The terrane puzzle: new perspectives on paleontology and stratigraphy from the North American cordillera

Special Paper 442 I .. • THE GEOLOGICAL SOCIETY OF AMERICA® The Terrane Puzzle: New Perspectives on Paleontology and...

Author: Robert B. Blodgett | George D. Stanley

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Special Paper 442

I

.. •

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

2008015342

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: [email protected]; [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)0212.3.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 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)0782.0.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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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)0252.3.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 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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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.

*E-mails: [email protected]; [email protected]; [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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Paleogeographic significance of Ediacaran cyclomedusoids

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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.

7.

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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

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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 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 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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Printed in the USA


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)0782.0.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)0252.3.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

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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.,

*[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 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 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 (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,


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

84

Boucot et al. MIDDLE PRAGIAN PALEOGEOGRAPHY AND LITHOFACIES CRATON

M

E

N

I TR

A

C NS

O

to L it

ra

O A P P

M

I

H

C

S

l

H

I

al

lim

ea t on es

c nd

U

B

P

R

V

e ton

AREA OF ACADIAN OROGENY

rtz os e

one

ost

dol

da

riti

Pe

? RP

e

l

s rite po va

d an

O

Q

NE

e ston dolo

d

MS

Shallow-subtidal limestone and dolostone

I NT

AL NT

sa nd s

C

A

CH

CN

?

Shelf edge

AR

ua

ND

LA

hert

Shelf edge

bti

DV

Strait

sinkholes

Su

D

NE

VA

B

Outer shelf argillite and turbiditic limestone

Flysch

VIN

BPRO

channel remnants and

Pe dolo ritidal ston e

C AS EAN IN

O

Estuarine

SU

AREA OF PROTOANTLER OROGENY

AVA TERRLONIAN ANE

AMERICAN

CE

NORTH

S BA

IN

AN

O CE

MS

M

0

500 KILOMETERS

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


Locality, field no., and USGS collection no.

85

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

86

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)962.0.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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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)882.0.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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northwestern Washington: Geological Society of America Bulletin, v. 67, no. 12, p. 1766. Fatka, O., and Vavrdová, M., 1998, Early Cambrian Acritarcha from sediments underlying the Devonian in Moravia (Měnín Borehole 1, Czech Republic): Czech Geological Survey Bulletin, v. 73, no. 1, p. 55–60. Fortey, R.A., and Mellish, C.J.T., 1992, Are some fossils better than others for inferring palaeogeography: Terra Nova, v. 4, p. 210–216, doi: 10.1111/ j.1365-3121.1992.tb00474.x. Greenman, C., Chatterton, B.D.E., Boucot, A.J., and Berry, W.B.N., 1977, Coarse Silurian(?) and Devonian detrital rocks in northwestern Washington: Evidence of Silurian(?) and Devonian tectonic activity, 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. 467–480. Isaacson, P.E., Measures, E.A., McFaddan, M.D., and Dorobek, S.L., 1989, Coral-stromatoporoid carbonate buildup succession, Jefferson Formation (Late Devonian), central Idaho, U.S.A., in H. Geldsetzer, N.P. James, and G.E. Tebbutt, eds., Reefs: Canada and adjacent areas: Canadian Society of Petroleum Geologists Memoir 13, p. 471–477. Isaacson, P.E., Hladil, J., Jian-wei, S., Kalvoda, J., and Grader, G., 1999, Late Devonian (Famennian) glaciation in South America and marine offlap on other continents, in Feist, R., Talent, J.A., and Daurer, A., eds., North Gondwana: Mid-Paleozoic terranes, stratigraphy and biota: Jahrbuch der Geologischen Bundesantstalt (Wien), v. 54, p. 239–257. Johnson, J.G., and Klapper, G., 1978, Devonian brachiopods and conodonts from central Oregon: Journal of Paleontology, v. 52, no. 2, p. 295–299. Johnson, J.G., and Trojan, W.R., 1982, The Tecnocyrtina brachiopod fauna (?Upper Devonian) of central Nevada: Geologica et Palaeontologica, v. 16, p. 119–150. Kindle, E.M., 1908, The fauna and stratigraphy of the Jefferson Limestone in the northern Rocky Mountain region: Bulletins of American Paleontology, v. 4, no. 20, p. 81–117. Laird, W.M., 1947, An Upper Devonian brachiopod fauna from northwestern Montana: Journal of Paleontology, v. 21, no. 5, p. 453–459. Le Herissé, A., and Gouvernnec, R., 1995, Biogeography of upper Llandovery and Wenlock acritarchs: Review of Palaeobotany and Palynology, v. 86, p. 11–133. Li Jun, 1987, Ordovician acritarchs from the Meitan Formation of Guizhou province, southwest China: Palaeontology, v. 30, p. 613–634. Nautiyal, A.C., 1977, The paleogeographic distribution of Devonian acritarchs and biofacies belts: Journal of the Geological Society of India, v. 18, no. 2, p. 535–561. Nolan, T.B., Merriam, C.W., and Williams, J.S., 1956, The stratigraphic section in the vicinity of Eureka, Nevada: U.S. Geological Survey Professional Paper 276, 77 p. Park, C.F., Jr., and Cannon, R.S., Jr., 1943, Geology and ore deposits of the Metaline quadrangle, Washington. U.S. Geological Survey Professional Paper 202, 81 p. Playford, G., 1993, Miospores and organic-walled microphytoplankton characteristic of strata contiguous with the Devonian-Carboniferous boundary: Comptes Rendus, Seventh International Congress on the Carboniferous– Permian, Buenos Aires, v. 1, p. 127–160. Ross, G.M., Parrish, R.R., and Winston, D., 1992, Provenance and U-Pb geochronology of the Mesoproterozoic Belt Supergroup (northwestern United States): Implications for age of deposition and pre-Panthalassa plate reconstructions: Earth and Planetary Science Letters, v. 113, p. 57–76, doi: 10.1016/0012-821X(92)90211-D. Roux, A., 1991, Ordovician algae and global tectonics, in Riding, R., ed., Calcareous algae and stromatolites: Berlin, Springer-Verlag. Servais, T., Brocke, R., Fatka, O., Le Herissé, A., and Molyneux, S.G., 1997, Value and meaning of the term acritarch: Acta Universitatis Carolinae Geologica, v. 40, p. 631–643. Sorauf, J.E., 1972, Middle Devonian coral faunas (Rugosa) from Washington and Oregon: Journal of Paleontology, v. 46, no. 3, p. 426–439. Stewart, J.H., and Suczek, C.A., 1977, Cambrian and latest Precambrian paleogeography and tectonics in the western United States, 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. 1–18. Tappan, H., 1980, The Paleobiology of plant protists: San Francisco, W.H. Freeman. Tongiorgi, M., Yin, L., and DiMillia, A., 1995, Arenigian acritarchs from the Daping section (Yangtze Gorges area, Hubei Province, Southern China)

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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

Printed in the USA


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 1000

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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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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

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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 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

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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.

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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.

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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).

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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 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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tonic implications: Geology, v. 31, no. 4, p. 335–338, doi: 10.1130/00917613(2003)0312.0.CO;2. Reed, J.C., Jr., 1993, Map of the Precambrian rocks of the conterminous United States and some adjacent parts of Canada (plate 1), in Reed, J.C., Jr., Bickford, M.E., Houston, R.S., Link, P.K., Rankin, D.W., Sims, P.K., and Van Schmus, W.R., Precambrian: Conterminous U.S.: Boulder, Colorado, Geological Society of America, Geology of North America, v. C-2, plate 1. Ross, G.M., and Villeneuve, M., 2003, Provenance of the Mesoproterozoic (1.45 Ga) Belt basin (western North America): Another piece in the preRodinia paleogeographic puzzle: Geological Society of America Bulletin, v. 115, no. 10, p. 1191–1217, doi: 10.1130/B25209.1. Sims, P.K., editor, Anderson, J.L., Bauer, R.L., Chandler, V.W., Hanson, G.N., Kalliokoski, J., Morey, G.B., Mudrey, M.G., Jr., Ojakangas, R.W., Peterman, Z.E., Schulz, K.J., Shirey, S.B., Smith, E.I., Southwick, D.L., Van Schmus, W.R., and Weiblen, P.W., 1993, The Lake Superior region and Trans-Hudson orogen, in Reed, J.C., Jr., Bickford, M.E., Houston, R.S., Link, P.K., Rankin, D.W., Sims, P.K., and Van Schmus, W.R., eds., Precambrian: Conterminous U.S.: Boulder, Colorado, Geological Society of America, Geology of North America, v. C-2, p. 11–120. Smith, D.R., Noblett, J., Wobus, R.A., Reinhard, A., Unruh, D., and Chamberlain, K.R., 1999, A review of the Pikes Peak batholith, Front Range, central Colorado: A “type example” of A-type granitic magmatism, in Boyd, D.W., Lillegraven, J.A., and Snoke, A.W., eds., Proterozoic magmatism of the Rocky Mountains and environs, Part I: Rocky Mountain Geology, v. 34, no. 2, p. 289–312. 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. Stewart, J.H., Amaya-Martínez, R., Stamm, R.G., Wardlaw, B.R., Stanley, G.D., Jr., and Stevens, C.H., 1997, Stratigraphy and regional significance of Mississippian to Jurassic rocks in Sierra Santa Teresa, Sonora, Mexico: Revista Mexicana de Ciencias Geológicas, v. 14, no. 2, p. 115–135. Stewart, J.H., Gehrels, G.E., Barth, A.P., Link, P.K., Christie-Blick, N., and Wrucke, C.T., 2001, Detrital zircon provenance of Mesoproterozoic to Cambrian arenites in the western United States and northwestern Mexico: Geological Society of America Bulletin, v. 113, no. 10, p. 1343–1356, doi: 10.1130/0016-7606(2001)1132.0.CO;2. Stewart, J.H., Poole, F.G., Harris, A.G., Repetski, J.E., Wardlaw, B.R., Mamet, B.L., and Morales-Ramírez, J.M., 1999, Neoproterozoic(?) to Pennsylvanian inner-shelf, miogeoclinal strata in Sierra Agua Verde, Sonora, Mexico: Revista Mexicana de Ciencias Geológicas, v. 16, no. 1, p. 35–62. Timmons, J.M., Karlstrom, K.E., Dehler, C.M., Geissman, J.W., and Heizler, M.T., 2001, Proterozoic multistage (ca. 1.1 and 0.8 Ga) extension recorded in the Grand Canyon Supergroup and establishment of northwest- and north-trending tectonic grains in the southwestern United States: Geological Society of America Bulletin, v. 113, no. 2, p. 163–180, doi: 10.1130/0 016-7606(2001)1132.0.CO;2. Van Schmus, W.R., and Bickford, M.E., co-editors, Anderson, J.L., Bender, E.E., Anderson, R.R., Bauer, P.W., Robertson, J.M., Bowring, S.A., Condie, K.C., Denison, R.E., Gilbert, M.C., Grambling, J.A., Mawer, C.K., Shearer, C.K., Hinze, W.J., Karlstrom, K.E., Kisvarsanyi, E.B., Lidiak, E.G., Reed, J.C., Jr., Sims, P.K., Tweto, O., Silver, L.T., Treves, S.B., Williams, M.L., and Wooden, J.L., 1993, Transcontinental Proterozoic Provinces, in Reed, J.C., Jr., Bickford, M.E., Houston, R.S., Link, P.K., Rankin, D.W., Sims, P.K., and Van Schmus, W.R., eds., Precambrian: Conterminous U.S.: Boulder, Colorado, Geological Society of America, Geology of North America, v. C-2, p. 171–334. Wooden, J.L., and DeWitt, E., 1991, Pb isotopic evidence for the formation of Proterozoic crust in the southwestern United States, in Ernst, W.G., ed., Metamorphism and crustal evolution of the western United States: Englewood Cliffs, New Jersey, Prentice Hall, Rubey Volume VII, p. 69–86. MANUSCRIPT ACCEPTED BY THE SOCIETY 14 DECEMBER 2007

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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.

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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

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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

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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

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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

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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.

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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).

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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.

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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.

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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

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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 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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Caruthers and Stanley

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

159

Age

GH

Keku Strait

Alexander

Hound Island Volcanics

Early Norian

MI 0087

SWK

Keku Strait

Alexander


Early Norian

MI 0074 MI 0056

BS

Keku Strait

Alexander


Early Norian

CPE

Keku Strait

Alexander


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


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.

162


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


163

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


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.


173

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

174


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


Wrangellia Terrane Southern LC PP MI 8302 MI 9947

X

O

O

X


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


X

X

X X


12

X

X


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


O

X

Early Norian


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


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


TS MI 9913

Specimen Count


177

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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


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.

181

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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

183

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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Katvala and Stanley

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


56°58.991'

133°53.455'

Permian?

MI 0042

–

Cannery


56°59.547'

133°52.395'

Permian?

MI 0043

–

Cannery


57°00.678'

133°51.515'

Permian?

MI 0044

–

Cannery


57°01.796'

133°51.384'

Permian?

Keku Volcanics

56°55.069'

134°11.039'

Cretaceous? Late Triassic

MI 0045 MI 0046

–


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

–


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?


Top Cathedral Falls

56°54.103'

133°43.281'

Late Triassic

MI 0066

1932


Portage Pass

56°55.218'

133°52.491'

Late Triassic

MI 0067

1889


Hamilton Island Southeast

56°54.559'

133°51.118'

Late Triassic

MI 0068

–


Floating Unknown

56°52.838'

134°02.406'

Late Triassic

MI 0069

–


Cornwallis Peninsula East—June 22

56°56.215'

134°15.180'

Late Triassic

MI 0070

1906


Cornwallis Peninsula East, M1906

56°56.099'

134°14.534'

Late Triassic

MI 0071

–


Kuiu Island East

56°52.922'

134°01.397'

Late Triassic

MI 0072

–


Kuiu Island East-A

56°52.726'

134°01.336'

Late Triassic

MI 0073

–


K(?) and K(not)

56°52.770'

134°02.203'

Late Triassic

MI 0074

2136


Southwest of Kousk Island

56°53.690'

134°00.508'

Late Triassic

MI 0075

–


Hamilton Island West—site 3

56°55.653'

133°54.135'

Late Triassic

MI 0076

–



56°55.885'

133°54.636'

Late Triassic

MI 0077

–



56°55.920'

133°54.644'

Late Triassic

MI 0078



56°56.260'

133°55.125'

Late Triassic

MI 0079

1886– 1887 1899


Hound Island West, M1899

56°52.798'

133°56.805'

Late Triassic

MI 0080

1923



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



56°52.017'

133°56.075'

Late Triassic

–


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


Hamilton Island Southwest

56°54.696'

133°52.330'

Late Triassic


Hound Island East

56°52.635'

133°56.011'

Late Triassic

MI 0084

–

MI 0085

MI 0087

1890/ 1921 1900– 1901 1912


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


57°02.296'

134°00.652'

Permian

MI 0090

–



57°02.448'

134°00.729'

Late Triassic

MI 0091

–


57°02.599'

134°00.805'

MI 0092

–

Cannery?/ Hamilton Island Limestone Hound Island Volcanics


57°02.893'

134°00.878'

Permian?/Late Triassic Late Triassic

MI 0093

–

Cannery?


57°03.763'

133°59.112'

Permian?

MI 0094

–



57°02.682'

134°01.154'

Late Triassic

MI 0095

–


56°55.040'

134°11.011'

Late Triassic

MI 0096

–


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


HL

–

64

halobiid bivalves

HL


65

halobiid bivalves

HL


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


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


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.

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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


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


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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Katvala and Stanley

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-


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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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


C

95

CPE-C4

primitius?

C

Metapolygnathus primitius?

C

96

CPE-C5

primitius

C


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


C

99

FL-C11

primitius

C


C

99

FL-C12

primitius

C


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,


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

204

Katvala and Stanley

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


205

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,

206

Katvala and Stanley

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


207

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

208

Katvala and Stanley

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,


209

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.


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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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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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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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Stanley et al.

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

231

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 (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

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