list). 1987b Weyla (Lywea) meeki nov. nom.—Damborenea: 186, 189, text-fig. 26. 1992 Weyla (Weyla) unca (Philippi)—Damborenea, pl. 115, fig. 16a–b. 1994 Weyla (Lywea) unca (Philippi 1899)—Aberhan: 46,
pl. 23, fig. 5; pl. 24, fig. 2; pl. 25, fig. 2. 1994 Weyla (Lywea) aff. unca (Philippi)—Aberhan: 46, pl. 24, fig. 1. 1998a Weyla (Lywea) unca (Philippi 1899)—Aberhan: 128, pl. 16, figs. 4, 6, 7, 9; pl. 18, figs. 7, 10, 13; pl. 19, figs. 4, 6; text-fig. 10. 1998b Weyla (Lywea) unca (Philippi)—Aberhan, fig. 4F. Material: One fragmented internal mold of a right valve (MB.M.4707) from the upper Hettangian–lower Sinemurian, unit 16, four fragmented right valves, two left valves, and two external molds of left valves (ERNO-8193 to 8194, MB.M.4708 to 4713) from the lower Sinemurian, unit 18, and two fragmented right valves (MB.M.4714 to 4715) from the upper Sinemurian, unit 19, Sierra del Álamo; two right valves (MB.M.4716 to 4717) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: In contrast to other species of Weyla described above, Weyla (Lywea) unca is biconvex at all ontogenetic stages. The right valve is more convex than the left valve. The right valve is higher than long, whereas the left one is longer than high. The ribs, 12 to 13 in number, and the interspaces are V-shaped in cross section on both valves. They decrease in strength toward the ventral margin. The interspaces are of the same width as the ribs. On the flanks of the ribs, a fine radial striation and subtle growth lines can be seen.
Figure 10. (A–B) Weyla (Lywea) unca (Philippi, 1899); (A) latex cast of an external mold of a left valve, × 1, ERNO-8193. (B) latex cast of an external mold of a left valve, × 1, ERNO-8194; all lower Sinemurian of Sierra del Álamo. (C–F) Modiolus (Modiolus) giganteus Quenstedt, 1857. (C) internal mold of a right valve, × 1, ERNO-8195; (D–F) steinkern; (D) right valve view, (E) dorsal view, (F) left valve view; × 1, ERNO-8196; all lower Pliensbachian of Sierra de Santa Rosa. (G) Modiolus (Modiolus) cf. baylei (Philippi, 1899); latex cast of an external mold of a right valve, lower Pliensbachian of Sierra de Santa Rosa, × 1, ERNO-8197. (H) Modiolus (Cyranus) hillanus (J. Sowerby, 1818). Plaster cast of an external mold of a left valve, lower Sinemurian of Sierra del Álamo, × 1, ERNO-8198. (I–J) Frenguelliella poultoni H. A. Leanza, 1993. (I) latex cast of an external mold of a left valve, × 2, ERNO-8199; (J) latex cast of an external mold of a right valve, × 2, ERNO-8200; all lower Sinemurian of Sierra del Álamo. (K) Groeberella sp. A; composite mold of a left valve, × 1.5, ERNO-8201, lower Pliensbachian of Pozos de Serna. (L–M) Prosogyrotrigonia sp. A. (L) latex cast of an external mold of a left valve, × 1, ERNO-8202; (M) latex cast of an external mold of a right valve, × 2, ERNO-8203; all lower Sinemurian of Sierra del Álamo. (N) Lucinidae gen. et sp. indet. A; composite mold of a right valve, × 1, ERNO-8204, lower Pliensbachian of Sierra de Santa Rosa. (O–R) Lucinidae gen. et sp. indet. B. (O) internal mold of a left valve, × 2, ERNO-8205, upper Hettangian–lower Sinemurian of Sierra del Álamo; (P) internal mold of a right valve, × 2, ERNO-8206, lower Pliensbachian of Sierra de Santa Rosa; (Q) slab with internal molds of an articulated specimen in butterfly position, × 2, ERNO-8207, upper Hettangian–lower Sinemurian of Sierra del Álamo; (R) latex cast of a slab with several densely packed valves, × 2, ERNO-8208, upper Hettangian–lower Sinemurian of Sierra del Álamo.
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Remarks: W. unca is the only species of Weyla described here that is biconvex. The specimens figured by Damborenea (1987b) as Weyla (Lywea) meeki are considered to belong to W. unca (see Aberhan, 1998a, p. 122). W. meeki (proposed by Damborenea, 1987b, to replace Pecten acutiplicatus Meek) has a plano-convex shell (Aberhan 1998a) and possibly is another junior synonym of W. alata (see above). This is the first record of this species from the Lower Jurassic of Mexico. Subclass Isofilibranchia Iredale, 1939 Order Mytiloida Férussac, 1822 Family Mytilidae Rafinesque, 1815 Genus Modiolus Lamarck, 1799 Type species. Mytilus modiolus Linné, 1758. Subgenus Modiolus Modiolus (Modiolus) giganteus Quenstedt, 1857 Figs. 10C–10F. 1857 Modiola gigantea sp. nov.—Quenstedt: 439, unnumbered text-fig. 1942 Modiolus gigantoides sp. nov.—A. Leanza: 182, pl. 11, fig. 3. 1987a Falcimytilus ? gigantoides (A. Leanza 1942)—Damborenea: 85, pl. 3, figs. 5–7; text-figs. 21, 22. 1992 Falcimytilus (?) gigantoides (Leanza)—Damborenea, pl. 117, fig. 3. 1994 Modiolus (Modiolus) giganteus Quenstedt 1857—Aberhan: 49, pl. 27, figs. 6–9 (see for synonymy list). 1997 Falcimytilus sp.—Damborenea and González-León: 191, fig. 5.11. Material: One steinkern, one external and three internal molds of left valves, and four internal molds of right valves (ERNO-8195 to 8196, MB.M.4738 to 4744) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa; one internal mold of a right valve (MB.M. 4745) from the lower Pliensbachian, Pozos de Serna. Description: Modiolus (M.) giganteus is medium-sized, moderately inflated, and equivalved with a subtriangular to sickle-shaped outline. Its height almost equals its length. The dorsal margin is straight to slightly convex and somewhat longer than half the total shell length. The posterior margin is convex and meets the dorsal margin with an obtuse angle of about 130°. The ventral margin is slightly concave, whereas the anterior margin is very short and convex. The umbo is prosogyrate and terminal. A faint umbonal ridge extends from the umbo to the postero-ventral corner of the shell. The strength of the umbonal ridge varies. A shallow sulcus, which lies anteriorly of the umbonal ridge, meets the ventral margin at its greatest concavity. The outer surface of the shell is covered with regularly spaced growth lines. Remarks: For a detailed description of the species, affinities to other representatives of Modiolus, and comments on differences in size between Lower Jurassic specimens from South America and Middle Jurassic specimens from Europe see Damborenea (1987a) and Aberhan (1994).
Modiolus (Modiolus) cf. baylei (Philippi, 1899) Fig. 10G. cf. Mytilus scalprum Goldf.—Bayle & Coquand: 15, pl. 7, figs. 3–4. cf. 1899 Modiola baylei sp. nov.—R. A. Philippi: 48, pl. 24, fig. 8 (copy from Bayle and Coquand 1851, pl. 7, figs. 3–4), cf. 1987a Modiolus baylei R. Philippi 1899 ?—Damborenea: 91, pl. 4, figs. 2, 3a–b, 4–5 (see for detailed synonymy list). cf. 1994 Modiolus (Modiolus) baylei (Philippi 1899)—Aberhan: 47, pl. 26, figs. 1–4. 1997 Modiolus cf. M. baylei Philippi, 1899—Damborenea and González-León: 191, fig. 5.10. Material: One articulated specimen (MB.M.4707) from the upper Hettangian–lower Sinemurian, unit 16, Sierra del Álamo; one fragmented composite mold of a left valve (MB.M.4746) from the lower Pliensbachian, middle member, section 3, and one external mold of a right valve (ERNO-8197) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: Modiolus (Modiolus) cf. baylei has a very elongated, narrow shell with an umbonal ridge. The umbo is prosogyrate and subterminal. The height is less than half of the length of the shell. The dorsal margin is straight and nearly half as long as the shell. The posterior margin is slightly convex. The ventral margin is straight and becomes slightly concave at the point where it meets the umbonal ridge. The surface of the shell is covered with concentric growth lines. Together with radial striae on the ventral side of the umbonal ridge, these growth lines produce a delicate reticulate pattern on the shell surface. Remarks: The specimens closely match those described and illustrated as M. cf. baylei in Damborenea and GonzálezLeón (1997, fig. 5.10) from the same locality. Compared with M. baylei Philippi (Damborenea, 1987a; Aberhan, 1994), they are more elongated, and therefore are referred to M. baylei with reservation. Subgenus Cyranus Hodges, 2000 Type species. Modiola hillana J. Sowerby, 1818, designated by Hodges, 2000. Modiolus (Cyranus) hillanus (J. Sowerby, 1818) Fig. 10H. 1818 Modiola Hillana sp. nov.—J. Sowerby: 21, pl. 212, fig. 2. 1915 Modiola hillana Sow.—Jaworski: 419. 1925 Modiola hillana Sow.—Jaworski: 66. 1987a Modiolus cf. thiollierei (Dumortier 1869)—Damborenea: 91, pl. 3, fig. 8. 2000 Modiolus (Cyranus) hillanus (J. Sowerby, 1818)—Hodges: 57, pl. 5, figs. 18–24; text-figs. 61, 62 (see for extensive synonymy list). Material: One external mold of a left valve (ERNO-8198) from the lower Sinemurian, section 18, Sierra del Álamo. Description: Modiolus (Cyranus) hillanus is a medium-sized
Early Jurassic bivalves of the Antimonio terrane mytilid with an elongated to trapezoidal outline (L = 35 mm, H = 21 mm). Its dorsal margin is straight and half as long as the shell. The posterior margin is convex and well rounded and meets the dorsal margin at a very obtuse angle. The form of the ventral margin is straight to slightly convex. The umbo lies subterminal. A slightly curved umbonal ridge extends from the umbo to the postero-ventral corner of the shell. The outer surface of the shell is covered with faint, regularly spaced concentric growth lines. Remarks: This is the first record of M. (C.) hillanus from Mexico. A very similar species is Modiolus (Cyranus) ventricosus Roemer 1836, in which Hodges (2000) also included Modiolus thiollierei Dumortier. According to Hodges, M. hillanus can be distinguished from M. ventricosus by a greater oblique height–oblique length ratio and a more strongly curved body. Modiolus (M.) scalprum J. Sowerby, as described by Aberhan (1994) from the Lower Jurassic of Chile, can be distinguished from M. hillanus by the straighter, narrower, and more elongated outline. Damborenea (1987a, p. 91, pl. 3, fig. 8) figured a specimen from the Pliensbachian of Rio Atuel, Argentina, which has been described by Jaworski (1915, 1925) as Modiola hillana Sow., and placed it into M. cf. thiollierei. In our view, this specimen corresponds much better to M. (C.) hillanus described herein and therefore is considered the first record of M. hillanus from South America. In comparison with the other species of Modiolus described herein, M. hillanus is less elongated than M. cf. baylei, and is not as curved and sickle-shaped as M. giganteus. Subclass Heteroconchia Hertwig, 1895 Order Modiomorphoida Newell, 1969 Family Myoconchidae Newell, 1957 Genus Myoconcha J. de C. Sowerby, 1824 Type species. Myoconcha crassa J. de C. Sowerby 1824, by monotypy. Myoconcha neuquena A. Leanza, 1940 Figs. 13A–13C. 1929 Myoconcha cf. Valenciennesi Bayle and Coquand— Jaworski: 8, pl. 1, fig. 8a–b (non Bayle and Coquand). 1940 Myoconcha neuquena sp. nov.—A. Leanza: 126, pls. 1–2. 1997 Myoconcha neuquena Leanza, 1940—Damborenea and González-León: 192, figs. 5.12–5.14 (see for synonymy list). Material: Three specimens (MB.M.4826 to 4828) from the lower Pliensbachian, middle member, section 3, and 16 specimens (ERNO-8209 to 8210, MB.M.4829 to 4842) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Most of the specimens are articulated and preserved with shell. Description: Myoconcha neuquena is medium- to largesized (L = 82–145 mm), thick-shelled and equivalved with a mytiliform outline. The umbo is narrow, acute, and terminal. A rather prominent, acute posterior ridge runs from the umbo
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to the postero-ventral corner. This ridge makes the shell look slightly quadrate in cross section. The greatest inflation of the shell is at about mid-length. The ventral margin is nearly straight and meets the posterior margin with a narrow curve of about 90°. The posterior margin is slightly convex and oblique. The dorsal margin is evenly convex. The surface is covered with numerous regularly spaced, commarginal folds. Remarks: Damborenea and González-León (1997) describe M. neuquena from the same locality. Their specimens are similar in many characters, such as size, ornamentation and state of preservation. Order Trigonioida Dall, 1889 Family Trigoniidae Lamarck, 1819 Genus Frenguelliella A. Leanza, 1942 Type species. Trigonia inexspectata Jaworski, 1915. Frenguelliella poultoni H. A. Leanza, 1993 Figs. 10I, 10J, and 11. 1929 Trigonia cf. inexspectata Jaw.—Jaworski: 7, pl. 1, fig. 4. 1979 Frenguelliella sp. B.—Poulton: 18, pl. 1, fig. 10. 1993 Frenguelliella poultoni sp. nov.– H. A. Leanza: 26, pl. 2, figs. 3–6. Material: Three external molds of right valves and four external molds of left valves (ERNO-8199 to 8200, MB.M.4749 to 4753) from the lower Sinemurian, unit 18, Sierra del Álamo; one composite mold of a right valve (MB.M.4757) from the upper Sinemurian, lower member, section 3, Sierra de Santa Rosa; one internal and two external molds of right valves and one external mold of a left valve (MB.M.4758 to 4761) from the lower Pliensbachian, Pozos de Serna; one articulated specimen, one steinkern, and internal molds of one left valve and one right valve (MB.M.4762 to 4765) from the lower Pliensbachian, middle member, section 3, and one external mold of a left valve and one steinkern (MB.M.4768 to 4769) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: Frenguelliella poultoni is inequilateral, small- to medium-sized and has an oval to subrectangular outline. The umbo is situated mesially. The postero-dorsal margin is slightly concave. The postero-dorsal corner forms an obtuse angle. The posterior margin is straight. The posteroventral corner of the shell forms an angle of about 90°. The ventral margin is well curved and passes gradually into the anterior margin. The flank is covered with 8–15 simple, concentric ribs, which start at the anterior margin and end shortly before the distinct marginal carina. The marginal carina wears faint beads. The area is divided by a subtle median groove and covered with fine commarginal lamellae. Remarks: Our specimens apparently are identical to Trigonia cf. inexspectata described by Jaworski (1929) from the same locality at Sierra de Santa Rosa. This specimen is an incomplete external mold of a left valve, which superficially resembles Frenguelliella inexspectata (Jaworski, 1915,
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Figure 11. Scatter diagram showing the length-to-height ratio of Frenguelliella poultoni (H.A. Leanza, 1993).
p. 377, pl. 5, fig. 2) from the Lower Jurassic of Argentina but also shows some differences, such as a smaller number of concentric ribs and a less distinct marginal carina. Because of the poor preservation of the specimen, Jaworski referred it to T. inexspectata with reservation and hoped for a better determination when better-preserved material would become available. Later, Poulton (1979) described a similar specimen as Frenguelliella sp. B from the lower Sinemurian of the Laberge Group, Yukon Territory, Canada. Finally, Leanza (1993) described this species from the Pliensbachian of the Neuquén Basin, Argentina, under the new name Frenguelliella poultoni. He separated F. poultoni from Frenguelliella tapiai Lambert, 1944, which occurs in the same formation, by the subquadrate shape of F. tapiai and the greater number of costae on its flanks, and by the style of ornamentation of the area. Trigonia chubutensis, first described by Feruglio (1934, pl. 4, figs. 9, 11) from the Lower Jurassic of Rio Genua, Patagonia, is very similar to F. poultoni in the style of ornamentation on the flank and on the area, but the latter species has a more prominent marginal carina and a smaller area. Family Groeberellidae Pérez, Reyes, and Damborenea, 1995 Genus Groeberella H. A. Leanza, 1993 Type species. Myophoria neuquensis Groeber, 1924. Groeberella sp. A Fig. 10K. Material: One composite mold of a left valve (ERNO-8201) from the lower Pliensbachian, Pozos de Serna. Description: Groeberella sp. A is medium-sized, subtrigonal to oval in outline, and has an opisthogyrate, mesial umbo. The posterior margin is short and slightly concave. The ventral margin is divided by the second costa from the anterior
into a posterior part, which is slightly concave, and an anterior part, which is straight. The anterior margin is short and slightly convex. The postero-dorsal margin is not preserved. The flank wears two strong radial costae. The marginal carina and the escutcheon carina are very prominent and stronger than the two radial costae on the flank. Other characters cannot be observed. Remarks: The genus Groeberella is known from the Lower and Middle Jurassic of Argentina and Chile (see Pérez et al., 1995, and H. A. Leanza, 1993, for details) and this is the first record from Mexico. So far, there is only one species that has been assigned to this genus, Groeberella neuquensis (Groeber, 1924). This species has a conspicuously larger shell than the Mexican specimen, and it carries commarginal growth lines, which cover the whole shell and develop into thick transverse rugae on the marginal carina and the second costa. H. A. Leanza (1993) and Pérez et al. (1995) both described some other specimens of the genus Groeberella that cannot be assigned to G. neuquensis. These specimens differ from G. neuquensis by their smaller size, a thinner shell, and a different outline. Family Prosogyrotrigoniidae Kobayashi, 1954 Genus Prosogyrotrigonia Krumbeck, 1924 Type species. Prosogyrotrigonia timorensis Krumbeck, 1924. Prosogyrotrigonia sp. A Figs. 10L and 10M. Material: One external mold of a left valve and one external mold of a right valve (ERNO-8202 to 8203) from the lower Sinemurian, unit 18, Sierra del Álamo. Description: Prosogyrotrigonia sp. A is small, inequilateral, suboval in outline, and slightly longer than high. Its umbo is small and prosogyrate. The dorsal margin is slightly concave. The posterior margin is slightly convex and meets the evenly convex ventral margin at an obtuse angle. The anterior margin is also evenly convex. The flank is covered with regular commarginal riblets (about 11 per cm in a specimen with L = 23 mm and H = 16 mm). The interspaces between these ribs are concave and about twice as wide as the ribs. Their width increases gradually with the growth of the shell. Instead of a marginal carina, the border between shell flank and area is marked by a slight angulation in the surface of the shell. Additionally, flank and area are separated by a change in ornamentation, as the area is covered with weaker and more numerous commarginal riblets. In early growth stages, the ribs on the area are similar in number and strength to the ones on the flank. In one juvenile specimen (L = 9 mm, H = 7.5 mm), only the most ventral rib splits into two weaker riblets at the border between flank and area. On an even smaller specimen, all ribs on flank and area are similar in strength and number. The escutcheon is long and narrow. Remarks: Prosogyrotrigonia sp. A shows all characteristic elements of the genus, such as absence of a marginal carina and change in ornamentation. So far, representatives of this
Early Jurassic bivalves of the Antimonio terrane
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Figure 12. Scatter diagram showing the length-to-height ratio of Lucinidae gen. et sp. indet. B (circles) and other lucinids (squares) from the Jurassic of North and South America, based on typical specimens illustrated in the literature. 1—Lucina goliath Gottsche; 2—Lucina neuquensis Weaver; 3—Lucina bellona d’Orbigny; 4—Lucina liasina Steinmann; 5—Lucina payllalefi A. Leanza; 6—Lucina huayquimili A. Leanza; 7—Lucina lotenoensis Weaver; 8—Lucina chubutensis Wanish; 9—Lucina plana v. Zieten; 10—Lucina atacamensis Möricke; 11—Lucina laevis Gottsche; 12—Lucina intumescens Gottsche; 13—Lucina feruglioi Wanish; 14—Lucina argentina Behrendsen; 15—Lucina dosiniaeformis Gottsche; 16—Lucina problematica Terquem.
genus are known from Timor and Japan. The type species, Prosogyrotrigonia timorica Krumbeck (1924, pl. 19, figs. 19–22), has been described from the Rhaetian of Timor. This species differs from the Mexican specimens in a more rectangular outline and finer ribs on the flank. Prosogyrotrigonia inouyei (Yehara) has been described, among others, by Kobayashi and Mori (1954, p. 157, pl. 15, figs. 3–5) and Hayami (1959, p. 70, pl. 7, figs. 12–14) from the Lower Jurassic of Japan. In comparison with the Mexican specimens, P. inouyei is nearly as long as high and has fewer and much coarser ribs. Very similar to the Japanese species is a specimen from the Hettangian of the Yukon Territory, Alaska, described as Prosogyrotrigonia (?) sp. cf. P. inouyei (Yehara) by Frebold and Poulton (1977, p. 96, pl. 2, figs. 5–9). The here described Prosogyrotrigonia sp. A seems to be conspecific to as yet undescribed material from the Hettangian and Sinemurian of northern Chile (pers. observation). Order Veneroida H. Adams and A. Adams, 1856 Family Lucinidae Fleming, 1828 Lucinidae gen. et sp. indet. A Fig. 10N. 1997 Lucinidae gen. et sp. indet.—Damborenea and GonzálezLeón: 192, fig. 9.2. Material: Composite molds of one right and two single valves, one internal mold of a left valve, and one steinkern (ERNO-8204, MB.M.4771 to 4774) from the lower Pliensbachian, middle member, section 3, Sierra de Santa Rosa; one external mold of a left valve (Mb.M.4780) from the lower Pliensbachian, Pozos de Serna.
Description: This lucinid bivalve is large, equivalved and sub-equilateral, only slightly inflated, and subquadrate in outline. The posterior margin is truncated. The anterior margin is slightly concave and meets the ventral margin at an obtuse angle. The ventral margin is slightly convex. The shell is covered with numerous, regularly spaced commarginal ribs (~10 ribs per cm). The inner margin seems to be smooth. Remarks: Damborenea and González-León (1997, p. 194) described a similar specimen from the same locality in Mexico and, in their remarks, mentioned affinities to the genus Mesomiltha Chavan. These authors also discussed affinities to related lucinid taxa. In addition, Lucina atacamensis Möricke, 1894 (53, pl. 4, fig. 10a–b) from the middle and upper Lower Jurassic of Chile can be mentioned, which differs in a more strongly inflated shell and an ornamentation that consists of weak commarginal growth lines. Lucinidae gen. et sp. indet. B Figs. 10O–10R and 12. Material: Twelve internal and 13 external molds of left valves, nine internal and four external molds of right valves (ERNO-8205, ERNO-8207 to 8208, MB.M.4795 to 4825), with only rare shell preservation from the upper Hettangian–lower Sinemurian, unit 16, Sierra del Álamo; one internal mold of a left valve (ERNO-8206) from the lower Pliensbachian, middle member, section 3, Sierra de Santa Rosa. Description: Lucinidae gen. et sp. indet. B is small (see Fig. 12), inequilateral, equivalved, and oval to subquadrate in
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outline. It is slightly inflated, and longer than high (Fig. 12). The anterior margin is slightly convex, and the ventral margin is well rounded and passes gradually into the well-rounded posterior margin. The postero-dorsal margin is slightly concave; the antero-dorsal margin is straight. The umbo is orthogyrate and is situated at about one-third of the shell length from the anterior. The escutcheon is lanceolate, the lunule small and narrow. The external molds show fine, regular, commarginal growth lines, which cover the whole surface. Except for these growth lines, the shell surface is smooth. The specimens have an integripalliate pallial line. Other internal characters cannot be observed. Remarks: Lucinidae gen. et sp. indet. B is a new representative of the family Lucinidae in South America. As hinge characters are not known, an exact assignment at the genus level is not possible. Many lucinids have been described from the Jurassic of America (see also Fig. 12), mostly from the Middle and Upper Jurassic of South America, but only a few are similar to Lucinidae gen. et sp. indet. B in outline and ornamentation. Strongest affinities exist to Lucina feruglioi Wanish de Carral Tolosa (1942, p. 53, pl. 5, figs. 2a–b) from the Lower Jurassic of Argentina, which agrees well in size and in the length-to-height ratio but is covered with numerous strong commarginal ribs. From the same locality, Wanish de Carral Tolosa (1942, p. 54, pl. 5, fig. 3) described Lucina chubutensis, which has an almost circular outline and also strong commarginal ribs. Lucina huayquimili A. Leanza (1942, p. 193, pl. 18, figs. 3, 5) from the Lower Jurassic of Argentina, which probably belongs to the genus Mesomiltha, is medium-sized, sub-orbicular, slightly longer than high, only slightly inflated, and wears concentric riblets. It differs mainly by its larger size and its ornamentation from the Mexican specimens. L. payllalefi A. Leanza (1942, p. 192, pl. 18, fig. 2) from the Lower Jurassic of Argentina carries numerous commarginal ribs and is higher than long. Lucina problematica Terquem (1855, p. 119, pl. 20, fig. 7) from the Hettangian of France is similar in size, outline, and length-to-height ratio, but has numerous prominent, regular, and narrow costae. Gottsche (1878) described several lucinids from the Middle Jurassic of Argentina. Lucina laevis Gottsche (1878, p. 27, pl. 5, fig. 9) also has a smooth shell surface but is nearly circular, and thus has a lower length-to-height ratio than the Mexican specimens. Similarly, Lucina intumescens Gottsche (1878, p. 27, pl. 5, fig. 10) is almost circular in outline and strongly inflated, and is covered with faint concentric riblets. Lucina dosiniaeformis Gottsche (1878, p. 28, pl. 6, fig. 13) is small, circular, has pointed umbones, and is ornamented with numerous commarginal riblets. A species that also has been described by Steinmann (1929, p. 79, fig. 92) and Cox (1956, p. 1184, pl. 128, fig. 4) from the Middle Jurassic of Peru is Lucina goliath Gottsche (1878, p. 28, pl. 5, fig. 11). With its height and length exceeding 50 mm, this species is conspicuously larger than the Mexican material. It is ornamented by numerous commarginal riblets.
Lucina plana von Zieten (1833, p. 96, pl. 72, fig. 4a–b) is a common species in the Lower and Middle Jurassic of Europe (e.g., Quenstedt, 1856: 319, pl. 44, fig. 4; Gottsche, 1878, p. 27, pl. 5, figs. 7, 8; Palmer, 1966, pl. 1, figs. 3–5). It is medium-sized, has strong commarginal ribs and fine radial riblets, and its length-to-height ratio is slightly lower than that of the material described herein. Lucina argentina Behrendsen (1891, p. 415, pl. 25, fig. 3) has first been described from the Upper Jurassic of Argentina. It is small, suboval, equilateral, and has relatively strong growth lines and a lower length-to-height ratio than the Mexican species. Haupt (1907, p. 218) and Weaver (1931, p. 350) also describe L. argentina from the Upper Jurassic and the Lower Cretaceous of Argentina. Lucina atacamensis Möricke (1894, p. 53, pl. 4, figs. 10a–b) from the Lower Jurassic of Chile is suboval in outline, medium-sized, strongly inflated, and carries concentric growth lines, which increase in prominence toward the posterior part of the shell. Its length-to-height ratio is only slightly lower than that of Lucinidae gen. et sp. indet. B. Lucina liasina Steinmann (1929, p. 67, fig. 71) from the Hettangian of Peru is medium-sized, suboval, and has a relatively small and acute umbo. It is ornamented with numerous regular commarginal growth rugae, and its length-toheight ratio is larger than that of the specimens described herein. L. neuquensis Haupt (1907, p. 217, pl. 10, figs. 3 a–b) from the Upper Jurassic of Argentina is relatively large and has a circular outline and prominent concentric ribs. Lucina bellona d’Orbigny is known from the Middle Jurassic of Europe (e.g., Morris and Lycett, 1853, pl. 67, pl. 6, figs. 18–18a) and the Lower and Middle Jurassic of Argentina, Mexico, and Peru (e.g., Cox, 1956, p. 1184, pl. 128, fig. 2). It is relatively large and has a circular outline and well-defined commarginal lamellae. L. corbisoides d’Orbigny var. loteonensis Weaver (1931, p. 351, pl. 39, fig. 241) from the Lower Cretaceous of Argentina has a circular outline and is ornamented with fine commarginal growth lines. A specimen figured by Damborenea (1992, pl. 119, fig. 8) from the Tithonian of Argentina has a slightly lower length-to-height ratio than the Mexican specimens. Family Astartidae d’Orbigny, 1844 Genus Neocrassina Fischer, 1886 Type species. Astarte obliqua Deshayes 1830 (= Cypricardia obliqua Lamarck 1819). Neocrassina gueuxi (d’Orbigny, 1850) Figs. 13D–13F and 14. 1850 Astarte gueuxi sp. nov.—d’Orbigny, p. 216, n° 80. 1853 Astarte consobrina sp. nov.—Chapuis and Dewalque, p. 149, pl. 22, fig. 3. 1856 Astarte psilonoti sp. nov.—Quenstedt, p. 45, pl. 3, fig. 14. 1864 Astarte dentilatum sp. nov.—Etheridge, p. 112, text-figs. 5–7.
Early Jurassic bivalves of the Antimonio terrane 1867 Astarte cammertonensis sp. nov.—Moore, p. pl. 17, fig. 3. 1869 Astarte fontis sp. nov.—Dumortier, p. 268, pl. 30, figs. 12–14. 1907 Astarte gueuxi—Thévenin, p. 25, pl. 8, figs. 8–10. 1966 Astarte gueuxi d’Orbigny—Palmer, p. 75, pl. 2, figs. 1–4 (see for comments on synonymy). 1997 Neocrassina? sp.—Damborenea and González-León, p. 194, fig. 9.1. 1998 Astarte gueuxii—Harper et al., p. 356, fig. 1B–C. Material: Eleven right valves and 10 left valves (ERNO-8212 to 8213, MB.M.4843 to 4861) from the lower Sinemurian–upper Hettangian, unit 16, two right valves and one left valve (MB.M.4866 to 4868) from the lower Sinemurian, unit 18, and three articulated specimens and three left valves (MB.M.4869 to 4874) from the upper Sinemurian, unit 19, Sierra del Álamo; one right valve and two left valves (MB.M.4875 to 4877) from the lower Pliensbachian, middle member, section 3, Sierra de Santa Rosa; one articulated specimen, two left valves, and two right valves (ERNO-8211, MB.M.4881 to 4884) from the lower Pliensbachian, Pozos de Serna. Specimens are preserved as molds or with shell. Description: Neocrassina gueuxi is inequilateral and of medium size. Its outline varies from oval and suboval to subquadrate. The umbo is placed at about one-third of the shell length from the anterior end and is prosogyrate. The anterior margin is short and straight to slightly concave. The ventral margin is slightly convex. The posterior margin is slightly convex. The dorsal margin is straight to slightly convex. The shell of adult specimens is covered with at least 25 regularly spaced, acute, commarginal ribs, which are separated by more or less equally wide interspaces. The intervals between the ribs are narrower on the anterior than on the postero-ventral part of the shell. A lunule and a small escutcheon are present. The details of the hinge are not completely preserved in both valves, but the most significant features of the left valve can be recognized. It has two well-developed, oblique, trigonal cardinal teeth (2 and 4b) and a posterior lateral (PII). The nymph is narrow. The inner part of the ventral margin is slightly crenulated. Remarks: This species has been described as Neocrassina? sp. by Damborenea and González-León (1997) from the same locality in the Sierra del Álamo. According to Damborenea and González-León (1997, p. 194), an exact classification was not possible because significant internal features were missing. On the basis of hinge characters observed in our specimens this identification can be confirmed. These authors mentioned a high similarity in shape and size to Neocrassina aureliae (Feruglio), a very common species in the Lower Jurassic of the Andes of South America (Feruglio, 1934, pl. 4, figs. 14–16; Wanish de Carral Tolosa, 1942, p. 47, pl. 4, figs. 1–2; A. Leanza, 1942, p. 190, pl. 17, figs. 4–7). However, the shape of N. aureliae varies from oval-triangulate to oval and slightly elongated, but never reaches such an
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elongated and subquadrate shape as in N. gueuxi. Furthermore, the umbo of N. aureliae is situated in a more anterior position, and the posterior margin is descending more strongly. Astarte gueuxi, described by Palmer (1966) from the Middle Lias Day’s Shell Bed in Dorset, Great Britain, exhibits a variability very similar to the material described herein. He plotted length against height as a percentage of length into a scatter diagram and concluded that A. gueuxi is a highly variable species, and many similar species are younger synonyms of this species (see also Fig. 14). These include Astarte fontis Dumortier (1869, p. 30, figs. 12–14) from the Middle Lias of Saint-Julien, France, and Astarte psilonoti Quenstedt (1856) from the Lias of Germany. Another similar and widespread species is Astarte andium Gottsche from the Lower and Middle Jurassic of Argentina (Gottsche, 1878, p. 29, pl. 7, fig. 8; Tornquist, 1898, p. 170, pl. 20, fig. 11; Wanish de Carral Tolosa, 1942, p. 49, pl. 4, figs. 4–5; Rangel, 1978, p. 29, pl. 4, fig. 1). As for N. aureliae, the outline of the shell does not agree with that of N. gueuxi. Furthermore, Riccardi et al. (1990, p. 88) mentioned that Lower Jurassic records of A. andium belong to N. aureliae. Astarte clandestina Gottsche (1878, p. 30, pl. 7, fig. 11) from the Middle Jurassic of Paso del Espinacito, Argentina, most closely resembles the material described herein. A difference to N. gueuxi is that the inner part of the ventral margin is not crenulated. Astarte keideli Wanish de Carral Tolosa (1942, p. 51, pl. 4, fig. 6) from the Lower Jurassic of Argentina is very similar in shape, but has fewer and stronger commarginal ribs than N. gueuxi. In Astarte cf. antipodum Giebel, Jaworski (1929, p. 7, pl. 1, fig. 7) described another astartid from the Sierra de Santa Rosa area. It has a sub-orbicular outline and 11 strong commarginal ribs, which are separated by much broader intervals, and therefore has little similarity with N. gueuxi. Jaworski (1915) described A. fontis Dumortier from the middle Lower Jurassic of South America. He mentioned a high similarity to the original material of Dumortier (1869) from the Middle Liassic of Saint-Julien, France. However, he did not figure his material. A. Leanza (1942) recognized a similarity of Jaworski’s material to his specimens of N. aureliae from the Lower Jurassic of Piedra Pintada, and included it into the synonymy list. A thorough revision of the concerned species is necessary to resolve this issue. This is the first report of N. gueuxi from America. Family Cardiniidae Zittel, 1881 Genus Cardinia Agassiz, 1841 Type species. Unio listeri J. Sowerby, 1817, subsequent designation by Opinion 292 of the International Commission on Zoological Nomenclature. Cardinia concinna (J. Sowerby, 1819) Figs. 13G and 13H. 1819 Unio concinnus sp. nov.—J. Sowerby: 43, pl. 223, figs. 1–2.
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1833 Unio concinnus Sowerby—v. Ziethen: 80, pl. 60, figs. 2–5. 1862 Cardinia concinna—Chapuis: 79, pl. 16, fig. 3. 1867 Cardinia concinna (Sowerby, spec.)—Dumortier: 207, pl. 47, figs. 2–3. 1885 Cardinia concinna (Sowerby)—Choffat: 2, pl. 3, fig. 4. 1934 Cardinia concinna (Sowerby)—Rosenkrantz: 94, pl. 4, fig. 1. 1975 Cardinia concinna (J. Sowerby 1819)—C. P. Palmer: 26, pl. 3, figs. 1, 2. 1977 Cardinia sp. aff. C. concinna (J. Sowerby)—Frebold and Poulton: 96, pl. 2, fig. 14. 2002 Cardinia concinna (J. Sowerby 1819)—Gahr, pl. 4, fig. 5. Material: 2 right valves and one left valve (ERNO-8214 to 8215, MB.M.4885) from the upper Hettangian–lower Sinemurian, unit 16, Sierra del Álamo. Description: Cardinia concinna is medium-sized, elongated-ovate, and slightly inflated. Its thick shell is twice as long as high. The umbo is slightly prosogyrate and situated at one-fifth of the total shell length from the anterior end. The dorsal margin is slightly convex and meets the posterior margin with a well-rounded, obtuse angle. The posterior margin is slightly convex and passes gradually into the evenly rounded ventral margin. The anterior margin is straight to slightly convex. The surface of the shell is ornamented with regular and very distinct growth lines. Although the shells are slightly eroded, the usual cardiniid dentition with a single, elongated cardinal tooth below the umbo, and the anterior and posterior lateral tooth are identifiable in one right valve from the Sierra del Álamo. Other internal features are not preserved.
Figure 13. (A–C) Myoconcha neuquena A. Leanza, 1940. (A) articulated specimen, right valve view, × 1, ERNO-8209; (B–C) articulated specimen; (B) right valve view, (C) left valve view; × 1, ERNO-8210; all lower Pliensbachian of Sierra de Santa Rosa. (D–F) Neocrassina gueuxi d’Orbigny, 1850. (D) latex cast of an external mold of a right valve, × 2, ERNO-8211, lower Pliensbachian of Pozos de Serna; (E) left valve, × 1, ERNO-8212; (F) latex cast of an internal mold of a left valve, × 1, ERNO-8213; all upper Hettangian–lower Sinemurian of Sierra del Álamo. (G–H) Cardinia concinna (J. Sowerby, 1819). (G) right valve, × 1, ERNO-8214; (H) left valve, × 1, ERNO-8215; all upper Hettangian–lower Sinemurian of Sierra del Álamo. (I) Cardinia? sp. A; latex cast of an external mold of a left valve, × 1, ERNO-8216, upper Hettangian–lower Sinemurian of Sierra del Álamo. (J) Protocardia (Protocardia) truncata (J. de C. Sowerby, 1827); composite mold of a right valve, × 1.5, ERNO-8217, lower Sinemurian of Sierra del Álamo. (K) Protocardia (Protocardia) luggudensis (Troedsson, 1951); right valve view of a composite mold of an articulated specimen, × 1.5, ERNO-8218, lower Pliensbachian of Sierra de Santa Rosa. (L) Protocardia striatula (J. de C. Sowerby, 1829); composite mold of a left valve, × 1.5, ERNO-8219, lower Pliensbachian of Sierra de Santa Rosa. (M) Protocardia sp. A; composite mold of left valve, × 1.5, ERNO-8220, lower Pliensbachian of Pozos de Serna. (N–O) Isocyprina ancatruzi (A. Leanza, 1942); steinkern. (N) dorsal view, (O) right valve view; × 1, ERNO-8221, lower Pliensbachian of Sierra de Santa Rosa.
Figure 14. Scatter diagram showing the ratio of height-length to length of Neocrassina gueuxi d’Orbigny, 1850.
Remarks: C. concinna is primarily known from the Lower Jurassic of Europe, especially from Great Britain (see Palmer, 1975, for details), and this is the first report of the genus from Mexico. This species seems to be somewhat variable, as the material described by J. Sowerby (1819, p. 43, pl. 223, figs. 1–2) is characterized by a well-rounded posterior margin, whereas, for example, Goldfuss (1837, p. 181, pl. 132, fig. 2a–b) and Chapuis (1862, p. 79, pl. 16, fig. 3b) illustrated forms with a somewhat truncated posterior margin. The material from Mexico corresponds well to Sowerby’s material. From North America, only one poorly preserved specimen has been described so far by Frebold and Poulton (1977) from the Hettangian of Yukon Territory. Cardinia andium (Giebel) (e.g., Damborenea, 1992, pl. 117, fig. 1), which is a common and typical species of the Lower Jurassic of Argentina and Chile, is less elongated and therefore has a much more rounded outline. Cardinia chubutensis Wanish de Carral Tolosa (1942, p. 44, pl. 3, fig. 3) has a similar ornamentation but is more elongated, and the posterior margin is truncated. In Cardinia densestriata Jaworski (1915, p. 375, pl. 5, fig. 1), the umbo is situated at about one-third of the total length from the anterior end, the intervals between the growth lines are narrower than in C. concinna, and the outline is less oval. Cardinia? sp. A Fig. 13I. Material: One external mold of a left valve (ERNO-8216) from the upper Hettangian–lower Sinemurian, unit 16, Sierra del Álamo; one external mold of a right valve (MB.M.4886) from the lower Pliensbachian, Pozos de Serna. Description: Cardinia? sp. A is small-sized, equivalved and inequilateral, subquadrate in outline, and moderately inflated. The umbones are small and are situated at onethird of the total shell length from the anterior end. The dorsal margin is short and straight to slightly convex. The posterior margin is straight to slightly convex and obliquely
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slopes to the postero-ventral corner. The ventral margin is straight, the anterior margin slightly convex. The shell is covered with seven to eight thick, commarginal, imbricate folds with upturned, rounded edges. Internal characters are not preserved. Remarks: Because of the lack of internal features, the generic assignment to Cardinia is questionable and based on the great resemblance of our specimens to Cardinia rugulosa Tate (1875, p. 508, text-fig. 3) from the lower Pliensbachian of Great Britain. Following the first description by Tate, Palmer (1975, p. 32, pl. 5, fig. 5) described a new specimen that corresponds very closely to our material. The specimens of Tate (1875), although similar in ornamentation and size, differ somewhat in being more elongated and having a convex posterior margin. C.? sp. A differs from the here described C. concinna by its smaller size and its different ornamentation and outline. Family Cardiidae Lamarck, 1809 Genus Protocardia von Beyrich, 1845 Type species. Cardium hillanum J. Sowerby, 1813, subsequent designation by Herrmannsen, 1847. Subgenus Protocardia Protocardia (Protocardia) truncata (J. de C. Sowerby, 1827) Fig. 13J. 1827 Cardium truncatum sp. nov.—J. de C. Sowerby: 102, pl. 553, fig. 3. 1852 Cardium truncatum Sowerby—Quenstedt: 540, pl. 45, fig. 19. 1871 Cardium truncatum Sow.—Phillips: 136, pl. 8, fig. 28. 1936 Protocardia truncata Phil. (Sow.)—Kuhn: 267, pl. 12, fig. 22. 1948 Protocardia truncata (J. de C. Sowerby)—Wilson: 18, fig. 5D. 1951 Protocardia truncata (Sow.)—Troedsson: 180, pl. 14, figs. 3–5. Material: One external mold of a right valve (MB.M.4887) from the upper Hettangian–lower Sinemurian, unit 16, and two composite molds of right valves and two external molds of left valves (ERNO-8217, MB.M.4888 to 4890) from the lower Sinemurian, unit 18, Sierra del Álamo. Description: Protocardia (P.) truncata is small, subquadrate, equivalved, and sub-equilateral. The height and length are almost equal. The valves are moderately inflated. The umbones are well rounded, slightly prosogyrate, rising above the hinge-line. The anterior and posterior margins are straight to slightly convex. The ventral margin is well curved. An umbonal ridge extends from the umbo to the posteroventral corner. The surface of the shell is covered with fine, commarginal growth lines. The posterior part of the shell is ornamented with radial ribs. Four to six of these ribs are situated anterior of the umbonal ridge, and about the same number of ribs occur posterior to the umbonal ridge. The first one or two are
very faint. The ribs become gradually fainter posteriorly and disappear before reaching the posterior margin, where the surface is smooth. The interspaces between the ribs are very narrow. The inner part of the postero-ventral margin is slightly crenulated where the ribs meet the margin; otherwise it is smooth. Other internal features are unknown. Remarks: So far, P. truncata has not yet been described from North or South America. It is known from the Lower and Middle Jurassic of Europe (Sweden, Great Britain, Germany). Among other species of Protocardia described from the Lower Jurassic of America, Cardium (Protocardium) appressum Gabb, 1877 (286, pl. 40, fig. 17) from Peru differs from P. truncata by a larger size and a more prominent umbo. Affinities to Protocardia striatula (J. de C. Sowerby), which also is known from Canada and Argentina, are discussed below. Related species from the European Jurassic include Protocardia (P.) oxynoti Quenstedt, 1856 (110, pl. 13, fig. 46), which is similar in outline but has a much fainter umbonal ridge, and only two or three ribs are situated anterior to the ridge. Protocardia (P.) philippiana Dunker (1847, p. 116, pl. 17, figs. 6a–c) has a more trigonal outline, a weak umbonal ridge, and no radial ribs anterior to the ridge. Protocardia (Protocardia) striatula (J. de C. Sowerby, 1829) Fig. 13L. 1829 Cardium striatulum sp. nov.—J. de C. Sowerby: 576, pl. 553, figs. 1–2. 1931 Protocardia striatula (Sowerby?) Phillips—Weaver: 353, pl. 38, figs. 236–237. 1957 Protocardia striatula (Phillips)—Frebold: 13, pl. 3, figs. 11–12. 1964 Protocardia striatula (Phillips)—Frebold, pl. 7, fig. 15. Material: Composite molds of one left valve and one right valve (ERNO-8219, MB.M.4897) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: Protocardia (P.) striatula is medium-sized, orbicular, and equivalved. The shell is very globose and nearly as high as long. Its umbones are well rounded, orthogyrate, and protrude considerably beyond the hinge line. The anterior margin and the ventral margin are evenly rounded. The posterior margin is only slightly convex. A faint post-umbonal ridge leads from behind the umbo to the postero-ventral corner. The shell is ornamented with regular, fine, commarginal ribs. Additionally, the posterior part of the shell is covered with 11 to 15 radial ribs. About eight of these ribs are situated anterior of the post-umbonal ridge, the first one or two being more subtle than the others. The radial ribs do not reach the postero-dorsal margin. Internal characters cannot be observed. Remarks: The Mexican specimens correspond well to the material described from the Toarcian of western Canada (Frebold, 1957, 1964) and from the Bajocian of Argentina (Weaver, 1931). They are somewhat larger than European specimens described by Benecke (1905, p. 228, pl. 17, figs. 1–6) from the Liassic of Alsace-Lorraine and by Gahr (2002,
Early Jurassic bivalves of the Antimonio terrane p. 126, pl. 4, figs. 10–11) from the Toarcian of Spain. Protocardia (P.) substricklandi Tornquist (1898, p. 38, pl. 10, fig. 2) from the Middle Jurassic of Argentina is very similar to P. striatula, but its ornamentation is weaker. In comparison with P. truncata (see above), P. striatula attains a larger size, is more inflated, and exhibits a conspicuous commarginal ornamentation. Protocardia (Protocardia) luggudensis (Troedsson, 1951) Fig. 13K. 1951 Anisocardia luggudensis sp. nov.—Troedsson: 185, pl. 14, fig. 1a–c. Material: One internal mold of a right valve (MB.M.4893) from the upper Sinemurian, unit 19, Sierra del Álamo; one composite mold of an articulated specimen and one internal mold of a right valve (ERNO-8218, MB.M.4896) from the lower Pliensbachian, middle member, section 3, and two steinkerns (MB.M.4894 to 4895) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: Protocardia (P.) luggudensis is mediumsized, trigonal in shape, equivalved, and sub-equilateral. The shell is slightly higher than long. The valves are highly inflated. The umbones are well rounded, very prominent, and slightly prosogyrate. The anterior margin is straight to slightly convex. The ventral margin is evenly convex. The posterior margin is convex. A weak post-umbonal ridge runs from behind the umbo to the postero-ventral corner. The surface of the shell is covered with fine, commarginal striae. The posterior part of the shell, behind the post-umbonal ridge, wears about five faint radial ribs. Internal characters are not preserved. Remarks: P. luggudensis is a poorly known species that, until now, was only known from the lower Pliensbachian of Sweden (Troedsson, 1951). Troedsson (1951) incorrectly attributed it to Anisocardia because of its hinge characters and its great similarity to Anisocardia globosa (Roemer, 1839) from the Upper Jurassic of Europe (e.g., Arkell, 1934, p. 272, pl. 36, figs. 3–7). In comparison with P. luggudensis, A. globosa is more strongly inflated and has a more asymmetric umbo. Furthermore, as P. luggudensis exhibits posterior radial ribs, a typical feature of the genus Protocardia, it belongs to Protocardia, and not to Anisocardia. P. luggudensis differs from the here described P. striatula by its highly inflated shell and posterior radial ribs, which are restricted to the part of the shell posterior to the umbonal ridge. P. truncata has a less inflated shell, and the posterior ribs nearly cover the whole posterior half of the shell. P. sp. A (see below) differs by its elongated outline. Protocardia (Protocardia) sp. A Fig. 13M. Material: One composite and one internal mold of left valves (ERNO-8220, MB.M.4898) from the lower Pliensbachian, Pozos de Serna.
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Description: Protocardia (P.) sp. A is oval-elongated in outline, inequilateral, and clearly longer than high. The umbo is small and orthogyrate. The posterior margin is slightly convex. The ventral margin is well rounded and passes gradually into the evenly convex anterior margin. The antero-dorsal margin and the postero-dorsal margin are slightly convex. The surface of the shell is covered with fine, regular, commarginal growth lines. The posterior part of the shell, behind a smooth umbonal ridge, is ornamented with numerous fine, radial striae that extend from the umbo to the posteroventral part of the shell. Remarks: Because of the presence of posterior radial ribs, the material can confidently be referred to Protocardia, even if its elongated form is quite unusual for the genus. Most species of Protocardia are higher than long or have a lower length-toheight ratio than P. (P.) sp. A. A similar specimen has been described as Protocardia (P.) sp. A by Fürsich et al. (2000, p. 114, pl. 12, fig. 13) from the Callovian and Oxfordian of Kachchh, India. It resembles the Mexican material but is larger and has a more distinct commarginal ornamentation. Family Arcticidae Newton, 1891 Genus Isocyprina Röder, 1882 Type species. Cardium cyreniforme Buvignier, 1852; subsequent designation by Cossmann, 1921. Isocyprina ancatruzi (A. Leanza, 1942) Figs. 13N and 13O. 1915 Venilicardia (Cyprina) cornuta d’Orb.—Jaworski: 387, pl. 6, figs. 1–4. 1942 Cypricardia ancatruzi sp. nov.—A. Leanza: 188, pl. 15, fig. 3; pl. 17, figs. 1–3. Material: Ten steinkerns (ERNO-8221, MB.M.4899 to 4907) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: Isocyprina ancatruzi is trigonal to trigonally suboval, medium-sized, inequilateral, equivalved, and moderately inflated. The umbo is broad, slightly enrolled, prosogyrate, and is situated at one-fifth of the shell length from the anterior end. A weak posterior carina runs from behind the umbo to the postero-ventral corner of the shell. Posterior to this carina, the shell is less inflated, and is slightly concave in cross section. The dorsal margin is convex and oblique, and meets the posterior margin at a very obtuse angle. The posterior margin is short, straight to slightly convex, and oblique. It meets the ventral margin at nearly a right angle. The ventral margin is convex and meets the straight anterior margin at a right to slightly obtuse angle. The surface of the shell is smooth except for regular growth lines. No internal features are observable because of the moderate preservation. Remarks: I. ancatruzi was reported by Jaworski (1915) and Wanish de Carral Tolosa (1942) as Venilicardia cornuta
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d’Orbigny, a species that is known from the Upper Jurassic of northern France, northern Germany, and the Swiss Jura Mountains. Jaworski (1915, p. 391) provided good illustrations and a very detailed description of the hinge features. Subsequently, Damborenea (in Riccardi et al., 1990, p. 88) referred to it as Isocyprina ancatruzi, a view that is followed here. Order Pholadomyoida Newell, 1965 Family Pholadomyidae Gray, 1847 Genus Pholadomya G. B. Sowerby, 1823 Type species. Pholadomya candida G. B. Sowerby, 1823, subsequent designation by Gray, 1847. Subgenus Pholadomya Pholadomya (Pholadomya) fidicula J. de C. Sowerby, 1826 Figs. 15A–15C. 1819 Lutraria lirata sp. nov.—J. Sowerby: 47, pl. 225, figs. 1–2. 1826 Pholadomya fidicula nom. nov.—J. de C. Sowerby: 86. 1874 Pholadomya fidicula, Sow.—Moesch: 25, pl. 8, figs. 4–7; pl. 9, figs. 6–8 (see for synonymy list). 1965 Pholadomya fidicula Sowerby—Alencáster and Buitrón: 36, pl. 9, figs. 1–4. 1997 Pholadomya cf. P. fidicula (J. Sowerby, 1819)—Damborenea and González-León: 195, fig. 9.3. 2004 Pholadomya (Pholadomya) fidicula J. de C. Sowerby 1826—Aberhan: 121, pl. 2, figs. 1–6 (see for extensive synonymy list). Material: Seventeen composite molds of articulated specimens (ERNO-8222 to 8223, MB.M.4908 to 4922) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Remarks: Ph. fidicula was recently described by Aberhan (2004) from the Lower Jurassic of Chile and no further description is necessary here. This is the most abundant species of Pholadomya in the study area. Compared with Pholadomya (Ph.) idea (see below), it has a more elongated outline, more numerous radial ribs, and is less inflated. Pholadomya (Ph.) cf. voltzi (see below) has a more rounded and less elongated outline and wears only eight very faint ribs. For further discussion see Aberhan (2004). Pholadomya (Pholadomya) idea d’Orbigny, 1850 Figs. 15D–15F. 1833 Pholadomya ambigua Sowerby—von Zieten: 86, pl. 65, fig. 1. 1850 Pholadomya idea, d’Orb., 1847—d’Orbigny: 7e étage, p. 216. 1915 Pholadomya ambigua Sow.—Jaworski: 423. 1991 Pholadomya idea d’Orbigny—Poulton: 32, pl. 4, figs. 4–12, 14–15. 1997 Pholadomya cf. P. ambigua (J. Sowerby, 1819)—Damborenea and González-León: 195, fig. 9.4. 2004 Pholadomya (Pholadomya) idea d’Orbigny 1850—Aberhan: 124, pl. 3, figs. 4, 6; pl. 4, fig. 1; pl. 10, fig. 6.
Material: Seven composite molds of articulated specimens (ERNO-8224, MB.M.4923 to 4928) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Remarks: The recent description of Ph. idea by Aberhan (2004) perfectly fits the Mexican specimens. Ph. (Pholadomya) cf. ambigua described by Damborenea and González-León (1997) from the same locality in the Sierra de Santa Rosa is included in Ph. idea herein. Although similar in style of ornamentation, the length-to-height ratio of Ph. ambigua J. Sowerby is close to 1, and thus the shell is noticeably less elongated than that of Ph. idea. Pholadomya (Pholadomya) cf. voltzi Agassiz, 1842 Fig. 15G. cf. 1842 Pholadomya Voltzii sp. nov.—Agassiz: 122, pl. 3c, figs. 1–6 (non-figs. 8–9). cf. 1874 Pholadomya Voltzi, Ag.—Moesch: 20, pl. 6, figs. 2–3; pl. 9, figs. 1, 3 (see for short synonymy list). cf. 2004 Pholadomya (Pholadomya) voltzi Agassiz 1842—Aberhan: 128, pl. 1, figs. 1–3; text-fig. 3A–D. Material: One composite mold of a left valve (ERNO-8225) from the lower Toarcian, Sierra de La Jojoba. Description: Pholadomya (Ph.) cf. voltzi is medium-sized, oval to slightly elongated in outline, and moderately inflated. The umbones are well rounded and narrow. It has a slightly concave postero-dorsal margin and a well curved ventral and anterior margin. The posterior margin is only weakly convex. The shell is ornamented with rounded, commarginal
Figure 15. (A–C) Pholadomya (Pholadomya) fidicula J. de C. Sowerby, 1826. (A) right valve view of a composite mold of an articulated specimen, × 1, ERNO-8222; (B–C) composite mold of an articulated specimen; (B) right valve view, (C) left valve view; × 1, ERNO-8223; all lower Pliensbachian of Sierra de Santa Rosa. (D–F) Pholadomya (Pholadomya) idea d’Orbigny, 1850; composite mold of an articulated specimen. (D) right valve view, (E) dorsal view, (F) left valve view; × 1, ERNO-8224, lower Pliensbachian of Sierra de Santa Rosa. (G) Pholadomya (Pholadomya) cf. voltzi Agassiz, 1842; Composite mold of a left valve, × 1, ERNO-8225, lower Toarcian of Sierra de La Jojoba. (H) Goniomya (Goniomya) sp. A; composite mold of a right valve, × 1, ERNO-8226, lower Pliensbachian of Sierra de Santa Rosa. (I–J) Osteomya dilata (Phillips, 1829). (I) composite mold of a right valve, × 1, ERNO-8227; (J) latex cast of an external mold of a right valve, × 1, ERNO-8228; all lower Sinemurian of Sierra del Álamo. (K–M) Pachymya? sp. A; composite mold of an articulated specimen. (K) left valve view, (L) dorsal view, (M) right valve view; × 1, ERNO-8229; all lower Pliensbachian of Sierra de Santa Rosa. (N) Ceratomya concentrica (J. de C. Sowerby, 1825); composite mold of a right valve, × 1, ERNO-8230, lower Pliensbachian of Sierra de Santa Rosa. (O) Pleuromya uniformis (J. Sowerby, 1813); steinkern, right valve view, × 1, ERNO-8234, lower Pliensbachian of Sierra de Santa Rosa. (P–R) Ceratomya petricosa (Simpson, 1855). (P) internal mold of a left valve, × 1, ERNO-8231; (Q) internal mold of a right valve, × 1, ERNO-8232; (R) internal mold of a right valve, × 1, ERNO-8233; all lower Pliensbachian of Pozos de Serna. (S) Pleuromya uniformis (J. Sowerby, 1813); steinkern, left valve view, × 1, ERNO-8234, lower Pliensbachian of Sierra de Santa Rosa.
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growth rugae, which are regularly spaced. Seven radial, straight, very faint ribs lead from the umbo to the ventral margin. These ribs cover the whole surface of the shell except for the posterodorsal part, where only commarginal rugae are present. Remarks: This is the first record of Ph. cf. voltzi from Mexico. Although there is only one specimen available for study, it is clearly distinguishable from the other species of Pholadomya described herein. Compared with specimens from Europe and northern Chile (see synonymy list) it is less elongated and therefore has been assigned to Ph. voltzi with reservation. Genus Goniomya Agassiz, 1841 Type species. Mya angulifera J. Sowerby, 1819; subsequent designation by Herrmannsen, 1847. Subgenus Goniomya Goniomya (Goniomya) sp. A Fig. 15H. Material: One composite mold of a right valve (ERNO-8226) from the lower Pliensbachian, middle member, section 3, Sierra de Santa Rosa. Description: Goniomya (G.) sp. A is medium-sized, elongated, and moderately inflated. The umbones are orthogyrate and are situated mesially. The antero-dorsal and postero-dorsal margins are slightly concave. The short and evenly curved anterior margin passes smoothly into the slightly convex ventral margin. The posterior margin is nearly straight. A weak diagonal ridge extends from the umbo to the postero-ventral corner. The form of the escutcheon and the lunule are unknown. The ornament consists of commarginal ribs that cover the anterior half of the shell. At a length shortly anterior to the umbo, they deviate and run toward the ventral margin. Having passed the midline of the shell, they deviate again and run obliquely toward the postero-dorsal margin, forming a more or less rectangular angle at the second bend. The ribs disappear when they reach the diagonal ridge. Posterior to this ridge, only commarginal growth lines can be seen. Remarks: This is the first report of Goniomya from the Lower Jurassic of Mexico. A different species, Goniomya (G.) calderoni Castillo and Aguilera (1895, p. 9, pl. 5, figs. 17–18), was described from the Upper Jurassic of Mexico. Goniomya sp. A has a conspicuous ornamentation that resembles a mathematical root sign. This ornamentation has not yet been described in any other species of Goniomya, which usually have a trapezoidal or v-shaped ornamentation. With only one specimen available, intraspecific variations of morphological characters are unknown, and erection of a new species must await more material. Genus Osteomya Moesch, 1874 Type species. Mya dilata Phillips, 1829. Osteomya dilata (Phillips, 1829) Figs. 15I–15J. 1829 Mya dilata sp. nov.—Phillips: 155, pl. 11, fig. 4.
1991 Osteomya dilata (Phillips 1829)—Yin and Fürsich: 156, pl. 11, fig. 2 (see for synonymy list). 2004 Osteomya dilata (Phillips 1829)—Aberhan: 134, pl. 6, figs. 4–5. Material: One composite and one external mold of right valves with partial shell preservation (ERNO-8227 to 8228) from the lower Sinemurian, unit 18, Sierra del Álamo. Description: Osteomya dilata is medium-sized, elongatedovate, and moderately inflated. Its umbones are broad and opisthogyrate and are situated about two-fifths of the shell length from the anterior end. The postero-dorsal margin is strongly concave, the antero-dorsal margin straight to slightly concave. The ventral margin is slightly convex and passes with a slight curve into the also slightly convex anterior margin. The posterior margin is truncated. A weak posterior ridge extends from the umbo to the postero-ventral corner of the shell. The surface of the shell wears faint, irregularly spaced, commarginal folds. Remarks: This is the first record of the genus from Mexico. The studied specimen, although laterally deformed, agrees well with specimens of O. dilata recently documented from the Middle Jurassic of Kachchh, western India (Pandey et al., 1996), and the Lower Jurassic of northern Chile (Aberhan, 2004). Genus Pachymya J. de C. Sowerby, 1826 Type species. Pachymya gigas J. de C. Sowerby, 1826. Pachymya? sp. A Figs. 15K–15M. 1997 Pachymya? sp.—Damborenea and González-León: 195, fig. 9.5. Material: One steinkern and one composite mold of an articulated specimen (ERNO-8229, MB.M.4929) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: Pachymya? sp. A is medium- to large-sized, inequilateral and has an elongated, subrectangular form, with the postero-dorsal margin reaching nearly the same height as the umbones. The umbones are narrow and orthogyrate, and are situated at about one-fifth of the shell length from the anterior end. The shell is moderately inflated, with the greatest inflation in the antero-dorsal part of the shell. A smooth posterior ridge extends from the umbo to the postero-ventral corner of the shell. The postero-dorsal margin is slightly concave and meets the posterior margin at an obtuse angle. The posterior margin is convex and gaping. It passes into the nearly straight ventral margin with a slight curve. The anterodorsal margin descends rapidly to the short and well-curved anterior margin. The escutcheon is long and narrow and bordered by smooth ridges. The surface of the shell is covered with commarginal, irregularly spaced growth lines. Remarks: The available specimens certainly belong to the same species that was described as Pachymya? sp. by Damborenea and González-León (1997) from the same locality. These authors also discussed affinities to various
Early Jurassic bivalves of the Antimonio terrane species of the genus Homomya Agassiz. Because of the presence of a very faint umbonal carina, the specimens are doubtfully assigned to Pachymya. Pachymya? sp. A from the Sinemurian of Chile (Aberhan, 2004, p. 138, pl. 9, figs. 1, 6) also has a very low umbonal carina, but is more strongly inflated, the umbones are in a more anterior position, and the ventral margin is convex rather than straight. Family Ceratomyidae Arkell, 1934 Genus Ceratomya Sandberger, 1864 Type species. Isocardia excentrica Roemer, 1836. Ceratomya concentrica (J. de C. Sowerby, 1825) Fig. 15N. 1825 Isocardia concentrica J. de C. Sowerby: 147, pl. 391, fig. 1. 1934 Ceratomya concentrica (J. de C. Sowerby)—Arkell: 315, pl. 43, fig. 10 (see for extensive synonymy list). 1997 Ceratomya sp.—Damborenea and González-León: 196, fig. 9.6. Material: One composite mold of a right valve (ERNO-8230) from the lower Pliensbachian, middle member, section 3, and one internal mold of a left valve (MB.M.4930) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: Ceratomya concentrica is small to mediumsized, slightly inequilateral, and has an oval outline. Its umbones are large, tumid, prosogyrate, and strongly incurved, and are situated about one-fourth to one-third of the shell length from the anterior end. The shell is slightly longer than high and well inflated. The posterior part is flattened and slightly elongated. The posterior margin is straight to slightly convex and meets the evenly convex ventral margin with an angle of about 100°. The ventral margin passes gradually into the well-curved anterior margin. The antero-dorsal margin is short and slightly concave. The surface of the shell is covered with faint, regularly spaced, commarginal undulations, which cover the entire surface of the shell. Remarks: C. concentrica has not been identified from North or South America before. The present specimens are very similar to those described and figured by Arkell (1934) from the Upper Jurassic of England. Many other authors also described C. concentrica from the Middle and Upper Jurassic of Europe (e.g., Scholz, 2005). Only Bernad (1997, p. 24, pl. 3, fig. 10) and Gahr (2002, pl. 5, fig. 8) figured C. concentrica from the Lower Jurassic of Spain. Ceratomya sp. of Damborenea and González-León (1997) from the same locality in Mexico, is identical in ornamentation and outline to the specimens described in this study, and is therefore included into the synonymy list. Similar in ornamentation is ?Ceratomya sp. of Alencáster (1977, p. 160, fig. 14) from the Upper Jurassic of Mexico, but it differs in being slightly higher than long. Ceratomya sp. A, which has been described by Aberhan (2004, p. 138, pl. 7, fig. 2) from the Pliensbachian of Chile, differs from the Mexican specimens by a higher length-height ratio.
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Ceratomya petricosa (Simpson, 1855) Figs. 15P–15R. 1855 Venus petricosa sp. nov.—Simpson: 121. 1867 Isocardia liassica sp. nov.—Moore: 217, pl. 7, fig. 3. 1876 Ceromya petricosa, Simpson—Tate and Blake: 408, pl. 14, fig. 1a+b. 1928 Ceratomya petricosa (Simpson)—Cox: 244, pl. 18, fig. 4a+b. 1973 Ceratomya petricosa (Simpson) 1855—Palmer: 254, pl. 1, fig. 12. Material: Internal molds with rare shell preservation of 13 right valves and six left valves and one steinkern (ERNO-8231 to 8233, MB.M.4931 to 4947) from the lower Pliensbachian, Pozos de Serna. Description: Ceratomya petricosa is medium-sized, equivalved, inequilateral, and has a suboval outline. The umbones are prominent, prosogyrate, and strongly involute. The anterior margin is convex and passes gradually into the straight to slightly convex ventral margin. The posterior margin is slightly convex and meets the ventral margin at an obtuse angle. The shell has its strongest inflation in its anterior half. The shell, which is relatively thin, is ornamented with numerous narrow, fine, commarginal ribs. In internal molds, only commarginal growth lines can be seen. Remarks: Hitherto, C. petricosa was known only from the Lower Jurassic of Great Britain. Simpson (1855) and Cox (1928) mentioned radiating rows of very small pustules on the surface of C. petricosa. This is a common feature of anomalodesmatan bivalves in general, but, as in the Mexican specimens, it often is not preserved. C. petricosa differs from C. concentrica by its more elongated outline and less inflated shell. Additionally, C. petricosa has finer and more numerous ribs. Family Pleuromyidae Dall, 1900 Genus Pleuromya Agassiz, 1843 Type species. Mya gibbosa J. de C. Sowerby, 1823. Pleuromya uniformis (J. Sowerby, 1813) Fig. 15O and 15S. 1813 Unio uniformis sp. nov.—J. Sowerby: 83, pl. 33, fig. 4. 1935 Pleuromya uniformis (J. Sowerby)—Arkell: 325, pl. 45, figs. 1–13 (see for extensive synonymy list). 1977 Pleuromya tellina Agassiz—Alencáster: 161, figs. 16a–c. 2004 Pleuromya uniformis (J. Sowerby 1813)—Aberhan: 146, pl. 8, figs. 8–14 (see for extensive synonymy list). Material: One steinkern (MB.M.4955) from the lower Pliensbachian, middle member, section 3, and two steinkerns (ERNO-8234, MB.M.4956) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: Pleuromya uniformis is medium-sized, inequilateral, and elongated-ovate. The umbones are broad, orthogyrate, and are situated at about one-third of the shell length from the anterior end. Its postero-dorsal margin is long and straight to slightly concave, whereas the antero-dorsal margin is very short and descends toward the well-curved
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TABLE 1. DISTRIBUTION OF EARLY JURASSIC BIVALVES PRESENT IN SONORA DURING THE HETTANGIAN–SINEMURIAN AND/OR PLIENSBACHIAN Taxon Hettangian– Pliensbachian Sinemurian Palaeoneilo elliptica W Parallelodon cf. hirsonensis Grammatodon (G.) sulcatus
E E
END
Gervillella araucana
E ( N)
E
Gervillella leesi
E ( N)
END
Cercomya (Capillimya) peruviana
END
Plagiostoma schimperi
END
Pinna (P.) cf. folium
W
Antiquilima (A.) cf. nodulosa Oxytoma (O.) cf. inequivalvis
E W
Entolium (E.) corneolum
E
E
Agerchlamys wunschae
E (T)
E (T)
Eopecten velatus
W
Weyla (W.) alata
E
Weyla (W.) bodenbenderi
E
E
Weyla (W.) titan
E
E
Weyla (Lywea) unca
E
E
Modiolus (M.) giganteus
E
Modiolus (M.) cf. baylei Modiolus (Cyranus) hillanus
E
E W
Pholadomya (Ph.) fidicula
E
Pholadomya (Ph.) idea
W
Pholadomya (Ph.) cf. voltzi
W
Pleuromya uniformis
W
Ceratomya concentrica
END
Ceratomya petricosa Osteomya dilata
W E
Groeberella sp. A Frenguelliella poultoni Protocardia (P.) truncata
END E (N)
Protocardia (P.) striatula Protocardia (P.) luggudensis
E
W E END
W
Neocrassina gueuxi
W
W
Cardinia concinna
W
Isocyprina ancatruzi
END
Myoconcha neuquena Lucinidae gen. et sp. indet. B
E E
END
END
Note: E—East Pacific; END—endemic to Sonora during that particular time slice; N—restricted to North America during the particular time; T—trans-temperate; W—widespread. See text for details.
anterior margin. The posterior gape is narrow. The ventral margin is straight to smoothly convex and passes gradually into the anterior margin and the convex posterior margin. The surface of the shell is covered with regularly spaced, commarginal growth lines. Remarks: This is the first record of the genus from the Lower Jurassic of Mexico. Many authors commented on the great variability in shape in P. uniformis (see Aberhan, 2004,
and references therein), and the Mexican specimens fit well into the known morphological range of this species. Family Laternulidae Hedley, 1918 Genus Cercomya Agassiz, 1843 Type species. Cercomya pinguis Agassiz, 1843. Subgenus Capillimya Crickmay, 1936 Type species. Capillimya capillifera Crickmay, 1936. Cercomya (Capillimya) peruviana Cox, 1956 Figs. 6D–6F. 1956 Cercomya peruviana sp. nov.—Cox: 1185, pl. 128, fig. 7. 1992 Cercomya peruviana Cox—Damborenea, pl. 118, fig. 6. 2004 Cercomya (Capillimya) peruviana Cox 1956—Aberhan: 148, pl. 9, figs. 7–9. Material: One composite mold of an articulated specimen (ERNO-8168) from the lower Pliensbachian, middle member, section 4, Sierra de Santa Rosa. Description: Cercomya (Capillimya) peruviana is rostrate, inequilateral, and only slightly inflated, and has an elongated, tapering posterior end with a narrow posterior gape. The umbones are narrow, orthogyrate, and are situated mesially. The antero-dorsal margin is slightly convex, the postero-dorsal margin concave. The ventral margin is slightly undulated and passes laterally into the convex anterior margin. The escutcheon is long and narrow and bordered by fine ridges. The anterior part of the shell is covered with strong and rounded commarginal folds. They terminate at a shallow sulcus, which extends from the umbo to the ventral margin. Behind this sulcus, the surface of the shell is covered with commarginal growth lines and faint radial ribs, which become stronger toward the posterior end of the shell. These two ornamental elements produce a subtle reticulate pattern, especially on the middle part of the shell. Remarks: This is the first record of the genus from the Lower Jurassic of Mexico. C. peruviana shows the typical rostrate shape, which separates representatives of Cercomya from other species of the Anomalodesmata. It differs from Cercomya undulata (J. Sowerby, 1827), which has been described by Aberhan (1994, p. 21, pl. 6, figs. 7–9) from the Toarcian of Chile, by the shallow sulcus extending from the umbo to the ventral margin, and fine, distinct riblets covering the posterior part of the shell. A comparison with other species of Cercomya has recently been performed by Aberhan (2004). EARLY JURASSIC BIVALVE PALEOBIOGEOGRAPHY Early Jurassic bivalves of western North and South America are fairly well known. This offers the opportunity to analyze the bivalves of the Antimonio terrane documented herein in a paleobiogeographic context. Following the description of distributional patterns, we quantitatively compare the Antimonio bivalves with contemporaneous faunas along the eastern paleo-Pacific margin and discuss implications for the paleogeographic position of the Antimonio terrane during Early Jurassic times.
Early Jurassic bivalves of the Antimonio terrane
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TABLE 2. LIST OF STUDIED AREAS AND MAIN SOURCES OF DATA Area Brooks-Mackenzie Basin
Source
Western Canada Sedimentary Basin
Imlay, 1967; Poulton, 1991; collections of the Geological Survey of Canada, Calgary and Ottawa Frebold, 1957, 1964, 1966, 1969; Aberhan, 1998a
Wrangellia
Whiteaves, 1900; Aberhan, 1998a
Stikinia
Lees, 1934; Thomson and Smith, 1992; Stanley and McRoberts, 1993; Aberhan, 1998a
Northern Chile
Bayle and Coquand, 1851; Möricke, 1894; Philippi, 1899; Covacevich and Escobar, 1979; von Hillebrandt, 1980; Pérez, 1982; Aberhan, 1994; Damborenea, 1996 Behrendsen, 1891; Burckhardt, 1900; Jaworski, 1914; Weaver, 1931; Leanza, 1942; Escobar, 1980; Manceñido and Damborenea, 1984; Damborenea, 1987b, 1992, 1993, 1996, 2002; Riccardi et al., 1988, 1991; Damborenea and Manceñido, 1992 Jaworski, 1929; Damborenea and González-Léon, 1997; present study
Neuquén Basin
Sonora
Biogeographic Affinities With respect to their paleogeographic distribution we assigned all taxa that have been resolvable at species level (including open nomenclature) to one out of three broad distributional categories (Table 1): endemic to Sonora, eastern Pacific, and widespread. It should be noted that geographic ranges of taxa may change through time, leading to a different categorization for the same taxon. Although the distributional patterns vary slightly between the Hettangian–Sinemurian and the Pliensbachian, the general proportions of these categories remain constant. Species apparently restricted to the eastern Pacific margin of North and/or South America form about half of the total fauna. Documentation of these taxa in Sonora provides further evidence of the paleolatitudinal continuity of species such as Gervillella araucana, Antiquilima cf. nodulosa, Entolium corneolum, Frenguelliella poultoni, various species of Weyla, and Pholadomya fidicula. Known from western North and South America and New Zealand—which was fairly close to the South Pole during the Early Jurassic—Agerchlamys wunschae was previously considered a bipolar eastern-Pacific form. However, because of the new lowpaleolatitude Sonoran record of this species, it is now better categorized as a trans-temperate taxon (see also Kauffman, 1973). Widespread taxa are those that, in addition to occurring in the eastern Pacific, are reported from other regions, particularly Europe. These include species that were only known from northwestern Europe and/or western Tethys until now (e.g., Modiolus hillanus, Ceratomya petricosa, Protocardia truncata, Neocrassina gueuxi), but also species with a more widespread distribution (e.g., Palaeoneilo elliptica, Oxytoma cf. inequivalvis, Pholadomya cf. voltzi). About one-fourth (Pliensbachian) to one-third (Hettangian–Sinemurian) of the described species belongs to the category of widespread taxa. Species that appear to be limited to Sonora at the indicated time are classified as endemic. These taxa either are known only from Sonora, as is the case for Groeberella sp. A and Lucinidae gen. et sp. indet. B, or occur also in other regions earlier or later. For example, Grammatodon sulcatus is recorded from the Sinemurian of northern Chile and Sonora, but in the Pliensbachian is only known from Sonora. Conversely, Cercomya peruviana seems
to have originated in Sonora in the Pliensbachian and subsequently dispersed to northern Chile (Toarcian to Aalenian), and to Peru and the Neuquén Basin (Bajocian) (Aberhan 2004). For Protocardia luggudensis (endemic to Sonora during the Hettangian–Sinemurian) and for Ceratomya concentrica and Plagiostoma schimperi (endemic during the Pliensbachian), these are the first records in America and also the oldest occurrences globally. About one-seventh (Hettangian–Sinemurian) and one-fourth (Pliensbachian) of the described species belong to the endemic category. Comparison with Other Eastern Pacific Faunas To document faunal similarities of the Antimonio terrane to other eastern Pacific regions we utilized all taxa of the bivalve order Pectinoida. This group was selected because its Early Jurassic members are rather common and widespread in the area, well studied, and latitudinally distinct, and they have already proved useful in similar paleobiogeographic analyses (Aberhan 1998b, 1999). The seven regions compared herein are northern Chile (22° to 31°S present-day latitude) and the Neuquén Basin (31° to 41°S present-day latitude) from South America, and, from western North America, the Brooks-Mackenzie Basin, the Western Canada Sedimentary Basin, the allochthonous Stikinia and Wrangellia terranes of Canada, and the Antimonio terrane of Sonora (Table 2; Fig. 16). Wrangellia and Stikinia are the two largest terranes of the Canadian Cordillera. Their time of accretion to the North American craton and interactions with other terranes prior to docking are debated, but it seems that Wrangellia had not yet collided and Stikinia at best had just started to collide with the craton by Middle Jurassic time (e.g., Monger et al., 1982; van der Heyden, 1992; Rowley, 1992). Biogeographic patterns suggest that no significant latitudinal shifts of Wrangellia and Stikinia occurred from Sinemurian to Pliensbachian times, but substantial post-Pliensbachian northward displacement relative to the craton was inferred (e.g., Aberhan 1998b, 1999; Smith et al. 2001). However, this is in conflict with reinterpretations of paleomagnetic data that suggest that the allochthonous terranes were in much the same latitudinal positions relative to the craton as they are today (Irving and Wynne 1990; Vandall and Palmer 1990).
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Figure 16. Present-day position of selected allochthonous terranes (gray shading) and autochthonous sedimentary basins (oblique hatching) along the western margin of North America.
Paleobiogeographic Results Not surprisingly, cluster analysis of Hettangian–Sinemurian data (Table 3) shows a high degree of faunal similarity between northern Chile and the Neuquén Basin, two autochthonous areas located immediately next to each other (Fig. 17). A high degree of similarity also exists between Wrangellia and Stikinia, suggesting that their geographic distance was relatively small at that time. At a lower similarity level the Brooks-Mackenzie Basin and the Western Canada Sedimentary Basin cluster together. The Antimonio terrane of Sonora has strongest links to Wrangellia and Stikinia. In the Pliensbachian, faunal ties again are strong between Wrangellia and Stikinia. Also, northern Chile and the Neuquén Basin are grouped in the same cluster. In contrast to the Hettangian–Sinemurian, however, the Antimonio terrane now forms a cluster together with the Neuquén Basin and particularly northern Chile (Fig. 17). Paleogeographic Interpretations Assuming that the similarity between faunas roughly reflects their geographic distance from each other, the relatively high faunal concordance of the Antimonio terrane with Wrangellia and Stikinia during the Hettangian–Sinemurian argues in support of relative spatial proximity of these three regions. In the Pliensbachian, geographic
distance to Wrangellia had increased and the Antimonio terrane was closer to the Andean basins of Argentina and Chile. There are two possible, mutually non-exclusive mechanisms to explain these patterns: First, Stikinia and Wrangellia might have moved northward from Hettangian–Sinemurian to Pliensbachian times and thus lost their proximity to Sonora. Substantial northward movement of Wrangellia and Stikinia has indeed been postulated previously (see above). However, biogeographic data suggest that the paleolatitudinal position of Wrangellia and Stikinia remained largely stable relative to the craton from Sinemurian to Pliensbachian times, and northward movement only started in the Toarcian–Early Aalenian (Aberhan 1999). Alternatively, the Antimonio terrane was displaced toward the south, thereby moving closer to northern Chile and the Neuquén Basin, and at the same time losing similarity to Wrangellia and Stikinia. This scenario gains support from a direct comparison of similarity coefficients between the Antimonio terrane and autochthonous regions (Table 4). Dice and Jaccard coefficients between the Antimonio terrane and Argentina/Chile are markedly higher in the Pliensbachian than during the HettangianSinemurian. However, a southward shift of the Antimonio terrane already starting in the Pliensbachian is in contrast to the conclusions of Campa and Coney (1983) and Marzolf (2003) that a displacement of this region did not occur until the Middle or Late Jurassic, when the Antimonio Group, together with the Caborca block, was displaced southward along the Mojave-Sonora megashear. Caveats Occurrence-based data stored in the Paleobiology Database (http://paleodb.org) indicate a generally very good correlation between percent similarity of contemporaneous marine invertebrate faunas from different regions and their geographic proximity (measured as great circle distance) during the Phanerozoic (A.I. Miller, 2005, personal commun.). Nevertheless, it can be argued that the assumption of a rough correspondence between the similarity of faunas and their geographic distance from each other is an oversimplification that neglects other important factors that control faunal distribution, such as sedimentary facies and prevailing paleocurrents. With regard to facies, a wide spectrum of lithologies is represented in the upper Hettangian to Pliensbachian of Sonora, comprising fine-grained and coarse-grained siliciclastics, carbonates, and mixed siliciclastic-carbonate sediments (see above). Most pectinoid bivalves from Sonora were notably eurytopic with respect to the substrate, and occurred predominantly in fine-grained siliciclastics and carbonates, bioclastic limestones, and calcareous sandstones. A similarly weak dependence on substrate conditions was found in most pectinoids from the other regions analyzed here (Aberhan, 1998b). An exception is the Western Canada Sedimentary Basin, the Pliensbachian samples of which are limited to black shales and dark calcareous mudstones, reflecting low-diversity, oxygen-deficient benthic environments. Altogether, however, the represented range of upper Hettangian to Pliensbachian lithologies and environments is comparable for the analyzed regions.
TABLE 3: GEOGRAPHIC AND TEMPORAL DISTRIBUTION OF EARLY JURASSIC PECTINOID BIVALVES ALONG THE EASTERN MARGIN OF THE PALEO-PACIFIC Taxon Hettangian–Sinemurian Pliensbachian Otapiria neuquensis Damborenea, 1987b Chi; Neu Neu Otapiria pacifica Covacevich and Escobar, 1979 Otapiria tailleuri Imlay, 1967
Chi; Neu WCSB
Lupherella boechiformis (Hyatt, 1894)
Sti
Oxytoma (O.) inequivalvis (J. Sowerby, 1819) Palmoxytoma cygnipes (Young and Bird, 1822)
Chi; Neu; WCSB; Bro; So
Chi; Neu; Wra; Sti; WCSB; Bro
Chi; Neu; WCSB; Bro
Wra; Bro
Arctotis (A.) sp. A
WCSB
Meleagrinella ferniensis (McLearn, 1924)
Sti
Meleagrinella cf. oxytomaeformis Polubotko, 1968 Meleagrinella sp.
a
Asoella asapha (Leanza, 1942) Placunopsis radiata (Phillips, 1829) b
?terquemiid gen. et sp. nov.?
Wra
Wra
Bro
Bro
Neu
Neu
Chi; Neu; Sti
Chi; Neu; Sti
Wra
Kolymonectes carlottensis (Whiteaves, 1900)
Wra; Sti
Kolymonectes weaveri Damborenea, 1997
Neu
Kolymonectes staeschei (Polubotko, 1968)
WCSB; Bro
c
Kolymonectes sp.
Neu
Entolium (E.) corneolum (Young and Bird, 1828) Entolium (E.) cf. lunare (Roemer, 1839)
Chi; Wra; Sti; WCSB; Bro; So
Chi; Neu; Wra; Sti; Bro; So
Chi; Neu
Chi; Neu
Entolium (E.) mapuche Damborenea, 2002
Neu
Agerchlamys wunschae (Marwick, 1953)
Chi; Wra; WCSB; So
d
Agerchlamys sp. A
Neu; So Wra
Camptonectes (C.) auritus (Schlotheim, 1813)
Neu
e
Camptonectes (C.) subulatus (Münster, 1836)
Neu; WCSB; Bro
Neu; WCSB
Canadonectites paucicostatus Aberhan, 1998a
Wra
Wra
Chlamys (Ch.) textoria (Schlotheim, 1820)
Chi; Neu; Wra; Sti; WCSB
Chi; Neu; Wra; Sti
Chlamys valoniensis (Defrance, 1825)
Chi; Neu; Sti; WCSB; Bro
Eopecten hartzi (Rosenkrantz, 1957) Eopecten velatus (Goldfuss, 1833)
Sti; WCSB
Neu
Chi; Neu
Chi; Neu; So
Ochotochlamys cf. bureiensis Sey, 1984
WCSB
Ochotochlamys aequistriata Aberhan, 1998a Ochotochlamys sp.
Wra; Sti; WCSB
f
Neu
Pseudopecten (P.) equivalvis (J. Sowerby, 1816)
Chi; Neu
g
Radulonectites sosneadoensis (Weaver, 1931) Radulonectites sp. A
Wra; Sti; WCSB
h
pectinid gen. et sp. indet.
Chi; Neu; Sti; Bro Bro
i
Weyla (W.) alata (von Buch, 1838) Weyla (W.) bodenbenderi (Behrendsen, 1891) Weyla (W.) titan (Möricke, 1894) Weyla (Lywea) unca (Philippi, 1899) Weyla (Lywea) yukonensis Aberhan, 1998a
Chi Chi; Neu; Sti; Wra; WCSB; So
Chi; Neu; Sti; Wra; So
Chi; Wra; Sti; So
Chi; Neu; Wra; Sti; So
So
Chi; So
Chi; Neu; Wra; Sti; So
Chi; Neu; Sti; So
Sti; Wra; WCSB
Note: Bro—Brooks-Mackenzie Basin; Chi—northern Chile; Neu—Neuquén Basin; So—Sonora; Sti—Stikinia; WCSB— Western Canada Sedimentary Basin; Wra—Wrangellia. Taxonomic notes: a—Meleagrinella sp. in Poulton (1991, pl. 6, figs. 8-10; pl. 9, figs. 15-16; pl. 11, figs. 12-13); b— ?terquemiid gen. et sp. nov.? in Aberhan (1998a, pl. 10, figs. 20-21); c—Kolymonectes sp. in Damborenea (2002, pl. 2, figs. 16-19); d—Agerchlamys sp. A in Aberhan (1998a, pl. 12, figs. 6, 7, 9); e—Includes Entolium(?) sp. and Camptonectes (C.) sp. of Poulton (1991, pl. 6, figs. 31, 33-37); f—Ochotochlamys sp. in Damborenea (2002, pl. 9, figs. 16-22; text-figs. 8j-–); g—Includes Camptonectes (Camptochlamys) sp. of Poulton (1991, pl. 11, figs. 17-21); h— Eopecten(?) sp. in Poulton (1991, pl. 6, fig. 38); i—pectinid gen. et sp. indet. in Aberhan (1994, pl. 20, fig. 3).
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Figure 17. Dendrograms (unweighted pair-group method using arithmetic average with the Jaccard coefficient as distance measure) of geographic units defined by the distribution of eastern Pacific pectinoid bivalves during the Hettangian–Sinemurian and the Pliensbachian. WCSB—Western Canada Sedimentary Basin.
Other factors, such as oceanic current systems and climatic regimes clearly influenced distribution patterns in the geological past as they do today. However, we are unaware of any Early Jurassic changes in these factors that could explain the observed shift in faunal similarity. For example, oceanic conditions in the eastern paleo-Pacific are expected to have been relatively stable and comparable to those of the modern Pacific (Parrish, 1992). Sonora would have been under the influence of a stable so-called paleo-California current, flowing toward the equator in midnorthern to low paleolatitudes (see also discussion in Aberhan, 1999). Nevertheless, we consider our paleogeographic results as preliminary because (1) they are based on a small sample of organisms; and (2) too few paleontological data are currently available from key areas in the southwestern United States, which can be regarded as autochthonous since the Jurassic. CONCLUSIONS This study provides a detailed account of the taxonomy and biogeographic affinities of Early Jurassic bivalves from the Anti-
monio terrane of Sonora, Mexico, and considerably expands our previous state of knowledge. Fifty taxa are described and illustrated, 34 of which are reported from this region for the first time. Almost half the taxa are represented by very few (one to three), often poorly preserved specimens. As a consequence, continued fieldwork is expected to further increase sampled Early Jurassic bivalve diversity in this region, and to lead to a more precise identification of those taxa, which could only be determined at the supra-specific level herein. Clearly, more comprehensive faunal data should be used in the future to corroborate or revise our tentative conclusion that southward displacement of the Antimonio terrane had already commenced in the Pliensbachian. Thus, a broader range of faunal groups as well as areas should be included in the analysis. For example, it is critical to include data from the southwestern United States, in particular California, Oregon, and Nevada, which play a key role in paleogeographic reconstructions of western North American terranes. To cite but one example, the New York Canyon sections in the Gabbs Valley Range, Nevada, are known for their diverse Early Jurassic fauna of ammonites and bivalves, but the latter are still very insufficiently documented. ACKNOWLEDGMENTS We would like to thank John Marzolf, József Pálfy, and George Stanley for accompanying us in the field, József Pálfy for supplying biostratigraphic information, and Wolfgang Stinnesbeck for contributing Early Jurassic bivalves from Pozos de Serna. Photographs were taken by Waltraud Harre and Carola Radke. Dieter Korn and Henning Scholz assisted with the figures and cluster analysis. Susana Damborenea and Franz T. Fürsich are thanked for their helpful reviews of the manuscript. This study was supported financially by a grant from the Deutsche Forschungsgemeinschaft (Ab 109/2 to Aberhan), which is acknowledged with gratitude. REFERENCES CITED Aberhan, M., 1994, Early Jurassic Bivalvia of northern Chile, Part I, Subclasses Palaeotaxodonta, Pteriomorphia, and Isofilibranchia: Beringeria, v. 13, p. 3–115. Aberhan, M., 1998a, Early Jurassic Bivalvia of western Canada, Part I, Subclasses Palaeotaxodonta, Pteriomorphia, and Isofilibranchia: Beringeria, v. 21, p. 57–150. Aberhan, M., 1998b, Paleobiogeographic patterns of pectinoid bivalves and the Early Jurassic tectonic evolution of western Canadian terranes: Palaios, v. 13, p. 129–148, doi: 10.2307/3515485.
Table 4. Diversity and similarity coefficients between Sonora and various regions along the eastern paleo-Pacific margin during the Hettangian–Sinemurian and the Pliensbachian, expressed as Dice (D) and Jaccard coefficient (J) Hettangian–Sinemurian Pliensbachian Sonora Sonora 6=7 6=7 Neuquén Basin D = 20 J=9 Neuquén Basin D = 50 J = 20 6 = 14 6 = 20 N. Chile D = 50 J = 20 N. Chile D = 67 J = 25 6 = 14 6 = 13 Wrangellia D = 59 J = 23 Wrangellia D = 42 J = 17 6 = 11 6 = 11 Stikinia D = 47 J = 19 Stikinia D = 53 J = 21 6 = 11 6 = 11 WCSB D = 32 J = 14 WCSB D = 15 J=7 6 = 13 6=5 Brooks-M. Basin D = 14 J=7 Brooks-M. Basin D = 31 J = 13 6=8 6=5 Note: 6—number of taxa per region. WCSB—Western Canada Sedimentary Basin; Brooks-M.—Brooks-Mackenzie.
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The Geological Society of America Special Paper 442 2008
Dinosaurs of Alaska: Implications for the Cretaceous origin of Beringia Anthony R. Fiorillo Museum of Nature and Science, P.O. Box 151469, Dallas, Texas 75315, USA ABSTRACT Fossils within accreted terranes are typically used to describe the age or origin of the exotic geologic blocks. However, accretion may also provide new pathways for faunal exchange between previously disconnected landmasses. One such landmass, the result of accretion, is Beringia, that entity encompassing northeastern Asia and northwestern North America and the surmised land connection between the two regions. The present concept of Beringia as a Quaternary subcontinent includes a climatic component in the form of glacial advances and retreats driving changes in sea level. These changes may have facilitated exchanges of marine biota between the Pacific Ocean and Arctic Basin, or exchanges of terrestrial faunas and floras between Asia and North America. The Beringian ecosystem includes specializations of the flora and fauna, especially in the vertebrate fauna. A review of tectonic reconstructions and the striking taxon-free parallel patterns in data on the Cretaceous and Quaternary fauna and flora suggest that a generalized concept of Beringia should be formally extended back in time to the Cretaceous. A significant shift in emphasis of defining variables occurs with this extension. Climate, in the form of meteorological phenomena, and geologic history are important variables in the previously recognized definition of Beringia. The extension of Beringia into the Cretaceous implies that Beringia is rooted in its accretionary rather than its climatic history; in other words, the geographic pattern as the result of tectonics is the defining parameter for Beringia. Keywords: Cretaceous, dinosaurs, Alaska, Beringia.
INTRODUCTION Many papers that discuss paleontology within the context of accreted terranes typically use those fossils to address questions centered on the origin of particular geologic blocks or the
timing of events such as collisions with other blocks. One of the consequences of accretionary tectonics is that the movement of geologic blocks can create new connections between landmasses that promote redistributions of biota. In contrast to the more traditional studies of paleontology and accreted terranes, this review
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[email protected] Fiorillo, A.R., 2008, Dinosaurs of Alaska: Implications for the Cretaceous origin of Beringia, in Blodgett, R.B., and Stanley, G.D., Jr., eds., The terrane puzzle: New perspectives on paleontology and stratigraphy from the North American Cordillera: Geological Society of America Special Paper 442, p. 313–326, doi: 10.1130/2008.442(15). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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focuses on the consequences of the relevant accretionary tectonics in the high latitudes of northeast Asia and northwest North America to the paleobiology and biogeography of the dinosaur fauna known from Alaska. The accretionary history of this region provided the opportunity for faunal exchange for many of the major groups of dinosaurs found in Late Cretaceous rocks in western North America. The patterns of biogeography and the paleoecology of Arctic dinosaurs during the Cretaceous are strikingly similar to the large-scale patterns of faunal exchange and paleoecology of later Tertiary and Quaternary mammals. These similarities suggest that the concept of Beringia as a phenomenon restricted to the Tertiary and Quaternary is too narrow. Beringia is a concept that is at least reasonably familiar to most students of earth history. Originally proposed as an icefree corridor between northeastern Asia and northwestern North America that served as a refugium during the Quaternary for Arctic plants (Hulten, 1937), a strict adherence to the original proposal remains a rich source of discussion (e.g., Abbott and Brochmann, 2003). However, the concept has grown and changed (Guthrie, 1982, 2001; Hopkins, 1967, 1996; Hopkins et al., 1982, Repenning, 1987, 1990; Repenning et al., 1987; West, 1996; O’Neill, 2004), to the point of even becoming a political entity (Schaaf, 1993). The concept of a land bridge connection that occurred intermittently and was likely responsible for the exchanges of a
broader suite of terrestrial organisms between these two modern landmasses through the Tertiary and Quaternary is well established in the science of earth history (Flerow, 1967; Hopkins, 1967, 1996; Colinvaux, 1980; Cwynar and Ritchie, 1980; Guthrie, 1982, 2001; Hopkins et al., 1982, Repenning, 1987, 1990; Repenning et al., 1987; West, 1996; Abbott and Brochmann, 2003; O’Neill, 2004). This same region referred to as Beringia by Cenozoic workers is also the most likely route for dinosaur migrations between central Asia and western North America in the Cretaceous. In explaining biogeographic patterns for dinosaurs, other workers (e.g., Russell, 1993; Sereno, 2000) have made vague references to Beringia (or the Bering Land Bridge) without critical examination of the geological implications of these biogeographical observations. The purpose of this review is threefold: first to summarize the accretionary history of the major blocks that contain the Late Cretaceous dinosaur record of Alaska; second, to synthesize the current understanding of these Alaskan dinosaurs; and third, to compare the characteristics of a Pleistocene Beringia with those inherent in a proposed extension of Beringia to the Cretaceous. The term Beringia, and its implications, can be more formally extended into the Cretaceous, perhaps as early as the AptianAlbian. This extension demonstrates Beringia is rooted in its accretionary history rather than in a climatic history. Further, this is not an arbitrary extension. Whereas the previous concept of
Figure 1. Location of the areas containing Campanian-Maastrichtian dinosaur localities in Alaska that are discussed in this report. 1—Aniakchak National Monument; 2—numerous localities along the Colville River; 3—Denali National Park.
Dinosaurs of Alaska Beringia was based in large part on climatic conditions, specifically meteorological phenomena, the new concept proposed here emphasizes the role of tectonics expressed through geography. Further, a Cretaceous origin for Beringia has profound implications for the large-scale paleobiological patterns and processes for high-latitude terrestrial ecosystems. DISCOVERY OF DINOSAURS IN ALASKA— BACKGROUND In contrast to the dinosaur fauna from elsewhere in North America, Alaska has only recently achieved recognition as a source of dinosaur records. Work along the lower Colville River (Fig. 1) since the mid-1980s has documented a series of prolific fossil sites and quarries, the vast majority of which are CampanianMaastrichtian in age. The Cretaceous on the North Slope of Alaska is represented by thousands of meters of sedimentary rocks (Mull et al., 2003) (Fig. 2), and Roehler and Stricker (1984), Parrish et al. (1987), and Fiorillo (2004) all suggested that a rich older Cretaceous fossil record will eventually emerge. Given the well-established Campanian-Maastrichtian dinosaurian record compared to the fragmentary earlier record, patterns of biogeography and paleoecology of dinosaurs as they relate to a Cretaceous Beringia are largely drawn from this youngest Cretaceous interval. The fauna recovered from quarry excavations and accumulated river bar and bank float includes specimens of chondrichthyan and osteichthyan fishes, large and small theropods, a hypsilophodontid, a pachycephalosaur, ceratopsian and hadrosaurian dinosaurs, as well as multituberculate, marsupial, and placental
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mammals (Nelms 1989; Gangloff, 1994, 1998b; Rich et al., 1997; Fiorillo and Gangloff, 2000, 2001; Gangloff et al., 2005). This growing Campanian-Maastrichtian record of the dinosaur fauna of Alaska is listed in Table 1. Dinosaur bones were first collected by Robert Liscomb of Shell Oil Company from the lower Colville River in 1961 near Ocean Point during regional fieldwork preliminary to oil exploration. The collections were misidentified as mammoth bones and remained in storage for over two decades. Richard V. Emmons of Shell Oil brought these bones to the attention of Louis Marincovich (1997, personal commun.) of the U.S. Geological Survey (USGS), who in turn brought showed them to Charles Repenning, also of the USGS, in 1983. Repenning, a trained vertebrate paleontologist, identified them as dinosaurian, and thus was the first person to recognize the presence of dinosaur bones in Alaska. Shell Oil Company donated the original collection by Liscomb to the University of Texas where Davies (1987) provided the first formal description of dinosaurs from the North Slope. In the mid-1980s Marincovich organized combined field parties from the USGS, the University of California at Berkeley, and the University of Alaska–Fairbanks, which were able to locate the source of Liscomb’s collections at a locality near Ocean Point that has become known as the Liscomb Quarry or Liscomb Bonebed. The University of California Museum of Paleontology (UCMP) and the University of Alaska Museum (UAM) continued fieldwork in this area from 1987 to 1990 and demonstrated the abundant occurrence of dinosaur and associated vertebrate fauna within the Prince Creek Formation (Clemens and Nelms, 1993; Clemens, 1994; Gangloff, 1994, 1998b).
Figure 2. Generalized Cretaceous stratigraphy for regions discussed here. Stratigraphy for Aniakchak National Monument (number 1 in Fig. 1), the Colville River (number 2 in Fig. 1), and Denali National Park (number 3 in Fig. 1) taken from Detterman et al. (1981), Mull et al. (2003), and Csejtey et al. (1992), respectively, although the age for the dinosaurbearing part of the Cantwell Formation is considered Late Cretaceous rather than Paleocene (Ridgway, et al., 1997). The most complete dinosaur record is from the nonmarine Campanian-Maastrichtian rock units but additional dinosaur remains, predominantly footprints of ornithopods and theropods, have been mentioned in the Nanushuk Formation (Roehler and Stricker, 1984; Witte et al., 1987; Gangloff, 1998a; Spicer, 2003; Fiorillo et al., 2005).
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Fiorillo TABLE 1. CAMPANIAN–MAASTRICHTIAN DINOSAUR FAUNA FROM ALASKA Saurischia Theropoda Dromaeosauridae Dromaeosaurus albertensis Saurornitholestes langstoni Troodontidae Troodon formosus Tyrannosauridae Ornithischia Ornithopoda Hadrosauridae Edmontosaurus sp. Marginocephalia Pachycephalosauridae indet. Ceratopsia Ceratopsidae Pachyrhinosaurus sp. Note: List compiled from Gangloff (1998b), Gangloff et al. (2005), Fiorillo and Gangloff (2000, 2001), Nelms (1989), and Fiorillo and Gangloff (2003).
The Prince Creek Formation is a thick sequence of nonmarine, largely clastic sedimentary rocks. Previously these sites were considered to be within the upper tongue of the Prince Creek Formation, formally known as the Kogosukruk Tongue. The Kogosukruk Tongue of the Prince Creek Formation was considered to overlie the Schrader Bluff Formation, a marine incursion, with the lower tongue of the Prince Creek Formation, the Tuluvak Tongue forming the base of this sequence (Chapman et al., 1964). This review follows the recently revised lithostratigraphy by Mull et al. (2003) in restricting the Prince Creek Formation to the rocks overlying the Schrader Bluff Formation and underlying the Paleocene Sagavanirktok Formation. On the basis of pollen, the age of the Prince Creek Formation in the vicinity of Ocean Point is generally considered to be close to Maastrichtian in age (Fredriksen, 1990). Potassiumargon dates from tephra beds through the local section yield dates that range from 68 to 71 Ma with the weighted mean being 69 Ma (Conrad et al., 1990). Dates from 40Ar/39Ar and single sanidine crystals place the Liscomb Quarry very close to the CampanianMaastrichtian boundary (Gradstein et al., 2004). Pollen data (Fiorillo et al., 2008) from another extensive bonebed, the KikakTegoseak Quarry (Fiorillo and Gangloff, 2003), located some 80 km upriver from the Liscomb Quarry, suggest that this quarry is also near the Campanian-Maastrichtian boundary. For this interval of time on the North Slope, the mean temperature has been reported to range from 2–4 °C for the coldest monthly mean to 10–12 °C for the warmest month mean (Parrish et al., 1987). Unconformably overlying the Prince Creek Formation in the northern third of the region is the Pliocene to Holocene Gubik Formation, which is attributed to a number of discrete
marine transgressions (Feder et al., 2003). Underlying and interfingering with the Prince Creek Formation are the marine beds of the Upper Cretaceous Schrader Bluff Formation (SantonianMaastrichtian) and the Lower Cretaceous Nanushuk Group, which is exposed upriver (Mull, 1985; Mull et al., 2003). Paleomagnetic studies of these rocks suggest a latitudinal range of 67–85°N during the Late Cretaceous (Witte et al., 1987; Besse and Courtillot, 1991). The Prince Creek Formation consists largely of fluviatile sediments shed by the Brooks Range to the south. By the Late Cretaceous the Brooks Range had reached its maximum elevation and was likely being reduced by erosion. Paleogeographic reconstructions of the Prince Creek Formation place the Brooks Range up to hundreds of kilometers away from the dinosaur quarries. Dinosaur-bearing beds that have been found in organicrich siltstones have been interpreted as overbank or crevasse splay deposits (Phillips, 1990, 2003), whereas those found in conglomeratic sandstones have been interpreted as channel lag deposits (Gangloff, 1998b). The Prince Creek Formation was deposited in a well-vegetated coastal lowland environment. In addition to these North Slope finds, Cretaceous dinosaurs have been found in a handful of other localities around the state. An ankylosaur skull from Campanian-Maastrichtian rocks in south-central Alaska is the only documented dinosaur bone of this age found outside of the region of northern Alaska (Gangloff, 1995). During the summer of 2001, a joint Dallas Museum of Natural History and National Park Service paleontological survey of Aniakchak National Monument (Fig. 1) yielded the first evidence of dinosaurs in southwestern Alaska, a series of three-toed footprints attributable to hadrosaurs (Fiorillo and Parrish, 2004). These tracks are in the Chignik Formation (Fig. 2), a Campanian-Maastrichtian sequence of rocks similar in age and depositional setting to the Prince Creek Formation on Alaska’s North Slope. CRETACEOUS ACCRETIONARY HISTORY OF DINOSAUR-BEARING GEOLOGIC BLOCKS IN ALASKA Lawver and Scotese (1990) provide a detailed discussion of the tectonic models related to the evolution of the Canada Basin, the most widely accepted of which proposes counterclockwise rotation of the geologic block that now comprises much of northern Alaska. The counterclockwise rotation of this block is related to a spreading ridge that ultimately led to the creation of the Canada Basin. The pivot point for this rotation is thought to be centered in the vicinity of the current Mackenzie River Delta. Similarities between Jurassic invertebrate microfossils contained within this northern Alaska block and those found in Prince Patrick Island in the Canadian Arctic suggest that this rotation began no earlier than the Jurassic (Mickey et al., 2002) (Fig. 3). By the Cretaceous, the resulting arrangement of plates then formed a connection between northern Alaska and the Russian Far East, an arrangement first discussed in the early 1980s (Churkin and Trexler, 1980, 1981).
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Figure 3. Polar projection of the paleogeography of northern Asia and northern North America during the Campanian-Maastrichtian. Base map by the PLATES Project University of Texas Institute of Geophysics. Legend: A—Prince Patrick Island; B—Wrangellia composite and South Margin composite terranes; C—Arctic Alaska/Arctic Alaska composite terrane.
Names for the northern block of Alaska have varied, but this review follows the nomenclature for the various blocks used in the tectonic reconstructions of the Arctic region by Lawver et al. (2002). Therefore, the northern Alaska block is referred to here as the Arctic Alaska/Arctic Alaska composite (Fig. 3). Forming the northern boundary of the land connection between Asia and North America, the geologic terranes of interest are the Arctic Alaska/Arctic Alaska composite, the Chukotka, and the YukonKoyukuk Basin. The southern boundary of this land connection is defined by the Wrangellia composite and the Southern Margin composite. During the Middle Jurassic (180 Ma; Lawver et al., 2002) the arrangement of these plates presented a substantial oceanic barrier between northeastern Asia and northwestern
North America. Nevertheless, if sea level was low enough, faunal exchange between these two landmasses may have been possible with a connection across the pole. During the Aptian (120 Ma; Lawver et al., 2002), Chukotka had slid southward such that a Bering style land bridge, at the levels of resolution available, was approached. Dinosaurian data from the Cedar Mountain Formation (Aptian-Albian) in Utah (Cifelli et al., 1997) suggest that the late Albian or early Cenomanian may have been when a functional land connection arose between Asia and North America. By the Turonian (90 Ma; Lawver et al., 2002) geologic blocks were positioned such that there almost certainly was a dry land connection between Asia and North America. Extensive Upper Cretaceous bentonites in northern Alaska appear to have their
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origin in the Russian Far East (Spicer, 2003) and these are likely the result of the emplacement of these terranes into the configuration yielding this dry land connection. A pathway for terrestrial organisms, of course, implies a coeval interruption in the passage of marine biota. In his review of Danian mollusks from the Prince Creek Formation, Marincovich (1993) concluded that the youngest shallow-marine Pacific-Arctic connection must have closed soon after the late early Albian (110–105 Ma), and that the North American Western Interior Seaway had already ceased to be an exchange gateway for shallow-marine mollusks from the south after the early Maastrichtian. Further suggestion of a landmass connection between these two continents is the emplacement of appropriately aged continental igneous rocks from St. Lawrence Island near the present Bering Strait (Amato et al., 2003). For southern Alaska, the Wrangellia composite and South Margin composite terranes are placed well north of 60° by the Late Cretaceous. The Chignik Formation, containing the dinosaur footprints mentioned above, are part of the Peninsular terrane, the structural unit that encompasses much of southwestern Alaska and is a member of the Wrangellia composite terrane. Paleomagnetic reconstruction based on the Upper Cretaceous and lower Tertiary volcanic rocks of this terrane suggest that the Chignik Formation was deposited at approximately the current latitude (Hillhouse and Coe, 1994). In contrast, coeval sedimentary rocks of the Peninsular terrane elsewhere in southern Alaska provide more ambiguous paleomagnetic results, suggesting positions as far as 3000 km south of the present position of the rock sequence (Hillhouse and Coe, 1994). Nevertheless, the argument that the terrane, at least the portion of concern to this study, was near its present position during the Late Cretaceous is stronger for two reasons. First, the resolution of the data for this position is higher and, second, these more precise data were collected in the vicinity of the Aniakchak National Monument, rather than much farther east (Hillhouse and Coe, 1994) where the data giving southern paleomagnetic results were collected. The current latitude of the tracksite is 56°–57°N. Therefore, the tracksite from the Chignik Formation in Aniakchak National Monument is within geographic boundaries of Beringia. In summary, the geologic blocks that currently comprise the geographic region referred to as Beringia reached their appropriate configuration sometime in the Cretaceous, likely ca. 100 Ma. The connection is almost certainly firm by the Turonian (Lawver et al., 2002). Figure 3 shows the paleogeography of northern Asia and North America for the Late Cretaceous. It is now appropriate to examine the biological characteristics that underlie the concept of Beringia in context of their applicability to the Cretaceous. Such a comparison will suggest that the origin of Beringia is rooted in its accretionary history rather than a meteorological one. PALEOBIOLOGY OF HIGH-LATITUDE DINOSAURS IN ALASKA Cretaceous northern high-latitude dinosaurs raise two basic questions. First, did they live in this environment year-round or
did they migrate to lower latitudes? And second, if they were year-round residents of the high latitudes what adaptations, if any, were needed? The recovery of dinosaur remains from the ancient high latitudes of North America has been problematic given current understanding of dinosaur physiology. To resolve these problems, several workers supported the hypothesis that these animals were long-distance, seasonal migrants, with migratory ranges of several thousands of kilometers (Hotton, 1980; Brouwers et al., 1987; Parrish et al., 1987; Currie, 1989). These arguments used a modern analog, caribou (Rangifer tarandus), to support their respective conclusions. Histological data suggest that the juvenile hadrosaurs from one quarry in this region were older than one year in age (Fiorillo and Gangloff, 2001). In addition, studies on caribou and hadrosaurs from the North slope, which compared relative sizes of juveniles and adults, suggest, on the basis of qualitative energetics, that the hadrosaurs had not attained sufficient body size for long-distance migration. These North Slope hadrosaurs have therefore been reinterpreted as year-round residents of the Cretaceous high latitudes (Fiorillo and Gangloff, 2001). Given this consideration, one group of animals, the small theropods, has been examined more closely for an adaptive response to this environment (Fiorillo and Gangloff, 2000). Such an adaptive response has been shown for this group in the most dominant meat-eating dinosaur of the North Slope is Troodon formosus, an animal known for its exceptionally large eyes compared to skull size (Fiorillo and Gangloff, 2000). This animal is known all the way south to west Texas, but in the more southern occurrence of dinosaurs, Troodon formosus is very rare (e.g., 6% in Montana; Fiorillo and Currie, 1994). In Alaska Troodon formosus outnumbers all other theropods combined by a ratio of nearly 2:1 (Fiorillo and Gangloff, 2000). This higher frequency in Alaska indicates that Troodon formosus was pre-adapted to the low light conditions of the Arctic where it thrived (Fiorillo and Gangloff, 2000). This adaptive advantage further expressed itself with an increase in body size for Troodon formosus, presumably the result of increased access to food resources due to the competitive edge (Fiorillo, 2008). PALEOECOLOGY OF ALASKAN DINOSAURS Recent work in the coeval Chignik Formation of southwestern Alaska has shown the presence of several dinosaur tracks attributable to hadrosaurs (Fiorillo and Parrish, 2004). These tracks, combined with known dinosaurian localities in northern Alaska, indicate the widespread existence of a high-latitude terrestrial ecosystem during the Cretaceous capable of supporting large-bodied herbivores. Although this ecosystem has been inferred previously, the existence of these tracks extends the known geographic range of this terrestrial ecosystem (Fiorillo and Parrish, 2004). Floral reconstructions of this ecosystem suggest one or more gradients in floral diversity. Northern Alaska during the
Dinosaurs of Alaska Campanian-Maastrichtian consisted of a canopy of conifers and an understory of ginkgoes, flowering plants, ferns, and cycads (Spicer, 1987, 2003; Spicer and Parrish, 1986; Parrish and Spicer, 1988). A recent well-illustrated summary of the megafloral remains of the Cretaceous of northern Alaska has been provided by Spicer (2003). Similarly, in southwestern Alaska, the conifer cf. Metasequoia was an important canopy species in the forest. Compared with the Campanian-Maastrichtian of the North Slope, the angiosperm understory in southern Alaska is far more diverse. Only one angiosperm leaf form has been identified from the Late Cretaceous of the Colville River region (Parrish and Spicer, 1988), in contrast to 12 in the Chignik from Aniakchak Bay (J. Parrish, 2003, personal commun.). Besides megafloral information from these two regions, preliminary pollen data from the North Slope—specifically from the Kikak-Tegoseak Quarry (Fiorillo and Gangloff, 2003; P. Zippi, 2003, personal commun.; Fiorillo et al., 2008)—and from Aniakchak National Monument (P. Zippi, 2004, personal commun.) in southwestern Alaska are becoming available. Additionally, similar data are being compiled for the lower part of the Cantwell Formation (Fig. 2), a Campanian-Maastrichtian fluviatile rock unit, in Denali National Park in central Alaska (Trop, 1996; P. Zippi, 2004, personal commun.). Comparison of pollen data of these three regions shows variability in the fossil flora across Alaska during the Campanian-Maastrichtian (Table 2). Further, the northern and southern localities are more similar to each other than either is to the central Alaska locality, suggesting coastal influences on northern and southern localities and a more continental habitat in the interior. Taphonomic work has shown that there are two dominant species of herbivores, within this ecosystem, the hadrosaur Edmontosaurus sp. (Fiorillo and Gangloff, 1999) and the ceratopsian Pachyrhinosaurus sp. (Fiorillo and Gangloff, 2003), although only an indeterminate hadrosaur has been reported from southern Alaska (Fiorillo and Parrish, 2004). These taxa are best known from remarkably rich bonebeds along the Colville River (Fiorillo and Gangloff, 1999; 2003) where concentrations of bone can exceed 100 bones per square meter. These dense bonebeds demonstrate the gregarious nature of both animals in the ancient high latitudes.
TABLE 2. COMPARISON OF CAMPANIAN–MAASTRICHTIAN POLLEN GENERA BETWEEN THREE AREAS IN ALASKA Aniakchak Denali KikakNational National Tegoseak Monument Park Quarry Kikak-Tegoseak Quarry 62 31 100 Note: The pollen lists are from Trop (1996) and P. Zippi (2004, personal commun.). Comparisons made using Simpson’s Coefficient of Similarity (C/N1 × 100, where C equals the number of shared taxa between two assemblages and N1 equals the total number of taxa in the smaller sample). Aniakchak National Monument is in southwestern Alaska; Denali National Park is in central Alaska; and Kikak-Tegoseak Quarry is on the North Slope of northern Alaska.
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If these keystone herbivores were year-round residents, how did they survive, given that megafloral remains for the Prince Creek Formation suggest a depauperate flora (Spicer, 2003)? The key may be in the fossil pollen record, which, in addition to the variance across the region, suggests an abundance of herbaceous plants (P. Zippi, 2004, personal commun.). In summary, the paleoecological setting for the CampanianMaastrichtian of Alaska was a widespread terrestrial ecosystem capable of supporting large populations of herbivorous dinosaurs. Judging from the dominance of their remains, two of these taxa, a hadrosaur and a ceratopsian, were likely keystone species for the ecosystem and also likely were gregarious. The floral record suggests that plant composition varied across the region, possibly reflecting marine versus continental influences. PALEOGEOGRAPHY OF ALASKAN DINOSAURS Many of the taxa of Late Cretaceous dinosaurs in North America are strikingly related to forms found in Asia, lending support to a northern migration route between northeast Asia and northwest North America (e.g., Russell, 1993; Pasch and May, 1997; Currie, 2000; Norman and Sues, 2000; Sereno, 2000; Godefroit et al., 2003). Various models have been proposed to describe the pattern for the distribution of dinosaur taxa on these two respective landmasses. Rather than go into detail through each taxonomic group, for the purposes of discussion (see below) it will suffice to highlight three different models for three different dinosaur groups. The first model, based on detailed taxonomic analysis of the marginocephalians, requires at least three separate dispersal events to account for the geographic distribution of the related taxa (Sereno, 2000) (Fig. 4). A simpler model for faunal exchange is for the Ornithopoda, which suggests an origin for the group in Asia (Godefroit et al., 2003) (Fig. 5) but barriers to further faunal exchange during the Cretaceous, as evidenced by the increasing taxonomic diversity within the group during this period (Norman and Sues, 2000). Although a similar model of dispersal from Asia seems likely for the theropod group the Troodontidae, Currie (2000) has pointed out that the biogeographic data for the group are ambiguous and that the Troodontidae could have originated on either continent (Fig. 6). From these studies it should be clear that no simple biogeographic model adequately addresses the biogeographic patterns found in the Cretaceous dinosaurs of Asia and North America. Rather, the pattern of faunal exchange is complex, episodic, and in some cases bi-directional. These basic characteristics of faunal exchange have direct parallels in the Quaternary patterns of terrestrial faunal exchange in Beringia. DISCUSSION—A CRETACEOUS BERINGIA? A hypothesized migration route for dinosaurs between Asia and North America has received only vague, ad hoc references by vertebrate paleontologists (e.g., Russell, 1993; Sereno, 2000)
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Figure 4. Faunal exchange model for marginocephalians between Asia and North America in the Cretaceous. (sensu Sereno, 2000). At least three separate dispersal events between Asia and North America, marked by arrows. This model is parallel to that proposed for the exchange of microtine rodents (Repenning, 1967, 1987). Arrows indicate direction of faunal exchange. Base map by the PLATES Project, University of Texas Institute of Geophysics.
without a fuller discussion of the implications of such faunal exchanges. Parallels can be drawn between the attributes of Beringia as a relatively recent biogeographic phenomenon and comparable attributes found in the Late Cretaceous of Alaska, thereby providing a critical assessment of a Cretaceous origin for Beringia based on accretionary tectonics rather than restricting Beringia to a climatic phenomenon. These attributes are shown in Table 3. Understanding the history of Beringia is important because it shapes the ecological understanding of a vast high-latitude geographic region. More specifically, understanding Beringia elucidates the reactions of species and communities to different
climates, as well as the nature of faunal exchanges between two major landmasses. In 1990, Presidents Mikhail Gorbachev and George H.W. Bush endorsed the establishment of an international park spanning the Bering Strait, creating a political component to Beringia that also attempted to define the geographic the extent of region. The objective of the park was to develop a program to bring Russian and American scientists together to study traditional lifeways, biogeography, and landscape history on the Seward and Chukotka Peninsulas. The program would be divided into research components such as ethnoarchaeology, anthropology, historic architecture, geology, paleoecology and wildlife biology (Schaaf, 1993).
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Figure 5. Faunal exchange model for the Ornithopoda between Asia and North America in the Cretaceous (sensu Godefroit, 2003). This model is parallel to that proposed for many Quaternary vertebrate taxa such as bison, cervids, mammoths, sheep, and goats (Repenning, 1967; Flerow, 1967). Arrow indicates direction of faunal exchange. Base map by the PLATES Project, University of Texas Institute of Geophysics.
Certainly the concept of Beringia is beyond these two peninsulas. Literature on Beringia (e.g., Hopkins, 1967; Hopkins et al., 1982; Repenning, 1987; West, 1996) suggests a consensus definition of Beringia based on a combination of factors: geography, climate, and biology. Hopkins (1967) observed that the geography of Beringia was determined by the combination of the internal dynamics of the planet combined with external dynamics such as the atmosphere and climate. He further pointed out that one of the most striking aspects of Beringia was the ability of organisms to find various pathways to different continents. Rather than the political entity outlined by the Russian–United States agreement, a more typical geographic definition of Beringia
is the region that extends east from the Kolyma River in Siberia to the Mackenzie River in Canada (e.g., Hopkins, 1996). The initial definition of Beringia as put forth by Hulten (1937) has somewhat more ambiguous boundaries but encompasses this same general region. Hulten recognized floral similarities between northeastern Asia and northwestern North America and proposed the name Beringia for the presumed landmass that resulted from a drop in sea level due to glacial advance. As such, the concept originated in the modern distribution of plants and their relationship to Pleistocene glaciation. Vertebrate paleontologists have since similarly recognized the importance of a dry land connection between the two landmasses for its role in shaping fossil faunas. The migrations
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Figure 6. Faunal exchange pattern proposed for the Troodontidae between Asia and North America in the Cretaceous (sensu Currie, 2000) in which the origin of the taxon remains unresolved. Arrow indicates directions of faunal exchange. Base map by the PLATES Project, University of Texas Institute of Geophysics.
of terrestrial vertebrates have been shown to be bidirectional but over time increasingly biased toward Asian taxa moving to North America (Repenning, 1967; 1987). Moreover, these migrations were episodic in character (Repenning, 1987). To date, most discussion of Beringia is focused on its role in Plio-Pleistocene history. However, a few authors, most notably Hopkins (1967), have argued that Beringia has a greater antiquity. Hopkins (1996) and O’Neill (2004) have summarized the development of the Beringian concept and its characteristics. With respect to this concept of Beringia, Beringia is composed of a complex system of vegetative zones (e.g., Anderson and Brubaker, 1994) and supported by several gregarious keystone species such as bison, horses, and mammoths. These recent
summaries of Beringia (Hopkins, 1996; O’Neill, 2004) describe how these keystone species survived on a postulated productive ecoregion referred to as the tundra-steppe, arctic-steppe, or mammoth steppe. In apparent conflict with the reconstruction is the depauperate fossil pollen record of the Quaternary, leading some palynologists (Colinvaux, 1980; Cwynar and Ritchie, 1980) to argue instead for an inhospitable polar desert and for faunal migrations occurring only during periods of glacial warming. Guthrie (1982) offered a solution to this productivity paradox (sensu O’Neill, 2004) by suggesting that the volume of herbaceous pollen indicated sufficient vegetation to accommodate the diverse herbivorous vertebrate fauna that moved between the two continents.
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TABLE 3. COMPARISON OF BROAD FEATURES DEFINING BERINGIA AS A QUATERNARY CONCEPT WITH THOSE FEATURES FOUND IN THE CAMPANIAN–MAASTRICHTIAN OF ALASKA Quaternary Campanian–Maastrichtian Land connection between continents Land connection between continents Bidirectional pattern of faunal exchange Bidirectional pattern of faunal exchange Complex system of vegetative zones Pollen variance from northern to interior to southern Alaska Supported gregarious keystone species Bonebeds show gregarious herbivores Bison, horse, mammoth Hadrosaurs, ceratopsians Supported populations year-round Likely supported populations year-round Productivity paradox Productivity paradox Floral data suggest impoverished system Megafloral data suggest taxonomically limited system Abundance of pollen suggests increased herbaceous Pollen data suggest abundant herbaceous vegetation, flora in past, possibly a food source for mega-herbivores possibly a food source for keystone herbivores
Inspection of the Cretaceous terrestrial ecosystem shows a similar paradox. Bonebeds in the Cretaceous rocks of northern Alaska show concentrations of large herbivores existing in a high-latitude environment year-round (Fiorillo and Gangloff, 2001). Further, like the Quaternary pollen record, the megafloral record for the Campanian-Maastrichtian suggests a taxonomically very limited flora (Spicer, 2003). Nevertheless, as Guthrie (1982) suggested for the Quaternary, the apparent abundance of herbaceous pollen in the Alaskan Cretaceous solves a similar paradox during that time. Inherent to these large-scale parameters defining Beringia is the role of climate. Whereas an underlying assumption regarding the Pleistocene has been that Beringia was simply colder than today, there is now recognition of no-analog differences with the present climate (Guthrie, 2001). In his review of Beringian climate, Guthrie (2001) suggested that cloud cover in the Pleistocene was vastly different than observed today, leading to increased aridity for Beringia, which in turn was a limiting determinant for several forms of terrestrial fossil vertebrates. Such a model may account for the distribution of fossil organisms in a wide swath of geography that includes Beringia, but more relevant to this discussion is the acknowledgment by Beringian workers that climate, even within the strictest temporal definition, likely showed significant variability. Thanks to the work of many (e.g., Parrish and Spicer, 1988; Spicer and Parrish, 1990; DeConto et al., 1999; Upchurch et al., 1999), it is well established that the overall climate for the northern high latitudes during the Cretaceous was generally milder than that observed today. Within the context of this cooler modern-day climate, climate change through the Quaternary was the basis for fluctuations in sea level as related to glacial advances and retreats. These fluctuations provided opportunities for organismal exchanges between Asia and North America during the Quaternary. However, Woodburne and Swisher (1995) reexamined the issue of intercontinental overland dispersals. Using high-resolution geochronology they were able to demonstrate that the correlation between sea-level changes and dispersal of land mammals through the Tertiary is not as tightly constrained as previously thought, that there is equal likelihood of dispersal occurring without eustatic change in sea level. They argued
that tectonic factors are equally responsible. They concluded, therefore, that climate change is not a major determinant in land mammal dispersal. Thus the eustatic mechanism driving faunal exchange is not needed for the Quaternary of Beringia, and is not required for the extension of Beringia into the Cretaceous. Certainly the role of tectonics has been invoked previously with respect to influencing faunal exchange. Repenning (1990) pointed out the role of tectonic uplift in altering the migration pattern of the Meadow Mouse (Microtus) into North America, suggesting that the uplift of the Chugach and St. Elias Mountains of southeastern Alaska altered the local climate of the region. This, in turn, changed the path of faunal exchange for Microtus from a coastal route to a more inland route. Jerzykiewicz (1996) suggested that climate change, in the form of the drying of the landscape of central Asia, was what initiated the migration of Asian dinosaurs to North America. The basis for this suggestion stemmed from the observation of the changes in landscape, drainage systems, and climate. In central Asia, block faulting transformed large perennial lakes into semiarid steppes with ephemeral drainage, whereas in western Canada, tectonic collision and related thrust faulting changed the Alberta Basin into an inland basin isolated from the sea. Jerzykiewicz (1996) drew somewhat from Matthew (1915) and Barrell (1916) to follow that climate changes are an important factor in the migration and regional distribution of land vertebrates, and that the evolution of land vertebrates was a response to recurrent periods of aridity. These recurrent periods produced a progressive diversification of the land surface, and evolution requires isolation of habitat and competition from mingling faunas. This proposed climate model would only work, however, if all migrations were unidirectional, and, as others have pointed out (e.g., Sereno, 2000; Currie, 2000), dinosaurian migrations were bidirectional. Evidently, then climate, in the form of meteorological phenomena, may have been significant in Beringia, but it probably did not play a defining role in its origins. The extension of Beringia into the Cretaceous demonstrates that Beringia as driven by climatic history in the Plio-Pleistocene was necessarily preceded by an earlier phase driven by accretionary events in the Cretaceous. A Cretaceous origin for Beringia, in turn, has profound implications for large-scale paleobiological
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patterns and processes for high-latitude ecosystems. The paleontological evidence for Beringia serving as a north-south pathway between Pacific and Arctic marine biotas at particular points in geologic time and the evidence for its providing an east-west conduit between Asiatic and North American terrestrial biotas at other times should eventually complement one another. Marincovich’s (1993) interpretation of the origins of the northern Alaska mollusk fauna in the Prince Creek Formation, for example, is consistent with this review’s extension of the Beringia concept to the Cretaceous. Further alignments between marine paleontological (e.g., Ogasawara, 2002; Dyke et al., 1996; Feder et al., 2003) and terrestrial paleontological evidence can be expected to refine our understanding of how and when Beringia made the transition from an accretionary-process–driven landform in the late Mesozoic to a region or subcontinent largely shaped by Plio-Pleistocene climatic processes. In summary, then, the concept of Beringia, as suggested by Hopkins (1967), was determined by the combination of the internal dynamics of the planet combined with external dynamics such as the atmosphere and climate. Enough data are now known to suggest that the origin of Beringia lies in the Cretaceous and that the defining determinant for Beringia is in the accretionary tectonic history. Given the differences in climate between Cretaceous Alaska and Quaternary Beringia, further study of these parallels through time will provide insight into what is meant by particular ancient ecosystems and perhaps even into the origins of the Arctic.
his enthusiastic discussions regarding Beringia. I also thank R.B. Blodgett, C. Mull, and A. Pasch for their respective constructive reviews of an earlier version of this manuscript. I, however, assume all responsibility for the conclusions drawn here. I also thank C. Repenning for sharing the unpublished manuscript by L. Marincovich with L. Jacobs who in turn brought it to my attention. Thanks as well to L. Lawver and L. Gahagan and the PLATES Project, University of Texas Institute of Geophysics, for providing me the base map for Figures 3–6. Finally, I thank the editors of this volume for their invitation to participate. The National Science Foundation, grant OPP-0424594, and the Jurassic Foundation provided financial support for research on the North Slope. The National Park Service, Alaska Region, provided financial and logistical support for work in Aniakchak National Monument and Denali National Park. I also gratefully acknowledge the support of the Museum of Nature and Science, Dallas, Texas, the University of Alaska Museum of the North, American Airlines, Whole Earth Provision Company, Arco Alaska, Phillips Petroleum, Alyeska Pipeline Service Company, the Barrow Arctic Science Consortium, VECO Polar Resources, and the Arctic Management Unit of the Bureau of Land Management for additional logistical support. This paper is dedicated to Charles A. Repenning, first for his contributions to Beringian science and biochronology of PlioPleistocene mammals and secondly for being the first person to recognize the significance of Alaskan dinosaurs and starting the rush north for more.
CONCLUSIONS
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This review relates accretionary tectonics in the high latitudes of northeast Asia and northwest North America to the emerging understanding of the paleobiology and biogeography of dinosaurs in Alaska. The patterns of biogeography and the paleoecology of Arctic dinosaurs during the Cretaceous are similar to the large-scale patterns of the mosaic floral pattern, the patterns of faunal exchange of later Tertiary and Quaternary mammals, and their paleoecology observed in a more recent geologic definition of Beringia. The consequences of these patterns are that we need to expand the concept of Beringia from being a strictly late Cenozoic phenomenon. Instead, the term Beringia, and its implications, should be formally extended into the Cretaceous. This extension demonstrates that Beringia is rooted in its accretionary history rather than in a climatic history. A Cretaceous origin for Beringia has profound implications for the large-scale paleobiological patterns and processes for high-latitude ecosystems.
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MANUSCRIPT ACCEPTED BY THE SOCIETY 14 DECEMBER 2007
Printed in the USA
Contents Introduction 1. Paleogeographic significance of Ediacaran cyclomedusoids within the Antelope Mountain Quartzite, Yreka terrane, eastern Klamath Mountains, California Nan Lindsley-Griffin, John R. Griffin, and Jack D. Farmer 2. Silurian-bearing terranes ofAlaska Constance M. Soja 3. Silurian
Ga~·tropodafrom
the Alexander terrane,
southeast Alaska David M. Rohr and Robert B. Blodgett 4. Provenance, depositional setting, and tectonic
implications of Silurian polymictic conglomerates in Alaska's Alexander terrane Constance M. Soja and Lena Kmtikov 5. Devonian brachiopods of south westernmost Laurentia: Biogeographic affinities and tectonic significance Arthur J. Boucot, Forrest G . Poole, Ricardo AmayaMartinez, Anita G . Harris, Charles A. Sandberg, and William R. Page
11. Conodont biostratigraphy and facies correlations in a Late Triassic island arc, Keku Strait, southeast Alaska Erik C. Katvala and George D. Stanley Jr. 12. Stratigraphy ofthe Triassic Martin Bridge Formation, Wallowa terrane: Stratigraphy and depositional setting George D . Stanley, Jr. , Christopher A . McRoberts, and Michael T. Whalen 13. Late Triassic (Carnian-Norian) mixed carbonate-
volcaniclastic facies of the Olds Ferry terrane, eastern Oregon and western Idaho Todd A. LaMaskin 14. Early Jurassic bivalves of the Antimonio terrane (Sonora, NW Mexico): Taxonomy, biogeography, and paleogeographic implications Annemarie Scholz, Martin Aberhan, and Carlos M . Gonzalez-Leon 15. Dinosaurs ofAlaska-· Implications for the Cretaceous
origin ofBeringia Anthony R. Fiorillo
6. Devonian brachiopods from northeastern Washington:
Evidence for a non-allochthonous terrane and Late Devonian biogeographic update Peter E. Isaacson 7. Paleobiogeographic affinities ofEmsian (late
Early Devonian) gastropods from Farewell terrane (west-central Alaska) Jiri Fryda and Robert B. Blodgett 8. Significance of detrital zircons in Upper Devonian
ocean-basin strata ofthe Sonora allochthon and Lower Permian synorogenic strata of the Mina Mexico foredeep, central Sonora, Mexico Fonest G. Poole, George E. Gehrels, and John H. Stewart 9. The jlora, fauna, and sediments of the Mount Dall
Conglomerate (Farewell terrane, Alaska, USA) David Sunderlin ~·ilicified shallow-water corals and other marine fon·ils from Wrangellia and the Alexander terrane, Alaska, and Vancouver Island, British Columbia Andrew H. Camthers and George D. Stanley Jr.
10. Late Trim;sic
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