Formation and applications of the sedimentary record in arc collision zones

- T HE GEOLOGICAL SOCIETY OF AMERICA® -

Special Paper 436

Formation and Applications of the Sedimentary Record in Arc Collision Zones Edited by

Amy E. Draut U.S. Geological Survey USGS Pacific Science Center 400 Natural Bridges Drive Santa Cruz, California 95060 USA Peter D. Clift School of Geosciences University of Aberdeen Aberdeen, AB24 3UE UK and DFG-Research Centre Ocean Margins Geowissenschaften, Universität Bremen Klagenfurter Strasse 28359 Bremen Germany David W. Scholl Emeritus Senior Scientist U.S. Geological Survey 345 Middlefield Road Menlo Park, California 94025 USA and Stanford University (Emeritus) USA

Special Paper 436 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 Editor: Marion E. Bickford and Donald I. Siegel Library of Congress Cataloging-in-Publication Data Formation and applications of the sedimentary record in arc collision zones / edited by Amy E. Draut, Peter D. Clift, David W. Scholl. p. cm. Includes bibliographical references and index. ISBN 978-0-8137-2436-2 (pbk.) 1. Island arcs. 2. Rocks, Carbonate—Analysis. 3. Convergent margins. 4. Plate tectonics. 5. Geology, Stratigraphic. I. Draut, Amy E. II. Clift, P. D. (Peter D.) III. Scholl, David W. QE511.2.F67 2008 551.1′36—dc22 2007048410 Cover: Oligocene shelf sandstones and shales of the paleo-Chinese passive margin exposed on the northern coast of Taiwan. The rocks are buried and deformed during the ongoing collision of the Luzon Arc with the passive margin of southern China. Towards the north of Taiwan the mountains experience gravitational collapse and extension, culminating in subsidence and formation of the Ilan Plain basin and Okinawa Trough. Photo by Andrew Lin.

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Contents

Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v 1. Preservation of forearc basins during island arc–continent collision: Some insights from the Ordovician of western Ireland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1 Paul D. Ryan 2. Basin formation by volcanic arc loading . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 Dave Waltham, Robert Hall, Helen R. Smyth, and Cynthia J. Ebinger 3. Cenozoic arc processes in Indonesia: Identification of the key influences on the stratigraphic record in active volcanic arcs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27 Robert Hall and Helen R. Smyth 4. Carbonate-platform facies in volcanic-arc settings: Characteristics and controls on deposition and stratigraphic development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 Steven L. Dorobek 5. Sediment waves in the Bismarck Volcanic Arc, Papua New Guinea . . . . . . . . . . . . . . . . . . . . . . . 91 Gary Hoffmann, Eli Silver, Simon Day, Eugene Morgan, Neal Driscoll, and Daniel Orange 6. The Lichi Mélange: A collision mélange formation along early arcward backthrusts during forearc basin closure, Taiwan arc-continent collision . . . . . . . . . . . . . . . . . . . . . . . . . . . 127 Chi-Yue Huang, Chih-Wei Chien, Bochu Yao, and Chung-Pai Chang 7. Oblique subduction in an island arc collision setting: Unique sedimentation, accretion, and deformation processes in the Boso TTT-type triple junction area, NW Pacific . . . . . . . . . . 155 Yujiro Ogawa, Yoshihiro Takami, and Sakiko Takazawa 8. The West Crocker formation of northwest Borneo: A Paleogene accretionary prism . . . . . . . . 171 Joseph J. Lambiase, Tan Yaw Tzong, Amelia G. William, Michael D. Bidgood, Patrice Brenac, and Andrew B. Cullen 9. Temporal changes in the composition of Miocene sandstone related to collision between the Honshu and Izu Arcs, central Japan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 185 Koichi Okuzawa and Ken-ichiro Hisada 10. Cenozoic volcanic arc history of East Java, Indonesia: The stratigraphic record of eruptions on an active continental margin. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 199 Helen R. Smyth, Robert Hall, and Gary J. Nichols

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11. New constraints on the sedimentation and uplift history of the Andaman-Nicobar accretionary prism, South Andaman Island . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 223 R. Allen, A. Carter, Y. Najman, P.C. Bandopadhyay, H.J. Chapman, M.J. Bickle, E. Garzanti, G. Vezzoli, S. Andò, G.L. Foster, and C. Gerring 12. Post-collisional collapse in the wake of migrating arc-continent collision in the Ilan Basin, Taiwan. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 257 Peter D. Clift, Andrew T.S. Lin, Andrew Carter, Francis Wu, Amy E. Draut, T.-H. Lai, L.-Y. Fei, Hans Schouten, and Louis Teng 13. The Guerrero Composite Terrane of western Mexico: Collision and subsequent rifting in a supra-subduction zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 279 E. Centeno-García, M. Guerrero-Suastegui, and O. Talavera-Mendoza 14. Tectonic architecture of an arc-arc collision zone, Newfoundland Appalachians . . . . . . . . . . . 309 Alexandre Zagorevski, Cees R. van Staal, Vicki McNicoll, Neil Rogers, and Pablo Valverde-Vaquero 15. The Catalina Schist: Evidence for middle Cretaceous subduction erosion of southwestern North America . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 335 M. Grove, G.E. Bebout, C.E. Jacobson, A.P. Barth, D.L. Kimbrough, R.L. King, Haibo Zou, O.M. Lovera, B.J. Mahoney, and G.E. Gehrels 16. Sedimentary response to arc-continent collision, Permian, southern Mongolia . . . . . . . . . . . . 363 C.L. Johnson, J.A. Amory, D. Zinniker, M.A. Lamb, S.A. Graham, M. Affolter, and G. Badarch 17. Links among mountain building, surface erosion, and growth of an accretionary prism in a subduction zone—An example from southwest Japan . . . . . . . . . . . . . . . . . . . . . . . . 391 Gaku Kimura, Yujin Kitamura, Asuka Yamaguchi, and Hugues Raimbourg

Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 405

Preface The Sedimentary Record in Arc Collision Zones Sediment deposited in volcanic arc settings yields the best available record of active-margin evolution spanning long periods of geologic time. The tectonic, erosional, and magmatic geochemical histories preserved in arc sedimentary records, and particularly in arc collision zones, hold the key to understanding multiple, inter-connected geologic processes: the formation and destruction of continental crust, changes in plate configuration and rates of plate motion, subduction-zone deformation and associated seismogenesis, and, as is becoming increasingly apparent, links between tectonically driven rock uplift and climate. Arc sedimentary and volcanic records from many locations have been used to decipher tectonic evolution of modern and ancient convergent margins (e.g., Dickinson et al., 1982; Scholl et al., 1983; Tatsumi et al., 1983; Ingersoll, 1983; Harbert et al., 1986; Dewey and Ryan, 1990; Gill et al., 1994; Fackler-Adams et al., 1997; Trop et al., 2003; Busby et al., 2005; Clift et al., 2005; Huang et al., 2006), and geochemical studies of arc sediment and volcaniclastic deposits have done much to illuminate possible origins of continental crust (e.g., Pearcy et al., 1990; Miller and Christensen, 1994; Draut et al., 2002). Though much has been learned from these records and their applications thus far, many important questions remain regarding controls on formation of the sedimentary record in arc environments, where tectonically generated topography is typically complex. A greater understanding of how and where accommodation space is created, and of depositional and reworking processes in arc settings, is essential for determining the ability of arc sediment to preserve accurately the tectonic and magmatic signals for which they are commonly studied (e.g., Marsaglia and Ingersoll, 1992; Busby and Ingersoll, 1995; Draut and Clift, 2006). Better comprehension of how arc sediment forms and evolves will facilitate new and more advanced applications of arc stratigraphy as well as, perhaps, inhibiting inferences for which the sedimentary record provides insufficient or inconclusive evidence. This compilation of papers resulted indirectly from a Geological Society of America Penrose Conference held in October 2005, co-sponsored by the British Sedimentological Research Group (Geological Society of London) and the International Association of Sedimentologists. Approximately half of the papers in this volume were authored by participants in that meeting, which focused on lessons in tectonics, climate, and eustasy gained from the stratigraphic record preserved in arc collision zones. The first set of papers in this collection focuses on formation and evolution of the sedimentary record in arc settings and arc collision zones, concentrating on modern intra-oceanic examples. Two modeling papers examine basin formation around intra-oceanic arcs: Ryan presents a model indicating that forearc-basin topography can evolve as a function of eclogite formation, which causes basin subsidence (forming accommodation space) and may explain the problematic preservation of forearc basins during arc-continent collision and orogeny. Waltham et al. model flexural subsidence due to loading of arc and forearc lithosphere by a volcanic arc, and find that the predicted loads are sufficient to account for much or all of basin subsidence within 100 km of active arc volcanoes, a process suggested to be particularly common in older, mature arcs. A related study of Indonesian arc stratigraphy by Hall and Smyth identifies key influences on development of the sedimentary record in active volcanic arcs, including complexities of basin stratigraphy such as extension, subduction polarity reversal, collision events, and fluctuating accommodation space associated with variable loading of the lithosphere by the arc itself. Dorobek presents a comprehensive analysis of evolving carbonate sedimentation processes around volcanic arcs through time. He shows that although voluminous volcanic activity is harmful to reef growth, most of the volcaniclastic output of arc volcanoes is trapped in basins. This allows carbonate reefs to act as sensitive recorders of vertical tectonic motions in forearc regions. Hoffmann et al., using new high-resolution seismic images from Papua New Guinea, examine the reworking of volcaniclastic sediment around volcanic arcs; their study demonstrates that variations in sediment thickness caused by the formation of current-driven bedforms warrant caution when comparing bed thicknesses across sedimentary sections sampled in cores. Huang et al. explore mélange development in the active arc-continent collision zone of Taiwan. Ogawa et al. review the tectonic development of the unique trench-trench-trench triple junction between the Honshu and Izu Arcs, now located south of Tokyo. They also present a physical analogue model showing formation of synthetic and antithetic faulting and topography within this oblique arc collision zone and evaluate stress partitioning between the basement and overlying sediment cover. v

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The second half of the volume presents new applications of arc sedimentary records. Lambiase et al. use new sedimentology, micropaleontology, and structural analysis of turbidite sequences in Borneo to interpret them as an accretionary prism and examine in detail the syn-depositional deformation structures that occur within such a system. Okuzawa and Hisada have reconstructed the tectonic evolution of collision between the Izu and Honshu Arcs, Japan, using provenance of pre- and syn-collisional Miocene clastic deposits. They focus on heavy-mineral studies of clinopyroxene, garnet, and spinel to show a switch in provenance in the trench fill during the middle Miocene driven by initial uplift and erosion of the Honshu Arc after collision with the Izu arc massif. Smyth et al. trace the complete eruptive history of a well preserved Indonesian Eocene–Miocene volcanic arc from initiation to termination, and present evidence for contamination of volcanism by Archean–Cambrian continental crust underneath the arc. These authors also note that the volume of acidic arc volcanism has previously been underestimated because weathered acidic volcaniclastic material was misinterpreted as terrigenous erosion products. Allen et al. use sediment provenance, 40Ar-39Ar mica dating, and fission-track data to constrain the uplift history and sources of sediment entering the accretionary prism of the Sunda subduction zone as recorded in the Andaman Islands. This study demonstrated that the rising Himalaya did not contribute greatly to this accretionary wedge, but that erosion switched from being local, arc-derived to continental Myanmar around 40 Ma. Clift et al. use a variety of geological and geophysical data sets to reconstruct exhumation and subsidence rates along the Ilan Basin, Taiwan (western limit of the Okinawa Trough), to show that subduction polarity reversal associated with a migrating, steady-state oblique collision between the Luzon arc and the passive margin of southern China causes rapid formation of deep marginal basins along reactivated detachment faults that are filled by detritus from the orogen. In a comprehensive summary paper, Centeno-Garcia et al. demonstrate the utility of arc lithofacies and isotopic signatures to reconstruct the highly complex tectonic evolution of western Mexico, including multiple stages of terrane collision and rifting. Zagorevski et al. present new U-Pb ages constraining timing of volcanism, sedimentation, and the age of Ordovician arc-arc collision during the closure of the Iapetus Ocean and formation of the Appalachian-Caledonide orogen, developing a tectonic model analogous to modern arcarc collisions in the southwestern Pacific. Grove et al. use new detrital zircon U-Pb age data to propose a subduction-erosion origin for portions of the Catalina Schist in the western United States, in contrast to the long-held theory that the metamorphic complex formed entirely during nascent subduction. The final two papers provide new insights into possible tectonic-climate coupling in arc-collisional settings: Johnson et al. use provenance, paleoclimatic, and stratigraphic data to construct a regional synthesis of Permian arccontinent collision in Eurasia, encompassing the timing of tectonic events and implications for climatic changes associated with orographic effects of regional uplift. Kimura et al. investigate intriguing feedback mechanisms among orogenesis, surface erosion via the Asian monsoon, growth of the Nankai Trough accretionary prism by incorporation of sediment eroded from the orogen, and generation of large earthquakes that promote additional sediment erosion. It is expected that the utility and applications of arc sedimentary records will expand considerably with upcoming research initiatives such as those of the Integrated Ocean Drilling Program (IODP) that will address processes in modern arc settings, including the Nankai Trough, Izu-Bonin forearc, and the collision zone between the Yakutat terrane and North America, among others. New advances in analytical techniques will improve age controls on tectonic evolution of active margins and, it is hoped, offer greater insight into the links and feedback processes among collision-related uplift, precipitation, sediment flux to subduction zones, and seismic activity with its associated societal hazards. Amy E. Draut Peter D. Clift David W. Scholl

Preface

REFERENCES CITED Busby, C.J., and Ingersoll, R.V., eds., 1995, Tectonics of sedimentary basins: Blackwell Science, 592 p. Busby, C.J., Bassett, K., Steiner, M.B., and Riggs, N.R., 2005, Climatic and tectonic controls on Jurassic intra-arc basins related to northward drift of North America, 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. 359–376. Clift, P.D., Pavlis, T., Debari, S.M., Draut, A.E., Rioux, M., and Kelemen, P.B., 2005, Subduction erosion of the Jurassic Talkeetna-Bonanza arc and the Mesozoic accretionary tectonics of western North America: Geology, v. 33, p. 881–884. Dewey, J.F., and Ryan, P.D., 1990, The Ordovician evolution of the South Mayo trough, western Ireland: Tectonics, v. 9, p. 887–901. Dickinson, W.R., Ingersoll, R.V., Cowan, D.S., Helmond, K.P., and Suczek, C.A., 1982, Provenance of Franciscan graywackes in coastal California: Geological Society of America Bulletin, v. 93, p. 95–107, doi: 10.1130/00167606(1982)93<95:POFGIC>2.0.CO;2. Draut, A.E., Clift, P.D., Hannigan, R.E., Layne, G., and Shimizu, N., 2002, A model for continental crust genesis by arc accretion: Rare earth element evidence from the Irish Caledonides: Earth and Planetary Science Letters, v. 203, p. 861–877, doi: 10.1016/S0012-821X(02)00931-7. Draut, A.E., and Clift, P.D., 2006, Sedimentary processes in modern and ancient oceanic arc settings: evidence from the Jurassic Talkeetna Formation of Alaska and the Mariana and Tonga arcs, western Pacific: Journal of Sedimentary Research, v. 76, p. 493–514, doi: 10.2110/jsr.2006.044 Fackler-Adams, B.N., Busby, C.J., and Mattinson, J.M., 1997, Jurassic magmatism and sedimentation in the Palen Mountains, southeastern California: Implications for regional tectonic controls on the Mesozoic magmatic arc: Geological Society of America Bulletin, v. 109, p. 1464–1484, doi: 10.1130/0016-7606(1997)109<1464:JMASIT>2.3.CO;2. Gill, J.B., Hiscott, R.N., and Vidal, P., 1994, Turbidite geochemistry and evolution of the Izu-Bonin Arc and continents: Lithos, v. 33, p. 135–168, doi: 10.1016/00244937(94)90058-2. Harbert, W., Scholl, D.W., Vallier, T.L., Stevenson, A.J., and Mann, D.M., 1986, Major evolutionary phases of a forearc

basin of the Aleutian terrace: Relation to North Pacific tectonic events and the formation of the Aleutian subduction complex: Geology, v. 14, p. 757–761, doi: 10.1130/00917613(1986)14<757:MEPOAF>2.0.CO;2. Huang, C.-Y., Yuan, P.B., and Tsao, S.J., 2006, Temporal and spatial records of active arc-continent collision in Taiwan: A synthesis: Geological Society of America Bulletin, v. 118, p. 274–288, doi: 10.1130/B25527.1. Ingersoll, R.V., 1983, Petrofacies and provenance of Late Mesozoic forearc basin, northern and central California: The American Association of Petroleum Geologists Bulletin, v. 67, p. 1125–1142. Marsaglia, K.M., and Ingersoll, R.V., 1992, Compositional trends in arc-related, deep-marine sand and sandstone: A reassessment of magmatic-arc provenance: Geological Society of America Bulletin, v. 104, p. 1637–1649, doi: 10.1130/0016-7606(1992)104<1637:CTIARD>2.3.CO;2. Miller, D.J., and Christensen, N.I., 1994, Seismic signature and geochemistry of an island arc: A multidisciplinary study of the Kohistan accreted terrane, northern Pakistan: Journal of Geophysical Research, v. 99, p. 11623–11642, doi: 10.1029/94JB00059. Pearcy, L.G., Debari, S.M., and Sleep, N.H., 1990, Mass balance calculations for two sections of island arc crust and implications for the formation of continents: Earth and Planetary Science Letters, v. 96, p. 427–442, doi: 10.1016/0012821X(90)90018-S. Scholl, D.W., Vallier, T.L., and Stevenson, A.J., 1983, Arc, forearc, and trench sedimentation and tectonics; Amlia corridor of the Aleutian Ridge, in Watkins, J.S., and Drake, C.L., eds., Studies in continental margin geology: American Association of Petroleum Geologists Memoir 34, p. 413–439. Tatsumi, Y., Sakuyama, M., Fukuyama, H., and Kushiro, I., 1983, Generation of arc basalt magmas and thermal structure of the mantle wedge in subduction zones: Journal of Geophysical Research, v. 88, p. 5815–5825. Trop, J.M., Ridgway, K.D., and Spell, T.L., 2003, Sedimentary record of transpressional tectonics and ridge subduction in the Tertiary Matanuska Valley–Talkeetna Mountains forearc basin, southern Alaska, in Sisson, V.B., Roeske, S.M., and Pavlis, T.L., eds., Geology of a transpressional orogen during ridge-trench interaction along the North Pacific margin: Geological Society of America Special Paper 371, p. 89–118.

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The Geological Society of America Special Paper 436 2008

Preservation of forearc basins during island arc–continent collision: Some insights from the Ordovician of western Ireland Paul D. Ryan* Department of Earth and Ocean Sciences, National University of Ireland, Galway, Ireland

ABSTRACT A new model is proposed for the problematic preservation of an Ordovician forearc basin, which records a complete sedimentary record of arc-continent collision during the Grampian (Taconic) orogeny in the west of Ireland. The South Mayo Trough represents an arc and forearc complex developed above a subduction zone in which the slab dipped away from the Laurentian passive margin. The collision of this arc with Laurentia caused the Middle Ordovician Grampian orogeny. However, the South Mayo Trough, in the hanging wall of this collision zone, remained a site of marine sedimentation during the entire process. Early sediments show derivation from an island-arc complex, an ophiolitic backstop, and polymetamorphic trench sediments. These are conformably overlain by marine deposits derived from a more evolved arc complex and an emerging juvenile orogen. This transition is dated as being coeval with the Grampian metamorphism of the Laurentian footwall. The problem remains as to why subsidence continued in a basin on the hanging wall. It is proposed that the suppression of the expected topography is due to the nature of the Laurentian continental margin. Geophysical and geological evidence suggests that this was a volcanic margin during Neoproterozoic rifting. It is argued that the subduction of this margin caused the formation of eclogites, which reduced its buoyancy. Simple numerical models are presented which show that this is a viable mechanism for the suppression of topography during early stages of arc-continent collision and hence for the preservation of forearcs. Keywords: forearc, preservation, Grampian Orogeny, arc-continent collision, eclogite.

INTRODUCTION

(Ryan and Dewey, 1991); the early Eocene of Kamchatka (Konstantinovskaia, 2000, 2001); the Neoproterozoic of the TuvaMongolia Massif (Kuzmichev et al., 2001) and of the Anti-Atlas of Morocco (Thomas et al., 2002). This process can result in slab break-off and a reversal of subduction polarity (McKenzie, 1969) such as in South Mayo (Ryan and Dewey, 1991) or Taiwan (Teng, 1996; Clift et al., 2003) and the creation of an active continental

Island arc–continent collision is an important mechanism for continental growth. However, proposed models for this process are varied and complex. A forearc can collide with a passive continental margin (Fig. 1A), initiating orogeny at that margin, as suggested for the Middle Ordovician of South Mayo, Ireland *[email protected]

Ryan, P.D., 2008, Preservation of forearc basins during island arc–continent collision: Some insights from the Ordovician of western Ireland, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 1–9, doi: 10.1130/2008.2436(01). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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margin (Fig. 1B). Alternatively, a new subduction zone can initiate elsewhere in the plate to the rear of the arc (Fig. 1C), as has been proposed for the late Archaean Wutaishan orogeny (Wang et al., 1996). The forearc can be either thrust in its entirety over the footwall, the preferred model for South Mayo (Ryan and Dewey, 1991; Dewey and Mange, 1999; Clift et al., 2004), or become part of an imbricated stack driving subsidence in a foreland basin developed upon the footwall as in Taiwan (van der Werff, 1995). Alternatively, modeling shows that if the overriding oceanic plate fails in the backarc or arc region, the arc and forearc can be wholly or partly subducted (Fig. 1D) if they are not sufficiently buoyant (Konstantinovskaia, 2000; Boutelier et al., 2003). This leads to an early orogenic event, with little of the arc complex preserved and perhaps ultrahigh pressure metamorphism in the continental margin. A forearc can also collide with an active margin (Fig. 1E), a model used for the collision of the Kohistan arc with Asia during the Late Cretaceous (Khan et al., 1997). This should produce a climactic orogenic event following a phase of crustal growth by addition of subduction-related magmatism. However, an alternate model is that the Kohistan backarc collided with an active Asian margin (Rolland et al., 2000; O’Brien, 2001; Robertson and Collins, 2002). In all such models the preservation potential of the forearc is low. If obducted, it lies at high structural levels in the hanging

Figure 1. Sketches showing various proposed mechanisms affecting the forearc region during island arc–continent collision: (A) Initial precollision configuration. (B) Obduction of the arc and forearc, followed by subduction polarity reversal. (C) Obduction of the arc and forearc, followed by initiation of a new subduction zone elsewhere in the system, which may have either polarity. (D) Weakness in the arc or backarc leads to the partial or complete subduction of the forearc and arc without reversal of sense of subduction with accompanying ultrahigh pressure metamorphism of the continental margin. (E) Two opposing magmatic arcs, or an island arc and a magmatic arc collide. (F) Neither forearc is obducted, and a transform boundary is created.

wall of an orogeny and is likely to be removed rapidly by erosion, especially if slab break-off occurs. Alternatively, it may be subducted or dismembered by strike-slip faulting. This contribution discusses the geology of western Ireland, where an Ordovician island arc is believed to have collided with the Laurentian passive margin (Ryan and Dewey, 1991), causing the Grampian orogeny (475–462 Ma) prior to the final closure of the Iapetus Ocean in Late Silurian times (420–416 Ma). Remarkably, a significant proportion of the forearc, including a possible forearc basin (Dewey and Ryan, 1990), the South Mayo Trough, is preserved in spite of the complex having been in the hanging wall of the earliest orogenic phase. It will be argued that the geology of the footwall to this system, the Laurentian continental margin, was critical in the preservation of the South Mayo forearc. ORDOVICIAN GEOLOGY OF WESTERN IRELAND The main geological divisions of western Ireland are shown in Figure 2. The rocks to the north of the Fair Head–Clew Bay Line have a Grenvillian basement (Daly, 1996, 2001; Daly and Flowerdew, 2005), overlain by Neoproterozoic sediments of the Dalradian Supergroup, which were deposited from pre–720 Ma to post–595 Ma (Condon and Prave, 2000). Mafic dikes cut the basement, postdate all Grenvillian fabrics, and are isotopically

Insights from the Ordovician of western Ireland

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Figure 2. (A) Geology of western and northwestern Ireland. (B) Geology of South Mayo. Fm.—Formation; Gp.—Group; Cplx.—complex; FCL—Fair Head–Clew Bay Line; SUF—Southern Upland Fault.

and geochemically matched with pre-deformational dikes in the overlying Dalradian (Daly and Flowerdew, 2005). Both basement and cover were affected by the Grampian orogeny between 475 Ma and 462 Ma (Daly and Flowerdew, 2005; Flowerdew et al., 2005). In the extreme south the Dalradian Supergroup contains mafic volcanics with crossite (Gray and Yardley, 1979). This region is commonly interpreted as making up part of the Laurentian passive margin to the Iapetus Ocean, which opened in latest Neoproterozoic times, the Dalradian sediments having been deposited during rifting and breakup of Rodinia (Strachan and Holdsworth, 2000). The Clew Bay Complex lies to the south of the Fair Head– Clew Bay Line and is associated with a strong magnetic lineament (Max and Riddihough, 1975; Max et al., 1983) and is regarded

as the westward continuation of the Highland Boundary Fault in Scotland. The complex is highly sheared and comprises two main components: a sedimentary mélange underlying most of Clew Bay (Dewey and Ryan, 1990) and a dismembered ophiolite in the south (Ryan et al., 1983). Heavy mineral (Dewey and Mange, 1999) and isotopic studies (Chew, 2003) support a mature continental provenance for the sediments. The Clew Bay Complex has been variously interpreted as part of the Dalradian (Phillips, 1973), a lateral equivalent to the Highland Border Complex of Scotland (Harper et al., 1989), a shear carpet formed beneath advancing ophiolitic nappes (Dewey and Shackleton, 1984), or an accretionary complex above a south-verging subduction zone (Dewey and Ryan, 1990). The age of the complex is uncertain, although a Cambrian–Ordovician age is likely on provenance

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grounds (Dewey and Mange, 1999), and it shares the same structural history as the Dalradian to the north with the S2 fabric in both dated at ca. 460 Ma (Chew et al., 2003). Provenance studies also suggest a stratigraphic linkage between this complex and the South Mayo Trough by Early to Middle Ordovician times (Dewey and Ryan, 1990). The South Mayo Trough (Fig. 2B) contains a thick sequence of Lower to Middle Ordovician sediments that are preserved in a large syncline, the Mweelrea-Partry Syncline. The northern limb contains an Arenig-Llanvirn sequence that exceeds 9 km in thickness in which early sediments show derivation from a primitive island arc complex, an ophiolitic backstop, and polymetamorphic trench sediments. These are conformably overlain by marine clastic sequences derived from a more evolved arc complex and an emerging juvenile (Grampian) orogen (Dewey and Mange, 1999). Associated volcanic complexes, which lie mainly in the southern limb, are of primitive arc affinities near the base (Ryan et al., 1980; Clift and Ryan, 1994; Draut et al., 2004) and show an increasing proportion of Laurentian-derived melt upsection (Draut et al., 2004). Dewey and Ryan (1990) interpreted the South Mayo Trough as a forearc basin on the grounds that it lay between an arc (volcanic complexes) and an accretionary prism (the Clew Bay Complex) which in turn had been thrust over a passive continental margin, Laurentia, whose leading edge had been subducted. It is remarkable that the South Mayo Trough, interpreted as lying in the obducted hanging wall, apparently provides a continuous record of sedimentation during Grampian deformation of the Laurentian margin, interpreted as lying in the footwall. The Dalradian rocks of Connemara (Fig. 1) are in fault contact with the South Mayo Trough (Dewey et al., 1970; Ryan and Archer, 1978) and were probably emplaced outboard of the oceanic terrane of South Mayo by Late Ordovician lateral motion (Hutton, 1987). The southern boundary is a SE-verging thrust, the Mannin Thrust (Leake et al., 1983), which emplaced migmatitic metasediments and intrusives over evolved arc rocks dated at 474.6 ± 5.5 Ma (Friedrich, 1998; Draut and Clift, 2002) that show chemical similarities to volcanics within the South Mayo Trough. This suggests that Connemara lay structurally beneath an evolving arc by Llanvirn times (Draut and Clift, 2002). U/Pb studies of syntectonic gabbros and later granitic melts date D2 of the Grampian orogeny at 473 Ma, and post-D3 rapid uplift that took place before 462 Ma (Friedrich et al., 1999a, b). The Grampian (Taconic) orogeny was a Middle Ordovician event that affected the Laurentian margin prior to final closure of the Iapetus Ocean, resulting from the collision of an arc with the Laurentian margin, followed by a subduction polarity reversal (Dewey and Shackleton, 1984; Ryan and Dewey, 1991). Island arc–continent collision began in Arenig times (Draut and Clift, 2001), with hard collision being reflected in an increase of continental component in the arc melt (Draut and Clift, 2001; Draut et al., 2004). Provenance studies in the South Mayo Trough suggest initiation in an oceanic forearc, with later sediments recording the stripping of an evolved arc and a juvenile orogen (Dewey and Mange, 1999; Clift et al., 2004). This change in provenance was

coeval with the emplacement of the D2 gabbros in Connemara at 473 Ma (Dewey and Mange, 1999). Radiometric evidence from both North Mayo and Connemara suggests that the entire event took <20 m.y., with initial collision taking ~10 m.y. and subduction flip a further 5 m.y. (Dewey and Mange, 1999; Soper et al., 1999; Draut and Clift, 2001). Evidence for the reversal of subduction polarity beneath the Laurentian margin after the Grampian orogeny is found in Middle to Upper Ordovician and Silurian strata of Connemara and South Mayo. The South Connemara Complex, to the south of the Southern Upland Fault (SUF, Fig. 2), was probably formed in a Late Ordovician trench of opposite polarity adjacent to the Laurentian margin (Ryan and Dewey, 2004). Silurian deposits in north Connemara and South Mayo (Fig. 2) contain abundant arc detritus (Menuge et al., 1995; Dewey and Mange, 1999) and probably formed in a supra-subduction zone setting (Williams and Harper, 1991; Menuge et al., 1995). These deposits are mostly shallow marine and of 2.5–3 km in thickness, suggesting only modest stretching of a normally buoyant lithosphere without topography (equivalent to a β factor of ~1.25). Critically, where the Silurian sedimentary and volcanic sequences rest unconformably upon the Ordovician rocks, the earliest slaty cleavage, if present, cuts both sequences. Structural and stratigraphic evidence, therefore, suggests that prior to Silurian deposition the Ordovician sequence of South Mayo was only modestly deformed, of anchimetamorphic grade, and near sea level (Dewey and Ryan, 1990). If the arc-continent collision model is accepted for South Mayo, a mechanism must, therefore, have operated to suppress the topography of the forearc during and after this event. ISOSTATIC CONSEQUENCES OF ARC-CONTINENT COLLISION The likely topographic response of an arc-continent collision is investigated in two ways. Firstly, a flexural isostatic model is used to estimate the uplift of the hanging wall for a 10 m.y. period from 485 to 475 Ma (Dewey and Mange, 1999; Draut and Clift, 2002), when arc-continent collision was thought to have taken place. Secondly, a simple model, assuming Airy isostasy, is used to estimate the likely uplift during slab break-off, which is assumed to have led to the emplacement of mafic plutons and quartz-diorite gneisses in Connemara between 475 and 470 Ma (Friedrich et al., 1999a, b). The aim is to investigate the conditions that are required to keep the forearc basin beneath sea level during both of these events. The flexural response of the lithosphere during collision, but prior to slab break-off, is estimated using a two-dimensional (2D) kinematic flexural model. The present across-strike width of the rocks affected by the Grampian orogeny (the Grampian and NorthWestern terranes of Murphy et al., 1991) is ~100–140 km. Rocks to the north in Donegal were metamorphosed to greenschist facies (Pitcher and Berger, 1972), suggesting that the Grampian orogen extended farther north. Strike swing and post-Grampian collapse (Alsop, 1991) make it difficult to estimate the original width of

Insights from the Ordovician of western Ireland the orogen, but allowing for such factors, Dewey and Mange (1999) suggest 200 km as a reasonable value. It is assumed that an arc complex, including South Mayo, was emplaced over the Laurentian margin along a thrust that nucleated upon the subduction zone. A hanging wall of 210 km was moved over a flat 28 km in depth (ca. 0.8 GPa) for 100 km and then up a ramp at an angle of 14.2° for a further 110 km (Fig. 3). A specific gravity of 2200 kg m–3 was assumed for sediments in the flexural basin. A value of 2840 kg m–3, the mean density of diorite, was used for both the arc material of the hanging wall and a continental basement containing up to 30% mafic material for the footwall. The thrust was moved at 20 mm.yr –1 so that the deformation front moved 200 km in 10 m.y. The actual rate of overthrusting may have been greater (Dewey and Mange, 1999). The hanging wall was moved along flow lines parallel to the ramp–flat surface. Erosion was modeled using equation 5 below, and the model was compensated using flexural isostasy (Watts, 2001) after each increment of movement. Different elastic thicknesses

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(Te) of 2.0 km (Fig. 3A) and 20.0 km (Fig. 3C) were used to estimate the response of flexurally weak and flexurally strong systems, respectively. The exact position of the South Mayo forearc basin within such a system is unknown but is taken as being between 100 and 200 km south of the thrust tip. The model with low Te (Fig. 3A) predicts some 3 km of uplift in the forearc region accompanied by ~1.5 km of erosion and a flexural basin on the Laurentian margin with 0.5 km of sediment (Fig. 3B). The model with high Te (Fig. 3C) predicts some 6 km of uplift in the forearc region accompanied by 3–4 km of erosion and a flexural basin on the Laurentian margin with 3.4 km of sediment (Fig. 3D). No Ordovician flexural basin is recorded in Ireland, nor is there any evidence for one farther north offshore on the WIRELINES seismic reflection section (Klemperer et al., 1991). Thus it is concluded that the overthrusting of the arc hanging wall over the continental margin footwall was likely to have taken place in a low Te regime. There still remains the problem that if the assumptions made in this simple mechanical model

A B

C D

Figure 3. (A) Flexural isostatic model in which a hanging wall of 28 km is emplaced at 20 mm.yr –1 over a continental footwall. The heavy upper line represents the erosion surface after 10 m.y. at the end of the model run. The dashed line represents the thrust surface. The erosion constant is 5 m.y.–1, and the elastic thickness (Te) is 2 km. (B) Subsidence-time curve for the foreland basin produced at the toe of the thrust system in (A). (C) A similar model to that in (A) but with Te = 20 km. (D) Subsidence-time curve for the foreland basin produced at the toe of the thrust system in (C).

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are reasonable (see caption for Fig. 3), the South Mayo Trough was likely to have been above sea level prior to slab break-off at ca. 475 Ma. A simple technique is employed to estimate the possible topography associated with subsequent slab break-off. Figure 4A shows a typical tectonic cartoon of the process of island arc–continent collision at the point of subduction reversal (based upon Dewey and Mange, 1999). Airy compensation is assumed in the light of the previous flexural analysis. Topography (h) with respect to the summit of the oceanic ridges is given by (see Dewey et al., 1993):

h=

( ρ a − ρ l ) ( lz − h ) , ρc

(1)

where ρc, ρa, ρl are the specific gravities of the crust, asthenosphere, and lithosphere, respectively, and lz is the thickness of the lithosphere. Where ρl may be estimated from:

ρl =

[(Cz − h) ρc + (lz − Cz ) ρm ], [lz − h] •

•

(2)

where Cz is the thickness of the crust and ρm the specific gravity of the mantle. The following values were assumed for the specific gravities: ρc = 2840 kg.m–3; ρm = 3330 kg.m–3; and ρa = 3281 kg. m–3. The value for ρa places the upper surface of the lithosphere with Cz = 30 km and lz = 125 km at sea level. If topography above sea level (d) is corrected for the mean depth of the mid-ocean ridges we obtain:

d + 2.64 = h .

(3)

When (2) and (3) are substituted in (1) and simplified, we get:

d=

[Cz (ρm − ρc )+ lz (ρa − ρm )] − 2.64. ρa

(4)

Equation 4 can be used to estimate the topography of any geodynamic cartoon if the thicknesses of crust and mantle are measured at regular intervals across the model (Fig. 4). Although this may differ from dynamic topography, it provides a simple

A

B

Figure 4. (A) Isostatic consequences of an arc-continent collision and subduction polarity flip model taken from Dewey and Mange (1999). The heavy solid line in the upper topographic grid shows the maximum topography produced by this model, assuming Airy compensation; the dashed line shows the likely topography after 10 Ma erosion at an erosion constant of 5 m.y.–1. (B) Isostatic consequences of both the model in Figure 4A and the thrust models from Figure 3A, with the geology of South Mayo speculatively positioned onto the hanging wall, assuming lower crustal-phase changes. The heavy lines in the upper topographic grid show the likely maximum topography for the model in Figure 3A, assuming the lower crust below 40 km is either 100% eclogite (solid line) or 50% eclogite (dashed line). The lighter lines show the maximum topography of the tectonic model in Figure 4A, assuming that the lower crust below 40 km is either 100% eclogite (solid line) or 50% eclogite (dashed line).

Insights from the Ordovician of western Ireland check on the admissibility of any such cartoon. The height of a feature zt originally at z0, allowing for subsequent erosion during a given interval (t m.y.), can be estimated using:

⎛ z t = z0 ⎜ 1 − e ⎜ ⎝

ρc . k .t ρm

⎞ ⎟, ⎟ ⎠

(5)

where k is an erosion constant (m.y.–1). Values for k vary from 1 m.y.–1 for Midwestern basins in the United States to 5 m.y.–1 for the Moroccan Rif and 20 m.y.–1 for the New Zealand Alps (estimated from data in Goudie, 1995). The topography associated with slab break-off and the remaining topography after 10 m.y. are shown in Figure 4A. Approximately 5 km of material would have been removed from the forearc region, assuming an erosion constant k = 5 m.y.–1 (Fig. 4A). This analysis suggests that the South Mayo Trough, or any other similar forearc complex, would have been unlikely to survive arc-continent collision followed by slab break-off unless the topography was depressed by some other factor. It is unlikely that this was due to some dynamic process, as the South Mayo region became a site of Silurian marine sedimentation following the Grampian orogeny in basins with modest β factors (see above), suggesting that the crust was never far above sea level in the post-Grampian interval. The model suggests uplift rates in the forearc region on the order of 2 mm.yr –1 to 5 mm.yr –1 during collision and slab break-off. Similar values of 5 mm.yr –1 are recorded in Sumba where a forearc basin is overriding the Australian margin (van der Werff, 1995); this is a region that has been suggested as an analogy for the Grampian orogeny (Dewey and Mange, 1999). However, if such rates were sustained over the duration of the Grampian orogeny, it is unlikely that the South Mayo Trough would have been preserved. Some feature of the Grampian collision zone must, therefore, have depressed the topography of this orogen. Three possibilities exist. Firstly, South Mayo may not have been involved in the Grampian arc-continent collision, but the abundance of evidence reviewed above makes this seem highly unlikely. Secondly, slab pull may have initially reduced the topography dynamically. However, the emplacement of mafic plutons during the principal D2/D3 Grampian deformation in Connemara is generally believed to have been related to delamination of the downgoing Iapetan slab at the time of, not after, peak deformation. A third possibility is that the nature of the continental basement in the footwall may have buffered rapid uplift. ROLE OF THE LAURENTIAN BASEMENT Deep reflection seismic studies (WIRELINE, Klemperer et al., 1991) show that the basement north of the Fair Head–Clew Bay Line off the shore of western and northwestern Ireland is anomalously reflective between ~1.2 s two-way traveltime (TWTT) and the Moho at ~10 s TWTT. Klemperer et al. (1991) attributed this to the presence of considerable volumes of basaltic

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material within the basement of either Tertiary or Neoproterozoic age. Mafic dikes and sills intruded in Argyll Group rocks occur throughout the Dalradian sequence and are particularly prevalent in Donegal, composing up to 20% of the surface geology in some regions (Pitcher and Berger, 1972). Correlation with the Tayvallich lavas of Scotland suggest that they were emplaced at ca. 595 Ma (Halliday et al., 1989). The Dalradian in northwestern Ireland is associated with positive, moderate-frequency magnetic anomalies (Armstrong et al., 1998), which are interpreted as having been caused by these mafic bodies. In Donegal Bay, Carboniferous deposits overstep the Dalradian metasediments, and the associated broad, negative magnetic anomalies mask the pattern associated with the Dalradian rocks. This relationship would not have occurred if the mafic rocks had been post-Carboniferous. Furthermore, the WIRELINE survey did not pass close to any known Tertiary seamounts. Thus it is most likely that the basement reflectivity is associated with the Neoproterozoic basalt swarms emplaced during rifting in Argyll Group times and that in this region the Laurentian margin was a volcanic margin. The presence of voluminous basalt within the deep crust provides a mechanism for depressing topography during convergence (Dewey et al., 1993). If mafic rocks convert to eclogite, their specific gravity increases by up to 20%, and other continental rocks on converting to their eclogite facies equivalents will show increases in specific gravity of between 5% and 10% (Dewey et al., 1993). This is likely to occur below 40 km in a refrigerated system such as a subducting continental margin. The effect of any such conversion to eclogite facies for the lower crust is to increase the overall specific gravity ρc' of the crust as follows:

ρc' =

ρe ∗ Ez + ρc ∗ (Cz − Ez ) , Cz

(6)

where ρe is the density of the eclogitized crust, and Ez is the thickness of the eclogitic layer. Figure 4B shows the effect of substituting the modified values for ρc' into equation 4 on the assumption that all the crust in the collision zone beneath 40 km is eclogitized and that this region has a basaltic component of either 100% or 50%. The analysis is performed for two possible scenarios. The first is the model of Dewey and Mange (1999), shown in Figure 4A, which assumes that the crust on the leading edge of the Laurentian margin is thickened during convergence. The second is based upon the flexural model shown in Figure 3A, which assumes that the Laurentian crust is thick beneath the overriding forearc but thins toward the ocean-continent boundary. The first model predicts that the forearc region will have a mean topography of a few hundred meters above sea level. The frontal region of the arc would be ~4 km below sea level as a result of the negative buoyancy attributable to the thickened eclogite-facies lower crust in this region. The presence of lherzolite within the lithospheric mantle would also depress topography as the depth of

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Ryan

the mantle slab increases and the aluminous phases convert from spinel to garnet below 90 km. However, this effect is small. If a 10% lherzolite component containing 10% spinel is assumed, the maximum depth in this model (Fig. 4B) would increase by 30 m on the southern side of the forearc basin. The second model suggests that the forearc region could be ~200 m below sea level but that much of the frontal arc region would be above sea level if there was 100% eclogite. The second model seems to provide a better fit to the geology of South Mayo, where sedimentation is believed to have continued during collision and slab break-off. However, given the uncertainties in any such analysis, it is difficult to differentiate between the two models. The important point to note is that significant production of eclogite within the footwall to the collision zone can depress the topography sufficiently to allow sedimentation to continue in the forearc region during arc-continent collision. DISCUSSION AND CONCLUSIONS It is suggested that the production of significant volumes of eclogite at the base of the footwall during arc-continent collision may depress the topography sufficiently to preserve the forearc within the hanging wall. This process is likely to be important only along volcanic margins or those that contain older basic massifs. In this context, it is worthy of note that in the Anti-Atlas orogen, where a forearc basin (Sarhro Group) has been preserved during arc-continent collision, the continental basement contains dike swarms and mafic volcanics (Ifzwane Suite) related to rifting (Thomas et al., 2002). This analysis has the limitations of any 2D static model. It cannot, for instance, distinguish between the effects of slab break-off along the length of a continental margin or gradual tearing (Clift et al., 2003). However, any out-of-plane effects of a tearing slab should assist in depressing topography. There also remains the problem that a root of eclogites has not been recorded in any deep geophysical survey of western Ireland. However, the Grampian orogen underwent extensive orogenic collapse in the Late Ordovician (Alsop, 1991; Clift et al., 2004) during which time such a root would have been exhumed and retrogressed to amphibolite. The buoyancy gained during the decompression of the eclogites could have buffered topography and limited subsidence during this phase (Dewey et al., 1993). Such a mechanism may account for the cessation of sedimentation within the South Mayo Trough. An alternative explanation would be to apply the model of Jull and Kelemen (2001) and suggest that the crustal eclogitic root beneath Laurentia became detached and sank into the mantle in late post-Grampian times. This would have driven rapid uplift in the arc region of the model in Figure 4A and in the accretionary complex in the model of Figure 4B. Both regions were sites of Silurian sedimentation (Fig. 2) onto a low-metamorphic-grade basement, making such a process unlikely, at least beneath the South Mayo Trough.

REFERENCES CITED Alsop, G.I., 1991, Gravitational collapse and extension along a mid crustal detachment: The Lough Derg Slide, northwest Ireland: Geological Magazine, v. 128, p. 345–354. Armstrong, G.D., Brown, C., and Mars, J.E., 1998, The Irish continental crust revealed by potential field methods: Technical Program Extended Abstracts, Society of Exploration Geophysicists, 68th Annual Meeting, New Orleans, Louisiana, September 1998; CD-ROM and Extended Abstracts Volume. Boutelier, D., Chemenda, A., and Burg, J.-P., 2003, Subduction versus accretion of intra-oceanic volcanic arcs: Insight from thermo-mechanical analogue experiments: Earth and Planetary Science Letters, v. 212, p. 31–45, doi: 10.1016/S0012-821X(03)00239-5. Chew, D.M., 2003, Structural and stratigraphic relationships across the continuation of the Highland Boundary Fault in western Ireland: Geological Magazine, v. 140, p. 73–85, doi: 10.1017/S0016756802007008. Chew, D.M., Daly, J.S., Page, L.M., and Kennedy, M.J., 2003, Grampian orogenesis and the development of blueschist facies metamorphism in western Ireland: Geological Society [London] Journal, v. 160, p. 911–924. Clift, P.D., and Ryan, P.D., 1994, Geochemical evolution of an Ordovician island arc, South Mayo, Ireland: Geological Society [London] Journal, v. 151, p. 329–342. Clift, P.D., Schouten, H., and Draut, A.E., 2003, A general model of arc continent collision and subduction polarity reversal from Taiwan and the Irish Caledonides, in Larter, R., and Leat, P., eds., Intra-Oceanic Subduction Systems; Tectonic and Magmatic Processes: Geological Society [London] Special Publication 219, p. 81–98. Clift, P.D., Dewey, J.F., Draut, A.E., Chew, D.M., Mange, M., and Ryan, P.D., 2004, Rapid tectonic exhumation, detachment faulting and orogenic collapse in the Caledonides of western Ireland: Tectonophysics, v. 384, p. 91–113, doi: 10.1016/j.tecto.2004.03.009. Condon, D.J., and Prave, A.R., 2000, Two from Donegal: Neoproterozoic episodes on the northeast margin of Laurentia: Geology, v. 28, p. 951–954, doi: 10.1130/0091-7613(2000)28<951:TFDNGE>2.0.CO;2. Daly, J.S., 1996, Pre-Caledonian history of the Annagh Gneiss Complex, northwestern Ireland, and correlation with Laurentia-Baltica: Irish Journal of Earth Sciences, v. 15, p. 5–8. Daly, J.S., 2001, The Precambrian, in Holland, C.H., ed., The Geology of Ireland (2nd edition): Edinburgh, Scottish Academic Press, p. 7–45. Daly, J.S., and Flowerdew, M.J., 2005, Grampian and late Grenville events recorded by mineral geochronology near a basement cover contact in north Mayo, Ireland: Geological Society [London] Journal, v. 162, p. 163–174, doi: 10.1144/0016-764903-150. Dewey, J., and Mange, M., 1999, Petrography of Ordovician and Silurian sediments in the western Ireland Caledonides: Tracers of a short-lived Ordovician continent arc collision orogeny and the evolution of the Laurentian Appalachian Caledonian margin, in MacNiocaill, C., and Ryan, P.D., eds., Continental Tectonics: Geological Society [London] Special Publication 164, p. 55–107. Dewey, J.F., and Ryan, P.D., 1990, The Ordovician evolution of the South Mayo Trough, western Ireland: Tectonics, v. 9, p. 887–901. Dewey, J.F., and Shackleton, R.M., 1984, A model for the evolution of the Grampian tract in the early Caledonides and Appalachians: Nature, v. 312, p. 115–121, doi: 10.1038/312115a0. Dewey, J.F., McKerrow, W.S., and Moorbath, S., 1970, The relationship between isotopic ages, uplift and sedimentation during Ordovician times in western Ireland: Scottish Journal of Geology, v. 6, p. 133–145. Dewey, J.F., Ryan, P.D., and Andersen, T.B., 1993, Orogenic uplift and collapse, crustal thickness, fabrics and metamorphic phase changes: The role of eclogites, in Prichard, H.M., et al., eds., Magmatic Processes and Plate Tectonics: Geological Society [London] Special Publication 76, p. 325–343. Draut, A.E., and Clift, P.D., 2001, Geochemical evolution of arc magmatism during arc-continent collision, South Mayo, Ireland: Geology, v. 29, p. 543– 546, doi: 10.1130/0091-7613(2001)029<0543:GEOAMD>2.0.CO;2. Draut, A.E., and Clift, P.D., 2002, The origin and significance of the Delaney Dome Formation, Connemara, Ireland: Geological Society [London] Journal, v. 159, p. 95–103.

Insights from the Ordovician of western Ireland Draut, A.E., Clift, P.D., Chew, D.M., Cooper, M.J., Taylor, R.N., and Hannigan, R.E., 2004, Laurentian crustal recycling in the Ordovician Grampian Orogeny: Nd isotopic evidence from South Mayo, western Ireland: Geological Magazine, v. 141, p. 195–207, doi: 10.1017/S001675680400891X. Flowerdew, M.J., Daly, J.S., and Whitehouse, M.J., 2005, 470 Ma granitoid magmatism associated with the Grampian Orogeny in the Slishwood Division, NW Ireland: Geological Society [London] Journal, v. 162, p. 563–575, doi: 10.1144/0016-784904-067. Friedrich, A.M., 1998, 40Ar/39Ar and U-Pb geochronological constraints on the thermal and tectonic evolution of the Connemara Caledonides [Ph. D. thesis]: Cambridge, Department of Earth, Atmospheric and Planetary Sciences, Massachusetts Institute of Technology, 228 p. Friedrich, A.M., Hodges, K.V., Bowring, S.A., and Martin, M.W., 1999a, Geochronological constraints on the magmatic, metamorphic, and thermal evolution of the Connemara Caledonides, western Ireland: Geological Society [London] Journal, v. 156, p. 1217–1230. Friedrich, A.M., Bowring, S.A., Martin, M.W., and Hodges, K.V., 1999b, Shortlived continental magmatic arc at Connemara, western Irish Caledonides: Implications for the age of the Grampian orogeny: Geology, v. 27, p. 27– 30, doi: 10.1130/0091-7613(1999)027<0027:SLCMAA>2.3.CO;2. Goudie, A.S., 1995, The Changing Earth: Rates of Geomorphological Processes: London, Blackwell, 302 p. Gray, J.R., and Yardley, B.W.D., 1979, A Caledonian blueschist from the Irish Dalradian: Nature, v. 278, p. 736–737, doi: 10.1038/278736a0. Halliday, A.N., Graham, C.R., Aftalion, M., and Dymoke, P., 1989, The depositional age of the Dalradian Supergroup: U-Pb and Sm-Nd isotopic studies of the Tayvallich Volcanics, Scotland: Geological Society [London] Journal, v. 146, p. 3–6. Harper, D.A.T., Williams, D.M., and Armstrong, H.A., 1989, Stratigraphical correlations adjacent to the Highland Boundary Fault in the west of Ireland: Geological Society [London] Journal, v. 146, p. 381–384. Hutton, D.H.W., 1987, Strike slip terranes and a model for the evolution of the British and Irish Caledonides: Geological Magazine, v. 124, p. 405–425. Jull, M., and Kelemen, P.B., 2001, On the conditions for lower crustal convective instability: Journal of Geophysical Research, v. 106, p. 6423–6446, doi: 10.1029/2000JB900357. Khan, M.A., Stern, R.J., Gribble, R.F., and Windley, B.F., 1997, Geochemical and isotopic constraints on subduction polarity, magma sources and palaeogeography of the Kohistan Arc, northern Pakistan: Geological Society [London] Journal, v. 154, p. 935–946. Klemperer, S.L., Ryan, P.D., and Snyder, D.B., 1991, A deep seismic reflection transect across the Irish Caledonides: Geological Society [London] Journal, v. 148, p. 149–164. Konstantinovskaia, E.A., 2000, Geodynamics of an Early Eocene arc–continent collision reconstructed from the Kamchatka Orogenic Belt, NE Russia: Tectonophysics, v. 325, p. 87–105, doi: 10.1016/S0040-1951(00)00132-3. Konstantinovskaia, E.A., 2001, Arc–continent collision and subduction reversal in the Cenozoic evolution of the Northwest Pacific: An example from Kamchatka (NE Russia): Tectonophysics, v. 333, p. 75–94, doi: 10.1016/ S0040-1951(00)00268-7. Kuzmichev, A.B., Bibikova, E.V., and Zhuravlev, D.Z., 2001, Neoproterozoic (~800 Ma) orogeny in the Tuva-Mongolia Massif (Siberia): Island arc–continent collision at the northeast Rodinia margin: Precambrian Research, v. 110, p. 109–126, doi: 10.1016/S0301-9268(01)00183-8. Leake, B.E., Tanner, P.W.G., Singh, D., and Halliday, A.N., 1983, Major southward thrusting of the Dalradian rocks of Connemara, western Ireland: Nature, v. 305, p. 210–213, doi: 10.1038/305210a0. Max, M.D., and Riddihough, R.P., 1975, Continuation of the Highland Boundary Fault in Ireland: Geology, v. 3, p. 206–210, doi: 10.1130/00917613(1975)3<206:COTHBF>2.0.CO;2. Max, M.D., Ryan, P.D., and Inamdar, D.D., 1983, A magnetic deep structural geology interpretation of Ireland: Tectonics, v. 2, p. 431–451. McKenzie, D.P., 1969, Speculations on the consequences and causes of plate motions: Geophysical Journal of the Royal Astronomical Society, v. 18, p. 1–32. Menuge, J.F., Williams, D.M., and O’Connor, P.D., 1995, Silurian turbidites used to reconstruct a volcanic terrain and its Mesoproterozoic basement in the Irish Caledonides: Geological Society [London] Journal, v. 152, p. 269–278.

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Murphy, F.C., Anderson, T.B., Daly, J.S., Gallagher, V., Graham, J.R., Harper, D.A.T., Johnston, J.D., Kennan, P.S., Kennedy, M.J., Long, C.B., Morris, J.H., O’Keefe, W.G., Parkes, M., Ryan, P.D., Sloan, R.J., Stillman, C.J., Tietzsch-Tyler, D., Todd, S.P., and Wrafter, J.P., 1991, An appraisal of Caledonian suspect terranes in Ireland: Irish Journal of Earth Sciences, v. 11, p. 11–41. O’Brien, P.J., 2001, Subduction followed by collision: Alpine and Himalayan examples: Physics of the Earth and Planetary Interiors, v. 127, p. 277– 291, doi: 10.1016/S0031-9201(01)00232-1. Phillips, W.E.A., 1973, The pre-Silurian rocks of Clare Island, Co. Mayo, Ireland, and the age of metamorphism of the Dalradian in Ireland: Geological Society [London] Journal, v. 129, p. 585–606. Pitcher, W.S., and Berger, A.R., 1972, The Geology of Donegal: A Study of Granite Emplacement and Unroofing: New York, John Wiley, 435 p. Robertson, A.H.F., and Collins, A.S., 2002, Shyok Suture Zone, N Pakistan: Late Mesozoic-Tertiary evolution of a critical suture separating the oceanic Ladakh Arc from the Asian continental margin: Journal of Asian Earth Sciences, v. 20, p. 309–351, doi: 10.1016/S1367-9120(01)00041-4. Rolland, Y., Pecher, A., and Picard, C., 2000, Middle Cretaceous back-arc formation and arc evolution along the Asian margin: The Shyok Suture Zone in northern Ladakh (NW Himalaya): Tectonophysics, v. 325, p. 145–173, doi: 10.1016/S0040-1951(00)00135-9. Ryan, P.D., and Archer, J.B., 1978, The Lough Nafooey fault: A Taconic structure in western Ireland: Geological Survey of Ireland Bulletin, v. 2, p. 255–264. Ryan, P.D., and Dewey, J.F., 1991, A geological and tectonic cross-section of the Caledonides of western Ireland: Geological Society [London] Journal, v. 148, p. 173–180. Ryan, P.D., and Dewey, J.F., 2004, The South Connemara Group reinterpreted: A subduction-accretion complex in the Caledonides of Galway Bay, western Ireland: Journal of Geodynamics, v. 37, p. 513–529, doi: 10.1016/ j.jog.2004.02.018. Ryan, P.D., Floyd, P.A., and Archer, J.B., 1980, The stratigraphy and petrochemistry of the Lough Nafooey Group (Tremadocian), western Ireland: Geological Society [London] Journal, v. 137, p. 443–458. Ryan, P.D., Sawal, V.K., and Rowlands, A.S., 1983, Ophiolitic mélange separates ortho- and para-tectonic Caledonides in western Ireland: Nature, v. 302, p. 50–52, doi: 10.1038/302050a0. Soper, N.J., Ryan, P.D., and Dewey, J.F., 1999, Age of the Grampian Orogeny in Scotland and Ireland: Geological Society [London] Journal, v. 148, p. 173–180. Strachan, R.A., and Holdsworth, R.E., 2000, Late Neoproterozoic (<750 Ma) to Early Ordovician passive margin sedimentation along the Laurentian margin of Iapetus, in Woodcock, N., and Strachan, R., eds., Geological History of Britain and Ireland: London, Blackwell, p. 73–87. Teng, L.S., 1996, Extensional collapse of the northern Taiwan mountain belt: Geology, v. 24, p. 949–952, doi: 10.1130/0091-7613(1996)024<0949: ECOTNT>2.3.CO;2. Thomas, R.J., Chevallier, L.P., Gresse, P.G., Harmer, R.E., Eglington, B.M., Armstrong, R.A., de Beer, C.H., Martini, J.E.J., de Kock, G.S., Macey, P.H., and Ingram, B.A., 2002, Precambrian evolution of the Sirwa Window, Anti-Atlas Orogen, Morocco: Precambrian Research, v. 118, p. 1– 57, doi: 10.1016/S0301-9268(02)00075-X. van der Werff, W., 1995, Cenozoic evolution of the Savu Basin, Indonesia: Fore-arc basin response to arc-continent collision: Marine and Petroleum Geology, v. 12, p. 247–262, doi: 10.1016/0264-8172(95)98378-I. Wang, K.Y., Li, J.L., Hao, J., Li, J., and Zhou, S., 1996, The Wutaishan orogenic belt within the Shanxi Province, northern China: A record of late Archaean collision tectonics: Precambrian Research, v. 78, p. 95–103, doi: 10.1016/0301-9268(95)00071-2. Watts, A.B., 2001, Isostasy and Flexure of the Lithosphere: Cambridge, England, Cambridge University Press, 458 p. Williams, D.M., and Harper, D.A.T., 1991, End-Silurian modifications of Ordovician terranes in western Ireland: Geological Society [London] Journal, v. 148, p. 165–171.

MANUSCRIPT ACCEPTED BY THE SOCIETY 24 APRIL 2007

Printed in the USA

The Geological Society of America Special Paper 436 2008

Basin formation by volcanic arc loading Dave Waltham Robert Hall Helen R. Smyth Cynthia J. Ebinger SE Asia Research Group, Department of Geology, Royal Holloway University of London, Egham, Surrey, TW20 0EX, UK

ABSTRACT This paper quantifies the flexural subsidence expected from loading by a volcanic arc. The resulting mathematical model shows that the arc width should grow with time and that the subsidence beneath the load can be estimated from the observed arc width at the surface. Application of this model to the Halmahera Arc in Indonesia shows an excellent fit to observations if a broken-plate model of flexure is assumed. The model also gives an excellent fit to data from East Java, also in Indonesia, where it is possible to forward model gravity anomalies. In particular, the depth, location, and width of the depocenter-associated gravity low are accurately reproduced, although the model does require a high density for the volcanic arc (2900 kg m–3). This may indicate additional buried loads due, for example, to magmatic underplating. Our main conclusion is that loads generated by the volcanic arc are sufficient to account for much, if not all, of the subsidence in basins within ~100 km of active volcanoes at subduction plate boundaries, if the plate is broken. The basins will be asymmetrical and, close to the arc, will contain coarse volcaniclastic material, whereas deposits farther away are likely to be volcaniclastic turbidites. The density contrast between arc and underlying crust required to produce the Indonesian arc basins means that they are unlikely to form in young intraoceanic arcs but may be common in older and more mature arcs. Keywords: volcanic arc, loading, flexural subsidence, Indonesian arcs.

mechanism responsible for basin formation. Dewey (1980) suggested that these basins result from extension caused by rollback of the subduction hinge; i.e., the location of subduction moves progressively away from the overriding plate and so drags the overriding plate with it. Another suggestion is that extensional stresses are set up by secondary mantle currents created by subduction (Toksöz and Bird, 1977). It is also possible that basins result from loading by magmatic underplating, as has been suggested for loading of oceanic lithosphere by volcanic islands (Watts et al., 1985). Another possibility (Bahlburg and Furlong,

INTRODUCTION This paper investigates whether the isostatic load of volcanic arcs could be responsible for a significant fraction of the subsidence frequently observed in volcanic arc settings. Arc volcanism occurs in the overriding plate of a subduction system by melting of the mantle wedge owing to introduction of volatiles carried beneath it by the subducting plate. Basins commonly form both trenchward of the arc (the forearc) and behind the arc (the backarc). However, there is no consensus concerning the precise

Waltham, D., Hall, R., Smyth, H.R., and Ebinger, C.J., 2008, Basin formation by volcanic arc loading, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 11–26, doi: 10.1130/2008.2436(02). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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1996; Smith et al., 2002) is that the arc volcanism itself produces significant near-surface loads leading to the formation of flexural basins. This last suggestion is the one pursued in this paper, although we emphasize that basins form in arcs by a number of mechanisms, and we are not proposing that all basins form in this way. The model has particular relevance to basins that form very close to the volcanic arc. The proximal parts of such basins will be characterized by coarse volcanic debris of primary volcanic and sedimentary origin, typically terrestrial to shallow marine, and may include mass-flow deposits of different types. The idea that the lithosphere flexes in response to volcanic loads is not new, although it has most commonly been applied to loading of oceanic lithosphere by island chains associated with hot-spot volcanism (Watts and Cochran, 1974; Watts et al., 1997). Bahlburg and Furlong (1996) applied a continuous-plate flexure model to subsidence in the Ordovician foreland of northwestern Argentina. Their model used geographically extensive volcanic loads and was able to produce >8 km of subsidence. In a similar study using the same modeling algorithm (attributed to Slingerland et al., 1994), Smith et al. (2002) modeled subsidence of the Abiquiu Embayment in the Rio Grande Rift, southwestern USA. The Rio Grande study also employed a volcanic load that filled most of the resulting basin, but the calculations of Smith et al. (2002) accounted for only 800 m of subsidence in that area. Other authors (e.g., Karner and Watts, 1983; Nunn et al., 1987) applied similar models to continental subsidence resulting from thrust-emplaced loads on an unbroken plate. These studies showed that gravity profiles were consistent with flexural subsidence, although, in most cases, additional subsurface loads were required to explain the observed subsidence. In this paper we produce a semianalytical model relating the observed surface width of the loading volcanic arc to the resulting subsidence. We also look at the consequences of assuming a broken-plate rather than a continuous-plate model for the underlying lithosphere. Our approach principally differs from that employed by others in that it predicts the volcanic load from first principles, thus providing constraints on loading that are independent of those obtained, for example, from gravity modeling. Such an approach is difficult for geometrically complex thrust-emplaced loads (Nunn et al., 1987) but, as shown here, is relatively straightforward for a volcanic arc load. EVIDENCE FOR ARC LOADING IN INDONESIA In many volcanic arcs, thick sequences of sedimentary and volcaniclastic rocks in basins are close to, and on both sides of, the volcanic arc. Much, if not all, of the material in the basins is derived from the arc itself, and much of it is very coarse. In Indonesia (Fig. 1), many basins near modern volcanic arcs have proved rich in hydrocarbons and consequently have been the target of exploration work. The North, Central, and South Sumatra Basins, and the offshore Northwest Java and East Java Basins are examples. However, these basins are typically >100 km from the active volcanic arc and have been described as backarc basins

(e.g., Busby and Ingersoll, 1995), although they are not floored by oceanic crust and lack many features of backarc basins (Smyth et al., 2007, this volume). Nor are they retro-arc foreland basins formed in response to thrusting and dynamic subsidence driven by subduction (e.g., DeCelles and Giles, 1996), as pointed out by Moss and McCarthy (1997). The Sunda Shelf basins have many characteristics of rift basins (Cole and Crittenden, 1997; Hall and Morley, 2004). However, the basins we are concerned with are much closer to the arc itself. Seismic lines cross many of these basins, but it is rare for them to approach close to the volcanic arc because the volcanic-rich sequences are normally considered to be relatively unprospective as far as hydrocarbon source rocks are concerned. Also, potential reservoir rocks commonly have problems with loss of porosity and permeability owing to breakdown of unstable grains, and close to the arcs, seismic data are difficult to acquire and interpret. We have studied a number of Indonesian volcanic arcs in the field and have worked on the sequences in basins close to the volcanic arcs. The ages and histories of the basins suggest that their development is related closely to the development of the volcanic arc. We have chosen two different arcs in Indonesia for which the history of the basin and the arc is known, and where volcanic loading is a plausible explanation of at least some of the basin subsidence. These are South Halmahera and East Java (Fig. 1), where the arcs are built on different types of crust. South Halmahera Halmahera has a long volcanic arc history extending from the Late Jurassic, almost entirely intraoceanic (Hall et al., 1995). The present arc is in its final stages of activity, as subduction has nearly eliminated the Molucca Sea, and the Sangihe and Halmahera Arcs (Fig. 2), on the west and east sides of the Molucca Sea, are actively colliding. The Molucca Sea plate dips east under Halmahera and west under the Sangihe Arc in an inverted U-shape (McCaffrey et al., 1980). Seismicity shows ~200–300 km of lithosphere subducted beneath Halmahera, whereas the Benioff zone associated with the west-dipping slab can be identified to a depth of at least 600 km beneath the Sangihe Arc. The modern arc (Figs. 1 and 2) resumed activity in the Quaternary, after a brief decline in volcanic activity, and the axis of the arc moved ~50 km west from its late Miocene to Pliocene position (Hall et al., 1988b, 1995). Both these arcs are built on older volcanic arcs active during Cretaceous and early Cenozoic time. The Quaternary and Neogene Halmahera Arcs have a chemical character typical of intraoceanic arcs (Morris et al., 1983; Forde, 1997). There is no evidence of continental crust beneath them except at the southernmost end of the volcanic arc (Morris et al., 1983), on Bacan and Obi, where movement on strands of the Sorong fault zone brought slivers of the continental crust beneath the arc in the last few million years (Hall et al., 1995; Ali et al., 2001). Mapping of the Halmahera Arc shows that the basement is ophiolitic, formed in an early Mesozoic intraoceanic arc (Hall et al., 1988a; Ballantyne, 1992), overlain by Cretaceous, Eocene,

Basin formation by volcanic arc loading Manila Trench

Hainan

Indochina

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Malay peninsula

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a nd Su

C Su ent ma ral tra

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gh rou helf or T ul S m i T Sah

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Figure 1. Major basins of the southern Sundaland margin and location of the Halmahera and Java Arcs.

and Oligocene arc volcanic rocks. Shallow Miocene marine carbonates unconformably overlie all the older rocks. The Neogene Halmahera Arc became active at ca. 11 Ma at its southern end on Obi (Hall et al., 1995). In SW Halmahera, volcanic activity began a little later as magmatism propagated north. Basins formed on each side of the arc. In South Halmahera, activity in the volcanic arc, and subsidence on each side of the arc, began in the late Miocene at ca. 8 Ma. To the west, turbidites and debris flows were deposited in the forearc, and to the east (Fig. 2), a marine basin developed in Weda Bay (Hall, 1987; Hall et al., 1988b; Nichols and Hall, 1991). On the western side of the SW arm of Halmahera (Fig. 2) there was subsidence and deposition of at least 1000 m of submarine slope deposits in the forearc (Nichols and Hall, 1991). In the forearc, subsidence significantly exceeded the supply of material. In contrast, on the backarc side, sediment supply broadly kept pace with subsidence. On land in the SW arm of Halmahera there are between 2800 and 3800 m of sedimentary rocks in the basin east of the arc (Nichols and Hall, 1991). They were deposited close to the arc in shallow water and rest unconformably on

pre-Miocene volcanic rocks or locally on shallow-water lower to middle Miocene limestones. The sequence fines up from fan-delta conglomerates into sandstones, mudstones, and limestones. Shallow marine deposition continued into the Pliocene. The depth of water at the time of deposition of the upper part of the sequence is uncertain but was no more than a few hundred meters. The Miocene–Pliocene sequence interpreted from the seismic sections (Fig. 3) across Weda Bay varies in thickness. It rests unconformably on a karstified limestone surface and is up to 2000 m thick in local depocenters but is typically ~1000 m thick. Hence the basin was markedly asymmetric, with the greatest thickness of sediment adjacent to the Halmahera Arc on the western side of the basin. The local depocenters offshore, and the fan deltas onshore, are thickest close to interpreted volcanic centers. There is little evidence of faulting associated with basin formation on seismic lines (Fig. 3), and there is no indication of rifting. The lower-middle Miocene limestones represent an approximate sea-level datum and probably covered the whole area; close to the arc they were removed by erosion before deposition of the basin sequence. The Halmahera Basin was actively

200

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nch Tre

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e pin ilip Ph

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Figure 2. (A) Location and major tectonic features of the Molucca Sea region. Small, black-filled triangles are modern volcanoes. Bathymetric contours are at 200, 2000, 4000, and 6000 m. Large barbed lines are subduction zones, and small barbed lines are thrusts. (B) Cross section across the Halmahera and Sangihe Arcs on section line B. Thrusts on each side of the Molucca Sea are directed outward toward the adjacent arcs, although the subducting Molucca Sea plate dips east beneath Halmahera and west below the Sangihe Arc. (C) Inset is the restored cross section of the Miocene–Pliocene Weda Bay Basin of SW Halmahera on section line C, flattened to the Pliocene unconformity, showing estimated thickness of the section.

0

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3 Figure 3. Uninterpreted seismic lines and interpreted sections across parts of the sedimentary sequences deposited at the western edge of Weda Bay, Halmahera. Locations of lines A and B are shown on inset map C. There is little evidence of faulting on the sections and no indication of a rift character to the basin. The Miocene–Pliocene arc-derived sequence is deposited on top of middle Miocene shallow marine limestones and is terminated by an intra-Pliocene unconformity. The sequence thickens toward the Miocene–Pliocene volcanic arc. TWT—two-way traveltime.

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subsiding from the late Miocene until the middle to late Pliocene, an interval of ~5 m.y. Based on field measurement of sections and biostratigraphic dating, Nichols and Hall (1991) estimated that during deposition of this interval the western side of the basin, adjacent to the arc, subsided at least 2.8 km, representing an average subsidence rate of at least 47 cm/1000 yr. Subsidence rates for the eastern part of the basin, ~100 km from the arc, were between 17 and 34 cm/1000 yr. At ca. 3 Ma, sedimentary rocks of this basin were thrust west over the Neogene arc, and later there was thrusting of the forearc from the west. The Quaternary Halmahera Arc is built unconformably upon all these rocks. The thrusting is a result of collision of the Sangihe and Halmahera Arcs. In Weda Bay there has been recent subsidence, and its deepest parts are almost 2000 m below sea level. Part of this subsidence is probably due to thrust loading to the west, but part may result from movements along splays of the Sorong Fault that appear to control the form of the present-day depocenter (Nichols and Hall, 1991).

o

o

110 E

112 E

East Java East Java (Fig. 4) is situated on the continental margin of Sundaland. There has been subduction to the south of Java, along the Java Trench, since the early Cenozoic (Hall, 2002). The basement of most of East Java has previously been interpreted as arc and ophiolitic material accreted to the continental margin in the Late Cretaceous, but our work (Smyth et al., this volume) has shown that there is old continental crust beneath the Southern Mountains of East Java. Our work in East Java also suggests that the present northward subduction of Indian-Australian lithosphere began in the middle Eocene. The oldest Cenozoic rocks resting on older basement are terrestrial conglomerates without volcanic material, but a short distance above these rocks volcanic debris appears in middle Eocene sediments and increases in abundance upsection (Smyth, 2005). There is a record of two volcanic arcs (Fig. 4) in East Java (Smyth et al., this volume). An early Cenozoic arc formed in the

o

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Approximate location of the Oligocene-Miocene volcanoes of the Southern Mountains Arc

\ o

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Deposits of the Southern Mountains Volcanic Arc Modern volcanoes of the Sunda Arc Limit of the strong negative gravity anomaly defined by -150 µms-2 contour

o

8S

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Approximate line of section (after Pertamina, 1996)

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N East Java Sea: Edge of the Sunda Shelf

Kendeng Basin: Thick sequence behind Eocene to Early Miocene arc

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Southern Mountains Arc: Eocene to Early Miocene volcanic arc

5 km

Basement

100 km

Figure 4. Location of the Eocene to early Miocene volcanic arc in the Southern Mountains of East Java, the Kendeng Basin, and the modern arc. Inferred volcanic centers of the Southern Mountains Arc are shown with open triangles, and active volcanoes are shown with solid triangles. A cross section of the Kendeng Basin along the line indicated is based on Pertamina (1996). The Kendeng Basin sequence is interpreted to thicken toward the Southern Mountains volcanic arc.

Basin formation by volcanic arc loading Southern Mountains and was active from the middle Eocene to the early Miocene (ca. 42–18 Ma). It erupted material ranging in composition from andesite to rhyolite (Smyth, 2005). The arc formed a chain of volcanic islands during the early Cenozoic that were initially similar in character to volcanoes of the present-day Izu-Bonin-Mariana Arc and the Aleutian Islands, except that behind the arc was a broadly shallow marine shelf and no deep oceanic backarc basin. The volcanoes were terrestrial, but some of their products were deposited close to the arc in a marine setting, and they formed separate small islands. The Oligocene volcanic centers are well preserved and have a spacing similar to volcanoes of the modern arc on Java (Smyth et al., this volume). They are predominantly andesitic and have been described as the “Old Andesites” (van Bemmelen, 1949). However, the volcanoes erupted considerable volumes of more siliceous material in explosive Plinian eruptions, and this material was dispersed widely as ash. The volume of siliceous material has been overlooked in descriptions of volcanic activity on Java. More details of the stratigraphy are given in Smyth et al. (this volume).

17

Immediately behind, and to the north of, the Southern Mountains Arc is the deep Kendeng Basin (Fig. 4). The basin is long (at least 400 km) and narrow (100–120 km) and trends east-west, parallel to the Southern Mountains Arc. The basin is characterized by a strong negative Bouguer gravity anomaly (Fig. 5), which exceeds −580 μms–2, and extends from west to east. The basin formed during the middle Eocene (Untung and Sato, 1978). The Kendeng Basin succession is not well exposed but contains much volcanic debris carried north from the arc. The oldest rocks are not seen in situ but are sampled by mud volcanoes currently erupting through the basin sequence. They are terrestrial and shallow marine rocks similar to those deposited close to the arc during the Eocene (de Genevraye and Samuel, 1972). The Kendeng Basin succession records a deepening of the basin with time, and during the Oligocene thick sequences of volcaniclastic turbidites and pelagic mudstones were deposited, suggesting that subsidence exceeded the supply of material. Seismic lines across the northern parts of the depocenter show that the Kendeng Basin sequence thickens toward the Southern Mountains Arc and is ~3 km thick in the north (Fig. 4; Pertamina,

Figure 5. Bouguer gravity anomaly map of East Java and location of the modeled profile discussed in the text.

18

Waltham et al.

1996). Untung and Sato (1978) suggested that the deeper parts of the basin contain ~6 km of section. Gravity calculations suggest there may be as much as 10 km of sediment in its thickest parts. The increase in thickness of the sequence toward the arc and the importance of volcanic debris suggest that most of the Kendeng Basin fill was derived from the arc. Because the Kendeng Basin is so poorly exposed, and there are no seismic lines crossing the entire basin, it is impossible to assess the role of faulting in the basin’s development. South- and north-dipping thrusts shown on interpreted sections across the basin (Pertamina, 1996) could be reactivated normal faults. To the north of the Kendeng Basin was a carbonate and clastic shelf during the early Cenozoic. This was the edge of the Sundaland continent. During the Eocene to early Miocene there were elongate emergent ridges, trending roughly NE-SW and oblique to the arc, separated by depressions. These ridges contain terrestrial clastic sediments and coals at their base, thought to be Eocene, overlain by shallow marine Eocene to Miocene clastic sediments and platform carbonates. There are no reports of volcanic material in the shelf sequences, although reported clay layers may be fine-grained volcanogenic air-fall tuff deposited far from the arc. At the shelf edge there is volcanic material, including clays, zircons, and volcanic quartz (Smyth, 2005). The basin edge at the southern edge of the shelf is partly exposed in a foldthrust belt where there was 10–30 km of shortening, suggested to have occurred during the Pliocene (de Genevraye and Samuel, 1972) or late Miocene. The Southern Mountains Arc extends to the south coast of Java, and there has been almost no hydrocarbon exploration offshore directly south of the coast, so little is known about this region. Marine geophysical studies show ~1 km of sedimentary cover of unknown age on basement in small forearc basins (Kopp et al., 2006). The Southern Mountains Arc has been elevated and tilted since the early Miocene and now dips uniformly to the south at ~30° (Smyth, 2005). On land the arc is ~40 km wide. Arc activity ceased in the early Miocene, followed by a period of little or no volcanic activity (Smyth et al., this volume). Volcanic arc activity resumed in the late Miocene ~50 km north of the Southern Mountains Arc, and the modern volcanoes are built on the Kendeng Basin (Smyth et al., this volume). The products of the late Miocene to Holocene arc are more basic than the older arc and are predominantly basaltic andesites to andesites (e.g., SoeriaAtmadja et al., 1994; van Bemmelen, 1949). Causes of Basin Formation In South Halmahera and East Java there is association between volcanic arc activity and basin formation. Both areas were emergent or close to sea level at the time of basin initiation. Subsidence began as the volcanic arc became active. Both basins have asymmetrical profiles, with the greatest thickness of material close to the volcanic arc. Clastic input came from the volcanic arc, and both basins are dominated by volcanic debris. In

Halmahera, local depocenters contain sequences that are thickest close to inferred volcanic centers. Some crustal extension is likely in both volcanic arcs and is probably required in order to allow magma to reach the surface. However, in Halmahera there is little evidence for extensional faulting in the basin sequences, and no significant faulting is seen on seismic lines. In neither basin is there evidence for a typical rift character, nor is there evidence for significant crustal extension. In neither region are there oceanic backarc basins. There is no evidence for thrusting before or during sedimentation. Thrusting, unrelated to arc development, occurred in both areas after the basins formed and filled. In Halmahera, arc-arc collision caused thrusting, first from the backarc side and then from the forearc side, directed toward the volcanic arc. In East Java there was mainly northward-thrusting of the Kendeng Basin at its northern edge some time after volcanic arc activity ceased. All these observations suggest that volcanic activity contributed in some way to basin formation, possibly through loading by the volcanic arc itself or possibly by weakening of the plate, or by a combination of both. MATHEMATICAL MODEL In this section we develop a simple analytical model for describing how load-generated subsidence is controlled by an accumulating volcanic arc. The aim is to produce the simplest possible model capable of testing whether the observed subsidence is compatible with the likely size of load. The primary assumptions of our model are that subsidence is caused by a volcanic arc–generated line load and that the subsidence directly under the load is proportional to the size of the load, i.e., s = kV,

(1)

where s is subsidence, k is a constant, and V is the load (Fig. 6). For flexure of a uniform, unbroken beam, k is given by (Turcotte and Schubert, 2002, equations 3.127, 3.131, and 3.135) k = π/(2xbΔρ1g),

(2)

where xb is the flexural-bulge to arc-center distance, Δρ1 is the density contrast between the basin fill and the asthenosphere, and g is the acceleration from gravity. For a broken plate the equivalent formula is (Turcotte and Schubert, 2002, equations 3.127, 3.141, and 3.144) k = 3π/(4xbΔρ1g).

(3)

The resulting basin is assumed to be filled by sediments up to a horizontal surface and also filled by the volcanic arc but with a triangular (in cross section) subaerial load caused by the current volcanic edifice itself. The load therefore consists of two components, a load resulting from the density excess of the buried volcanic arc plus a load resulting from the subaerial volcanic edifice. Hence

Basin formation by volcanic arc loading

19

Xb W

Sediment fill S Flexing Lithosphere

Asthenosphere

Figure 6. Schematic diagram showing concept of basin formation by volcanic arc loading. W—arc width; xb—flexural bulge to arc-center distance; s—subsidence.

s ⎧ ⎫ V = ⎨Δρ2 ∫ wds' + ρβw2 ⎬ g , 0 ⎩ ⎭

(4)

where Δρ2 is the density excess of the volcanic arc over the remaining basin fill, w(s) is the arc width as a function of subsidence, ρ is the volcanic arc density, and β is given by β = 0.5 tan(θ) = h/W,

edifice) is unlikely to be available, and so a theoretical model for arc width as a function of subsidence is required. Combining equations 1 and 4, differentiating with respect to s, and using the boundary condition that w(0) = 0 produces, 1 = Aw + 2Bww′,

(6)

where the prime indicates differentiation and

(5)

A = kgΔρ2,

(7)

where θ is the volcano slope, h is the final volcano height, and W the final arc width. Note that equation 4 implies that there was no significant deformation of the volcanic arc during basin subsidence. Given a well-constrained cross-sectional geometry, equations 1 through 4 alone would be sufficient to test the concept that basins can be formed by flexural loading resulting from a volcanic arc. In practice the required information (e.g., detailed geometry of the arc deposits beneath the present-day volcanic

B = kgβρ,

(8)

with k given by equation 2 or 3. The solution to ordinary differential equation 6 is (Appendix) w = (1/A) f(A2s/B),

(9)

where f(x) is the function shown in Figure 7A. Hence, the width of the arc increases rapidly in the early stages of subsidence (i.e.,

A f 0.9 0.8 0.7 0.6 0.5 0.4 0.3 0.2 0.1 0

0

1

2

B

5

4

3

x

6

Arc Width (km) 0

5

10

15

20

25

30

35

40

0

4 Continuous-plate (CP)

6

Co

Broken-plate (BP) BP sediment density +10%

nti

nu

ou

s-p

e 0%

0%

+1

+1

ity

h dt

ns

wi

de

sin

Ba

Arc

BP basin width +10% BP slope +10%

10

e

at pl

BP arc density +10%

8

lat

n-

ke

o Br

Subsidence (km)

2

12

Figure 7. (A) Loading function derived in Appendix. Arc width increases with subsidence along a suitably scaled version of this curve. (B) Sensitivity analysis. The highest dashed line shows the result for the continuous-plate model, equation 9, and the average values from Table 1. The solid line shows the result for the broken-plate model, equation 9, and the average values from Table 1. The other curves show the result of increasing each parameter by 10% using the broken-plate model.

Basin formation by volcanic arc loading when A2s/B is small) and asymptotically approaches a maximum width of 1/A as subsidence becomes large. Figure 7B illustrates the effect of arc width on basin subsidence predicted from equation 9. The “standard” simulation uses the average values of the parameters given in Table 1 and assumes a broken-plate model (i.e., it uses equation 3 rather than equation 2). This simulation shows that highly significant subsidence can be obtained from reasonable parameter values given sensible widths for a volcanic arc. Figure 7B also shows the sensitivity to parameter uncertainties. Each of the curves is obtained by increasing a single parameter 10% above its standard value. The result of using a continuous rather than a broken plate is also shown. Figure 7B demonstrates that the most important factor is whether or not the plate can be considered broken. Other factors have relatively minor influences, although uncertainties in the arc density are, perhaps, the most significant factor. However, it should be noted that a 10% uncertainty in arc density is unlikely, whereas a 10% uncertainty in volcanic slope or in basin width is realistic. In the following section, observations on the geometry and gravity profiles from two Indonesian arcs are used to test the model developed in this section.

Although the backarc basin has been thrust westward, the very coarse character of the Neogene fan-delta deposits suggests they were deposited within a few kilometers of the arc. The seismic lines show that the basin deposits thin eastward, away from the arc, and suggest that the basin width was between 60 and 80 km. The estimated thickness of the sequence close to the Neogene arc is between 2800 and 3800 m (Nichols and Hall, 1991) and up to 2 km offshore farther from the arc. In addition to this information, the mathematical model requires densities for the basin fill, mantle, and volcanic arc. Based upon Hamilton (1979), we assume a basin-fill density of 2300 ± 100 kg/m3, a mantle density of 3400 kg/m3, and a volcanic arc density of 2700 kg/m3. These densities, together with the observed basin width of 70 ± 10 km, subsidence of 3300 ± 500 m, and β = 0.07 ± 0.02 (i.e., volcano slopes of 8 ± 2°), give a predicted arc width of 23 ± 5 km for a broken-plate model, in good agreement with the observed volcano diameters. The continuous plate model predicts an arc width of 28 ± 7 km. Hence, the observed subsidence and arc width are compatible with basin formation by arc loading and suggest that a broken-plate model fits the data better than a continuous-plate model of lithospheric flexure. The parameters used, and the results obtained, for the Halmahera example are summarized in Table 1. In our second test case we do not, unfortunately, have good constraints upon basin subsidence, but, on the other hand, we do have a gravity profile that can be directly compared with a simulated profile based upon our mathematical model of arcloaded flexure. In East Java the volcanoes are much larger than those on Halmahera. Several of the present-day volcanoes have elevations >3000 m, and all except two of East Java’s modern volcanoes are between 1600 and 3300 m high. The volcanoes typically have diameters of 50–55 km at their base. The modern arc began activity ~10–12 m.y. ago. As in Halmahera, the size of the volcanic centers of the Southern Mountains Arc is more difficult to assess. They have a similar distribution and spacing to the modern volcanoes. Several of the centers can be mapped up to 40 km north of the coast, and the present steep coastline is not the southern limit of the volcanoes. Thus, an arc width of 50 km, similar to the diameter of the modern volcanoes, is reasonable and consistent with an interval of arc activity of ~23 m.y. The basin width can be determined with greater certainty; it is between 100 and 120 km from the arc margin to the shelf edge, and there is estimated to have been 10–30 km of shortening on thrusts at the shelf edge.

APPLICATION TO INDONESIAN ARCS The modern Halmahera volcanoes have elevations up to 1730 m and diameters at sea level of their base of between 10 and 25 km, typically ~20 km for those on land. Most of the active volcanoes with smaller diameters are offshore volcanic islands (Makian, Tidore, and Ternate) rising from sea-floor depths of several hundred meters, and therefore their diameter at the base is not known. A diameter of 20 km for all the volcanoes at their base is reasonable. The volcanoes of the Quaternary arc have been active for no more than 2 m.y. and probably less. The dimensions of the Neogene volcanic centers are not known precisely, partly because of the nature of the exposure in rainforest terrain, and partly because they are overthrust by backarc and forearc basin rocks. The distribution of their products identified during mapping of the islands at 1:250,000 scale suggests that they were larger than the Quaternary volcanoes, consistent with their longer period of activity from 8 to 3 Ma. A width of the arc of 20 km is therefore a reasonable value. Like the Neogene arc, the exact width of the Halmahera backarc basin is uncertain. Its eastern edge is in Weda Bay at water depths of 2 km, and detailed seismic lines cover only the western side of the offshore basin.

TABLE 1. PARAMETERS USED IN THE FLEXURAL MODEL FOR HALMAHERA Parameter Minimum Maximum Comments 3 Basin fill density (kg/m ) 2400 2200 Minimum density gives maximum arc width 3 Volcanic arc density (kg/m ) 2700 2700 Andesite Mantle density (kg/m3) 3400 3400 Standard Basin width (km) 60 80 Gives estimated elastic thickness of 4.5–7.1 km Subsidence (m) 3250 3350 See text Slope (°) 10 6 Minimum slope gives maximum arc width Arc width (km) Arc width (km) Arc width (km)

10 17 21

21

25 28 34

O bs e r v e d Predicted broken-plate model Predicted continuous-plate model

22

Waltham et al.

Independent estimates of basin thickness are unavailable for this area, but gravity data (see Fig. 3) can be used to test the model. The Bouguer anomaly data along an East Java profile is shown in Figure 8. Figure 9 shows the same data after a regional trend was removed owing to the gravity signature of a subducting lithospheric plate deep beneath Java. The gravity modeling was performed by an add-on to our main arc-load modeling program, which simply summed contributions from a large number of small, rectangular, constant-density slabs using well-known expressions (e.g., see Telford et al., 1990). The small density

A

1400

contrast between the mantle lithosphere and the asthenosphere was ignored, and a single density was assumed for the entire mantle. A forward model of gravity over an arc-loaded, flexing, broken elastic lithosphere is also shown in Figure 9. Flexure and gravity model parameters are given in Table 2. Note that β = 0.05 implies a modern volcano height of 2500 m and a slope of 6°. Using the parameters listed in Table 2, the mathematical model predicts a subsidence of 6.9 km beneath the arc. The key features of Figure 9 are the gravity low owing to the basin depocenter, but also the gravity high at the south end

S

N

S

N

1200

Bouguer Anomaly (gu)

1000 800 600 400 200 0 -200 -400 -600

B

-800 900 800

Topography (m)

700 600 500 400 300 200 100 0

50

100

150

200

250

Distance (km) Figure 8. Bouguer anomaly (A) and topographic profile (B) for East Java for the line shown in Figure 5.

300

Basin formation by volcanic arc loading 600

23

N

S

Bouguer Anomaly (gu)

400 200 0 -200 -400 -600 -800 -1000 -1200 50

100

150

200

250

300

350

Distance (km) Figure 9. Bouguer anomaly from Figure 8 after removal of regional trend. The solid line is the predicted gravity anomaly based on a filled basin produced by subsidence caused by volcanic arc loading of a broken elastic plate.

TABLE 2. PARAMETERS USED IN THE GRAVITY AND FLEXURAL MODELING OF THE EAST JAVA BASIN Parameters Units 3 Basin fill density (kg/m ) 2200 3 2900 Volcanic arc density (kg/m ) 3 Crustal density (kg/m ) 2700 3 40 Crustal thickness (kg/m ) Mantle density (kg/m3) 3400 Basin width (km) 150 Arc width (km) 50 Slope (°) 6 Subsidence pred icted (m) 6900

of the profile between 50 and 150 km, which is due to the positive effect of the volcanic arc itself. The remaining discrepancies between modeled and observed gravity profiles occur at distinct topographic highs and would be removed by suitable topographic corrections. The basin-fill density is at the low end of figures given by Hamilton (1979). The crustal thickness of 40 km is possibly too high, although Hamilton (1979) shows a speculative thickness beneath the Java magmatic arc of >30 km, and the value is well within the plausible range for a mature arc built on continental crust (e.g., Kay and Mpodozis, 2001). However, the arc density has been set rather high, and this may reflect the need for further loads to produce the observed gravity high near the arc. DISCUSSION The key prediction of the mathematical model of basin formation by volcanic arc loading, as presented here, is that there

is a simple relationship between arc width at the surface and the amount of subsidence. This model is strongly supported by analysis of the Halmahera Arc, whose width agrees closely with that predicted by our model from the observed subsidence. Our model is further supported by the strong similarity between the observed Bouguer anomaly on East Java to that predicted by a model of basin subsidence above a broken plate. The gravity profiles agree closely in both the depth and width of the anomaly produced by the basin. For Halmahera, the subsidence observed can be accounted for entirely by arc loading if a broken-plate model is assumed. This is consistent with weakening of the arc crust by magmatic heating. For East Java, the modeling suggests that an additional load is required to produce the observed subsidence. One obvious contributor could be crustal extension, and we cannot rule out extensional faulting, as the Kendeng Basin is poorly exposed and there are no seismic lines crossing the basin. However, the large volumes of highly acidic volcanic products that erupted in East Java in the Southern Mountains Arc suggest significant differentiation of arc magmas before eruption, which should have produced large volumes of basic cumulates (S. Sparks, 2005, personal commun.). The relatively high density used in the model would be consistent with large volumes of dense cumulates at deep crustal levels. Magmatic underplating by such cumulates would have provided an additional load as in volcanic islands (Watts et al., 1985). Probably the biggest uncertainty in the modeling and interpretation is the density contrast between crust and arc. Little is

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Waltham et al.

known about the crustal structure and thickness of any Indonesian arc. The model assumes a normal continental crust for the arc basement with a load similar in composition to andesites. This is reasonable in East Java, where the basement is continental crust and preserved volcanic centers are predominantly andesites. It is also likely, based on observations of the character of small areas of exposed basement and oil-company drilling of the basement, that accreted arc and ophiolitic material includes serpentinites, which would lower the average crustal density. For Halmahera, Milsom et al. (1996) concluded that the crust was at least 20 km thick and had a bulk density approaching that of continental rather than oceanic crust. Local gravity highs are associated with Paleogene arc volcanic rocks rather than the ophiolites, and nowhere do the ophiolites have the high gravity fields associated with classic ophiolites. Although much of Halmahera is underlain by ophiolitic rocks, significant serpentinization of the ultramafic parts could account for the low crustal density. The products of the modern arc (Morris et al., 1983) and the Neogene arc (Forde, 1997; Macpherson et al., 2003) are almost entirely basaltic andesites and andesites.

volcanoes and the underlying crust, and therefore the volcanoes do not act as a load. In continental margin arcs such as Java, and long-lived intraoceanic arcs such as Halmahera, there is a larger density contrast between the deeper crust and the eruptive products of the arc. In these cases the volcanoes act as a significant load and produce flexural basins close to the arc. The absence of deep basins in arcs in some parts of Indonesia and elsewhere in the world may be due to the small density contrast between volcanoes and the underlying crust in those regions. APPENDIX Equation 6 is 1 = Aw + 2Bww′.

(A1)

Changing the variable to x = A2s/B

(A2)

f = Aw

(A3)

1 = f + 2 f f ′,

(A4)

and scaling using CONCLUSIONS The role of volcanoes in the development of basins in arc regions has generally been ignored, overlooked, or not considered important. This is possibly because most studies have been concerned with larger true backarc basins and forearc basins both of which are formed much farther from the arc. However, in several Indonesian arcs there are thick sequences of sedimentary and volcaniclastic rocks in deep basins very close to the arc. The mathematical modeling shows that volcanic loading can make a contribution to basin subsidence in these settings, and may be the primary cause. Volcanic loading can account for basins close to the arc but is not relevant to basins >~100 km from the arc. The basins produced by arc loading are asymmetrical, and basin sequences are thickest close to the arc. These are likely to be dominated by very coarse, terrestrial and shallow marine, primary volcanic and volcaniclastic rocks that may include a variety of mass-flow deposits. Locally and intermittently the supply of material from the arc may exceed subsidence, and there may be rapid vertical and lateral variations in grain size and changes from shallow marine to terrestrial deposits. Farther from the arc, subsidence typically exceeds supply, and the more proximal deposits pass laterally into deeper water sedimentary deposits such as turbidites. The distribution and character of the material, particularly close to the arc, has some similarities to rocks described by Draut and Clift (2006) from the Jurassic Talkeetna Formation of Alaska. However, in the Alaskan Jurassic arc and in the modern Mariana and Tonga Arcs, accommodation space is not likely to have been created by volcanic loading but had already existed. This is because in young intraoceanic settings the volcanoes rise from great depths above the ocean floor. In young intraoceanic arcs there is also likely to be little density difference between the arc

then gives

where f ′ now indicates differentiation with respect to x. Equation A4 is a nonlinear ordinary differential equation (ode), which does not have an analytical solution in terms of elementary functions. However, at small values of f the second term on the right side dominates, so that 1 ~2 f f ′ with a solution of f = x0.5. At large values of f the first term dominates, and the solution asymptotically approaches f = 1. The function f can be estimated using finite differencing with f ~(fi + fi+1)/2

(A5)

f ′ ~(fi+1 – fi)/Δx,

(A6)

and

which, after substitution into (A4), yields 2

f i+1

Δx − Δx ⎛ Δx ⎞ f i + Δx , = + ⎜ ⎟ + fi 2 − 4 2 ⎝ 4 ⎠

(A7)

which, together with the boundary condition that f0 = 0, allows the function in Figure 2 to be calculated. The resultant function has the required properties that f ~x0.5 for small x and f ~1 for large x.

Basin formation by volcanic arc loading ACKNOWLEDGMENTS Our work on Indonesian arcs has been supported at different times by the University of London Central Research Fund, NERC, and the Royal Society, but mainly by a consortium of oil companies whose membership has changed with time. We thank colleagues in Indonesia at the Geological Research and Development Center Bandung (now the Geological Survey of Indonesia), LIPI, and Institut Teknologi Bandung for help with field work. We thank Chris Elders and Andi Salahuddin for discussion of Weda Bay seismic interpretation.

REFERENCES CITED Ali, J.R., Hall, R., and Baker, S.J., 2001, Palaeomagnetic data from a Mesozoic Philippine Sea Plate ophiolite on Obi Island, Eastern Indonesia: Journal of Asian Earth Sciences, v. 19, p. 535–546, doi: 10.1016/S13679120(00)00053-5. Bahlburg, H., and Furlong, K.P., 1996, Lithospheric modeling of the Ordovician foreland basin in the Puna of northwestern Argentina: Tectonophysics, v. 259, p. 245–258, doi: 10.1016/0040-1951(95)00129-8. Ballantyne, P.D., 1992, Petrology and geochemistry of the plutonic rocks of the Halmahera ophiolite, eastern Indonesia, an analogue of modern oceanic forearcs, in Parson, L.M., et al., eds., Ophiolites and Their Modern Oceanic Analogues: Geological Society [London] Special Publication 60, p. 179–202. Busby, C.J., and Ingersoll, R.V., 1995, Tectonics of Sedimentary Basins: Cambridge, Massachusetts, Blackwell, 579 p. Cole, J.M., and Crittenden, S., 1997, Early Tertiary basin formation and the development of lacustrine and quasi-lacustrine/marine source rocks on the Sunda Shelf of SE Asia, in Fraser, A.J., et al., eds., Petroleum Geology of SE Asia: Geological Society [London] Special Publication 126, p. 147–183. DeCelles, P.G., and Giles, K.A., 1996, Foreland basin systems: Basin Research, v. 8, p. 105–123, doi: 10.1046/j.1365-2117.1996.01491.x. de Genevraye, P., and Samuel, L., 1972, The geology of Kendeng Zone (East Java): Proceedings of Indonesian Petroleum Association 1st Annual Convention, Jakarta, p. 17–30. Dewey, J.F., 1980, Episodicity, sequence and style at convergent plate margins, in Strangeway, D.W., ed., The Continental Crust and Its Mineral Deposits: Geological Association of Canada Special Paper 20, p. 553–573. Draut, A.E., and Clift, P.D., 2006, Sedimentary processes in modern and ancient oceanic arc settings: Evidence from the Jurassic Talkeetna Formation of Alaska and the Mariana and Tonga Arcs, Western Pacific: Journal of Sedimentary Research, v. 76, p. 493–514, doi: 10.2110/jsr.2006.044. Forde, E.J., 1997, The geochemistry of the Neogene Halmahera Arc, eastern Indonesia [Ph.D. thesis]: University of London, 268 p. Hall, R., 1987, Plate boundary evolution in the Halmahera region, Indonesia: Tectonophysics, v. 144, p. 337–352, doi: 10.1016/0040-1951(87)90301-5. Hall, R., 2002, Cenozoic geological and plate tectonic evolution of SE Asia and the SW Pacific: Computer-based reconstructions, model and animations: Journal of Asian Earth Sciences, v. 20, p. 353–434, doi: 10.1016/S13679120(01)00069-4. Hall, R., and Morley, C.K., 2004, Sundaland Basins, in Clift, P., et al., eds., Continent-Ocean Interactions within the East Asian Marginal Seas: American Geophysical Union Geophysical Monograph 149, p. 55–85. Hall, R., Audley-Charles, M.G., Banner, F.T., Hidayat, S., and Tobing, S.L., 1988a, Basement rocks of the Halmahera region, Eastern Indonesia: A Late Cretaceous–Early Tertiary arc and fore-arc: Geological Society [London] Journal, v. 145, p. 65–84. Hall, R., Audley-Charles, M.G., Banner, F.T., Hidayat, S., and Tobing, S.L., 1988b, Late Paleogene–Quaternary geology of Halmahera, Eastern Indonesia: Initiation of a volcanic island arc: Geological Society [London] Journal, v. 145, p. 577–590.

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Hall, R., Ali, J.R., Anderson, C.D., and Baker, S.J., 1995, Origin and motion history of the Philippine Sea Plate: Tectonophysics, v. 251, p. 229–250, doi:10.1016/0040-1951(95)00038-0. Hamilton, W., 1979, Tectonics of the Indonesian region: U.S. Geological Survey Professional Paper 1078, 345 p. Karner, G.D., and Watts, A.B., 1983, Gravity anomalies and flexure of the lithosphere at mountain ranges: Journal of Geophysical Research, v. 88, p. 10,449–10,477. Kay, S.M., and Mpodozis, C., 2001, Central Andean ore deposits linked to evolving shallow subduction systems and thickening crust: GSA Today, v. 11, no. 3, p. 4–9, doi: 10.1130/1052-5173(2001)011<0004: CAODLT>2.0.CO;2. Kopp, H., Flueh, E.R., Petersen, C.J., Weinrebe, W., Wittwer, A., and Meramex Scientists, 2006, The Java margin revisited: Evidence for subduction erosion off Java: Earth and Planetary Science Letters, v. 242, p. 130–142, doi: 10.1016/j.epsl.2005.11.036. Macpherson, C.G., Forde, E.J., Hall, R., and Thirlwall, M.F., 2003, Geochemical evolution of magmatism in an arc-arc collision: The Halmahera and Sangihe arcs, eastern Indonesia, in Larter, R.D., and Leat, P.T., eds., Intraoceanic Subduction Systems: Tectonic and Magmatic Processes: Geological Society [London] Special Publication 219, p. 207–220. McCaffrey, R., Silver, E.A., and Raitt, R.W., 1980, Crustal structure of the Molucca Sea collision zone, Indonesia, in Hayes, D.E., ed., The Tectonic and Geologic Evolution of Southeast Asian Seas and Islands, Part 1: American Geophysical Union Geophysical Monograph 23, p. 161–178. Milsom, J., Hall, R., and Padmawidjaja, T., 1996, Gravity fields in eastern Halmahera and the Bonin Arc: Implications for ophiolite origin and emplacement: Tectonics, v. 15, p. 84–93, doi: 10.1029/95TC02353. Morris, J.D., Jezek, P.A., Hart, S.R., and Gill, J.B., 1983, The Halmahera island arc, Molucca Sea collision zone, Indonesia: A geochemical survey, in Hayes, D.E., ed., The Tectonic and Geologic Evolution of Southeast Asian Seas and Islands, Part 2: American Geophysical Union Geophysical Monograph 27, p. 373–387. Moss, S.J., and McCarthy, A.J., 1997, Discussion—Foreland basin systems: Basin Research, v. 9, p. 171–176. Nichols, G.J., and Hall, R., 1991, Basin formation and Neogene sedimentation in a backarc setting, Halmahera, eastern Indonesia: Marine and Petroleum Geology, v. 8, p. 50–61, doi: 10.1016/0264-8172(91)90044-2. Nunn, J.A., Czerniak, M., and Pilger, R.H., Jr., 1987, Constraints on the structure of Brooks Range and Colville Basin, Northern Alaska, from flexure and gravity analysis: Tectonics, v. 6, p. 603–617. Pertamina, B.P.P.K.A., 1996, Petroleum geology of Indonesian basins, principles, methods and application. Volume IV: East Java Basins, 107 p. Slingerland, R., Harbaugh, J.W., and Furlong, K., 1994, Simulating Clastic Sedimentary Basins: New York, Prentice Hall, 220 p. Smith, G.A., Moore, J.D., and McIntosh, W.C., 2002, Assessing roles of volcanism and basin subsidence in causing Oligocene–Lower Miocene sedimentation in the northern Rio Grande Rift, New Mexico, USA: Journal of Sedimentary Research, v. 72, p. 836–848. Smyth, H.R., 2005, Eocene to Miocene basin history and volcanic history in East Java, Indonesia [Ph.D. thesis]: University of London, 476 p. Smyth, H.R., Hall, R., and Nichols, G.J., 2008, Early Cenozoic volcanic arc history of East Java, Indonesia: The stratigraphic record of eruptions on a continental margin in a tropical setting, in Draut, A.E., et al., eds., Lessons from the Stratigraphic Record in Arc Collision Zones: Geological Society of America Special Publication 436 (this volume). Soeria-Atmadja, R., Maury, R.C., Bellon, H., Pringgoprawiro, H., Polvé, M., and Priadi, B., 1994, Tertiary magmatic belts in Java: Journal of Southeast Asian Earth Sciences, v. 9, p. 13–17, doi: 10.1016/0743-9547(94)90062-0. Telford, W.M., Geldart, L.P., and Sheriff, R.E., 1990, Applied Geophysics: Cambridge, UK, Cambridge University Press, 770 p. Toksöz, M.N., and Bird, P., 1977, Formation and evolution of marginal basins and continental plateaus, in Talwani, M., and Pitman, W.C., eds., Island Arcs, Deep Sea Trenches and Back Arc Basins: American Geophysical Union Maurice Ewing Series 1, p. 379–393. Turcotte, D.L., and Schubert, G., 2002, Geodynamics (2nd edition): Cambridge, UK, Cambridge University Press, 456 p. Untung, M., and Sato, Y., 1978, Gravity and geological studies in Java, Indonesia: Geological Survey of Indonesia and Geological Survey of Japan, Special Publication 6, 207 p.

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van Bemmelen, R.W., 1949, The Geology of Indonesia: Nijhoff, The Hague, Government Printing Office, 732 p. Watts, A.B., and Cochran, J.R., 1974, Gravity anomalies and flexure of the lithosphere along the Hawaiian-Emperor seamount chain: Royal Astronomical Society Geophysical Journal, v. 38, p. 119–141. Watts, A.B., ten Brink, U.S., Buhl, P., and Brocher, T., 1985, A multichannel seismic study of lithospheric flexure across the Hawaiian-Emperor seamount chain: Nature, v. 315, p. 105–111, doi:10.1038/315105a0.

Watts, A.B., Peirce, C., Collier, J., Dalwood, R., Canales, J.P., and Henstock, T.J., 1997, A seismic study of lithospheric flexure in the vicinity of Tenerife, Canary Islands: Earth and Planetary Science Letters, v. 146, p. 431– 447, doi: 10.1016/S0012-821X(96)00249-X.

MANUSCRIPT ACCEPTED BY THE SOCIETY 24 APRIL 2007

Printed in the USA

The Geological Society of America Special Paper 436 2008

Cenozoic arc processes in Indonesia: Identification of the key influences on the stratigraphic record in active volcanic arcs Robert Hall Helen R. Smyth SE Asia Research Group, Department of Geology, Royal Holloway University of London, Egham, Surrey, TW20 0EX, UK

ABSTRACT The Indonesian region includes several volcanic island arcs that are highly active at the present day, and also contains a record of Cenozoic volcanic activity owing to subduction of oceanic lithosphere at the margins of SE Asia. As a result of long-term subduction, there is a high regional heat flow, and a weak crust and lithosphere, as identified in other subduction zone backarcs. The stratigraphic record in the Indonesian region reflects a complex tectonic history, including collisions, changing plate boundaries, subduction polarity reversals, elimination of volcanic arcs, and extension. The arcs have not behaved as often portrayed in many arc models. They mark subduction but were not continuously active, and it is possible to have subduction without magmatism. Subduction hinge retreat was accompanied by significant arc volcanism, whereas periods of hinge advance were marked by reduction or cessation of volcanic activity. Growth of the region occurred in an episodic way, by the addition of ophiolites and continental slivers, and as a result of arc magmatism. In Indonesia, relatively small amounts of material were accreted from the downgoing plate during subduction, but there is also little evidence for subduction erosion. During collision the arc region may fail, resulting in thrusting, and the weakest point is the position of the active volcanic arc itself. Volcanic arcs shift position suddenly, and arcs can disappear during collision by overthrusting. Arcs are geologically ephemeral features and may have very short histories in comparison with most well-known older orogenic belts. The stratigraphic record of the basins within arc regions is complex. Because of a weak lithosphere the character of sedimentary basins may be unusual, and basins are commonly very deep and subside rapidly. There is a high sediment flux. The volcanic arc itself influences the stratigraphic record and basin development. The load imposed by the volcanic arc causes flexure and provides accommodation space. The volcanic arc thus can form the basin and supply most of its sediment. Tropical processes influence the mineralogy and apparent maturity of the sediment, especially volcanogenic material. A complex stratigraphy will result from the waxing and waning of volcanic activity. Keywords: Cenozoic, Sunda Arc, Banda Arc, Sangihe Arc, Halmahera Arc, Indonesia, stratigraphic record. Hall, R., and Smyth, H.R., 2008, Cenozoic arc processes in Indonesia: Identification of the key influences on the stratigraphic record in active volcanic arcs, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 27–54, doi: 10.1130/2008.2436(03). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Hall and Smyth

INTRODUCTION The size of Indonesia, the abundance of volcanic activity, and its long volcanic history make it an important area for understanding arc processes and influences of volcanic arcs on the stratigraphic record. Hamilton (1988) commented that the arcs of Indonesia–southern Philippines–western Melanesia are in “the region of greatest modern variety and complexity,” but the region remains understudied and often overlooked. The long history of subduction beneath Indonesia has influenced arc development. Subduction has produced a thin and warm lithosphere beneath the upper plate, which is unlike that of older stable continents and their margins. These features have influenced the development of the Indonesian arcs, which differ in many ways from textbook examples of arcs described from other parts of the world. They were formed at the margins of a large, weak subduction backarc region and consequently seem to have been unusually responsive to plate boundary forces. Because of their youth it is not possible to study the deep levels of the arcs, but the arc history that is revealed by their stratigraphy offers insights into the development of older orogens and older volcanic arcs. In this paper we identify some important features of Indonesian volcanic arcs, particularly those that influence the stratigraphic record of arc activity. Because the geology of Indonesia is not familiar to many readers, we begin our discussion with some

10°N

Andaman Sea

background on the region, and a summary of the history of several of the most important Cenozoic arcs in Indonesia. We then discuss particular features of the record of arc activity and attempt to identify those that may be of general interest in understanding arc processes and those that may be useful in interpreting the record of activity in older arcs elsewhere in the world. Our discussion of the stratigraphic record in Indonesian arcs is based mainly on studies carried out by the SE Asia Research Group in several different arcs over many years, especially in Java and Halmahera, but also in arcs of Sumatra, Sulawesi, and Borneo. These studies have included extensive field work and related work documented in theses, but some of the results are unpublished and in places we may appear to make statements unsupported by easily accessible literature; we hope this will be forgiven, and we felt it was justified in our attempt to give an account of the region in a paper of reasonable length. Indonesia and Large Explosive Volcanic Events Indonesia is a large country that includes >18,000 islands and stretches >5000 km from west to east (Fig. 1). Wallace (1869) tried to draw attention to its size in his book The Malay Archipelago by comparing the island of Borneo to the British Isles, and subsequently Umbgrove (1938) and van Bemmelen (1949) compared the region to areas more familiar to North American

South China Sea Philippines

5°N

Malay Peninsula

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Sangihe Celebes Sea Borneo

Nias

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ts

0°

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Str ai sar Ma kas

Indian Ocean

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Banda Sea

Flores Sea

Timor

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0

500

1,000

Northwest Shelf Australia

kilometers 100°E

110°E

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Figure 1. Indonesia shown on a digital elevation model (DEM), based on Shuttle Radar Topography Mission (SRTM) data merged with the Sandwell and Smith (1997) bathymetry.

Cenozoic arc processes in Indonesia and European readers. Figure 2 is a modern attempt to show the size of the region by comparing Indonesia to the conterminous United States. According to the Smithsonian Global Volcanism Program (Smithsonian, 2006), Indonesia includes 150 Holocene volcanoes; 95 of these could be considered active because they have erupted since A.D. 1500. Indonesia has a long record of volcanic activity, recording subduction related to northward movement of the Indian-Australian plate and westward motion of the Pacific plates. Cenozoic subduction of Indian-Australian oceanic lithosphere beneath most of Indonesia began after Australia separated from Antarctica and moved rapidly northward from ~45 m.y. ago (Hall, 2002). With a long history of volcanic activity, and a well-known explosive character in historical times, it is surprising that the only exceptionally large eruptions known from Indonesia in the Cenozoic are those of Toba. Of the Holocene volcanoes, 32 have records of very large eruptions with a volcanic explosive index (VEI) >4; 19 of these have erupted in the past 200 yr, including Tambora in 1815 (VEI = 7) and Krakatau in 1883 (VEI = 6). There are only 4 volcanoes in the Smithsonian global data set of large-volume Holocene explosive eruptions with a VEI of 7; the other 3 erupted in prehistorical times and pre-date 4000 B.C. Tambora, on the island of Sumbawa, is known for its impact on global climate, and its 1815 eruption resulted in the Northern Hemisphere’s “year without a summer” in 1816, when crops failed, causing famine and population movements (Harington, 1992; Zeilinga de Boer and Sanders, 2002). The 74 ka eruption of Toba (Chesner and Rose, 1991; Chesner et al., 1991), on the

29

island of Sumatra, was the largest in the last 2 m.y., and its effect on climate must have been even more devastating. It has been speculated that the eruption led to accelerated glaciation and a bottleneck in human evolution (Ambrose, 2003; Rampino and Self, 1992, 1993a, b). For the last 36 m.y., 42 large explosive events are recorded in a recent global compilation (Mason et al., 2004); 33 of these large explosive events are recorded from North America, whereas Toba alone is situated in a humid tropical climate. The record has been interpreted to indicate that there were two pulses in Cenozoic global volcanism: one between 36 and 25 Ma, and the other since 13.5 Ma. However, Indonesia remains a relatively remote area, which is difficult to explore because of the absence of infrastructure, difficult terrain, and rain forest cover. We suggest that the absence of exceptionally large eruptions in the stratigraphic record from Indonesia has more to do with preservation and sampling than with volcanic activity. As we report below, there is indeed evidence of large explosive events in Indonesia during the Cenozoic, and no doubt more of these remain to be discovered. The lack of information on large eruptions in Indonesia is not dissimilar to the general lack of information on Indonesia’s volcanic history. This partly reflects the absence of modern detailed studies, but ideas about arcs may be unduly influenced by results from more accessible or better exposed regions, from ocean drilling, and from geochemical studies. However, as with large eruptions the stratigraphic record in Indonesia has some valuable insights to reveal; this paper identifies some of them.

5000 km Figure 2. Indonesia and Malaysia compared to the USA. The box is 60° from east to west and 25° from north to south.

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Hall and Smyth

ARC ACTIVITY IN INDONESIA

active since the Eocene (e.g., Garwin et al., 2005; Hall, 2002; Hamilton, 1977) and is the product of subduction of the IndianAustralian plate beneath the Sundaland margin. Sundaland (Fig. 4) is the continental core of SE Asia, comprising Indochina, the Thai-Malay Peninsula, Sumatra, Java, Borneo, and the shallow marine shelf (Sunda Shelf) between them, and formed by amalgamation of continental blocks during the Triassic. There is a Proterozoic basement (Liew and McCulloch, 1985; Liew and Page, 1985) intruded by Permian–Triassic granites formed by subduction and postcollisional thickening of the continental crust (Hutchison, 1989, 1996). In Borneo the oldest rocks known are Paleozoic metamorphic rocks intruded by Mesozoic granites. From Sumatra to Borneo the continental core is surrounded by Mesozoic ophiolitic and arc igneous rocks, and possible fragments of continental material, accreted during the Cretaceous. In Sumatra and West Java the Sunda Arc is built in part on the continental margin of Sundaland. In East Java it is constructed on arc and ophiolitic rocks accreted at the Mesozoic active margin and in part on a continental fragment that was added by a Cretaceous collision. In Sumatra a record of subduction-related magmatism extends back into the early Mesozoic or late Paleozoic (McCourt et al., 1996), but in Java the oldest record of subduction is Cretaceous. The Sundaland Cretaceous active margin is interpreted to have run the length of Sumatra into West Java and then turned northeast into SE Borneo (Hamilton, 1979). In Central Java and SE Borneo there is evidence for Cretaceous high pressure–low temperature subduction-related metamorphism (Parkinson et

The distribution of modern-day volcanoes in Indonesia shows clearly that most volcanic activity is related to subduction (Fig. 3). This is also emphasized by the distribution of seismicity (Cardwell and Isacks, 1978; Cardwell et al., 1980; Engdahl et al., 1998; England et al., 2004). Global positioning system (GPS) observations (Bock et al., 2003; Kreemer et al., 2000; Rangin et al., 1999), seismicity, and volcanic activity indicate rapid plate movements and considerable tectonic complexity. Present-day Indonesian volcanic activity can be regarded as the products of four separate volcanic arcs (Fig. 3): the Sunda Arc, stretching from Sumatra through Java to the east, the horseshoeshaped Banda Arc of eastern Indonesia, the Halmahera Arc of the North Moluccas, and the Sangihe Arc, extending from Sulawesi into the southern Philippines. Here we give first a brief summary of the Cenozoic history of each of these arcs and the character of the crust on which they are constructed to provide the background to our commentary on Java, east Indonesia, and the Moluccas. More detailed discussion of each of these arcs, literature references, and animated reconstructions of the tectonic development are in Hall (2002) and Garwin et al. (2005). Java: The Sunda Arc Indonesia is situated on the southern margin of the Eurasian plate at the edge of Sundaland (van Bemmelen, 1949; Hamilton, 1979; Hutchison, 1989; Katili, 1975). The Sunda Arc has been

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Seram

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95°E

Obi

ait

ra

Sulawesi

115°E

Timor 120°E

125°E

130°E

Figure 3. Present-day volcanoes in Indonesia (Smithsonian, 2006) and the principal volcanic arcs discussed in the text: Sunda Arc, Banda Arc, Halmahera Arc, and Sangihe Arc. Indonesia is shaded in gray, and the area of Malaysia is unfilled.

Manila Trench

Cenozoic arc processes in Indonesia

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pine

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nc Tre da Sun

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Molucca Suture



Celebes Sea

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10°N

g Trou

Sahul Shelf Coral Sea

AUSTRALIA 130°E

140°E

20°S

Figure 4. Wide plate boundary zone of Indonesia. The light shaded area is the zone of collision between the Eurasian, Indian-Australian, and Pacific–Philippine Sea plates, which can be considered as a wide suture zone, within which there are smaller sutures shown in darker shading. Lines show incremental growth of the Indonesian region at different stages by addition of fragments of continental, arc, and ophiolitic rocks. Modified from Hall and Wilson (2000).

al., 1998). In Indonesia there is almost no Paleocene record, and it appears that most of Sundaland was emergent. All published plate tectonic reconstructions show subduction beneath Java before 45 Ma, but there is no record of volcanic activity in Java in the early Cenozoic (Smyth et al., this volume). We suggest there was a cessation of subduction beneath Java in the Late Cretaceous following the Cretaceous collision of a continental fragment at the Sundaland margin (Smyth et al., this volume). Our recent and current work in Java suggests that subduction resumed in the middle Eocene, forming a volcanic arc in the Southern Mountains that ran the length of Java. At this time, ca. 45 Ma (Fig. 5), the rate of Australia-Antarctica separation increased, and Australia began to move northward relatively rapidly. Since then there has been continuous subduction of Indian Ocean lithosphere beneath the Sunda Arc. During the Eocene and Oligocene the position of the volcanic arc remained broadly fixed, and from Sumatra to Sulawesi abundant volcanic activity accompanied northward subduction of the Indian-Australian

plate. In the Java-Sulawesi sector of the Sunda Arc, volcanism greatly diminished during the early and middle Miocene, although northward subduction of Indian-Australian lithosphere continued (Fig. 5). Following Australian continental collision in east Indonesia, the Sunda subduction hinge advanced northward as a result of counterclockwise rotation of the Borneo-Java part of the Sundaland margin, and this is interpreted to have led to termination of magmatism, despite continued subduction, because of the absence of fresh material replenishing the mantle wedge (Macpherson and Hall, 1999, 2002). At the end of the middle Miocene at ca. 10 Ma, volcanic activity resumed in the Java-Sulawesi sector of the Sunda Arc, after the termination of Borneo-Java rotation. Magmatism began again in Java with increased vigor, but in a position north of the Paleogene arc, forming the modern arc (Fig. 5). During most of the Cenozoic since the Eocene, Java was a chain of volcanic islands, and only in the last 5 m.y. has the island emerged as an extensive area of land.

tra

ma

Su

nd

va

25 Ma

120°E

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pin es –

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t P hil ip

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ah er a

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Ha

West Philippine Basin

Izu

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as

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rian

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a

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tr ma Su

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Bismarck Sea

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Halmahera

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Cenozoic arc processes in Indonesia

33

Figure 5. Reconstructions of the Indonesian region, based on Hall (2002). 45 Ma: The Pacific plate was moving broadly NW and subducting beneath the east Asian margin and the eastern margin of the Philippine Sea plate, forming the Izu-Bonin-Marianas Arc. Spreading centers linked the West Philippine Sea Basin and the Celebes Sea. Australia was moving NNE on the Indian-Australian plate, and northward subduction produced the Sunda Arc and the East Philippines–Halmahera Arc. The proto–South China Sea was subducting southward beneath the north Borneo–Luzon margin. The parts of the Pacific and Indian-Australian plates shown in blue without anomalies have been subducted since 45 Ma. 25 Ma: The East Philippines–Halmahera–South Caroline Arc collided with the Australian margin in New Guinea. Collision between Australia and SE Asia began with initial contact of the Sula Spur, the continental promontory east of the Bird’s Head of New Guinea, and East Sulawesi. These events caused major reorganization of plate boundaries. Active spreading in the South China Sea was caused by southward subduction of the proto–South China Sea. 15 Ma: Subduction continued beneath the Sunda Arc, but arc volcanic activity ceased in Java between ca. 20 and 10 Ma as the hinge advanced, owing to collision of the Australian margin in Sulawesi. Arc terranes in New Guinea moved westward in a wide left-lateral strikeslip zone. Locking of splays at the western end of the fault zone induced eastward subduction of the Molucca Sea; westward subduction of the Molucca Sea beneath the Sangihe Arc had begun at ca. 25 Ma. 5 Ma: Volcanic activity resumed in Java at ca. 10 Ma and has continued until the present. Hinge rollback into the Banda Embayment south of the Bird’s Head produced the arcuate subducted slab beneath the Banda Arc and led to backarc spreading in the Banda Sea. The Molucca Sea was in the process of elimination by subduction at its west and east sides beneath the Sangihe and Halmahera Arcs.

Eastern Indonesia: The Banda Arc East of Java and Borneo, Indonesia can be considered to be a wide and complex suture zone (Fig. 4), and even today it can be tectonically described only in terms of several small plates and multiple subduction zones (Hall, 2002; Hall and Wilson, 2000; Hamilton, 1979). Eastern Indonesia is the product of a long period of extension, subduction, and collision and has a complex basement; there is some continental crust but also much more arc and ophiolitic basement in comparison with western Indonesia. The Banda Arc is the horseshoe-shaped arc, which today extends east from Flores to Buru (Fig. 3), passing through Timor and Seram and includes both an outer nonvolcanic arc and an inner volcanic arc. Although there was a volcanic arc in eastern Indonesia during the Paleogene, and parts of this and possibly older arcs may be found in the highest nappes of Timor and other Banda islands, we do not term them the Banda Arc. The older volcanic arc rocks formed part of the Paleogene Sunda Arc. The Banda volcanic arc is young and has been active only for ~10 m.y. (Abbott and Chamalaun, 1981; Honthaas et al., 1999), and the horseshoe-shaped Banda Arc is a phenomenon of the same time interval. It formed by subduction of an embayment within the northward-moving Australian plate (Charlton, 2000; Hall, 1996, 2002; Hamilton, 1979). The arc developed within the collision zone after the Australian margin collided with the former active margin of the Sunda Arc. From the Eocene the Sunda Arc can be traced east through Sulawesi into the East Philippines and Halmahera Arcs in the Pacific. During the Paleogene (Fig. 5) subduction was northward beneath Sulawesi and the East Philippines until the first collision of Australian crust, the Sula Spur, with SE Asia in the early Miocene ~25 m.y. ago (Fig. 5). The collision caused the long subduction system to separate into two parts. West of Sulawesi, northward subduction continued in the Sunda Arc, but to the east subduction ceased, and the Australia–Philippine Sea plate boundary became a strike-slip system. Between the two, the Australia–SE Asia collision formed a mountain belt in East Sulawesi. By the late middle Miocene, at ca. 12 Ma, convergence in East Sulawesi could no

longer be accommodated by orogenic contraction. At this time the oldest oceanic lithosphere in the Indian Ocean, of Late Jurassic age, arrived at the eastern end of the Java Trench. This area of old crust north of the NW Shelf of Australia formed an embayment in the Australian margin, the proto–Banda Sea. Because of its age and thickness, the Jurassic ocean lithosphere fell away rapidly, causing the Banda subduction hinge to roll back rapidly to the south and east, inducing massive extension in the overlying plate. In western Sulawesi this first induced extensional magmatism, which began at ca. 11 Ma (Polvé et al. 1997). As the hinge rolled back into the Banda Embayment it led to formation of the Banda Arc and the opening of the North Banda Sea, as described by Hamilton (1988). Further rollback caused the opening of the Flores Sea and later the South Banda Sea (Fig. 5). South of the Bird’s Head microcontinent the rollback of the subduction hinge resulted in collision of the Banda volcanic arc at ca. 3 Ma with the Australian margin in the region of Timor and the cessation of volcanic activity in this segment of the arc. After collision, convergence ceased south of the volcanic arc, and new plate boundaries developed north of the arc between Flores and Wetar and to the north of the South Banda Sea. Within the Bird’s Head microcontinent there has been significant shortening and probable intracontinental subduction within the last 3 m.y. at the Seram Trough. Moluccas: The Halmahera and Sangihe Arcs The Halmahera and Sangihe Arcs both have long volcanic histories. The modern Halmahera Arc is constructed on older arcs, of which the oldest known is an intraoceanic arc formed in the Pacific in the Mesozoic (Hall et al., 1988a, 1995) presumably built on older oceanic crust. The modern Sangihe Arc is constructed on arcs formed at the Pacific margin in the early Cenozoic (Hall, 2002); again the deepest parts are built on older oceanic crust (Evans et al., 1983). Both of the currently active arcs formed during the Neogene. They are unusual in that they are the only arcs in the world that are currently colliding. The pre-Eocene history of these arcs is not well known. At 45 Ma (Fig. 5) the Halmahera Arc was far out in the western

34

Hall and Smyth

Pacific but was situated on the southern margin of the Philippine Sea plate beneath which there was northward subduction of Indian-Australian lithosphere. Between 45 and 25 Ma the Philippines-Halmahera Arc remained at approximately the same latitude above a north-dipping subduction zone north of Australia. At ca. 25 Ma (Fig. 5) there was the most important Cenozoic change in plate boundaries in the region. Arc-continent collision between the East Philippines–Halmahera Arc and the New Guinea margin terminated northward subduction of oceanic lithosphere north of Australia, and a major left-lateral strike-slip boundary developed in northern New Guinea. Arc terranes were translated westward within this strike-slip system. The arc terranes at the southern edge of the Philippine Sea plate moved in a clockwise direction along the northern New Guinea margin within the leftlateral Sorong strike-slip zone. At the western end of the leftlateral fault system there was subduction beneath the Sangihe Arc, and shortening and uplift in East Sulawesi. The Philippine Sea plate moved northward as it rotated clockwise, accompanied by complex strike-slip faulting and minor subduction at its western edge in the Sangihe Arc and Philippines (Fig. 5). The Molucca Sea double subduction system was initiated at ca. 15 Ma (Fig. 5). Subduction on the west side of the Molucca Sea in the Sangihe Arc had started soon after the 25 Ma plate reorganization, but the oldest Neogene volcanic rocks in the Halmahera Arc have ages of ca. 11 Ma (Baker and Malaihollo, 1996). Initiation of east-directed Halmahera subduction probably resulted from the locking of one of the strands of the left-lateral Sorong fault zone at the southern edge of the Molucca Sea. In eastern Indonesia the

Molucca Sea has been eliminated by subduction at both its eastern and western sides (Fig. 5). The Sangihe–North Sulawesi Arc is now thrusting over the Halmahera Arc in the Molucca Sea, as discussed below. The central Molucca Sea mélange wedge and ophiolites represent the forearc basin and basement of the Sangihe Arc, which will soon have completely overridden the Halmahera Arc (Hall, 2000). RESULTS OF LONG-TERM SUBDUCTION There is a great contrast between present tectonic activity, manifested by seismicity and volcanism at the margins, and the apparent interior stability of Sundaland, which has led some authors to describe the core as a shield or craton (e.g., BenAvraham and Emery, 1973; Gobbett and Hutchison, 1973; Tjia, 1996). It is true that most of the areas of extreme relief, high elevations, and actively rising mountains are in eastern Indonesia, but the core is certainly not a craton. Most of the Sunda Shelf is flat and close to sea level (Fig. 1), and even the large island of Borneo has generally low elevations with the exception of the isolated 4 km peak of Mount Kinabalu in the north. However, the picture of Indonesia as an active volcanic margin surrounding a stable continental region is misleading. There has been significant deformation within Sundaland during the Cenozoic. Heat flow, seismic tomography, and geological observations indicate that the Sundaland continental core north of Indonesia is unusual. Sundaland has high surface-heat-flow values (Fig. 6), typically more than 80 mW/m2 (Artemieva and Mooney, 2001; Hall

Figure 6. Contoured heat-flow map for SE Asia, based on the database of Pollack et al. (1990, 1993) and oil company compilations (Kenyon and Beddoes, 1977; Rutherford and Qureshi, 1981).

Cenozoic arc processes in Indonesia and Morley, 2004). At the margins of Sundaland, high heat flow is related to subduction-related processes and magma rise, but in the interior this is not a result of arc magmatism as sometimes suggested (e.g., Nagao and Uyeda, 1995). High heat flow values are recorded in the interior Sundaland basins from the Gulf of Thailand to west Borneo >800 km from the active volcanoes of the Sunda Arc and similar distances from oceanic crust of the South China Sea. The hot interior of Sundaland appears to be the consequence of high upper-crustal heat flow from radiogenic granites and their erosional products, the insulation effects of thick sediments, and a high mantle heat flow (Hall and Morley, 2004). P and S wave seismic-tomography models (Fig. 7) show that Sundaland is an area of low velocities in the lithosphere and underlying mantle (e.g., Bijwaard et al., 1998; Lebedev and Nolet, 2003; Ritsema and van Heijst, 2000; Widiyantoro and van der Hilst, 1997) in contrast to Indian and Australian continental lithosphere to the NW and SE, which are colder, thicker, and stronger. Low mantle velocities are commonly interpreted in terms of elevated temperature, and this is consistent with regional high heat flow, but they may also partly reflect elevated mantle volatile contents, partial melting, or seismic anisotropy (Lebedev and Nolet, 2003). The high heat flows seen regionally across Sundaland, and the generally low mantle velocities observed in tomographic models, suggest mantle heat-flow values on the order of 40 mW/m2 (Hall and Morley, 2004), which are at the high end of the range estimated globally (Artemieva and Mooney, 1999). Thus they suggest that the region is underlain by a thin and weak lithosphere. Similar lithosphere has been identified in other subduction zone backarcs (Hyndman et al., 2005). Subduction has produced a different mantle beneath the upper plate, associated with a lithosphere that is not like that of older stable continents and their margins. The high heat flow and thin, weak lithosphere are the consequence of long-term subduction at the Sundaland margins,

35

notably at the Indonesian arcs. These features have influenced arc development in some ways not identified in many arc models. ISLAND ARC MODELS Many models of island arc development were formulated in the 1960s and 1970s during the rapid period of development of the plate tectonic theory (e.g., Dewey, 1980; Dewey and Bird, 1970; Dickinson, 1973, 1974, 1977; Dickinson and Seely, 1979; Karig, 1971), and they are still commonly reproduced in textbooks. Much of the subsequent work on island arcs in the past 30 yr has focused on geochemistry and magmatic processes, as it seems to be widely thought that our knowledge of tectonic processes in volcanic arcs is complete. A common arc model (Fig. 8) suggests that after the initiation of subduction (a problem in itself), a volcanic arc develops, typically between 70 and 100 km from the subduction trench (Dickinson, 1973, 1977; Karig and Sharman, 1975); this width increases with the age since arc initiation. An accretionary prism formed by progressive transfer of material from the downgoing slab to the upper plate develops near the trench (Scholl et al., 1986). As time proceeds the accretionary prism becomes larger and wider (Dickinson, 1973, 1977), underthrusting leads to elevation of the arcward part of the prism which may form a forearc high, and the arc-trench gap progressively widens. At the same time, debris from the volcanic arc is deposited in the forearc in basins in front of the arc, and the size and width of these basins increase with time. Continued arc magmatism is interpreted to lead to thickening of the crust beneath the arc. The backarc region is usually not shown in the tectonic models; in some arcs backarc basins are formed, floored by oceanic crust, whereas in others no oceanic backarc basins are formed. The reasons for the formation of backarc basins, which are found mainly in the western Pacific, are still not clear. However, these basins are some distance from the arc, typically >100 km. Little

Figure 7. Depth slices through S20RTS S-wave tomographic model (Ritsema and van Heijst, 2000) for SE Asia. High shear velocities are represented by blue, and low shear velocities by red, with an intensity that is proportional to the amplitude of the shear velocity perturbations. The range in shear velocity variation is given below each map.

36

A

Hall and Smyth Volcanic arc Oceanic crust Trench

Forearc basin

B

Structural in folding (kneading)

crystalline Offscraped oceanic deposits SEDIMENT SUBDUCTION

sedimentary

C Subduction complex

ocean crust crystalline

Lithosphere

small accretionary body

Forearc basin

D Asthenosphere crystalline Forearc basin SUBDUCTION EROSION

SEDIMENT SUBDUCTION

30 km

E CRATONIC MASSIF

100 km

SUBDUCTION EROSION

50 km

Figure 8. Some models for arc development. (A) The arc-trench gap increases in width as progressive accretion occurs at the trench (modified from Dickinson, 1977). As discussed in the text, this type of model is based in part on observations from Indonesia, but new mapping suggests a less important role for subduction accretion. (B, C, D, E) The range of tectonic processes at subduction margins, from accretion to erosion, modified from Scholl et al. (1980).

attention has been given to the area directly behind the volcanic arc, and most arc-derived debris is usually shown to move into the forearc rather than into the area behind the arc. In recent years, in many accounts of arcs in other regions, and in modeling of arcs, the arc model has become more sophisticated. Partly as a result of discoveries made by the Ocean Drilling Program, features such as subduction erosion (e.g., Kay et al., 2005; Ranero and von Huene, 2000; von Huene and Scholl, 1991) and changing slab dips (e.g., Funiciello et al., 2003; Kay et al., 1987; Vannucchi et al., 2004) have been interpreted as important influences on arc development (Fig. 8). However, there has been no ocean drilling of the Indonesian arcs, and few offshore investigations of the forearcs; we are still dependent on field studies and the

knowledge acquired by hydrocarbon exploration in Indonesia to interpret arc history and tectonic processes. The Indonesian arcs differ from these models in a number of ways. Arc magmatism may be discontinuous despite continuous subduction. In many Indonesian arcs there is little evidence of major addition of new material by accretion at the trench, but equally there is little evidence for subduction erosion. There are sudden movements of the position of the volcanic arc, and there may be evidence of significant thrusting within the arc itself, not always with obvious cause; collisions, subduction accretion, subduction erosion, or changing angles of subduction do not offer solutions. In general, in the Indonesian arcs there is evidence of a staccato development, possibly related to processes

Cenozoic arc processes in Indonesia of arc-continent collision but in some cases with no clear reason. Observations of seismicity on the basis of well-located hypocenters (Engdahl et al., 1998; England et al., 2004) show that along the modern arc there are sectors that are aseismic or significantly less seismically active in comparison with other sectors. This may indicate differences in coupling of the downgoing and overriding plate at the subduction zone, and variations in coupling and decoupling in time as well as space could contribute to a discontinuous tectonic history. This could also be related to dynamics of flow in the mantle wedge (S. Lamb, 2006, personal commun.; Oncken et al., 2006). Large-scale dynamic processes have undoubtedly influenced the stratigraphic record of the Indonesian arcs. However, smaller scale effects are also observed. The growth of arc volcanoes has in some parts of the region contributed to basin development by flexural loading. Furthermore the products of the Indonesian volcanic arcs have been modified by processes related to eruption and weathering in a tropical setting. Both these features may have relevance to the interpretation of ancient arcs and their stratigraphic record, as discussed below. We begin with the larger scale effects and move on to more subtle features of arc stratigraphy. Magmatism and Subduction The association between subduction and magmatism is well known. Early plate tectonic models (e.g., Oxburgh and Turcotte, 1970) attributed magmatism to frictional heating at the Benioff zone, melting of the subducting slab, and compressional shortening of the overriding plate in the arc-trench gap. Melts are now considered to originate primarily in the mantle wedge above the subduction zone and are thought to result from the input of volatiles that lower the mantle solidus. In Indonesia, almost all Cenozoic magmatism seems to have been influenced by subduction, although it has not always been the direct result of melting of the mantle wedge above the subducting slab. In some cases a subduction signature is inherited from previous subduction events. The geochemistry of igneous rocks (e.g., Pearce and Peate, 1995) is often used to infer the tectonic setting in ancient arcs, but if a subduction signature is inherited, volcanic arcs may be mistakenly identified in ancient orogenic belts as subduction related when they are not. Macpherson and Hall (1999, 2002) highlighted areas in Indonesia where there has been arc-like magmatism but subduction is not active, such as south Sulawesi (Polvé et al., 1997) or northern Java, where volcanoes are very far from the active arc and have an ultra-potassic character (Edwards et al., 1994; Macpherson, 1994). In these places the character of magmatism reflects older subduction events that have enriched the mantle. However, the association between magmatism and subduction is commonly not straightforward. Some arc models portray compressional arcs with active magmatism, and contraction is often thought to be typical of active arcs, probably because the subduction model causes us to think of plate convergence. In the Indonesian arcs, contraction is commonly associated with

37

cessation of arc activity. Hamilton (1988, 1995) pointed out that subduction zones are often portrayed as static systems with the subducting slab rolling over a stationary hinge and sliding down a slot fixed in the mantle, but this is not a good model. He emphasized the necessity of viewing subduction in a dynamic way and argued that it is most commonly driven by gravity acting on the subducting slab, and most subduction zones are characterized by retreat of the hinge with time, or rollback. Extension of the overriding plate is known to be common in arc settings (e.g., Chase, 1978; Dewey, 1980; Elsassar, 1971; Hamilton, 1988, 1995) and may be the normal condition for a volcanic arc. For the period since 25 Ma, tectonic models (Hall, 1996, 2002) suggest that hinge retreat (Fig. 9) in most subduction zones in SE Asia and the western Pacific was accompanied by significant arc volcanism (Macpherson and Hall, 1999, 2002) and in some zones by marginal basin formation. Extension in arcs may be essential at regional and local scales to induce melting, to provide pathways for magmas, and to vary the rate of supply of volatile components to the mantle wedge. In particular, hinge retreat allows replenishment of the mantle wedge by an inflow of hot, undepleted mantle (Andrews and Sleep, 1974; Furukawa, 1993). This can also cause ablation of the base of the overriding plate (Furukawa, 1993; Iwamori, 1997; Rowland and Davies, 1999), and the mantle that upwells may melt by decompression. These processes can maintain steady-state arc magmatism at a relatively constant distance from the trench over prolonged periods. Where subduction is characterized by a fixed hinge, or by hinge advance, magmatism may decline or cease, although subduction continues. A fixed hinge may result from coupling of the subducting and overriding plate or from collision, causing the two plates to move in the same direction as the slab is being subducted. In these cases, magmatism may cease, although subduction continues, simply because the mantle wedge is not being replenished, becomes depleted, and can no longer melt. For the period since 25 Ma (Fig. 9), hinge advance in SE Asia and the western Pacific was accompanied by reduction in magmatism (Macpherson and Hall, 1999, 2002). The Sunda Arc provides a good example. From Java eastward the hinge advanced from ca. 20 to 10 Ma, and magmatic activity declined. In East Java, Paleogene volcanic activity culminated with a short-lived explosive phase (Smyth et al., this volume) and then declined significantly or terminated; magmatism resumed at ca. 12–10 Ma. There is no indication that subduction ceased or slowed during the period of reduced magmatism, and the Indian-Australian plate continued to move northward relative to Eurasia at a similar rate to that during the Paleogene. The reduction in magmatism can be understood as the result of hinge advance driven by the collision of Australia in eastern Indonesia (Hall, 1996, 2002) and manifested by counterclockwise rotation of the Borneo region (Fuller et al., 1999). When rotation of Borneo ceased, the hinge retreated again and magmatism resumed, especially in the Banda Arc, where dramatic retreat of the subduction hinge required a massive influx of fertile mantle beneath the overriding plate.

25-5 Ma Japan 25-5 Ma

u

ky

u Ry

Izu-Bonin 25-5 Ma

Phil

Hinge Advance

ippi nes

Mariana 25-15 Ma Mariana 15-5 Ma

Sangihe

Sunda 25-5 Ma

Halmahera 15-5 Ma

Solomons New Hebrides 10-5 Ma

Sunda-Banda 20-10 Ma

A: Andaman Sea BS: Banda Sea Bi: Bismarck Sea M: Mariana Trough Wd: Woodlark Basin

5-0 Ma Japan

u

ky

u Ry

ippi renc ne T

M

Hinge Retreat

Mariana

ihe

Halmahera

North Sulawesi

Sunda

Hinge Advance

h

ng

Sa

A

Izu Bonin

Phil

Manila Trench

Hinge Retreat

BS da

Ban

Bi

New Britain Solomons

Wd

Figure 9. Movements of subduction hinges in the Neogene, modified from Macpherson and Hall (2002). Regional plate reorganizations occurred at ca. 25 Ma and 5 Ma. Major coastal outlines are shown for reference. For subduction zones shown without shading, there was no significant movement of the hinge. Bold letters indicate areas of young marginal basin formation.

Cenozoic arc processes in Indonesia Subduction Accretion and Subduction Erosion A feature of many early arc models is an increase in the width of the arc-trench gap with increasing age, commonly attributed to accretion, and some of these models were based on observations from Indonesia (e.g., Dickinson, 1973; Moore and Karig, 1980). In contrast, in many arcs the role of subduction erosion in the development of the arc has become widely recognized as important (e.g., Scholl et al., 1980; von Huene et al., 2004). In Indonesia these contrasting processes are difficult to assess, because the age and nature of the crust between the arc and the trench in Indonesian arcs is unexplored or poorly known. There is a great need for detailed exploration of these parts of all arcs. However, observations on land suggest that neither subduction accretion nor erosion has had a significant impact on arc history. Change in Width of Arc-Trench Gap with Age There is good evidence of Early Cretaceous subduction and widespread evidence from Sumatra to SE Borneo of middle to Late Cretaceous collisions accompanied by ophiolite emplacement (Miyazaki et al., 1998; Parkinson et al., 1998; Wakita, 2000). As noted above, there is almost no Paleocene stratigraphic record in Indonesia, and the whole region seems to have been emergent. In many parts of Indonesia, poorly dated, typically terrestrial rocks interpreted to be Paleocene or Eocene rest unconformably on older basement rocks. In East Java the oldest sedimentary rocks that rest unconformably on ophiolitic basement are middle Eocene or older and lack volcanic debris (Smyth, 2005; Smyth et al., this volume). There is little evidence for latest Cretaceous to early Eocene volcanic activity in most of the Sundaland margin between Sumatra and Sulawesi. The Cenozoic stratigraphic record, commonly volcanogenic, begins in the middle Eocene, and in Java the current phase of subduction along the arc began after 45 Ma, when Australia began to move rapidly northward. Although it is commonly assumed that subduction at the Sunda Trench in the Late Cretaceous (e.g., Metcalfe, 1996; Hall, 2002) continued into the Cenozoic, we question this assumption. Hamilton (1988) stated that the modern subduction system in Java was inaugurated no earlier than late Oligocene time. We, too, consider that the present subduction began only in the early Cenozoic, although we now know of evidence for subduction beneath Java before the late Oligocene (Smyth, 2005; Smyth et al., this volume). However, there is no evidence in Java to support the suggestion of late Paleogene southward subduction or an arc collision in or with Java in the Paleogene (Hamilton, 1988). In the Sunda Arc between Sumatra and Java the arc-trench gap is rather large, between 300 and 350 km (Fig. 1), and subduction began in the early Cenozoic. If the arc-trench gap increased progressively with age, there should be evidence of movement of either the volcanic arc or the trench during the Cenozoic. In East Java the volcanic arc remained in the same position from the Eocene to the early Miocene (Smyth, 2005; Smyth et al., this volume). It then ceased activity, and a new arc formed ~10 m.y. later, 50 km to the north, where it has remained since the late Miocene.

39

Near the trench there is some evidence for a young accretionary complex constructed against an older backstop (Fig. 10), interpreted to be an older accretionary complex (Kopp et al., 2002, 2006; Schlüter et al., 2002), but this is only ~50 km wide. Schlüter et al. (2002) suggest there was a progressive growth of the arc-trench gap south of Sumatra since the Paleocene, but the age control on this development is poor. In contrast, Kopp et al. (2006) suggest there has been subduction erosion at the Sunda margin south of Java, and if this is correct, the arc-trench gap must have been wider in the past. The seismic lines on which these interpretations are based do not show a progressive growth of the accretionary zone but rather a young accretionary complex juxtaposed against an older one (Fig. 10). This relationship could equally be interpreted in terms of an abrupt change in the subduction history, and we note that on land there is evidence for such a change; the volcanic arc ceased activity at ca. 20 Ma and resumed activity ~10 m.y. later, 50 km north of its previous position. Why this happened is not clear. At present the available data are simply not adequate to distinguish different models. However, investigations of the Sumatra forearc in the Nias region do provide the basis for testing the interpretation of arc growth by continuous accretion. Nias, at the north end of the Sumatra forearc (Fig. 1), is commonly cited as the classic example of a forearc high elevated by accretion, following the work of Moore and Karig (1980) on the basis of their mapping of part of the island. Later mapping of the whole island and other islands of the forearc does not support this model (Samuel, 1994; Samuel et al., 1995, 1997; Samuel and Harbury, 1996). There is little evidence on Nias for the progressive addition of material. The ophiolitic basement includes Cretaceous rocks and locally is overlain unconformably by a thin Eocene marine cover. Most of the island consists of a thick sequence of Oligocene to lower Miocene deep marine clastic sedimentary rocks and a thick sequence of lower to upper Miocene shallow marine clastic sedimentary rocks. There are a few limestones and some tuff layers in the sequence. The island became emergent in the Pliocene. The ophiolitic basement was in place by the Eocene, and the deep marine sedimentary rocks above the basement were deposited on it. However, the sedimentary rocks are not material scraped off the downgoing plate but represent the fill of a forearc basin, later inverted (Fig. 11), in which most material was carried from the direction of Sumatra. The Oligocene to early Miocene development of the outer forearc was characterized by extension, followed by early Miocene inversion. Mélanges are not subduction accretion complexes but are diapiric in origin and formed by mud volcanism that continues today. The forearc islands have a complex history of local uplift and subsidence to which addition of material at the trench has contributed, but the Sumatra forearc does not fit very well to the model of continuous widening and accretion interpreted from the early land-based studies. Mélanges Mélanges were formed at numerous stages in the development of arcs in Indonesia (Hall and Wilson, 2000; Hamilton,

40

Hall and Smyth NE

SW 0

JAVA SEA

A

Outer Arc High

Mentawai Forearc Basin

TWT (s)

2 4

INDIAN OCEAN 10°S 100°E

6

C

Java

Trenc h

Mentawai Wrench Fault

Accretionary Wedge I

110°E

Accretionary Wedge II Sunda Trench

8 10

B

A

Continental crust 50 km

Oceanic crust of Indian Ocean Plate ACCRETIONARY DOMAIN

SW

0 Trench infill

Depth (km)

Forearc Strata

FOSSIL

ACTIVE FRONTAL PRISM

-10

NE

OUTER HIGH

Sub

Oceanic mantle

SHELF

OCEANIC-TYPE CRUST

Shallow Mantle

duct

-30

B

ing S

lab 50 km

-40 S

OUTER FOREARC HIGH

N

FOREARC BASIN

0 1

TWT (s)

2

Basin Sediments: Hemipelagic / Pelagic sediment and Volcanic Ashes

Slope Deposits Trench

3

Vertical Movement and Doming

Normal Faulting

Post-sedimentary buckling

4 5 6

C

40 km

Figure 10. Cross sections across the Sunda forearc, based on marine geophysical investigations. (A) At the south end of Sumatra the accretionary wedge is interpreted to be ~300 km wide and constructed against continental crust beneath the forearc basin (Schlüter et al., 2002). (B) Interpreted tectonic units of the Sunda margin south of West Java (Kopp et al., 2002), where the accretionary domain is interpreted to have an active part and an older part, juxtaposed against oceanic-type crust beneath the forearc basin. (C) Section south of Central Java (Kopp et al., 2006) shows an uplifted forearc high, covered by thin sediments interpreted as evidence of subduction erosion. The age of the unconformity is unknown, but the thin sediments above the unconformity suggest a young age. A similar unconformity on land is late Miocene to Pliocene, which developed after northward thrusting within Java.

1979). Marine geophysical studies suggest that mélanges are formed at Indonesian subduction zones (Kopp et al., 2002; Schlüter et al., 2002), but they are also formed at later stages during collision (Silver and Moore, 1978). In the central Molucca Sea (Fig. 12) mélanges were formed at two stages. Mélanges reported from Talaud (Moore et al., 1981) and present on Mayu were not formed during the present arc-arc collision but are older rocks forming part of the pre-Neogene basement of the Sangihe forearc. Presumed mélanges of the present collision complex are all submarine and constitute part of the bathymetrically shallow

and seismically incoherent volume of sediment in the central Molucca Sea. Mélanges have also formed far from subduction zones without collision (Samuel et al., 1997). In Timor, mélange formation is attributed by some authors mainly to thrusting related to collisional processes (Harris et al., 1998) and by others mainly to diapiric processes expressed as modern eruption of mud volcanoes (Barber et al., 1986). Active mud volcanism in East Java (Smyth, 2005), behind both the Paleogene and modern arcs, erupts rock samples from deeper parts of the basin directly north of the modern arc and brings to the surface blocks

Cenozoic arc processes in Indonesia

Mujoi Sub-basin

Lahewa Sub-basin

Mola Basement High

Gomo Sub-basin

SEDIMENT & MELANGE COMPLEX

MID-MIOCENE– EARLY PLIOCENE Coast

41

OLIGOCENE– EARLY MIOCENE SEDIMENT

LATE PLIOCENE– PLEISTOCENE Coast

MID-MIOCENE– EARLY PLIOCENE BASEMENT

NE

SW 0

SW Indian Ocean

Accretionary wedge

Trench

Nias

Forearc basin

V=H

10 Km

NE

Sumatran fault zone, magmatic arc & backarc

Continental crust 0

V=H

60

Km

Moho

+

Overriding continental lithosphere

Figure 11. Cross sections across the Sumatra forearc, modified from Samuel and Harbury (1996). Mapping of Nias by Samuel (1994) shows that the area between Sumatra and Nias was in extension for much of the Cenozoic and that most of the island’s sediments were derived from Sumatra. The island has not emerged as a forearc high as the result of progressive subduction accretion. V=H—vertical and horizontal scales are equal.

of the oldest parts of the basin sequence entrained in overpressured muds. This is clearly not related to subduction because the Benioff zone is >100 km below the site of mud volcanism. The mélanges at the surface, and the presumed equivalents beneath the surface, are the result of overpressures related to rapid deposition of thick sedimentary sequences, hydrocarbon generation, and young thrusting very far from the subduction zone. Mélanges and Ophiolites The emplacement of ophiolites is often linked to the processes of mélange formation and accretion. Between the Sangihe and Halmahera Arcs (Fig. 12) is a collision complex in which there are ophiolites and mélanges. The Talaud Ridge at the center of the Molucca Sea was interpreted on the basis of marine geophysics as a mélange wedge (McCaffrey et al., 1980; Silver and Moore, 1978) including slices of the Molucca Sea lithosphere (Fig. 13). These rocks are exposed on Talaud and on two tiny islands along the ridge, Mayu and Tifore. The ophiolites were interpreted by McCaffrey (1983, 1991) as part of the subducted Molucca Sea plate, which has been emplaced by the collision

of the two arcs (Fig. 13). However, on Talaud a post–middle Miocene sedimentary sequence rests unconformably on ophiolites that are middle Eocene or older (Moore et al., 1981). The middle Miocene to Pleistocene rocks are tuffaceous sandstones, siltstones, and shales with intercalations of limestone, marl, and conglomerate in which the sediments are dominated by volcanic debris. They were deposited in deep water by turbidity currents. On Talaud, Neogene strata are in many places little deformed, and they are moderately to strongly deformed locally and typically in narrow zones adjacent to mélange rocks or ophiolites. On Mayu, coastal exposures of highly indurated mélange yield Eocene ages (Baker, 1997). All of the mélange is probably older than middle Miocene, and none appears to be related to the present collision. The similarity in character and structure of the ophiolitic rocks to those of the basement complex on Halmahera (Hall et al., 1988a) suggest that much of the mélange formed during Eocene or older deformation events. The interpretation of the ophiolites as part of the subducted Molucca Sea lithosphere requires some complex arguments (McCaffrey, 1991) to account

42

Hall and Smyth Philippine Fault

Philippine Trench

EURASIAN PLATE Sulu Sea

PHILIPPINE SEA PLATE

Mindanao

00

18

km Celebes Sea

Cotobato Trench Sangihe Arc

Halmahera Arc Halmahera

MOLUCCA SEA PLATE

SORONG FA ULT ZONE

Sulawesi

New Guinea

Banda Sea AUSTRALIAN PLATE

1800 km

Figure 12. Molucca Sea region, showing the converging Halmahera and Sangihe Arcs, displayed in a three-dimensional (3D) diagram that represents, in simplified form, the geometry of the converging plates in east Indonesia. Modified from Hall et al. (1995).

for their elevation instead of their disappearance by subduction. In contrast, if the ophiolites are interpreted as part of the Eocene or older basement (Fig. 13), as they are in Halmahera, and they formed the basement to the forearc, the gravity model is much easier to understand. Where the Halmahera forearc and arc have been significantly overthrust the Sangihe forearc has been jacked up. The wide Molucca Sea collisional complex is composed of the accretionary wedges of both arcs. The basement of the Sangihe forearc is exposed where it thrusts over this wedge. The ophiolitic rocks of the central Molucca Sea are not part of the Molucca Sea plate but are the basement of the Sangihe forearc (Hall, 2000). The ophiolites were the basement before the middle Miocene. Accretion or Erosion? In many Indonesian arcs, there seems to have been relatively little growth of the forearc region by the addition of material from the downgoing plate. Many so-called accretionary complexes are in fact dominated by material derived from the arc. At the northwestern end of the Sunda Arc there is undoubtedly sediment at the distal end of the Bengal Fan (Curray, 1994) that is being added to the overriding plate, although most of this is probably mud-dominated sediment (Stow et al., 1989). Farther south, south of south Sumatra and Java, the amount of sedimentary cover on the subducting plate is much thinner, yet the arc-trench

gap is just as wide. The role of the upper plate in providing material is well illustrated at the eastern edge of the Philippine Sea plate in the Nankai margin, where classic fold and thrust structures are formed in the thick “accretionary wedge” (Mikada et al., 2005; Taira et al., 1991). The sediment in the upper part of the prism is dominated by clastic material carried along the trough from the Japan–Izu-Bonin Arc collision zone, and relatively small amounts of material are contributed from hemipelagic sediments on the subducting slab (Karig and Ingle et al., 1975; Klein and Kobayashi et al., 1980; Taira et al., 1991). In Indonesia the role of the structurally upper plate is well illustrated by the eastern Makassar Strait, where there has been no subduction, yet accretionary-style structures have developed in sand-dominated sedimentary wedges that have built out from Sulawesi (Puspita et al., 2005). These structures closely resemble the structures developed in subduction accretionary complexes, and major eastward subduction during the Cenozoic has been suggested by some authors (e.g., Charlton, 2000; Guntoro, 1999). However, most authors have interpreted the West Sulawesi margin to be either a passive margin with oceanic crust to the west (e.g., Bergman et al., 1996; Cloke et al., 1999; Hall, 1996; Hamilton, 1979) or a rifted margin with the thinned continental crust to the west (e.g., Calvert, 2000; Situmorang, 1982). Recent offshore seismic data show that there has been no subduction (Fraser et al., 2003; Puspita et al., 2005) and suggest that the Makassar Strait is floored

Accretionary complex Accreted sediments and crust SNELLIUS RIDGE of Halmahera forearc Backarc basin

Forearc basin 10 20

v

v v v v v v v v v v v v v v v v v v v v v Arcv crust v v v v v v v v v v v v v v

Kilometers

30

EURASIAN PLATE Sangihe Arc

40

v

v v v v v v v v v v v v v

Forearc crust

Fo

rea

ru rc c

st

v v v

v v v v

Arc crust v v v v

v

v v v v v

v v

MINDANAO

Backarc crust v

v v

v

Talaud

PHILIPPINE SEA PLATE Halmahera Arc

50

Sangihe mantle

60

A Snellius Ridge

Sangihe

Halmahera mantle

B

70 80

A

MOLUCCA SEA PLATE

Morotai ES

I

90

SU LA W

100

Mayu

HA LM AH ER A

0

1

TALAUD

SANGIHE

C s

s

NW 0 10 20

SANGIHE v v v v v v v v v v v v v v v Arc crust v v v v v v v v v v v v v v

MOROTAI

Forearc basin

Forearc basin v v v v

SE

Accretionary complex

v

Forearc crust

Forearc crust

v

v v v v v v v v v v v Arc crust v v v v v v v v v v v v v v v v

Backarc basin

v

v v v

v v

Backarc v v v v v v v v

crust

Kilometers

30

EURASIAN PLATE Sangihe Arc

40

PHILIPPINE SEA PLATE Halmahera Arc

50 60

Halmahera mantle

Sangihe mantle

70

MOLUCCA SEA PLATE

B

80 90 100

100 50 Calculated

0

50

100

-50

Talaud-Mayu 150 Ridge

200

-0.59

-0.59

-0.43

MGAL

0

Observed

-100 -150

250 km

-1.83

-0.43

0.40

10 20

C

0.44

WNW 100

Sangihe Arc

0

km

100

200

30 Halmahera Arc

ESE 300

0 10 20

2.86

Kilometers

30 40 50

3.40 g/cm3

60 70 80

3.34

90

Mantle density

+0.20

+0.05

+0.05

+0.20

D

100

Figure 13. Cross sections across the Molucca Sea (Hall, 2000) drawn at the same vertical and horizontal scales (A, B) in comparison with those based on gravity modeling (C, D) by McCaffrey et al. (1980). In section A at the latitude of Talaud, the entire arc and forearc of the Halmahera Arc has been overridden by the Sangihe forearc. Ophiolites interpreted as part of the Sangihe forearc basement are exposed in the Talaud Islands. Farther south (section B), only part of the forearc has been overridden, but the Halmahera Arc in Morotai was overridden by its own backarc in an earlier thrusting episode. In contrast, McCaffrey et al. (1980) interpret the ophiolites as part of the subducted Molucca Sea plate.

44

Hall and Smyth

by extended continental crust (Nur’aini et al., 2005; Puspita et al., 2005). The structures are foreland-type fold and thrust belts, and developed as westward thrusting has progressed since the Pliocene, all the material having been derived from Sulawesi. There certainly has been addition of material to the continental margin of Sundaland during the Cenozoic (Fig. 4). However, the additions occurred during relatively short periods of time, at widely spaced intervals, and were related to the collision of continental fragments at the Sundaland margin. The material added included arc and ophiolitic rocks, which may in some cases represent backarc basins within the continental margin or the remnants of material subducted during the arc-continent collision. In Indonesia, arc growth seems to have occurred discontinuously rather than by continuous steady-state addition of material at the trench that led to gradual widening of the arc-trench gap. Disappearance of Arcs Westward subduction of the Molucca Sea beneath the Sangihe Arc probably began in the early Miocene. Eastward subduction of the Molucca Sea plate beneath Halmahera began in the middle Miocene. The double subduction zone was initiated at this time, forming a new plate, the Molucca Sea plate, separate from the Philippine Sea plate. The oldest volcanic rocks dated from the Halmahera Arc are 11 Ma in Obi at its southern end, and they become younger to the north (Baker and Malaihollo, 1996). The Molucca Sea was eliminated from south to north, and the two forearcs began to collide. The result of the collision was that the Halmahera Arc system (meaning the entire region between the trench and backarc region) failed repeatedly, with thrusting in different directions at different stages in the collision. First, the backarc was thrust over the volcanic arc, and later the forearc was thrust toward the volcanic arc. In south Halmahera the backarc region was thrust onto the forearc, in places entirely eliminating the Neogene arc. At the southern end of the Halmahera Arc on Obi the arc was thrust onto the forearc (Ali and Hall, 1995; Ali et al., 2001). After west-vergent thrusting, volcanism in the Halmahera Arc resumed between Bacan and north Halmahera. At the south end of the arc on Obi, and at the north end from Morotai northward, volcanism ceased. In the northern Molucca Sea the Sangihe forearc was then thrust east onto the Halmahera forearc and arc. In the region between Morotai and the Snellius Ridge, parts of the Neogene Halmahera Arc and forearc have now disappeared. Farther south, this east-vergent thrusting carried the Halmahera forearc onto the flanks of the active Halmahera Arc, and pre-Neogene rocks of the Halmahera forearc basement are now exposed in islands of the Bacan group and off the coast of northwest Halmahera. Cross sections drawn across the present-day collision zone from south to north can also be considered to display the sequence of events in time, and a series of sections illustrating the earlier stages in the collision can be inferred from the geology of the Halmahera Arc (Fig. 14). An interesting consequence of the collision is that the Halmahera Arc is being progressively overridden by the Sangihe Arc, and in the northern Molucca Sea almost

the whole of the forearc and volcanic arc have disappeared. This process is not complete, and it is likely that much, if not all, of the Halmahera Arc will disappear beneath the Sangihe Arc. This process is certainly not subduction erosion, as intended by most authors (Scholl et al., 1980; von Huene and Scholl, 1991; von Huene et al., 2004). However, neither is it subduction accretion, as usually envisaged. The arc-trench gap of the Sangihe Arc has increased since the two arcs came into collision, but the arctrench gap of the Halmahera Arc has diminished. Arc Movements and Changing Slab Dips Observations in Indonesian arcs raise the question of whether there is such a thing as a compressional arc. Arcs may certainly be compressed, but when contraction occurs volcanic arc activity ceases. Magmatism may resume when compression ceases, but the stratigraphic record suggests it does not persist during contraction. In the Banda Arc, magmatic activity terminated in the Wetar sector of the volcanic arc at ca. 4 Ma, when arc-continent collision began in Timor, and was never resumed during stacking of nappes (Audley-Charles, 2004). In the Halmahera Arc, volcanic activity ceased as arc-arc collision began but resumed in a new location shortly afterward. The repeated failures in different places in the entire arc system in Halmahera suggest that intermittent thrusting was followed by periods of relaxation in which arc magmatism resumed. Such abrupt movements in the position of the volcanic arc are not uncommon and may follow periods when an arc is compressed, but they may be due to other tectonic causes. Sudden shifts in position of the volcanic arc could be related to shortening of the entire arc system (forearc, arc, and backarc), which may involve thrusting of the backarc or the forearc toward the arc, the magmatic arc itself being the weakest point. Field observations in Halmahera show that the site of thrusting was close to the former active arc. Experimental and numerical modeling studies also show that the active arc is the weakest point in the whole arc system (Shemenda, 1994; Tang and Chemenda, 2000; Tang et al., 2002). These studies show that failure occurs directly beneath the arc. It can take place on faults dipping toward the trench or arc, and this depends on the thickness of subducted crust and flexural rigidity of the upper plate. Since the middle Pliocene (Fig. 14) the Halmahera Arc system failed more than once close to the site of the active volcanic arc, presumably reflecting its weakness owing to mineralogy and magmatism. Active volcanism has most recently resumed in a restricted area at the center of the arc chain, over a distance of ~200 km, in comparison with the 700 km length of the early Pliocene arc. The arc-arc collision explains why the arc volcanism may have ceased and then resumed as stresses changed, but this does not really account for the change in position of the arc. The Quaternary volcanoes are situated ~50 km west of the Pliocene centers (Fig. 15). Hall et al. (1988b) speculated that the shift may be related to a steepening dip of the subducted slab beneath Halmahera. A highly complex geometry of subducted slabs lies

Talaud ridge

0 Ma Philippine Sea plate Halmahera Arc and forearc overridden by Sangihe forearc Molucca Sea plate sinks deeper Emergence of ophiolitic basement of Sangihe forearc

0 Ma

Halmahera forearc overridden by Sangihe forearc Molucca Sea plate sinks Local emergence of islands in ‘collision complex’

0 Ma

Halmahera Arc fails again Overthrusting of forearc region: detachments deep within basement Minor backthrusting at back of Sangihe forearc

Halmahera Arc

Sangihe Arc

2 Ma

Molucca Sea plate Volcanic activity ceases in Halmahera arc Arc completely overridden by backarc region

Figure 14. Cross sections across the Molucca Sea drawn at same vertical and horizontal scales to illustrate the sequence of convergence of the Halmahera and Sangihe Arcs since 2 Ma. The lowermost section is inferred from geological mapping on land. The upper three sections are drawn at different latitudes across the Molucca Sea from south (bottom) to north (top). They can be considered to represent the sequence of events in the last ~2 m.y. that have led to the structure at the latitude of Talaud, where convergence is most advanced. Collision has resulted in the almost complete elimination of the Halmahera Arc and forearc at the latitude of Talaud.

A

129oE

M O RO TA I

127oE

o

2N

Modern volcanoes of the Halmahera Arc Pleistocene to modern Halmahera Arc

Molucca Sea

Miocene to Pliocene Arc HALMAHERA

0o

Weda Bay

0

50

100

BACAN Km

B

o

o

110 E

112 E

Modern volcanoes of the Sunda Arc Oligo-Miocene volcanic centers of the Southern Mountains Arc

7oS

EAST JAVA

SUND

A ARC

SOUT MOUN HERN TAIN ARC S

8oS

0

50

100

Km

Figure 15. Abrupt movements of the volcanic arcs in (A) Halmahera (ca. 2 Ma) and (B) Java (ca. 10 Ma) took place, as discussed in the text.

Cenozoic arc processes in Indonesia beneath the Halmahera and Sangihe Arcs and the southern Philippines (Fig. 12). The west-dipping Philippine slab has now arrived at depths where it would hit the subducted Halmahera slab and increase its dip, thus shifting the arc to the west if melting occurs at a constant depth. Macpherson et al. (2003) showed that differences in the geochemistry of Quaternary lavas in comparison with Neogene lavas of the Halmahera Arc are consistent with an increased sediment flux interpreted to indicate an increase in dip of the subducted slab. In East Java, activity in the Southern Mountains Arc terminated in the early Miocene after a period of >20 m.y. at the same location. About 10 m.y. later a new arc formed, 50 km north of the older arc, that has remained at the same position since the late Miocene (Fig. 15). Why the volcanic arc moved north is not known. There is no evidence for a collision at the Java Trench. There was Neogene contraction in Java, although it is not well dated. It is possible that the slab dip remained at the same angle, the depth to the Benioff zone remained constant, and therefore contraction of the arc meant that when volcanism resumed the new arc formed at the same distance from the trench. Several possible explanations are suggested by Smyth et al. (this volume), but all are speculative. The abrupt change in position of the volcanic arc is a feature of several Indonesian arcs, and the causes are not known. However, we do not know of examples of gradual movement of the arc with increasing age of the magmatic arc, as might be expected from many arc models. This could mean that slab dip gradually declines as the arc-trench gap increases, so that the arc remains in the same position. Obviously this idea is difficult to test, as there is no way of measuring slab dip in the past. The dip of slabs in most Indonesian arcs is typically steep where the slab reaches >~200 km in depth. The depth to the Benioff zone beneath active arcs is also much more variable than previously thought (England et al., 2004). Perhaps magmatic activity essentially pins the arc in one position until it ceases, and then there are a number of factors which could control the location of the resumption of volcanic activity, including thickness or strength of the crust, preexisting weaknesses in the crust, or changing dip of the subducting slab. At present, all are conjectures and illustrate how incomplete our knowledge of arcs remains. Volcanic Arc Loading Volcanoes exert a load. Since most basic to intermediate arc volcanoes are denser than average continental crust, longlived volcanism should contribute to subsidence. Several of the Indonesian arcs have deep filled basins close to, and both in front of and behind, the arc. Few of these basins have been explored seismically, and their deep structure is unknown. However, there are indications that arc loading may have contributed to their formation, because the basin sequences thicken toward the arc, and the timing of basin development is closely linked to activity in the arc. As discussed by Nichols and Hall (1991) and Smyth et al. (this volume), the basins are not what are normally

47

considered backarc basins. They are found behind the arc but are much closer to it than typical backarc basins. They are not underlain by newly formed oceanic crust, nor are they characterized by obvious extension. They are filled with volcanic arc material, mainly reworked as sediments. The basins formed in a setting that may have been extensional (as noted above, possibly the typical condition of a volcanic arc) or neutral, using the terminology of Dickinson (1995) and Busby and Ingersoll (1995), but certainly not compressional. They are not retro-arc basins, and although there may be evidence of thrusting, this occurred after the basin formed. East Java and Halmahera provide two examples. In Java the deep Kendeng Basin behind the arc is filled with a thick sequence of volcanic and sedimentary rocks mostly derived from the arc itself. The basin formed in the middle Eocene, when the volcanic arc began its activity. Modeling by Smyth (2005) suggested that at least part of the accommodation space was created by the load of the volcanic arc. The basin in Weda Bay (Nichols and Hall, 1991) behind the Halmahera Arc has been investigated as part of oil exploration, and seismic lines close to the arc show no indication of fault control on subsidence (see Waltham et al., this volume). Again, a likely important contribution to subsidence is flexural loading by the volcanic arc. This suggestion is explored in more detail by Waltham et al. (this volume) in a mathematical model. They conclude that volcanic loading can make a contribution to basin subsidence in arc settings within ~200 km of the arc, which may be the primary cause. In continental margin arcs such as Java, and long-lived intraoceanic arcs such as Halmahera, there is a significant density contrast between the deeper crust and the eruptive arc products. In these cases the volcanoes form a load that can produce or contribute to the formation of flexural basins close to the arc. Elsewhere the absence of deep basins in arcs may be due to the small density contrast between volcanoes and the underlying crust. Sediment Character in Indonesian Arcs All the Indonesian arcs have been in an equatorial position throughout the Cenozoic (Hall, 2002), and as a result the arc products have formed in a tropical climate. Tropical processes have several effects on rocks and grains within them. Deep tropical weathering of well-jointed rocks can lead to a high degree of in situ rounding of material that eventually falls out of the outcrop (Fig. 16). The onion-skin type of weathering of outcrops is common in Indonesia, and rounded clasts are formed at all scales from huge boulders to pebbles. The rounded clasts may be incorporated in volcaniclastic and sedimentary rocks without any transport whatsoever. Weathering of material also leads to rapid destruction of unstable rock fragments and minerals. This is particularly important for volcanic-derived material that contains many unstable minerals and rock fragments such as mafic minerals, feldspars, and clays. The interpretation of transport history and maturity based on grain shape and light mineral modes can be very misleading in tropical settings. Two recent studies of

48

Hall and Smyth

Figure 16. Tropical weathering of rocks, resulting in rounding of clasts before they are even removed from the outcrop. Rounded clasts are produced at all scales from small pebbles to boulders. Both photographs show rocks exposed by quarrying of the Pendul Diorite, Jiwo Hills, East Java.

sedimentary rocks deposited close to active margins in the Indonesian region illustrate the effects of tropical weathering (van Hattum, 2005; Smyth, 2005). In both cases the tectonic setting is known, but if these examples had come from much older orogenic belts the interpretation could well be very different. There was subduction of the proto–South China Sea beneath north Borneo (Hall, 2002; Hazebroek and Tan, 1993; Hutchison, 1996; Tongkul, 1991) between the Eocene and early Miocene. A large sedimentary fan, the Crocker Fan, formed at this active margin. All sandstones are quartz-rich, and their compositions plotted on conventional ternary diagrams suggest they are mature recycled orogenic products, but there are some anomalies (van Hattum et al., 2003; van Hattum, 2005). The oldest sandstones show the greatest compositional maturity, and although younger sandstones plot as recycled orogenic sediments they are less mature, which is inconsistent with their derivation by recycling of older sedimentary rocks as suggested by Hutchison et al. (2000) and William et al. (2003). The sandstone textures are in marked contrast to their apparent compositional maturity (van Hattum, 2005; van Hattum et al., 2006). They are immature, typical first-cycle sandstones; grains are angular to subangular, and there are few clasts that suggest recycling. The sandstones are poorly sorted and have a muddy matrix and very low porosity, and they contain abundant euhedral and subhedral zircon grains typical of first-cycle sandstones. Subrounded and (rare) rounded grains are less abundant. Tourmaline also occurs predominantly as unabraded grains. The shapes and lack of abrasion of zircons and tourmalines indicate that long-distance transport is unlikely. Zircon and tourmaline typically make up >70% of the heavy mineral assemblages in the Crocker Fan sandstones. Their abundance indicates erosion from acid plutonic rocks, and their shapes suggest a nearby source area. Apatite would normally also be abundant in material derived from acid plutonic rocks, but it is commonly absent, and those grains present are typically pitted or partially dissolved. This suggests that the abundance of apatite

reflects weathering rather than provenance. Apatite is stable during burial but is susceptible to acidic weathering, which is common in humid tropical settings (Morton, 1984). Plots of chemical indices of weathering and alteration (Fig. 17) show that very few of the sandstones are similar in composition to the fresh source rock (van Hattum, 2005). The composition has changed between erosion and deposition. The tropical setting of the sandstones needs to be considered before provenance can be interpreted. In East Java, Eocene to lower Miocene quartz-rich sandstones plot on commonly used ternary diagrams as recycled orogenic or cratonic interior–derived sediments (Fig. 18), but this interpretation is even more misleading than that for the Borneo sandstones. Examination of quartz grains shows that many have a volcanic origin, and a volcanic provenance is supported by the abundance of fresh euhedral zircons. The sandstones have previously been interpreted as eroded from a continental Sundaland source because they are rich in quartz and well sorted, but careful examination reveals abundant evidence of their volcanic origin on the basis of textures, light mineral constituents, quartz character, clay mineralogy, and zircon character and ages (Smyth, 2005; Smyth et al., this volume). These quartz-rich rocks are not the product of recycling of material derived from continental source regions but are the result of explosive volcanic events such as the eruption of crystal-rich magma, followed by air-fall sorting and subsequent epiclastic reworking (Cas and Wright, 1987; Walker, 1972). In tropical settings, intense weathering rapidly causes the breakdown of labile minerals, mineral aggregates, and lithic fragments. The redeposited sediment will be rich in resistant minerals such as quartz and heavy minerals like zircon, and it has a higher percentage of quartz per unit volume than the source material. The high rates of weathering observed in tropical settings can also have a significant influence on the preservation of unconsolidated volcanic deposits. The potential for preserving volcaniclastic deposits on steep terrestrial slopes is low unless they are rapidly overlain by lavas or buried under younger arc

Cenozoic arc processes in Indonesia

100

100

Strongly weathered

80 70 60

Inc

95 90

rea

sin

gc

he

mi

ca

rity

80

50 40

lm

atu

85

SiO2%

CIW (100)[Al2O3/(Al2O3+CaO+Na2O)]

90

49

HUMID

75 70

30

65

Fresh (source) rock

20

ARID

60

10

55

Strongly altered

0

0

10

20

30

40

50

60

70

80

90

100

CIA (100)[Al2O3/(Al2O3+CaO+Na2O+K2O)]

50 0

5

10

15

20

25

Al2O3+K2O+Na2O

Figure 17. (A) Chemical index of weathering (CIW; Harnois, 1988) plotted against chemical index of alteration (CIA; Nesbitt and Young, 1984). (B) Chemical maturity, expressed as a function of percentage of SiO2 and total percentage of Al2O3 + K2O + Na2O. After Suttner and Dutta (1986) for Sabah sandstones from van Hattum (2005).

deposits. The breakdown and recycling of soft, nonwelded ashes can be very rapid in terrestrial settings with intense precipitation. Lahars are common hazards on many of Java’s volcanoes such as Merapi (Lavigne et al., 2000). Preservation is more likely if the volcanic material is deposited in a marine setting below the storm-wave base, but even this material may be apparently more mature than would be expected in a nontropical environment because of eruptive sorting of material in the ash cloud. The East Java sandstones are well sorted, quartz-rich, and clearly deposited in shallow marine settings, but they were derived from volcanic material produced by explosive eruptions that occurred a short period before their deposition. We know of other quartzrich sandstones in Indonesia that we and others have previously interpreted as having a continental provenance, and our experience of sandstones in Borneo and East Java now causes us to consider more carefully if some of these interpretations may have been wrong. Quartz character and heavy mineral studies (Smyth, 2005; Smyth et al., this volume) are of particular importance for reassessment. In older orogenic belts, similar misinterpretations could easily be made, especially if the ancient climate is not considered or known. CONCLUSIONS: LESSONS FOR ANCIENT ARCS The Indonesian arcs differ in many ways from those commonly portrayed in many arc models. Are the Indonesian arcs typical, or are they one member of a spectrum of arcs? The absence of well-documented major eruptions in the Cenozoic stratigraphic record of Indonesian arcs is likely to reflect their tropical setting and the lack of detailed studies in difficult terrain and relatively

inaccessible areas. An extremely large eruption, possibly on the scale of Toba, terminated the Eocene to early Miocene phase of arc activity in the Southern Mountains of East Java (Smyth, 2005) at ca. 20 Ma, and other such events will surely be recognized in the future. The tropical setting means that some features may be typical only of tropical arcs—for example, the rapidly enhanced maturity of sandstones and enrichment in quartz in volcanogenic debris deposited in arc basins. Nonetheless, in ancient arcs it is still necessary to consider the likely climatic setting before interpreting provenance, both in terms of transport distances and source regions. Other features, such as the importance of extension in arcs, the influence of volcanic arc loading on sedimentary basin development, the critical role of hinge movement in magmatism, and the weakness of the arc, have more general applicability. Certain features of Indonesian arcs could also be of general relevance but could simply reflect particular differences between tectonic processes in SE Asia and those of other arc regions—for example, the absence of a relationship between arc-trench gaps and duration of magmatic arc activity, emplacement of ophiolites, and formation of mélanges. The long history of subduction in the region has resulted in a thin, warm, and weak lithosphere, and although the region is not typical of continental crust it does share many features with other subduction zone backarcs (Hyndman et al., 2005). The behavior of Indonesian arcs may be more representative of arcs in the upper plate of long-lived subduction margins in which the lithosphere becomes unusually responsive to small changes in plate boundary forces. Indonesian arcs appear to have had relatively short lives in comparison with many ancient arcs, but this is probably a

50

Hall and Smyth

A

B

Q CRATON INTERIOR

Continental Block Recycled Orogen

TRANSITIONAL CONTINENTAL

Magmatic Arc Q - Total free quartz F - Feldspar L - Total lithic fragments

DISSECTED ARC

BA

SE

ME

NT

UP

LIF T

RECYCLED OROGEN

TRANSITIONAL ARC

UNDISSECTED ARC

F

1mm

L

D

C

200 µm

200 µm

Figure 18. (A) Quartz sandstones from East Java plotted on a QFL ternary diagram, which could be interpreted as indicating a reworked continental orogenic provenance. The sandstones are actually first-cycle products of rapid reworking of volcanogenic material. Examples of volcanic quartz grains from the Jaten Formation, Pacitan, East Java. (B) Bipyramidal quartz grain. (C) Scanning electron microscope (SEM) image of a quartz shard. (D) SEM image of rounded melt embayments in quartz.

reflection of their young age and the greater ease with which events of different arcs can be resolved in the relatively recent past. Pre-Cenozoic arcs with lives of tens of millions of years may include several phases of arc activity, and several different arcs. The complexity of tectonic evolution in the region shows that a great deal can happen in a 45 m.y. time span, and much of this complexity could be missed in the study of ancient arcs. This could be an explanation of the apparently compressional arc; short-lived contractional events in older arcs may appear to be contemporaneous with arc magmatism when in fact they punctuate arc activity. The history of the Neogene Halmahera Arc shows that an arc can be formed and destroyed in a short

time and may ultimately leave little trace in the stratigraphic record. Even in the young arcs of Indonesia, dating of events is not good, and more and better dating is required to understand and interpret arc processes. We also need to know more about the deep structure of arcs. For example, very little is known about the crust between the arc and trench in any Indonesian arc, and almost nothing is known of the thickness and character of the crust beneath the volcanic arc. Most arcs in the western Pacific are similarly poorly known. At the very least, the record of the Indonesian arcs shows that some features of arc models need to be questioned and that our understanding of tectonic processes in volcanic arcs needs to be improved.

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Printed in the USA

The Geological Society of America Special Paper 436 2008

Carbonate-platform facies in volcanic-arc settings: Characteristics and controls on deposition and stratigraphic development Steven L. Dorobek* Department of Geology & Geophysics, Texas A&M University, College Station, Texas 77843, USA

ABSTRACT Shallow-marine carbonate facies from volcanic-arc settings provide an important, but commonly overlooked, record of relative sea-level change, differential subsidenceuplift, paleoclimate trends, and other environmental changes. Carbonate strata are thin where volcanic eruptions are frequent and voluminous, unless shallow, bathymetric highs persist for long periods of time and volcaniclastic sediment and erupted materials are trapped in adjacent depocenters. Carbonate platforms and reefs can attain significant thickness, however, if subsidence continues after volcanic activity ceases or the volcanic front migrates. The areal extent of shallow-marine carbonate sedimentation is likewise affected by differential tectonic subsidence, although carbonate platforms are most laterally extensive during transgressive to highstand conditions and when arc depocenters are filled with sediment. Tectonic controls on shallow-marine carbonate sedimentation in arc depocenters include (1) coseismic fault displacements and associated surface deformation; (2) longwavelength tectonic subsidence related to dynamic mantle flow, flexure, lithospheric thinning, and thermal subsidence; and (3) large-scale plate deformation related to local conditions of subduction. Depositional controls on carbonate sedimentation in arc depocenters include (1) the frequency, volume, and style of volcanic eruptions; (2) accumulation rates for siliciclastic-volcaniclastic sediment; (3) the frequency, volume, and dispersal paths of erupted material; (4) (paleo)wind direction, which influences both carbonate facies development directly and indirectly by controlling the dispersal of volcanic ash and other pyroclastic sediment, which can bury carbonate-producing organisms; (5) the frequency and intensity of tsunami events; and (6) volcanically or seismically triggered mass-wasting events, which can erode or bury carbonate strata. Regarding platform morphologies in arc-related settings, (1) fringing reefs or barrier reef systems with lagoons may develop around volcanic edifices throughout the long-term evolution of volcanic arcs; (2) local reefs and mounds may build on intrabasinal, fault-bounded highs within underfilled forearc, intra-arc, and backarc basins; (3) isolated platforms with variable platform margin-to-basin transitions are

*Current address: Maersk Oil & Gas, Esplanaden 50, DK 1263 Copenhagen, Denmark. Dorobek, S.L., 2008, Carbonate-platform facies in volcanic-arc settings: Characteristics and controls on deposition and stratigraphic development, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 55–90, doi: 10.1130/2008.2436(04). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Dorobek common in “underfilled” and tectonically active depocenters; and (4) broad ramps and rimmed carbonate shelves are typically found in tectonically mature and sediment-filled depocenters. Keywords: carbonate platforms, reefs, carbonate facies, subsidence, basins, depocenters, carbonate stratigraphy.

INTRODUCTION Shallow-marine carbonate facies may constitute an important part of the stratigraphy that forms in volcanic-arc settings, although they have received relatively minor attention from geoscientists (cf. Dickinson, 1995, 2001; Smith and Landis, 1995; Soja, 1996; Wilson and Bosence, 1997; Nunn, 1998a; Wilson, 2002; Wilson and Lokier, 2002; Wilson and Vecsei, 2005; Bosence, 2005). Shallow-marine carbonate facies in volcanic-arc settings are important recorders of sea-level changes, differential subsidence across arc depocenters, paleoclimate, and other environmental conditions. Thus, careful analysis of these strata is critical for reconstructing the complicated geological histories of volcanic-arc depocenters. Significant petroleum accumulations are also stored in carbonate reservoirs in some volcanic-arc depocenters (e.g., Indonesia backarc basin; Sharaf et al., 2005), although these rather limited examples may be associated with special petroleum-system conditions. This paper presents a review of the tectonic and depositional controls that influence shallow-marine carbonate systems associated with volcanic arcs along convergent margins. Examples of carbonate systems from modern and ancient settings are used to illustrate key principles, although it is important to note that relatively few, thoroughly documented examples exist from which to draw conclusions or construct predictive models. Detailed stratigraphic information for carbonate successions that formed in a variety of arc depocenters during different times in Earth history is simply not available. Instead, the main objective of this paper is to describe how tectonic deformation, other depositional controls, and siliciclastic and volcaniclastic sedimentation can influence shallow-marine carbonate facies development in arc settings. More complete documentation of carbonate successions from modern arc settings and from the rock record will undoubtedly lead to refinement of the basic models presented here, but I hope this paper serves as a general guide to the analysis of arcrelated carbonate systems. TECTONIC CONTROLS ON ARC-RELATED CARBONATE SYSTEMS Tectonic deformation along convergent plate margins strongly influences shallow-marine carbonate depositional systems in volcanic-arc depocenters. Differential tectonic subsidence or uplift in any tectonic setting can be attributed to (1) local, fault-controlled deformation, which can involve fault systems

that only cut upper crustal levels to truly lithosphere-scale fault systems; and (2) long-wavelength, basin-scale differential subsidence and uplift, which typically involve lithosphere-scale deformation processes (e.g., flexure, lithospheric thinning, or thermal subsidence) or dynamic mantle flow (i.e., dynamic topographysubsidence; cf. Gurnis, 1990a, b, 1991). Convergent margins with active subduction, however, are unique tectonic settings where coseismic displacements along great subduction faults may cause long-wavelength (>100 km) surface deformations that can significantly affect depositional systems on the upper (or overriding) plate. This style of differential tectonic subsidence and uplift is exclusive to active subduction zones and was dramatically illustrated during the December 2004 Sumatran earthquake (Fig. 1). These various scales of tectonic deformation typically interact simultaneously and in highly complex ways in volcanic-arc settings, making the subsidence and uplift histories of arc depocenters difficult to characterize, predict, and understand. Other tectonic characteristics of volcanic-arc settings can also affect carbonate sedimentation, such as large-scale aspects of subduction, plate motion, and plan-view geometries of plate margins. The following section first introduces large-scale tectonic controls on arc depositional systems and progresses to consider more local tectonic deformation and its effects on deposition in arc depocenters. Intra-Oceanic Island Arcs versus Continental-Margin Arc Systems Whether a volcanic arc forms as an intraoceanic or a continental-margin arc system is of first-order importance for shallow-marine carbonate systems. Intraoceanic island arcs form on the leading edge of an overriding oceanic plate. In contrast, continental-margin arc systems form where a slab of oceanic lithosphere is subducted beneath an upper plate of continental lithosphere. Arc-related basins along both types of plate margins may be similar, although the initial topography of the overriding plate and rheological differences between oceanic and continental lithosphere of the upper plate impose different conditions on basin development. Volcanism associated with intraoceanic arc systems initially occurs in submarine environments. Successive submarine eruptive events cause progressive shallowing of volcanic edifices until they eventually build to within the photic zone, and shallow-marine carbonate deposition can begin, which may take place after millions of years of eruptive activity. In

Carbonate-platform facies in volcanic-arc settings

14°N

Extensional, Contractional, Strike-Slip, or Neutral Arc Settings

Andaman Islands 12°N

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um Sumatra at

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28 March 2005 epicenter

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Figure 1. Distribution of coseismic uplift and subsidence across the Andaman Islands and Western Indonesian forearc after the December 2004 and March 2005 major earthquakes. Displacement along the subduction fault between the subducting Indian Ocean plate and the overriding Indonesian forearc propagated rapidly northward from the epicenters (indicated by stars) toward the Andaman Islands. Areas of coastal uplift are indicated by the stippled pattern. Areas of coastal subsidence are indicated by dark shading. Dashed line represents interpreted subsidence hinge line. Data obtained from SAR interferometry (InSAR) analyses. Modified from Japan Geographical Survey Institute Web site.

Convergent margins can be classified according to the regional styles of deformation that affect the upper plate (cf. Dewey, 1980). Various sectors of the upper plate can be characterized by extensional, contractional, strike-slip, or very minor (i.e., neutral) deformation. Extensional deformation can develop within (1) the thermally weakened basement of the arc massif, (2) the backarc region, or (3) the forearc region, especially where there is significant tectonic erosion during subduction (Clift et al., 1998). Most deforming volcanic arcs have at least local strikeslip fault zones, even where contractional or extensional styles of deformation are dominant. In compressional arcs, thrust and reverse faults typically form in the backarc region on the upper plate; these faults are typically antithetic to the dip of the subducting slab. Compressional retroarc foreland basin settings are not considered here, although carbonate facies in extensional backarc basins are discussed in a later section. It also is important to note that some dip-trending sections across subduction margins show compressional deformation in one segment of an arc system but other styles of deformation in other parts of the system. For example, the Sumatran forearc region is generally considered to be an “accretionary” forearc (Clift and Vannucchi, 2004), yet the arc itself is cut by the Sumatran Fault, which is a major, through-going, strike-slip fault system that extends along the entire length of Sumatra. Although the Sumatran forearc region is generally thought to be accreting, active zones of transtensional faulting cut at high angles across the forearc basin and partition the Sumatran forearc into multiple depocenters. These closely juxtaposed styles of tectonic deformation are described here simply to demonstrate that it can be inappropriate to make general characterizations about the overall tectonic style of a particular arc system without examining long-term deformation and subsidence histories across the entire system, from the trench to the backarc region, and understanding the kinematic history of the deformation. Accurate, large-scale kinematic analyses may be impossible in strongly tectonized, ancient arc settings. Margin Geometries, Relative Plate Motion, and Changes in Subduction

contrast, continental-margin arc systems may actually begin as subaerial settings and require either a significant eustatic rise or tectonic subsidence to submerge them to photic depths, when shallow-marine carbonate deposition can begin. Although both oceanic and continental lithosphere will bend in response to the growth of surface loads like volcanic arc systems, they have different flexural rigidities so their flexural response will differ. Continental lithosphere also may have more inherent strength variations than oceanic lithosphere because of its typically long and complicated geologic history. These variations in strength and the presence of preexisting faults can influence basin development in continental-margin arc systems.

Plate-motion vectors and the plan-view geometry of continental margins (i.e., before and during convergence and collision) also influence subduction or arc collision along specific margin segments (Dewey, 1980; Thomas, 1983; Bradley, 1989). The initiation of subduction or collision usually will be diachronous along strike where the motion of one plate is highly oblique to another plate margin. If either plate is rotating about a vertical pole, then the rate of subduction or shortening caused by collision typically increases with distance from the pole of rotation. The presence of salients or recesses along either plate margin will also influence which segments along either margin are first to collide.

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The angle of convergence or subduction also strongly influences the styles and distribution of deformation within the overriding plate. Where the relative motion of both plates is nearly perpendicular to the trench axis, and trench rollback is slower than the velocity at which the overriding plate approaches the subduction zone, compressional styles of deformation should dominate in the overriding plate (cf. Dewey, 1980). In contrast, oblique subduction characteristically produces some type of transtensional, transpressional, or pure strike-slip deformation in the overriding plate. The style and geographic location of upperplate deformation (e.g., within the arc-trench gap, the volcanic arc, or in the backarc region) and volcanic activity depend on (1) the angle of convergence; (2) the dip of the subducted slab; (3) whether any seamounts, active spreading-ridge segments, fracture zones, or other rough surface features on the subducted slab arrive at the subduction zone; (4) the degree of mechanical coupling between the subducting and overriding plates; and (5) the location and orientation of lateral strength variations or preexisting fault zones in the overriding plate, which might be reactivated during convergence. Other mechanical or rheological attributes of plates or microplates involved along the convergent margin may influence deformation in the overriding plate but probably in less significant ways. The timing of collision and styles of deformation associated with convergent margins will change if plate-motion vectors vary over time. Changes in the relative motion of plates also cause intraplate stress trajectories to be reoriented over time, and consequently the styles and patterns of deformation within plate interiors should evolve, especially where preexisting basement fault systems are reactivated by in-plane stress. Magmatic activity in the overriding plate will also be affected by changes in the nature of subduction (cf. Otsuki, 1989). Where the subducted slab has a shallow dip, the location of the volcanic front generally is farther inboard from the trench, or the margin may lack substantial volcanic and magmatic features altogether where “flat-slab” subduction occurs. The overriding plate may also be deformed where the dip of the subducted plate becomes very shallow (nearly horizontal) and there is mechanical coupling between the subducted plate and the underside of the overriding plate (e.g., Neogene deformation of the Chilean-Argentinean foreland; Jordan et al., 1983; Jordan and Allmendinger, 1986). If the dip of the subducted slab steepens over time, trench rollback and extension may occur across much of the arc (Uyeda and Kanamori, 1979; Otsuki, 1989). The plan-view geometry of the overriding plate may also control curvature of the subducting plate, downdip from the subduction zone, which in turn may influence deformation across backarc regions in the upper plate (cf. Cahill and Isacks, 1992). Where arc-continent or continent-continent collision has occurred, subduction will usually become very limited or may cease altogether. In these cases, the subduction zone may jump farther outboard and also change dip (or polarity) so that continued plate convergence can be accommodated by continued subduction of oceanic lithosphere. Where continental plates or arc

microplates meet along a convergent margin, some blocks might be “extruded,” or pushed laterally away from the zone of convergence (e.g., Tapponnier et al., 1982; Coward, 1990; Maynard et al., 1997), although the amount of extrusion and lateral translation of blocks away from the collision zone varies. Significant shortening may continue to occur along the convergence zone, but lateral translation of fault-bounded blocks may be required to accommodate additional convergence once the crust within the collision zone cannot be tectonically thickened anymore. Regional strike-slip-fault networks that bound the translated blocks accommodate the extrusion. The sense of slip on these faults may change over time as one block progressively plows into the adjoining plate and more inboard blocks move laterally to accommodate additional shortening in the collision zone. Tectonic Drivers of Long-Wavelength Differential Subsidence in Arc Settings Volcanic-arc settings are characterized by rapidly changing and complexly distributed subsidence patterns, which reflect the interactions of multiple tectonic drivers that operate along subduction margins. These drivers of differential long-wavelength subsidence generally involve deformation or long-term rheological changes of the entire lithosphere, and include flexure, thermal subsidence, magmatic underplating, and thinning in extensional arcs (Fig. 2). Dynamic mantle flow can also influence long-wavelength subsidence patterns in the upper plate (i.e., dynamic topography; Gurnis, 1990a, b, 1991). Long-wavelength subsidence patterns will influence the distribution, plan-view dimensions, and ultimate thickness of carbonate platforms, with the most widespread development of carbonate facies generally occurring during waning stages of subduction, when tectonic subsidence rates are progressively slowing and arc depocenters are more likely to become filled with sediment. Deformation in the forearc region, especially within the subduction complex, can result in other types of long-wavelength subsidence or uplift that are not related to lithospheric deformation. The subduction complex can deform in multiple ways, including complex internal thrusting, backthrusting, normal faulting, and gravitational sliding. These different types of deformation are thought to occur as the subduction complex tends to maintain a wedge shape or “critical taper” (Davis et al., 1983; Dahlen et al., 1984). Underplating and subduction erosion also can cause long-wavelength surface deformation across the top of the subduction complex (Fig. 3). Deformation of the subduction complex can, in turn, cause differential subsidence in the outer part of the forearc basin because the subduction complex commonly forms the seaward limit of this depocenter. This differential subsidence affects shallow-marine carbonate deposition in the outer part of the forearc basin when the basin is filled with sediment and the basin floor is within the photic zone. Along convergent margins where the subduction complex is absent or poorly developed (e.g., with sediment-starved trenches or where subduction erosion dominates), depositional

Carbonate-platform facies in volcanic-arc settings

59

DS Fsc

DS Fsc

Tva

Ffbf

Fbabf

(little or no thermal subsidence of old trapped oceanic crust)

A

Tva

Ffbf

0

100 km

Tbaoc Tbaoc

Fbabf

0

B

100 km

DS

DS Fsc

Fsc

Tva

Frfb

Ffbf

C

0

Tra

Tva

Tba

Ffbf

200 km

D

0

Basin Fill

Dynamic Subsidence

Continental or Arc Crust

Flexural Subsidence

Oceanic Crust

Thermal Subsidence

200 km

Figure 2. Causes of long-wavelength (>100 km) tectonic subsidence across the overriding plate for different types of convergent margins. Flexural subsidence patterns (indicated by dashed black lines): Fsc—flexural subsidence of subducted plate caused by load of the subduction complex; Ffbf—flexural subsidence of forearc region caused by sediment load within the forearc basin; Fbabf—flexural subsidence of backarc region caused by sediment load within the backarc basin; Frfb—flexural subsidence of retroarc foreland basin region caused by sediment load and bordering thrust belt. Thermal subsidence patterns (indicated by dotted lines): Tva—thermal subsidence caused by cooling magmatic arc, after arc becomes extinct; Tbaoc—thermal subsidence of oceanic crust created along backarc spreading center; Tra—thermal subsidence caused by cooling of remnant magmatic arc in extensional backarc settings; Tba—thermal subsidence caused by cooling of rifted backarc continental or arc crust. DS—long-wavelength dynamic subsidence across retroarc settings (indicated by solid black line). In-plane stress may also affect any flexural subsidence patterns, depending on whether stress is compressive or tensile (in-plane stress effects are not shown here). Note how various tectonic subsidence mechanisms may interfere with subsidence patterns caused by other drivers. Patterns are highly generalized. Any of these drivers of long-wavelength subsidence can be counter-affected by shortening or tectonic inversion across the overriding plate. See text for additional discussion. (A) Intra-oceanic volcanic arc with old oceanic crust trapped in backarc region. There may be little or no thermal subsidence of the old, trapped oceanic crust. (B) Backarc region underlain by newly created oceanic crust formed by backarc spreading axis. (C) Compressional backarc region (retroarc foreland basin system) underlain by continental crust. (D) Extensional backarc region underlain by rifted continental crust. Synrift subsidence due to crustal thinning is not shown; only postrift thermal subsidence is shown for the extensional retroarc basin.

gradients from the arc to trench tend to be steep for long periods of time, and shallow-water carbonate deposition will be limited mostly to fringing reefs and narrow platform systems around the volcanic arc. Fault-Related Surface Deformation in Arc Depocenters Displacement along individual faults in arc settings can cause shorter wavelength surface deformation. This surface deformation dramatically influences carbonate deposition if the deformation occurs in shallow-water environments. Although

a thorough description of surface deformations caused by fault displacements is beyond the scope of this paper, typical coseismic surface deformations are described in Table 1. Important aspects of these surface deformations include the following: 1. Coseismic surface deformations are geologically instantaneous and can amount to several meters where fault displacements are greatest, which is typically found at the midpoint of fault surfaces along which displacement occurs. 2. Recurrence intervals of earthquakes that cause significant surface deformation in arc depocenters may be several

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A

uplift of trench-slope break caused by back-rotation of imbricate thrust sheets; local uplift patterns controlled by individual thrust slivers

B

uplift of subduction complex over zone of underplating; uplift geometry controlled by geometry of duplex system

C

uplift of trench-slope break caused by back-thrusting as subduction complex attempts to maintain critical taper; local fault-related folding

decades to several thousand years. Even with comparatively long recurrence intervals, surface subsidence and uplift can accrue to many tens or even hundreds of meters within a few thousand or tens of thousands of years, which will have obvious effects on carbonate facies development. Longer earthquake recurrence intervals may allow carbonate platforms to “recover” and fill new accommodation space that was created during an earthquake. 3. Coseismic displacements on great subduction faults can cause large areas of carbonate platforms and reefs in arc depocenters to subside or to be instantly uplifted above sea level (Fig. 1). Coastal and intertidal environments will be most greatly affected. These events generally affect the forearc region only, where mechanical coupling between the subducting and overriding plates is strongest. 4. Coseismic fault displacements can generate tsunamis or cause mass wasting that can obliterate or bury large areas of shallow-marine carbonate deposition. 5. Some faults may creep aseismically, and although their related surface deformations may slowly develop, significant amounts of uplift or subsidence can still accrue over time. Aseismic fault creep may be most common in subduction complexes where faults cut weakly lithified, highly sheared, fluid-rich materials. Active faults in arc depocenters can control carbonate-platform morphology, platform margin-to-basin profiles, and lateral facies changes during deposition. Active faults can segment preexisting carbonate platforms or control the location of carbonate

backthrusting

D

regional trench-slope adjustment and local fault-controlled subsidence

extensional faulting

E

subsidence caused by tectonic erosion or sediment subduction; compaction has similar effect

tectonic erosion beneath subduction complex

sediment subduction

Figure 3. Possible mechanisms for internal deformation and evolution of subduction complexes, with associated short- to intermediatewavelength surface deformation. Local surface deformations commonly are expressed as complex patterns of uplift, which are superimposed on longer-wavelength patterns of tectonic subsidence. Figure modified from Dickinson (1995). (A) Accretionary offscraping of successive imbricate thrust panels (Seely et al., 1974). For growing subduction complexes, back-rotation of imbricate thrusts causes progressive uplift of crest of subduction complex (i.e., trench-slope break region). There also may be local fault-related folding, which creates antiformal surface deformations. (B) Selective subduction and underplating of thrust duplexes (Sample and Fisher, 1986). Underplating at the base of the subduction complex can cause uplift over the zone of underplating. Structural styles related to underplating are often poorly imaged on seismic profiles but probably are most likely expressed as a duplex zone. (C) Backthrusting of subduction complex toward forearc basin (Silver and Reed, 1988). Uplift of trench-slope break and distal part of forearc basin may occur as the subduction complex maintains a critical taper. Backthrusting in arcward side of subduction complex generally causes uplift in trench-slope break region, with local surface deformation caused by fault-related folding. (D) Denudational normal faulting. Both regional subsidence and local fault-controlled subsidence may occur across trench slope as the subduction complex maintains a critical taper via extensional deformation. (E) Subduction erosion along décollement beneath subduction complex (cf. von Huene et al., 1982). Subduction erosion or sediment subduction along underside of subduction complex creates subsidence across subduction complex. Compaction of subduction complex may have similar effects.

Carbonate-platform facies in volcanic-arc settings Table 1. TYPICAL COSEISMIC SURFACE DEFORMATIONS FOR DIFFERENT FAULT TYPES Fault type Fault length Surface upliftinvolved subsidence Normal faults 5–50 km 0–3 m Reverse/thrust faults 5–100 km 0–5 m Subduction faults 100–1000+ km 0–5 m Strike-slip faults: 1–50 km 0–3 m Transfer fault 50–500+ km 0–3 m Major intraplate or plate boundary

61

The bathymetry of arc depocenters can be rapidly modified by subaerial or submarine deposition of volcanic materials or large-scale mass-wasting events. Arc settings are also seismically active during active subduction, and large-magnitude earthquakes commonly generate significant ground displacements, so slope failure occurs frequently. Although erosional and depositional modifications to shallow-water depositional profiles are relatively common in arc depocenters, they are (1) episodic and thus difficult to predict and correlate in stratigraphic successions, (2) variable in lateral extent, and (3) involve highly variable amounts of rock materials.

and stratal patterns (Caron et al., 2004). Elongated, tide- and storminfluenced balanid (barnacle) shoals grew on local antiformal highs near the crest of the subduction complex. In contrast, arc-attached, low-gradient shelf and ramp profiles were built along the arc side of the forearc basin. Carbonate strata are more sheetlike along the flank of the arc, are generally transgressive in character, and were dominated by epifaunal bivalves. These characteristics are thought to reflect a higher flux of sediment from the adjacent arc. The transport mechanisms and volume of material involved in sediment gravity flows will control the size of carbonate platforms or reefs that might be buried during a single event. The frequency of catastrophic depositional events within a particular segment of the arc will determine whether adjacent carbonate platforms and reefs can recover from a series of events. If relatively low volumes of volcanogenic material are involved in any catastrophic event, and the events occur with low frequency or are widely dispersed along strike, then carbonate platforms and reefs may recover and even utilize lava flows or volcaniclastic facies as substrates. Mass wasting can also occur at any time along steep volcanic edifices, although these catastrophic events are triggered more often by volcanic earthquakes or by inflation of the volcano’s surface as magma ascends into subsurface magma chambers prior to an eruption. Antecedent drainage networks on the flanks of a volcano may control dispersal of low to moderate volumes of volcanogenic material, although dispersal of large-volume, density-stratified pyroclastic flows are generally less affected by antecedent topography. The flow paths of only the basal, highdensity part of pyroclastic flows may be diverted by antecedent topography, such that the resultant coarse-grained deposits are confined to ravinement floors on the flanks of a volcano. These high-density flows then spread laterally, where they become unconfined in lower hillslope regions where antecedent valleys are less incised and surface gradients flatten. In contrast, the lowdensity upper parts of pyroclastic flows are commonly stripped from the basal, high-density part of the flows. The low-density parts of pyroclastic flows can flow over topographic obstructions and lay down more areally extensive deposits. Low-density pyroclastic flows can also flow for several kilometers across seawater. Lava flow paths will also be controlled by antecedent topography, unless the volume of lava exceeds the ability of valleys and hillslope ravinements to contain the flows.

Sediment Flux to Arc Depocenters

Underfilled versus Filled Arc Depocenters

Sediment flux from the volcanic arc or uplifted arc-massif rocks can have many effects on carbonate production in adjacent depocenters. High sediment flux can completely overwhelm carbonate-producing benthic organisms. Lesser amounts of sediment flux might lead to mixed carbonate-siliciclastic-pyroclastic successions. The caliber of sediment supply from volcanic-arc and arc-massif rocks can also determine the types of carbonate sediment produced. For example, late Pliocene carbonate facies in the New Zealand forearc show major lateral changes in sediment types

Sedimentary basins are often described as “underfilled” or “filled” (cf. Covey, 1986; Flemings and Jordan, 1990; Jordan, 1995), which describes whether depositional systems deposit sediment into a depocenter or carry sediment past a former depocenter. Although a thorough discussion of this concept is beyond the scope of this paper, it is useful for understanding where and when widespread carbonate platforms might form in arc depocenters. Although unique conditions of subduction can determine subsidence and basin-filling patterns in arc systems, arc depocenters

facies tracts wherever displacement rates exceed the ability of carbonate facies to fill newly generated accommodation space on either side of a fault. Carbonate lithofacies deposited around faults that were active during deposition may contain greater amounts of gravity-flow deposits, tsunamites, or other catastrophic-event beds. DEPOSITIONAL CONTROLS ON CARBONATE SEDIMENTATION IN ARC SETTINGS The accumulation of noncarbonate sediments and eruptive products has important effects on carbonate deposition in arc depocenters. Various depositional and erosional processes that are characteristic of arc settings can also affect carbonate deposition. The relative influence of these factors on carbonate deposition may depend on the developmental stage of the arc system. Erosional and Depositional Modifications to Bathymetric Profiles

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are generally more likely to become filled (1) during waning stages of subduction when tectonic subsidence slows down, (2) in arc depocenters with high sediment accumulation rates, or (3) in arc depocenters that begin near base level at the onset of subduction. Previously deposited siliciclastic sediment and erupted materials in filled depocenters can provide broad substrates for shallow-marine carbonate deposition. Areally extensive carbonate platforms most commonly form when rising sea level floods these substrates. In addition, tectonic surface deformations become more important for shallow-water carbonate depositional systems in filled depocenters. If a filled depocenter is submerged to water depths that are within wave base and typical tidal ranges (<20 m water depth), even subtle tectonic surface deformations will significantly affect shallow-water carbonate depositional systems.

system, whereas the leeward side of the arc may be dominated by muddier, low-energy carbonate facies. Volcanic arcs can also have significant orographic influence on precipitation, with windward sides of the arc being wetter overall than the leeward sides. Higher precipitation on the windward side of an arc may mean greater chemical and physical weathering, with a greater flux of sediment, especially finer grained sediment, to adjacent marine depocenters. High weathering, erosion, and transport of sediment to the windward side of an arc can greatly suppress carbonate production. Prevailing winds can also greatly influence the dispersal of erupted pyroclastic material, especially volcanic ash. Leeward depocenters may have much greater accumulations of volcanic ash (Sigurdsson et al., 1980), which also can suppress carbonate production.

Oceanographic Circulation Nutrient Supply The types of carbonate sediment that accumulate in arc depocenters will be influenced by the location of the arc system with respect to global climatic zones or oceanographic circulation systems. Cool-water carbonate facies (James, 1997) will be more typical of high-latitude arcs or where cold, upwelling water reaches shallow-water parts of an arc system. Cool-water carbonate deposits typically consist of the skeletal remains of heterotrophic and light-independent biota that are dominated by suspension feeders (e.g., bryozoans) and filter feeders (e.g., barnacles), which are extremely sensitive to the amount of suspended sediments and dissolved nutrients in the water column (Henrich and Freiwald, 1995). Thus, extensive cool-water carbonate production may be difficult to establish close to high-latitude arc systems because of high sediment flux from the arc. Some arc systems may also have unique physiography with respect to oceanographic circulation so that anomalous water masses either enter oceanic passageways or upwelling cold-water masses impinge on shallow-marine settings and influence carbonate facies development over large parts of the arc system. For example, the present-day Gulf of California, Mexico, is a backarc setting where active seafloor spreading has created an open passage at the southern end of the gulf. Seasonal variations in upwelling cause latitudinal variations in nutrient levels within the Gulf of California (Halfar et al., 2004). In the southern part of the gulf, carbonate production is oligotrophic-mesotrophic and coral reef–dominated. The central gulf is dominated by mesotrophic-eutrophic, red-algal sediments, and the northernmost parts of the gulf are dominated by eutrophic, molluscan-bryozoan sediment assemblages. Wind Direction Paleowind direction directly influences the distribution of high-energy, wave-agitated carbonate-facies belts within carbonate platforms (Eberli and Ginsburg, 1989). Volcanic arcs can build significant topographic relief such that if the arc lies in a climatic belt with strong prevailing winds, wave-agitated facies tracts may preferentially build on the windward side of the arc

High nutrient levels in seawater suppress carbonate production (Hallock and Schlager, 1986). A high nutrient flux from arc systems may be related to intense chemical weathering of arc rocks, eruption products, or, in the case of modern settings, anthropogenic input, and can have major deleterious effects on carbonate production. Upwelling of cold, nutrient-rich deep water can also suppress carbonate production or influence carbonate-sediment types in some arc settings (Halfar et al., 2004). CARBONATE DEPOSITION WITHIN ARC DEPOCENTERS Shallow-marine carbonate platforms and reefs can form in forearc, intra-arc, and backarc settings. Various drivers and expressions of tectonic subsidence and uplift affect each setting and may vary as the arc system evolves. Temporal and spatial changes in differential tectonic subsidence have the greatest effects on carbonate sedimentation when arc depocenters are nearly filled to sea level. Depositional processes and sediment flux from the arc system also vary between arc depocenters, which also can influence carbonate deposition. The carbonate-platform classification scheme of Read (1985) is used throughout this paper. Stratigraphic intervals that record progressive surface deformation during deposition are growth strata (cf. Vergés et al., 2002). Carbonate growth strata are deposited around or on top of actively growing tectonic structures and can be used to reconstruct both the kinematic history of individual structures and basin-scale patterns of tectonic subsidence and uplift. Carbonate Platforms and Reefs across the Forearc Region Many forearc regions are characteristically too deep for shallow-marine carbonate sedimentation, especially during early stages of subduction and basin development. Thus, extensive shallow-marine carbonate platforms are not likely to develop

Carbonate-platform facies in volcanic-arc settings across the forearc region until deformation or basin filling has provided suitable substrates for carbonate deposition. Using the classification scheme of Seely (1979) and Dickinson and Seely (1979) for forearc settings, and assuming that the forearc region is at the proper (paleo)latitude, carbonateplatform development is possible where (1) the crest of the subduction complex forms a shoal water or emergent ridge (Figs. 4, 5); (2) the forearc basin is largely filled with sediment so that a broad, shallow, shelf-like substrate extends from the volcanic arc toward the trench-slope break (Fig. 5); (3) the arc massif in front of the volcanic arc is shallow and provides a suitable substrate for carbonate sedimentation; and (4) material flux (i.e., volcaniclastic sediment, ash fall, lava flows) from the volcanic arc does not suppress shallow-marine carbonate deposition across the forearc region. These conditions are similar to those required for carbonateplatform development in other tectonically active basins. Siliciclastic and volcanic flux must be sufficient to fill depocenters to sufficient levels so that regional substrates are available for carbonate sedimentation under the proper sea-level conditions (generally high-frequency transgressive events) but not overwhelm and terminate carbonate production. Until a forearc basin is nearly filled with sediment and a suitable shelf-like profile is constructed, only structurally high parts of the subduction complex or frontal part of the volcanic arc will be shallow enough for carbonate-platform development. Carbonate Platforms and Reefs on Subduction Complexes The first area where shallow-water carbonate facies may develop across forearc settings is the subduction complex (Figs. 4, 6). Most carbonate bodies associated with ancient subduction complexes consist of highly recrystallized and sheared pelagic carbonate facies that have been tectonically accreted and incorporated into the subduction complex. Carbonate-capped seamounts on the subducting plate or carbonate strata on continental or arc microplates may also be obducted or accreted to the frontal part of subduction complexes (e.g., Permian Nabeyama Formation, central Japan, Minoura, 1992; Daiichi-Kashima Seamount, Japan Trench, Cadet et al., 1987; Shin Tani, 1989;

Trench Slope

Trench Fill

Slope Basins

? ? 20

Subducting Slab of Oceanic Crust

Subduction Complex

10 km 10 20 30 40 50 km

Eratosthenes Seamount, Mediterranean Sea, Galindo-Zaldívar et al., 2001). Tectonic uplift and growth of the subduction complex, however, can also raise parts of the subduction complex into the photic zone and possibly even above sea level, creating substrates for shallow-marine carbonate deposition. Reefs and associated carbonate facies can fringe emergent parts of the subduction complex, such as in emergent islands of the western Indonesian forearc. Variably sized isolated platforms can also form on shallow, submerged parts of subduction complexes. Carbonate-platform sequences that form on subduction complexes are rarely more than a hundred meters thick, and apart from fringing reef systems, the platforms commonly have ramplike profiles that mimic the depositional gradients across the top of the subduction complex. Carbonate-platform sequences on subduction complexes typically do not thicken or steepen over time and then evolve into rimmed platform morphologies, which reflects the relatively short life of the platforms, the rapidly changing patterns of uplift and subsidence across the top of the subduction complex, and the deleterious effects of volcaniclastic material supplied by the volcanic arc. Carbonate platforms and reefs on top of subduction complexes may be rare in the rock record because (1) the arcward side of the subduction complex may not become shallow enough for carbonate-platform development until the subduction complex has become tectonically uplifted or the forearc basin has become nearly filled with sediment, which are conditions likely to develop only for short periods of time in the development of forearc regions with highly accretionary subduction complexes (Clift and Vannuchi, 2004); or (2) thin carbonate-platform strata that formed on the subduction complex are misinterpreted as allochthonous units that have been tectonically incorporated into the subduction complex. Recognizing carbonate facies that formed on shallow to emergent parts of highly deformed subduction complexes requires careful age dating and identification of depositional contacts (typically angular unconformities). Older carbonate sequences that form on top of, but are subsequently incorporated into, actively deforming subduction complexes will likely have highly sheared contacts with other lithotectonic units in the subduction complex.

Arc - Trench Gap Trench Slope Break

Volcanic Front Forearc Basin

Eroded arc rocks Pluton

?

?

? ?? ??

63

Arc Massif

Arc Magma Source

Figure 4. Schematic cross section of the forearc region. Stippled pattern represents undeformed or only mildly deformed forearc basin fill. The forearc basin shown here is a largely filled depocenter. Crust beneath the forearc basin may be highly variable in composition but is shown in this diagram as mafic, probably oceanic crust. Some forearc regions may not have the well-developed subduction complex that is depicted in this cross section. Modified from Dickinson (1995).

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UNDERFILLED

FILLED

Tsb

Tsb Tr

Tr

A. Sloped (slope basin)

B. Sloped (slope basin)

Tsb

Tsb Tr

Tr C. Ridged (submerged ridge)

D. Terraced (deep marine)

Tsb

Tsb Tr

Tr E. Ridged (shoal-water ridge)

F. Shelved (shallow marine)

Tsb

Tr G. Ridged (emergent ridge)

Tsb

Tr H. Benched (terrestrial)

Figure 5. Morphologic variants of forearc basins, with basin fill stippled and vertically exaggerated for clarity. Left-side diagrams show underfilled forearc basins (most common for island arcs); right-side diagrams show filled forearc basins (most common for continentalmargin arc systems). Symbols: dashed lines—sea level; Tr—trench; Tsb—trench-slope break (coincides with shelf-slope break for shelved forearcs). Modified from Seely (1979) and Dickinson and Seely (1979); from Dickinson (1995).

Quaternary examples. Quaternary carbonate reefs and associated facies fringe emergent parts of modern subduction complexes in the forearc regions of Sumatra (Simeulue, Nias, Siberut, Sipora, and Enggano islands), the Solomon Islands (Tetepare, Rendova, Ranongga islands), and the Lesser Antilles (Barbados). There is little published information on the thickness and lithologies of offshore Holocene reefs and related facies in these examples. Many modern subduction complexes are partially emergent. Pliocene to Quaternary reefs and other shallow-water facies may crop out as a series of emergent coastal terraces across the crest of the subduction complex. These terraces record both eustatic sea-level changes and progressive uplift of the subduction complex (Mesolella et al., 1969; Wheeler and Aharon, 1991; Mann et al., 1998; Nunn, 1998a, b, 2000; Dickinson, 2001). In fact, uplift

rates for subduction complexes can be derived by comparing the elevations of well-dated, subaerially exposed Quaternary reef terraces to estimates for eustatic sea-level curves for Quaternary time (Fig. 7). “Chimneys” of precipitated carbonate minerals have also been documented from dredge samples and by observations made from submersible vehicles across the surface of many modern subduction complexes (Kulm et al., 1986; Kulm and Suess, 1990; Haggerty, 1991; Hattori et al., 1995). These irregular columns of precipitated carbonate are associated with fluid seeps that discharge from permeable fault zones that cut to the top of the deforming subduction complex (cf. Moore and Vrolijk, 1992). The seeps provide nutrients for chemosynthetic fauna, including a complex microbial community. Abiogenic and biologically mediated carbonate precipitation occurs as the cold, methane-rich pore fluids discharge from the episodically transmissive faults. These carbonate chimneys are not limited to photic depths and can form anywhere along the top of the subduction complex as long as it is above the carbonate compensation depth. Volumetrically, however, carbonate chimneys represent only a small fraction of the total sediment accumulation within the subduction complex. Ancient examples. Pre-Quaternary examples of shallowmarine carbonate facies deposited on subduction complexes are rarely described in the literature. Ancient seep-related carbonate chimneys are reported in Eocene strata of the Barbados subduction complex (Larue and Suess, 1985). One example is provided by Upper Cretaceous(?) to middle Eocene fringing reefs and carbonate ramp sequences that formed at intermittent times during development of the Central American forearc, Costa Rica (Lundberg, 1982; Seyfried et al., 1991; DiMarco et al., 1995; Kolarsky et al., 1995). The Upper Cretaceous(?) (Campanian–Maastrichtian) Barra Honda platform is a cryptic, <100-m-thick, mud-rich platform sequence with rudist reefs on its windward side that formed on remnants of an underlying Cretaceous arc-forearc complex (Seyfried et al., 1991). DiMarco et al. (1995) suggest that the Barra Honda platform strata may be largely Paleocene in age. Kolarsky et al. (1995) also interpreted the basement beneath the Paleogene forearc region of southern Costa Rica and adjacent Nicaragua and Panama as uplifted Cretaceous oceanic crust or seamounts. These contrasting interpretations are mentioned here to illustrate how difficult it can be to interpret the history of these complex settings. Regardless, the Barra Honda platform underwent episodes of shallow submergence and carbonate accumulation, followed by intermittent subaerial exposure and karstification, which have been attributed to superimposed eustatic fluctuations and tectonic uplift caused by accretionary growth of an underlying subduction complex (Seyfried et al., 1991). Continued development of the Central American arc-forearc system during latest Cretaceous to Paleocene time was associated with further structural and bathymetric differentiation of the subduction complex (or “outer arc” of Seyfried et al., 1991), forearc basin, and volcanic arc. Foraminifer-rich ramp facies prograded from both sides of the subduction complex during Late Cretaceous to late Paleocene uplift

Carbonate-platform facies in volcanic-arc settings SUBDUCTION COMPLEX DEPOZONE platform types: • isolated platforms & buildups substrates: • bathymetric crest of subduction complex • may be smaller buildups within subduction complex; probably more common in arcward parts • may extend arcward into forearc basin if distal part of forearc basin is filled

2 km 1

65

FOREARC BASIN DEPOZONE platform types: • small isolated buildups & reefs on structural highs • may be faulted or structurally complicated by later deformation events substrates: • intrabasinal structural highs • transpressional to transtensional structures may be typical

? ? ?

VE ~ 10 10 20 30 40 50 km

Oceanic crust ‘Transitional’ forearc crust Arc massif Subduction complex Trench fill Forearc basin fill Carbonate platform/reef facies

ARC-FLANK & FOREARC BASIN DEPOZONE platform types: • small isolated buildups & reefs • arc-fringing platforms & reefs • typically thin (<100 m thickness) • ramps to low-relief rimmed shelf profiles • platforms tend to become wider as forearc basin fills with sediment substrates: • faulted arc basement • volcaniclastic substrates during later stages of filling

Figure 6. Schematic cross section of forearc region, showing generalized types and locations of carbonate platforms. The forearc basin shown here is largely filled with sediment. Underfilled conditions would characterize the forearc basin during early stages, with smaller platforms and buildups on structural highs or building from shallow, submerged parts of the arc massif. Crust beneath forearc basin may be highly variable in composition, but is shown in this diagram as mafic, probably oceanic crust. Note how carbonate platforms across the forearc region tend to become more areally extensive over time as the forearc basin fills with sediment or as the subduction complex grows, both of which provide progressively wider substrates for shallow carbonate sedimentation. Basin model modified from Dickinson (1995).

of the outer arc. Eocene mixed carbonate-siliciclastic strata (Mal País Formation) contain foraminiferal limestone units up to several tens of meters thick. Strata of the Mal País Formation also overlie, with angular unconformity, older strata that draped basement rocks of the Nicoya Peninsula. Nicoya basement rocks are thought to represent an exhumed subduction complex by some workers (cf. Lundberg, 1982). Carbonate Platform and Reefs in Forearc Basins Carbonate platforms and reefal buildups make up much of the stratigraphic fill in some forearc basins. If the forearc basin is at the proper (paleo)latitude, tropical carbonate facies can build away from the seaward side of the volcanic arc or arc massif (Fig. 6). This is commonly the bathymetrically shallowest part

of the arc-trench system when subduction is active and the volcanic arc is being constructed by volcanic activity. Carbonateplatform facies can accumulate to several hundred meters thick and extend for tens to several hundred kilometers along strike. Volcaniclastic strata typically separate platform sequences within the forearc basin fill and reflect pulsed eruptive activity or episodic reorganization of dispersal systems that deliver volcaniclastic sediment to the forearc basin (Beaudry and Moore, 1985; Dickinson, 1995). Carbonate deposition in forearc basins is influenced by processes that operate at different time scales. At relatively short time scales (i.e., 105–106 yr), thin fringing reefs and narrow platform systems can form around volcanic edifices at virtually any time in the history of a volcanic arc and forearc basin, as long as

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submerged parts of volcanic features are within the photic zone and sediment flux from the arc does not overwhelm carbonate production. Development of laterally extensive and relatively thick carbonate platforms, however, requires both longer periods of time, probably on the order of several million years, long-term tectonic subsidence, and the continued availability of shallowwater substrates. These conditions are more likely to develop during later stages of forearc-basin evolution, when tectonic deformation and arc volcanism are diminishing, forearc subsidence rates are slowing, and the forearc basin is largely filled with sediment (cf. Dickinson, 1995). Proximity to the volcanic arc means that the forearc basin will be greatly influenced by sediment flux from the arc. During earlier phases of forearc-basin evolution, volcanic eruptions and seismic events may be more frequent. Catastrophic eruptions or mass wasting of volcanic edifices can instantly terminate carbonate systems in forearc basins, either by burying carbonate-

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producing organisms or by incorporating previously deposited carbonate strata into mass flows. Platform types. The locations and morphologies of carbonate platforms found in forearc basins will be strongly controlled by the evolving topography-bathymetry across the basin, which depends on the evolutionary stage of the basin. In immature, underfilled forearc basins, fringing reefs and narrow, steep-sided platforms may form only along the trench-facing side of the arc and arc massif. During intermediate stages of forearc basin evolution, isolated mounds and reefal buildups might form on remnant or actively forming structural highs across any part of the forearc basin; intrabasinal structural highs are most common in forearc basins that are segmented by transcurrent strike-slip faults. For example, transtensional deformation of the late Miocene Hellenic forearc basin on Crete created half-grabens where footwall highs became sites for local patch-reef development within the basin (ten Veen and Postma, 1999).

Carbonate-platform facies in volcanic-arc settings As the forearc basin is filled with sediment and the frontal part of the arc massif is eroded, regional depositional gradients generally flatten, and broad substrates develop where shallowmarine carbonate systems can develop. During mature, filled stages of forearc-basin development, carbonate systems commonly construct ramp or low- to moderate-relief shelf profiles that prograde into the forearc basin from erosional remnants of the volcanic arc (cf. Beaudry and Moore, 1985; Dickinson, 1995). Carbonate sedimentation, however, may be terminated at any time across a forearc basin during major eruptive events or during massive hillslope failure that force carbonate-producing environments to be reestablished elsewhere within the basin. Short- to long-term tectonic uplift of the arc and arc massif may cause carbonate environments to become subaerially exposed and either terminate carbonate deposition or force the carbonate factory to move elsewhere. Local seep-related carbonate buildups have also been identified in some ancient forearc basins. Examples include the Jurassic Fossil Bluff Group, Antarctica (Kelly et al., 1995) and the Eocene Joes River mélange, Barbados (Larue and Suess, 1985). Quaternary examples. Many modern forearc basins from tropical settings are bordered by extensive fringing reefs and carbonate platforms that built away from the volcanic arc or arc massif. For example, Quaternary fringing reefs and narrow rimmed platforms are found along segments of the eastern Indonesian forearc basin, where sediment flux from the volcanic arc is relatively low. In contrast, modern carbonate facies are much less extensive in forearc depocenters along the western Indonesian archipelago because there is greater sediment flux from the volcanic arc. Carbonate deposition is suppressed in the forearc basins of Sumatra and Java because the modern volcanic arc has sufficient relief to cause significant orographic effects on precipitation. The winter monsoon blows eastward across the forearc basins of Sumatra and Java and brings heavy rainfall to southwest-facing slopes of the western Indonesian volcanic arc. Meteoric runoff actually lowers salinity in the Java Sea to ~30 ppt during the winter monsoon season. The seasonal decrease in salinity and increase in the amount of suspended sediment delivered to the shallow shelf areas around Sumatra and Java suppress carbonate sedimentation. Thus, modern fringing reefs are geographically limited to the northwestern and southeastern tips of Sumatra and a few offshore islands within the proximal forearc basin. Subsurface data suggest that volcaniclastic flux from the arc had less influence on carbonate sedimentation during early Miocene time, when carbonate shelves prograded away from the Sumatran arc massif (Beaudry and Moore, 1985), which might reflect a weaker monsoon at that time. Modern microatolls (i.e., massive, ring-shaped coral heads up to a few meters in diameter) provide a high-resolution record of subtle uplift and subsidence across the Sumatran forearc basin (Fig. 8; Sieh et al., 1999; Zachariasen et al., 1999). Modern coral heads in shallow-water areas surrounding the outer ridge of the Sumatran forearc basin grow upward and outward until they reach annual lowest low tide (Zachariasen et al., 1999), at which point

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they can only grow outward. Relative sea-level changes at any given location across the Sumatran forearc region then shape the internal growth structure of these microatolls. Coral heads growing on stable substrates develop nearly flat upper surfaces (Sieh et al., 1999). When subjected to gradual submergence, the living coral colony along the outer parts of these microatolls develops a raised rim, whereas their older, dead interiors become depressions. Microatolls subjected to gradual emergence typically develop conical shapes because growth is limited to progressively lower and outer parts of the expanding structure. Careful analysis of growth bands within coral heads from the western coast of Sumatra and the outer-arc islands of western Indonesia showed temporal and regional patterns of submergence and emergence across the Sumatran forearc over the last several hundred years, with microatoll uplift during the giant earthquake of 1833 serving as a key correlation horizon (Sieh et al., 1999; Zachariasen et al., 1999). The microatoll record suggests rapid submergence over several decades prior to the 1833 earthquake. Rates of pre-1833 submergence ranged from 5 to 11 mm/yr and increased trenchward across the Sumatran forearc (Zachariasen et al., 1999). The microatolls record coseismic uplift during the 1833 earthquake and can be observed at six sites along a 125-km length of the subduction zone, with uplift on the order of 1–2 m. Similar to the interseismic submergence patterns, coseismic uplift also apparently increased toward the trench. The great Sumatran earthquakes of December 2004 and March 2005 produced significant, instantaneous uplift and subsidence across the Sumatran forearc (Fig. 1). The differential coseismic subsidence and uplift were expressed over a 300 km distance from the trench axis and >12° of latitude (>2000 km); modern coral reefs and coastal deposits were uplifted tens of centimeters to >2 m locally. A regional, arc-parallel hinge line apparently separates uplifted from subsided parts of the forearc, with uplifted regions on the trenchward side of the forearc basin. This uplift may be preserved in the rock record as a regional disconformity within shallow-water strata. It would be difficult to recognize whether these disconformable surfaces in forearc strata were related to coseismic surface deformation or high-frequency sea-level falls. Regardless, the recent Sumatran earthquake provides an important reminder of the rapid uplift and subsidence that can occur in forearc settings. Other modern forearc basins where Quaternary carbonate facies are found include (1) the forearc basin along the southeastern side of North Island, New Zealand (Kamp and Nelson, 1988), and (2) the island of Espiritu Santo, Vanuatu (Wells, 1988; Johnson and Greene, 1988; Greene et al., 1988; Cabioch et al., 1998). North Island, New Zealand. In the modern forearc basin of North Island, New Zealand, thin veneers of carbonate sediment are being deposited over only ~3% of the forearc region (Kamp and Nelson, 1988). This limited area of modern shallow-water carbonate sedimentation reflects the combined deleterious effects of the relatively cool-water conditions of offshore eastern New Zealand, the high-energy wave-swept conditions of this forearc setting, and the flux of siliciclastic sediment from North Island.

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Figure 8. (A) Schematic cross sections that depict the effect of relative sea-level changes on coral growth. The outermost solid-black growth band depicts the most recent coral growth band. Solid gray bands are dead coral growth bands that grew under the same sea-level conditions as the living coral. Dashed growth bands are dead bands that grew under different, prior sea-level conditions. (a) Hemispherical coral head below highest level of survival (HLS), with a new skeleton accreting in a concentric fashion on the outside of the head. (b) The same coral head under the same sea-level conditions after it has reached HLS. No additional upward growth is possible, although lateral accretion continues along the sides of the coral head that are below HLS. (c) “Hat” morphology of a coral head that grew up to HLS, but then underwent a drop in HLS. The part of the coral exposed above HLS has died. Lateral growth below the new HLS develops a lower outer rim around a higher center. The elevation difference between the two flats is a measure of the amount of emergence. (d) “Cup” morphology of a microatoll that underwent a sea-level rise after the coral had been growing at HLS. The coral grows upward toward the new HLS, constrained only by its growth rate. Upward and outward growth over the old HLS surface produces a raised outer rim, indicative of submergence. The elevation of the new HLS is not recorded by the coral until the coral grows up to it. From Zachariasen et al. (1999). (B) Cross section of a microatoll from the Indonesian forearc. The X-rayed thin slab reveals a clear record of annual growth bands that expanded radially outward (from left to right) at ~1 cm/yr. The HLS of the coral during the past 35 yr is recorded in the topography of the coral’s upper surface. The arrows track the apparent rise of sea level in the 1960s and its subsequent apparent fall. Growth patterns in corals such as these can provide high-resolution records of true sea-level change or the effects of incremental coseismic subsidence or uplift. From Sieh et al. (1999).

Carbonate-platform facies in volcanic-arc settings Beyond the modern shelf, carbonate sediments (<0.5 m thick) cover the tops of isolated, elliptical banks (10–50 km2 in area) that rise from the slope to depths of 150–500 m (e.g., Madden Banks). The banks may represent the submarine expression of surface deformation on top of the subduction complex. Carbonate sediments that form thin veneers on these submarine ridges consist of coarse skeletal sand (dominantly a cool-water bivalveforaminifer assemblage with lesser coral, brachiopod, gastropod, and echinoid fragments). Patchy mosaics of coarse skeletal sand are also found on the shallow shelf updip from the slope banks. These skeletal-sand accumulations on the modern shelf largely consist of bryozoans, corals, barnacles, bivalves, and benthic foraminifers and are associated with local patches (a few to tens of square kilometers in area) of gravelly sediment on rugged, current-swept seafloor topography. Espiritu Santo, Vanuatu. Holocene carbonate sedimentation around the island of Espiritu Santo, Vanuatu, occurs above actively uplifting volcanic-arc basement (cf. Cabioch et al., 1998). The Cenozoic tectonic history of Espiritu Santo is complex (Meffre and Crawford, 2001). The Vanuatu (or New Hebrides) archipelago can be divided into Western, Central, and Eastern Belts (Greene et al., 1988), which record discrete and geographically separated phases of volcanic activity. The Western Belt, which includes Espiritu Santo, consists largely of upper Oligocene to middle Miocene calc-alkaline lavas and associated volcaniclastic rocks that formed the volcanic arc above a westward-dipping but currently extinct subduction zone to the east. Subduction along the eastern side of Vanuatu ceased during late Miocene time and switched to the western side of the archipelago, where an eastward-dipping subduction zone now exists. Late Miocene to Pliocene volcanic rocks of the Eastern Belt record the volcanic arc that formed above the east-dipping subduction zone. Late Pliocene–Holocene volcanic rocks of the Central Chain record another westward shift of the volcanic arc toward the New Hebrides Trench. Remnants of the older Oligocene–Miocene volcanic arc of the Western Belt

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now form the basement rocks of the present-day Vanuatu forearc region. A subduction complex is absent along this margin owing to its young age and the limited amount of sediment delivered to the trench. The central part of the Western Belt, including Espiritu Santo, has been actively uplifting throughout Quaternary time owing to subduction of the d’Entrecasteaux Ridge at this location. The uplift history of Espiritu Santo is recorded by a series of emergent and notched Holocene reef terraces that were deposited on the uplifting Oligocene–Miocene volcanic substrate (Fig. 9; Cabioch et al., 1998). Although we know the present-day tectonic setting of Espiritu Santo, and the Holocene reef terraces should be described as having formed in a forearc setting, it might be difficult to make the same tectono-stratigraphic interpretation if similar carbonate terraces were found in the rock record. More likely, these reef terraces would be described as “intra-arc basin” deposits (see discussion below about intra-arc carbonate facies), simply because they are found stratigraphically above deformed, arc-related rocks. Seismic profiles from offshore Espiritu Santo also show drowned Quaternary carbonate reefs and reef caps (1– 6 km across, some reefs up to 300 m relief, reef tops at ~300 m water depth) and “mounded” sediment bodies (375–450 m water depth) (Fig. 10; Johnson and Greene, 1988; Greene et al., 1988), which formed on submerged and complexly deformed parts of the relict arc basement. The locations of some reefs and shelf margins are probably fault controlled. These highly variable patterns of Quaternary carbonate sedimentation around Espiritu Santo reflect the complex differential subsidence and uplift that can occur across young, incipiently formed forearc regions. In this case, Espiritu Santo represents subaerial remnants of an old volcanic arc that were deformed and became basement for the forearc region of a younger volcanic arc. It is worth noting that Quaternary carbonate facies apparently are extremely rare within forearc basins along the westfacing, convergent margins of North and South America, even

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Figure 9. Cross section of the uplifted Tasmaloum fringing reef, Vanuatu, showing the 6, 8, 10, 12, 14, 16, 18, and 20 ka time lines. Only the most characteristic dates are reported (those in italics are 14C dates converted to calendar years). A 6 ka reef flat is now 40 m above sea level, which shows the rapid uplift rates that have characterized this forearc region over the last 6000 yr. From Cabioch et al. (1998).

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Figure 10. (A) Main geologic and tectonic features of the New Hebrides Arc. SMFZ—Santa Maria fracture zone; ABFZ—Aoba fracture zone; AMFZ—Ambrym fracture zone; EFZ—Epi fracture zone; VFZ—Vulcan fracture zone. Modified from Greene et al. (1988). (B) Main geologic and tectonic features of the central New Hebrides Arc, Espiritu Santo region. Basement rocks of Espiritu Santo consist of late Oligocene to middle Miocene volcanic and volcaniclastic rocks erupted from a paleo-volcanic arc (Western Belt rocks) that formed above a now-extinct, westward-dipping subduction zone. During late Miocene time, however, the subduction zone switched to the western side of the New Hebrides Arc (present-day New Hebrides Trench), and an east-dipping subduction zone developed beneath arc rocks of the Western Belt, which includes Espiritu Santo. The active volcanic arc to this subduction zone consists of volcanic centers that make up the Central Chain, to the east of the Western Belt. Thus, Western Belt rocks now constitute the basement of the modern forearc region of the Central Chain, and Espiritu Santo is in a transitional tectonic state. Depocenters that formed during extension of the older Oligocene–Miocene arc (Western Belt rocks) might be considered intra-arc basins. The volcanic centers of the Western Belt are no longer active, however; so the present-day depocenters should be considered part of the forearc of the Miocene–Holocene volcanic arc (i.e., the Central Chain). Modified from Greene et al. (1988). (C) Interpreted seismic-reflection profile across the East Santo and North Aoba basins, New Hebrides arc system. Note Quaternary pinnacle reefs–buildups along the outer, fault-controlled margin of the East Santo shelf. Modified from Greene et al. (1988).

those margin segments that lie within tropical climatic zones. The limited amounts of modern carbonate sedimentation along these margins probably reflects the strong coastal upwelling systems that bring cold, nutrient-rich bottom water onto the shallow, narrow shelf areas that border the western sides of these forearc basins. In contrast to modern continental margins, extensive Paleozoic carbonate platforms formed along the eastern side of the paleo–Pacific Ocean because marginal oceanographic barriers (e.g., island arcs, orogenic belts, and microplates) may have protected the carbonate platforms from the deleterious effects of upwelling (Whalen, 1995). Quaternary carbonate reefs and shallow-platform facies are also found in some relict forearc basin settings where subduction is no longer active but where the basin’s bathymetry and structural features were inherited from previous times of active subduction. In some ways, these basins might be considered forearc successor basins. For example, the Indispensable Basin of the eastern Solomon Islands has been interpreted as either a relict forearc basin or the leading edge of the partially subducted Ontong-Java Plateau (Vedder and Bruns, 1989). The Ontong-Java Plateau resisted southwestward-directed subduction when it collided with the Solomon Islands arc in middle to late Miocene time, forcing a change in subduction polarity and initiation of a younger, northeastward-directed subduction zone beneath the southwestern side of the Solomon Islands (Vedder and Bruns, 1989). Tectonic structures and remnant bathymetry along the relict forearc side of the Solomon Islands continue to influence modern sedimentation there, even ~15 m.y. after the change in subduction polarity. For example, seismic profiles along the southwestern slope of the Indispensable Basin show several drowned Pleistocene(?) reefs at ~300–500 m water depths along a steep, faulted(?) escarpment just southeast of Santa Isabel Island (Niem, 1989). The steep bathymetric gradient across the forearc region and the low relief and relatively small surface area of the relict volcanic arc suggest that this relict forearc basin will probably remain underfilled for a long time into the future, thus limiting the potential surface area for shallow-marine carbonate sedimentation. Bathymetric profiles around offshore parts of Viti Levu, Fiji, are structurally controlled and record older Neogene to possibly Quaternary deformation. Faults that cut the older Neogene forearc basin and

frontal arc massif influenced the location, dimensions, and morphology of Neogene to Holocene reefs across the area (Fig. 11; Rodd, 1993). Ancient examples. Ancient examples of forearc-basin carbonate sequences include (1) Paleogene carbonate platforms on transpressional ridges that transected the forearc basin of Costa Rica (Krawinkel and Krawinkel, 1996); (2) Oligocene to Miocene reefs and platform facies (Tau, Wailotua, and Qalimare Limestones) in the Outer Melanesian forearc basin on the island of Viti Levu, Fiji (Hathway, 1994, 1995; Rodd, 1993); (3) Miocene Rethymnon Formation, eastern Crete (Pomoni-Papaioannou et al., 2003); (4) Miocene mounded carbonate buildups to carbonate shelf systems in several sub-basins of the Indonesian forearc (Beaudry and Moore, 1985; Matson and Moore, 1992; van der Werff, 1996); and (5) Pliocene “cool-water” skeletal limestones, Hawke’s Bay, New Zealand (Kamp et al., 1988; Ballance, 1993; Caron et al., 2004). The present-day Terraba belt represents the inverted Tertiary forearc region of Costa Rica (Krawinkel and Krawinkel, 1996). This basin underwent three stages of basin evolution: (1) an initial transpressional phase during Late Cretaceous–Eocene time, which created the transpressional ridges that served as substrates for carbonate-platform facies; (2) a transtensional phase during Oligocene to early Miocene time, when the forearc region was affected by transtensional fault systems and was subsiding rapidly, causing the carbonate platforms to drown; and (3) a second transpressional phase of deformation from middle Miocene to Holocene time. Thus, tectonic deformation across the Costa Rican forearc basin initially created the substrates for carbonate sedimentation but ultimately caused the platforms to drown. Upper Oligocene to Miocene carbonate intervals exposed on the island of Viti Levu, Fiji, and imaged on seismic profiles from adjacent offshore areas consist of thin reefs or platform sequences that are interstratified with volcaniclastic strata (Figs. 11, 12; Hathway, 1994, 1995; Rodd, 1993). The Qalimare Limestone, which crops out on Viti Levu, is at least 300 m thick and consists of coral-algal reefs and mounded facies that formed on the edge of a shallow platform in front of the arc massif. Seismic profiles across the Bligh Water Basin from the northern offshore shelf of Viti Levu have intrabasinal highs that served as substrates for

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Figure 11. (A) Geologic map of Fiji. Most of the platform areas are shallow-water carbonate environments, although some platform areas are presently at subphotic depths (i.e., drowned or incipiently drowned). Also note the steep, typically fault-bounded edges of most platform areas, as indicated by solid black lines. The islands that make up Fiji represent the eroded and faulted remnants of an older volcanic arc–forearc basin system. Bathymetry in meters. From Rodd (1993). (B) Geologic cross section across Fiji and adjacent platform areas. This section shows various isolated platforms and buildups in the relict Bligh Water forearc basin. Note how faults controlled platform locations, dimensions, and morphology. From Rodd (1993).

local late Oligocene to middle Miocene reef development within the basin. The Miocene Rethymnon Formation (<60 m thick) in the Apostoli Basin, eastern Crete, consists of skeletal-rhodalgal limestone facies that were deposited on a warm, temperate ramp that developed along the arcward side of the Neogene forearc basin in Crete. Skeletal and rhodalgal grain types in the Rethymnon Formation are indicative of nontropical carbonate facies, but the presence of large benthic foraminifers suggests warm, temperate conditions (Pomoni-Papaioannou et al., 2003). The Rethymnon Formation onlaps the underlying Pandanassa Formation (alluvial fan and fluvial floodplain deposits) and Apostoli Formation (shoreface sandstone, Herostegina sands, and gray-marl facies), suggesting general shallowing of the Apostoli Basin and possible uplift along the arcward flank of the basin during deposition of the Rethymnon Formation. Seismic profiles across several late Oligocene–Holocene sub-basins of the western Indonesian forearc show carbonate buildups and shelflike profiles of Miocene age that formed on the arcward side of these forearc depocenters (Beaudry and Moore, 1985; Matson and Moore, 1992). Most of the isolated buildups appear as mounded features that likely formed on structural highs during transgressive events. The shelflike carbonate profiles may also record deposition during generally transgressive conditions. Both the carbonate buildups and shelflike platform sequences are overlain by progradational siliciclastic strata supplied from the arcward side of the basin during regressive events. Pliocene strata in the forearc basin on North Island, New Zealand (Fig. 13), include coarse-grained, shallow-water, skeletal-sand sheets, which locally contain steeply dipping clinoform

packages (Kamp et al., 1988). These Pliocene skeletal sand bodies formed as outer forearc basin deposits. The axes of the sand bodies are oriented subparallel to faults and deformation fabrics within the subduction complex farther east, which suggests that the sand-body locations and orientations may have been controlled by complex surface deformations that developed above parts of the subduction complex that lie beneath the outer forearc basin. The skeletal sand bodies were probably deposited during active broad wavelength folding of the forearc basin, as indicated by (Fig. 13): (1) successively older Pliocene limestone strata having progressively steeper dips and greater elevations (>300 m above sea level) along the eastern flank of the basin; (2) successively younger Pliocene carbonate strata, shifted laterally toward the central axis of the forearc basin, especially along the more uplifted eastern flank of the basin; (3) older Pliocene carbonate sand bodies concentrated along the eastern and western flanks of the forearc basin, grading laterally into siliciclastic facies in the basin axis, whereas younger carbonate intervals are thinner and more sheetlike in character; and (4) unconformities within the Pliocene section that merge, becoming composite unconformities on the flanks of the forearc basin (Kamp and Nelson, 1987, 1988; Caron et al., 2004). General tectono-stratigraphic model for carbonate facies in forearc regions. Although detailed studies of carbonate strata deposited across forearc regions are relatively rare, previous studies of the tectono-geomorphologic evolution of forearc regions provide insight into the possible stratigraphic evolution of carbonate systems that might form there. The evolution of carbonate platforms and reefs across the forearc region will be influenced by (Fig. 6):

Sea level

Isolated platform on faulted high

arc uplift = increased sediment flux to forearc basin

c Ar

s ma

s if

Figure 12. Schematic tectonic and depositional model for southwestern Viti Levu during early to middle Miocene time. View approximately toward present northeast. Note how fault-controlled isolated platforms of the lower to middle Miocene Qalimare Limestone were deposited on the frontal side of the volcanic arc, which was rising isostatically because of crustal thickening. This in turn caused more erosion of the arc and greater flux of epiclastic sediment to the forearc basin. From Hathway (1994).

74

Dorobek Age (Ma)

N.Z. Stages

LOCALITIES & LITHOSTRATIGRAPHY, HAWKE’S BAY, NEW ZEALAND Limestone

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Tt

Subduction Complex

Figure 13. Pliocene stratigraphy across the forearc basin of North Island, New Zealand. Measured stratigraphic sections at specific locales across the forearc basin are shown in the upper part of this figure; stratigraphic thickness in meters. Note the association of limestone with siliciclastic units and the progressive displacement of successively younger limestone units toward the basin axis. The cross section does not show all the structures distributed across this forearc basin and implies only broad, synformal folding over time. Most Pliocene limestone was deposited along the outer, trenchward side of the basin, although younger carbonate units shifted progressively westward over time, toward the arc. Also note the progressive decrease in dip of successively younger units (see lower cross-sectional view). These growth stratal patterns reflect local surface deformation above fault-propagation folds and more regional differential subsidence/uplift across the North Island forearc basin, which is expressed as the progressive westward shift of shallow-marine carbonate sedimentation across the basin. New Zealand stage names and their relative ages include: Wn—Nukumaruan (late PliocenePleistocene); Wm—Mangapanian (late Pliocene); Wp—Waipipian (middle Pliocene); Wo—Opoitian (early Pliocene); Tt—Tertiary. No vertical exaggeration. From Kamp and Nelson (1988).

1. Whether the subduction complex is accretionary or nonaccretionary. Accretionary subduction complexes grow laterally as trench sediments are added to the frontal part of the deforming wedge. Thus, accretionary subduction complexes commonly widen and shallow over time and provide ever-increasing potential surface areas for shallow-water carbonate sedimentation. 2. Differential uplift or subsidence across the forearc region. For example, if a formerly submerged subduction complex is uplifted and becomes emergent, any preexisting carbonate platforms or reefs across the crest of the subduction complex will become subaerially exposed, and carbonate depositional environments will be forced to

shift laterally off the flanks of the subduction complex. The total area available for shallow-water carbonate sedimentation may actually decrease, especially if uplift creates steep bathymetric gradients across the top of the subduction complex. If the upper surface of the subduction complex is below photic depths at the beginning, however, uplift may bring the surface into shallow-water depths and allow a carbonate factory to develop. If the top of the subduction complex is already emergent or at shallow water depths, continued uplift of the subduction complex may produce progradational to downstepping stratal patterns within carbonate platforms or fringing reefs that cap the subduction complex. In contrast,

Carbonate-platform facies in volcanic-arc settings regional subsidence of the subduction complex (owing to subduction erosion, flexure of the upper plate, or other tectonic causes) might cause preexisting carbonate platforms and reefs either to drown or to backstep toward any remaining shallow-water parts of the subduction complex. Similar complex chronostratigraphic and facies relationships may develop above other types of actively growing intrabasinal structural highs within the forearc basin or along the proximal arcward side of the forearc basin. 3. Transcurrent faults that can partition differential subsidence or uplift across the forearc region. Transcurrent faults that cut the forearc region are most common where oblique convergence occurs (e.g., Western Indonesian forearc). Transcurrent faults can partition parts of the forearc into antiformal highs or fault-bounded horst and graben segments, which control the location of platforms or isolated reefs (cf. Krawinkel and Krawinkel, 1996). Given the complex subsidence and uplift histories of forearc regions, carbonate-platform and reef strata that form across the forearc region during active subduction will likely be thin (<100 m thick) and may be bounded by either subaerialexposure surfaces or drowning surfaces with complex chronostratigraphic relationships. The thickness and compositions of interstratified noncarbonate facies will be highly variable, depending on: 1. The evolving bathymetry and patterns of differential subsidence or uplift across the forearc region, which may change rapidly over relatively short time scales. It is possible for the forearc region to transition repeatedly between phases of subsidence and uplift, and regions of subsidence and uplift may be closely juxtaposed in space and time. The amount of time that lapses between episodes of uplift or subsidence is another critical factor that determines whether episodic deformation forces carbonate facies tracts to relocate to other shallow-water parts of the forearc region. 2. The flux of volcanogenic and siliciclastic sediment into the forearc basin. Rapidly prograding mixed carbonatesiliciclastic shelf facies may build from the volcanic arc into a forearc basin during later stages of basin evolution when the volcanic arc has built significant relief, the basin is filled with sediment, and the basin is slowly subsiding or uplifting. Progradational siliciclastic facies commonly will be supplied by a series of alluvial and deltaic point sources, where drainage networks reflect local climate and the irregular to ridgelike topography of the volcanic arc. In humid settings, the arc may act almost like a line source of sediment, with many local drainages supplying volcaniclastic sediment to the forearc basin, especially along windward-facing slopes of high-elevation archipelagos. Shallow-marine carbonate sedimentation may be suppressed where there is a large flux of volcanogenic material (either as eruptive

75

products or as epiclastic sediment) to the forearc. Siliciclastic deposits can serve as substrates for shallow-marine carbonate sedimentation during transgressive events if sediment supply is relatively low. Shallow-marine carbonate strata may provide an excellent record of even subtle patterns of differential subsidence or uplift across the forearc region. In many cases, dip-trending transects through mature forearc basins show stratal patterns that indicate that the basins behaved more or less as broad, synformal depocenters, with smaller-scale, structural irregularities superimposed on this broader wavelength deformation. Across the area of the trench-slope break, deformation of the underlying subduction complex typically creates a broad-wavelength (tens of kilometers) uplift that is oriented subparallel to the trench axis. As previously described, wedge accretion and related tectonic deformation may bring the top of the subduction complex into shallow water depths and allow a carbonate factory to develop, or continue until the subduction complex becomes emergent and the area for shallow-water carbonate sedimentation may actually decrease. Uplift of the subduction complex may also force progradational to downstepping stratal patterns within carbonate-platform strata or fringing reefs that were deposited on the crest or flanks of the shallow subduction complex. In contrast, regional subsidence of the forearc region may cause carbonate platforms and reefs either to drown, backstep, or shift to any remaining shallow-water parts of the forearc region. Smaller-scale seafloor irregularities across the trench-slope break may be the surface expression of fault-propagation folds or other surface deformations across the top of independently moving, fault-bounded slivers within the subduction complex (e.g., Pliocene forearc basin, North Island, New Zealand; Kamp and Nelson, 1987, 1988; Kamp et al., 1988). Reactivation of preexisting fault zones in forearc basement rocks or syndepositional deformation of older forearc-basin strata can also create broad wavelength folds that influence facies patterns and growth stratal patterns within more axial parts of the forearc basin. Intrabasinal forearc highs may (1) become substrates for shallow-marine carbonate facies if they reach shoalwater depths, (2) influence dispersal of sediment gravity flows shed from adjacent shallow-water platform areas if the highs are more deeply submerged, or (3) be eroded by marine or subaerial processes as they are uplifted to progressively shallower depths or possibly even above sea level. Thus, the broad-wavelength seafloor characteristics of intrabasinal fold axes would have similar effects on sedimentation as surface deformations across the trench-slope break. Lastly, transcurrent faults can partition the forearc region into separate fault-bounded depocenters with intervening highs. Transcurrent faults may strongly partition differential subsidence or uplift across the forearc region and are especially common where oblique convergence occurs. Vertical offset on the transcurrent faults may create steep bathymetric gradients that control the location of platform margins or isolated reefs on the intervening highs between the faults.

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Carbonate Platform and Reefs within Intra-Arc Basins Soja (1996) reviewed island-arc carbonate sequences, and although the tectonic setting of some ancient examples cited therein is somewhat ambiguous, most probably formed within intra-arc basins or were at least built on volcanic edifices. Islandarc carbonate sequences may characteristically have endemic faunas, because many volcanic arcs are biogeographically isolated from more laterally extensive, land-attached carbonate platforms, where faunal radiation is more likely to occur (Soja, 1992, 1996). It is not clear if a tendency toward endemism has any effect on carbonate-facies types, stratal architecture, or platform-reef morphology. Volcanic arcs may also serve as “refuges” for reef-building faunas when local conditions become stressed or during global extinction events. For example, while most reef-building organisms became extinct during Late Triassic time, a few coral species seemed to have survived into Early Jurassic time in volcanic-arc settings that were subsequently accreted to western North America (cf. Stanley and McRoberts, 1993). Carbonate reefs and shallow-marine facies commonly fringe volcanic edifices within many intra-arc basins, especially if the volcanic arc is at the proper (paleo) latitude. Carbonate reefs and associated facies may be more common on the windward side of volcanic archipelagos because reef-building organisms prefer to construct frameworks on the windward side of islands and landmasses, and abundant ashfall deposition on the downwind side of the arc may inhibit carbonate sedimentation. Thick, areally extensive carbonate-platform sequences are apparently rare in intra-arc basins because (1) these basins typically have steep bathymetric profiles when they are underfilled, so there is little available substrate for shallow-marine carbonate sedimentation; (2) there is abundant volcanic activity that inhibits carbonate sedimentation; (3) there commonly are frequent phases of rapid subsidence or uplift, during which time carbonate platforms either might be drowned or terminated by subaerial exposure; and (4) frequent seismic events may trigger massive failure along the flanks of volcanic edifices, which can either incorporate and destroy large stretches of preexisting platforms and reefs or bury shallowmarine carbonate environments with mass-flow deposits. The thickest and most areally extensive carbonate platforms and reefal buildups within intra-arc basins probably form during late stages of basin evolution or following the last phases of volcanic activity when (1) volcanic edifices are likely to be subsiding rather than being uplifted or undergoing active volcanic construction; (2) the basins are more likely to be filled with volcaniclastic sediment, which will provide additional shallow-water substrates for carbonate accumulation; (3) volcanic topography has been subdued by erosion so that a wider surface area is available for carbonate sedimentation during sea-level highstands; and (4) volcanic eruptions and seismic events have ceased or are much less frequent than during earlier stages of basin development. Carbonate platforms and reefs can form during active phases of arc volcanism as long as volcanic activity constructs features

that are submerged and within the photic zone, and the volume of volcanogenic material dispersed to intra-arc basins is not great enough to suppress carbonate production (cf. Bardintzeff et al., 1985). Polymict breccias, however, dominate the stratigraphy in many intra-arc basins and may provide the only record of past carbonate sedimentation. These breccias can be the dominant lithofacies in intra-arc basin successions that are >5 km thick and consist largely of volcanic and volcaniclastic clasts, with lesser amounts of sandstone, mudstone, and shallow-marine carbonate facies. Platform- and reef-derived clasts (with dimensions up to several hundred meters) within these sedimentary breccias may provide the only evidence for former shallow-marine carbonate systems in many intra-arc basins (cf. Soja, 1990). Intra-Arc Basin Classification and Tectonic Models Intra-arc basins are depocenters within volcanic arcs and are underlain by arc-massif crust. Smith and Landis (1995) classified intra-arc basins as volcano-bounded, fault-bounded, or hybrid types. Volcano-bounded intra-arc basins are depocenters found between constructional volcanic edifices. Volcano-bounded basins associated with continental-margin subduction systems typically have only thin sediment fills, which Smith and Landis (1995) suggested are due to the generally high-standing character and low subsidence rates of their continental-crust substrates. In contrast, volcano-bounded basins from intraoceanic settings may accumulate thick volcanogenic strata, possibly because they subside more rapidly or begin as deep-water depocenters on oceanic crust. Erosion rates, however, are generally much higher in continental-margin arc systems than in intraoceanic arc systems, because continental-margin arc systems are largely subaerial features, whereas volcanic highs and intra-arc basins may be largely submerged in intraoceanic arc systems. Thus, erosion of volcanic highs also affects the amount of sediment delivered to intra-arc basins and the rate at which they are filled. Fault-bounded intra-arc basins are defined by fault networks. Fault patterns within most volcanic arcs and first-motion studies of intra-arc earthquakes indicate that extension or strike-slip deformation causes most fault-bounded intra-arc basins to form. Fault-bounded intra-arc basins typically have bounding faults that trend parallel to the volcanic arc, although some basins and their main bounding fault systems trend transverse to the volcanic arc. Fault-bounded intra-arc basins also may subside more rapidly than volcano-bounded depocenters (Smith and Landis, 1995). Transverse oblique-slip fault systems that delineate some intra-arc basins may accommodate block rotations along the arc axis, such as in the Aleutian forearc region (Geist et al., 1988). The curved trend of the Aleutian arc system, the oblique subduction that occurs along this convergent margin, and the strong coupling between the subducting and overriding plates may explain the block rotations. Differential subsidence of narrow crustal blocks between transverse faults may create submarine channels that funnel sediment gravity flows to the forearc basin or trench slope. Arc-parallel strike-slip faults can also accommodate significant lateral displacement of arc segments on either side of the

Carbonate-platform facies in volcanic-arc settings fault zone (e.g., Semangko or Sumatra Fault, Beck, 1983; Philippine Fault, Karig et al., 1986; Sarewitz and Lewis, 1991). Hybrid intra-arc basins have elements of both volcano- and fault-bounded basins. Hybrid basins are defined by Smith and Landis (1995) as depocenters with fault-defined margins, but where most of the topographic or bathymetric relief is defined by constructional volcanic features. Arc massifs and associated intra-arc basins typically undergo complex deformation histories and differential subsidence as a consequence of changes in subduction parameters (dip angle, obliquity, and rate of subduction), dynamic mantle flow, changes in surface loads, and magmatic activity. These large-scale tectonic and magmatic processes and resultant patterns of uplift and subsidence within the volcanic arc change over time scales on the order of 5–20 m.y. (Smith and Landis, 1995). As a result of the constantly changing tectonic and magmatic conditions that characterize the long-term history of most volcanic arcs, intra-arc basins may transition rapidly between volcano-bounded, fault-bounded, and hybrid types. All styles of deformation have been recognized within volcanic arcs and their related intra-arc basins. Extensional and strikeslip structures are generally more common within the arc massif than contractional structures, unless arc-continent collision has occurred. The history of deformation within intra-arc settings is poorly preserved because it is obscured and overprinted by multiple eruptive and deformation events, thick sediment accumulations may deeply bury and obscure fault networks, or the arc and its related basins are eventually incorporated into collisional orogenic belts and original tectono-stratigraphic relationships are destroyed. Many arc systems also record flips in subduction polarity such that former forearc regions may become backarc or interarc settings, or vice versa. Preexisting faults within the arc massif also may be reactivated by a different sense of displacement during subduction-polarity flips or other changes in subduction characteristics. These complicated tectonic histories may make it difficult to characterize the geometry, distribution, and kinematics of fault networks that develop within a volcanic arc. Extensional deformation is common where trench rollback causes the volcanic arc to split apart. An active seafloor-spreading ridge may form if the arc is completely stretched apart, leaving a remnant arc on the trailing side of the spreading axis and an active arc on the other side. In these cases, intra-arc rift basins evolve into an extensional backarc setting. Intra-arc rifting may occur faster than in intracontinental rift settings, with correspondingly much faster rates of synrift subsidence (Yamaji, 1990). Some volcanic arcs, such as the Marianas arc system, have been repeatedly split by intra-arc rifting and seafloor spreading (Hamilton, 1979). Other arcs have been split apart during a flip in subduction polarity; changes in subduction polarity may not occur synchronously along all segments of the convergent margin. These factors may create problems with basin classification and understanding of the tectonic processes that cause complex spatial and temporal patterns of uplift and subsidence. Examples of Cenozoic volcanic arcs that have been rifted apart during changes in subduction polarity include the Solomon Islands (Johnson, 1979; Kroenke,

77

1984; Bruns et al., 1989) and Vanuatu archipelagos (Carney et al., 1985; Greene et al., 1988). The fault-bounded basement highs that delineate intra-arc basins are parallel or oblique to arc segments. These structural highs typically serve as (1) minor sediment source areas if they are subaerially exposed, (2) substrates for shallow-marine sedimentation if they are only slightly submerged, or (3) bathymetric obstructions to sediment gravity flows if the structural highs are more deeply submerged. Although seismic data are limited from recently rifted arcs like the Solomon Islands and Vanuatu, the arc-parallel or arc-crossing highs that separate intra-arc depocenters within these arcs probably represent oblique transfer zones. Oblique dip-slip displacement on the faults that bound the intraarc highs allows extensional strain to be transferred between normal fault systems within the arc. These intra-arc fault-bounded highs may influence sedimentation within the arc for a long time after they form, especially if extension stops before the arc is completely rifted apart. Transtensional, transpressional, and pure strike-slip styles of deformation within the arc and intra-arc basins are more likely to develop where there is oblique subduction. As the convergence direction becomes more oblique or subparallel to the trend of the subduction system, pure strike-slip fault networks should become more common. Along some modern arc systems with highly oblique subduction angles, regional strike-slip faults may cut through and translate long segments of the volcanic arc (e.g., Sumatran or Semagang Fault along the island of Sumatra, western Indonesia; Philippine Fault, Philippine archipelago). Regional shortening within most volcanic arcs typically does not develop until arc-continent or continent-continent collision occurs and magmatic activity within the arc has long ceased. The limited amount of shortening within many volcanic arcs (and backarc regions) may indicate that mechanical coupling between the overriding and subducting plates diminishes with increasing distance from the trench (cf. Uyeda and Kanamori, 1979). Platform Types Most carbonate sequences within intra-arc basins consist of either fringing reefs constructed on volcanic or plutonic basement or barrier reefs with relatively narrow (<10 km wide) back-reef and lagoonal facies. These fringing reefs and narrow platforms can build relatively steep depositional profiles, as indicated by mountainside exposures of ancient examples and bathymetric profiles across modern intra-arc basins. Fringing reefs also may form within collapsed calderas, although these probably are minor, short-lived accumulations that are terminated or blasted away by subsequent eruptions. Carbonate platforms and reefs also form in fault-bounded intra-arc basins where fault patterns and displacement histories strongly influence platform dimensions, the location of platformmargin facies tracts, and platform morphology. Small, isolated platforms and pinnacle reefs are found on fault-bounded highs within some intra-arc basins of Cenozoic island arcs in the South Pacific. Local fault-bounded basement highs that border some of

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these intra-arc basins serve as substrates for larger isolated platforms. For example, Woodhall (1985) suggested that conjugatefault patterns controlled the locations of volcanic vents and Quaternary reefs on the Exploring Isles of the northern Lau Ridge. Other platform types are rare. Ramp-like to nonrimmed platforms and scattered mounded buildups may develop on the leeward sides of volcano-bounded intra-arc basins. Ramps may still have relatively steep depositional gradients owing to the steepness of the underlying, volcanically constructed antecedent topography. Leeward carbonate ramps may be relatively thin in the rock record because volcanic ash from active volcanoes preferentially accumulates in leeward intra-arc depocenters and suppresses or terminates carbonate sedimentation (cf. Eldredge and Kropp, 1985). Intra-arc platforms and reefs exhibit complex patterns of stratigraphic development, because intra-arc basins are characterized by complex spatial and temporal patterns of differential subsidence and uplift during their evolution. During active tectonic deformation, platforms and reefs may rapidly expand or shrink over short time scales (<100,000 yr), especially during icehouse times when high-amplitude glacio-eustatic sea-level changes are more likely to occur (cf. Read, 1995). Intra-arc fault systems commonly are steeply dipping and develop significant surface scarps (100 to >1000 m), and volcanic edifices also commonly have dramatic bathymetric gradients, both of which limit the area over which carbonate facies tracts can shift laterally during relative sea-level changes. During late stages of basin development, when intra-arc depocenters are more likely to be filled, carbonate platforms and reefs may utilize volcanogenic successions as substrates and develop into more areally extensive systems. Differential subsidence and uplift across arc terranes, however, can still be significant during later stages of basin evolution and can cause complex stratigraphic development within laterally extensive carbonate platforms. Quaternary examples. Quaternary intra-arc carbonate platforms and fringing reefs are found in the Solomon Islands, Fiji, and the Kadavu island group (Nunn and Omura, 1999; Nunn, 2000); Vanuatu (Macfarlane et al., 1988); Lau Ridge (Woodhall, 1985); Tonga Ridge (Cunningham and Anscombe, 1985); the Indonesian archipelago (Sumatra, Java, Bali, islands; Wells, 1988); the submerged Grenadine Bank and other, locally drowned pinnacles and isolated buildups on volcanic substrates elsewhere in the Lesser Antilles (D’Anglejan and Mountjoy, 1973; Bouysse, 1984; Dey and Smith, 1989); the Ryukyu Islands of southwest Japan (Iryu et al., 1995; Nakamori et al., 1995; Sagawa et al., 2001); and various parts of the Philippines. In many cases, Quaternary reefs and narrow fringing carbonate platforms have been uplifted to form a series of notched terraces that are exposed near the modern coastlines of these islands. These uplifted Quaternary limestone terraces attest to the rapid uplift rates (average rates up to 1 cm/yr) that characterize many modern volcanic-arc settings. Many Quaternary reefs or shallow-marine carbonate facies have developed in successor intra-arc basins (e.g., Quaternary

carbonate facies from the Grenadine Bank and some northwestern islands, Lesser Antilles, D’Anglejan and Mountjoy, 1973; Bouysse, 1984; Dey and Smith, 1989; Quaternary Ucuna Limestone of the Neogene Takalau Limestone Group, Lau Ridge, Woodhall, 1985; Quaternary carbonate units from islands of the Western and Eastern Belts of Vanuatu, Macfarlane et al., 1988). In a successor intra-arc basin the volcanic arc has been dormant or extinct for a long time, yet regional topography-bathymetry is still strongly influenced by the inherited tectonic-volcanic geomorphology. In some cases, carbonate facies may be deposited within an arc segment where the volcanoes are extinct or dormant, even though other parts of the volcanic archipelago may be active. Long periods of geologic time (10 m.y. or more) may also pass between major eruptive episodes. This is the case in the Grenadine Islands, Lesser Antilles, where intra-arc carbonate facies form a submerged and incipiently drowned carbonate platform (the Grenadine Bank) that is currently at 20–40 m water depths. The Grenadine Bank probably formed during a Pleistocene lowstand of sea level on eroded remnants of the Lesser Antilles volcanic arc (D’Anglejan and Mountjoy, 1973; Dey and Smith, 1989). Active volcanoes are found within the same archipelago on the islands of Saint Vincent and Grenada, which are to the north and south, respectively, of the Grenadine Bank. Given that the Grenadine Bank is the same general distance from the subduction trench as are the active volcanoes on Saint Vincent and Grenada, renewed volcanic activity should be expected at some future time there. The lack of recent volcanic activity within the Grenadine Islands, however, and the prevailing easterly wind allow extensive modern carbonate sedimentation to occur on the eastern part of the Grenadine Bank (D’Anglejan and Mountjoy, 1973; Dey and Smith, 1989). The Kallinago Depression in the northern Lesser Antilles (Fig. 14) represents a volcano-bounded intra-arc basin that formed when volcanism shifted from an outer (trenchward) Eocene to early Miocene arc to an inner late Miocene to Holocene arc farther west. The Kallinago Depression diminishes and merges southward into a single arc near the island of Guadeloupe. Isolated, early(?) Pleistocene buildups (e.g., Luymes Bank; Fig. 14) have formed on the steep, volcanic substrates surrounding this relatively young active arc. The Lau Ridge is another example where Quaternary carbonate facies are accumulating in a remnant intra-arc setting. The Lau Ridge represents the remains of a late Oligocene–early Miocene volcanic arc that was left behind as a trailing, extinct arc when the Lau Basin opened by seafloor spreading and separated the Tonga Ridge from the Lau Ridge. Emergent to shallowly submerged parts of extinct volcanic centers that compose the Lau Ridge continue to serve as substrates for Holocene carbonate sedimentation (mostly skeletal sand facies; Woodhall, 1985). In the Vanuatu (New Hebrides) arc, Oligocene–Miocene volcanic and volcaniclastic rocks of the Western Belt of the archipelago represent products from a volcanic arc that was rifted apart during middle Miocene (ca. 14 Ma) time. Renewed arc volcanism jumped eastward to form the mostly late Miocene

Carbonate-platform facies in volcanic-arc settings

A’

A Anguilla

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Figure 14. Map of northern Lesser Antilles and Kallinago Depression and interpreted seismic section. The Kallinago Depression is a volcano-bounded intra-arc basin that formed between the inactive Eocene–early Miocene arc and younger late Miocene–Holocene arc of the Lesser Antilles. Note how volcanic axes converge southward to a single volcanic arc (i.e., Martinique). Interpreted seismic profile (location of section A–A′ shown in upper left of base map), showing slightly deformed sedimentary fill within the Kallinago Depression (profile from Bouysse, 1984). Luymes Bank is a drowned carbonate buildup whose top is at ~150–200 m water depth at present. Bathymetry in meters. Modified from Smith and Landis (1995).

to late Pliocene Eastern Belt of the Vanuatu arc. Continued rifting between the two belts ultimately led to the youngest volcanic centers within the Central Belt, which formed volcanically constructed and fault-bounded highs during late Pliocene to Holocene time (Macfarlane et al., 1988). Pliocene to Holocene carbonate facies form fringing reef complexes around extinct volcanoes of the Western and Eastern Belts and around active volcanoes of the Central Belt. Subduction polarity, however, apparently changed during Miocene time from west-dipping to east-dipping subduction, and the subduction zone switched from the eastern to the western side of the arc. Given the complex tectonic history of the New Hebrides arc, it is difficult to classify some of the depocenters as either intra-arc or forearc basins because many are in a protracted, transitional phase of development. Ancient examples. Ancient examples of intra-arc carbonate platforms, small banks, and reefs include (1) the Silurian Heceta Formation, Alexander Terrane, southeastern Alaska (Soja, 1990); (2) the Middle Devonian Kennett Formation, Lower Carboniferous Bragdon and Baird Formations, and Lower Permian McCloud Limestone, Eastern Klamath Mountains, California (Demirmen and Harbaugh, 1965; Watkins, 1985, 1993a, b; Watkins and Flory, 1986); (3) Upper Cretaceous lenticular reef bodies in the

Clifton Member of the Mount Peace Formation, Lucea Inlier, western Jamaica (Grippi, 1980); (4) the Oligocene(?) Paumbapa Formation on Sumba Island and offshore, Indonesia (Fortuin et al., 1997; Rutherford et al., 2001); (5) local mounded buildups that are scattered throughout the Tertiary section on islands and offshore structures of the Tonga Ridge (Cunningham and Anscombe, 1985; Herzer and Exon, 1985); (6) Miocene to Pliocene reefal limestone and carbonate-platform facies (Tokalau Limestone Group) on various islands of the Lau Ridge (Woodhall, 1985); (7) lower to middle Miocene skeletal limestones (Tari Formation and Qalimare Limestone) on Viti Levu, Fiji (Hathway, 1994, 1995); and (8) the Miocene Lelet Limestone, which was deposited on the eroded remnants of the New Ireland volcanic arc, Papua New Guinea (Stewart and Sandy, 1988). These intra-arc carbonate facies typically are complexly interstratified with pyroclastic sands and tuffs, volcaniclastic breccias, coarsegrained pebbly sands, siltstone, mudstone, and lava flows, and also are commonly cut by steeply dipping syn- and postdepositional faults. Other possible ancient examples of intra-arc carbonate sequences include a Lower Jurassic coral reef in the Telkwa Range, British Columbia (Poulton, 1989; Stanley and McRoberts,

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1993); thin, Lower Jurassic coral biostromes in the Chocolate Formation, Arequipa massif, Peru (Wells, 1953; Stanley and McRoberts, 1993); Middle Jurassic coral biostromes from the Wallowa Terrane, west-central Idaho, USA (Stanley and Beauvais, 1990); and Upper Triassic carbonate units (including coral reefs) of the Martin Bridge Formation, Wallowa Mountains, Oregon, USA (Stanley, 1982; Stanley and Senowbari-Daryan, 1986), and at Lime Peak, southern Yukon Territory, Canada (Reid and Ginsburg, 1986). Although the Mesozoic examples from North America are typically described as “island arc carbonates,” their specific paleogeographic settings remain somewhat ambiguous. Unlike Quaternary intra-arc carbonate platforms and reefs, for which only limited subsurface data are available, ancient sequences provide better constraints on the thickness and internal facies of intra-arc carbonate platforms and reefs. Most intra-arc carbonate sequences are several tens of meters thick, although some are several hundred meters thick. Shallow-marine, muddy carbonate facies of the Miocene Lelet Limestone on New Ireland, northeastern Papua New Guinea, may be an exception, as they reportedly are up to 1 km thick, with offshore equivalents of the Lelet Limestone ranging from 1 to 2 km thick (Exon et al., 1986; Stewart and Sandy, 1988). The lateral extent of ancient intra-arc carbonate facies is usually much more difficult to constrain than for Quaternary examples, because ancient intra-arc basins may be subjected to multiple phases of postdepositional deformation.

raised the arc substrate into the photic zone, thin carbonate packages may form rings around the volcanic substrates. With continued arc development (in either intraoceanic or continental-margin arc systems), intra-arc basins may evolve into fault-bounded depocenters, where faults then control the areal dimensions and locations of carbonate platforms and reefs as well as the location of facies tracts and the overall morphology of platform margin-to-basin transitions. Carbonate platforms will likely have steep, fault-controlled margins in profile and have linear segments in plan view, both of which would reflect the influence of underlying fault zones. The trend of many linear platform-margin segments within fault-bounded intra-arc basins might be predictable if the underlying fault zones develop with consistent orientations or as conjugate fault sets that reflect regional stress orientations. As tectonic and volcanic activity wane during later stages of basin development, and intra-arc basins progressively fill with sediment (largely volcaniclastic), carbonate platforms may become thicker and more areally extensive. Remnant volcanicand fault-controlled topography, however, may continue to influence the dimensions of platforms and the locations of platformmargin-to-basin facies transitions until enough basin-filling sediment has filled the antecedent topography.

General Tectono-Stratigraphic Model for Intra-Arc Carbonate Platforms and Reefs The different types of intra-arc basins influence the types of carbonate platforms and reefs that form in them. In general, the rapid rates of tectonic deformation in intra-arc settings, the likelihood of catastrophic eruptions, and the close proximity of highrelief source areas for siliciclastic sediment (i.e., the volcanic arc) cause carbonate sedimentation to be only intermittent and shortlived within intra-arc basins. Fringing reefs and narrow platforms are the likely dominant type of carbonate systems found in most intra-arc settings (Fig. 15). These rarely attain significant thickness (i.e., >100 m thick) before a catastrophic eruption or mass-wasting event buries the carbonate system. Mass-wasting events also might incorporate carbonate strata, causing large-scale disruption or transport of shallow-marine carbonate facies into deeper-water settings. Displaced blocks of reef and shallow-water carbonate facies within deep-water shale and volcaniclastic deposits may provide the only evidence of the former intra-arc platforms. There may be differences between intra-arc carbonate successions from intraoceanic and continental-margin arc systems. In intraoceanic arcs, shallow-marine carbonate facies will likely not appear within intra-arc basins until the volcanic arc has built itself into the photic zone. Volcano-bounded and hybrid intra-arc basins may dominate over fault-bounded basins in young intraoceanic volcanic arcs. The basal part of the basin-filling sequence in these basins will likely consist of deep-water volcaniclastic and pelagic-hemipelagic deposits. After volcanic activity has

Carbonate sedimentation also occurs in backarc settings on a variety of tectonic structures and bathymetric highs constructed by depositional processes. As in other arc depocenters, the areal extent, thickness, morphology, and internal stratigraphy of carbonate platforms and reefs will be determined by the tectonic history of the backarc and whether it is filled with sediment.

Carbonate Platforms and Reefs in Backarc Basins

Backarc-Basin Types and Tectonic Models Backarc basins form on the back side of subduction-related volcanic arcs (Fig. 16). Extensive geophysical surveys of many modern backarc regions over the last 20 yr or so have led to the recognition of three main types of backarc basins (Dickinson, 1974; Karig, 1983; Ingersoll, 1988; Marsaglia, 1995): (1) extensional backarc basins that form by rifting and seafloor spreading within or behind continental-margin or intraoceanic arcs, (2) remnant-ocean backarc basins that form by entrapment of old oceanic crust and are associated with intraoceanic subduction zones, and (3) compressional backarc basins, which are more commonly classified as retro-arc foreland basins (see discussion below). Marine backarc basins may be underlain by continental to transitional crust if they form by extension behind a continental-margin arc system. In contrast, backarc basins will be underlain by highly subsided oceanic crust if the basin forms along an intraoceanic subduction zone, and remnant oceanic crust is trapped behind the newly formed volcanic arc. Backarc basins also can be underlain by young oceanic crust if they form where rifting in the backarc region ultimately leads to seafloor spreading.

Carbonate-platform facies in volcanic-arc settings VOLCANO-BOUNDED INTRA-ARC BASINS platform types: • isolated platforms & buildups, commonly steep-sided • thin fringing platforms & reefs during early stages; complexly interstratified with volcanics, volcaniclastics • may be preferential development on windward side of arc • complex chronostratigraphic relationships between individual carbonate sequences • may be drowned platforms adjacent to actively growing platforms • platforms become larger & thicker as depocenters fill, but eruptions can have catastrophic effects on platforms substrates: • volcanic edifices & volcaniclastic fill

Volcano-bounded intra-arc basins

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FAULT-BOUNDED INTRA-ARC BASINS platform types: • isolated platforms & buildups • similar controls & relationships as for volcano-bounded basins (windward side, complex chronostratigraphy, etc.) • extensional deformation common; post-rift thermal subsidence will affect stratigraphic development • internal growth strata possible if platforms on actively deforming fault-bounded highs substrates: • volcanics, volcaniclastics, faulted arc basement or volcanics/volcaniclastics • platforms become larger & thicker as depocenters fill or as deformation wanes

Fault-bounded intra-arc basins

Figure 15. Types of carbonate platforms and buildups within intra-arc basins. Platform thicknesses, widths, and relief are not to scale. Only generalized platform types and buildups are shown. Isolated platforms and fringing platforms-reefs are most common. Stratigraphic relationships and depositional contacts with structural features and various geomorphologic surfaces across intra-arc basins are also highly generalized. Modified from Smith and Landis (1995) to emphasize carbonate-platform facies.

Extensional backarc basins are commonly associated with subduction zones where trench rollback occurs. Many extensional backarc basins actually begin as fault-bounded intra-arc basins, when a volcanic arc begins to rift apart. An extensional backarc basin might come into existence when there is the first evidence for a topographically expressed, but volcanically extinct, remnant arc that is separated from an active volcanic arc by a rift zone. There likely will be a distinct bathymetric or topographic axis to the newly formed extensional backarc basin at this point in its tectonic evolution. This definition for an extensional backarc basin does not depend on the amount of stretching across the rift zone or on whether rifting has ceased or seafloor spreading has begun. The remnant arc also can be both a source area for siliciclastic sediment and a substrate for shallow-marine carbonate sedimentation, although carbonate facies may be deposited only during early stages of the basin’s history, before the remnant arc has undergone significant thermal subsidence. Backarc basins that are underlain by trapped fragments of older oceanic crust have different tectonic and subsidence histories than those that form by backarc extension. Remnant-ocean backarc basins are associated with intraoceanic subduction zones, and their dimensions depend on where subduction is initiated within the overriding plate and the nearest bathymetric high

(commonly a continental margin) that defines the distal side of the basin (see below). Backarc regions typically are areas of overall crustal extension or strike-slip deformation. Crustal shortening is rare in most backarc regions except for some continental-margin arc systems where the subducting and overriding plates are strongly coupled (e.g., segments of the Andean and western Indonesian backarc regions and the Sea of Japan backarc region). In cases where significant crustal shortening occurs in the backarc region of continental-margin arc systems, the flexural basins that develop are more appropriately described as retro-arc foreland basins, which are not considered further in this paper. The styles and intensity of deformation in the backarc region depend largely on (1) the tectono-evolutionary stage (or maturity) of the arc system; (2) the relative motions (e.g., angle of convergence) between the subducting and overriding plates; (3) the age, thickness, dip, and crustal type of the subducting slab; (4) possible changes in subduction polarity during convergence; (5) whether trench rollback is possible; (6) whether a spreading ridge, seamount, or other large-scale seafloor features have been subducted; or (7) whether arc-continent or continent-continent collision has begun. Extensional backarc regions are dominated by normal-fault networks that are subparallel or slightly oblique to the trend of

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Dorobek old oceanic crust trapped by intraplate subduction zone

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continental crust beneath retro-arc foreland basin (compressional arc system)

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Figure 16. Cross sections through different types of convergent margins, showing crustal types of backarc regions. (A) Intraoceanic volcanic arc, with old oceanic crust trapped in backarc region. (B) Backarc region underlain by newly created oceanic crust created by spreading axis. (C) Compressional backarc region (retro-arc foreland basin system) underlain by continental crust. (D) Extensional backarc region underlain by rifted continental crust.

the volcanic arc. As in other extensional settings, oblique-slip transfer faults trend at high angles to basin-bounding normal faults and accommodate displacement transfer between offset or nonparallel normal faults. This style of extensional deformation is characteristic of the 700-km-long Izu-Bonin intraoceanic arc system, which is considered to be an example of an extensional arc system where an incipient backarc basin(s) is forming (Taylor et al., 1990, 1991; Klaus et al., 1992). The extensional sub-basins along the Isu-Bonin arc are typically asymmetric, 25–110-kmlong and 25–40-km-wide features. Most basin-bounding normal faults trend subparallel to the axis of the arc system. Similar to continental rift systems, oblique transfer faults and accommodation zones link oppositely dipping normal faults that bound the rift depocenters. Submarine volcanic centers appear to be

concentrated along bends in the rift system that are interpreted as accommodation zones. Modern examples of backarc extension behind continental-margin volcanic arcs include the Andaman Sea, the Okinawa Trough, and the Japan Sea. Highly oblique subduction from westernmost Indonesia northward to the continental margin of Myanmar, strong coupling between the subducting Indian-Australian plate and overriding Eurasian plate, and subduction of an aseismic ridge may explain the backarc rifting and seafloor spreading in the Andaman Sea (Eguchi et al., 1980). The Okinawa Trough is an incipient continental backarc basin and displays similar structural features as the Sumisu rift zone (Letouzey and Kimura, 1986). The Japan Sea is a fully developed continental-margin backarc basin, where contraction or transpression has occurred along its eastern side, and subduction may be in its beginning stages (Kikuchi et al., 1991; Tamaki, 1995; Okamura et al., 1995; Takano, 2002). Retro-arc foreland basins are flexural depocenters that result from crustal shortening on the continental side of the volcanic arc; examples include the Andean and western Indonesian “backarc” settings. In comparison, crustal shortening within backarc basins (sensu stricto) along intraoceanic convergent margins is apparently rare. The limited amount of shortening in the backarc region of most intraoceanic volcanic arcs probably reflects weak coupling between the subducting and overriding plates along these types of convergent margins. Platform Types The lateral extent and cumulative thickness of carbonate facies within backarc basins will vary, depending on the tectonic origin(s) and evolutionary stage of the basin (Fig. 17). For backarc basins that form by entrapment of old oceanic crust behind the arc, carbonate facies may be entirely limited to shallow submerged areas on the backarc side of the arc massif, and there may be no other shallow-water substrates for carbonate platforms across the backarc region. Most emergent or shallow-water substrates in these backarc basins will be constructed by volcanic activity. Continued growth of carbonate platforms and reefs in these backarc settings may be frequently interrupted by volcanic eruptions or mass-wasting events so that reefs and other carbonate facies must be reestablished after a catastrophic event destroys or buries them (Eldredge and Kropp, 1985). In contrast, extensional backarc basins will have bathymetric profiles that reflect stretching of the arc massif or other types of backarc lithosphere. During initial rifting of the volcanic arc, carbonate sequences may be deposited in intra-arc basins (see above). Where stretching continues to the point of whole lithosphere failure and a seafloor-spreading ridge develops, a backarc basin is clearly identifiable, and trailing rifted margins develop on either side of the basin. There is no consensus, however, as to when an intra-arc basin structurally, bathymetrically, or sedimentologically transitions into a backarc basin. Apparently there are no preserved examples of thick, longlived carbonate platforms within backarc basins. Backarc basins

INTRA-OCEANIC BACKARC SETTINGS

A shelf/ramp sequence constructed on older volcaniclastic strata

COMPRESSIONAL BACKARC SETTINGS

B

smaller isolated ramp sequence ramp sequence platforms, shoals, and reefs constructed on older constructed on foreland on antiformal highs above foreland basin strata side of basin back-arc thrust belt (commonly post-orogenic)

pinnacle reef or outer reef tract may be fault-controlled

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EXTENSIONAL BACKARC SETTINGS: C. Backarc intracontinental rifts D. Backarc intracontinental rifts leading to spreading and conjugate passive margins E. Intra-arc rifting leading to spreading

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more regional ramp/shelf sequences during post-rift stages; flexural onlap and post-rift thermal subsidence; substrates include post-rift fill and continental basement

may contain syn-rift platforms on fault-bounded highs if depocenters are marine

EXTENSIONAL BACK-ARC BASINS

D

may contain syn-rift platforms on fault-bounded highs if depocenters are marine, but more extensive ramps/shelves form during post-rift phase

thermally subsiding conjugate passive margin; similar controls on regional platform development as passive margins

may contain syn-rift platforms on fault-bounded highs if depocenters are marine, but more extensive ramps/shelves during post-rift phase

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Figure 17. Different types of backarc settings and regional carbonate platforms that may form in them. In all examples shown, the subduction zone and forearc region (not shown) are left of the volcanic arc (indicated by triangle). Horizontal scales are generalized; no vertical scale implied. Only general patterns of platform development are shown; internal stratigraphy and details about platform morphology will vary significantly in each backarc setting. (A) Intraoceanic volcanic arc, with old oceanic crust trapped in backarc region. Carbonate platforms and reefs can form on the backarc side of intraoceanic arcs. In general, substrates for these backarc carbonate systems will be constructional volcanic edifices, eroded arc-massif basement, or volcanic-volcaniclastic deposits. Local fault-controlled bathymetry may control platform dimensions and facies patterns. (B) Retro-arc foreland-basin system. Extensive carbonate sedimentation can occur in retro-arc foreland basins if they are slightly underfilled and the continental foreland is below sea level. Thus, carbonate sedimentation is probably more common during underfilled stages of retro-arc foreland-basin development. Carbonate ramps are common, especially along the distal (foreland) side of the basin. Various styles and intensities of deformation across the foreland can significantly affect carbonate sedimentation. Foreland-basin carbonate systems are discussed in more detail in later sections. (C) Extensional backarc region underlain by rifted continental crust. If rifted continental basement is flooded early, synrift carbonate platforms may form on fault-bounded basement highs. More extensive platforms may develop during late post-rift stages, when the extensional backarc basin is nearly filled and thermal subsidence dominates. (D) Backarc region underlain by newly created oceanic crust created by spreading axis. Where stretching in the backarc region progresses to the point of whole lithosphere failure, a seafloor spreading axis will develop. If stretching began within an intra-arc setting, the trailing remnant arc becomes extinct and undergoes relatively rapid thermal subsidence. The remnant arc can serve as a substrate for carbonate-platform development, although these are usually short-lived platforms. (E) If stretching began within an intra-arc setting, conjugate passive margins would develop on both sides of the newly created backarc ocean basin. Both conjugate margins should behave like other passive margins. Thus, carbonate platforms and reefs should become more areally extensive over time as the margins undergo thermal subsidence and progressively wider potential substrates become available for carbonate sedimentation.

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that are underlain by trapped oceanic crust are deep-water settings almost from their inception, so shallow-marine carbonate facies will not be deposited until the volcanic arc shoals to shallow-water conditions. Even after the volcanic arc has built itself up to shallow-water depths, backarc carbonate facies generally will consist only of fringing reefs or narrow rimmed-shelf platforms that are constructed on the steep-sided volcanic edifice. More areally extensive carbonate platforms can be constructed only after appropriate substrates for shallow-marine carbonate sedimentation have been constructed, which requires some combination of erosion of the volcanic edifice (by both subaerial and submarine processes) and deposition of volcaniclastic sediment and lava flows. Similarly, most carbonate sequences within extensional backarc basins consist of fringing reefs or relatively narrow ramps and rimmed shelves that are constructed (1) on the backarc side of the active volcanic arc and arc massif, (2) for a short time on the remnant arc as it drifts away from the active arc, or (3) on fault-bounded basement highs that partition extensional backarc regions into sub-basins. Postrift thermal subsidence ultimately submerges the drifting remnant arc to subphotic water depths so that any productive carbonate factories on the remnant-arc crust are soon terminated. Scattered pinnacle reefs and mounded buildups that are typically less than a few hundred meters thick may be constructed on synrift topography in backarc regions. These buildups apparently are drowned soon after they form because of the combined effects of rapid subsidence and short-term eustatic sea-level rise. Compressional backarc settings are more like retro-arc foreland basins, so they are not discussed here. Regardless, carbonate platforms and reefs are common in these settings, forming on fault-bounded highs that result from inversion tectonics. Isolated platforms and buildups are common morphologies. Quaternary examples. Quaternary carbonate platforms and reefs are found in the backarc basins of the eastern Indonesian archipelago, the Marianas, and Fiji. Quaternary carbonate reefs and platforms from modern backarc settings may overlap, either geographically or conceptually, with intra-arc carbonate systems, especially where fault-bounded intra-arc basins evolve into backarc basins during progressive stretching of a volcanic arc. Most Quaternary carbonate systems in backarc basins consist of (1) predominantly fringing reefs or barrier reef tracts with narrow back-reef “lagoons” that are constructed on the backarc side of the volcanic arc, (2) drowned pinnacle reefs or mounded carbonate buildups that also typically form on fault-bounded highs in the backarc region, or (3) relatively narrow platforms that are constructed on faulted backarc basement highs that either extend from the volcanic arc or serve as accommodation zones that partition extensional backarc regions into separate sub-basins. The present-day abundance of fringing reef tracts and drowned buildups reflects both the high-amplitude (i.e., >100 m) character of the post-Pleistocene sea-level rise and the typically steep structural or volcanic topography that was flooded by this rise in sea level. The obvious structural controls on platform dimensions and

locations reflect the ongoing faulting and underfilled character of many Quaternary backarc basins, which typically are in early stages of development (i.e., <5 m.y. old). Carbonate platforms and reefs are rare in more tectonically mature backarc basins, which have undergone more tectonic subsidence, and potential substrates for carbonate sedimentation are below photic depths. Ancient examples. Ancient examples of backarc carbonate platforms and reefs include (1) thin intervals of skeletal limestone in the dominantly siliciclastic lower member (150 m total thickness) of the middle Cretaceous Olvidada Formation, Baja California Norte, Mexico (Phillips, 1993); (2) Upper Triassic (Norian) coral buildups and carbonate foreslope(?) “conglomerates” (70 m total thickness), Puale Bay, Alaska Peninsula (Wang et al., 1988); and (3) possibly various Triassic carbonate-platform sequences from Western North America (Stanley, 1982). The interpretation of many of these examples as backarc-basin carbonate sequences is based largely on paleogeographic reconstructions where the relative position of the volcanic arc and the direction of subduction are known. Future work may ultimately show, however, that some of these examples formed in forearc or intra-arc basins. General Tectono-Stratigraphic Model for Carbonate Facies Deposited across Backarc Regions General tectono-stratigraphic models for carbonate platforms and reefs that form in backarc settings depend on whether the backarc basin is extensional or nonextensional (Fig. 17). Given the relative randomness of eruptive events and shifting dispersal patterns for volcaniclastic sediment, there may not be a generalized and recurrent stratigraphic succession that develops in backarc regions. Carbonate sedimentation, however, is probably more common, long-lived, and areally extensive during later stages of backarc basin evolution, when subsidence rates should be slower, volcanic activity is waning, deposition of volcaniclastic and epiclastic sediment has created broad substrates for carbonate platform development, and arc topography could be highly eroded so that less sediment is supplied to the backarc region. In remnant-ocean backarc basins that form where older oceanic crust is trapped behind an intraoceanic volcanic arc, carbonate facies will likely develop only after volcanic edifices have built themselves up to within photic depths (Fig. 17A). Narrow shelves or fringing reef systems may form along the backarc side of the newly formed volcanic arc. Carbonate platforms can become more areally extensive on the backarc side of these intraoceanic arcs when deposition of volcaniclastic sediment and lava flows have built broad, shallow-water substrates. Active volcanism and tectonic deformation of the volcanic arc will prevent thick carbonate accumulations from developing, but as volcanism and deformation wane, thicker platform sequences might develop. Thus, the oldest carbonate sediments in remnant-ocean backarc basins likely record when the growing volcanic arc first shallowed into photic depths. These initial deposits might be followed by complexly interlayered sequences of thin carbonate facies with thicker volcaniclastic deposits and lava flows. The thickest and perhaps most areally extensive carbonate sequences record the

Carbonate-platform facies in volcanic-arc settings end of volcanic and tectonic activity in remnant-ocean backarc basins. Rare(?) crustal shortening might occur at any evolutionary stage in remnant-ocean backarc basins and could generate surface deformations (probably related to fault-propagation folding) or cause regional flexural subsidence across the backarc region that might also influence carbonate-sedimentation patterns. In contrast, carbonate sedimentation in extensional backarc basins will likely be influenced by the fault systems that accommodate stretching (Fig. 17C–E). Initial stretching will likely occur within the hot and weak lithosphere that underlies the volcanic arc, so arc-related basement rocks may provide initial substrates for the earliest carbonate strata. Most carbonate platforms in modern extensional backarc basins consist of fringing reefs or narrow rimmed shelves that build from the steep tectonic, volcanic, or erosional topography on either side of the basin. There are no reported examples of synrift platforms or reefs within more axial, deep-water parts of modern extensional backarc basins, but they should be recognizable in ancient backarc basins by their association with extensional fault systems and by their internal growth stratal patterns. As stretching proceeds, the evolving synrift topography must influence the location, dimensions, and internal growth stratigraphy of carbonate platforms and reefs, just as in marine intracontinental rift settings. If extension progresses to the point at which a spreading axis develops, the rifted basement of the tectonically separated remnant and still-active arc terranes may begin to subside independently. The remnant arc will likely undergo faster rates of postrift subsidence than the active-arc side of the extensional backarc basin, because there will not be a nearby heat source that will prevent the remnant arc from thermally subsiding. In fact, continued volcanic activity, changes in the properties of the subducting slab, or various subduction-related geodynamic processes may even cause uplift along the active-arc side of the extensional backarc region. Carbonate sedimentation, however, may be most extensive in extensional backarc basins during their late synrift to early postrift (i.e., “drift”) stages of development when there are numerous potential substrates across the backarc region where fringing and isolated platforms can develop (cf. Dorsey and Kidwell, 1999). These isolated platforms will reflect the strong influence of remnant synrift topography, where fault-bounded highs are substrates for carbonate platforms and structural lows trap siliciclastic and volcaniclastic sediment from the adjacent active and remnant arcs. Later in the postrift history, when sediments have filled the remnant synrift topography, broad depositional surfaces may serve as substrates for carbonate platforms during relative sea-level rises. Along the remnant-arc side of the extensional backarc basin, the inactive arc will be erosionally beveled and continue to undergo thermal subsidence. Sediment supply from the remnant arc and subsidence rates should progressively decrease over time, which may also enhance development of broad carbonate platforms during later postrift phases. In contrast, carbonate sedimentation along the active-arc side of the basin may never become areally extensive or long-lived as long as there is continued flux of volcanogenic material to the backarc region.

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Tectonically Influenced Unconformities in Arc Depocenters Unconformity development across carbonate platforms in arc depocenters can affect the distribution of porosity and permeability within these platforms. Tectonically enhanced unconformities will be best developed within carbonate platforms that form where tectonic uplift may occur, such as: 1. The crest of accretionary subduction complexes, where the internal deformation history of the subduction complex controls the character of surface deformations that affect unconformity development. Disconformities that might cap shallowing-upward carbonate sequences on the crest of the subduction complex can grade laterally into conformable contacts toward more axial parts of the forearc basin (Caron et al., 2004). 2. Intra-arc basins, which might be uplifted as a volcanic arc grows. Entire isolated platforms or only minor segments of barrier- or fringing-reef systems may be capped by tectonically enhanced unconformities in intra-arc settings, depending on the scale of arc uplift and whether uplift patterns are affected by local faulting. Carbonate platforms in extensional and neutral backarc basin settings are not likely to have major, tectonically enhanced unconformities because these settings are generally undergoing subsidence, with little uplift throughout much of their history. In contrast, compressional backarc settings (i.e., retro-arc foreland basins) may have tectonically enhanced unconformities in the wedge-top, forebulge, and backbulge depocenters (cf. Dorobek, 1995; Jordan, 1995; DeCelles and Giles, 1996). SUMMARY Carbonate platforms and reefs provide sensitive records of differential subsidence and uplift, environmental conditions, and interactions with other depositional systems in arc settings. In general, detailed stratigraphic studies are lacking for all arcrelated depocenters because original stratal relationships are usually poorly preserved, physical access in many arc settings (both ancient and modern) is difficult, high-quality seismic profiles and extensive well data are not widely available, and actual stratigraphic relationships are highly complex in three dimensions. Thus, it is difficult to construct general stratigraphic models for any arc depocenter. Long-wavelength tectonic-subsidence patterns in arc systems will control the spatial distribution of shallow-water carbonate facies over long time scales and ultimately will control the thickness of carbonate strata. Local, fault-controlled differential subsidence will influence the sizes, morphology, and internal facies patterns of carbonate platforms that develop in any arc depocenter. Siliciclastic and volcaniclastic flux will come largely from the volcanic arc or arc-massif rocks. Carbonate facies may be better developed with greater distance from a volcanic arc, although carbonate facies may even accumulate along the flanks of the arc

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if drainage patterns divert siliciclastic diment elsewhere, volcanic eruptions are infrequent or are of low volume, or major windward-leeward effects are caused by arc topography. Carbonate-platform morphology can be highly variable across arc depocenters and is largely dependent on the availability of tectonic or depositionally constructed substrates. For tropical carbonate systems, these substrates generally need to be at photic depths. Cool-water carbonate facies in arc depocenters, however, might form over a broader range of water depths, but they will still be limited by the availability of suitable substrates. Other factors, such as nutrient levels and wind direction, also influence carbonate-facies development in arc depocenters, although the relative influence of these controls is highly specific to any particular arc setting. I hope this review paper serves as a starting point for analysis of these important systems, although more detailed studies, especially of relatively young or modern arc settings, are needed. Unraveling the history of more highly deformed, ancient examples will require the insight provided by studies from younger arc settings as well as careful tectonic reconstructions, structural restorations, accurate age dating, and detailed sedimentologic and stratigraphic analysis of the ancient examples. ACKNOWLEDGMENTS Much of the research for this paper was completed while I was on leave from Texas A&M University during the 2000–2001 academic year. ExxonMobil Upstream Research Company (EMURC) supported the research during my leave and provided an excellent work environment. I am indebted to colleagues at EMURC, both past and present, but most notably Jim Markello, who was a major proponent for the study. Comments and suggestions by Dan Bosence, John Reijmer, Sam Rice, Moyra Wilson, and co-editor Peter Clift improved the manuscript and are most appreciated. REFERENCES CITED Ballance, P.F., 1993, The New Zealand Neogene forearc basins, in Ballance, P.F., ed., South Pacific Sedimentary Basins, Sedimentary Basins of the World 2: Amsterdam, Elsevier Science, p. 177–193. Bardintzeff, J.M., Brousse, R., and Gachon, A., 1985, Conditions of building coral reefs on a volcano: Mururoa in Tuamotu and Rurutu in Australes (French Polynesia), in Gabrie, C., and Salvat, B., eds., Proceedings of Fifth International Coral Reef Congress: Miami, Florida, Rosenstiel School of Marine and Atmospheric Science, University of Miami, p. 401–405. Beaudry, D., and Moore, G.F., 1985, Seismic stratigraphy and Cenozoic evolution of West Sumatra forearc basin: American Association of Petroleum Geologists Bulletin, v. 69, p. 742–759. Beck, M.E., Jr., 1983, On the mechanism of tectonic transport in zones of oblique subduction: Tectonophysics, v. 93, p. 1–11, doi: 10.1016/00401951(83)90230-5. Bosence, D.W.J., 2005, A new, genetic classification of carbonate platforms based on their basinal and tectonic setting in the Cenozoic: Sedimentary Geology, v. 175, p. 49–72, doi: 10.1016/j.sedgeo.2004.12.030. Bouysse, P., 1984, The Lesser Antilles island arc: Structure and geodynamic evolution, in Biju-Duval, B., Moore, J.C., et al., Initial Reports of the Deep Sea Drilling Project: Washington, U.S. Government Printing Office, v. 78, p. 83–103.

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Printed in the USA

The Geological Society of America Special Paper 436 2008

Sediment waves in the Bismarck Volcanic Arc, Papua New Guinea Gary Hoffmann Eli Silver Simon Day Eugene Morgan Earth Sciences Department, University of California, Santa Cruz, California 95064, USA Neal Driscoll Scripps Institution of Oceanography, La Jolla, California 92093, USA Daniel Orange* AOA Geophysics, Castroville, California 95012, USA

ABSTRACT In the Bismarck Volcanic Arc in Papua New Guinea, six fields of sediment waves were imaged with sonar. Sediment structures observed in seismic data and swath bathymetry are not unique and can result from predominantly continuous (bottom) currents, or episodic (turbidity) currents, or from deformation of sediment. Two of these wave fields overlap and appear to be of turbidity-current origin and modified by bottom currents, with one field unconformably overlying the other field. A field off the coast of Dakataua caldera displays an arcuate morphology, and a series of enclosed depressions within the field suggests creation by extensional deformation of rapidly deposited sediment. Scour features in side-scan imagery suggest turbidity-current activity, which also likely modifies the sediment waves. The wave field is isolated from hyperpycnal currents, however, suggesting that in the absence of a shelf, coastal erosion and small landslides can produce semiregular gravity-driven sediment flows that deposit in deep (>1400 m) water. In Kimbe Bay a fourth sediment-wave field also displays arcuate morphology and enclosed depressions within the field. This wave field is found within a bay >40 km from shore and also appears to have been formed by a combination of extensional deformation of sediment and energetic current activity. Two additional fields in Hixon Bay are fed by small and medium rivers (<~450 m3/s mean annual discharge) draining volcanoes and mountainous regions. One small field appears within a slide scar, suggesting that the initial topography of the scar provided the conditions for early sediment-wave growth. A much larger field is best explained by repeated hyperpycnal currents originating from the Pandi River. We cored a series of upward-fining, graded sequences consistent with a turbidity-current origin. Ages from these cores and measurements of relative thickness in sub-bottom imagery of the

*Present address: Black Gold Energy, Plaza Kemang Timor 22, Jakarta 12510, Indonesia Hoffmann, G., Silver, E., Day, S., Morgan, E., Driscoll, N., and Orange, D., 2008, Sediment waves in the Bismarck Volcanic Arc, Papua New Guinea, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 91–126, doi: 10.1130/2008.2436(05). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Hoffmann et al. field constrain deposition rates for the field and suggest that a large part of the Pandi River discharge must be bypassing the shelf and depositing on the sediment-wave field in deep water (>1200 m). These findings suggest that the sedimentary record in arc collision zones will be dominated by mass-wasting deposits very close to volcanoes, and by river discharge depositing in select, extent regions far from shore. Because sedimentation rates can vary by a factor of 2 between the two flanks of a sediment wave, care must be taken when comparing bed thickness across an entire sedimentary section. Keywords: sediment waves, Papua New Guinea, Bismarck Sea, marine geology, geomorphology.

INTRODUCTION More than half of the total suspended sediment supplied by rivers to the sea originates from small to medium mountain rivers (<~450 m3/s mean annual discharge; Mulder et al., 2003), especially those in tectonically active regions such as the Bismarck Volcanic Arc in Papua New Guinea (Milliman and Syvitski, 1992). At another active arc-continent collision, Taiwan, as much as 42% of the sediment discharge to the ocean occurs by hyperpycnal flows (>40 g/L; Warrick and Milliman, 2003), typically during flood events (Dadson et al., 2005). Large earthquakes can cause an increase in the occurrence of hyperpycnal flows and the percentage of sediment discharged at hyperpycnal concentrations for some rivers for short intervals, owing to landslides (Warrick and Milliman, 2003; Dadson et al., 2005). Under certain conditions, these hyperpycnal turbidity currents can result in deep-sea sediment waves (H.J. Lee et al., 2002; Wynn and Stow, 2002; Schwehr et al., 2007). The Bismarck Volcanic Arc differs from arc-collisional settings such as Taiwan and other regions where turbidity-current sediment waves have been identified (for example, the Var field southeast of France; Migeon et al., 2001) by having active volcanism. Widely dispersed tephra may increase the frequency of hyperpycnal discharges (Hayes et al., 2002), and ashfall from volcanic eruptions may also directly produce turbidity currents (Fiske et al., 1998). Bottom currents, rapid slope failure, and slow soft-sediment deformation, for instance caused by repeated earthquake loading, are also processes that have been suggested as being capable of producing or greatly modifying sediment-wave fields (Gardner et al., 1999; Lee and Chough, 2001; H.J. Lee et al., 2002; Wynn and Stow, 2002). O’Leary and Laine (1996), H.J. Lee et al. (2002), and Wynn and Stow (2002) provide criteria to distinguish among these various processes. Because large undulations (hundreds of meters to a few kilometers) are easily identifiable in sonar surveys of the seafloor, and hyperpycnal turbidity currents are greatly affected by climate, uplift, and seismicity (Mulder et al., 2003), turbidity-currentgenerated sediment-wave fields should be ideal targets for studying these factors in arc and arc-collisional environments. In the Bismarck Sea we imaged multiple sets of large sediment waves,

giving us the rare opportunity to examine the distinguishing criteria using multiple wave fields within the same data set. We can also test if these criteria can be consistently applied to differentiate wave-forming processes, and what sediment waves formed by these processes tell us about dispersal of sediment. We begin with a discussion of the processes that create sediment waves and morphologically similar features. One of these processes, turbidity currents, is surrounded by some controversy, and so we proceed with an examination of the term and its relation to hyperpycnal currents. We follow the discussion of sediment-wave-generation processes and turbidity currents with our specific observations and discuss them in terms of their distinguishing features and how they relate to climate and sediment transport in the Bismarck Sea. STUDY LOCATION Papua New Guinea has two seasons, the summer monsoon season and the winter trade-wind season. Although cyclonic events are uncommon, near-continuous tropical rain produces frequent landslides and consequent large sediment discharge to the coastal ocean from the mountains, as well as by floods any time of year (McAlpine and Keig, 1983; Pickup, 1984; Walsh and Nittrouer, 2003). The sediment load for the entire island of New Guinea is high, at ~1.7 × 109 t yr–1, roughly the estimated load of all North American rivers combined (Milliman, 1995). The Bismarck Volcanic Arc is located in northern Papua New Guinea and forms the southern boundary of the Bismarck Sea (Fig. 1). In the eastern half of the arc the Solomon Sea plate is subducting beneath the Bismarck Sea plate at the New Britain Trench (Johnson, 1979; Taylor, 1979). In the west, subduction has ceased as the Finisterre and Adelbert terranes, parts of a remnant Paleogene volcanic arc that lies in the Bismarck forearc, collide with the Australian plate and are being uplifted to form the Finisterre and Adelbert Mountains (Abbott et al., 1994; Pigram and Davies, 1987). The volcanism in the arc is related to subduction of the Solomon Sea plate and occurs behind the remnant Paleogene arc, which also includes the island of New Britain. The submarine environment surrounding the Bismarck Volcanic Arc was

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-3˚ ismic Bismarck Sea Se Schouten Islands

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an

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N

rita ew B

ench in Tr

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-8˚

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146˚

148˚

150˚

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Figure 1. Location map showing regional setting of the wave fields of the Bismarck Volcanic Arc. The New Britain Trench accommodates subduction of the Solomon Sea plate. New Britain and the Finisterre and Adelbert terranes are parts of a remnant Paleogene arc. The modern subduction-related volcanic arc is offset to the north of the remnant arc. From west to east, some of the modern volcanoes include the Schouten Islands, Manam, Karkar, Long, Tolokiwa, Umboi, Sakar, Ritter, Garove, Dakataua and the Willaumez Peninsula (WP), Pago (P), Karai (K), Gallosfulo (G), Bamus (B), Lolobau (L), Ulawun (U), and Rabaul (R), which lies on the northeast tip of the Gazelle Peninsula (GP) on the island of New Britain. The Willaumez Peninsula is composed of a north-south string of volcanoes. The locations of several other volcanoes are shown as small black circles. Major faults in the region include the Ramu-Markham Fault and the Wide Bay Fault (WBF). Two new faults were mapped during this survey: the Kimbe Bay Escarpment (KBE) in Kimbe Bay (KB), east of the Willaumez Peninsula, and the Torkoro Escarpment (TE), in the Hixon Bay region (HB) north of Lolobau. The outline of the debris flow from the 1888 collapse of Ritter Island is shown. The Sepik River is the largest river on the north coast of Papua New Guinea. It drains a large part of the New Guinea Fold and Thrust Belt. The Pandi River drains a large part of the Nakanai Mountains in New Britain, and the eastern flanks of Ulawun and Bamus volcanoes. The locations of Figures 2, 5, 9, and 14 are shown.

the target of this study. We mapped submarine volcanic collapse features, including deposits of the 1888 Ritter Island collapse (Johnson, 1987), for input into tsunami models (e.g., Ward, 2001). The locations of a selection of volcanoes in the arc are shown in Figure 1. The arc extends westward to the Schouten Islands between 143° E and 145° E, where it meets the Bismarck Sea seismic zone. It extends east to ~152° E at Rabaul caldera on the northeast tip of the Gazelle Peninsula. The study area includes the Schouten Islands in the west but goes only so far as the Hixon Bay area just west of the Gazelle Peninsula. The Gazelle Peninsula is moving northwest relative to the rest of the island of New Britain along the left-lateral Wide Bay Fault (Fig. 1; Madsen and Lindley, 1994). The Wide Bay Fault defines the eastern edge of the Hixon Bay region, where we mapped a new fault, the Torkoro Escarpment. The study area includes the deep-water (>500 m depth) environment around the volcanoes in the arc, including on either side of the Willaumez Peninsula, a string of volcanoes

that juts northward from New Britain into the Bismarck Sea, ending in Dakataua caldera. To the east of the Willaumez Peninsula is Kimbe Bay, which is bounded on three sides by large volcanoes. It is also the location of a newly mapped fault, the Kimbe Bay Escarpment. Collapse blocks were found near several volcanoes during the survey, including off Tolokiwa Island and off Dakataua caldera. At both these locations, sediment-wave fields were also mapped nearby. Sediment-wave fields were also mapped in Kimbe Bay and the Hixon Bay region (boxed regions of Fig. 1). These sediment-wave fields form the basis of the current paper. SEDIMENT-WAVE FORMATION AND DISTINCTIVE FEATURES Sediment waves, in the most common use of the term, refers to any undulating pattern that is caused by differential sediment deposition and erosion over time, analogous to dunes or antidunes,

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Hoffmann et al.

as opposed to undulating patterns that are caused by slope failure such as slow deformation or slide events (H.J. Lee et al., 2002; Wynn and Stow, 2002). Large undulations with wavelengths of 100 m to >10 km are observed in a variety of settings, including on channel levees, axes, and mouths or in canyons (Normark et al., 1980; Kidd et al., 1998; Nakajima and Satoh, 2001; Migeon et al., 2001; Lewis and Pantin, 2002); on the flanks of volcanic islands (Wynn et al., 2000); in troughs (Howe, 1996); and on plateaus and continental slopes (Lee and Chough, 2001; O’Leary and Laine, 1996). Short-wavelength (0.1–1 m) sediment waves have been observed in settings such as terrestrial rivers (Mohrig and Smith, 1996; Jerolmack and Mohrig, 2005a, 2005b; Coleman et al., 2005). Owing to imaging limitations in the deep-sea environment, however, we confine our current discussion to large sediment waves. Two processes are recognized as producing sediment waves as defined here, namely turbidity currents and bottom currents (Wynn and Stow, 2002; Schwehr et al., 2007). Waves of turbidity-current origin are created by numerous episodic events preferentially depositing sediment on the upstream flank of the waves in antidune fashion when the flow is in the Froude number regime of ~0.5–1.9 (Hand, 1974; Bowen et al., 1984; Wynn et al., 2000; H.J. Lee et al., 2002). The Froude number is the ratio between current speed and the speed of gravity waves in a density-stratified fluid, such as a sediment-laden-current layer in seawater. Recent research suggests that initiation of a turbidity-current sediment-wave field may require breaks in slope or preexisting undulations, but once started, these wave fields may grow on their own (Ercilla et al., 2002; H.J. Lee et al., 2002; Kubo and Nakajima, 2002). Deep-sea sediment waves caused by turbidity currents generally have wavelengths of 200 m to 7 km, and wave heights of 2–70 m, and occur on slopes ranging from 0.1° to 0.7° (Wynn and Stow, 2002; H.J. Lee et al., 2002). In profile, the waves migrate upstream and upslope because of preferential deposition on the upstream flank (Wynn and Stow, 2002; H.J. Lee et al., 2002). Internal reflections are commonly continuous between waves (H.J. Lee et al., 2002). Wave crests tend to be oriented perpendicular to the local slope, as turbidity currents travel downslope (Wynn and Stow, 2002; H.J. Lee et al., 2002; Schwehr et al., 2007). Wave sequences progressively thin downslope, wave heights can decrease downslope, and some reflections pinch out downslope in profile, since sediment concentration and deposition in turbidity currents progressively decrease downslope (Mulder et al., 1998). Thus deposition in the distal end will be from large events that occur only rarely (H.J. Lee et al., 2002). Wave crests tend to be linear or sinuous in plan view, with some bifurcation (Wynn and Stow, 2002; H.J. Lee et al., 2002). Waves formed by turbidity currents are often associated with other evidence of turbidity-current activity, such as channels, levees, and scour features. In their review, H.J. Lee et al. (2002) note that the overall cross-sectional shape of a sediment-wave field is commonly concave upward.

The generally accepted concept of bottom-current sedimentwave generation is that internal waves induced within a current in a density-stratified medium flowing over topography will differentially deposit and erode sediment on the seafloor (Flood, 1988). Sediment waves formed by bottom currents have similar dimensions as turbidity-current-generated waves (Wynn and Stow, 2002). Wave crests tend to be linear or sinuous with bifurcations (Wynn and Stow, 2002). Bedforms often migrate upcurrent but can also migrate downcurrent (Wynn and Stow, 2002). They are distinct from turbidity-current-generated sediment waves in that they can form on very low slopes or flat seafloor, and wave crests are often aligned obliquely to the slope because the currents are contour-parallel (Wynn and Stow, 2002). Both bottomand turbidity-current-generated waves tend to display asymmetry between flanks in backscatter imagery (H.J. Lee et al., 2002). Undulating patterns can also arise from deformation, however. These can occur from extension (O’Leary and Laine, 1996) or compression (Lee and Chough, 2001; Hill et al., 2004). These undulating patterns may be related to rapid slope failure, such as slumps (H.J. Lee et al., 2002), or to slow deformation caused by intermittent earthquake loading (Lee and Chough, 2001). These features can be up to 10 km in wavelength, and up to 100 m in wave height (Wynn and Stow, 2002), and they can occur on steep slopes or very shallow slopes (<0.5°; Lee and Chough, 2001). In profile, deformational features are commonly associated with faulting (O’Leary and Laine, 1996; H.J. Lee et al., 2002). Planform imagery of these features is rare, but wave crests may be arcuate, without bifurcation (Wynn and Stow, 2002). Bottom currents, turbidity currents, and deformation can all interact in the creation of sediment-wave fields (Howe, 1996; Faugères et al., 2002; H.J. Lee et al., 2002). In this manuscript we will try to tease out the dominant and subordinate processes responsible for creating these morphologic “sediment-wave” features. TURBIDITY FLOWS AND HYPERPYCNAL CURRENTS Considerable controversy exists regarding the nature of flow and deposits of turbidity currents, as well as the terms used in the literature that apply to various sediment transport regimes (Kneller and Buckee, 2000; Shanmugam, 2000; Mulder and Alexander, 2001; Mulder et al., 2001; Shanmugam, 2002; Mulder et al., 2002). For the sake of clarity in our present discussion, we note an overlap between the terms hyperpycnal current and turbidity flow, and we define our particular usage in this paper. Following Mulder and Alexander (2001), hyperpycnal current refers to sediment-laden river discharge that is denser than the body of water into which it enters. The flow can travel long distances in a semidiscrete layer, entraining fluid and possibly additional sediment if it is exerting shear stresses on the bed high enough to cause erosion (Skene et al., 1997; Mulder and

Sediment waves in the Bismarck Volcanic Arc Alexander, 2001; Mulder et al., 2003). For river discharge entering seawater, the sediment concentration needs to be at least ~40 kg m–3 for the flow to be negatively buoyant (Mulder and Syvitski, 1995; Warrick and Milliman, 2003). Depending on the river, hyperpycnal-discharge events can occur about once every 100 yr, or as often as once per year (Mulder et al., 2003). Turbidity flow refers to sediment suspended primarily by the upward component of turbulence (Lowe, 1982; Middleton, 1993; Stow et al., 1996; Shanmugam, 2000; Mulder and Alexander, 2001). Such flows contain suspended sediment concentrations up to 10% by volume (Shanmugam, 2000; Mulder and Alexander, 2001). They can be generated by a number of mechanisms, such as entrainment of sediment into surrounding seawater at the upper surface of a debris flow. Other mechanisms include hyperpycnal river discharge, and concentration processes in a low-density layer (Lowe, 1982; Middleton and Hampton, 1976; Parsons et al., 2001) such as a layer of tephra fallout during eruptions (Fiske et al., 1998). Coastal erosion during storm events that affect easily eroded sequences could also generate turbidity currents (Hayes et. al., 2002). Scully et al. (2002), for instance, showed that energetic water waves increase the capacity for critically stratified gravity flows to transport sediment. Combined with direct turbidity-current generation from ashfall and indirect formation from reworking of tephra, the presence of volcanoes may significantly increase the occurrence of turbidity-current generation in this setting over other settings. We will therefore use turbidity current to refer to a sustained turbidity flow generated by any mechanism, and hyperpycnal current to refer to turbidity currents generated owing to hyperpycnal discharge at river mouths. METHODS Multibeam bathymetric data were collected aboard the R/V Kilo Moana on cruise KM0419, using a hull-mounted SIMRAD EM-120 echo sounder. Backscatter data were collected using a towed MR1 side-scan sonar system. Side-scan data were gridded in a 16 m grid. Multibeam bathymetry and backscatter data were collected along the 1000 km length of the arc in a swath up to 100 km wide. Bathymetric data were median filtered, using a 500 m kernel. High-resolution Compressed High Intensity “Radar” Pulse (CHIRP) sub-bottom sonar data were collected with an Edgetech 1–6 kHz swept frequency with a 50 ms duration. The CHIRP data are unmigrated. This system transmitted approximately every 5–10 m along selected transects. Penetration was up to a few tens of meters. Twelve short (<2 m) gravity and piston cores were collected, primarily targeting volcanic-collapse debris flows identified in side-scan imagery. We examine four sediment-wave fields that were identified in the Bismarck Arc in order to determine the various generation mechanisms of the sediment-wave fields. Doing so will help us understand how material is transported through the arc system and how important slope failure is compared to bottom and

95

turbidity currents. We begin in the central part of the arc north of Tolokiwa, and then continue eastward to look at Dakataua, Kimbe Bay, and Hixon Bay (Fig. 1). TOLOKIWA WAVE FIELDS Tolokiwa Island volcano, near the center of the Bismarck Volcanic Arc (Fig. 1), has undergone at least one lateral collapse in the Holocene. The sediment waves occur in a field north of the island on the northern side of a broad channel ~8 km wide and 40 m deep that grades at ~0.2° to 0.5° from west to east into the basin (Fig. 2). Two sediment-wave fields overlap in this region. The first is composed of large sediment waves with wavelengths of 600 m to 1 km and wave heights of ~10 m to 40 m (Fig. 3). The second field is composed of smaller sediment waves with wavelengths of 100 m to 400 m and wave heights of a few meters (Fig. 3; Fig. 4C, E). Waves in both sets are identifiable in side-scan images as having a distinctly higher backscatter along the eastern flanks (Fig. 3). The field of large sediment waves is overlapped by the field of small sediment waves, with some small waves appearing between and overlying large waves (Fig. 3; Fig. 4C). The crests of the small waves are oriented slightly oblique to their local slope. This obliquity varies from wave to wave but is <30°. The crests of the large waves are perpendicular to the local slope. The high-backscatter eastern, downslope flanks of the waves correspond to the thinner, more steeply dipping beds of the waves (Fig. 4). The large sediment waves commonly show differential deposition, with the upslope side receiving a factor of 2 or more accumulated sediment than the downslope side. The crests of the sediment waves migrate upslope through time (Fig. 4A, B). The small waves lap onto the large waves (Fig. 4C, D), and they overlie parallel reflectors that begin at a depth of 5 m (Fig. 4E). DAKATAUA WAVE FIELD The Dakataua wave field resides on the northern flank of Dakataua caldera. The field is 24 km by 18 km and elongated N-S (Fig. 5). The average north-dipping slope of the field is 1.4°, with the upslope flanks of the waves dipping south between 1° and 1.5°, and the downslope flanks dipping north ~10°. The wave field as a whole forms a topographic high of up to 200 m relative to the flatter seafloor that surrounds the wave field. The seafloor to the east of the wave field is deeper than that to the west. This difference decreases northward, although north of the 240-mhigh ridge at the north end of the field the seafloor maintains an eastward-dipping component of ~0.15° (Fig. 5). Between the caldera and the wave field, the seafloor dips smoothly at ~5°. In the side-scan data the waves generally do not exhibit asymmetry; however, the downslope flanks nearest the caldera display high backscatter (Fig. 6). Scour features are apparent on the 5° slope leading into the wave field (Fig. 6).

Large Sed Waves all m S

Collapse Blocks

es av dW Se

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Fig 4

Fig 3

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No Data

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Figure 2. Bathymetry of Tolokiwa sediment-wave fields and surroundings. Two distinct fields overlap—one with wavelengths of 0.6–1 km, labeled large sed waves, and one with wavelengths of 200–400 m, labeled small sed waves. Locations of Figures 3 and 4 are shown. Bathymetric contour interval is 25 m, with annotated contours in bold every 200 m.

Ritter debris flow Fig 4

Large waves

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Small waves

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Figure 3. Side-scan mosaic (top) and bathymetry (bottom), contoured at 2 m intervals of Tolokiwa sediment-wave fields. Location of Figure 4 shown. The waves are apparent in the side-scan imagery as features with higher acoustic backscatter (darker shades) in the western flanks than eastern flanks. Insonification direction is north-south. Nadir lines run east-west at 5°02′ S and 5°06′ S. Black east-west lines are sea-surface reflection artifacts. In both parts of the figure, short dashed lines show local trend of slope. In side-scan image, several wave crests are denoted with short, thin lines to highlight the slight obliquity of the crest-normal to the slope. Dotted lines outline the sediment-wave fields. Note in the side-scan image that some small waves appear within the large waves in the region of overlap.

upslope migration

(Left) The upslope flank is 9.2 m thick, measured to the lowest strong reflection; the downslope flank, measured to the same reflection, is as thin as 4.5 m. These measurements assume a sound velocity of 1500 m/s. (Below) Small sediment waves onlapping large sediment waves

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E Note that reflections on the upslope sides of the waves thin or pinch out downslope

Parallel reflections beneath small wavelength sediment waves

(Left) Small sediment waves appearing in between two large sediment waves, and onlapping them.

C Figure 4. CHIRP profile of eastern end of sediment-wave field north of Tolokiwa Island where the field merges with the ripple field. Location shown in Figure 2. VE—vertical exaggeration; TWT—two-way traveltime. Enlarged boxes demonstrate particular features of note. All enlarged boxes are at the same scale, which is 4 times larger than and at the same vertical exaggeration as the base figure.

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Figure 5. Location map showing wave field in relation to Dakataua caldera and bathymetry (in meters). Contour interval is 25 m, with annotated contours in bold every 200 m. Locations of Figures 6, 7, and 8 are shown. A lineation, shown as the dash-dot line, runs through the field approximately N-S. The lineation ends slightly west of the 240-m-high ridge at the north end of the field.

150˚ 00'

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Figure 6. Side-scan mosaic of Dakataua sediment wave field. Locations of Figures 7 and 8 shown. Scale is 1.6 times that of Figure 5. High backscatter is dark. The two elongated regions of high backscatter in the field (~4°52′ S and 4°53′ S) are in the troughs between waves. Solid lines denote the slope break between wave crests and steep downslope flanks of waves. Dashed lines denote enclosed depressions. The Dakataua lineation is shown as the dash-dot line.

Sediment waves in the Bismarck Volcanic Arc The waves themselves are 1–4 km in length, with wave heights of ~50 m, as measured from the upslope trough to the crest. In plan view they are irregular, although generally arcuate, with crests largely continuous across the width of the field (Fig. 6). The crests are sinuous, and some are bifurcate (Figs. 6 and 7), and they also tend to be wider than the troughs between waves (Fig. 7). Many of the troughs between waves contain enclosed depressions (Figs. 6 and 7). Some of the depressions are associated upslope with steep slopes that display high backscatter, but many are not (Fig. 6), and the depressions are not associated with high backscatter scour. A lineation runs N-S through the field (Fig. 5), characterized by eastward-dipping slopes, widened depressions, and pinchedout wave crests (Fig. 7). The CHIRP failed to penetrate more than 1 m throughout this field (Fig. 8). KIMBE BAY WAVE FIELD Whereas the Tolokiwa and Dakataua waves are both isolated from large inputs of fluvial sediment, Kimbe Bay is fed on three sides by a multitude of small rivers that drain volcanoes. The Kimbe Bay wave field lies near the center of the bay, almost 50 km from shore, adjacent to the Kimbe Bay Escarpment. The Kimbe Bay wave field is >30 km long and 15 km wide (Fig. 9). In plan view the waves are generally arcuate, although irregular, with wavelengths varying from 1 to 4 km and wave heights reaching >80 m in some areas (Fig. 9). The average slope of the wave field is 1°. Within the field, many enclosed depressions attain a depth of ~100 m and are found between wave crests (Fig. 10). Some low-wave-height features (~10 m) are found west of the main sediment-wave field. The main wave field lies on a topographic high of up to 250 m relative to the sea floor to the west (Fig. 9). To the east, a crevice ~2 km wide and up to ~100 m deep separates the wave field from the Kimbe Bay Escarpment. To the south of the wave field is a topographic ridge at least 200 m high (Fig. 9). Several channels lead into the crevice between the wave field and the escarpment. One channel is associated with high backscatter as it cuts through the ridge east of the Kimbe Bay Escarpment (Fig. 11), then drops 200 m into a plunge pool (S.E. Lee et al., 2002), which is >100 m deep as measured from the downslope side of the channel (Fig. 9). The crevice increases in width northward, to nearly 6 km, where depressions appear (Fig. 9). The depressions and the escarpment appear as very high backscatter features in side-scan imagery (Fig. 11). The high backscatter regions within these depressions occur on the southern slopes, and the depressions attain depths up to 75 m (Fig. 12). A lineation is contrasted against the background sediment of the wave field, appearing as a higher backscatter region nearly 10 km long and 1 km wide (Fig. 11). The steeper, downslope sides of the sediment waves produce higher backscatter than the rest of the wave field (Fig. 11). The slopes of the high backscatter faces of the waves approach

101

10° (Fig. 10). The high backscatter of the downslope sides of the waves is associated in the CHIRP profiles with thinning sequences and upslope migration (Fig. 13). But in most cases (Fig. 13A, B, D) the upslope migration is apparent only on the steep downslope flanks, and only in a few areas can upslope migration be seen clearly across a whole wave, which might be due to cutting across some waves obliquely (Fig. 13C). Additionally, the character of the reflections varies from wave to wave, especially in terms of penetration depth. Some wave crests are imaged to a depth of ~8 m (Fig. 13B, C), whereas wave crests both upslope and downslope of these crests are imaged only to a depth of 2–3 m (Fig. 13A, D). HIXON BAY Hixon Bay allows us to test whether fluvial sedimentation is an important process far from shore in this system. It is several kilometers wide and the receptacle for the Pandi River (Fig. 14). This river drains ~800 km2 of the Nakanai Mountains and the eastern flanks of Ulawun and Bamus volcanoes (Fig. 1). Backscatter, Morphology, and Sub-bottom Profiles of the Hixon Bay Sediment-Wave Field A large sediment-wave field is found in the Tokoro Trough of Hixon Bay (Figs. 14–16). This trough attains depths >1900 m, and two channels feed into the trough from the Pandi River (Fig. 14). The sediment-wave field (Fig. 15) begins shortly below the point at which channel scour diminishes between Torkoro and Mele Reefs. A slide scar is present east of Lolobau Ridge and is shown in more detail in Figure 17. In side-scan images (Fig. 18), high backscatter scour features highlight the Pandi River channels. The slide scar shows a stippled pattern of high backscatter in the upper part, and the edges are highlighted as high backscatter streaks in the lower half of the slide scar. The sediment-wave field (Fig. 18) appears in backscatter as a pattern of alternating higher and lower backscatter, with regions of varying wavelengths. We divide the wave field into six groups on the basis of morphology (Fig. 16), and we summarize their characteristics in Table 1. The six groups of waves are also distinct from one another in sub-bottom imagery (Table 2). We correlate a transparent layer between stronger reflections across group 2 and into group 4 (Fig. 19). Reflection thickness variations above this layer indicate that wave-crest deposition rates decrease monotonically downslope. The thickness varies by a factor of 1.5 across the group 2 waves, and by nearly a factor of 2 from group 2 to group 4 (Figs. 19–21). We also measured the relative thickness of reflector sequences at the distal end of the wave field, taking advantage of continuous reflectors in group 4 (Fig. 22D). From the upslope side to the downslope side of Figure 22D the sequence thins by ~20%.

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Figure 7. Bathymetry of Dakataua wave field, contoured at a 4 m interval, with annotated contours in bold every 20 m. Location of CHIRP line depicted in Figure 8 shown. The downslope sides are consistently steeper than the upslope sides of the features. Note the bifurcation near the center of the CHIRP line shown in Figure 8 and elsewhere. The Dakataua lineation is characterized by eastward-dipping slopes, widened depressions, and pinched-out wave crests.

5m

CHIRP penetration is about 1 m. Diffractions give the illusory appearance of deeper penetration.

250 m

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Downslope, the subsurface remains largely impenetrable.

Figure 8. CHIRP profile of Dakataua wave field. Location shown in Figure 7. VE—vertical exaggeration; TWT—two-way traveltime. Detailed enlargements are at the same vertical exaggeration and 3.5 times the scale as the base figure. Both demonstrate the extremely limited penetration (~1 m) that characterizes this wave field.

Fig 12 K imb e Bay Escarp me nt

Depressions

Kimbe Wave Field

Fig 10

Plunge Pool

Fig 13

Crevice

Channels

10 km No Data

Figure 9. Bathymetry of the Kimbe Bay sediment-wave field. Contour interval is 25 m, with annotated contours in bold every 200 m. Side-scan mosaic of the same region at the same scale is shown in Figure 11. Motion along the newly discovered Kimbe Bay Escarpment is unknown. In the region between the escarpment and the wave field are several depressions associated with high backscatter. The outlines of these depressions correspond to regions of high backscatter. Detailed bathymetry and backscatter imagery of the pull-apart basins are shown in Figure 12. Locations of Figures 10 and 12 are shown. CHIRP line depicted in Figure 13 is shown.

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Figure 10. Detailed bathymetry of Kimbe Bay sediment-wave field. Contour interval is 5 m, with annotated contours in bold every 25 m. CHIRP line depicted in Figure 13 is shown; the scale of this line is 2.9 times that of Figures 9 and 11. Note the pronounced irregular morphology of the waves, with an overall arcuate pattern of alternating crests and troughs; superimposed on this is a chaotic assortment of peaks and valleys, resulting in highly sinuous structures.

150˚ 30' -4˚ 50'

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Fig 12

n tio ea Lin

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Plunge Pool

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Crevice

Fig 13

Channel Scour

-5˚ 05'

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Figure 11. Side-scan mosaic of the same region depicted in Figure 9, and at the same scale. Dotted outline of field same as in Figure 9. Locations of Figures 10 and 12 are shown. CHIRP line depicted in Figure 13 is shown. Low backscatter is light, and high backscatter is dark. High backscatter pattern within the wave field is associated with troughs. High-backscatter lineation is oriented 33.7° with respect to the general trend of the Kimbe Bay Escarpment. Solid lines denote the slope break between wave crests and steep downslope flanks of waves. Dashed lines denote enclosed depressions.

150˚ 38' -4˚ 50'

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Figure 12. Close-up side-scan mosaic and bathymetry of enclosed depressions in the crevice, and the backscatter lineation noted in Figure 11. Contour interval is 5 m, with annotated contours in bold every 25 m. Scale is 2.8 times that of Figures 9 and 11.

The waves display some downslope thinning (B) and apparent upslope migration on their downslope flanks (A and B). The irregularity of the morphology (see Figure 12) results in the CHIRP profiles imaging reflections from adjacent surfaces at different depths (A). seafloor reflections from different locations

B 10 m 500 m

upslope migration

A sequence thins downslope

upslope migration

N

S 2.20 30 m

1 km

2.30 T W T (s)

VE = 15

2.40

2.50

C

upslope migration

D

10 m 500 m

upslope migration Apparent upslope migration is present in most of the waves (C and D). In some cases this is true for the whole wave (C), but for most of the waves, upslope migration is only apparent on the steep, downslope flanks where it becomes difficult to distinguish between internal structure and diffractions (A, B, and D). Penetration varies widely from wave to wave, with some crests imaged to a depth of up to 6 to 8 m (B and C), and others only imaged to a depth of 2 to 3 m (A and D), suggesting sedimentation conditions are not consistent across the wave field. Figure 13. CHIRP profile across a part of the Kimbe Bay wave field. Location shown in Figures 9, 10, and 11. VE—vertical exaggeration; TWT—two-way traveltime. Detailed enlargements are at 4.2 times the scale and at the same vertical exaggeration as the base figure.

151˚ 15'

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No Data

16 00

No Data

We st T ork oro R

idg e

Torkoro Trough

-4˚ 40' ro R rko t To Eas

Tork oro Esca rpm ent

1 800

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00

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idg

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Mele Rf

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s

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No Data

00 Fig. 17

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Fig. 16

Torkoro Reef

5 km

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No Data

Lolobau Island

Hixon Bay

Ulawun Volcano

-5˚ 00' 151˚ 10'

151˚ 15'

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Pandi River

151˚ 25'

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New Britain

151˚ 30'

-5˚ 00' 151˚ 35'

Figure 14. Bathymetry of the Hixon Bay region. Contour interval is 25 m, with annotated contours in bold every 200 m. Several channels feed into the imaged region, including those from the Pandi River, which drains part of the Nakanai Mountains (see Fig. 1) and Ulawun volcano. The Pandi River channels fade out as they enter the region between the Mele Seamount and Torkoro Reef. This is where a major sediment-wave field begins. This field is shown in more detail in Figure 16, the location of which is shown here. A slide scar is observed on the slope east of West Torkoro Ridge and Lolobau Ridge. This scar is shown in more detail in Figure 17, the location of which is shown here.

151˚ 15'

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ne

ent

ls

r

carp m

idg e

Torkoro Trough

We st T ork oro R

Tork oro Es

co u

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Large sediment waves Lower slide scar Small sed wvs

Cha nne l sco

ro R rko t To Eas

ur

idg

Upper slide scar

e

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Ch an

Lolobau Trough

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Mele Rf Mele Seamount

r

u co

stippled pattern

ls ne an Ch

lower backscatter

Fig. 17

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Fig. 16

Torkoro Reef

-4˚ 55'

Lolobau Island

Hixon Bay

Ulawun Volcano

-5˚ 00' 151˚ 10'

5 km

Channel scour

Lolobau Ridge

151˚ 15'

151˚ 20'

Pandi River

151˚ 25'

-4˚ 55'

New Britain

151˚ 30'

-5˚ 00' 151˚ 35'

Figure 15. Hixon Bay area, with the location and scale the same as for Figure 14. Low backscatter is light, and high backscatter is dark. Regions without data are white. The slope of the Torkoro Escarpment is notably high in backscatter. The slide scar shows up clearly as a pattern of high backscatter. The backscatter pattern is distinctly different between the upper slide scar and the lower slide scar. The wave field within the box locating Figure 16 can be seen as a repetitive pattern of higher and lower backscatter. Both large sediment waves and small sediment waves can be seen. The large sediment waves are seen as a longer wavelength, lower contrast pattern, whereas the small sediment waves are seen as a shorter wavelength, higher contrast pattern. Channels are associated with high-backscatter scour.

151˚ 24'

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1650

Fig 22

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162

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Group 5

0 1375

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Torkoro Reef 151˚ 18'

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Figure 16. Bathymetry of the Hixon Bay main sediment-wave field. Contour interval is 5 m, with annotated contours in bold every 25 m. Sediment waves are divided into five groups, based on morphology. Locations of CHIRP profiles through sediment waves in Figures 19 through 22 are shown. Arrows highlight the channels that lead to the sediment-wave field. See text for details on the groups of waves. Core holes BB08 and BB09 are shown.

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151˚ 15'

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00

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region of undulating bathymetry

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ar

p

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Sc

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1275

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Fig

26

Figure 17. Detailed bathymetry of the slide scar in the Hixon Bay region. Contour interval is 5 m, with annotated contours in bold every 25 m. The slide scar and scarp are delineated with the dotted line. In the upslope portion of the scar, an undulating pattern is seen. A CHIRP line that cuts across these features is shown in Figure 25.

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0 120

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Gravity Cores from the Hixon Bay Sediment-Wave Field We collected two gravity cores to examine the nature of sediment being deposited and its rate of deposition. The core locations are shown in plan view in Figure 16 and in cross section in Figure 22. Core BB09 was taken in the trough between two waves, just upslope of a 1-km-wavelength, 20-m-high wave. The slope on which it is located dips <0.05° to the south. Core hole BB08 is 3.5 km downslope from core hole BB09 on the downslope flank of a wave near the very end of the field. The wave has a length of 1 km and a height of 10 m. The north-dipping flank of the wave on which core hole BB08 is located has a slope of 1°. Core BB09 is 164 cm long, and core BB08 is 85 cm long (Fig. 23). Depths in the cores are not corrected for compaction during collection. Core BB09 penetrated through a unit of gray clay containing pumice dropstones (~130 cm; Fig. 23). Core BB08 penetrated into, but not through, this same unit (~80 cm). In both cores a series of silty and sandy tephras is interbedded with fine silt, silty clay, and clay units (Fig. 23). The well-sorted fine-black-sand tephra at 105 cm depth in core BB09 correlates physically with the well-sorted fine-black-sand tephra at 53 cm depth in core BB08 (Fig. 23). If this correlation holds, then the sedimentation rate between the clay-pumice unit and this tephra unit was slightly

lower in core BB09 than in core BB08. Subsequently, however, sedimentation would have been more rapid by nearly an order of magnitude in core BB09 than in core BB08, because of the correlation of the well-sorted fine white and black sand unit at 65 cm depth in core BB09 and at 42 cm depth in core BB08 (Fig. 23). In both cases, this white and black sand unit is topped with a pale gray clay unit that is slightly thicker in core BB08. Radiocarbon dating of planktonic foraminifers in core BB08 indicates that the white and black sand unit was deposited between 373 and 242 yr ago. The pale gray clay unit above it then took more than 200 yr to emplace, at a rate of roughly 0.5 mm yr–1. Calibrated radiocarbon dates of planktonic foraminifers from both cores are shown plotted in Figure 24. The depths of the dates from each core have been normalized to the top of the well-sorted white and black sand layer we assume is from a single, discrete event affecting both sites. Ages in both Figure 23 and Figure 24 are given in years before 2004, when the cores were collected. Ages are taken from median probability dates, with upper and lower ranges shown. Sedimentation patterns are broadly similar at the sites of cores BB08 and BB09, but in detail there are important differences, in particular the thickness of the pale clay layer above the white and black sand layer, the relative thicknesses between the

151˚ 18'

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Fig 22

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high backscatter spots

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Group 6

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r ca s de Sli

Group 4

1

2 Fig

ou Gr

1 km Seamount

p3

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Fig

19

20

Fig

-4˚ 48'

-4˚ 48'

scour Group 1

Torkoro Reef 151˚ 18'

151˚ 20'

151˚ 22'

151˚ 24'

151˚ 26'

Figure 18. Side-scan mosaic of the central Hixon Bay region sediment-wave field. Dark tones are high backscatter. Region and scale are the same as for Figure 16. Locations of the CHIRP lines depicted in Figures 19 through 22 are shown. The different groups of waves in the sediment-wave field display different acoustic characteristics, with some groups displaying strong asymmetry between upslope and downslope flanks (groups 3 and 6), and others displaying weak or no asymmetry. High backscatter spots mark the slopes of a few group 4 waves nearest the slide scar.

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TABLE 1. BACKSCATTER AND MORPHOLOGICAL CHARACTERISTICS OF HIXON BAY SEDIMENT WAVES Group number Wave length Wave height Backscatter contrast Wave crests Channel (m) (m) 1 600–800 20–30 None Slightly sinuous Yes 2 ~1000 30–80 Faint Somewhat sinuous No 3 500, D.D. 20, D.D. High Linear or somewhat sinuous Yes 4 1000–2000 30–40 High contrast spots near slide scar Slightly sinuous No 5 ~1000 15–20 Faint Slightly sinuous No 6 100–500 10–20 Faint Linear or slightly sinuous No Note: Backscatter contrast refers to the flanks of individual waves. Channel denotes whether the group is associated with a channel directly upslope. D.D.—decreasing downslope.

TABLE 2. SUB-BOTTOM PROFILE CHARACTERISTICS OF HIXON BAY SEDIMENT WAVES Penetration depth Reflector character Upslope Continuous reflectors Figure reference (m) migration 1 <5 N.D. N.D. N.D. 19 2 Up to 20 Well stratified, D.T. Yes Between some waves 19, 20 3 <5 Primarily N.D. N.D. 20 diffractions 4 Up to 20, Well stratified, D.T. Yes Between some waves, 19, 21, 22 less at distal end at distal end 5 ~10 Well stratified, D.T. Yes Between some waves 21, 22 6 <5 Primarily N.D. N.D. 21 diffractions Note: D.T.—downslope thinning, meaning the sequence is observed to thin downslope; N.D.—no data, meaning the group was not imaged clearly in sub-bottom profiles. Group number

well-sorted fine-black-tephra layer and the white and black sand layer, and the number of upward-fining units above the white and black sand layer (Fig. 23). The heterogeneity between the cores is not surprising because the wave field as a whole is heterogeneous, and sedimentation patterns vary across the field as a whole.

In profile, one wave shows clear upslope migration (Fig. 25). The complex morphology and slopes of up to 4° within the slide scar prevent clear imaging of the field with a shallow-towed CHIRP system. The downslope end of the slide scar also appears to have been buried (Fig. 16). DISCUSSION

Hixon Bay Region Slide The upper part of the Hixon Bay slide contains undulating features, and sediment dispersed from slope failures within the slide scar may deposit on the distal end of the sediment-wave field (Fig. 17). The slide scar is apparent but diverges from a half-circle shape in its southern end. A clearly defined chute runs downslope toward Torkoro Trough, although part of the northern boundary of this chute is obscured by undulations that compose most of the upper part of the slide scar. These undulations are reminiscent of the sediment waves discussed above, with a wavelength of roughly 1 km and wave heights of up to 40 m. Either side of the slide chute appears as a channel ~5 m deeper than the center of the chute (Fig. 17). In side-scan imagery these channels are associated with high backscatter streaks (Fig. 15). The waveforms appear as alternating high and low backscatter regions, with the steeper, downslope flanks associated with high backscatter. Where the overall slope is flattest the contrast in backscatter is much less.

Based solely on the criteria given in O’Leary and Laine (1996), H.J. Lee et al. (2002), and Wynn and Stow (2002), all of the imaged wave fields fail simple classification. Nonetheless, they provide important clues to the origins of sediment-wave fields. Sediment waves, although well-recognized features, are not ubiquitous but form only under very specific conditions. Once the origin of a given wave field has been established, it yields important information on broader processes. Tolokiwa Sediment-Wave Fields The upslope migration and acoustic asymmetry observed in the large sediment waves northeast of Tolokiwa suggest a current origin. The dimensions of the waves and the average slope of the field are consistent with either a turbidity-current origin or a bottom-current origin. The region is seismically active, and so the low slopes (<0.5°) are also consistent with a deformational origin, but the fields are not immediately associated with a fault.

sequence thins downslope

NW

SE 1.75

TWT (S)

1.85

50 m

Group 1

VE = 8.2 1 km

Levee

1.95 2.05 Group 2 2.15

Group 4

2.25

upslope migration

10 m 200 m continuous reflections between features

Figure 19. CHIRP line across Hixon Bay sediment waves. Location shown in Figure 18. Detailed enlargements are at the same vertical exaggeration and at 5.1 times the scale of the central base figure. The black tick marks indicate a transparent layer between a series of reflections that is correlated throughout this part of the wave field. Angled lines highlight upslope migration in the wave field. This CHIRP line transects a small part of group 1, a small levee, group 2, and a small part of group 4. VE—vertical exaggeration; TWT—two-way traveltime. The thickness of deposits above the transparent layer increases upslope from ~5 m in group 4 to 6 m at the bottom of group 2 to 9 m at the top of group 2. This measurement is for the thickness at the wave crest and assumes a sound velocity of 1500 m s–1.

erosional surface

10 m

200 m

NW

SE 1.60 VE = 6.6

1.70

50 m

channel TWT (s)

1.80

1 km

levee 1.90 2.00 2.10

Group 2 Group 3

2.20

Figure 20. CHIRP line across the Hixon Bay sediment-wave field. Location shown in Figure 18. The line begins within the channel leading to group 3. It then crosses the levee of the channel at a highly oblique angle. The line then crosses the largest waves of group 3. Because the group 3 waves are small, diffractions dominate the CHIRP image. Penetration in the group 3 waves is also <5 m. The group 2 waves display upslope migration, as highlighted by the angled lines. VE—vertical exaggeration; TWT—two-way traveltime.

10 m 250 m

2.10

E

W

TWT (s)

2.20 outside of wave field

Group 6

2.30 50 m 2.40

VE = 13.9 1 km

Group 4

Group 5

10 m 250 m

Figure 21. CHIRP line across the Hixon Bay sediment waves. Location shown in Figure 18. This line cuts obliquely across the southern end of group 4, group 5, and the central part of group 6, and slightly beyond. The tick marks indicate a transparent layer that we interpret to be the same transparent layer seen in Figure 19. VE—vertical exaggeration; TWT—two-way traveltime. The group 4 waves appear to display upslope migration, as highlighted by the angled lines. Note the decreasing penetration from group 4 to group 5 to group 6.

B

Upslope migration

C

BB09

S

N

2.30 TWT (s)

50 m

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2.40 BB09

Group 4

Group 5

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BB08

Torkoro Trough

continuous reflections

A

D

BB08

10 m

500 m

Upslope migration Figure 22. CHIRP profile of Hixon Bay sediment wave field. Location shown in Figure 18. VE—vertical exaggeration; TWT—two-way traveltime. Core holes BB08 and BB09 are shown. This line crosses the group 5 waves perpendicular to their crests. Each of these waves displays clear upslope migration, as shown in the two enlargements (A and B). In the group 4 waves shown here, upslope migration can be seen (angled line in panels C and D). Also apparent is a downslope thinning of the sequence. Reflections in panel D are continuous and grade into parallel reflections in Torkoro Trough, and thin by ~20% from the upslope-most wave in panel D to the downslope-most wave. This thinning is also seen when comparing cores BB08 and BB09 (see Fig. 23).

Depth (cm) 0

age before 2004

BB09

description

description faintly bedded brownish-gray fine silt fining up to clay

thin dark tephra bearing silts and clay upward-fining silt to clay

faintly bedded brownish-gray fine silt fining up to clay

90

50

basal mixed fine sand faintly bedded brownish-gray fine silt fining up to clay faintly bedded brownish-gray fine silt fining up to clay

homogeneous brownish-gray silty clay 119

0

97

25

upward-fining silt to poorly sorted silty clay pale gray clay sorted fine white and black sand silty clay pale clay upward-fining fine-silt to silty clay well sorted fine black sand tephra

sand injection pale gray clay sorted fine white and black sand

75

age before 2004

upward-fining silty-clay to clay

faintly bedded brownish-gray fine silt fining up to clay

25

BB08

poorly sorted upward-fining clay and silt

brown mud

upward-fining fine-silt to silty clay silty clay fine silt tephra silt/fine sand brown-gray silty clay

187 242 373 437

50 437 573

75

gray clay with pumice dropstones

laminated silt and silty-clay

100

black sandy tephra gray clay tephra sand injection

100

black well sorted fine sand tephra interbedded clays and tephras

125

125 gray clay with pumice dropstones

pale gray well sorted silt tephra

150

pale gray homogenous clay

150

silty tephra

Figure 23. Cores BB09 and BB08 from the Hixon Bay region sediment-wave field, with calibrated radiocarbon ages of planktonic foraminifers. Both cores penetrated a gray clay unit containing pumice dropstones. Above this unit is a series of tephra deposits and silt, silty clay, and clay deposits. Both units contain a thin layer of well-sorted white sand with black grains (65 cm depth in core BB09, and 42 cm depth in core BB08), topped by a unit of pale gray clay. Ages (in years before 2004) from this pale gray clay unit indicate that it was deposited slowly in comparison with the upward-fining silt and clay units found above it in both cores. Depths are uncorrected for compaction of sediment during core collection.

0

-30

Top of pale clay layer (BB08)

-40

Top of well sorted white and black sand layer

Depth in BB09 (cm)

t axis)

Depth in BB08 (cm)

r-1 (righ

-20

axis) -1 left yr (

6 mm y

4 mm

-10

-50

-60

-70 Data from BB08

Data from BB09

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100

200

300 400 500 Age (years before 2004)

600

700

Figure 24. Depth-age plot of radiocarbon dates from planktonic foraminifers from cores BB08 and BB09 in the Hixon Bay sediment-wave field. Age is in years before 2004, when the cores were collected. Depths have been normalized to the top of the well-sorted white and black sand layer we observe in both cores, and which we assume is from a single, discrete event affecting both sites. Depth within core BB08 is shown on left axis. Depth within core BB09 is shown on right axis. Median ages are shown, with upper and lower ranges shown. For comparison, lines corresponding to example sedimentation rates for each axis are drawn. The sedimentation rate for core BB09 since the late nineteenth century is ~6 mm yr–1. The average sedimentation rate for core BB08 since the event that produced the white and black sand layer is about two-thirds the average sedimentation rate in core BB09 since the same event. Thus, 4 mm yr–1, using the scale of the left axis, is shown.

One wave in the field (right) shows apparent upslopemigration

1.50

Scarp

TWT (s)

1.60 30 m 1.70

VE = 11.9 1 km

1.80 1.90

5m

200 m

Figure 25. CHIRP line through Hixon Bay slide. Detailed enlargements are at the same vertical exaggeration and at 5 times the scale of the base figure. VE—vertical exaggeration; TWT—two-way traveltime. Upslope migration is clear only in one wave in this field. Location of figure shown in Figure 17.

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2 1.9

relative deposition thickness

1.8 1.7 1.6 1.5 1.4 1.3 1.2 1.1 1 0

1

2

3

4

5

6

7

distance (km) distance (km)

Figure 26. Plot of relative deposition thickness on wave crests in the Hixon Bay sediment wave field as a function of relative distance. Measurements are taken from seven wave crests shown in Figure 19, normalized to the thickness of the most distal wave crest measured. These are plotted as diamonds. Three measurements are taken from wave crests in the distal end of the waves shown in Figure 22, plotted as crosses. These measurements are also normalized to the most distal wave crest in the set. Distance is kilometers upslope from the most distal measurement in the set. The proximal measurement from Figure 22 (the cross at 5.5 km) is extrapolated to the wave crest from the site of core hole BB09 and assumes a relative sedimentation rate at the site of this core hole of 1.55 times the sedimentation rate at the site of core hole BB08. The plot shows that these two sets of measurements, which are separated by 10 km, are consistent with a downslope exponential decay of sedimentation rate from wave crest to wave crest. Short-dashed line shows an exponential curve with a characteristic decay length of 7.6 km.

We interpret that some of the small waves are younger features than the large waves, owing to the onlapping contact that the small waves have with the large waves. Nevertheless, other small waves interspersed with the larger waves could have formed concomitantly. Because of their acoustic asymmetry, and the oblique orientation of their wave crests with respect to local gradients, we interpret the small waves to be of bottom-current origin. The lack of evidence for turbidity-current activity in the vicinity of the wave fields is consistent with this interpretation. The origin of the large waves remains inconclusive. We suggest, however, that the consistent upslope migration more likely indicates a current origin than a deformational origin. The nearest possible source of turbidity currents to the field is debris flows from Tolokiwa, but the orientation of the waves is inconsistent with currents flowing from south to north. One alternative is that the large waves were also generated by bottom currents. In this case, the differences in morphology between the large waves and small waves would imply a difference in bottom-current

velocity and direction, and possibly a change in density stratification within the ambient seawater (Flood, 1988). However, the crests of the large waves are oriented perpendicular to the local slope, suggesting that they were formed by turbidity currents derived from shallow regions to the west. The small waves thus may have formed subsequent to or concomitantly with the large waves as they were modified by bottom currents. Dakataua Sediment-Wave Field Features in the Dakataua wave field suggest both a deformational origin and a turbidity-current origin. The decreasing wave height with distance from the caldera (Fig. 5), combined with the decreasing high backscatter scour features in side-scan imagery (Fig. 6), suggest that turbidity currents may play a role. The field is isolated from hyperpycnal currents, however. The wave field begins ~5 km from shore, south of a steep slope that drops more than 1200 m in this space. We suggest that coastal erosion by

Sediment waves in the Bismarck Volcanic Arc energetic waves during large storms could result directly in turbidity currents as sediment is concentrated in the surf zone during peaks of storm activity and then released during lulls in wave action (Scully et al., 2002). Energetic waves could also result in turbidity currents by destabilizing the slope of the caldera, resulting in numerous small slope failures of weakly consolidated sediment that rapidly entrains seawater to become a turbidity flow (Marr et al., 2001). Additionally, ashfall from eruptions can result in a hyperconcentrated sediment layer on the sea surface, which can result in turbidity currents (Fiske et al., 1998; Parsons et al., 2001). Lastly, turbidity flows could be triggered by earthquakes. Turbidity-current-generated sediment waves proximal to volcanic sources but lacking fluvial input are recognized elsewhere (Wynn et al., 2000). The arcuate and irregular morphology of the waves, as well as the large enclosed depressions in the field, suggest that deformation of some kind may play a role in the formation of these sediment waves. The Dakataua lineation (Fig. 7) may be a fault with some degree of downthrow to the east. We suggest that the Dakataua wave field is formed by some combination of turbidity currents and deformation. Lacking sub-bottom data, however, the origin of the features remains speculative. Kimbe Bay Sediment-Wave Field The Kimbe Bay wave field is far from a potential source of hyperpycnal currents. The highly irregular and relatively arcuate plan-view morphology and the many enclosed depressions in the field suggest a deformational origin. The waves display some subsurface characteristics of current activity, however (Fig. 13), in upslope migration of the wave crests. The presence of nearby channels and plunge pools indicates that turbidity currents are active in the vicinity of this field. The crevice to the east of the field may be related to motion along the Kimbe Bay Escarpment if this is an active fault, or the crevice may be the site of focused channel flow. The depressions in the crevice may be tectonically controlled, although without subsurface imagery we cannot confirm this interpretation. The lineation in the field (Fig. 11) may be the site of a splay fault related to the Kimbe Bay Escarpment. Like the Dakataua field, the Kimbe Bay field forms a topographic high relative to the surrounding seafloor, suggesting higher sedimentation rates within the field. The two fields are also similar in terms of wave height, wavelength, the presence of deep enclosed depressions within the fields, and irregular morphology. Both fields are closely related to signs of turbidity-current activity and possible faults. We suggest that the Kimbe Bay field is also formed by a combination of turbidity currents and deformation, although we cannot determine which process dominates. Hixon Bay Sediment-Wave Field The morphological heterogeneity of this sediment-wave field suggests that different parts of the field formed under different conditions. However, the groups we have defined are not

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entirely distinct from one another, with some wave crests continuous through the boundaries of two or more groups (Fig. 18). This overlap suggests that either the morphology predates deposition by currents or that currents interact with topography, depending on their magnitude, and that a continuum of magnitudes and slightly different flow paths will result in indistinct boundaries between flow conditions, giving rise to the variability in the morphology of the field. Hyperpycnal currents that originate from the Pandi River will tend to follow the channels leading toward the wave field, as seen by the high backscatter scour in these channels (Fig. 18). A turbidity-current origin for the Hixon Bay sediment waves is strongly supported by sub-bottom data. The similarity of reflections from wave to wave (Figs. 19–22), the progressive downslope thinning of reflection sequences (Figs. 19, 22), the continuity of reflections between waves (Figs. 19, 22), and the ubiquitous upslope migration of waves (Figs. 19–22) all argue against a deformational origin and in favor of a current origin. Torkoro Trough is enclosed on three sides (Fig. 14). Cores taken from the sediment-wave field show a series of upward-fining, graded units consistent with turbidites (Fig. 23). These observations argue against a bottom-current origin and in favor of a turbidity-current origin of the wave field. The heterogeneous flow conditions across the wave field may be largely the result of focused flow down the channels and different magnitudes of flow events. The high-backscatter regions within the slide scar (Fig. 25) and the high-backscatter spots on the upslope flanks of three sediment waves near the base of the slide scar (Fig. 18) suggest active sediment transport through the slide scar. Some of this sediment may reach the distal portion of the sediment-wave field. However, the undulating topography in the upslope part of the slide scar (Fig. 17) implies a site of active deposition. This deposition, together with the large, mountainous drainage area of the Pandi River in comparison with the catchments of rivers draining the east flank of Lolobau or the west flank of Ulawun, and also the absence of channels or turbidity scour leading toward the slide scar, all suggest that hyperpycnal discharge from the Pandi River is the dominant source for the upward-fining units in the cores. Slope failure and turbidity flows caused by rapid sedimentation and triggered by earthquakes also may explain some of the heterogeneity of the field. We conclude that the Hixon Bay sediment-wave field is of hyperpycnal-current origin, at least insofar as active growth is concerned. We cannot determine whether or not the field grew from initial topography created by some other process. Sediment Supply to the Hixon Bay Sediment-Wave Field The Hixon Bay sediment-wave field demonstrates that significant quantities of sediment are capable of reaching the Kimbe Bay wave field, >40 km from shore. The Hixon Bay sedimentwave field covers ~70 km2 (Fig. 16). The waves display a pattern of decreasing deposition downslope. Using a transparent layer

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that can be correlated through much of the field, we measured the relative thicknesses of deposits on seven wave crests (Fig. 19). Farther downslope in the distal end of the field, we measured three more by tracing continuous reflections (Fig. 22D) and by extrapolating the difference in sedimentation rates observed in cores BB08 and BB09 to the wave crest nearest core hole BB09 (Fig. 22C), taking advantage of continuous reflections between this core and the wave crest. Although flow conditions vary throughout the field, we assume that sediment is primarily, although not exclusively, originating as discharge from the Pandi River. The similarity of cores BB08 and BB09 (Fig. 23) and the similarity of reflection sequences from wave to wave in groups 2, 4, and 5 (Figs. 19–22) are consistent with this assumption. We assume no discontinuities in sedimentation rate between groups. That reflector thickness decreases steadily downslope from wave to wave, even between groups (Figs. 19, 22) is consistent with this assumption. The waves display upslope migration, with thicker upslope sediment deposits and thinner downslope sediment deposits. We assume that the thickness of sediment at a wave crest is approximately the average sediment thickness for a given wave. We normalize the seven measurements of thickness from Figure 19 to the thickness measurement of the wave crest farthest downslope in this set. We similarly normalize the three measurements of thickness from the distal end of the wave field (Fig. 22). Lastly, we assume a power-law distribution of hyperpycnal-current events from the Pandi River (Warrick and Milliman, 2003; Dadson, et al., 2005). Thus, deposition rates are greater in the upslope part of the wave field because this region receives sediment from many small, frequent events as well as from large, infrequent events, whereas the downslope part of the wave field receives sediment only from the large, infrequent events (H.J. Lee et al., 2002). A power-law distribution of events is expected to produce an exponential decay of sedimentation rates away from the source. We find that the two sets of normalized relative-thickness measurements are consistent with a single exponential decay rate (Fig. 26). Linear functions will also fit both sets of measurements. However, linear functions are not invariant to scalar multiplication. Because we are normalizing each set of measurements, a single linear function cannot fit both sets. The characteristic decay rate of an exponential function, however, is invariant to scalar multiplication. The characteristic decay length of this curve is ~7.6 km. That is, from wave crest to wave crest, sedimentation rates are inferred to decrease by half 7.6 km downslope from any given point in the field. Based on the above assumptions, we extrapolate this decay rate to the remainder of the field and calibrate it with sedimentation rates inferred from radiocarbon dating of cores BB08 and BB09 (Fig. 24). Because we are interested in modern discharge of the Pandi River, we use only the upper parts of the cores, estimating sediment accumulation in the distal end of the field at 4 mm yr–1 (Fig. 24). Note that this estimate does not account for compression of the cores during collection and so is a lower bound estimate. However, some of this is likely to be pelagic sedimentation, and

some of this may be related to currents flowing down the Hixon slide scar (Fig. 25). Thus, we conservatively halve this estimate of sediment flux arriving from the Pandi River. Integrating, the average deposition rate for the whole field is ~5 mm yr–1. Multiplying by the area of the field, this yields a volumetric deposition of 3.5 × 105 m3 yr–1. Assuming a bulk density of the sediment of 2200 kg m–3 results in a mass deposition rate of 7.7 × 105 t yr–1. If this estimate represents a significant portion of Pandi River discharge, then it places a testable lower bound on the sediment supply of this system. We cannot characterize the uncertainty in this estimation, but it suggests that a large part of the Pandi River sediment bypasses the Papua New Guinea shelf and is deposited in deep water. Because we observe sedimentation rates of >4 mm yr–1 nearly 40 km from shore, we conclude that fluvial sediment entering Kimbe Bay can easily account for the deposition we observe in the Kimbe Bay wave field, although probably not the morphology itself. The rivers entering Kimbe Bay are smaller than the Pandi River, on the whole, and 40 km from shore the Hixon Bay sediment-wave field is dying out, not just beginning, as is the case in Kimbe Bay. Thus, although the Kimbe Bay field may be dominantly of turbidity-current origin, it is probably not of hyperpycnal-current origin. We note that previous work has suggested that as much as 90% of sediment from the Sepik River may bypass the shelf as sediment gravity flows (Walsh and Nittrouer, 2003). This inference was based on a comparison of sediment discharge from the Sepik and measurements of depositional thickness on the shelf. The measured thickness of deposits only accounted for 10% of expected discharge. A similar argument was made by Dadson et al. (2005) regarding sediment discharged at hyperpycnal concentrations from the Choshui River in Taiwan. The Hixon Bay sediment waves present the opposite side of the story for the Pandi River. Results from this field suggest that on active tectonic margins with narrow shelves, hyperpycnal flows play an important role in delivering sediment to the ocean. This may be capable of explaining missing shelf deposits elsewhere, such as off the Sepik River, the Choshui River, and the Eel River (Walsh and Nittrouer, 2003; Dadson, et al., 2005; Sommerfield and Nittrouer, 1999). CONCLUSIONS We image six sediment-wave fields in the Bismarck Volcanic Arc. The two Tolokiwa fields appear to have formed by turbidity currents and were subsequently or concomitantly modified by bottom currents. The Kimbe Bay and Dakataua wave fields appear to be of deformational and turbidity-current origin. In Hixon Bay, two hyperpycnal-current-generated sediment-wave fields are observed. One is a small, apparently young field in its initial stages of growth. The other field is an older, more stable feature related to discharge from the Pandi River. Measurements from this wave field indicate that a large part of sediment discharge from the Pandi River is deposited in deep water.

Sediment waves in the Bismarck Volcanic Arc The evidence for turbidity currents in the vicinity of the Dakataua and Kimbe Bay wave fields, and that both fields are topographic highs relative to the surrounding seafloor, suggest focused deposition within the area of these wave fields. Between these fields and the Hixon Bay wave fields, sediment waves with turbidity-current components of formation appear to be a common feature in the Bismarck Volcanic Arc, particularly where sedimentation rates are relatively high. In these locations, deposition rates can vary greatly in the space of a few kilometers, both along the strike of the arc because of focused sedimentation, and across the strike because of the variation of sedimentation rates between the upslope and downslope flanks of a sediment wave. Thus, care must be taken when applying interpretations of a sedimentary section to the remainder of the arc environment. Between these fields and the Tolokiwa fields, we do not observe large sediment waves. Topographic relief in this region is lower than in eastern New Britain, and so we expect that sedimentation rates are significantly lower. Thus, sediment waves with turbidity-current contributions to formation may be an indicator of relatively high sedimentation rates, although the inverse may not be true. We are not aware of large turbidity-current sediment waves that are recognized in the sedimentary record. This is probably due largely to an inherent difficulty in recognizing such large, subtle features from outcrops, particularly if the sequence is at all deformed. Sections in which high sedimentation rates have been recognized thus may be good places for reexamining the sedimentary record for sediment-wave fields. ACKNOWLEDGMENTS Funding for this research was provided by the Marine Geology and Geophysics section of the U.S. National Science Foundation, grant OCE-0327004 to E. Silver and S. Ward, and grant OCE-0328278 to N. Driscoll. We are grateful for the advice and support of many people in carrying out this research, including Hugh Davies, James Robins, Bruce Appelgate and the Hawaii Mapping Research Group, Russell Perembo and Susan John of the University of Papua New Guinea for picking foraminifers, Michaele Kashgarian of the radiocarbon facility at Lawrence Livermore National Laboratory for assistance with dating, the Oregon State University coring team, and the captain and crew of the research vessel Kilo Moana. We thank A. Draut (Associate Editor), H. Lee, J. Sample, and K. Straub for their thorough and insightful reviews of this manuscript. Most of the figures for this paper were generated in part using Generic Mapping Tools. REFERENCES CITED Abbott, L.D., Silver, E.A., Thompson, P.R., Filewicz, M.V., Schneider, C., and Abdoerrias, 1994, Stratigraphic constraints on the development and timing of arc-continent collision in northern Papua New Guinea: Journal of Sedimentary Research, v. 64, p. 169–183. Bowen, A.J., Normark, W.R., and Piper, D.J.W., 1984, Modeling of turbidity currents on Navy submarine fan, California Continental Borderland: Sedimentology, v. 31, p. 169–185, doi: 10.1111/j.1365-3091.1984.tb01957.x.

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The Geological Society of America Special Paper 436 2008

The Lichi Mélange: A collision mélange formation along early arcward backthrusts during forearc basin closure, Taiwan arc-continent collision Chi-Yue Huang Department of Earth Sciences, National Cheng Kung University, Tainan, Taiwan, and Research Center of Ocean Environment and Technology, National Cheng Kung University, Tainan, Taiwan Chih-Wei Chien Department of Earth Sciences, National Cheng Kung University, Tainan, Taiwan Bochu Yao Guangzhou Marine Geological Survey, Ministry of Land and Resource, P.R.China, Guangzhou 510075 Chung-Pai Chang Center for Space and Remote Sensing Research, National Central University, Chungli 320, Taiwan

ABSTRACT Marine surveys show that the submarine Huatung Ridge extends northward to the Lichi Mélange in the southwestern Coastal Range, suggesting that formation of the Lichi Mélange is related to arcward thrusting of the forearc strata in the western part of the North Luzon Trough during active arc-continent collision off southern Taiwan. A new seismic survey along the 21° N transect across the North Luzon Trough in the incipient arc-continent collision zone further reveals that deformation of the Huatung Ridge occurred soon after sedimentation in the western forearc basin, whereas sedimentation was continuous in the eastern part of the remnant North Luzon Trough until the complete closure of the forearc basin approaching SE Taiwan. This suggests that the sequence in the Huatung Ridge can be coeval with just the lower sequence of the remnant-forearc-basin strata. Multiple lines of new evidence, including micropaleontology, clay mineralogy, and fission track analyses along the Mukeng River and its tributary key sections, are used to test this thrusting-forearc-origin hypothesis of the Lichi Mélange. In the SW Coastal Range the Lichi Mélange lies between the collision suture of Longitudinal Valley to the west and the Taiyuan remnant forearc basin to the east. A field survey indicates that the Taiyuan forearc-basin sequence and its volcanic basement were thrust westward over the Lichi Mélange along the east-dipping Tuluanshan Fault. The Lichi Mélange shows varying degrees of fragmentation of strata, mixing, and shearing. An apparently wide range of facies is present, from the weakly sheared broken formation facies, with discernible relict sedimentary structures, to the intensely sheared block-in-matrix mélange facies, with pervasively scaly foliation Huang, C.Y., Chien, C.W., Yao, B., and Chang, C.P., 2008, The Lichi Mélange: A collision mélange formation along early arcward backthrusts during forearc basin closure, Taiwan arc-continent collision, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 127–154, doi: 10.1130/2008.2436(06). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Huang et al. dipping to the SE. Sedimentological study reveals that the subangular to subrounded, fractured, matrix-supported metasandstone conglomerates in the pebbly mudstone layers are repeatedly found in the broken formation facies of the Lichi Mélange. Their composition and occurrence are identical to the deep-sea-fan conglomerate beds in the Taiyuan remnant-forearc-basin strata to the east. Benthic foraminiferal faunas are similar in the Lichi Mélange, regardless of the varying intensity of shearing and strata disruptions, and are compatible with the benthic foraminiferal fauna in the Taiyuan remnant-forearc-basin turbidites, supporting the interpretation that the protolith of the Lichi Mélange was originally deposited in the North Luzon Trough. Age determination of planktic microfossils further demonstrates that the Lichi Mélange is early Pliocene (3.5–3.7 Ma), implying that this mélange was deposited in a short time and that deformation occurred soon after its deposition. The early Pliocene age of the Lichi Mélange is coeval with just the lower part of the Taiyuan remnant forearc strata, and is much younger than the upper forearc sequence (3–1 Ma). Thus the Taiyuan coherent-forearc-basin strata (3.7–1 Ma) were deposited continuously in the remnant North Luzon Trough regardless of the deformation in its western part (the protomélange). This scenario is an analogue for the modern configuration of the Huatung Ridge–remnant North Luzon Trough off the southern Coastal Range in the active arc-continent collision zone north of lat 21° N. In addition to its kaolinite content (11–15%), the clay mineral composition of the Lichi Mélange is compatible with the Taiyuan remnant forearc turbidites. In the Coastal Range, kaolinites are found only in the volcanic rocks of the Tuluanshan Formation. This additional kaolinite in the Lichi Mélange could not have been derived from the exposed accretionary prism to the North Luzon Trough by sedimentary mass slumping, because no such volcanic rocks are now exposed in the accretionary prism west of the Coastal Range. Instead, they could have been derived from the Tuluanshan Formation when it was emplaced into the Lichi Mélange by thrusting during the last 1 Ma when the Luzon arc-forearc was accreted to form the southern Coastal Range. Thus the kaolinites of the volcanic arc rocks were redistributed into the Lichi Mélange by fluid flows along the ubiquitous geological fractures in the mélange, consistent with the field occurrences of the large, rootless, fault-bounded volcanic rocks of andesitic breccia, tuff, and agglomerates that were floating in the intensely sheared block-in-matrix mélange facies of the Lichi Mélange. Mélange is commonly considered to develop in the accretionary prism of a subduction zone. However, the Lichi Mélange in the SW Coastal Range originated from the thrust forearc strata, representing a unique forearc mélange for orogenic belts worldwide. The young age and wide distribution—especially the continuous offshoreonshore connection—of the Lichi Mélange provides a unique example for further research into active modern mélange-forming processes by forearc thrusting during progressive closure of the forearc basin in this active region of arc-continent collision. Keywords: Lichi Mélange, coastal range in eastern Taiwan, forearc collision mélange, initial arc-continent collision, forearc closure.

INTRODUCTION The Coastal Range in eastern Taiwan has resulted from the accretion of the Luzon Arc-forearc to the exhumed metamorphic basement of eastern Taiwan (Fig. 1) during the arc-continent collision in the last 2 m.y. (Chai, 1972; Biq, 1973; Huang et al., 2000, 2006). Now the location of initial arc-continent collision lies off southern Taiwan (Fig. 1; Huang and Yin, 1990; Huang et al.,

1992; Reed et al., 1992; Lundberg et al., 1997; Malavieille et al., 2002). Stratigraphic and geochemical studies have shown that the Coastal Range is composed of three accreted Miocene–Pliocene volcanic islands, three Plio-Pleistocene remnant forearc basins, two intra-arc basins, and the Pliocene Lichi Mélange (Fig. 2). The accreted Miocene–Pliocene volcanic islands (from north to south: Yuehmei, Chimei, and Chengkuangao) are composed of andesite, agglomerates, and tuff of the Tuluanshan Formation,

119 o E

120 o E

121 o E

122 o E

China V

V V

123 o E

V o

25 o N

25 N

an

ge

Okinawa

R

Trough

Taiwan

o

119 E

uk

III

yu

II

o

122 E

Tr e n c

h

23o N

Sea

SLT HTR

HP o

121 E

Ry

Philippine

TT o

120 E

N L T

a l ani M

21o N

South China Sea

Hengchun Ridge

ing p o K a lope S Fig. 10

Tr e n 22o N

24o N

North

L. C oa

stal

n er W est

ge Ran

ch

23o N

al

V Ra n . ge

C

r ent

Arc

24o N

Luzon

F o ot hil ls Hs ueh sh a n

Taiwan Strait

22o N

I

21o N

o

123 E

Figure 1. Shaded relief map showing regional tectonics inland and off the shore of Taiwan and lines of three E-W seismic profiles. The South China Sea oceanic crust is subducting eastward along the Manila Trench, developing the North Luzon Arc. The accretionary prism of the Manila subduction system extends from the submarine Hengchun Ridge northward to the Miocene deep-sea sediments in the Central Range–Hengchun Peninsula. The ophiolite-bearing Kenting Mélange in the Hengchun peninsula represents the subduction complex within the accretionary prism. Subsequently, the North Luzon Arc collided obliquely with the underthrusting Eurasian continent north of 21° N. The collision resulted in deformation of the western part of the North Luzon Trough as the Huatung Ridge backthrusted eastward, thereby closing the North Luzon Trough from south to north. The Luzon Arc and forearc are accreting on eastern Taiwan, as the Coastal Range and the Huatung Ridge further extend directly northward as the Lichi Mélange in the southernmost Coastal Range. I—intraoceanic subduction zone; II—initial arc-continent collision zone; III—advanced arc-continent collision zone; LM—Lichi Mélange; KM—Kenting Mélange; NLT— North Luzon Trough; HTR—Huatung Ridge; TT—Taitung Trough; SLT—Southern Longitudinal Trough; HP—Hengchun Peninsula; L.V.—Longitudinal Valley.

130

Huang et al.

TECTONOSTRATIGRAPHIC MAP OF THE COASTAL RANGE

o 24

Hualien v

Longitudinal Valley

E

v v

NG

Post-collision Sequence in the suture basin

RA

Forearc Basin Sequence

Yuehmeiv

RA

Loho Basin

v v

CE

Collision Complex Lichi Melange

v v v

TLS

lt

ei F au im

+ v + +

Ch

al V alle y f aul t

v

+ + +

+

Chimei

v

gitu

Chengkung Basin

v

v

Lon

Chingpu Basin

v v

v

din

Intra-Arc

v

v

v

+ + + +

v

v

v

v

KLS

v

v

v v

v v

v v v v v v

v

Yuehmei Volcanic island Chimei volcanic island Chengkuangao volcanic island

v

v

v

v

Luzon Volcanic Arc v

Shuilien

v

NT

Taiyuan Basin

v v

v v

L

Shuilien Basin

v v

Loho

v

fau

SEA

lt

v

an

v v v v v v v v v v v v v v v

Fig. 9B Tungho

PPI I LI

v v v

v

Fig. 8

Taiyuan

v v v v v v v v

v v

N

Fig. 9A 0

10 km

C en tr al R an ge L CCooaasst a .V. t al lRR ange ange

v

Lichi

o 23

PH

v v

o 23

Chengkung

v v v

v

Kuanshan

Chengkuangao v

Fuli

NE

Tu lu

an

sh

Antung

Lutao

Taitung

o 121 30'

Figure 2. Tectonostratigraphic map of the Coastal Range (modified from Huang et al., 1995). Rectangles denote locations of Figures 8 and 9A, B.

whereas the remnant forearc basins (Shuilien, Loho, and Taiyuan) and the intra-arc basins (Pliocene Chingpu basin on Chimei volcanic island and Pleistocene Chengkung basin on Chengkuangao volcanic island) are filled with turbidites (Takangkou and Chimei Formations) derived from the accretionary prism. (Figs. 2 and 3;

Chen et al., 1990; Huang et al., 1992, 1995, 2000). In the Coastal Range the chaotic Lichi Mélange, lying between the remnant Loho and Taiyuan forearc basins to the east and the collision suture of the Longitudinal Valley to the west (Fig. 2), has long been considered the key to understanding the arc-continent tectonic processes in Taiwan (Hsu, 1956; Biq, 1971, 1973; Wang, 1976; Page and Suppe, 1981). The Lichi Mélange occurs in a zone ~2 km wide with characteristics of block-in-matrix fabrics, preferred foliation in scaly argillaceous matrix, and extensional web and boudinage structures in the sandstone blocks (Fig. 4; Chen, 1997; Chang et al., 2000). The variously sized blocks (millimeters to kilometers) of different lithologies in the Lichi Mélange include the andesite suite (andesite, volcanic breccias, tuffs, and volcaniclastic turbidites), the ophiolitic suite (serpentine, gabbro, and pillow basalt), and the sedimentary suite (sandstone, sandstoneshale interbeds, shale, and limestone; Biq, 1971; Wang, 1976; Hsu, 1976; Liou et al., 1977; Page and Suppe, 1981). The andesite suite is derived from the Luzon Volcanic Arc, whereas the origin of the dismembered ophiolite blocks may represent the oceanic crust of either the South China Sea (Suppe et al., 1981; Chung and Sun, 1992) or the Philippine Sea plate beneath the Luzon forearc-Arc (Juan et al., 1980; Malavieille et al., 2002). On the other hand, the sedimentary blocks include the weakly lithified Pliocene turbidite units (tens of meters to a kilometer in size) with similar lithology, age, and sedimentary turbidite structures as the remnant-coherent-forearc-basin strata of the Coastal Range, and the angular, well-lithified, whitish quartz-rich, feldspathic sandstones of late Miocene age (meters to kilometers in size). The latter unit is similar to the deep-sea-fan sandstones in the upper part of the accretionary prism in the Hengchun Peninsula (southern extension of the Central Range; Cheng et al., 1984; Huang et al., 1997) and is not observed in the coherent-forearc-basin turbidites (Pliocene–Pleistocene age) of the Coastal Range. The origin of the Lichi Mélange has been variously interpreted (Fig. 5). It was first considered to be a subduction complex that developed in the former Manila Trench during subduction of the South China Sea oceanic crust in the Pliocene–Pleistocene (Fig. 5A; Biq, 1971, 1973) before the arc-continent collision. Then a sedimentary slumping olistostrome model was proposed (Fig. 5B; Wang, 1976; Ernst, 1977; Page and Suppe, 1981; Lin and Chen, 1986). The Lichi Mélange was regarded as a sedimentary olistostrome deposited in the western part of the North Luzon Trough forearc basin by eastward mass wasting from the exposed accretionary prism (the Central Range) during the Pliocene–Pleistocene arc-continent collision, thus displaying a facies change with the host non-sheared coherent-forearc-basin strata to the east (Fig. 5B). These two disparate models were proposed on the basis of on-land field surveys and mineralogical studies of the dismembered ophiolite blocks and large sandstone blocks in the Lichi Mélange without any constraint from regional marine geological considerations. Intensive marine surveys off southern Taiwan in the early 1990s began to provide detailed bathymetric maps and regional tectonic structures. These investigations do not support the tectonic subduction origin of the Lichi

Biostratigraphy

TIME (Ma)

EPOCH PLEISTOCENE

LITHOSTRATIGRAPHY

0.7

NN19

Collision Melange

RANGE

NORTH

MIDDLE

Shuillien Forearc Sequence

Chingpu Intra-arc Basin

Loho Forearc Sequence

Chimei Formation

1.9

Chimei 2.3

CN12

N21

Shuilien Conglomerate

2.7 3.1

(Arc collapse) (Intra-arc)

Conglomerate

Formation §¥§º Takangkou Fomation Formation

Takangkou Formation

Takangkou Formation

Tungho Ls. Tungho Ls.

3.9

(Arc collapse)

4.3

(Intra-arc) 4.7

B

NN12 A

Shuilien §Ê Takangkou

Chimei Formation

Lichi Melange

N19/20

C NN13

Takangkou Formation

3.5

NN15 CN11 NN14

Takangkou

Formation Formation

NN16

Chengkung Intra-arc Basin

Pinanshan Conglomerate

1.1

CN13

N18

5.1

6

15

(Chengkuangao Volcanic Island

Kangkou ¥§•¶©Ls.

5.5

NN11

SOUTH Taiyuan Forearc Sequence

1.5

CN10

PLIOCENE

Suture Basin

COASTAL

N22

NN18 NN17

MIOCENE

Longitudinal Valley

( Yuehmei Volcanic Island )

Tuluanshan Formation

( Chimei Volcanic Island )

Tuluanshan Formation

Figure 3. Tectonostratigraphy of the Coastal Range.

Figure 4. Occurrence of angular, sheared, whitish quartz-rich, feldspathic sandstone blocks in the intensely sheared block-in-matrix mélange facies (δ-grade) of the Lichi Mélange at sample location LM-4 (location shown in Fig. 9A). The large sandstone block with a black dot in the center is ~8 m wide.

A Subduction Model (Plio-Pleistocene)

Figure 5. Models proposed for the origin of the Lichi Mélange. (A) Tectonic subduction complex model (Biq, 1971, 1973; after Chang et al., 2001). (B) Sedimentary slumping olistostrome model (Page and Suppe, 1981). (C) Tectonic collision complex model (modified from Chang et al., 2001): (1) continuous turbidite sedimentation in the whole North Luzon Trough forearc basin (3.5–3.7 Ma); (2) arcward (eastward) backthrusting, resulting in formation of the proto–Lichi Mélange, similar to the present-day Huatung Ridge in the western part of the forearc basin; the backthrusting also resulted in emplacement of the deep-sea-fan sandstone from the accretionary prism in the Hengchun Peninsula to the proto–Lichi Mélange (3–2 Ma); (3) westward accretion and thrusting of the Luzon Arc-forearc to form the southern Coastal Range (<1 Ma) and formation of the modern Lichi Mélange (for details, see text). L—lower part of remnant forearc strata; U—upper part of forearc strata; 1—accretionary prism; 2—Lichi Mélange; 3—Luzon Arc; 4—Pinanshan Conglomerate in suture basin of the Longitudinal Valley. Blocks in the Lichi Mélange: a—deep-sea sandstone in accretionary prism of Hengchun Peninsula; b—oceanic crust of Luzon Arc-forearc; c—volcanic rocks of Luzon Arc.

Subduction Complex North Luzon Trough Luzon Arc forearc Basin Takangkou & Chimei Fms.

Lichi Melange

South China Sea ~

~ ~ ~~ ~ V

Tuluanshan Fm.

V

V

V

~

V

V

V

V

Arc volcanics

Huatung Basin V

V

Philippine Sea Plate

Eurasian Plate

B Slumping Olistostrome Model (Pliocene) Uplifted accretion ary prism (Central Range)

Coastal Range North Luzon Trough Forearc Basin SE

Erosion

NW

Sub

mar

ine s

lide

Mixed Terrain of

Zo Fa

ult

and Ophiolitic Rocks

coh ern t fore arc tur bid ites a t io n Form Ta ka ng ka ou

Lichi Melange

ne

Continent -derived Turbiditic Rocks

~ 2000 m

flow

V

V

V V

V

V

V

V

V

V

V

V

V

V

V

Tuluanshan

V

V

V

V

V

V

Formation

+ + +Chimei+ + + + + +

V V

V

V

V

V

V

V

V

V

Igneous complex

V

V

C Tectonic Collision Model (Plio-Pleistocene) Longitudinal Valley

Centr

Acc

retio nary Pr

(Huatung Ridg

4

North Luzon Arc Ch 1 (Remnant North Vo engk lca ua Luzon Trough) nic nga Isla o nd

e)

2

c b a

Remnant Forearc Basin

L

Suture basin

U

L

North Luzon Arc

on Trough

North Luz

Forearc Basin L

ism

Lichi Melange

~~~~~

Proto-Lichi Melange

Southern Longitudinal Trough

3.5 ~ 3.7 Ma

Hengchun Hengchun Peninsula Peninsula

U

~~

(1)

L

~ ~ ~ ~ ~~ ~ ~

Accre tiona ry Pri sm

Che n Volc gkuang anic a Islan o d

Taiyuan Trough

3 ~ 2 Ma

Hengchun Peninsula

Taiyuan Basin

~ ~~ ~ ~ ~~ ~ ~ ~ ~~ ~ ~ ~ ~

al Ra nge

(2)

E

Lichi Melange

Luzon Arc (Chengkuangao Volcanic Island)

Coastal Range (southern part)

Forearc Basin (Taiyuan Basin)

W

<1 Ma

Accretionary Prism Longitudinal Valley

Legend

(3)

Oce

ement

canic bas

st ~ Vol anic cru

U L

3

The Lichi Mélange Mélange, because the Lichi Mélange lies to the east off the accretionary wedge instead of within the accretionary prism (Huang and Yin, 1990; Reed et al., 1992; Huang et al., 1992, 2000; Liu et al., 1998; Malavieille et al., 2002). The marine surveys show a direct connection between the submarine Huatung Ridge and the Lichi Mélange in the southernmost Coastal Range (Fig. 6; Huang et al., 1992; Huang, 1993). This led to the proposal of a tectonic collision model, which suggests that the Lichi Mélange originated from the thrusting and shearing of the forearc strata, similar to the modern Huatung Ridge in the western part of the North Luzon Trough during initial arc-continent collision in the early Pliocene (Huang et al., 2000; Chang et al., 2000, 2001). The purpose of this study is to integrate independent lines of new evidence, including a new marine 240-fold multichannel seismic profile across the North Luzon Trough as well as on-land field observations and studies of the planktic and benthic foraminiferal fauna, clay mineral composition, and fission track patterns of zircon grains in the Lichi Mélange to compare with the Taiyuan remnant-forearc-basin strata to the east to constrain the origin of the Lichi Mélange. Most study samples were collected from the Mukeng River and its tributary south of Kuanshan in

133

the southwestern Coastal Range (Fig. 2), where the best sections of the Lichi Mélange are exposed, showing different grades of strata fragmentation, mixing, and shearing. The exposures here led to various interpretations of the origin of the Lichi Mélange in previous studies. STUDY METHODS Marine Surveys A new multichannel seismic survey along a 21° N transect across the North Luzon Trough off southern Taiwan (Figs. 1 and 7) was conducted in 1999, using the R/V Tanbao. The capabilities of this vessel included 240 channels, an I/O digital streamer, 12.5 m of the interval between channels, a 50 m shot distance, a 250 m offset, 2 ms of sample, and 10 s of recording. The energy source was a TI sleeve gun–airgun array. The total size was 3000 C.I., and the pressure was 2000 psi. The results of the survey were then integrated with two 6-channel seismic profiles (MW9006– 26 and MW9006–31, Fig. 1; Chang et al., 2001) previously conducted by the R/V Moana Wave in 1990 to show development of

Figure 6. Geological continuation from the Lichi Mélange in the southernmost Coastal Range to the submarine Huatung Ridge (deformed forearc unit by eastward thrusting) in the initial arccontinent collision zone (marine image adapted from Malavieille et al., 2002). PC—Pinanshan Conglomerate in Longitudinal Valley.

Two-way travel time (sec)

MW 9006-31 0.0

Southern Longitudinal Trough

W

1.0

Huatung Ridge

2.0

10km E 0.0 1.0

Taitung Trough

2.0

Lutao Volcanic Island

3.0 4.0

3.0 4.0 5.0

5.0

MW 9006-26 Southern W Longitudinal 0.0 Trough 1.0 Two-way travel time (sec)

Luzon Volcanic Arc

Luzon Volcanic Arc

Huatung Ridge North Luzon Trough

2.0 3.0

E 1.0 2.0 3.0

Lanhsu Volcanic Island

4.0 5.0

4.0 5.0

6.0

6.0

7.0

7.0

8.0

8.0

GMGS 973 Hengchun Ridge Accretionary Prism

W

8921

8961

9001

9041

9081

9121

9161

9201

9241

9281

9321

9361

9401

9441

9481

9521

9561

9601

9641

9681

9721

E

Two-way travel time (sec)

Luzon Volcanic Arc 3.0

4.0

5.0

6.0

7.0

8.0

Proto-Huatung Ridge 6 5 4 3 2 1

3.0

North Luzon Trough Forearc Basin

4.0

5.0

Batan Volcanic Island

6.0

7.0

8.0

Figure 7. Three seismic profiles, from north to south (see Fig. 1 for location), showing progressive formation of the Huatung Ridge and closure of the North Luzon Trough forearc basin. Line GMGS-973 further revealed syndeformational sedimentation in the North Luzon Trough, regardless of deformation of the proto–Huatung Ridge in the western part of the trough filled with six sequences of strata separated by five unconformities resulting from episodic arc-continent collisions.

The Lichi Mélange

Mukeng River sections (Fig. 8) and supplementary localities in the southernmost Coastal Range (Fig. 9). Both planktic and benthic foraminifers in 200 g of scaly argillaceous matrix (γ- to δgrade shearing) or the Bouma Te part of weakly sheared turbidite layers (β-grade) were identified. Distributions of the main benthic foraminifers were semiquantitatively plotted against stratigraphy according to their relative abundance. The results were then compared with the benthic foraminifers in the coherent-forearcbasin turbidites in the Taiyuan Basin (Chang, 1967). Multivariate cluster analysis was used to compare the similarity of benthic foraminifers within the Lichi Mélange. Benthic foraminifers that

Sector 3

(Fig. 12 A, B)

(Fig. 13 A, B)

(Fig. 14 A, B, C)

19 18 16

C

IIb

Ri

03 40

1 01

Mushi Bridge

85

2 10 06 g n ke 07 u 09 M 08 20 3 05 21

4

12 22

IV Nanshi Bridge

65

23 85

24

25

27

Sector 3 (Fig. 14 D, E)

C

IV

31 29

30

32

55

33

35

D 50

34 60

36 37

39 40

65

1

~5

Stress study sites of Chang et al. (2001)

Sector 1 (Fig. 12 C, D) - grade

H

al

26

III

II

A

N

Taiyuan Forearc Basin

11

I

28

5

5

13

Ra

N

02

65

60

r ve

ntr

4

N

65

Peinan

3

River

85

04

50

nge

B IIa

nge

A

I

Ra

N

D

15

han

2

17 14

Fault

N 500 m

Ce

N

Sector 2

Tu l u a n s

1

Sector 1

K

L

al

Raymond’s (1984) classification of different grades of strata fragmentation, mixing, and shearing—from coherent facies (αgrade), weakly sheared broken formation facies (β-grade) to strongly sheared dismembered facies (γ-grade), and intensely sheared mélange facies (δ-grade)—was followed for mapping the distribution and occurrence of the Lichi Mélange along the

ast

On-Land Study

Co

the Huatung Ridge along progressive closure of the North Luzon Trough by arc-continent collision.

135

Taitung

Raymond (1984)

45

38

Sector 2 (Fig. 13 C,D) - grade

- grade

Sandstone Block

Conglomerates

Igneous Block

Figure 8. Detailed geological map and sampling locations of the study area along the Mukeng River and its southern tributary sections. Raymond’s (1984) classification system in the right bottom corner is followed to map the distribution of various sheared units of the Lichi Mélange (modified from Chang et al., 2000). The section is divided into three sectors according to their exposure sequences from west to east. Field occurrence and distributions of major benthic foraminifers are shown in Figures 12–14. K—Kuanshan; L—Lichi; H—Hualien.

o

o

121 17'30''E

V

d in

an Tu l ua n sh

V

V

TF-7

V

g

V V

V

23 V

V

V V

g

V

L egend Gravel and sand (Holocene) Lichi Melange g With ophiolite Tungho Limestone (Late Pliocene) Takangkou Formation (Pliocene-Pleistocene) Formation V Tuluanshan (Miocene-Pliocene) Slate and schist (Eocene-Miocene)

V

o

TF-7

15.6 Ma 100 Ma

0.0 0 50 100 150 200 250 300 2.72 Ma 5.8 Ma

1.0

TF-9 0.5 92 Ma 50 100 150 200 250 300 0 1.0 2.08 Ma

TF-8 0.5

50

0

100 150 200 250 300

Age (Ma)

o

1 .0

88 Ma

18 Ma

LM-10

0 .5

0 .0

Tungho

0

50

o

121 15'E

100 150 200 250 300

0.0

23 N

V

0.5

V

V

50 3.0 Ma

Takangkou Formation

Taiyuan

62

1.0

0

V

V

TF-6

2.5 Ma 0.0

V

V

TF-6

0.0

V

V

Frequency

Lo

V

40.5 Ma 12.5 Ma 64.5 Ma

0.5

V

V 76

Frequency

V

V

V

i tu

V

V

1.0

Frequency

70

fa u

lt

Va ll

V

V

V

V

V

V

Frequency

g Yu n

al n tr

23

al

V

TF-9

TF-8

ng

V

V

LM-10

o

23 10'N

V

ey

Ce

V

V

fon

e ng

1 km

Ra

0

BV

V

f au

Fuli

Frequency

N

lt

121 15'E

121 17'30''E

100 150 200 250 300

Age (Ma)

A

N 30

42

58

65

33

40

50 40

65

F requency

Lichi Melange 1 .0

120 Ma

60 Ma

LM-1

0 .5 12.8 Ma

0

PEI NAN RIVER

48

72 42

1

2 Km

12 20

300

400

120 Ma

500

LM-2

100

200

300

400

500

F requency

9.76 Ma

Frequency

50 22

g

LM-3

0 .5 110 Ma 0 .0

0

50

100 150 200 250 300

1 .0 66 Ma

LM-4 0 .5 14.8 Ma 0 .0 0

Legend Pin an sh an Con glome r st e L ic h i Me lan ge S an d st on e L ime st on e Agglome r at ic an d e sit e B asalt G ab b ro S e r p e n t in e , Pe r id ot it e Tak an gk ou For mat ion

1 .0

Frequency

g

Blocks in melange

LM-4

200

0 .5

0

LM-2

g

100

9.6 Ma

1 .0

48

LM-3 LM-1

0

1 .0

0 .0

LM-5 60 50

Lichi

Frequency

0 .0 58

50

130 Ma

0 .5

0 .0

100 150 200 250 300

11 Ma

0

50

LM-5

100 150 200 250 300

Age (Ma)

Lichi Melange Figure 9. Fission track records of zircon grains separated from whitish quartz-rich feldspathic sandstone blocks in the Lichi Mélange (A) and from the coherent forearc turbidite layers of the Taiyuan remnant forearc basin (B).

The Lichi Mélange

137

Taiyuan remnant-forearc-basin turbidites were then compared with the samples collected from the nonmetamorphosed deep-sea sandstones in the accretionary prism of the Hengchun Peninsula (Fig. 10) to constrain the origin of the whitish sandstone blocks in the Lichi Mélange.

occur in <5 samples, and the samples that contain <50 specimens of benthic foraminifers in 200 g of sediments, were deleted during computation of the cluster analysis. Mineral compositions of the clay fraction (<2µm) of the collected samples were also identified by X-ray powder diffraction following the method described by Lin and Chen (1986). The results were then compared with those of Lin and Chen (1986), who had systematically documented the clay mineral compositions in the Lichi Mélange and the coherent-forearc-basin strata in the southern Coastal Range. In addition, fission tracks of the zircon grains separated from the Bouma Ta part of the Taiyuan remnant-forearc-basin turbidites were counted to compare with the results obtained from the whitish quartz-rich feldspathic sandstone blocks in the sheared Lichi Mélange near Fuli and the type locality of the Lichi Mélange along the Peinan River (Fig. 9). The detection and counting methods of the fission track age analysis described by Liu (1982) and Liu et al. (2001) were followed. The results of the zircon fission track records from both the Lichi Mélange and the

RESULTS Marine Surveys off the Southern Coastal Range The new 240-fold multichannel seismic profile (GMGS-973 in Fig. 1) across the North Luzon Trough (forearc basin) along the 21° N transect is shown in Figure 7. The profile shows syndeformational sedimentation in the North Luzon Trough. The forearc basin is filled by six sequences of strata separated by five unconformities. All of these six sequences in the western part of the North Luzon Trough are tilted upward and deformed gently by eastward thrust, whereas the strata in the eastern part of the trough remain flat, suggesting no deformation at all (Fig. 7).

1.0

O

5 km

Kt

5

Taiping Fm. Szekou Fm. Hengchun Limestone

~~ unconformity ~~ Ma

50-65 Ma

HP-4

0.5

0.0

Recent Alluvium

50

0

7

Hengchun Peninsula

O

120 40'E

Mt Mt

Kenting Melange (subduction complex) Mutan Fm. lower fan turbidites upper- middle fan conglomerates and sandstones

Ls

300

250

300

0.5

HP-3 0.0

Frequency

HP-1

Ma

Kt

250

Maanshan Fm.

~~ unconformity ~~

0

50 100 15-25 Ma

200

150

55-65 Ma

HP-2

0.5

0.0 0

Loshui Fm. middle fan sandstones

50

100

150

200

250

1.0 15-30 Ma

Frequency

22 00'N

Middle - Late Miocene

HP-2

Slope and trench-fill turbidites in accretionary prism

Hengchun

200

15-30 Ma 85-100 Ma

1.0

O

150

100

1.0

Frequency

Ls

Holocene

HP-3 HP-2

LEGEND

Pleistocene

Mt

Pliocene

Shallow - marine slope basin

HP-3 HP-4

Frequency

22 10'N

70-85 Ma

0.5

HP-1

O

120 50'E

0.0 0

50

100

150

200

250

Age (Ma) Figure 10. Fission track records of zircon grains separated from deep-sea-fan sandstones in the accretionary prism of Hengchun Peninsula. The fission-track age pattern of these deep-sea sandstones is similar to that obtained from the whitish quartz-rich feldspathic sandstone blocks in the Lichi Mélange (Fig. 9A).

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Geology of the Lichi Mélange Exposed along the Mukeng River The Lichi Mélange is exposed in the southwestern and southernmost parts of the Coastal Range (Figs. 1 and 2). The Lichi Mélange lies between the Longitudinal Valley to the west and the remnant forearc basins (Loho Basin and Taiyuan Basin) and volcanic rocks to the east (Fig. 2). In the northern Coastal Range the Lichi Mélange could have been eroded away by the river flowing northward along the eastern side of the Longitudinal Valley. The Longitudinal Valley, a 2–3-km-wide collision suture between the Central Range (accretionary prism and exhumed metamorphic basement) and the Coastal Range (accreted arc-forearc; Fig. 2) accommodates a part of the shortening across southeastern Taiwan (Yu et al., 1997). Between Antung and Kuanshan the volcanic basement beneath the Taiyuan forearc strata has been thrust westward over the Lichi Mélange along the Tuluanshan Fault (Fig. 2). Consequently, the Tuluanshan volcanics are exposed between the Lichi Mélange to the west and the remnant Taiyuan forearc basin to the southeast (Fig. 2). In the Coastal Range this is the only place where the volcanic rocks appear west of the remnant forearc basins, and they lie between the Lichi Mélange and the Taiyuan remnant forearc basin. Otherwise, the volcanic andesitic rocks are always observed to the east of remnantforearc-basin strata (Fig. 2; Hsu, 1956, 1976; Huang et al., 1995). The east-dipping Tuluanshan Fault runs in a NNE-SSW direction, representing the last accretion of the Luzon Arc-forearc to the exhumed Eurasian continent (eastern Central Range) as the southern part of the Coastal Range in the last 1 m.y. (Chang et al., 2000, 2001; Huang et al., 2000, 2006).

Figure 8 shows a detailed geological map of the Lichi Mélange exposed along the Mukeng River and its southern tributary in the southwestern Coastal Range (Chang et al., 2001). The coherent turbidite sequence (Takangkou Formation) of the Taiyuan remnant forearc basin is thrust westward over the Lichi Mélange along the Tuluanshan Fault (Fig. 8). West of this fault the study section shows various degrees (β-, γ-, and δ-grades of Raymond, 1984) of strata fragmentation, mixing, and shearing. The sequences with γ-grade and δ-grade shearing (dismembered and mélange facies) exhibit characteristics of block-in-matrix features (Fig. 4), with pervasively scaly foliation dipping to the southeast (measurements of foliations at five localities are shown in Figure 8; Chang et al., 2000) without discernible bedding (units I, II, III, and IV in Fig. 8), whereas the weakly disturbed (β-grade shearing or broken formation facies) sequences (units A, B, C, and D in Fig. 8) still preserve distinct Bouma turbidite sequences and clear bedding, which generally strike NE and dip SE. However, faults for various thickness are commonly observed in the intervals of different grades (centimeters thick in β-grade, several meters thick in γ-grade, and hundred meters thick in δ-grade) of shearing units and also along the boundaries between the different units (Fig. 11; Chang et al., 2001). From west to east the study profiles are divided into three sectors (Fig. 8) for convenient comparisons and discussion. Sector 1 includes a γ-grade unit and a β-grade unit (units I and A, respectively), whereas sector 2 is composed of two γ-grade units (units IIa and IIb) and two β-grade units (units B and C) in the north (Fig. 8). Units IIa and IIb are combined together as the thick δ-γ-grade unit (unit II) in the south. Sector 3 displays a γ-grade unit III, a thick and less disturbed (βgrade) unit D, and a thick (~500 m along the river), highly sheared

Figure 11. Ubiquitous shearing planes with fault gouges in the Lichi Mélange, especially in mélange facies unit IV (δgrade of Raymond, 1984; Fig. 8).

The Lichi Mélange (γ- and δ-grades) unit IV directly west of the Tuluanshan Fault (Fig. 8). Angular ophiolite blocks, predominantly of gabbro and whitish sandstone blocks, are found restrictedly in units I, II, III, and IV with dismembered facies or mélange facies. In contrast, pebbly mudstone layers composed of subangular to subrounded metasandstone conglomerates are found only in the less disturbed, broken formation facies of units A, C, and D (Fig. 8). Planktic Foraminifers and Age Determination of the Lichi Mélange Planktic foraminifers, including Sphaeroidinella dehiscens, Globorotalia tumida, G. multicamerata, Neogloboquadrina altispira, and Sphaeroidnellopsis seminulina, are commonly found in most of the samples from the Lichi Mélange along the Mukeng River and its tributaries. They show a persistent early Pliocene age (upper Zone N19-20), consistent with results obtained previously by calcareous nannoplankton (Zone NN15; Chi et al., 1981; Chi, 1982; Barrier and Müller, 1984). Benthic Foraminifers in Various Degrees of Sheared Units of the Lichi Mélange Distributions of benthic foraminifers in the various degrees of sheared strata in the Lichi Mélange along the Mukeng River and its tributary are shown in Figures 12–14. As a result of turbidite deposition, the benthic foraminifers in the Lichi Mélange are the mixed fauna of the indigenous deep-water taxa and the displaced shallow-marine species. The deep-marine fauna include the cosmopolitan species of the genera Bulimina, Uvigerina, Fontbotia, Pullenia, Hyalinea, Melonis, Gyroidinoides, and Cyclammina, living on the continental slope to abyssal plain of the present oceans. The shallow-marine fauna include species of the genera Amphistegina, Ammonia, Operculina, Asterorotalia, Pseudorotalia, Calcarina, Elphidium, Lenticulina, and Cellanthus, living in warm waters of the shallow shelf (<200 m). In each study sample, individuals of these shallow- and deep-marine genera make up 60%–80% of the total benthic foraminiferal specimens (Chien, 2003). For convenient discussion, the species with similar ranges of water depth or ecology are shown together in Figures 12–14. Clay Mineral Composition Clay mineral compositions of the sheared mudstone in the broken formation facies or scaly argillaceous matrix and the quartz-rich sandstone blocks in the mélange facies of the Lichi Mélange along the Mukeng River and its southern tributary are listed in the Appendix. The samples are all characterized by bearing illite (average, 51.5%), chlorite (average, 19.5%), mixedlayer clay minerals (average, 11.7%), and kaolinite (average, 11.1%). Smectite is minor (<3%). Kaolinite is always found in the samples of the Lichi Mélange regardless of the mudstones in the broken formation facies, the scaly argillaceous matrix, or the angular whitish sandstone blocks in the mélange facies.

139

DISCUSSION Significance of Marine Geology off the Southern Coastal Range in the Active Arc-Continent Collision Region Marine seismic investigations off southeastern Taiwan reveal the progressive closure of the forearc basin owing to the eastward (arcward) thrusting of the forearc basin strata to develop the Huatung Ridge (Fig. 7; Reed et al., 1992). Profile GMGS973 along the 21° N transect across the North Luzon Trough (Fig. 7), where the Luzon arc collides incipiently against the Eurasian continent, further shows that (1) active syndeformational sedimentation occurs in the forearc basin; once the sediments are deposited in the trough, the sequence in the western part of the forearc basin is deformed and then covered unconformably by the overlying sequence; (2) folding and thrusting in the western part of the forearc basin have occurred since the early history of forearc sedimentation; (3) sedimentation is continuous in the eastern part of the forearc basin regardless of active deformation in the west; and (4) the eastward thrusting accommodates the space between the backstop of the accretionary prism and the Luzon Arc. In compiling two 6-channel seismic profiles in the north (Fig. 7; MW9006–26 and MW9006–31; Chang et al., 2001), regional marine geology shows that the deformed Huatung Ridge forearc becomes larger and larger so that finally the remnant North Luzon Trough is closed at its northern terminus, approaching the southern Coastal Range (Fig. 6). Meanwhile, the deformed ridge crest is thrust upward so that sediments can no longer be deposited on the Huatung Ridge. Because the Huatung Ridge connects northward with the Lichi Mélange in the southernmost Coastal Range (Fig. 6), it is conceivable to consider that the Lichi Mélange has originated from the deformed forearc basin strata like the modern Huatung Ridge and to predict that the Lichi Mélange could have the following geological characteristics: A. The Lichi Mélange should lie west of the remnant forearc basin; B. The Lichi Mélange would contact the coherent forearc basin strata by faulting; C. The depositional time interval of the Lichi Mélange would be less than that of the entire forearc basin sequences, and the depositional age of the Lichi Mélange can be equivalent only with the lower sequence, but older than the upper sequence, of the remnant-coherent-forearc-basin strata; D. Regardless of shearing intensity and facies, the microfossils within the Lichi Mélange should be similar to and compatible with the fauna in the remnant forearc basin to the east; E. The depositional bathymetry of the Lichi Mélange would be similar to the present North Luzon Trough (1500– 3000 m); F. The relict sedimentary structures and the host sediment composition of the Lichi Mélange would be similar to those of the coherent-forearc-basin turbidites;

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04-0

200 m

04-2 04-3 04-4

03 100 m

I

Age

DEEP-SEA TAXA

Ammonia spp. Fontbotia wuellerstorfi Amphistegina spp. Asterorotalia spp. Calcarina calcar Pullenia bulloides Lenticulina spp. Hyalinea balthica Operculina spp. Pseudorotalia spp. Cellanthus craticulatus Melonis spp. Elphidium spp. Cibicidoides spp.

Uvigerina spp. Gyroidinoides spp. Cyclammina spp.

Sample Number

04-1

A

SHALLOW MARINE TAXA

N19 /20 ; NN15 (3.7 ~ 3.5 Ma)

300 m

Unit

Sector 1 4 00 m

B

Sheared grades

A

02

0m

01

Figure 12. (A) Examples of field occurrence and sampling localities of units I and A, respectively, in part of sector 1. (B) Distribution of major benthic foraminifers along the Mukeng River section.

G. The early eastward thrusting in the forearc basin in the initial arc-continent collision stage and the final westward thrusting of the Luzon Arc-forearc during the advanced arc-continent collision accreted to the Coastal Range could have resulted in accretions of the accretionary wedge materials west of the forearc basin and the volcanic arc rocks beneath the forearc basin to the highly sheared, block-in-matrix, dismembered facies and mélange facies of the Lichi Mélange; however, these blocks would not appear in the weakly sheared broken formation facies of the Lichi Mélange or the remnantcoherent-forearc-basin strata. If our different lines of new data agree well with these geological characteristics or predictions listed above, we would prefer that the Lichi Mélange originated along the arcward thrusts during forearc basin closure by active arc-continent collision. On the other hand, if the new data do not fit well with the predictions, we would consider the other possible origins—for

example, sedimentary slumping or mud diapirism—responsible for formation of the Lichi Mélange. Significance of Field Surveys and Age Determination Field surveys along the Mukeng River and its tributary show that the Lichi Mélange lies west of the Taiyuan remnant forearc basin and is thrust over by the coherent-forearc-basin strata along the Tuluanshan Fault (Fig. 8). Neither the Lichi Mélange nor the volcanic rocks are observed in the Taiyuan remnant forearc basin east of the Tuluanshan Fault, consistent with predictions A and B in the preceding subsection. Biostratigraphic study reveals that the protolith of the Lichi Mélange was deposited in a narrow age range within Zone N19-20 of planktic foraminifers or Zone NN15 (3.5–3.7 Ma) of calcareous nannoplankton (Figs. 12–14; Chi et al., 1981; Chi, 1982; Barrier and Müller, 1984). This age is coeval with the lower remnant forearc basin sequence (Fig. 3) but is older than the upper forearc

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141

29

Sample Number

28

1 00 m

A

20 0 m

29

I

Age

DEEP-SEA TAXA

Ammonia spp. Fontbotia wuellerstorfi Amphistegina spp. Asterorotalia spp. Calcarina calcar Pullenia bulloides Lenticulina spp. Hyalinea balthica Operculina spp. Pseudorotalia spp. Cellanthus craticulatus Melonis spp. Elphidium spp. Cibicidoides spp.

Uvigerina spp.Gyroidinoides spp. Cyclammina spp.

0m

27

SHALLOW MARINE TAXA

N19; / 20 NN15 (3.7 ~ 3.5 Ma)

Unit

Sector 1 300 m

D

Sheared grades

C

Figure 12. (C) Occurrence and imbricate structure of the fractured, subangular to subrounded metasandstone conglomerates in two pebbly mudstone layers of broken formation facies in unit A. (D) Distribution of major benthic foraminifers in part of sector 1 along the southern tributary section (detailed position shown in Fig. 8).

sequence (Zones N21–22 of planktic foraminifers, Chang, 1967; Zones NN16–19 of calcareous nannoplankton, <3.5–1.15 Ma, Horng and Shea, 1996). This indicates that (1) the turbidites in the proto–Lichi Mélange were deposited in a short time (~0.2 m.y.), (2) deformation (at ca. 3.5 Ma) of the proto–Lichi Mélange occurred as soon as these turbidites were deposited (3.5–3.7 Ma), and (3) after deformation of the proto–Lichi Mélange the young turbidites (3.5–1 Ma) were deposited continuously in the upper part of the Taiyuan remnant forearc basin, a scenario analogue with the modern configuration of the Huatung Ridge–remnant North Luzon Trough off the southern Coastal Range in the modern active arc-continent collision zone (Fig. 7). The results of biostratigraphic study are consistent with prediction C in the preceding subsection.

Significance of the Benthic Foraminiferal Study Figure 15 shows the result of cluster analysis of the benthic foraminifers in the Lichi Mélange. No samples of the mélange facies (δ-grade of shearing) appear in Figure 15 because of their strong depression or their broken or incomplete preservation from intensive compression or shearing, which prevents precise taxonomic determinations from the few fossil remains. Therefore, the samples collected from mélange facies are omitted from computation. The samples are clustered into two main groups. Up to 79% of the samples are clustered together as group A, whereas 21% of the samples (only 8 samples) are clustered as group B. Samples of group B were collected primarily from the weakly sheared,

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Huang et al.

500 m 400 m 300 m 20 0 m 1 00 m

B

Uvigerina spp. Gyroidinoides spp. Cyclammina spp.

11

10-1 10-2 10-3 10-4

09-1 09-2

IIb

DEEP-SEA TAXA

Ammonia spp. Fontbotia wuellerstorfi Amphistegina spp. Asterorotalia spp. Calcarina calcar Pullenia bulloides Lenticulina spp. Hyalinea balthica Operculina spp. Pseudorotalia spp. Cellanthus craticulatus Melonis spp. Elphidium spp. Cibicidoides spp.

Sample Number

Unit

C

SHALLOW MARINE TAXA Age

Sector 2

09-3 08-1 08-2 07-1 07-2 06-2 06-1 05

N19 / 20 ; NN15 (3.7 ~ 3.5 Ma)

B

Sheared grades

A A

0m

IIa

Figure 13. (A) Examples of field occurrence and sampling localities of unit C. (B) Distribution of major benthic foraminifers along the Mukeng River section in the north.

broken formation facies (β-grade) turbidites of unit C (6 samples) and two individual samples from the dismembered facies (γ-grade) of unit IIb (sample 6–2) in sector 2 and unit III (sample 38–2) of sector 3 (Fig. 8). Foraminifers in group B predominate in the indigenous deep-water species without significant shallow-marine assemblages (Fig. 13B), whereas samples of group A contain both the indigenous deep-water and the displaced shallow-water fauna (Figs. 12–14). Moreover, each group A and group B contains the samples collected from the units of the dismembered facies and the broken formation facies (Fig. 15).

The result implies that the benthic foraminifers in the Lichi Mélange are similar to each other regardless of different degrees of shearing intensity, consistent with prediction D in the first subsection under “Discussion.” Foraminifers in the Taiyuan remnant forearc sequence in the southern Coastal Range were studied by Chang (1967). Unfortunately, the abundance of each benthic foraminifer was reported semiquantitatively (R, <6; F, 6–10; C, >11–20; A, >21– 50). A comparison of the present work with the result of Chang (1967) can be made by cross-checking the faunal lists and

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143

Sheared grades

D

500 m

Unit

570 m

Sector 2

Age

C SHALLOW MARINE TAXA

DEEP-SEA TAXA

Ammonia spp. Fontbotia wuellerstorfi Amphistegina spp. Asterorotalia spp. Calcarina calcar Pullenia bulloides Lenticulina spp. Hyalinea balthica Operculina spp. Pseudorotalia spp. Cellanthus craticulatus Melonis spp. Elphidium spp. Cibicidoides spp.

Uvigerina spp. Gyroidinoides spp. Cyclammina spp.

Sample Number 37-1 37-2 37-3 37-4

C

30 0 m

400 m

36 35 33-1 33-2 33-3 33-4

200 m

34

32

N19 / 20; NN15 (3.7 ~ 3.5 Ma)

37-5

100 m

II 31

0m

30

Figure 13. (C) Examples of field occurrence and sampling localities of unit C. (D) Distribution of major benthic foraminifers in part of sector 2 along the southern tributary section (detailed position shown in Fig. 8).

relative abundance of the major taxa in both studies (Table 1). The results suggest that the benthic foraminifers in the Lichi Mélange (this study) and the Taiyuan remnant forearc basin turbidites (Chang, 1967) are similar: 1. The benthic foraminifers are all characterized by high diversity (Taiyuan forearc basin turbidites: 40 genera, 68 species, in Chang, 1967; Lichi Mélange: 86 genera and 167 species, this study) and low abundance of each species (85% species < 6 individuals/200 g of sediments in the Lichi Mélange and coherent forearc basin turbidites). The difference in species number between these two studies is due to the different classification system

used. Our study uses a newer classification of benthic foraminifers by Loeblich and Tappan (1988), whereas Chang (1967) followed an older classification of Loeblich and Tappan (1964). 2. Planktic foraminifers are abundant (>85% of total foraminiferal individuals), and benthic foraminifers are predominantly calcareous genera (>75%) in both Taiyuan forearc turbidites and the Lichi Mélange. 3. Foraminifers in both Taiyuan forearc turbidites and the Lichi Mélange are composed of indigenous deep-marine fauna and shallow-marine species (Table 1). The deeper marine taxa include Bulimina striata d’Orbigny, B. cf.

A

Sector 3 4 80 m

Unit

SHALLOW MARINE TAXA Age

C

Sheared grades

B DEEP-SEA TAXA

Ammonia spp. Fontbotia wuellerstorfi Amphistegina spp. Asterorotalia spp. Calcarina calcar Pullenia bulloides Lenticulina spp. Hyalinea balthica Operculina spp. Pseudorotalia spp. Cellanthus craticulatus Melonis spp. Elphidium spp. Cibicidoides spp.

Uvigerina spp.Gyroidinoides spp. Cyclammina spp.

Sample Number

D

15-2 15-1

2 00 m

300 m

16-1

III

16-2 16-3 16-4 16-5 18 17 14 14-1

12-1 12-2 12-3 12-4 12-5 12-6

0m

1 00 m

13-1 13-2

N19 / 20 ; NN15 (3.7 ~ 3.5 Ma)

40 0 m

19

Figure 14. (A, B) Examples of field occurrence and sampling localities of units III and D, respectively. (C) Distribution of major benthic foraminifers in part of sector 3 along the Mukeng River section in the north.

26 D Age

Sheared grades

SHALLOW MARINE TAXA

DEEP-SEA TAXA

Ammonia spp. Fontbotia wuellerstorfi Amphistegina spp. Asterorotalia spp. Calcarina calcar Pullenia bulloides Lenticulina spp. Hyalinea balthica Operculina spp. Pseudorotalia spp. Cellanthus craticulatus Melonis spp. Elphidium spp. Cibicidoides spp.

Uvigerina spp.Gyroidinoides spp. Cyclammina spp.

Sample Number

600 m

Unit

Sector 3 7 00 m

E

IV

400 m

25

30 0 m

24

200 m

D

N19 /20 ; NN15 (3.7 ~ 3.5 Ma)

500 m

26

0m

100 m

23

22

Figure 14. (D) Examples of field occurrence and sampling locality 26 in mélange facies of unit IV. (E) Distribution of major benthic foraminifers in part of sector 3 along the southern tributary of the Mukeng River (detailed position shown in Fig. 8).

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Huang et al.

0 + 8-1 8-2 9-1 10-4 6-2 10-3 38-2 10-1 12-2 15-1 11 13-2 4-0 4-1 4-2 37-4 37-5 5 16-1 16-2 15-1 18 37-3 2 4-4 28 27 16-4 38-5 31 16-5 17 19 25 36 23 24 38-1 14

Mudstone Shale Mudstone Mudstone Shale Mudstone Mudstone Mudstone Mudstone Sandstone Mudstone Sandstone Siltstone Sandstone Sandstone Shale Mudstone Mudstone Mudstone Mudstone Shale Mudstone Shale Mudstone Mudstone Mudstone Mudstone Mudstone Mudstone Mudstone Mudstone Mudstone Mudstone Mudstone Mudstone Mudstone Mudstone Mudstone Mudstone

Rescaled Distance Cluster Combine 5 10 15 20 +

+

+

+

25 +

Group B

rostrata Brady, Uvigerina probocidea Schwager, U. hispida Schwager, U. subperegrina Cushman and Kleinpell, Fontbotia wuellerstorfi (Schwager), Gyroidinoides orbicularis (d’Orbigny), Pullenia bulloides (d’Orbigny), Melonis nicobarensis (Cushman), and M. pompilioides (Fichtel and Moll). These deep-marine benthic foraminifers are cosmopolitan species living in water depths of 1500–3000 m of modern deep oceans (Van Morkhoven et al., 1986). The shallow-marine species are Amphistegina radiata (Fichtel and Moll); A. lessonii d’Orbigny; Operculina complanata (Defrance); O. ammonoides (Gronovius); Asterorotalia gaimardii inermis Billman, Hottinger and Oesterle; A. subtrispinosa (Ishizaki); A. yabei (Ishizaki); Pseudorotalia indopacifica (Thalmann); P. schroeteriana (Parker and Jones); Calcarina calcar d’Orbigny; Cellanthus craticulatus (Fichtel and Moll); Elphidium crispum (Linné); and Lenticulina limbosa (Reuss). Today, these shallow-marine benthic

Group A

Figure 15. Multivariate cluster analysis of the samples collected from the Lichi Mélange, based on benthic foraminiferal assemblage. The study samples are clustered into two main groups. Each group includes samples with various degrees of shearing. For details of both groups, see the text.

foraminifers are typical of the Kuroshio Current fauna living in the outer shelf facing the open ocean (Wang, 1985; Huang, 1989). 4. The indigenous deep-marine benthic foraminifers in the Lichi Mélange and the Taiyuan remnant forearc basin turbidites are also the major fauna recovered in the present hemipelagic muds of the North Luzon Trough (Huang, 1993), whereas the shallow-marine species are commonly found ubiquitously in the turbidite layers off southern and northeastern Taiwan, including the North Luzon Trough forearc basin and the Okinawa Trough backarc basin. They were displaced from the shallow shelf by turbidity currents (Huang et al., 2005; Chiu, 2005). The conclusions of this benthic foraminiferal study thus indicate that the host rocks of the Lichi Mélange were part of the North Luzon Trough forearc sequences before deformation. The conclusions agree well with predictions D and E in the first subsection under “Discussion.”

The Lichi Mélange

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TABLE 1: COMPARISON OF BENTHIC FORAMINIFERAL TAXA AND THEIR ABUNDANCE IN THE LICHI MÉLANGE (THIS STUDY) AND THE TAIYUAN FOREARC TURBIDITES (CHANG, 1967) Lichi Mélange Taiyuan Forearc Basin (this study) (Chang, 1967) Shallow-marine taxa: Shallow-marine taxa: Amphistegina radiata (Fichtel and Moll) Amphistegina radiata (Fichtel and Moll) Operculina ammonoides (Gronovius) Amphistegina lessonii d’Orbigny Pseudorotalia gaimardii compressiuscula (Brady) Operculina complanata (Defrance) Pseudorotalia schroeteriana (Parker & Jones) Operculina ammonoides (Gronovius) Elphidium sp. Ammonia takanabensis (Ishizaki) Lenticulina atlantica (Barker) Asterorotalia gaimardii inermis Billman, Lenticulina calcar (Linné) Hottinger and Oesterle Lenticulina iota (Cushman) Asterorotalia subtrispinosa (Ishizaki) Lenticulina nikobarensis (Schwager) Asterorotalia yabei (Ishizaki) Pseudorotalia indopacifica (Thalmann) Lenticulina spp. P. cf. schroeteriana (Parker & Jones) Calcarina calcar d’Orbigny Cellanthus craticulatus (Fichtel and Moll) Elphidium crispum (Linné) Lenticulina limbosa (Reuss) Lenticulina orbicularis (d’Orbigny) Lenticulina iota (Cushman) Lenticulina cf. cushmani (Galloway and Wissler) Lenticulina spp. Deep-marine taxa: Bulimina striata d’Orbigny Fontbotia wuellerstorfi (Schwager) Hyalinea balthica (Gmelin) Cibicidoides cf. crebbsi (Hedberg) Cibicidoides subhaidingerii (Parr) Gyroidinoides neosoldanii (Brotzen) Gyroidinoides orbicularis (d’Orbigny) Pullenia bulloides (d’Orbigny) Melonis nicobarensis (Cushman) Melonis pompilioides (Fichtel and Moll) Uvigerina asperula Czjzek Uvigerina hispida Schwager Uvigerina peregrina Cushman Uvigerina proboscidea Schwager Uvigerina subperegrina Cushman and Kleinpell Cyclammina japonica Asano kaiensis Fukuta and Shinoki Cyclammina tani Ishizaki Cyclammina cf. trullissata (Brady)

Sediment Composition Our field surveys recognize that three distinct end-member facies are present in the Lichi Mélange: weakly sheared broken formation facies (β-grade), strongly sheared dismembered facies (γ-grade), and highly sheared mélange facies (δ-grade). Like the coherent forearc basin turbidites east of the Tuluanshan Fault, the weakly sheared, broken formation facies still preserves distinct turbidite sedimentary structures like the Bouma turbidite sequence Ta–Te, and bedding without any angular sheared ophiolite or sedimentary blocks. The turbidite layers of the broken formation facies are primarily of the classical sandy flysch type (Figs. 12–14; C and D facies of Mutti and Lucchi, 1978). The basal part of these classical turbidite layers (Bouma Ta–Tc) is composed primarily of quartz, like the lower part of the coherent forearc sequences. Slate chips, which are commonly found in

Deep-marine taxa: Bulimina striata d’Orbigny Fontbotia wuellerstorfi (Schwager) Gyroidinoides orbicularis (d’Orbigny) Pullenia bulloides (d’Orbigny) Melonis pompilioides (Fichtel and Moll) Uvigerina (Siphouvigerina) ampullacea Brady Uvigerina hispida Schwager Uvigerina crassicostata Schwager Uvigerina excellens Todd Uvigerina proboscidea Schwager Uvigerina subperegrina Cushman and Kleinpell Cyclammina orbicularis Brady Cyclammina pusilla Brady Cyclammina tani Ishizaki

the upper coherent forearc basin turbidites (<3 Ma; Teng, 1982), are never observed significantly in the Lichi Mélange. However, lenses of pebbly mudstones (each layer 1–2 m thick; Fig. 12C) are found in the weakly sheared broken formation facies (units A, C, and D in Fig. 8). These subangular to subrounded, matrixsupported conglomerates (each pebble 5–30 cm in size) are composed primarily of metasandstone and are commonly fractured (Fig. 12C). The orientation of these conglomerates shows imbricate structure (Fig. 12C), representing typical channel fillings derived from the exposed Eocene metasandstone beds in the eastern Central Range. Similar pebbly mudstone layers with a composite thickness of 150 m can be traced in the lower part of the entire Taiyuan remnant forearc basin (Yao et al., 1988). Such feeder-channel conglomerate beds, generally called the Shuilien Conglomerate, occur in stratigraphic levels in the remnantforearc-basin sequences of the Coastal Range (Fig. 3).

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Therefore, sedimentological features, including occurrence and composition of conglomerates in the pebbly mudstone layers and the turbidite structures preserved in the broken formation facies, support the view of the Lichi Mélange as part of the North Luzon Trough forearc basin sequences in agreement with prediction F in the first subsection under “Discussion.” Significance of Clay Mineral Composition Clay mineral compositions in the Coastal Range were studied by Wang and Yang (1975), Lin and Chen (1986), Buchovecky and Lundberg (1988), and Yao et al. (1988). Lin and Chen (1986) documented systematically that the clay mineral compositions of the muddy matrix and sedimentary blocks in various facies of the Lichi Mélange were similar to each other in containing illite, chlorite, kaolinite, and a little semectite, which are similar to the clay mineral assemblages in the Taiyuan remnant-forearc-basin turbidites except for kaolinite. This was not expected, because sedimentological and micropaleontological studies all show that the Taiyuan forearc basin turbidites and the Lichi Mélange were derived from the same source and were all deposited in the same North Luzon Trough. Consequently, they should contain similar clay mineral composition just as they contain similar pebble composition and the benthic foraminiferal fauna. Lin and Chen (1986) claimed that the Lichi Mélange could have two sediment sources: a continental, low-grademetamorphosed sediment source (accretionary prism of the Central Range) to provide abundant illite and chlorite, and a weathered, low-latitude, high-rainfall archipelagic volcanic source for derivation of kaolinite by eastward sedimentary mass slumping. In contrast, the Taiyuan coherent-forearc turbidite sequences could have only a continental low-grade-metamorphosed sediment source (Yao et al., 1988; Buchovecky and Lundberg, 1988). Therefore, there is no kaolinite in the Taiyuan coherent-forearcbasin sequences. Our present work confirms the results of Lin and Chen (1986) that there is significant kaolinite in the Lichi Mélange, regardless of shearing intensity, and agrees that the Lichi Mélange and the Taiyuan remnant-forearc-basin strata were all deposited in the North Luzon Trough. However, marine geology off the Coastal Range shows that the North Luzon Trough forearc basin lies west of the Luzon Volcanic Arc (Figs. 1 and 2). If the volcanic arc has provided the weathered kaolinite to the forearc basin, as Lin and Chen (1986) claimed, the Taiyuan forearc basin turbidites in the eastern part of the North Luzon Trough would have received more kaolinite than the Lichi Mélange in the western part of the North Luzon Trough, unless part of the volcanic arc had been thrust to exhume the metamorphic Central Range in eastern Taiwan and exposed as the source for providing kaolinite to the Lichi Mélange in the western part of the North Luzon Trough but not to the remnant Taiyuan forearc basin strata in the east. However, no such Neogene volcanic arc rocks are found in the eastern Central Range today. Moreover, south of lat 23° N the Taiyuan remnant forearc basin and the Chengkuangao volcanic island were accreted as the southern Coastal Range during the

last 1 m.y. (Huang et al., 2006). At that time the configuration of the Central Range–Coastal Range thus should not have differed much from how it looks today. This implies that it may be improbable to consider that the kaolinite in the Lichi Mélange was derived from the volcanic arc in the eastern Central Range by eastward sedimentary mass slumping. In the Coastal Range, kaolinite is found only in the volcanic rocks of the Tuluanshan Formation (Wang and Yang, 1975), suggesting that the source of kaolinite in the Lichi Mélange could have been related to the tectonic disturbance of the Tuluanshan volcanic rocks, as follows: 1. Field surveys show that the volcanic rocks of the Tuluanshan Formation beneath the Taiyuan forearc basin are thrust westward over the Lichi Mélange along the Tuluanshan Fault (Fig. 8), and this is the only region in which the Tuluanshan volcanics are exposed between the Lichi Mélange and the Loho-Taiyuan remnant forearc basins; 2. The volcanic sequence of the Tuluanshan Formation never appears within the remnant forearc basins; 3. Large, isolated, rootless volcanic blocks of volcaniclastics, andesitic agglomerates, and tuffs are present in the Lichi Mélange from Loho to Lichi (Fig. 2; Hsu, 1956, 1976). These large blocks are fault bounded in the dismembered or mélange facies of the Lichi Mélange. Near Antung, hot springs are present along the fault lines where the fault-bounded tuff and andesitic agglomerates occur in the Lichi Mélange. We believe that such thrusting, fragmentation, and mixing of the volcanic rocks within the mélange facies have allowed incorporation of the kaolinites from the Tuluanshan Formation into the entire Lichi Mélange, but not to the overlying coherent-forearcbasin strata, by fluid flow (Fisher, 1996) along the ubiquitously sheared planes or fractures within the Lichi Mélange. Therefore, characteristic occurrence of clay mineral composition in the Lichi Mélange does not conflict with the conclusion that the protolith of the Lichi Mélange was deposited in the western part of the North Luzon Trough forearc basin. Instead, occurrence of characteristic kaolinite in the Lichi Mélange is strongly related to the structures in the southern Coastal Range, showing tectonic involvement of the volcanic basement beneath the forearc basin in the formation of the Lichi Mélange during closure of the North Luzon Trough, when the Luzon arc-forearc was accreted to the southern Coastal Range during the last 1 m.y. Consequently, the conclusion of this clay-mineral-composition study is consistent with field surveys and prediction G in the first subsection under “Discussion.” Significance of the Whitish Quartz-Rich Feldspathic Sandstone Blocks in the Lichi Mélange In the Lichi Mélange, some allochthonous blocks—such as the angular, whitish, well-lithified, quartz-rich, feldspathic sandstone blocks and ophiolite blocks—occur in the block-in-matrix dismembered facies and mélange facies. These blocks have not

The Lichi Mélange been observed in the weakly sheared broken formation facies of the Lichi Mélange or in the remnant-forearc-basin sequences. The origin of these allochthonous sandstone blocks is problematic. Like the kaolinite, these whitish angular sandstone blocks have been interpreted as being derived from the exposed accretionary prism by eastward sedimentary slumping toward the western North Luzon Trough, the site of the Lichi Mélange (Page and Suppe, 1981; Lin and Chen, 1986). However, no independent evidence supports this slumping hypothesis. Similar angular, sheared, quartz-rich sandstone blocks (1– 15 cm in size), embedded in the late Pliocene sheared mudstone (ca. 3 Ma), have been dredged from the modern Huatung Ridge but have never been collected from the remnant North Luzon Trough off the Coastal Range (Huang et al., 1992; Huang, 1993). The sandstone blocks and the associated mudstone in the Huatung Ridge contain not only illite and chlorite but also significant kaolinite similar to what we found in the Lichi Mélange (Huang et al., 1992). To understand the significance of these allochthonous sandstone blocks in the Lichi Mélange, we studied fission tracks of zircon grains, separated from the whitish quartz-rich feldspathic sandstone blocks in the Lichi Mélange, and also separated from the coarse grains in the Taiyuan remnant-forearc-basin turbidite layers. Fission track analyses of the zircon grains, separated from the angular, whitish sandstone blocks (>2 m) in the dismembered facies and the mélange facies in Fuli, and the type locality of the Lichi Mélange (Fig. 8), show a partial annealing feature with two age peaks concentrated at the early–middle Miocene (9–18 Ma) and the late Mesozoic (66–130 Ma; Fig. 9A). However, the zircon grains of the Taiyuan remnant-forearc-basin turbidites show a distinctly different fission track pattern: one that is completely reset in a tuff sample in the lower part of the forearc sequence (2.08 Ma in sample TF-8; Fig. 9B) or partially reset with a young peak of Pliocene age (2.5 Ma in sample TF-6; 3.0 Ma in sample TF-7; 2.7 Ma in sample TF-9) and an old tail of Mesozoic age (64–100 Ma; Fig. 9B). Although the fission track records of the zircon grains all show a similar partial annealing feature, there is a significant difference between them: a higher frequency of a young peak (<7 Ma) in the Taiyuan remnant-forearc-basin turbidites (Fig. 9B), but a higher frequency of an old peak (>66 Ma, Mesozoic) in the angular whitish sandstone blocks of the Lichi Mélange (Fig. 9A). The frequency for the Mesozoic age in the Lichi Mélange samples is as high as, or even higher than, the frequency for the Neogene age. On the contrary, the coherent forearc turbidites show a distinct late Neogene age peak but a very low frequency for a Mesozoic age. Taiwan has undergone subduction and collision tectonics since the Miocene. The fission track study of zircons has been used successfully to constrain the source of sediment provenance and geothermal history during the subduction and collision tectonics of Taiwan (Liu, 1982; Liu et al., 2001; Huang et al., 2006; Fuller et al., 2006). These studies have shown that fission tracks of zircon grains from the Miocene deep-marine turbidites of the

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accretionary prism (the western Central Range; Fig. 1) were partially reset, and the zircons separated from the exhumed metamorphic basement (eastern Central Range; Fig. 1) were totally reset. Our samples all show partial annealing characteristics, indicating that they were all derived from the accretionary prism. But the young peak age of the Taiyuan remnant-forearc-basin strata (2.7– 2 Ma) is much younger than that of the sheared whitish sandstone blocks of the Lichi Mélange (9–18 Ma), suggesting that not only the source, but also the mechanism, for emplacing these sheared angular sandstone blocks into the Lichi Mélange may have significantly differed from the turbidite sedimentation in the Taiyuan remnant forearc basin. Moreover, these angular, sheared, whitish quartz-rich sandstone blocks have not been observed in the Taiyuan remnant-forearc-basin strata, indicating that they may not have been delivered to the Lichi Mélange by turbidity currents as Taiyuan remnant-forearc-basin turbidite layers. To identify the source of these quartz-rich, feldspathic sandstone blocks in the Lichi Mélange, we also analyzed four samples collected from the deep-sea sandstones in the accretionary prism of the Hengchun Peninsula for comparison (Fig. 10). These samples also show a partial reset feature with two age peaks: a young Neogene peak (15–30 Ma) and an old Mesozoic peak (50–100 Ma; Fig. 10). The frequency of the Mesozoic peak is as high as the Neogene peak, similar to that found in the whitish sandstone blocks of the Lichi Mélange (Fig. 9A). These sandstone blocks in the Lichi Mélange were determined to be late Miocene in age (Chi, 1982; Barrier and Müller, 1984), consistent with the deep-sea-fan sandstones in the accretionary prism of the Hengchun Peninsula (Chang, 1966; Huang et al., 1997) but older than the Pliocene–Pleistocene Taiyuan remnant-forearc-basin strata and the Pliocene protolith of the Lichi Mélange. Three routes are possible through which the late Miocene deep-sea-fan sandstones in the late Miocene accretionary prism could have been transported to the Pliocene dismembered facies and mélange facies of the Lichi Mélange: (1) by erosion from the exposed accretionary prism (now the Central Range, before exhumation of the underthrust Eurasian continental materials during the last 2.5 m.y.; Huang et al., 2006) and then transported to the North Luzon Trough by turbidity flow longitudinally from north to south along the trough axis; (2) by erosion as (1) but transported eastward by mass slumping into the western part of the North Luzon Trough (Page and Suppe, 1981); or (3) by eastward thrusting to move the non-exposed late Miocene accretionary prism materials upward into the deformed forearc strata (proto–Lichi Mélange, like the modern Huatung Ridge) in the western part of the North Luzon Trough. It is improbable that the first and second routes could have caused the difference in the fission-track age patterns of zircons between the Lichi Mélange and the Taiyuan remnant forearc turbidites, because both processes would have provided similar zircon grains from the same source (the exposed accretionary prism) to the North Luzon Trough. Therefore, they should have similar fission-track age patterns in the eastern or the western part of the trough either by longitudinal turbidity flow or by eastward

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mass slumping. In contrast, the third route could have provided the zircon grains from different sources by different mechanisms independently to the Lichi Mélange and to the Taiyuan remnantforearc-basin turbidites. As the marine seismic profiles (Fig. 7) demonstrate, we believe that the blocks of whitish sandstone could have been thrust eastward from the non-exposed accretionary prism unconformably beneath the western North Luzon Trough forearc basin to the deformed proto–Lichi Mélange (Huatung Ridge) in the western part of the North Luzon Trough, instead of having been derived from the exposed accretionary prism by turbidity flows or mass slumping downward to the entire North Luzon Trough. By such eastward backthrusting, the late Miocene deep-sea sandstone is found restricted to the strongly sheared, block-in-matrix, dismembered facies and the mélange facies of the Lichi Mélange but not to the weakly sheared relict, broken formation facies of the Lichi Mélange or the Taiyuan remnantforearc-basin strata. This eastward thrusting mechanism is exactly like that of the Huatung Ridge, where we found similar angular, whitish sandstone blocks with kaolinite, or where the proto–Lichi Mélange is developing (Fig. 7), and is consistent with prediction G in the first subsection under “Discussion.” Source of the East Taiwan Ophiolite Dismembered ophiolite blocks are well known in the Lichi Mélange in the Coastal Range, eastern Taiwan, and are commonly known as the East Taiwan Ophiolite. However, the source and the mechanism for emplacing these oceanic-affinity blocks within the Lichi Mélange are still the subject of debate. Along the Mukeng River the number and size of these blocks (usually 0.1–2 m) are much less common and are smaller than those in the Kuanshan and Lichi areas (Fig. 1). Along the Mukeng River sections the East Taiwan Ophiolite blocks do not occur with the whitish sandstone blocks. However, they both appear to be restricted to the intensely sheared, dismembered facies or the mélange facies of units I, II, III, and IV (Fig. 8). In the Lichi area the larger East Taiwan Ophiolite blocks (several to hundreds of meters) do occur commonly with the whitish sandstone blocks (Liou et al., 1977; Page and Suppe, 1981). There are two possibilities for the source of the East Taiwan Ophiolite blocks: the South China Sea oceanic crust or the oceanic basement beneath the Luzon Arc-forearc of the Philippine Sea plate. If the East Taiwan Ophiolite was part of the South China Sea oceanic crust, its age would be younger than late Oligocene because this sea opened in 32–17 Ma (Taylor and Hayes, 1980). Microfossils recovered from a thin red shale intercalated within this ophiolite sequence indicated an age of 15 Ma (Zone NN5 of calcareous nannoplanktons; Huang et al., 1979), which would be close to the last phase of the opening of the South China Sea. This conclusion may favor an oceanic-crust source for the South China Sea (Suppe et al., 1981). However, no East Taiwan Ophiolite–like rock occurs in the present Central Range. These ophiolite blocks, presumably of a South China Sea source, could not have been derived from the exposed accretionary prism (Central Range)

by sedimentary slumping, because in the Lichi area the ophiolite blocks occur with the whitish sandstone blocks, which were improbably emplaced into the Lichi Mélange by mass slumping (discussed previously). Therefore, like the concurrent late Miocene angular sandstone blocks, the East Taiwan Ophiolite also could have been thrust eastward from the accretionary prism of the Hengchun Peninsula during the early stage of arc-continent collision. However, in the accretionary prism the ophiolite blocks are observed only in the Kenting Mélange of the Hengchun Peninsula (Fig. 10), and they differ from the East Taiwan Ophiolite in composition. For example, the chromatite that is commonly found in the Kenting Mélange of the Hengchun Peninsula (Chu et al., 1988) has never been observed in the East Taiwan Ophiolite. Unlike the deep-sea-fan sandstones that occur ubiquitously in the Hengchun Peninsula (Huang et al., 1997), the ophiolite blocks in the Kenting Mélange are present in a highly limited area (Fig. 10). Therefore, it would have been improbable for these ophiolite blocks to have been thrust eastward from the accretionary prism of the Hengchun Peninsula. From a regional tectonic view, we prefer the interpretation that the East Taiwan Ophiolite may represent the oceanic crust of the Philippine Sea Plate (Juan et al., 1980; Malavieille et al., 2002). These ophiolite blocks could have been westward thrust from the oceanic crust beneath the North Luzon Trough forearc basin or the Luzon Arc basement during the final westward accretion of the Luzon Arc-forearc to eastern Taiwan during the last 1 m.y. (Fig. 5C, 3). The age of the Luzon Arc is no later than middle Miocene, consistent with the age of the East Taiwan Ophiolite. The occurrence of ophiolitic rocks in the Banda forearc basement has been documented from the Bobonaro Mélange of the Timor arc-continent collision orogen by thrusting emplacement (Harris et al., 1998). It is also possible that the East Taiwan Ophiolite could have been developed by suprasubduction (Pearce et al., 1984) by which the lower crustal materials, such as serpentinite and peridotite, could be found in a forearc basin setting, similar to the modern Mariana forearc (Fryer et al., 1985). Global Significance of the Lichi Mélange: Comparison with the Bobonaro Mélange in the Timor Arc-Continent Collision Orogen Mélange is generally considered to be developed within the accretionary prism by scraping off the underthrusting oceanic crust and the overlying deep-sea marine sediments along the décollement fault in the subduction zone (Silver and Beutner, 1980; Cowan, 1985; Cloos and Shreve, 1988), and the mechanism responsible for formation of mélange could be tectonic thrusting, sedimentary slumping, or mud diapirism (Cowan, 1985; Barber et al., 1986). However, the Pliocene Lichi Mélange was developed in the forearc basin rather in the accretionary prism. The sedimentary mass-slumping process is unlikely as the major mechanism responsible for formation of the Lichi Mélange, as discussed previously. Today, two small mud volcanoes, each 100–200 wide, are found within the Lichi Mélange (Yang et al.,

The Lichi Mélange 2004). However, in comparison with the wide distribution of the Lichi Mélange (50 km long, 2 km wide), such small areas of mud volcanoes could not have played an important role in the development of the Lichi Mélange. Today, mélange can be found also in the Timor arc-continent collision region. The Bobonaro Mélange in Timor Island is characterized by active development, long extension from offshore to onshore for more than 2000 km, a young age (since 3 Ma), and good exposures. Other characteristics include a wide variety of rock types in block-in-matrix facies caused by polymechanisms of tectonic detachment, mud diapirism, and sedimentary slumping (Barber et al., 1986; Harris et al., 1998). In comparing the Lichi Mélange in the Coastal Range, eastern Taiwan, with the Bobonaro Mélange in Timor, of the Banda Arc region, it is interesting to note similarities and differences. For example, the Lichi Mélange and the Bobonaro Mélange are both of a young age (ca. 3 Ma), forming mélanges today in active arc-continent collision tectonics, traceable from onshore to offshore, involving forearc basement tectonics, and showing variable degrees of rock shearing, fragmentation, and mixing from weakly sheared broken formation facies to block-in-clay mélange facies. However, the Lichi Mélange developed in a forearc basin setting rather than in an accretionary prism setting, as did the Bobonaro Mélange. Therefore, no continental fragment of the lower plate, like the Australian pre- and post-rifting sequences in the Bobonaro Mélange, was involved during formation of the Lichi Mélange. On the other hand, the Luzon Volcanic Arc was involved during formation of the Lichi Mélange, but the Banda Volcanic Arc was not involved in formation of the Bobonaro Mélange. The microfossil composition of the Lichi Mélange (Pliocene) is much simpler than that of the Bobonaro Mélange (Jurassic to Pleistocene). Deep-sea pelagic sediments such as radiolarian ooze are present in the Bobonaro Mélange, but only hemipelagic muds and turbidites in the Lichi Mélange. Mud diapirism and sedimentary mass slumping were insignificant in the development of the Lichi

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Mélange, but they played an important role in the development of the Bobonaro Mélange. CONCLUSIONS New seismic surveys along the 21° N transect across the North Luzon Trough show a syndeformational sedimentation in the forearc basin. Episodic deformation has occurred in the western part of the North Luzon Trough since the early history of forearc sedimentation by the active collision between the Luzon Arc and the Eurasian continent. The deformation resulted in formation of the Huatung Ridge, which continues northward to the Lichi Mélange in the southernmost Coastal Range, suggesting a direct geological relationship between formation of the Lichi Mélange and the submarine Huatung Ridge. Multiple lines of new evidence, including on-land field surveys, planktic and benthic foraminiferal study, clay mineral composition, and zircon fission-track analyses, confirm that the Lichi Mélange was part of the North Luzon Trough forearc sequences (Fig. 5C, 1), representing a collision mélange formed along arcward (eastward) backthrusting during forearc closure when the Luzon arc collided incipiently with the Eurasian continent. The formation of the Lichi Mélange involved two thrust events: (1) early arcward backthrusting in the western part of the North Luzon Trough to form the proto–Lichi Mélange, like the Huatung Ridge, during progressive closure of the forearc basin in the initial arc-continent collision stage (ca. 3 Ma; Figure 5C, 2), followed by (2) westward thrusting of the Luzon Arc-forearc onto eastern Taiwan to form the present southern Coastal Range during the advanced arc-continent collision stage during the last 1 m.y. (Fig. 5C, 3). The Lichi Mélange provides a unique example in comparison with worldwide orogenic belts in that the mélange has developed by tectonic thrusting along closure of the forearc basin in the region of active arc-continent collision.

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APPENDIX

APPENDIX. CLAY MINERAL COMPOSITION (IN PERCENTAGES) OF CLAY FRACTIONS ALONG THE MUKENG RIVER AND ITS SOUTHERN TRIBUTARY SECTIONS IN THE LICHI MÉLANGE No.

Sample

Lithology

Illite

Chlorite

Kaolinite

Smectite

Mixed-layer

E

MK01-1-

Sheared mudstone

47

18

13

Trace to absent

22

E

MK02-

Sheared mudstone

43

20

9

Trace to absent

28

E

MK03-0-

Non-sheared silty mudstone

56

11

6

Trace to absent

27

E

MK03-3-

Slightly sheared mudstone

57

14

8

Trace to absent

21

E

MK03-4-

Sheared mudstone

48

19

12

Trace to absent

21

E

MK04-

Sheared mudstone

52

22

10

Trace to absent

16

C

MK06-1-

Sheared mudstone

47

14

11

Trace to absent

28

C

MK10-

Sheared mudstone

51

21

10

Trace to absent

18

B

MK12-

Non-sheared mudstone

56

18

10

Trace to absent

16

B

MK13-1-

Sheared mudstone

52

19

11

Trace to absent

18

D

MKN01-

Non-sheared mudstone

48

22

11

Trace to absent

19

A

MKN02-2-

Slightly sheared mudstone

47

19

12

Trace to absent

22

A

MKN02-4-

Slightly sheared mudstone

49

19

11

Trace to absent

21

A

MKN02-6

Sheared mudstone

59

12

11

Trace to absent

18

A

MKN03-1-

Sheared mudstone

53

21

11

Trace to absent

15

A

MKN04-

Sheared mudstone

56

21

10

Trace to absent

13

A

MKN05-2-

Non-sheared shale

53

16

10

Trace to absent

21

A

MKN05-4-

Non-sheared mudstone

44

23

14

Trace to absent

19

A

MKN05-6-

Sheared mudstone

47

20

10

Trace to absent

23

A

MKN07-

Sheared mudstone

39

17

15

Trace to absent

29

A

MKN08-

Non-sheared mudstone

63

19

14

Trace to absent

4

F

MKS01-

Non-sheared shale

67

21

8

Trace to absent

F

MKS02-

Non-sheared mudstone

52

14

13

F

MKS03-

Sheared mudstone

53

15

14

18

F

MKS03S-

Non-sheared siltstone

46

19

14

21

F

MKS04-2-

Non-sheared mudstone

64

23

7

6

F

MKS04-4-

Sheared mudstone

63

19

12

6 19

4 21

F

MKS05-1-

Non-sheared mudstone

53

18

10

F

MKS05-2-

Non-sheared mudstone

55

18

10

17

F

MKS06-

Non-sheared mudstone

52

19

12

17

G

MKS08-

Sheared mudstone

52

16

9

23

G

MKS09-

Mud in conglomerate layer

52

19

9

20

G

MKS10-

Sheared mudstone

53

18

11

18

TF-06a

Mudstone

54

24

6

16

TF-07

Mudstone

60

23

5

12

TF09

Mudstone

50

21

8

21

LM-1 mud

Mudstone

51

19

14

16

YF02C

Mudstone

53

18

6

23

YF03C

Mudstone

54

16

15

15

YF01C

Sandstone

28

50

15

7

LM-1

Sandstone

42

23

22

13

LM01

Sandstone

38

27

14

21

FK01C

Sandstone

42

15

17

16

Average:

51.5

1 9.5

11.1

17.8

For comparison, average clay mineral composition of Lin and Chen (1986):

43.7

9.8

13.7

3.7

28.9

The Lichi Mélange ACKNOWLEDGMENTS This study was financially supported by grants from the National Science Council (NSC93–2116-M-006–001; NSC94–2116M006–005) and the Research Center of Ocean Environment and Technology, NCKU, to C.Y. Huang, and grant G2000046705 from the National Important Basic Research and Development Projects to B. Yao. We appreciate the crew members of the R/V Tanbao for their great efforts during the seismic surveys across the North Luzon Trough off southern Taiwan. The authors also appreciate the constructive discussions with R. Harris and D. Reed in comparing the Lichi Mélange in eastern Taiwan with the Bobonaro Mélange in Timor. The manuscript was improved by comments from A. Basu and an anonymous reviewer. REFERENCES CITED Barber, A.J., Tjokrosapoetro, S., and Charlton, T.R., 1986, Mud volcanoes, shale diapirs, wrench faults and mélanges in accretionary complex, eastern Indonesia: American Association of Petroleum Geologists Bulletin, v. 70, p. 1729–1741. Barrier, E., and Müller, C., 1984, New observations and discussions on the origin and age of the Lichi Mélange: Geological Society of China Memoir, v. 6, p. 303–325. Biq, C., 1971, Comparison of mélange tectonics in Taiwan and in some other mountain belts: Petroleum Geology of Taiwan, no. 9, p. 79–106. Biq, C., 1973, Kinematic pattern of Taiwan as an example of actual continentarc collision: Taipei, Report of the Seminar on Seismology: US-ROC Cooperative Science Program, abstract, p. 21–26. Buchovecky, E.J., and Lundberg, N., 1988, Clay mineralogy of mudstones from the southern Coastal Range, eastern Taiwan: Unroofing of the orogen versus in-situ diagenesis: Acta Geologica Taiwanica, no. 26, p. 247–261. Chai, B.H.T., 1972, Structure and tectonic evolution of Taiwan: American Journal of Science, v. 272, p. 389–422. Chang, C.P., Angelier, J., and Huang, C.Y., 2000, Origin and evolution of a mélange: The active plate boundary and suture zone of the Longitudinal Valley, Taiwan: Tectonophysics, v. 325, p. 43–62, doi: 10.1016/S00401951(00)00130-X. Chang, C.P., Angelier, J., Huang, C.Y., and Liu, C.S., 2001, Structural evolution and significance of a mélange in a collision belt: The Lichi Mélange and the Taiwan arc-continent collision: Geological Magazine, v. 138, p. 633–651. Chang, L.S., 1966, A biostratigraphic study of the Tertiary in the Hengchun peninsula, Taiwan, based on smaller foraminifera (III: Southern Part): Proceedings of the Geological Society of China, no. 9, p. 53–63. Chang, L.S., 1967, A biostratigraphic study of the Tertiary in the Coastal Range, eastern Taiwan, based on smaller foraminifera (I: Southern Part): Proceedings of the Geological Society of China, no. 10, p. 64–76. Chen, C.H., Shieh, Y.N., Lee, T., Chen, C.H., and Mertzman, S.A., 1990, NdSr-O isotopic evidence for source contamination and an unusual mantle component under Luzon arc: Geochimica et Cosmochimica Acta, v. 54, p. 2473–2483, doi: 10.1016/0016-7037(90)90234-C. Chen, W.S., 1997, Mesoscopic structures developed in the Lichi Mélange during the arc-continent collision in the Taiwan region: Journal of the Geological Society of China, v. 40, p. 415–434. Cheng, Y.M., Huang, C.Y., Yeh, J.J., and Chen, W.S., 1984: The Loshui Formation: Deeper water sandstones in the Hengchun Peninsula, southern Taiwan: Acta Geologica Taiwanica, no. 22, p. 35–52. Chi, W.R., 1982, The calcareous nannofossils of the Lichi mélange and the Kenting mélange and their significance in the interpretation of plate tectonics of the Taiwan region: Ti-Chih (Geology), v. 4, p. 99–112. Chi, W.R., Namson, J., and Suppe, J., 1981, Stratigraphic record of plate interactions in the Coastal Range of eastern Taiwan: Geological Society of China Memoir, v. 4, p. 155–194. Chien, C.W., 2003, Forearc origin of the Lichi Mélange in Coastal Range, eastern Taiwan: Micropaleontological evidence [M.S. thesis]: Institute of Geosciences, National Taiwan University, 74 p.

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MANUSCRIPT ACCEPTED BY THE SOCIETY 24 APRIL 2007

Printed in the USA

The Geological Society of America Special Paper 436 2008

Oblique subduction in an island arc collision setting: Unique sedimentation, accretion, and deformation processes in the Boso TTT-type triple junction area, NW Pacific Yujiro Ogawa* Doctoral Program in Earth Evolution Sciences, University of Tsukuba, Tsukuba 305-8572, Japan Yoshihiro Takami College of Natural Science, University of Tsukuba, Tsukuba 305-8571, Japan Sakiko Takazawa* College of Natural Science, University of Tsukuba, Tsukuba 305-8571, Japan

ABSTRACT The NW corner of the Pacific Ocean is a place of unique Tertiary tectonism, which provides one of the clearest examples of arc-arc collision. Voluminous Cretaceous rhyolitic-granitic magmatism along the continental margin continues into the Paleogene. In contrast, Miocene island arc volcanism follows Eocene boninitic magmatism in the Izu-Mariana Arc, in association with the opening of backarc basins, including those in the Philippine and Japan Seas. The triple junction between the Eurasian, Philippine Sea, and Pacific plates arrived in the area south of Tokyo during the Miocene, just as the Japan Sea was opening. After the beginning of Philippine Sea plate subduction to the north, the Izu Island Arc began to collide obliquely with the Honshu Arc. As a result, this unique tectonic setting in the NW Pacific has produced a miniature Alpine-type orogenic belt (Tanzawa) in the collisional center, whereas in the eastern part of the Izu Arc sediment has been actively accreting in that forearc. Such settings have resulted in systematic accretionary prism formation from the early Miocene in the Boso-Miura peninsular area to the present in the Sagami Trough area. We modeled the tectonics by a simple sandbox experiment. Systematic fault and fracture patterns of the oblique subduction type are predicted to occur during arc-arc collision. Keywords: island arc collision, triple junction, forearc sedimentation, forearc sliver, accretionary prism, analogue modeling.

*Corresponding author, Ogawa: [email protected]. Present address, Takazawa: Ministry of Economy, Trade and Industry of Japan, Kasumigaseki, Tokyo 100-8901, Japan. Ogawa, Y., Takami, Y., and Takazawa, S., 2008, Oblique subduction in an island arc collision setting: Unique sedimentation, accretion, and deformation processes in the Boso TTT-type triple junction area, NW Pacific, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 155–170, doi: 10.1130/2008.2436(07). For permission to copy, contact editing@ geosociety.org. ©2008 The Geological Society of America. All rights reserved.

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INTRODUCTION Plate tectonic triple junctions provide special geologic and tectonic conditions in which unique topography, sedimentation, and deformation can form. Among the various types of triple junctions (McKenzie and Morgan, 1969), a trench-trench-trench (TTT) type of triple junction has been observed only off the Boso Peninsula, central Japan, in the northwestern corner of the Pacific Ocean. This is called the Boso triple junction (Ogawa et al., 1989; Seno et al., 1989) (Fig. 1). At present, another similar type of convergent triple junction is known nearby just north of Mount Fuji, which is called the Fuji triple junction. Thus, in a narrow area beneath the Tokyo region, two converging triple junctions make a very complicated tectonic setting (Sato et al., 2005). The geometry and tectonics of the Boso triple junction were discussed by Seno et al. (1989) and Huchon and Labaume (1989). The triple junction was originally called the off-Boso triple junction by Ogawa et al. (1989). The “central Japan triple junction” of Lallemant et al. (1996) is not an adequate term, because there are two triple junctions in central Japan. Seismic profiling and related submarine researches were performed during the JapanFrance KAIKO project (Renard et al., 1987; Nakamura et al., 1987; Pautot et al., 1987). After this project the principal results were published by Ogawa et al. (1989), Seno et al. (1989), Soh et al. (1988, 1990), and Lallemant et al. (1996), and many submarine and on-land research projects were conducted. Most important is the recognition that the northwestern part of the triple junction in the Sagami Trough is a place of active sedimentation, deformation, and accretion along the oblique subduction boundary caused

by the collision of the Izu Arc with the Honshu forearc. However, the geology and tectonics have not yet been fully explained with reference to the oblique subduction setting. In this paper we first review the submarine topography and geology along the present Sagami Trough area on the basis of the marine work of Kato et al. (1983), Kong et al. (1984), Ohkouchi (1990), and Iwabuchi et al. (1990), and many other works related to the KAIKO project as well as the terrestrial geology of Ogawa and Horiuchi (1978), Ogawa (1982), and Hanamura and Ogawa (1993), and others. We then present new evidence, including simple two-dimensional (2D) and three-dimensional (3D) analogue shear-box experiments for oblique subduction (Takami, 1999; Takazawa, 2003) in order to explain the uniqueness of the TTT type of the triple junction area in terms of the present sedimentation and deformation in an oblique subduction arc-arc collisional plate boundary. TECTONIC SETTING The northwestern Pacific Rim has a long tectonic history, dating from the Mesozoic and characterized by JurassicCretaceous Pacific plate subduction under NE Asia. The Japanese main islands are the products of such long-term subduction (Seno and Maruyama, 1984; Maruyama and Seno, 1986; Taira et al., 1989; Ogawa et al., 1997; Isozaki, 1996). Since the Oligocene and early to middle Miocene, when the Shikoku Basin, the easternmost backarc basin of the Philippine Sea plate, opened, the present Izu-Mariana Island Arc front has been moving from west to east just south off the shore of Honshu (Kobayashi, 1984,

Figure 1. Index map of the Izu Arc collision zone with the Honshu Arc since the middle Miocene (left) to the present (right). Izu forearc accretion occurs on the NW of the Boso triple junction. Plates A, B, and C are the Pacific, Philippine Sea, and Eurasian plates, respectively. Plate C, just on the NW side of the triple junction, was converted to the North American plate during the Quaternary, when the boundary between the North American and Eurasian plates extended to central Honshu. After Hanamura and Ogawa (1993).

Oblique subduction in an island arc collision setting 1995). After seafloor spreading in the Shikoku Basin ceased, the Izu Island Arc and its NE edge, coincident with the triple junction, arrived close to its present position. Subsequently the Shikoku Basin lithosphere began to subduct to the NE under Honshu (Seno and Maruyama, 1984). Consequently, the Izu Island Arc is now connected to central Honshu, but it also extends as far south as the Mariana Arc. The ridge of the Izu-Mariana Arc itself has collided with the Honshu Arc in central Honshu in several stages (Matsuda, 1978; Ogawa, 1985; Ogawa et al., 1985; Soh et al., 1988; Aoike, 1999; Sato et al., 1999). The timing of the first collision is synchronous with the opening of the Japan Sea at 15 Ma (Niitsuma, 1988; Takahashi and Saito, 1997). After that time the northeastern corner of the Philippine Sea plate has supplied Izu-Mariana forearc materials to the Honshu forearc that have been accreted to form the Miura-Boso accretionary prisms (Figs. 2 and 3) (Ogawa, 1982, 1985; Ogawa and Taniguchi, 1988; Hanamura and Ogawa, 1993; Takahashi and Saito, 1997; Yamamoto et al., 2000, 2005). The age of the Izu Arc extends at least to the Eocene (ca. 45 Ma; Seno and Maruyama, 1984). According to the age determination and spreading history of the Philippine Sea plate (Kinoshita, 1980; Okino et al., 1994; Kobayashi et al., 1995), the arc was once far to the south, in the Southern Hemisphere, and it rotated clockwise during its northward movement to its present position (Seno and Maruyama, 1984; Hall et al., 1995; Takahashi and Saito, 1997). Because of spreading in the backarc basins following rifting of the volcanic-island-arc ridges, the oldest parts of the Philippine Sea plate are now in both the westernmost and easternmost parts of the present Philippine Sea plate (Kobayashi, 1984, 1995; Kobayashi et al., 1995). The easternmost part of the Philippine Sea plate, the remnant of the leading edge of the Izu forearc at the trenches, is composed of ophiolitic rocks, ranging from Cretaceous to Miocene, but is dominated by Eocene boninitic rocks and younger island arc volcanic materials (Fujioka and Sakamoto, 1999; Ishiwatari et al.,

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2006). A similar stratigraphy is preserved as the remnant of the “Mineoka plate” and exposed onshore as the Mineoka Ophiolite in the Boso Peninsula (Fig. 4) (Ogawa and Taniguchi, 1988; Sato et al., 1999; Sato and Ogawa, 2000; Ogawa and Takahashi, 2004). Similarly, parts of the Philippine Sea plate are found in northern New Guinea as the North New Guinea plate (Seno, 1984; Lus et al., 2004), comprising Cretaceous to Eocene basaltic rocks and their pelagic cover (Ogawa and Taniguchi, 1987, 1988; Mohiuddin and Ogawa, 1998; Hirano et al., 2003; Ogawa and Sashida, 2005). Such oceanic plate history is now preserved in the Boso Peninsula and has been documented by the detailed structural and stratigraphic analysis of the Mineoka Ophiolite Belt (Hirano et al., 2003; Takahashi et al., 2003; Ogawa and Takahashi, 2004; Mori and Ogawa, 2005). The Mineoka Ophiolite Belt has lain in a forearc sliver fault zone since the middle Miocene, after which time there has been a systematic development of accretionary prisms south of the BosoMiura Peninsulas (Ogawa, 1983; Ogawa and Taniguchi, 1988; Lallemant et al., 1996). These prisms are formed of Izu forearc and paleo–Sagami Trough sediments derived from an area east of the Izu collision zone (Fig. 2) (Ogawa, 1982, 1985; Ogawa and Taniguchi, 1987, 1988; Hanamura and Ogawa, 1993; Yamamoto et al., 2000, 2005; Yamamoto and Kawakami, 2005). Thus the Sagami Trough is now the modern oblique subduction boundary at the northeasternmost corner of the Philippine Sea plate and has been a place of sedimentation and deformation between the Izu volcanic and Honshu continental arcs (Figs. 2 and 4). The Sagami Trough is also a place of large hazardous earthquakes in the historic and present ages, as typified by the Kanto earthquakes (Shishikura, 2003; Sato et al., 2005). The relative present motion of the Philippine Sea plate to the North American plate is toward N35°W at a rate of 2.5 cm/yr (Fig. 4; Seno, 1993). The Miura-Boso Peninsulas contain good outcrops of Miocene to Pliocene Izu forearc, deep-sea, volcaniclastic deposits that are now incorporated within accretionary prisms and covered

Figure 2. Schematic view of the model for the accretion of Izu forearc sediments to the Honshu forearc along the protoSagami Trough. View from the south. A voluminous amount of Izu volcaniclastic sediments plus some trench axial channel sediments from the Honshu Arc were deposited both in the Izu forearc and the proto–Sagami Trough, finally to be accreted to the Honshu forearc. Adapted from Hanamura and Ogawa (1993).

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Figure 3. Simplified trace map of the Tertiary formations on the Miura (M, west) and Boso (B, east) Peninsulas. Shaded areas represent lower Miocene and older sedimentary rocks. Dashed doubled lines approximately trace the present plate subduction–related faults. Neogene sediments to the north of the Mineoka Ophiolite Belt are composed of forearc basin sediments, whereas those to the south are mostly of accretionary prism material. After Ogawa and Taniguchi (1988).

unconformably with shallow-water, coarser grained, volcaniclastic sedimentary rocks of the forearc basin or the trench slope type (Hanamura and Ogawa, 1993; Yamamoto and Kawakami, 2005) (Figs. 2 and 3). The former Izu forearc sediments were generally deep-water deposits, formed in several kilometers of water depth (Kitazato, 1997), mostly hemipelagic (Soh et al., 1989, 1991; Stow et al., 1998, 2002), partly contourite (Lee and Ogawa, 1998), and include many intercalations of turbiditic or fall-type volcaniclastic deposits (tuff and lapilli, conglomerate, and sandstone) (Soh et al., 1989). All these sedimentary rocks are composed of Izu Arc volcanic rocks or their derivatives. Such sequences are divided into three prisms from north to south and from older to younger, passing from the Emi Group (Hirono and

Ogawa, 1998), to the Miura Group (Ogawa et al., 1989; Hanamura and Ogawa, 1993; Yamamoto and Kawakami, 2005; Yamamoto et al., 2005; Yamamoto, 2006), and finally to the Chikura Group (part of which may be forearc basin deposits) (Saito, 1992). Such accretionary prisms are highly deformed by dominant thrusts and folds, although most of the original sedimentary structures are preserved (Figs. 2 and 3). The reason for the good preservation is mostly due to the fact that the deformation occurred under semilithified conditions, which involved the most significant flow or shear along bedding planes and the concentration of strain in faulted or folded regions (Ogawa, 1982; Yamamoto et al., 2005; Yamamoto, 2006). As a result, the sediments between the faults and the deformed sections preserve

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140 E Figure 4. Tectonic map of the Sagami Trough. MP—Miura Peninsula; BP—Boso Peninsula; SB—Sagami Basin; MSTB—Middle Sagami Trough Basin; OBC—Okinoyama Banka Chain; SOT—So-oh Trough; BC—Boso Canyon; NB—North Basin; MF—Mogi Fan at the triple junction. Triangle shows the relative plate motion at present between NAM (North American plate), PHS (Philippine Sea plate), and PAC (Pacific plate). Relative motion of the Eurasian plate took place in the early Quaternary. After Seno and Maruyama (1984). Line A is the seismic profile shown in Figure 5, and the other lines are those of the KAIKO project, some of which are shown in Figure 6.

their original structures. Therefore, along the coast, exposures on the uplifted benches permit clear observation of sedimentary structures, including grading, lamination, bioturbation, and other features. Ogawa and his students have published descriptions of the terrestrial geology along the Miura-Boso Peninsula coasts (Hanamura and Ogawa, 1993; Yamamoto et al., 2000, 2005; Yamamoto and Kawakami, 2005; Yamamoto, 2006). The land geology is summarized as follows: 1. Thick forearc basin sedimentary rocks have accumulated since the early Miocene over the Mineoka Belt (in the north) and in the middle Boso Peninsula, whereas farther south several accretionary prisms have developed between the Izu forearc and Honshu forearc. The latter sediments were originally deposited in the deep sea, partly in the Izu forearc and partly along the boundary between the Izu and Honshu Arcs, in settings similar to the present Sagami Trough (Figs. 2–4).

2. Sedimentary rocks in the accretionary prisms are dominated by volcaniclastic sedimentary rocks (tuff, lapilli, and their derivatives) together with hemipelagic, muddy sedimentary rocks or contourite background deposits. 3. Sedimentary rocks in the accretionary prisms are deformed by many thrust faults but still preserve original sedimentary textures and structures because of their rapid uplift after accretion, although there has locally been much liquefaction, injection, slumping, and faulting. 4. Rotations of the accretionary prisms around a vertical axis are strong, and mostly on the NW side of the Miura-Boso Peninsulas and reaching on average 30°–50° clockwise rotation. Locally, however, rotation has been as much as 80°, owing to the collision between the Izu and Honshu Arcs (Kanamatsu et al., 1996). Collision has occurred (6–3 Ma) in the southern Boso Peninsula between the accretionary prism and the forearc or the overlying trench slope sedimentation (Yamamoto and Kawakami, 2005).

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TOPOGRAPHIC AND GEOLOGIC SUMMARY OF THE SAGAMI TROUGH The Sagami Trough is the present day plate boundary between the North American plate to the NE and the Philippine Sea plate to the SW. The latter is now subducting obliquely under the former at a rate of 2.5 cm/yr. The SE end of the Sagami Trough is coincident with the Boso triple junction on the IzuOgasawara (Bonin) trench floor (Fig. 4). The topography of the Sagami Trough is divided approximately into three segments, from the NW in the Sagami Bay area (Sagami Basin), to the middle Sagami Trough area (Middle Sagami Trough Basin), and finally the Boso triple junction area in the SE (Fig. 4). Each area has characteristic topography and internal structure (Nakamura et al., 1984, 1987; Pautot et al., 1987). Seismic profiling has been acquired, but direct observations by submersible are limited, and no drilling has yet been carried out. Sagami Bay Area The Sagami Bay area is occupied by a wide basin, the Sagami Basin, which is filled with a thick wedge of trench-fill sediments in the Sagami Trough (Kato et al., 1983; Kong et al., 1984; Ohkouchi, 1990) (Fig. 5). Pleistocene to Holocene sediments in the basin are not highly deformed but are thrust-bounded to the NE. A fold and thrust belt has developed on the NE side, where Miocene sedimentary rocks are deformed. Such Neogene

formations, ranging in age from early Miocene to Pleistocene, are overlain by multiple unconformities. Terrestrial geology and stratigraphy show that NW-trending folds and thrusts are more strongly deformed in comparison with the older formations (Kimura, 1976; Ogawa, 1982; Ohkouchi, 1990). Systematic younging of these formations toward the SW is recognized in the Miura and Boso Peninsulas, where accretionary prism development toward the SW is well documented (Ogawa and Taniguchi, 1988; Ogawa et al., 1989). The accreted sedimentary rocks either were originally deposited on the flanks of the Izu Arc or in the trench (Fig. 2) (Hanamura and Ogawa, 1993). The only structural difference in the Sagami Bay area is the formation of a dissecting fault pattern subperpendicular to the dominant NW lineament, which we suggest to be a systematic Riedel shear system, as described in the section “Riedel Shear Experiments.” The northeastern part of the Sagami Bay features the Okinoyama Bank Belt, trending from NW to SE as a trench slope break. A nearly parallel, slightly oblique en echelon pattern of NW-SE faults and subperpendicular NE-SW faults cut the entire trench slope area in the NE part of the basin (Fig. 6). The Okinoyama Bank Belt has steep cliffs along its SW boundary, and many Calyptogena (representative chemosynthetic clam) colonies have been known at the foot of these cliffs (Kanie, 1999). These en echelon faults are thought to be active and coincident with the Sagami Tectonic Line (Kimura, 1976) (Fig. 5), corresponding to the deformation front or splay fault of the Philippine Sea plate subduction (Kato et al., 1983; Sato et al., 2005).

Figure 5. Seismic profile of the Sagami Basin and its interpretation along line A in Figure 4, showing a large amount of trench-wedge movement in the Sagami Trough (Sagami Basin). The Sagami Tectonic Line (S.T.L.) SW of the Okinoyama Bank Chain (O.B.C.) is an offscraping or splay thrust. Adapted from Kato et al. (1983). A—Holocene sediments; B—Pleistocene; C—Pliocene; D—middle to upper Miocene; E—lower Miocene.

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Figure 6. Topographic features of the Sagami Basin and a new tectonic interpretation as the Riedel shears of each fault system (after Takami, 1999). Line A is the seismic profile shown in Figure 5. OIS—Oiso Spur; SMK—Sagami Knoll; MZK—Manazuru Knoll; MUK—Miura Knoll; MSK—Misaki Knoll; OYB—Okinoyama Bank; TKC—Tokyo Canyon; HIIs—Hatsushima Island; KMF—Kozu-Matsuda Fault; AGC—Ashigara Canyon.

Nakamura et al. (1984) attempted to explain the en echelon pattern with subperpendicular faults by proposing a special type of plate boundary. These workers proposed that “eduction” is occurring in this region, i.e., that the underlying plate (the Philippine Sea plate) does not subduct but is instead pulled away from the overriding plate (North American plate) with an extensional sense of motion. However, Kato et al. (1983) and Ohkouchi (1990) interpreted seismic profiles to indicate SW-vergent thrust faults in the area of the Okinoyama Bank Belt, forming an accretionary prism (Fig. 5). Ohkouchi (1990) suggests that all of Sagami Bay is characterized by a continuous anticline, based on seismic profile interpretation, and concludes that horizontal compression has occurred around this area. In addition, all the mechanistic solutions for earthquakes and tsunamis, and the resultant crustal movement from the

Taisho-Kanto earthquake in 1923 (M = 7.9), indicate that this boundary is not undergoing eduction but subduction (Ando, 1974; Aita, 1993; Sato et al., 2005). If so, this prompts the question of why dissection of the topography occurs in the Sagami Bay area and whether this is different from other areas. The 2D effects of oblique subduction can be analyzed by a simple sandbox experiment described in the section “Riedel Shear Experiments.” Middle Sagami Trough Area The middle Sagami Trough area is situated south off the shore of the Boso Peninsula as a NW-trending longitudinal basin (Middle Sagami Trough Basin in Fig. 4), demarcated by two faults, coincident with the NE Boso Escarpment and the SW Awa Canyon (Figs. 4 and 7). The interpretation of topography and

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seismic profiles obtained during the KAIKO project indicates that the convex topography of the basin-fill sediments between the two faults represents a modern accretionary prism within an oblique subduction setting (Tanahashi, 1986; Ogawa et al., 1989) (Figs. 7 and 8). This part of the present accretionary prism is colliding along its NW side with the Boso Peninsula, where a sharp thrust-fault scarp has developed in the So-oh Trough (Fig. 8). The Genroku-Kanto earthquake in 1703 (M = 8.2) occurred beneath this scarp, just off the Boso Peninsula (Shishikura, 2003). Many thrust faults in en echelon fashion were developed in the southern fault zone, whereas almost straight faults with dominant strikeslip motion were developed along the northern boundary of the basin (Fig. 8). A forearc basin lies on the NE side of the accretionary complex. The Sagami Trough canyon flows in a meandering fashion at the foot of the fault scarp (Boso Escarpment) along the northeastern boundary (Boso Canyon) (Figs. 4 and 8). The sediments in the basin were accreted from Izu forearc basin sediments. Some are slide deposits from the increase in the slope angle, causing slope instability, as the Izu Arc is colliding with the island of Honshu (Tokuyama et al., 1988). Within the

basin, oblique faults with superficial fractures are being systematically developed (Fig. 8). This set of fractures may be interpreted as R1 shears, as discussed in the section “Riedel Shear Experiments.” The 3D effect of oblique subduction can be successfully analyzed by a simple sandbox experiment. Boso Triple Junction Area The Boso triple junction area has a unique setting both in topography and internal structure (Ogawa et al., 1989; Seno et al., 1989; Iwabuchi et al., 1990). This area is not only tectonically but also topographically unstable. In a large sense the triple junction area is a large, deep triangular basin (Figs. 4 and 9), divided into three parts, one on the NW side, where a wide basin is up to 7400 m deep (North Basin in Fig. 4); another is the trench floor at 9400 m water depth, and the third is surrounded by steep slopes. The first basin has some deep sea terraces along which mud volcanoes and dissecting channels have developed (Ogawa et al., 1989). Canyons connecting to land on the northern side of the wide basin are channeled toward this basin but show significant

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Figure 7. Seismic profiles of the middle Sagami Trough Basin, showing the Izu forearc sediments (right) accreted to the left in the Sagami Trough (SGT). BOC—Boso Canyon; AWC—Awa Canyon; MYC—Miyake Canyon; twt—two-way traveltime. The Miyake Canyon is within the Izu Arc, and the Boso and Awa Canyons are in the Sagami Trough. Adapted from Ogawa et al. (1989). The bottom profile is from the KAIKO I Research Group (1986).

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meandering (Fig. 4). The canyons appear to be deformed and dragged by the downwarping of the basin basement (Soh et al., 1990), as they are not directly oriented toward the deepest part but show clockwise rotation. We suggest that older canyons were rotated by such downwarping (Fig. 9). The main channel in the negative basin flows out through a narrow gorge into the Izu-Ogasawara (Bonin) trench floor at a depth of 9400 m. At this point it forms the deep-sea Mogi Fan (Nakamura et al., 1987; Soh et al., 1988; Ogawa et al., 1989; Seno et al., 1989) (Figs. 4 and 9). The fan deposits were thrust and folded locally prior to accretion to the Izu forearc (Geological Survey of Japan, 1982; Soh et al., 1988; Ogawa et al., 1989). Huge deep-sea submarine slide bodies are recognized (Fig. 10). Surface topographies have many unstable slopes, shown by largescale concave and convex cliffs and scarps by submarine sliding. This was observed in the form of rough ridges and slides during unmanned submersible dives (KR99–11 Shipboard Scientific Party, 2000). In total the triple junction area is a wide, 100-km-scale triangular area of very unstable topography that is still in a state of active collapse. Because the present relative motion of the

surrounding three plate (Fig. 4) indicates that the present triple junction is unstable, the triple point is moving to the NW, changing the system of the junction. The topography (Figs. 9 and 10) indicates such instability, resulting in the unique sedimentation, accretion, and resultant erosion by gravitational instability. RIEDEL SHEAR EXPERIMENTS In order to better understand the tectonics around the Sagami Trough, we performed two types of simple sandbox tests: one in 2D to model Riedel shears caused by horizontal displacement, and another in 3D to simulate oblique subduction. 2D Riedel Shear-Box Test For 2D tests, Bartlett et al. (1981), Naylor et al. (1986), and many others have run simple shear-box tests and recognized systematic development of Riedel shears. Unfortunately, these experiments had no free zone on the displacement basement. We prepared a new experiment, using a rubber sheet, which was dislocated along a displacement basement, as shown in

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Figure 9. Bathymetric map showing the topography in and around the Boso triple junction, courtesy of JAMSTEC (JAMSTEC R/V Mirai and other data edited by M. Nakanishi, 2006, personal commun.). Strong lineaments from NW to SE from the Boso Peninsula side (B) continue to the middle part of the trench slope, gradually becoming obscure in the triangle area near the triple junction. Notice a trench floor submarine fan in the center, which is deformed by frontal thrust. M—Miura Peninsula; B—Boso Peninsula; I—Izu Peninsula.

Figure 11 (left), relative to the superjacent powder. The differences and characteristics of our experiment in comparison with conventional methods are as follows: (1) A rubber sheet is used as a cover for the dislocating board, and the center of the rubber sheet, 5–15 cm wide, is mobile (freely sheared) while the margins are settled by resin. (2) Quartz powder sediment, with a grain size averaging 8 µm in diameter, is set on this rubber in a layer varying from 1 to 4 cm thick. (3) The board is dislocated in a right-lateral sense (dextrally) by a continuously moving motor at a constant speed of 2 mm/s. The strain generated by this stress appears as en echelon folds. The strain of the rubber sheet in the en echelon folds is transferred to the overlying sediment. (4) The strain in the form of fractures on the powder surface during dislocation is recorded by camera. Because the Taisho-Kanto earthquake occurred on a right-lateral (dextral) oblique-slip fault in the Sagami Basin area, our experimental box was designed to efficiently mimic this type of horizontal deformation effect with minor vertical motions. Our experiment revealed several different results, described as follows. First, R1 (synthetic Riedel shear) with right-lateral

displacement occurred, which was followed by R2 (antithetic) shear with a left-lateral sense. Finally, P-shear (thrust shear) with right-lateral displacement occurred (Fig. 11, right). Theoretically R1 and R2 are associated as conjugate shear planes, but in fact the former is more strongly developed, and in general R2 rarely or only seldom occurs. The P-plane is not a theoretically resultant fracture but experimentally occurs as a result of the accumulation of material to produce horizontally compressive pressure into a thrust-component–bearing shear plane (Naylor et al., 1986). This stress-strain relation occurs instantaneously under simple shear. The en echelon folds in the rubber sheet produce a stress field in the overlying sediment that is different in each domain of the folded rubber. As a result, the stress field makes a conjugate set of shears, R1 and R2. The P-shears occur at the left-overstep part of R1, or around the R2 shear zones. This is explained as the result of compressional stress around these areas that gives the shears a thrust component. P-shear occurs slightly after R1 and R2. This is because the compressional stress increases according

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Figure 10. Three-dimensional (3D) topography at the triple junction (left), showing a 10-km-scale large submarine slide body that has collapsed onto the trench floor at 9000 m depth (KR99–11 Shipboard Scientific Party, 2000). The dive site of the remotely operated vehicle (ROV) Kaiko is shown by an arrow, with drawings (right) by Y. Ogawa from video views from the ROV of the collapsed slope. Sediments and rocks of the slope are composed of inclined strata of Pliocene age with minor blocks of middle Miocene age. V.E.—vertical exaggeration. Depths below sea level shown in meters.

to the displacement of the basement. The vertical displacement of R1 and R2 is also due to the compressional stress driven by the left-lateral (sinistral) overstep of R1. Sediment moved in the direction of both R1 and R2, after which sediment accumulated within the wedge between the two. This uplifts the sediment, and small-scale pressure ridges form. As displacement increases, R1 and R2 change their form into sigmoidal geometries because R1 and R2 are connected. The 85-km-long, 45-km-wide fault plane for the TaishoKanto earthquake fault in the Sagami Basin area (Ando, 1974; Aita, 1993) suggests that many subordinate fault planes are responsible for strain. In fact, at the time of that earthquake, the Sagami Tectonic Line corresponded to the main fault. In localities far from the epicenter at the southernmost tip of the Boso Peninsula and south of the Miura Peninsula, motion was recorded on exposed active faults. These facts indicate that plate motion has accommodated not only one fault plane but associated fault systems as well. Furthermore, two asperity parts have recently been discovered for the Taisho-Kanto earthquake (Sato et al., 2005) that reveal the diversity of the slip plane. This realization

indicates that the fault had some irregular concentrations of stress and strain. Thus the main fault of the Taisho-Kanto earthquake had diverse subfaults. Independent preexisting faults or newly formed faults both moved simultaneously. The subducted-landward, overriding plate is thought to have deformed by ductile processes outside the asperity region. This model leads us to think that the ductile rubber sheet, overlain by quartz powder, effectively mimics the system, with local folds being similar to the asperity regions. Comparison with the Topography of Sagami Bay In relation to plate rheology, Shimamoto (1989) showed that in a thin plate, as in an island arc, where dislocation is dispersed over a wide area, the plate boundary comprises a wide zone, and the island arc is deformed internally. In our study region the Sagami Basin area corresponds to the Izu Arc collision zone, and the northern part acts as the hanging wall of the subduction zone in which the North American plate undergoes complicated deformation and fracturing (dislocation) (Lallemant et al., 1996). In

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Figure 11. 2D Riedel shear box-test equipment with a 10-cm-wide free zone of rubber with right-lateral displacement (left), and one of the representative results (right), where R1, R2, and P show synthetic Riedel shear, antithetic Riedel shear, and thrust shear, respectively. After Takami (1999). D.T.—Dislocating Table; M.R.S.—Movable Rubber Sheet; U.M.R.S.—Unmovable Rubber Sheet.

addition, some convex topography reflects asperities in strain or stress concentrations (Sato et al., 2005). If such factors support these facts, then Riedel shears (R1 and R2) must develop as conjugate shear fractures in the superficial sediment layers above the hanging wall. We now compare examples of topography from the Sagami Bay area with our experimental results. Based on the experiments, the overriding plate was predicted to deform by ductile shear, resulting in conjugate Riedel shears, which we discuss below. The relative plate-movement direction along the Sagami Trough indicates that the Sagami Tectonic Line has oblique dextral dislocation (Fig. 4). If we assume that the strike of the Taisho-Kanto earthquake fault was N59°W (Aita, 1993), and this is considered to be the Y-shear (main shear) direction, then the Sagami Tectonic Line must be comparable to R1.

Our experiments indicate that R2 should also occur, and the direction corresponds well to the actual submarine topography. In order to corroborate the possibility of these fault systems, we plotted rose diagrams to graphically represent the directions of canyons and faults on the landward slope of Sagami Bay, as well as all the fractures in our experiments with <3 cm dislocation with all rubber widths (Fig. 12). It is clear that the directions in the two diagrams are similar. The internal angle between R1 and R2 in the experiments is ~10° smaller than the actual topography, but this may be due to the rotation of topography, as shown in experiments discussed below. It was not easy to infer the horizontal displacement sense of the topography just from bathymetric maps, but the faults around the Okinoyama Bank Chain are straightly aligned, suggesting

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a tectonic origin owing to high-angle, horizontal-displacement faults (Fig. 6). Seismic profiles indicate that the western boundary of the Miura Knoll (part of the Okinoyama Bank Chain) has a very steep dip (Fig. 5), suggesting horizontal displacement on a fault. As for the vertical displacement of the Sagami Tectonic Line, the evidence of the Taisho-Kanto earthquake shows that the Okinoyama Bank Chain was uplifted by a thrust component. The experiments are associated with small displacement, but the actual topography indicates uniform positive relief only on the eastern side of the Sagami Tectonic Line. The difference between the model and the observation may be due to our neglecting the vertical displacement of the Taisho-Kanto earthquake. The effect where the basement plate has a thrust-component–bearing oblique fault that affects the overlying sediment was modeled in 3D by Naylor et al. (1986)’s sandbox experiments (Mandl, 1988). This result shows that the basement dislocation may undergo R1 strain, with a thrust-faulting effect within an en echelon arrangement. Naylor et al.’s experiment predicts uplift along the eastern side of the Sagami Tectonic Line as a “flower structure.” The comparison of experiment and topography suggests that, although some of the submarine fault displacement is not known in terms of the sense of motion, the present topography of the Sagami Bay correlates well with the experiment in terms of the dominant fault and fracture directions. In addition, our study indicates that the submarine topography is mostly influenced by a regional dextral shear field. 3D Oblique Subduction Test As 3D experimental tests are rare, and oblique subduction models in particular had not previously been performed, we made a simple but effective test using a steel frame and stiff paper board (Takazawa, 2003). Wheat powder with colored-chalk layers and chocolate-powder mesh on the surface was prepared over the subduction boundary in order to mimic the trench wedge sediments. The inclined board (sloping 30° downward) was moved steadily by hand at variable degrees of obliqueness. Fractures on the surface and faults and folds inside were shown by cutting the model with a knife after the movement. For each different degree of obliqueness (0°, parallel to the trench, to 90° in 10° steps) the surface fracture development and internal structures were recorded by photograph. Representative results are shown in Figure 13. Results of 3D Tests Figure 13 shows the most representative case, in which the obliqueness angle was 30°. Surface fractures developed systematically. First, Riedel R1 shears formed, then locally R2 shears were seen, but these were relatively rare. Another direction of shear, interpreted as thrust shears (P-shears) formed at a later stage. Internally within the wedge, thrusts and their associated folds systematically developed (Fig. 13, left). In some cases the internal thrusts developed during the initial stage, which is similar to the accretionary prism model, because the thrust on

Figure 12. Rose diagrams of Riedel shear fractures of sandbox test with right-lateral displacement (above; see Fig. 11) and topographic lineament of the Sagami Basin (below; see Fig. 6). Each dominant direction indicates synthetic Riedel shears and antithetic Riedel shears of right-lateral displacement, respectively. After Takami (1999).

the ocean side formed at a later stage. In most cases, one or two back thrusts developed as the deformation proceeded, as part of a flower structure. This simple experiment clearly indicates that in cases of greater obliqueness (especially <40°), stronger Riedel shears develop, but the effect of thrust and fold development is more or less the same as in the case of smaller obliqueness (close to

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Figure 13. Results of 3D oblique subduction tests of right-lateral sense with 30° obliquity. Profile (left), and top surface fractures (right). Notice components of fractures on the surface of the oblique subduction; synthetic Riedel shears and thrust shears are distinct. Compare the experiment with the real examples from the Middle Sagami Trough Basin in Figure 8. After Takazawa (2003).

normal convergence). However, in the less oblique experiment, with convergence at more than 70°, no strong P-shears formed. The best fit to the Middle Sagami Trough Basin area in terms of the surface fracture system is for 20°–30° of obliqueness (Fig. 13). This corresponds to the natural case in the area, where measured obliqueness is ~30° (Figs. 4 and 8). In the So-oh Trough, between the Middle Sagami Trough area and the Sagami Bay area (NW corner in Fig. 8), just in front of the Boso Peninsula, obliqueness is almost normal, and no Riedel shears are observed on modern bathymetric maps. Therefore, we conclude that even a simple 3D oblique subduction model is successful in mimicking the natural features of the seafloor (Takazawa, 2003). CONCLUSIONS In this study we verified the origin of the fault system and related topography of the Sagami Trough and used experimental modeling to understand the development of Riedel shears in oblique subduction settings within an island arc collision zone in a TTT-type triple junction. In experiments, well-defined shears associated with horizontal displacement appeared both in 2D and 3D tests. The orientation of faults along the Sagami Trough, both in the northwestern part (Sagami Bay area) and in the middle part (Middle Sagami Trough Basin area) can be explained as the dominant two directions oblique to the main fault, which correlate well with R1 and R2 shears in sandbox experiments. The Sagami Tectonic Line can be understood as a plate-boundary thrust fault, where an oblique fault has a horizontal displacement with a thrust component. These results indicate that the fault system and topography can be attributed to Riedel shears, which generate wide regional

surface deformation driven by a dextral shear zone that is triggered by highly oblique subduction. More detailed constraints are needed to verify a direct link between the topography, geology, and tectonics. We suggest that direct observation and sampling at the seabed, supplemented by drilling and dredging, are needed to understand the triple junction structure in three dimensions. Only by doing this can our proposed model be rigorously tested. The different appearance of shear patterns in the Sagami Bay area and the Middle Sagami Basin area may be attributed to the different relationships between the basement and the overlying sediment cover. In the Sagami Basin a wide and thick accretionary wedge above an irregular basement (modeled with folds of the rubber sheet) reflects strain-asperity zones and is deformed by strain partitioning. As a result, both R1 and R2 shears are formed continuously. In contrast, in the case of the Middle Sagami Trough Basin the strain is transferred in a more regular fashion to the sediments from the subduction, forming a regular shear pattern. Further quantitative analysis by more sophisticated methods is needed to more fully understand this complicated region, but the simple experiments described here offer a plausible solution. ACKNOWLEDGMENTS We are grateful to Peter Clift and Amy Draut for inviting us to the 2006 Penrose Conference in Price, Utah. Peter Clift, Roland von Huene, and one anonymous reviewer kindly read the early version of the manuscript, and gave us critical suggestions for revision. Some bathymetric charts and figures were prepared through the courtesy of JAMSTEC and Masao Nakanishi.

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Printed in the USA

The Geological Society of America Special Paper 436 2008

The West Crocker formation of northwest Borneo: A Paleogene accretionary prism Joseph J. Lambiase* Tan Yaw Tzong* Amelia G. William* Department of Petroleum Geoscience, Universiti Brunei Darussalam, Tungku, Brunei BE 1410 Michael D. Bidgood Patrice Brenac GSS International, Aberdeenshire, UK Andrew B. Cullen Sabah Shell Sendirian Berhad, Miri, Malaysia

ABSTRACT An integrated structural, stratigraphic, and sedimentological analysis of the West Crocker formation in northwest Borneo suggests that it is best interpreted as an accretionary prism. The structural geology provides clear evidence of at least two episodes of syndepositional folding and thrust faulting. A probable Eocene age, indicated by foraminiferal and palynological assemblages, differs from the generally accepted Oligocene to early Miocene age and is consistent with deposition of the West Crocker formation during a phase of tectonism at the northwest Borneo margin. Sandstones within the West Crocker formation were deposited by high-density turbidity currents that constructed relatively small, progradational lobes in a slope apron environment, and trace fossil assemblages confirm bathyal water depths of ~1000 m or more. The composition of the sandstones, which contain abundant feldspars and lithic fragments, suggests that their provenance was the first-cycle product of an eroded orogenic belt, whereas immature textures indicate a short distance of transport. Keywords: turbidites, syntectonic, accretionary prism, Borneo.

INTRODUCTION The West Crocker formation is an informal name for distinguishing the western and northwestern outcrops of a Paleogene succession in northwest Borneo that constitutes the sand-rich Crocker Formation and its coeval shaly equivalent, the Temburong

Formation (e.g., Hutchison et al., 2000; Fig. 1A). The CrockerTemburong Formation has long been recognized as a thick succession of turbidites with northward-directed paleocurrents that was deposited in response to collisional tectonics associated with Late Cretaceous to early Miocene rifting in the South China Sea (Stauffer, 1967; Wilson, 1964; Tan and Lamy, 1990; Hutchison,

*Present address, Lambiase: Lambiase Geoscience Pte. Ltd., Singapore; e-mail: [email protected]. Present address, Tzong: Sabah Shell Sendirian Berhad, Miri, Malaysia. Present address, William: ExxonMobil Exploration and Production Malaysia Inc., Kuala Lumpur, Malaysia. Lambiase, J.J., Tzong, T.Y., William, A.G., Bidgood, M.D., Brenac, P., and Cullen, A.B., 2008, The West Crocker formation of northwest Borneo: A Paleogene accretionary prism, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 171–184, doi: 10.1130/2008.2436(08). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Figure 1. (A) Location of Borneo within southeast Asia and the distribution of the West Crocker and Temburong formations, and the Rajang Group (undifferentiated Crocker and Trusmadi Formations). (B) Location of the Bukit Melinsung (BM), Jalan Salaiman (JS), and Maju (MJ) outcrops; all are along main roads that are shown as dashed lines.

1996). Various authors have interpreted the unit as one major turbidite fan system derived from a relatively distant southerly source (e.g., Crevello, 2002; Tongkul, 1987). However, Hutchison (2005) notes that it is difficult to envision a single fan system persisting through the entire Paleogene to early Miocene age span of

the Crocker-Temburong Formation, and unconformities are suspected if not proven. Also, the structural style and burial history of the various parts of the formation differ significantly, suggesting that the unit comprises a number of successive fan systems, of which the West Crocker formation is the youngest and least deeply buried (Hutchison, 2005). The West Crocker formation lies directly below the Deep Regional Unconformity of Levell (1987), a profound angular unconformity that was generated during an orogenic event that uplifted Paleogene and older rocks, and marks the transition from that uplift to the rapid subsidence that accompanied deposition of the thick (up to 14 km), overlying Neogene shallow-marine clastic succession (Fig. 2). The Deep Regional Unconformity is generally regarded as being of late early Miocene age (e.g., Sandal, 1996), although it is diachronous, and some workers prefer a middle Miocene age (e.g., Hutchison et al., 2000). In the Luconia and Balingian provinces of Sarawak a regional unconformity of similar age separates cycle II and cycle III (Madon, 1999; Madon and Abolins, 1999) and generally marks the transition from extension to regional subsidence. There has been no direct age dating of the West Crocker formation, because the few microfossils that have been recovered are not age-diagnostic within the Eocene to Miocene (Hutchison, 2005). However, the West Crocker formation is regarded as a sandy facies that is stratigraphically equivalent to the shaly Temburong Formation (Fig. 2). The Temburong Formation is the youngest part of the Crocker Formation-Temburong Formation complex and was dated as early Oligocene to early Miocene by Wilson (1964), using foraminifera. Therefore, the West Crocker formation is presumed to lie directly below the Deep Regional Unconformity that marks the boundary with the overlying Neogene shallow-marine clastic sedimentary rocks (Fig. 2). It is generally agreed that the turbidites were deposited in a foreland basin that developed in response to uplift of the Rajang Group accretionary prism during the Sabah orogeny that began in the late Eocene (ca. 36 Ma; e.g., Tongkul, 1990; Hutchison, 1996), although the paleogeography of northwest Borneo and the exact timing of events during the early Tertiary are poorly constrained. The thickness of the West Crocker succession is uncertain partly because postdepositional thrust faulting makes stratigraphic reconstruction between outcrops nearly impossible; Wilson (1964) suggested that at least 6000 m of section is present, whereas Tongkul (1987) argues for ~1000 m. By the late Oligocene the West Crocker formation was being thrust over an area of buried continental shelf and attenuated continental crust called Dangerous Grounds (Hutchison, 2004; Hutchison et al., 2000). The SW-directed underthrusting of Dangerous Grounds during subduction generated NW-SE compression that was responsible for the complex structures of the Eocene to middle Miocene sedimentary rocks in western Sabah, including the West Crocker strata (Tongkul, 1990, 1991). Apatite fission-track data indicate that the West Crocker sedimentary rocks were buried to 4–8 km and were exhumed during development of the middle Miocene regional unconformity (Hutchison et al., 2000).

West Crocker formation of northwest Borneo

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Figure 2. Early Tertiary stratigraphy of northwest Borneo, with suggested revisions to the scheme of Hutchison et al. (2000).

Present Study This study investigated the sedimentology, stratigraphic architecture, and structural style of the West Crocker formation turbidites in order to develop a tectono-stratigraphic model for their deposition. Three outcrops were selected for detailed study: Bukit Melinsung, Jalan Salaiman, and Maju (Fig. 1B). The first two were chosen because they are much less deformed than most. One of them has the thickest continuously exposed vertical succession, whereas the other represents the most continuous lateral exposure. Approximately 350 m of vertical section is exposed continuously in the Bukit Melinsung outcrop (Fig. 1B), although lateral continuity is limited to ~50 m. About 70% of the succession is very fine to granular sandstone, and the remaining 30% is shale. Individual beds in the Jalan Salaiman outcrop near Kota Kinabalu (Fig. 1B) can be traced laterally for ~500 m. They

consist of very coarse to coarse grained, 1–6-m-thick sandstones interbedded with thinner shales in a 130-m-thick, continuously exposed vertical succession. The Maju outcrop was selected because it covers a large area and has a wide variety of structures in >800 m of exposed section, representing >400 m of stratigraphic succession (Fig. 1B). The succession has been deformed by multiple folding and faulting events that have generated thrust faults, normal faults, and a wide range of fold types. SEDIMENTOLOGY AND STRATIGRAPHY The outcrops were logged to define grain size trends, sedimentary structures, the nature of bed contacts, trace fossil distribution, and stacking patterns. Pollen and foraminifer assemblages were determined from the interbedded shales for the

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interpretation of age and depositional environment, respectively, and the sandstones were analyzed petrographically to evaluate gross composition. Sedimentary Facies The outcrops comprise three facies, which are designated as thick sandstones, thin interbedded sandstones and shale, and shale. Thick Sandstones Thick sandstones generally occur as amalgamated beds, although a few are isolated between thin shales. Amalgamated units generally consist of 4–25 sandstone beds with a total thickness of 6–33 m (Fig. 3A). Within the amalgamated units,

individual beds generally thicken and coarsen upward. Their basal contacts are straight and sharp, and relatively few are scoured. Flute casts, although not common, indicate a northerly paleoflow direction when corrected for structural dip (Fig. 3B). Load structures are moderately common. Within each unit the lower sandstone beds are relatively thin (0.3–1.0 m) and generally grade upward from coarse or very coarse sand to fine sand, laminated silt, and finally shale. Most beds are massive at the base and pass upward into parallel lamination (Fig. 3C), although some have convoluted bedding that passes upward into parallel lamination. Sandstone beds 1–6 m thick, rarely up to 10 m thick, occur near the top of the amalgamated units, although Hutchison (2005) reports that 10-m-thick beds are more common in a few other outcrops. The thick sandstones generally are coarse to very coarse sand, massive, and

Figure 3. (A) Amalgamated sandstones in the Jalan Salaiman outcrop. (B) Flute casts at the base of a bed in the Maju outcrop. (C) Typical West Crocker formation turbidite sandstone, with a scoured base and a sharp upper contact with the overlying shale. The sandstone is massive, with very coarse sand and granules near the base, and grades into parallel laminated, medium sand. The unit comprises Bouma divisions Ta, Tb, Td, and Te. (D) Climbing ripples in the siltstone cap of a turbidite sandstone.

West Crocker formation of northwest Borneo structureless. Grain size distribution is uniform throughout but grades into medium sand a few centimeters from the top of the bed. Parallel lamination and, less commonly, contorted lamination occur locally. The uppermost parts of the massive sand bodies locally have thin laminations of carbonaceous material and are rippled or wavy-bedded. Siltstones and shales that cap the sandstones are usually laminated and contain abundant carbonaceous material; isolated ripple cross-stratification and climbing ripples occur locally in thin sandy interbeds (Fig. 3D). Generally, the sandstones in the amalgamated units have flat, straight, and sharp upper contacts; the few that appear to have gradational upper contacts and scoured basal contacts generally are in the lower part of the unit. Thin Interbedded Sandstones and Shales The interbedded sandstone and shale facies comprises monotonous alternations of 4–8-cm-thick sandstones and shale within a 0.5–1.5-m-thick unit. The sandstone layers generally are fine- to medium-grained sand. However, some beds grade upward from coarse- to medium-grained sand to parallel-laminated sand and silt; some layers have discrete and climbing ripples or parallel lamination, commonly with a high content of carbonaceous material. Most of the sandstones have sharp basal and top contacts and are capped by thin, homogeneous and laminated shale layers, although a few have gradational contacts. Shale Shale units are homogeneous and range from 0.5 to 5.0 m in thickness. The lower parts of the shale units usually contain thin interbeds of fine to silty sand that exhibit straight, sharp contacts and ripple cross-lamination. Trace Fossils and Microfossils Microfossils Microfossil analysis yielded minimal information about the age and depositional setting of the West Crocker formation because the shale samples contain only a few agglutinating foraminifera (1–6 individuals per sample). All are either Haplophragmoides spp. or Recurvoides spp., except for one possible Trochammina sp., that are normally indicative of a dysaerobic, deep-water setting and an undifferentiated Tertiary age. Pollen recovery generally was poor, with many of the palynomorphs consisting of reworked Cretaceous forms or those with an undifferentiated Tertiary age such as Dicolpopollis spp. However, the presence of Inaperturopollenites spp., Liquidambar spp., Tiliaepollenites spp., and several Palmae taxa suggests an early Eocene to early Oligocene age. Trace Fossils A relatively diverse assemblage of trace fossils occurs in the sandstones of the three outcrops, but the number of beds with visible traces is low, and each locality tends to have a low diversity, and some are even monospecific. The traces that have been

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identified are Ophiomorpha, Thalassinoides, Terebellina, Planolites, Paleophycus that are 0.3–0.5 cm in diameter, Asterosoma, and Rhizocorallium. Ophiomorpha range from 1.0 to 2.5 cm in diameter and are up to 1.3 m long. Tajul Anuar Jamaluddin (1989) identified a similar trace fossil assemblage, plus Crossopodia, in other West Crocker formation outcrops. Trace fossils are rare within most of the shales and are mostly Helminthoida and Megagrapton of the deep-water Nereites ichnofacies (Pemberton et al., 1992) as well as Planolites (?) that are 0.3–0.5 cm in diameter. However, the apparent low abundance of trace fossils in the finer grained facies is clearly a product of intense, tropical surface weathering. A recently excavated area of the Jalan Salaiman outcrop revealed several bedding plane surfaces covered with abundant, high relief Megagrapton, Nereites, and Ophimorpha rudis traces that are indicative of bathyal water depths (Fig. 4; Pemberton et al., 1992). When the site was revisited 6 weeks later, the traces were barely visible. Petrography The modal composition of 36 sandstones was analyzed in thin section. They generally contain 40%–60% quartz, 15%–25% feldspar, and 10%–15% clay matrix (Fig. 5A). The feldspar fraction is mostly microcline and orthoclase, with a few occurrences of plagioclase. Lithic rock fragments account for 5%–20% of most samples but reach nearly 30% in a few examples. Thus, the sandstones are mostly subarkoses with a few sublitharenites and litharenites (Fig. 5B; Pettijohn et al., 1987). Twenty of the samples plot within the field that defines a continental block as the source of sediment, whereas the other 16 appear to have been derived from a recycled orogen (Fig. 5B; Dickinson and Suczek, 1979). The overall distribution is consistent with sandstones derived from older sediment that was the first-cycle product of an eroded orogenic belt, in this case almost certainly the Rajang Group. The sandstones are texturally immature, as all samples are poorly sorted and most grains are angular to subangular, although subrounded grains are moderately common. The matrix is made of very fine grains and clay minerals; much of the clay fraction comprises authigenic clays formed during diagenesis or surface weathering. Similarly, most of the feldspars in the framework constituents have been replaced by sericite, kaolinite, muscovite, and a few unidentifiable authigenic clay minerals. The petrographic analysis indicates that the sandstones were originally very immature and contained a considerable amount of feldspar and depositional matrix before diagenesis and surface weathering. Stratigraphy Generally, the three sedimentary facies stack into coarseningupward parasequences that range in thickness from 25 to 65 m. The basal shale in each parasequence is overlain by interbedded sand and shale, followed by an amalgamated sandstone unit at the top (Fig. 6). The sandstone beds within the amalgamated units thicken and coarsen upward, which is a typical progradational

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Figure 5. (A) Thin section of a typical poorly sorted, texturally immature West Crocker formation sandstone with some component grains labeled as quartz (Q), polycrystalline quartz (PQ), microcline (M), and feldspar (F). (B) Gross composition of selected West Crocker formation sandstones, based on the classification scheme of Pettijohn et al. (1987). The provenance fields for recycled orogens and continental blocks are from Dickinson and Suczek (1979).

Figure 4. Trace fossils in the shale beds of the West Crocker formation belong to the bathyal Nereites ichnofacies (Pemberton et al., 1992). They include (A) Megagrapton, (B) Nereites, and (C) Ophiomorpha rudis.

stacking pattern (Ricci-Lucchi and Valmori, 1980; Normark, 1978; Macdonald, 1986; Walker, 1992). Bukit Melinsung is the only outcrop with a vertical succession that is relatively undeformed and thick enough to discern stratigraphic trends among parasequences. There, the lower four parasequences thicken upward, with a corresponding increase in the proportion of sandstone, indicating that they comprise a single progradational parasequence set (Fig. 7). The parasequence set is incomplete, because a relatively large fault forms its upper boundary. The upper four parasequences are much sandier than the lower four. However, they also generally thicken and coarsen upward and make up part of another progradational parasequence set (Fig. 7).

West Crocker formation of northwest Borneo

Figure 6. Facies stacking pattern of a parasequence in the Jalan Salaiman outcrop. Relative grain size is schematic and not to scale.

STRUCTURAL STYLE All of the West Crocker formation outcrops show some degree of structural deformation. Most are cut by at least several, mostly thrust, faults and are folded to some degree but are not intensely deformed. However, the Maju outcrop covers a large area and exposes the temporal relationships among the various structures and thus offers a unique opportunity to analyze the structural development of the West Crocker formation. The outcrop consists of three stratigraphic successions; two are separated by a thrust fault across which the offset cannot be determined because it is more than the exposed length of the fault plane, and the other boundary is stratigraphic (Fig. 8). The oldest of the three successions forms the hanging-wall of the

Figure 7. Generalized sequence stratigraphy of the Bukit Melinsung outcrop, revised from Mohd. Khalid Jamiran (1999). Note that the thicker sandstones are units composed of stacked beds that thicken and coarsen upward within the unit, whereas individual beds fine upward. Several faults and folds within the succession are not shown; a relatively large thrust fault forms the upper boundary of the lower parasequence set. Relative grain size is schematic and not to scale.

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Figure 8. Schematic structural cross-section of the Maju outcrop, showing the major stratigraphic subdivisions. Note that the hanging-wall of the major thrust fault is more highly deformed than the footwall. Locations of the photographs in Figures 9 and 10 are as indicated. After Tan (2005).

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large thrust fault, and the younger two successions are in the footwall. The hanging-wall succession has undergone at least three phases of deformation, of which the older two are compressional and the third extensional, whereas the footwall successions were affected only by the second compressional event and the extensional phase. Hanging-Wall Succession Approximately 70 m of the hanging-wall succession is exposed laterally in near-dip and near-strike surfaces that are a total of 175 m long and range from 8 to 12 m high. The succession includes all three of the sedimentary facies described previously. The first phase of deformation was generated by NE-SW compression that created two thrust faults that verge southwest and a relatively large fold that was later refolded during a subsequent NW-SE compressional event. The refolded fold is a key structure for determining the relative timing of the deformational events and is bounded by faults to the west and south. It is a tight recumbent, refolded fold with two fold axes, the older of which reflects the first, NE-SW compressional event (Fig. 9A). The rock mass in the core of the tight fold comprises thick, disrupted, overturned sandstone beds, and the remainder is breccia. Structures formed during the first compressional event are restricted to the hanging-wall succession, suggesting that the deformation took place before deposition of the footwall successions and that these structures were synsedimentary at the formation time scale. The second, NW-SE phase of compression generated the second fold axis in the refolded fold and a series of high-angle, NW-verging thrust faults that cut across the folded beds, plus a number of other thrust faults and folds. Parasitic folds are commonly associated with the larger folds. Generally, thicker, competent sandstone beds are gently folded, whereas shalerich intervals have well-developed detachment folds, indicating a close relationship between lithology and fold geometry that is pervasive in the outcrop. Angular, 5–30-cm-diameter sandstone clasts within a sheared shale matrix form an incohesive, foliated fault breccia on some of the faults. Part of the hanging-wall succession is characterized by complex folds and sandstone blocks within sheared shale that are interpreted as a slump (Fig. 9B). Several small-scale (1–5 m), isolated folds are within the slump. A few minor normal faults with minimal displacement were generated by the third and final phase of deformation. All the young normal faults across the entire outcrop have small displacements and are viewed as the product of postorogenic gravitational collapse after compression rather than the reflection of a significant, extensional tectonic event (Tan, 2005). Lower Footwall Succession The lower footwall succession is 75 m thick and covers an area 8 to 10 m high and ~300 m laterally on a near-strike exposure

Figure 9. Structural features in the hanging-wall and lower footwall successions of the Maju outcrop. See Figure 8 for locations. (A) Part of the large refolded fold in the hanging-wall. The apparent anticline on the left side of the figure is actually an overturned syncline; the younging direction is toward the bottom of the photo. (B) Highly deformed succession in the hanging-wall that is interpreted as a slump; the younging direction is variable and locally indeterminable. Some of the sandstones are outlined with a white line for easier visualization of the structure. (C) Parasitic folds in a shale bed above gently folded sandstones that form an anticline in the lower footwall succession. The younging direction is toward the top of the photo.

surface. All three sedimentary facies are present within the succession that was deformed by the second, NW-SE–oriented compressional event and later extensional faulting. Structures generated by NW-SE compression dominate the outcrop and especially the footwall successions. Indeed, the entire outcrop can be interpreted as part of the southeast limb of a large, disharmonic anticline with three other folds representing relatively large-scale parasitic folds on the anticline (Tan, 2005).

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The lower footwall succession is a gently folded anticline within a thick sandstone-dominated succession. There are a number of parasitic folds that are restricted to shale-rich intervals up to 7 m thick, and a series of thrust faults that cut the succession (Fig. 9C). The thrust faults are more abundant in the NW limb than in the SE limb; this may indicate that the compressional stress was NW to SE directed, which coincides with the South China Sea spreading direction (Tan and Lamy, 1990). This suggests that at least some of the deformation in the West Crocker formation may be the product of regional tectonics, and further implies that the West Crocker formation may be age-equivalent to the early Oligocene to early Miocene (32–16 Ma) South China Sea spreading (Briais et al., 1993). Some of the thrust faults exhibit drag folding, and others preferentially follow shale-rich bedding planes. The extensional deformation phase is represented by several minor normal faults that cut the succession. Upper Footwall Succession The ~275 m of upper footwall succession strata also includes all three sedimentary facies and is exposed in near-strike and near-dip surfaces that total ~550 m in length and are 8 to 35 m high. The basal unit is a thick sandstone that directly overlies the gently folded anticline of the lower footwall succession. The basal unit also has been folded, but the sandstone beds pinch out against the flank of the underlying fold, creating an onlap geometry with a dip difference of 11° across the onlap surface (Fig. 10A). The onlap angle far exceeds depositional dip in similar turbidite systems (Kumar and Slatt, 1984), which indicates that the underlying fold had developed, or was growing, when the overlying sandstones were deposited, and is clear evidence of synsedimentary deformation. Dips of up to 12° or 13° are common on growing structures in unconsolidated sediments in modern turbidite systems (Morley, 2007), so there is no need to invoke burial and/or consolidation before folding of the underlying strata, or a significant amount of time, between folding and deposition of the onlapping sandstones. The lower part of the upper footwall succession dips consistently at 52°–60° and is cut by numerous thrust faults that were generated by the NW-SE compressional event, some of which form a complex fault zone. Near the top of the succession is a large-scale fold with a NW-SE–striking fold axis in which dips increase upsection to the point at which the beds eventually become overturned. The folding caused a distinct thickness change in a shale unit that is detached from an overlying sandstone bed in which there is no thickness change; this demonstrates the competency of the sandstones relative to the shales (Fig. 10B). There are also associated small-scale detachment folds within an overturned thin interbedded sandstone and shale unit (Fig. 10C). The fold is cut by several thrust faults, some of which are east verging and some of which have associated breccia. A number of normal faults that represent the final, extensional phase of deformation cut the succession and the older structures.

Figure 10. Structural features in the upper footwall succession of the Maju outcrop. See Figure 8 for locations. (A) Onlap relationship between the upper and lower footwall successions. A white line highlights the onlap surface. (B) Abrupt lateral change in shale thickness adjacent to a fold detachment, as indicated by the white arrows. The younging direction is to the right of the photograph. (C) Small-scale detachment fold in the overturned beds associated with a larger scale anticline. The arrow indicates the younging direction.

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TECTONO-STRATIGRAPHIC MODEL

Age and Tectonic Setting

Depositional Model

Age of the West Crocker Formation The age of the West Crocker formation is poorly constrained because few age-diagnostic microfossils have been recovered. Many workers assign an Oligocene to early Miocene age to the formation (e.g., Hutchison et al., 2000; Hutchison, 2004; Tongkul, 1990, 1991), although some acknowledge that it may be as old as Eocene (e.g., Rangin et al., 1990; van Hattum et al., 2003). Pollen data from this study suggest that the three outcrops are Eocene deposits, although the age of the entire West Crocker formation could be variable. A preliminary palynological analysis, using identical analytical techniques, also indicates an Ecoene age for the unconformably overlying Melingan Formation in Sabah (Fig. 2). It is important to note that in Sabah the Melingan Formation consists of steeply dipping, high-energy storm beds that are markedly different from the gently dipping, low-energy, wave-dominated sandstones with minor limestones in Sarawak and Brunei that were dated as early Miocene by Wilson (1964). This suggests that either the Melingan Formation has a broader age range than previously recognized (Fig. 2) or that the Sabah and SarawakBrunei outcrops belong to different stratigraphic units. Assuming that the stratigraphic relationship between the West Crocker and Melingan formations as defined by previous workers (e.g., Liechti et al., 1960; Wilson, 1964; Tate, 1974) and illustrated in Figure 2 is accurate, then the West Crocker formation probably is Eocene to Oligocene rather than Oligocene to lower Miocene, and the stratigraphy should be revised as suggested in Figure 2. This view indicates that most of the West Crocker formation is not a sandy facies equivalent of the shaly Temburong Formation, which resolves the problematic paleogeographic facies distribution whereby the sandy fans of the West Crocker formation lie N-NE of the Temburong Formation shales in a more distal position.

The sedimentary structures found in the West Crocker formation indicate that the sandstones were deposited by turbidity currents, as noted in previous studies (e.g., Stauffer, 1967; Tongkul, 1990). Most beds contain partial Bouma sequences with divisions Ta, Tb, and Td; Ta and Td; Ta and Tc; Ta and Te; or just Ta. These Ta–Tc dominant, massive, structureless but graded amalgamated sandstone beds with sharp and straight upper and lower contacts are typical of high-density turbidites (e.g., Stow et al., 1996; Lowe, 1982; Walker, 1992; Stow and Johansson, 2000). The generally structureless and massive turbidite sandstones indicate rapid deposition, probably by collapse fallout from high-density turbidity currents (Stow and Johansson, 2000) caused by an abrupt decrease in slope angle along the transport path. Parallel lamination and climbing ripples near the top of the sandstones and in the shale caps are the consequence of rapid and continuous fallout of finer grains during the final stages of deposition. The sharp bed contacts are the product of sediment bypassing and sequential sedimentation of different particle populations from turbulent high-density flows transporting very coarse to fine grains in suspension (Lowe, 1982; Stow et al., 1996). The very immature texture, angularity, and poor sorting of the sandstones suggest a nearby sediment source and a short transport distance, as does the presence of angular to subangular feldspars that are easily destroyed by extensive transport (Carozzi, 1993; Harwood, 1988). Many of the trace fossils belong to the Cruziana ichnofacies that is indicative of a shelfal environment with water depths normally <200 m (Pemberton et al., 1992; Frey and Pemberton, 1984). However, those in the shale, Ophiomorpha rudis, Nereites, Helminthoida, and Megagrapton, are deep-water traces of the Nereites ichnofacies (Pemberton et al., 1992). In such cases, the assemblage in the shales is the most reliable water depth indicator, confirming that the West Crocker turbidites probably were deposited in >1000 m of water (G. Pemberton, 2004, personal commun.). All the sedimentary characteristics are most consistent with deposition on a slope apron (Macdonald, 1993; Reading and Richards, 1994). Sands were deposited rapidly at the toe of the slope, where they formed small-scale lobes; the relatively thin sandstone beds (nearly all are <6 m, and most are <2 m thick) suggest that the depositional lobes were not large but had radii in the range of 1–5 km or less (Stow and Johansson, 2000). This results in a surface slope on the lobes of <1° even for the thickest parasequence of 65 m, which corresponds well with other sand-rich slope aprons, e.g., the middle Tonkawa Sandstone of the Anadarko Basin in the U.S. Midcontinent (Kumar and Slatt, 1984). The coarsening and thickening upward nature of most of the parasequences suggests that the lobes were progradational.

Tectonic Setting Several aspects of the structural evolution of the Maju outcrop indicate that at least part of the West Crocker formation underwent synsedimentary deformation, which differs from the traditional view that all the deformation is postdepositional (e.g., Tongkul, 1990, 1991; Hutchison, 1996, 2005). The foremost aspect is that two different phases of deformation affected the hanging-wall succession, whereas only one of them deformed the footwall. The onlap relationship between the lower and upper footwall successions (Fig. 8) not only confirms synsedimentary deformation but suggests that compressional tectonics was ongoing while most, if not all, of the units were deposited. Therefore, the compressional deformation of the deep-water turbidites is older than the uplift of the West Crocker formation. It is generally accepted that early Tertiary compression was related to subduction along the northwest Borneo margin, coupled with seafloor spreading related to the opening of the

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South China Sea (e.g., Tongkul, 1990; Hutchison, 2005). In Tongkul’s (1990) model, NW-SE compression was generated by SE-directed underthrusting of the Dangerous Grounds during late Oligocene (ca. 25 Ma) subduction. Both phases of compressional deformation must have been completed by the time that subduction ceased; many workers think that this occurred in the middle Oligocene (ca. 30 Ma; e.g., Tongkul, 1990), although some believe that it persisted to the early middle Miocene (Hutchison, 1996). Structures generated by NW-SE compression are abundant throughout the West Cocker formation (Stauffer, 1967; Tongkul, 1990, 1991; Hutchison, 1996, 2005), but the older, NE-SW compressional phase has not been documented previously. The compressional tectonic setting, synsedimentary deformation, and slope apron depositional model suggest that the West Crocker formation is an accretionary prism. Sediments were supplied episodically from uplifted Rajang Group strata in the active orogen and transported down a relatively steep slope to the adjacent basin. They accumulated at the base of the slope in a series of small fans that were deformed by at least two phases of syndeformational, compressional folding and thrust faulting. DISCUSSION The probable Eocene accretionary prism interpretation for the West Crocker formation not only provides the best reconciliation of the structural data and sedimentological data collected in the present study, but it overcomes many of the problems and inconsistencies of previous interpretations. One of the important sedimentological issues is the size and type of fan system that the West Crocker formation represents. Many workers interpret the West Crocker formation as a large submarine fan that was fed from a point source to the SW, relatively far from the West Crocker depocenter (e.g., Stauffer, 1967; Tongkul, 1987; Hutchison, 1996). This argument is partially based on paleogeographic reconstructions that suggest that marine waters covered present-day Sabah prior to the uplift that produced the Deep Regional Unconformity (e.g., Hall, 2002), thereby precluding a sediment source to the east. However, the thin sandstone beds and small lobes of the West Crocker formation are too small for a fan associated with a relatively large river (Reading and Richards, 1994). Another important issue is the provenance area for the West Crocker formation sediment. Van Hattum et al. (2003) concluded that the textural immaturity of the West Crocker formation reflects a sediment source from nearby, first-cycle orogenic sandstones, possibly the uplifted Rajang Group in Sarawak. These precursor sandstones probably were the erosional products of granites in the Schwaner Mountains in southern Borneo (Fig. 1A). Van Hattum et al. (2003) also attribute the subordinate, ophiolite-derived material in the West Crocker formation (Stauffer, 1967) to a secondary provenance area to the east. The samples analyzed in the present study contain considerably more feldspar (15%–25% as opposed to 0%–13%) and more

lithic fragments than van Hattum et al. (2003) report. This supports their observation that the West Crocker formation is compositionally less mature than the rest of the Crocker Formation and also suggests that an easterly provenance area was relatively more important. Similarly, William et al. (2003) conclude that the West Crocker formation was derived from nearby, uplifted Rajang Group sediments because of the composition and immature texture, although they prefer an easterly provenance area based on petrographic analysis, regional sedimentary systems, and stratigraphic relationships. Synsedimentary deformation is more compatible with an Eocene rather than with an Oligocene to early Miocene age for the West Crocker formation. The younger age would require that strong compression, and probably subduction, continued into the early Miocene (ca. 20 Ma), whereas many workers (e.g., Tongkul, 1990) contend that subduction ended in the middle Oligocene (ca. 30 Ma). An Eocene age also provides ample time for the estimated 4–8 km of burial that the West Crocker underwent before it was uplifted and exposed in the middle Miocene (Hutchison, 2005); the younger age necessitates that both the burial and uplift occurred in the early middle Miocene and were, therefore, extremely rapid. The West Crocker formation lies below a regional middle Miocene unconformity beneath which the Oligocene to early Miocene geology is characterized by extension in the Natuna Sea (Darman and Sidi, 2000), Sarawak (Mat-Zin, 1997; Madon, 1999), and the Dangerous Grounds off the shore of western Sabah (Milsom et al., 1997). In offshore western Sabah there is pre–middle Miocene extension, and many late Miocene anticlines cored with middle Miocene sedimentary rocks are interpreted as an inversion of older structures (Bol and van Hoorn, 1980; Madon et al., 1999). The problem posed by development of a major thrust belt synchronous with regional extension has been answered previously by having extension in the South China Sea induce subduction beneath Borneo, thereby causing the Sabah orogeny (Hutchison, 1996; Hall, 2002). An Eocene age for the West Crocker formation eliminates this problem and calls into question models for the region’s early Neogene tectonic evolution. CONCLUSIONS An integrated analysis of the structural geology and sedimentology of the West Crocker formation concludes that the abundant sandstone beds were deposited from high-density turbidity currents and formed small, prograding lobes on a deep-water slope apron. The composition and immature texture of the sands necessitates a short transport distance, suggesting that uplifted Rajang Group sediment was the source of the West Crocker formation. Micropaleontological data and stratigraphic relationships indicate that at least this part of the West Crocker formation is Eocene rather than Oligocene to early Miocene. Nearly all of the structural deformation is compressional and synsedimentary; the West Crocker formation is a syntectonic deposit, probably an accretionary prism.

West Crocker formation of northwest Borneo ACKNOWLEDGMENTS The authors wish to thank George Pemberton for assistance with trace fossil identifications and Nurhayati and Rhodora Tiamsing for help with drafting. Mohd. Khalid Jamiran contributed to the interpretation of stratigraphic architecture, and the manuscript benefited greatly from discussions with Stefan Back and comments by reviewers Charles Hutchison, Peter Clift, and an anonymous reviewer. Amelia William and Tan Yaw Tzong wish to thank ExxonMobil Exploration and Production Malaysia and Universiti Brunei Darussalam, respectively, for financial support. REFERENCES CITED Bol, A.J., and van Hoorn, B., 1980, Structural styles in western Sabah offshore: Geological Society of Malaysia Bulletin, v. 12, p. 1–16. Briais, A., Patriat, P., and Tapponnier, P., 1993, Updated interpretation of magnetic anomalies and seafloor spreading stages in the South China Sea: Implications for the Tertiary tectonics of Southeast Asia: Journal of Geophysical Research, v. 98, p. 6299–6328. Carozzi, A.V., 1993, Sedimentary Petrography: Englewood Cliffs, NJ, Prentice Hall, 263 p. Crevello, P.D., 2002, The great Crocker submarine fan: A world class foredeep turbidite system: Jakarta, 28th Indonesia Petroleum Association Proceedings, v. 1, p. 377–407. Darman, H., and Sidi, F.H., 2000, An Outline of the Geology of Indonesia: Jakarta, Indonesian Association of Geologists, 192 p. Dickinson, W.R., and Suczek, C.A., 1979, Plate setting and sandstone composition: American Association of Petroleum Geologists Bulletin, v. 63, p. 2164–2182. Frey, R.W., and Pemberton, S.G., 1984, Trace fossil facies models, in Walker, R.G., ed., Facies Models (2nd edition): Geological Association of Canada, p. 189–207. Hall, R., 2002, Cenozoic geological and plate tectonic evolution of SE Asia and the SW Pacific: Computer-based reconstructions, model and animations: Journal of Asian Earth Sciences, v. 20, p. 353–431, doi: 10.1016/S13679120(01)00069-4. Harwood, G., 1988, Microscopic techniques: II. Principles of sedimentary petrography, in Tucker, M., ed., Techniques in Sedimentology: Oxford, UK, Blackwell Science, p. 108–173. Hutchison, C.S., 1996, Geological evolution of South East Asia: Kuala Lumpur, Geological Society of Malaysia, 350 p. Hutchison, C.S., 2004, Marginal basin evolution: The southern South China Sea: Marine and Petroleum Geology, v. 21, p. 1129–1148, doi: 10.1016/ j.marpetgeo.2004.07.002. Hutchison, C.S., 2005, Geology of Northwest Borneo: Amsterdam, Elsevier, 421 p. Hutchison, C.S., Bergman, S.C., Swauger, D.A., and Graves, J.E., 2000, A Miocene collisional belt in north Borneo: Uplift mechanism and isostatic adjustment quantified by thermochronology: Geological Society [London] Journal, v. 157, p. 783–793. Kumar, N., and Slatt, R.M., 1984, Submarine-fan and slope facies of Tonkawa (Missourian–Virgilian) sandstone in deep Anadarko Basin: American Association of Petroleum Geologists Bulletin, v. 68, p. 1839–1856. Levell, B.K., 1987, The nature and significance of regional unconformities in the hydrocarbon-bearing Neogene sequences offshore West Sabah: Geological Society of Malaysia Bulletin, v. 21, p. 55–90. Liechti, P., Roe, F.N., Haile, N.S., and Kirk, H.J.C., 1960, The Geology of Sarawak, Brunei and the Western Part of North Borneo: British Borneo Geological Survey Bulletin, v. 3, 360 p. Lowe, D.R., 1982, Sediment gravity flows II. Depositional models with special reference to the deposits of high-density turbidity currents: Journal of Sedimentary Petrology, v. 52, p. 279–297. Macdonald, D.I.M., 1986, Proximal to distal sedimentological variation in a linear turbidite trough: Implications for the fan model: Sedimentology, v. 33, p. 243–259, doi: 10.1111/j.1365-3091.1986.tb00534.x.

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Macdonald, D.I.M., 1993, Controls on sedimentation at convergent plate margins: International Association of Sedimentologists Special Publication 20, p. 225–257. Madon, M., 1999, Geological setting of Sarawak, in Petroleum Geology and Resources of Malaysia: Kuala Lumpur, Petronas, p. 273–290. Madon, M., and Abolins, P., 1999, Balingian Province, in Petroleum Geology and Resources of Malaysia: Kuala Lumpur, Petronas, p. 343–367. Madon, M., Leong, K.M., and Anuar, A., 1999, Sabah Basin, in Petroleum Geology and Resources of Malaysia: Kuala Lumpur, Petronas, p. 501–542. Mat-Zin, I.C., 1997, Tectonics, evolution and sedimentation of the Sarawak Basin: Geological Society of Malaysia Bulletin, v. 41, p. 41–52. Milsom, J., Holt, R., Ayub, D.B., and Smail, R., 1997, Gravity anomalies and deep structural controls at the Sabah-Palawan margin, South China Sea, in Petroleum Geology of Southeast Asia: Geological Society [London] Special Publication 126, p. 417–427. Mohd, Khalid Jamiran, 1999, Sedimentology and sequence stratigraphy of Bukit Melinsung, Papar, Sabah [M.S. thesis]: Universiti Brunei Darussalam, 50 p. Morley, C.K., 2007, Growth of folds in a deepwater setting: Basin Research (in press). Normark, W.R., 1978, Fan valleys, channels, and depositional lobes on modern submarine fans; characters for recognition of sandy turbidite environments: American Association of Petroleum Geologists Bulletin, v. 62, p. 912–931. Pemberton, S.G., MacEachern, J.A., and Frey, R.W., 1992, Trace fossil facies models: Environmental and allostratigraphic significance, in Walker, R.G., and James, N.P., eds., Facies Models: Response to Sea Level Change, Geotext 1: Geological Association of Canada, p. 47–72. Pettijohn, F.J., Potter, P.E., and Siever, R., 1987, Sand and Sandstone (2nd edition): New York, Springer-Verlag, 571 p. Rangin, C., Bellon, H., Bernard, F., Letouzey, J., Muller, C., and Sanudin, T., 1990, Neogene arc-continent collision in Sabah, Northern Borneo, Malaysia: Tectonophysics, v. 183, p. 305–319, doi: 10.1016/0040-1951(90)90423-6. Reading, H.G., and Richards, M., 1994, Turbidite systems in deep-water basin margins classified by grain size and feeder system: American Association of Petroleum Geologists Bulletin, v. 78, p. 792–822. Ricci Lucchi, F., and Valmori, E., 1980, Basin-wide turbidites in a Miocene, over-supplied deep-sea plain: A geometrical analysis: Sedimentology, v. 27, p. 241–270, doi: 10.1111/j.1365-3091.1980.tb01177.x. Sandal, S.T., ed., 1996, The Geology and Hydrocarbon Resources of Negara Brunei Darussalam: Bandar Seri Begawan, Syabas, 243 p. Stauffer, P.H., 1967, Studies in the Crocker Formation, Sabah, in Stauffer, P.H., ed., Geological Papers 1966: Geological Survey of the Borneo Region of Malaysia, Bulletin 8, p. 1–13. Stow, D.A.V., and Johansson, M., 2000, Deep-water massive sands: Nature, origin and hydrocarbon implications: Marine and Petroleum Geology, v. 17, p. 145–174, doi: 10.1016/S0264-8172(99)00051-3. Stow, D.A.V., Reading, H.G., and Collinson, J.D., 1996, Deep seas, in Reading, H.G., ed., Sedimentary Environments and Facies (3rd edition): Oxford, UK, Blackwell Science, p. 395–453. Tajul Anuar Jamaluddin, 1989, Struktur sedimen dalam Formasi Crocker di kawasan Tamparuli, Sabah: Geological Society of Malaysia Bulletin, v. 24, p. 135–157. Tan, D.N.K., and Lamy, J.M., 1990, Tectonic evolution of the NW Sabah continental margin since the Late Eocene: Geological Society of Malaysia Bulletin, v. 27, p. 241–260. Tan, Y.T., 2005, Structural geology of the Maju outcrop, West Crocker Formation, Kota Kinabalu, Sabah, Malaysia [M.S. thesis]: Universiti Brunei Darussalam, 72 p. Tate, R.B., 1974, Palaeo-environmental studies in Brunei: Brunei Museum Journal, v. 3, p. 285–305. Tongkul, F., 1987, The sedimentology and structure of the West Crocker Formation in the Kota Kinabalu area, Sabah, in GEOSEA IV Proceedings: Jakarta, Indonesia Association of Geologists, p. 135–156. Tongkul, F., 1990, Structural style and tectonics of Western and Northern Sabah: Geological Society of Malaysia Bulletin, v. 27, p. 227–239. Tongkul, F., 1991, Tectonic evolution of Sabah, Malaysia: Journal of Southeast Asian Earth Sciences, v. 6, p. 395–405, doi: 10.1016/07439547(91)90084-B. van Hattum, M.W.A., Hall, R., and Nichols, G.J., 2003, Provenance of northern Borneo sediments: Jakarta, 29th Indonesia Petroleum Association Proceedings, v. 1, paper IPA03-G-016.

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Wilson, R.A.M., 1964, The geology and mineral resources of the Labuan and Padas Valley area, Sabah, Malaysia: Geological Survey of the Borneo Region of Malaysia, Memoir 17.

MANUSCRIPT ACCEPTED BY THE SOCIETY 24 APRIL 2007

Printed in the USA

The Geological Society of America Special Paper 436 2008

Temporal changes in the composition of Miocene sandstone related to collision between the Honshu and Izu Arcs, central Japan Koichi Okuzawa* Research Center for Deep Geological Environments, National Institute of Advanced Industrial Science and Technology, Tsukuba 305-8567, Japan Ken-ichiro Hisada Graduate School of Life and Environmental Sciences, University of Tsukuba, Tsukuba 305-8572, Japan ABSTRACT The Izu Arc has been colliding with the Honshu Arc in central Japan since ca. 15 Ma. In order to understand the provenance changes related to this collision, we studied lower to middle Miocene sandstones in and around the collision zone by analyzing their framework composition and the chemistries of detrital clinopyroxene, garnet, and chromian spinel. Sandstone deposited in the trench and forearc basin of the Honshu Arc prior to collision includes grains of detrital garnet and chromian spinel, which originated mainly from granites and low pressure-temperature (P-T) metamorphic rocks, and forearc peridotite, respectively, parts of the Honshu Arc. The forearc and trench-fill sandstones differ in terms of their framework composition; sedimentary lithics are more abundant in the forearc sandstone than in the trench. The two groups of sediments were supplied from different parts of the Honshu Arc. The lower part of the clastic sequence deposited within the Izu Arc is composed mainly of volcaniclastic rocks and yields detrital clinopyroxenes that originated from the Izu Arc. In contrast, the upper part is similar to the lower Miocene trench-fill deposits in terms of its framework composition and the chemistry of detrital garnet and chromian spinel. This reflects a change in provenance triggered by the initial contact of the Izu Arc and the trench between the Eurasian and Philippine Sea plates. The lower part of the middle Miocene trench-fill that was deposited following initial contact is also similar to the lower Miocene trench-fill. The upper part, however, resembles lower Miocene sedimentary rocks of the forearc basin. This suggests that the transport path was changed by collision. During the initial stages of collision between the Honshu and Izu Arcs, the Honshu Arc was preferentially uplifted, and therefore supplied most of the detritus to the collision zone. Keywords: sandstone, provenance, framework composition, detrital, clinopyroxene, garnet, chromian spinel, Honshu Arc, Izu Arc, collision. *Present address: Institute for Geo-Resources and Environment, National Institute of Advanced Industrial Science and Technology, Tsukuba 305-8567, Japan; [email protected] Okuzawa, K., and Hisada, K., 2008, Temporal changes in the composition of Miocene sandstone related to collision between the Honshu and Izu Arcs, central Japan, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 185–198, doi: 10.1130/2008.2436(09). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Okuzawa and Hisada Although the Izu Collision Zone is a suitable area for studies that seek to understand the nature of sediments generated in arcarc collision settings, there are few petrographical studies of sediments from the collision zone. According to Soh (1986), sandstone and conglomerate from upper Miocene to Pliocene trenchfill sedimentary rocks are composed mainly of sedimentary lithics rather than volcanic lithics. In contrast, sandstone from the lower Miocene accretionary complex adjacent to the collision zone is conspicuously rich in quartz and feldspar (Watanabe and Iijima, 1990). It appears that a change in provenance occurred during the early to middle Miocene (ca. 13–17 Ma) at approximately the same time that the collision zone formed. Several types of detrital heavy minerals have been examined in various provenance studies to determine the source rocks of the sediment. Chemical analysis of detrital heavy minerals has proved to be a powerful tool in provenance studies (e.g., Morton, 1991). Chromian spinel, garnet, and clinopyroxene have also been used to infer the nature of the sediment source rocks (e.g., Hisada and Arai, 1993; Takeuchi, 1994; Cawood, 1983). Chromian spinel generally originates from mafic to ultramafic

INTRODUCTION

150°

140°

120°

A

130°

The Izu Arc, an oceanic island arc, has been colliding with the Honshu Arc, a continental arc, in central Japan since ca. 15 Ma (Fig. 1); this zone of collision is termed the Izu Collision Zone (Taira et al., 1989). There are four main blocks of accreted Izu Arc crust: the Koma, Misaka, Tanzawa, and Izu Blocks (Fig. 2). Each block was accreted to the Honshu Arc at different times since ca. 15 Ma (Niitsuma and Akiba, 1985; Amano, 1986, 1991; Aoike, 1999). The accreted blocks are associated with thick clastic sequences showing coarsening-upward trends (e.g., Ito, 1985, 1987) and regression-type facies (e.g., Huchon and Kitazato, 1984). These sequences are interpreted to have been deposited in the trench between the Honshu Arc and the accreted blocks (Amano, 1986, 1991; Koyama, 1991, 1993). The collision history of each block was examined in previous studies on the basis of the lithology and geochemistry of the accreted blocks, sequence stratigraphy, and the geochronology and deformation history of trench-fill, as summarized by Amano et al. (1999) and Aoike (1999).

50°

American

Plate

Plate

140°E 38°N

100 km

North Eurasia

138°E

B

40°

Japan Sea

HG ns

h

c

Hd 36°N

Izu

Pacific

Arc

Plate

Sb 30°

Philippine Sea 140°

130°

Plate

Ch

Low P/T type metamorphic rocks Ophiolite

Boso Pen.

L MT

Mn

Fig. 2

Outer zone

High P/T type metamorphic rocks

Ry

rc

Permian accretionary complex

Hy

Izu a

Cretaceous to Tertiary accretionary complex Jurassic accretionary complex

Sh

Mn Inner zone

Ho

r uA

Circum-Izu massif serpentinite

34°N

Forearc peridotite (Kurosegawa Belt)

Continental block

Figure 1. Geologic outline of the Honshu Arc. (A) Plate tectonic setting of Japan. (B) Geologic map of basement rocks of the central Honshu Arc. Modified from Isozaki (1996). MTL—Median Tectonic Line; Hd—Hida Belt; HG—Hida Gaien Belt; Mn—Mino Belt; Ry—Ryoke Belt; Sb—Sambagawa Belt; Ch—Chichibu Belt (including Northern Chichibu, Kurosegawa, and Southern Chichibu Belts); Sh—Shimanto Belt (including Setogawa, Kobotoke, Hayama, and Mineoka Belts); Hy—Hayama; Mn—Mineoka.

Collision between the Honshu and Izu Arcs 138°E

187

139°E

N

Koma Group Ko

bo

tok

eB

elt

ISTL

Study area

Belt

Mt. Fuji

ga

Shim

Sa

anto

Mt. Hakone

mi

35°N

Suruga Trough

gh

u Tro 20km

Trench fill sediments

Ryuso Group

+

Figure 2. Geologic map of the Izu Collision Zone. Compiled from Tsuchi (1986), Koyama (1993), Sugiyama (1995), and Martin and Amano (1999). ISTL—Itoigawa-Shizuoka Tectonic Line. The study area is indicated by the three rectangular boxes.

Kurami Group

Setogawa

Izu Block

Izu

Oigawa Group

Belt

Tanzawa Block

Collision

Setogawa Group

Misaka Block

Zone

Intrusive rocks

Koma Block

Cretaceous-Oligocene

rocks; accordingly, the occurrence of detrital chromian spinel can be used to constrain the timing of emplacement of mafic or ultramafic rocks. The chemistry of spinel can also provide clues to the tectonic setting in the source region (e.g., Arai, 1992; Barnes and Roeder, 2001). Garnet is commonly found within metamorphic rocks and granite. Garnet has a complicated chemistry, and the chemical composition of metamorphic garnet depends upon both the parent rock and the metamorphic grade (Miyashiro, 1953). Clinopyroxene occurs as a stable phase in almost all types of igneous rocks, as well as various rocks that formed under conditions of both regional and contact metamorphism (Deer et al., 1982). Although chemical analysis of detrital heavy minerals is a useful tool for understanding provenance, there has previously been no example applying it to the sedimentary rocks in the Izu Collision Zone.

This paper describes the framework composition, mineralogy, and mineral chemistry of sandstone from pre- to syncollisional successions within the Izu Collision Zone. The chemistry of detrital minerals was compared with the mineral chemistry of potential source rocks proximal to the study area in order to identify provenance changes caused by arc-arc collision. OUTLINE OF REGIONAL GEOLOGY The geology of the central Honshu Arc is characterized by a zonation of geological units (Fig. 1B). The central part of the Honshu Arc is divided into the Inner and Outer Zones, which are separated by the Median Tectonic Line (MTL, Fig. 1B). The Inner Zone is composed of four belts that constitute a continental block (Hida Belt), an ophiolite (Hida Gaien Belt), a Jurassic

Okuzawa and Hisada Izu Collision Zone

Setogawa Belt

20

western part

central & eastern parts

Ks

Koma Group

Oigawa Group

40

50

Early Late

Eocene

35

Late

Oligocene

25

30

Mm

Kurami Group

Early

Miocene

15

Middle

10

E. Middle

accretionary complex (Mino Belt), and low pressure-temperature (P-T) metamorphic rocks (Ryoke Belt; Figure 1B; Isozaki, 1996). Cretaceous to early Tertiary granites (e.g., Ryoke Granite) and felsic volcanic rocks (e.g., Nohi Rhyolite) are widely distributed throughout the Inner Zone. The Outer Zone consists of a high P-T metamorphic accretionary complex (Sambagawa Belt), a Jurassic accretionary complex (Northern Chichibu Belt), an ophiolite (Kurosegawa Belt), a Jurassic accretionary complex (Southern Chichibu Belt), and a Cretaceous to Tertiary accretionary complex (Shimanto, Setogawa, Kobotoke, Hayama, and Mineoka Belts; Figure 1B; Isozaki, 1996). The Tertiary accretionary complex (Setogawa, Kobotoke, Hayama, and Mineoka Belts) includes a number of serpentinite bodies, which are called the Circum–Izu Massif serpentinites (Arai, 1991); these four accrectionary complex belts are also termed the Circum–Izu Massif Serpentine Belt (Arai, 1991). These serpentinite bodies are accompanied by gabbro and basaltic rocks of the mid-ocean-ridge basalt (MORB) type and intraplate type (e.g., Ogawa and Taniguchi, 1988; Hirano et al., 2003), and are considered to have formed as crust generated by seafloor spreading in a backarc environment (Arai, 1991). Alternatively, Arai and Okada (1991) and Arai (1994) proposed that the Circum–Izu Massif serpentinites were initially heterogeneous in their petrological characteristics. Arai (1994) proposed that the incipient Circum–Izu Massif serpentinites were formed in an arc environment. The serpentinites were later amalgmated with gabbro and basaltic rocks (e.g., Ogawa and Taniguchi). The study area covers parts of the Setogawa Belt and the Izu Collision Zone (Fig. 2). An early Miocene to early middle Miocene accretionary complex is distributed in the Setogawa Belt and is composed of three stratigraphic units, the Setogawa, Oigawa, and Ryuso Groups (Fig. 2). In the western part of the Setogawa Belt, the Setogawa Group is covered by the lower Miocene Kurami Group and younger sedimentary rocks, which were deposited in a forearc basin (Fig. 2). The Izu Collision Zone comprises accreted Izu Arc crust and trench-fill material (Osozawa et al., 1990; Hiroki and Matsumoto, 1999; Amano, 1991). The four lower to middle Miocene groups were targeted in this study in order to reconstruct the provenance changes associated with the early stage of Honshu-Izu collision: the Setogawa, Oigawa, and Kurami Groups of the Setogawa Belt and the Koma Group of the Izu Collision Zone. The Ryuso Group in the Setogawa Belt was excluded from this study because the group is composed mostly of volcanic rocks (Sugiyama, 1995) and is inadequate for the framework composition work. The Setogawa Group comprises basalt, limestone, chert, tuffaceous mudstone, and alternating sandstone and mudstone (Fig. 3; Sugiyama, 1995). The tuffaceous mudstone contains tectonic blocks of serpentinite, which is a component of the Circum-Izu Massif serpentinites (Arai, 1991), intercalated with conglomerate containing clasts of serpentinite. The Oigawa Group consists mainly of alternating sandstone and mudstone with minor amounts of conglomerate and sandstone (Sugiyama, 1995). The ages of the clastic rocks in the Setogawa and Oigawa Groups are early Miocene (24–19 Ma) and early to early middle

Age (Ma)

188

Setogawa Group

clastic rocks chert

tuffaceous mudstone limestone

volcanic rocks

Figure 3. Correlation diagram of the stratigraphic units targeted in this study. Ks—Kushigatayama Subgroup; Mm—Momonoki Subgroup. Compiled from Koyama (1993), Sugiyama (1995), and Hiroki and Matsumoto (1999).

Miocene (ca. 22–15 Ma), respectively (Fig. 3; Sugiyama, 1995). A total of seven sandstone samples were collected from the lower Miocene accretionary complex: three from the Setogawa Group, and four from the Oigawa Group. The Kurami Group crops out in the SW part of the Setogawa Belt (Fig. 2), where it unconformably overlies the Setogawa Group (Fig. 3). The Kurami Group consists of basal conglomerate, sandstone, alternating sandstone and mudstone, tuffaceous mudstone, and mudstone in ascending stratigraphic order (Ibaraki, 1986; Hiroki and Matsumoto, 1999). The group was deposited in the late early Miocene (18–15.5 Ma; Hiroki and Matsumoto, 1999) and represents forearc basin sediments (Hiroki, 1995; Hiroki and Matsumoto, 1999). Three sandstone samples were collected from the lower part of the Kurami Group. The Koma Group consists of two subgroups: the lower part is the Kushigatayama Subgroup, and the upper is the Momonoki Subgroup (Fig. 3; Kosaka and Tsunoda, 1969), with the Kushigatayama Subgroup being subdivided into lower and upper parts (Koyama, 1991, 1993). The lower part of the Kushigatayama Subgroup is composed mainly of tholeiitic intermediate to mafic volcaniclastic rocks and lava flows (Shimazu and

Collision between the Honshu and Izu Arcs Ishimaru, 1987). The upper part consists mainly of intermediate to mafic tuff intercalated with mudstone, sandstone, and conglomerate. The Momonoki Subgroup consists of alternating sandstone and mudstone, conglomerate, mudstone, siliceous mudstone, and felsic tuff (Koyama, 1991, 1993). The sequence shows a coarsening-upward trend (Aoike, 1999). Conglomerate clasts within the Momonoki Subgroup are derived from both the Honshu and Izu Arcs (Koyama, 1993). The Kushigatayama Subgroup was deposited within the Izu Arc, whereas the Momonoki Subgroup represents trench-fill sediment (Amano, 1991; Aoike, 1999). The Kushigatayama and Momonoki Subgroups were deposited at 16–15 Ma and 15–13.5 Ma, respectively, and are in faulted contact (Aoike, 1999). We collected a total of nine sandstone samples from the Koma Group: three from the lower part of the upper Kushigatayama Subgroup, two from the uppermost Kushigatayama Subgroup, and two each from the lower and upper Momonoki Subgroup.

189

A

Qm

Setogawa Group Oigawa Group Kurami Group

CI

Kushigatayama Subgp. Momonoki Subgp.

QR

TC Mixed

TR

Type A DA

Type B

BU TA

LR

Type C

F

B

Lt Ls

METHODS We mostly collected medium-grained sandstone samples from the Setogawa, Oigawa, Kurami, and Koma Groups. Thin sections were stained for K-feldspar and point-counted for 500 points per sample, following the Gazzi-Dickinson method (Ingersoll et al., 1984). Heavy liquid separation was performed to collect clinopyroxene, garnet, and chromian spinel from the selected 13 samples. A total of 100 g of each specimen was crushed for heavy liquid separation. Separated heavy minerals were then mounted and thin sectioned. The chemistry of these grains was analyzed using electron probe microanalysis (JXA8621 Super Microprobe; JEOL, Tokyo) at the Research Facility Center for Science and Technology, University of Tsukuba, Japan. Operating conditions were 20 kV accelerating voltage, 10 nA specimen current, and ~10-μm beam diameter. As no zoning was detected microscopically, we analyzed the center of each grain. The results of all the microprobe analyses were published previously by Okuzawa (2004).

UA

Qm: monocrystalline quartz F: total feldspar Lt: total lithics Ls: sedimentary lithics Lvf: felsic volcanic lithics Lvm: intermediate to mafic volcanic lithics

Type B Type A Type C

Lvf

Lvm

Figure 4. Framework composition data for sandstone analyzed in the present study. (A) Qm-F-Lt diagram. (B) Ls-Lvf-Lvm diagram. Compositional fields in Qm-F-Lt diagram are after Dickinson et al. (1983). CI—craton interior; TC—transitional continent; BU—basement uplift; QR—quartzose recycled; TR—transitional recycled; LR—lithic recycled; DA—dissected arc; TA—transitional arc; UA—undissected arc.

RESULTS Framework Composition of Sandstone Sandstone in the study area can be divided into types A, B, and C on the basis of framework composition, as indicated by the fields shown in Figure 4. Sandstone of the Setogawa and Oigawa Groups, the uppermost part of the Kushigatayama Subgroup (sample 02102504, Table 1), and the lower part of the Momonoki Subgroup (samples 02102101 and 02102207, Table 1) belongs to type A. This type of sandstone consists mainly of monocrystalline quartz (14%–41%), felsic volcanic lithics (11%–31%), and plagioclase (10%–25%), with minor amounts of K-feldspar (1%–11%; Table 1). Type A sandstone plots in the field of dissected to transitional magmatic arc (Dickinson et al., 1983; Fig. 4A).

Sandstone of the Kurami Group, the uppermost Kushigatayama Subgroup (sample 02102506, Table 1), and the upper Momonoki Subgroup (samples 02102208 and 02102303, Table 1) are classified as type B. This type of sandstone is composed mainly of felsic volcanic lithics (13%–36%), monocrystalline quartz (12%–22%), sedimentary lithics (mainly mudstone: 10%–19%), and plagioclase (7%–17%), with minor K-feldspar (0%–5%) and polycrystalline quartz (0%–6%; Table 1). Type B sandstone is richer in sedimentary lithics than that of type A (Fig. 4B) and plots in the field of transitional magmatic arc (Dickinson et al., 1983; Fig. 4A). Sandstone from the lower part of the upper Kushigatayama Subgroup (samples 02102501, 02102502, and 02102503, Table 1) belongs to type C. This sandstone consists mainly of

190

Okuzawa and Hisada TABLE 1. FRAMEWORK COMPOSITIONS OF SANDSTONE SAMPLES Sample no.

Qm

Qp

K-fel.

Pl.

Lvf

Setogawa

02092607 02093005 02093008

36.6 38.0 14.8

0.4 0.0 0.0

11.2 8.6 1.4

15.4 18.6 23.8

15.6 13.8 29.8

0.0 0.0 0.0

Oigawa

02122603 02122604 02122606 02122609

22.6 35.2 41.2 36.4

0.6 0.8 0.0 0.2

2.0 3.4 7.0 4.2

15.0 10.0 15.4 16.2

26.6 30.8 10.8 20.2

Kurami

02122803 02122807 02122810

22.4 11.6 21.0

4.6 2.2 3.6

4.0 5.0 5.0

11.2 12.0 7.0

Koma (Kushigata-yama Subgroup)

02102501 02102502 02102503 02102504 02102506

1.6 0.0 9.6 14.2 17.6

0.8 0.0 0.6 0.6 0.4

0.2 0.0 2.8 6.8 0.4

8.8 34.2 25.2 24.6 17.2

Group

2.8 0.8 2.0

If and Aut. 0.4 0.0 1.4

Mtx and Cem. 16.7 19.8 26.0

100.0 100.0 100.0

0.4 0.4 2.6 0.2

1.0 0.4 1.2 1.8

0.0 0.0 0.0 0.0

31.2 19.0 21.8 20.4

100.0 100.0 100.0 100.0

0.0 0.0 0.0

0.8 1.2 0.4

0.0 0.0 0.2

0.0 0.0 0.0

15.4 36.8 17.6

100.0 100.0 100.0

0.0 0.0 0.0 0.0 0.0

5.6 1.2 0.0 1.2 0.6

4.4 9.8 0.0 1.0 0.2

0.0 0.0 9.2 0.0 0.0

31.6 15.8 15.0 21.8 36.8

100.0 100.0 100.0 100.0 100.0

Other lithics

H.m.

S.m.

0.2 0.0 0.4

0.0 0.0 0.0

0.8 0.4 0.4

0.0 0.0 0.0 0.0

0.6 0.0 0.0 0.4

0.0 0.0 0.0 0.0

22.4 20.8 26.8

0.0 0.8 0.0

19.2 9.6 18.4

27.0 15.2 27.8 23.6 13.2

17.4 23.6 4.0 1.4 0.6

2.6 0.2 5.8 4.8 13.0

(Momonoki Subgroup)

Lvi

Ls

02102101 27.6 0.0 1.6 15.4 23.4 0.0 0.6 0.0 02102207 36.6 0.2 5.2 12.4 23.6 0.6 3.8 0.0 02102208 13.8 0.2 1.8 12.6 21.4 5.8 13.2 0.0 02102303 19.4 5.8 4.8 9.4 35.8 0.2 10.0 0.0 Note: Qm—monocrystalline quartz; Qp—polycrystalline quartz; K-fel.—K-feldspar; Pl.—plagioclase; Lvf—felsic volcanics; Ls—sedimentary lithics; H.m.—heavy minerals; S.m.—secondary minerals; If and Aut.—intrabasinal and Cem.—matrix and cement.

Total (%)

0.2 0.4 0.0 30.8 100.0 0.2 0.2 0.0 17.2 100.0 0.8 1.0 0.2 29.2 100.0 0.4 0.4 0.2 13.6 100.0 volcanics; Lvi—intermediate to mafic fragments and authigenic minerals; Mtx

Sandstone from the lower part of the upper Kushigatayama Subgroup (type C sandstone) yields abundant detrital clinopyroxene grains. Almost all of these grains are colorless to light green and are angular to subrounded. We obtained chemical analyses of 30 detrital clinopyroxenes. Twenty-eight of the analyzed grains are augite (Fig. 5A). They are Ca-rich and Fe-poor, with low contents of TiO2 (0.1–0.4 wt%), Cr2O3 (0.0–0.1 wt%), Na2O (0.2–0.4 wt%), and total Al (total Al as Al2O3: 0.6–1.5 wt%). The other two grains are augite and diopside, and are characterized by a higher Cr content (Cr2O3 >0.5 wt%) and total Al (3.1–4.2 wt%), with TiO2 and Na2O contents of 0.3 wt% and 0.1–0.2 wt%, respectively.

Eighty-eight detrital garnets were analyzed from the Oigawa Group, of which 70 are classified as pyrope-rich almandines (Pyr = 5–45 mol%), 10 are spessartine-rich almandines (Sps = 9–40 mol%), six are Ca-rich almandines (Grs + And = 14–23 mol%), and two are almandine-rich spessartines (Sps = 52 and 62 mol%; Fig. 6A). Forty-nine detrital garnets were obtained from the Kurami Group, of which 24 are spessartine-rich almandines (Sps = 7–44 mol%), 20 are pyrope-rich almandines (Pyr = 8–41 mol%), three are Ca-rich almandines (Grs + And = 14–17 mol%), one is an almandine-rich spessartine (Sps = 45 mol%), and one is an Mgrich spessartine (Sps = 53 mol%; Fig. 6B). Sixty-one detrital garnets were analyzed from the Koma Group. Their compositional range is similar to that of garnets from the Kushigatayama and Momonoki Subgroups (Fig. 6C). These grains are mostly pyrope-rich almandines (Pyr = 10–45 mol%) and spessartine-rich almandines (Sps = 6–45 mol%), with just seven grains of Ca-rich almandines (Grs + And = 6–37 mol%) and one of andradite-rich grossular (Grs = 89 mol%).

Chemistry of Detrital Garnet

Chemistry of Detrital Chromian Spinel

A total of 269 chemical analyses of detrital garnets were obtained from the Setogawa, Oigawa, Kurami, and Koma Groups, with the exception of the lower part of the upper Kushigatayama Subgroup. Seventy-one grains of detrital garnet were analyzed from sandstones of the Setogawa Group: 49 grains are classified as pyrope-rich almandine (Pyr = 7–45 mol%), 15 grains are spessartine-rich almandines (Sps = 9–47 mol%), 3 grains are Ca-rich almandines (Grs + And = 11–27 mol%), 3 grains are almandinerich spessartines (Sps = 43–63 mol%), and one is a Ca-rich spessartine (Sps = 40 mol%; Fig. 6A).

All the sandstone samples that contain detrital garnets also yield reddish-brown to black detrital chromian spinels. We obtained a total of 324 chemical analyses of chromian spinel grains. The chemical characteristics of detrital chromian spinels from the Setogawa, Oigawa, Kurami, and Koma Groups are similar. The spinels from the four groups are characterized by a wide range in Cr# [= Cr/(Cr+Al) atomic ratio], from 0.2 to 1.0 (Fig. 7). These spinels show a negative trend in terms of the relationship between Mg# [= Mg/(Mg+Fe2+) atomic ratio] and Cr# (Fig. 7). Values of Fe3+# [= Fe3+/(Cr+Al+Fe3+) atomic ratio] are mostly

felsic to mafic volcanic fragments (32%–44%) together with plagioclase (9%–34%; Table 1). Type C sandstone plots in the field of undissected and transitional magmatic arc (Dickinson et al., 1983; Fig. 4A). Chemistry of Detrital Clinopyroxene

Collision between the Honshu and Izu Arcs

A

Diopside

Hedenbergite

DIOPSIDE

HEDENBERGITE

L. Kushigatayama Subgroup AUGITE PIGEONITE

CLINOENSTATITE

CLINOFERROSILITE

Enstatite

Ferrosilite

B 0.04

Ti

Orogenic calcalkali basalts

Orogenic tholeiitic basalts 0

0

0.1

0.2

0.3

Al total Figure 5. Chemical composition data for detrital clinopyroxenes analyzed in the present study. The subdivision scheme (A) is after Morimoto (1988). The range of volcanic rocks of the lower Kushigatayama Subgroup is surrounded by broken lines (Shimazu and Ishimaru, 1987). The fields of orogenic calc-alkali and tholeiitic basalts (B) are after Leterrier et al. (1982).

<0.1, whereas TiO2 wt% is dominantly in the range 0.0–0.5, although some grains record Fe3+# values >1.0. PROVENANCE OF PRE-COLLISION SEDIMENTARY ROCKS Trench Sedimentary Rocks Sandstone of the Setogawa and Oigawa Groups (labeled Type A in Fig. 4) that was deposited in the trench prior to collision (Aoike, 1999) contains fragments of granite, felsic volcanic rock, and sedimentary rocks such as chert and mudstone. Pre–early Miocene granitic rock, felsic volcanic rock, and chert have not been found in the Izu Arc. Possible sources for these rock fragments include Cretaceous to early Tertiary granites and felsic volcanic rocks from the Inner Zone of the Honshu Arc, as well as a Jurassic to Cretaceous accretionary complex in the Honshu Arc.

191

Detrital garnets from the Setogawa and Oigawa Groups have similar chemical characteristics (Fig. 6A), and are mostly spessartine-rich almandine and pyrope-rich almandine (Fig. 6A). Spessartine-rich almandine is found in granitic rocks and low-grade to amphibolite facies metamorphic rocks (Fig. 6A; Nanayama, 1997). Low and high P-T metamorphic rocks and granites from the Honshu Arc (Fig. 1) in the area proximal to the Setogawa Belt are candidates as source rocks for the spessartine-rich almandine. Intrusion of granitic rocks in the Izu Collision Zone postdates sedimentation of the Setogawa and Oigawa Groups, and they cannot be source rocks for the detrital garnet. Most of the detrital spessartine-rich almandines plot in the same field as garnets from the granites and low P-T metamorphic rocks (Fig. 6A). Accordingly, we conclude that the detrital spessartine-rich almandines were probably derived from these rocks. Detrital pyrope-rich almandines from the Setogawa and Oigawa Groups plot in the fields of amphibolite, pegmatite, lowgrade metamorphic rocks, and granulite facies metamorphic rocks (Fig. 6A; Nanayama, 1997). Thirty-five grains of the almandine garnet analyzed from the lower Miocene Setogawa and Oigawa Groups have relatively high Mg contents (MgO >10 wt%). The Mg/Fe ratio of Ca-poor garnet is known to increase with increasing metamorphic grade (e.g., Coleman et al., 1965). Granulites that contain garnets with Mg contents >10 wt% have not been found in the Honshu Arc or in the Izu Arc (e.g., Teraoka et al., 1998), although pyrope-rich almandines have been reported from Jurassic to Cretaceous sandstones of the Honshu Arc (Adachi and Kojima, 1983; Teraoka et al., 1999). Such garnets have also been described from the Sino-Korean Massif of mainland Asia (e.g., Jiang, 1988), which is likely to have been the source of the detrital pyrope-rich garnets in the lower Miocene Setogawa and Oigawa Groups. Because detrital garnets are rare in the sandstones of the Setogawa and Oigawa Groups, they may have been reworked from pre-Miocene sandstones. The chemistry of detrital garnets from the middle Miocene trench-fill sedimentary rocks of the Koma Group is similar to that of detrital garnets from the lower Miocene Setogawa and Oigawa Groups (Figs. 6A, C). However, the Sea of Japan had begun to open in the middle Miocene (>15 Ma; e.g., Otofuji et al., 1985), and it is therefore difficult to envisage that the Sino-Korean Massif supplied detrital garnet to the study area at this time. Pyroperich garnets of the Koma Group were probably reworked from pre–middle Miocene sandstones. Detrital chromian spinels obtained from trench sedimentary rocks of the Setogawa and Oigawa Groups are characterized by a wide range in Cr#, low Fe3+#, and low TiO2 content, although some have a higher TiO2 content (>1.0 wt%; Fig. 7A). A high Cr#, coupled with low Ti and Fe3+ concentrations, is characteristic of chromian spinels from forearc peridotite (Bloomer and Fisher, 1987; Ishii et al., 1992; Parkinson and Pearce, 1998) or boninites (Arai, 1992; Barnes and Roeder, 2001). Chromian spinels with low TiO2 content plot mainly within the field of forearc peridotite (Fig. 7A[i]). Consequently, forearc peridotite

192

Okuzawa and Hisada

A Lower Miocene trench sedimentary rocks

B Lower Miocene forearc basin sedimentary rocks

Setogawa Group (N=71) Oigawa Group (N=88) Alm

Alm

Pyr

GNF

Sps

Kurami Group (N=49) Pyr

GNF

Sps ECF

ECF PG+Met.

PG+Met.

PG+Low Met.

PG+Low Met. Skarn

Skarn GNF

Alm

GNF

ECF

APF

Pyr

Gro+And

Alm

APF

C Middle Miocene collision zone sedimentary rocks Alm

Pyr

GNF

Sps

GNF

Alm

APF

ECF

Gro+And

Granites and metamorphic rocks in the Ryoke Belt

Sps: spessartine Alm: almandine Pyr: pyrope Gro: grossular And: andradite

PG+Met.

Skarn

Pyr

Metamorphic rocks in the Sambagawa Belt

ECF

PG+Low Met.

ECF

Koma Group (N=61) Momonoki Subgroup Kushigatayama Subgroup

PG: pegmatite Low Met.: low metamorphic rock APF: amphibolite facies ECF: eclogite facies GNF: granulite facies.

+And Pyr Gro+And

Figure 6. Chemical composition data for detrital garnets. Garnets from (A) lower Miocene trench sedimentary rocks; (B) lower Miocene forearc basin sedimentary rocks; (C) middle Miocene collision zone sedimentary rocks. Pegmatite and four metamorphic fields with solid lines are after Nanayama (1997). The fields of Sambagawa Belt metamorphic rocks are compiled from Higashino et al. (1981, 1984), Enami (1983), Sonobe and Takasu (2000), and Sakurai (2000). The field of granites and metamorphic rocks in the Ryoke belt is compiled from Ono (1969, 1975a, b, 1976, 1977, 1981), Asami and Hoshino (1980), Asami et al. (1982), Nureki et al. (1982), Seo et al. (1982), Shiba (1982, 1989), Wang (1985), Takagi and Nagahama (1987), Enami (1988), and Miyazaki (1999).

is a possible source rock for most of the detrital chromian spinels that have low TiO2 concentrations. TiO2-rich (>1.0 wt%) detrital chromian spinels are derived from a different source rock than that yielding the low-Ti spinels. High-Ti spinels with intermediate Cr# are found in intraplate magmas (Arai, 1992; Barnes and Roeder, 2001). Intraplate basalt is therefore a suitable source rock for the high-Ti detrital spinels.

Several bodies of mafic and ultramafic rocks occur close to the study area: forearc peridotites of the Kurosegawa Belt, peridotites and basaltic rocks of the high P-T Sambagawa Metamorphic Belt, and backarc basin peridotites and basaltic rocks of the Circum–Izu Massif Serpentine Belt (Fig. 1). All of these mafic and ultramafic rocks are exposed within the Honshu Arc. There is no possible source rock for detrital chromian spinel known in the Izu Arc. Basalts from the Circum–Izu Massif Serpentine

A Lower Miocene trench sedimentary rocks 1

ii)

0

Setogawa Group (N=125) Oigawa Group (N=111)

iii)

Cr

5

TiO2wt%

Cr# [Cr/(Cr+Al)]

i)

0

1

Mg# [Mg/(Mg+Fe2+)]

Cr# [Cr/(Cr+Al)]

1

Fe3+

Al

B Lower Miocene forearc basin sedimentary rocks 1

ii) Forearc peridotite

iii)

Kurami Group (N=43)

Cr

5 Intraplate basalts of Circum-Izu Massif Serpentine Belt TiO2wt%

Cr# [Cr/(Cr+Al)]

i)

Kurosegawa serpentinite Circum-Izu Massif Serpentinites

0

Mg# [Mg/(Mg+Fe2+)]

0

1

Cr# [Cr/(Cr+Al)]

1

Fe3+

Al

C Middle Miocene collision zone sedimentary rocks 1

ii) 5

Koma Group (N=86) Momonoki Subgroup Kushigatayama Subgroup

iii)

Cr

TiO2wt%

Cr# [Cr/(Cr+Al)]

i)

Sambagawa greenstone and peridotite 0

Mg# [Mg/(Mg+Fe2+)]

1

0

Cr# [Cr/(Cr+Al)]

1

Al

Fe3+

Figure 7. Chemical composition data for detrital chromian spinels analyzed in the present study. The range of the Circum–Izu Massif serpentinites is enclosed by thin lines (Arai et al., 1990). The range of the intraplate basalts in the Circum–Izu Massif Serpentine Belt is surrounded by a thick line (Okuzawa and Hisada, 2004). The range of the forearc peridotite is surrounded by a dotted line (Bloomer and Fisher, 1987; Parkinson and Pearce, 1998). The range of the Kurosegawa serpentinite is circled by a broken line (Okuzawa et al., 2004). The range of the peridotite and greenstone in the Sambagawa Belt is surrounded by a dotted thin line (Uesugi and Arai, 1999).

194

Okuzawa and Hisada

Belt have intraplate and MORB-type ocean floor signatures (e.g., Ogawa and Taniguchi, 1988; Hirano et al., 2003) and are interpreted as having been formed within a backarc setting or a mid-ocean-ridge setting (Hirano et al., 2003). However, Arai and Okada (1991) and Arai (1994) reported an occurrence of high Cr# detrital spinels from tectonic blocks of serpentine-rich sandstone within tectonic blocks of the Circum–Izu Massif serpentinites. Although the sedimentary age of the serpentine-rich sandstone is unsettled, Arai (1994) concluded that those sandstones were deposited immediately after emplacement of the Circum–Izu Massif serpentinites, and that these incipient serpentinites had petrological characteristics of peridotite formed in an arc setting. Many low-TiO2 detrital chromian spinels from the Setogawa and Oigawa Groups do not plot within the field of spinels from the Kurosegawa forearc peridotite, but instead within the field of spinels from the recent Circum–Izu Massif serpentinites (Fig. 7A[i], [ii]). There are three possible combinations of source rocks for the low-TiO2 detrital spinels: (1) ancient Circum–Izu Massif serpentinites that supplied all the low-TiO2 detrital spinels found in the Setogawa and Oigawa Groups, (2) ancient Circum–Izu Massif serpentinites and other forearc peridotites and/or boninites that supplied the low-TiO2 detrital spinels, and (3) other forearc peridotites and boninites that supplied all the low-TiO2 detrital spinels. In the Boso Peninsula (Fig. 1) the lower Miocene Mineoka Group and middle Miocene Sakuma Group contain detrital spinels with low Ti content and a wide range of Cr# (Cr# = 0.3–0.9: Okuzawa and Hisada, 2004). These authors concluded that the Circum–Izu Massif serpentinites had supplied detrital chromian spinels to both groups, because both groups include conglomerate containing clasts of serpentinite, gabbro, and basalts, which probably originated from the ophiolitic rocks in the Circum–Izu Massif Serpentine Belt. On this basis, the first scenario above appears to be the most plausible. Most of the high-TiO2 chromian spinels found in the Setogawa and Oigawa Groups plot within the field of alkaline basalt from the Circum–Izu Massif Serpentine Belt (Fig. 7A[ii]). Peridotites and basalts in the high P-T Sambagawa Metamorphic Belt also contain high-TiO2 spinels (Fig. 7A[ii]; Uesugi and Arai, 1999), however, most of the detrital spinels have a lower Fe3+# than spinels in the peridotite and basalt of the Sambagawa Belt (Fig. 7A[iii]). The high-TiO2 detrital chromian spinels were probably derived from intraplate basalts of the Circum–Izu Massif Serpentine Belt. The detrital chromian spinels with higher values of Mg# and Cr# also plot outside the fields of the forearc peridotites and Circum–Izu Massif serpentinites in an Mg#-Cr# diagram (Fig. 7A[i]). Intraplate basalt from the Circum–Izu Massif Serpentine Belt contains some spinels with higher values of Mg# and Cr# than those of the forearc peridotites and Circum–Izu Massif serpentinites (Ishida et al., 1990); however, most of the detrital chromian spinels have lower TiO2 contents than spinels in the intraplate basalts. According to Barnes and Roeder (2001), chromitites within ophiolite contain chromian spinels with high Mg# and Cr# and low Fe3+# and TiO2 contents. On this basis, some of the detrital spinels analyzed in the present study appear

to be derived from chromitite-bearing ophiolite. It is known that the Circum–Izu Massif serpentinites include a small amount of chromitite (Kitahara, 1954). In summary, the framework composition of sandstones and the chemistry of detrital chromian spinels and garnets indicate that prior to collision the trench sediments were supplied from the Honshu Arc. Forearc Basin Sedimentary Rocks Sandstone of the upper lower Miocene Kurami Group (labeled Type B in Fig. 4) was deposited in the forearc basin of the Honshu Arc and has a different framework composition from that of lower Miocene trench sedimentary rocks: The forearc basin sandstone is richer in sedimentary lithics in comparison with the trench sandstone, although the assemblage of detrital grains (Table 1) and the chemical characteristics of detrital chromian spinels and garnets (Figs. 6 and 7) are similar. Marsaglia and Ingersoll (1992) concluded that very little compositional variation is related to forearc basin geometry or to submarine transport from the forearc into the trench. The compositional differences recorded between lower Miocene trench and forearc basin sedimentary rocks indicate that detritus from the two groups of rocks was derived from different provenances. The forearc basin sedimentary rocks appear to have had their source from a different part of the Honshu Arc in comparison with the trench-fill sediments. The detritus of the Setogawa Group may have been axial currents along the trench axis, because the framework composition of the Setogawa Group sandstone resembles the trench sandstone of the lower Miocene Mineoka Group in the Boso Peninsula (Watanabe and Iijima, 1990; Okuzawa, 2004). CHANGES IN PROVENANCE FOLLOWING COLLISION The lower part of the upper Kushigatayama Subgroup sandstone comprises sandstone of type C (see above; Fig. 4). This rock contains intermediate to mafic volcanic lithics (Fig. 4B) and abundant detrital clinopyroxene. A volcanic arc is therefore a likely provenance for this rock type. We used the discrimination diagram of Leterrier et al. (1982) to infer the tectonic setting of the source rock on the basis of detrital clinopyroxene geochemistry (Fig. 5B). Two detrital clinopyroxenes with Cr2O3 >0.5 wt% are excluded from analysis because they might have been derived from different source rocks. Plots of detrital clinopyroxenes from the lower part of the upper Kushigatayama Subgroup form a tight cluster on the discrimination diagram, with most grains plotting in the field of orogenic tholeiitic basalts (Fig. 5B). On this basis, we conclude that detrital clinopyroxenes from the Kushigatayama Subgroup were derived from orogenic tholeiitic basalts. Lower Miocene volcanic rocks from the Honshu Arc consist of both tholeiitic and calc-alkaline rocks (e.g., Shuto et al., 1988). Basalts from the Circum–Izu Massif Serpentine Belt

Collision between the Honshu and Izu Arcs have intraplate and MORB-type ocean floor geochemical signatures (e.g., Ogawa and Taniguchi, 1988; Hirano et al., 2003). The chemistry of clinopyroxene in these rocks is different from that of orogenic tholeiitic basalts. Clinopyroxenes within high P-T metamorphic rocks from the Sambagawa Belt show omphacite and aegirine-augite compositions (Higashino et al., 1981; Sakurai, 2000) and are therefore different from detrital clinopyroxene grains in the upper Kushigatayama Subgroup. Volcanic rocks from the lower Kushigatayama Subgroup are island arc tholeiites and contain clinopyroxenes that are mainly diopside and augite, with minor chrome diopside (Fig. 5A; Shimazu and Ishimaru, 1987). The chemical characteristics of the detrital clinopyroxenes from the lower part of the upper Kushigatayama Subgroup correspond with those of clinopyroxene-bearing volcanic rocks from the lower Kushigatayama Subgroup. Fujioka and Saito (1992) reported chemical analyses of detrital clinopyroxenes from Oligocene to Pleistocene volcaniclastic sediments and sedimentary rocks sampled during Ocean Drilling Program (ODP) Leg 126 at the Izu Arc. Detrital clinopyroxenes from the upper Kushigatayama Subgroup analyzed in the present study are similar to those from Oligocene to Miocene sedimentary rocks sampled from ODP Leg 126 sites. In summary, detritus from the lower part of the upper Kushigatayama Subgroup appears to have been derived from the Izu Arc (Fig. 8A). Sandstone from the uppermost Kushigatayama Subgroup and the lower Momonoki Subgroup is classified within our type A category (see above; Fig. 4) and is similar to sandstone within the lower Miocene trench-fill (Fig. 8B) in terms of the chemistry of detrital heavy minerals (Figs. 6 and 7). The change from type C to type A that is observed stratigraphically upward within the Kushigatayama Subgroup is interpreted as a change in provenance related to the arrival of the Izu Arc at the Honshu Trench. This indicates that the basin within the Izu Arc in which the Kushigatayama Subgroup was deposited had moved close to the trench by 16–15 Ma (Aoike, 1999). Sandstone of the upper Momonoki Subgroup is classified as type B and contains detrital heavy minerals of similar chemistry to those within the lower Miocene Kurami Group, deposited in the forearc basin (Figs. 6 and 7). As discussed above, sandstone of type B was supplied from a different source area of the Honshu Arc from that of type A (Fig. 8C). The change in provenance recorded within the Momonoki Subgroup appears to have been caused by a change in the passage of detritus to the trench owing to progressive collision between the two arcs. Conglomerate within the upper Momonoki Subgroup contains volcanic clasts derived from both the Kushigatayama Subgroup and the Honshu Arc (Koyama, 1993); however, judging from the dominance of sedimentary lithics over intermediate to mafic volcanic lithics in sandstone (Fig. 4A) and the lack of detrital clinopyroxene, sediments of the Momonoki Subgroup were probably derived mainly from the Honshu Arc. During the early stages of collision within the Izu Collision Zone, the Honshu Arc appears to have been preferentially uplifted and thereby supplied more detritus to the trench in comparison with sources in the Izu Arc.

A

195

Lower part of the upper Kushigatayama Subgroup (eve of the collision) Type A

rc

Type B

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Upper part of the upper Kushigatayama Subgroup and lower Momonoki Subgroup (begining of collision) arc

shu

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. Gp

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Upper Momonoki Subgroup (late stage of collision of Koma Block)

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Figure 8. Temporal changes to framework composition of the Koma Group sandstone at three stages: (A) Lower part of the upper Kushigatayama Subgroup. (B) Upper part of the upper Kushigatayama Subgroup and the lower Momonoki Subgroup. (C) Upper Momonoki Subgroup. See the text for explanation of types A, B, and C sandstones.

CONCLUSIONS To understand the change in provenance resulting from collision between the Izu and Honshu Arcs, we studied the framework composition of sandstone and the chemical composition of detrital clinopyroxene, garnet, and chromian spinel from lower to

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middle Miocene coarse-grained sedimentary rocks deposited in and around the Izu Collision Zone. Sediments that were deposited prior to collision were derived from the Honshu Arc, including Cretaceous to early Tertiary granites, low P-T metamorphic rocks, Jurassic to Cretaceous accretionary complexes, and forearc peridotite and/or backarc basin ophiolite. Sedimentary rocks of the pre-collisional trench and forearc basin of the Honshu Arc differ in terms of the framework composition of sandstone. Sediments of the lower part of the upper Kushigatayama Subgroup were derived from volcanic rocks of the Izu Arc, whereas those of the uppermost Kushigatayama Subgroup and the lower Momonoki Subgroup resemble lower Miocene trench fill. This change in provenance was caused by the arrival of the Izu Arc at the Honshu Trench. Sandstone of the upper part of the Momonoki Subgroup is similar to lower Miocene sedimentary rocks deposited in the forearc of the Honshu Arc. This suggests that the passage of detritus to the trench changed following collision. In the early stages of collision between the Izu and Honshu Arcs the Honshu Arc was preferentially uplifted and consequently provided the bulk of the detritus supplied to the trench. ACKNOWLEDGMENTS We appreciate the informative and helpful discussions with Shoji Arai on the chemistry of chromian spinel and with Yujiro Ogawa on tectonics in the Boso Peninsula. Comments on the manuscript by Peter D. Clift, Andy Morton, Kantaro Fujioka, and an anonymous reviewer are acknowledged. We extend our thanks to Teruo Ohtsubo, Naoto Hirano, Jun-ichiro Kuroda, and Shoichiro Tokumine for their help in the field work. Thanks are due to Norimasa Nishida for his help with microprobe analysis. REFERENCES CITED Adachi, M., and Kojima, S., 1983, Geology of the Mt. Hikagedaira area, east of Takayama, Gifu Prefecture, central Japan: Journal of Earth Sciences, Nagoya University, v. 31, p. 37–67. Amano, K., 1986, Southern Fossa Magna as multiple collision belt: Earth Monthly (Chikyu), v. 8, p. 581–585 (in Japanese). Amano, K., 1991, Multiple collision tectonics of the South Fossa Magna in Central Japan: Modern Geology, v. 15, p. 315–329. Amano, K., Martin, A.J., Tanakadate, H., Kanaguri, S., Yoda, N., and Aizu, T., 1999, Tectonics and basin formation in the arc-arc collision zone: The example of the South Fossa Magna in central Japan: Structural Geology (Journal of the Tectonic Research Group of Japan), no. 43, p. 11–20 (in Japanese with English abstract). Aoike, K., 1999, Tectonic evolution of the Izu Collision Zone: Research Report of the Kanagawa Prefectural Museum of Natural History, v. 9, p. 111–151 (in Japanese with English abstract). Arai, S., 1991, The Circum–Izu Massif peridotite, central Japan, as back-arc mantle fragments of the Izu-Bonin arc system, in Peters, Tj., et al., eds., Ophiolite Genesis and Evolution of the Oceanic Lithosphere: Dordrecht, Netherlands, Kluwer Academic Publishers, p. 801–816. Arai, S., 1992, Chemistry of chromian spinel in volcanic rocks as a potential guide to magma chemistry: Mineralogical Magazine, v. 56, p. 173–184, doi: 10.1180/minmag.1992.056.383.04. Arai, S., 1994, The Circum–Izu Massif Serpentine Belt: Geoscience Reports of Shizuoka University, v. 20, p. 175–185, in Japanese with English abstract.

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Association of Mineralogists, Petrologists, and Economic Geologists, v. 82, p. 382–394 (in Japanese with English abstract). Shuto, K., Takimoto, T., Sakai, A., Yamazaki, T., and Takahashi, T., 1988, Geochemical variation with time of the Miocene volcanic rocks in northern part of the Northeast Japan arc: Journal of the Geological Society of Japan, v. 94, p. 155–172 (in Japanese with English abstract). Soh, W., 1986, Reconstruction of Fujikawa trough in Mio-Pliocene age and its geotectonic implication: Kyoto, Japan, Memoirs of the Faculty of Science, Kyoto University, Series of Geology and Mineralogy, v. 52, p. 1–68. Sonobe, M., and Takasu, A., 2000, Major and trace element chemistry of the garnets within the Sambagawa pelitic schists in the Asemigawa area, central Shikoku, Japan: Geoscience Report of Shimane University, v. 19, p. 151–166. Sugiyama, Y., 1995, Geology of the northern Setogawa Belt in the Akaishi Mountains and the formation of the Setogawa accretionary complex: Bulletin of the Geological Society of Japan, v. 46, p. 177–214 (in Japanese with English abstract). Taira, A., Tokuyama, H., and Soh, W., 1989, Accretion tectonics and evolution of Japan, in Ben-Avraham, Z., ed., Evolution of the Pacific Ocean Margin: New York, Oxford University Press, p. 100–123. Takagi, H., and Nagahama, H., 1987, The Ryoke Belt in the Hiki Hills, northeastern marginal area of the Kanto Mountains: Journal of the Geological Society of Japan, v. 93, p. 210–215 (in Japanese with English abstract). Takeuchi, M., 1994, Changes in garnet chemistry show a progressive denudation of the source areas for Permian-Jurassic sandstones, Southern Kitakami Terrane, Japan: Sedimentary Geology, v. 93, p. 85–105, doi: 10.1016/0037-0738(94)90030-2.

Teraoka, Y., Suzuki, M., and Kawakami, K., 1998, Provenance of Cretaceous and Paleogene sediments in the Median Zone of Southwest Japan: Bulletin of the Geological Society of Japan, v. 49, p. 395–411 (in Japanese with English abstract). Teraoka, Y., Okumura, K., Suzuki, M., and Kawakami, K., 1999, Clastic sediments of the Shimanto Supergroup in Southwest Japan: Bulletin of the Geological Society of Japan, v. 50, p. 559–590 (in Japanese with English abstract). Tsuchi, R., 1986, 1:200,000 Geologic Map of Shizuoka Prefecture: Shizuoka, Japan, Shizuoka Prefecture. Uesugi, J., and Arai, S., 1999, The Shiokawa peridotite mass in the Mikabu belt, central Japan, as a cumulate from intra-plate tholeiite: Memoirs of Geological Society of Japan, no. 52, p. 229–242 (in Japanese with English abstract). Wang, G., 1985, A Ca-Mn-Fe garnet found in the Ryoke metamorphic rocks at the Wazuka area, Kyoto Prefecture: Journal of the Japanese Association of Mineralogists, Petrologists, and Economic Geologists, v. 80, p. 459–462 (in Japanese with English abstract). Watanabe, Y., and Iijima, A., 1990, Evolution of the Tertiary SetogawaKobotoke-Mineoka turbidite fills: Tokyo, Journal of the Faculty of Science, University of Tokyo, Section II, v. 20, p. 425–441.

MANUSCRIPT ACCEPTED BY THE SOCIETY 24 APRIL 2007

Printed in the USA

The Geological Society of America Special Paper 436 2008

Cenozoic volcanic arc history of East Java, Indonesia: The stratigraphic record of eruptions on an active continental margin Helen R. Smyth* Robert Hall Gary J. Nichols SE Asia Research Group, Geology Department, Royal Holloway University of London, Egham TW20 0EX, UK

ABSTRACT The stratigraphic record of volcanic arcs provides insights into their eruptive history, the formation of associated basins, and the character of the deep crust beneath them. Indian Ocean lithosphere was subducted continuously beneath Java from ca. 45 Ma, resulting in formation of a volcanic arc, although volcanic activity was not continuous for all of this period. The lower Cenozoic stratigraphic record on land in East Java provides an excellent opportunity to examine the complete eruptive history of a young, well-preserved volcanic arc from initiation to termination. The Southern Mountains Arc in Java was active from the middle Eocene (ca. 45 Ma) to the early Miocene (ca. 20 Ma), and its activity included significant acidic volcanism that was overlooked in previous studies of the area. In particular, quartz sandstones, previously considered to be terrigenous clastic sedimentary rocks derived from continental crust, are now known to be of volcanic origin. These deposits form part of the fill of the Kendeng Basin, a deep flexural basin that formed in the backarc area, north of the arc. Dating of zircons in the arc rocks indicates that the acidic character of the volcanism can be related to contamination of magmas by a fragment of Archean to Cambrian continental crust that lay beneath the arc. Activity in the Southern Mountains Arc ended in the early Miocene (ca. 20 Ma) with a phase of intense eruptions, including the Semilir event, which distributed ash over a wide area. Following the cessation of the early Cenozoic arc volcanism, there followed a period of volcanic quiescence. Subsequently arc volcanism resumed in the late Miocene (ca. 12–10 Ma) in the modern Sunda Arc, the axis of which lies 50 km north of the older arc. Keywords: East Java, Indonesia, stratigraphic record, Cenozoic arc volcanism.

*Present address: Cambridge Arctic Shelf Programme (CASP), 181A Huntingdon Road, Cambridge CB3 0DH, UK Smyth, H.R., Hall R., and Nichols, G.J., 2008, Cenozoic volcanic arc history of East Java, Indonesia: The stratigraphic record of eruptions on an active continental margin, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 199–222, doi: 10.1130/2008.2436(10). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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INTRODUCTION The island of Java lies within the Indonesian archipelago between the landmasses of Eurasia and Australia, on the margin of the Eurasian plate. The southeastern part of this plate is known as Sundaland (Fig. 1), which is the Mesozoic continental core of SE Asia (e.g., van Bemmelen, 1949; Hamilton, 1979). Java is located on Sundaland’s southern margin. Many studies of volcanic arcs are based on well-exposed areas of arcs active long ago, such as the Cretaceous Allistos Arc in Baja California (Busby et al., 2006) and the Cambrian– Ordovician Lough Nafooey Arc of the Irish Caledonides (e.g., Draut and Clift, 2001). There are few studies of younger arcs in areas where active volcanism continues today. One reason for this is that younger arc products cover older arc sequences. Furthermore, many modern arcs are in tropical areas, such as Indonesia, the Philippines, and the Western Pacific, that are difficult to work in, commonly are not well exposed, and are remote. The modern volcanoes on Java form part of the Sunda Arc, which is well known for volcanic activity, including the famous nineteenth century eruptions of Krakatau and Tambora, and the Pleistocene eruption of Toba. However, despite a long volcanic record in Java, owing to subduction of Indian Ocean lithosphere during the Cenozoic, little is known of its pre-Pleistocene arc history. Since the middle Eocene (ca. 45 Ma) there has been northward subduction of the Indo-Australian plate at the Java Trench to the south of Java (Hall, 2002). Consequently, Java is essentially a volcanic island, and a long history of arc volcanism is recorded in its Cenozoic stratigraphy. Products of both active and ancient volcanic arcs can be observed. An east-west–trending chain of >30 modern volcanoes, forming part of the Sunda Arc,

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creates the central spine of the island of Java (Fig. 2). An arc of older, Eocene to Miocene volcanoes is parallel to and south of the Sunda Arc and is known as the Southern Mountains Arc (Smyth et al., 2005). Exposures of basement rocks are rare in East Java, and before this study the basement was thought to have formed during the Cretaceous, when fragments of arc and ophiolitic material were accreted to the southern margin of Sundaland (Wakita, 2000). There are no continental basement rocks at the surface or reported from drilling in East Java. The basement is often referred to as “transitional,” as to the north and west in Sundaland the basement is continental, and in the east near to Flores, it is oceanic (e.g., Hamilton, 1979). The Cenozoic stratigraphic record preserved on land in East Java provides an excellent opportunity to examine the eruptive history of a young, well-preserved volcanic arc. The presentday arc, active since the middle to late Miocene (ca. 12–10 Ma; Soeria-Atmadja et al., 1994), is immediately obvious, but the importance of the older Southern Mountains Arc has previously been underestimated. The relatively basic volcanic products of this arc, the “Old Andesites” (van Bemmelen, 1949), are widely known because they are well exposed and form prominent topographic features. However, the acidic products of the arc, which are widespread, have been largely overlooked. As a result of the work reported here, abundant quartz-rich sedimentary rocks in East Java are now known to be the primary and epiclastic products of acidic volcanism but were previously interpreted as continental sediments (e.g., Harahap et al., 2003). The acidic nature of the erupted material is a reflection of the character of the underlying crust, which as a result of new zircon dating is now thought to include a fragment of Archean continental crust. The full cycle of the Paleogene Southern Mountains Arc, from initiation to termination, can be documented in the stratigraphy of East Java. This paper reports the stages of arc development and the sequence and timing of events within the arc, determined by detailed field investigations accompanied by provenance analysis and isotopic dating. Most of the middle Miocene was marked by a period of volcanic quiescence that followed the termination of Paleogene arc volcanism in the early Miocene (ca. 20 Ma). Near the end of the middle Miocene (ca. 12–10 Ma) the modern Sunda Arc began activity in its present position, 50 km north of the Southern Mountains Arc. Possible explanations for the shift in position of arc volcanism are discussed later in this paper.

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Figure 1. Current plate tectonic setting of western Indonesia. The margin of Sundaland (Hamilton, 1979) is defined by a black dashed line. The black arrows show relative plate motion of boundaries—upper plate movement relative to lower (McCaffrey, 1996). The study area is enclosed by a rectangle.

East Java can be subdivided into three parts, broadly parallel to the elongation of the island, representing (1) the early Cenozoic Southern Mountains Arc, (2) a deep basin north of the arc, and (3) a marine shelf north of the basin (Figs. 3 and 4). The modern arc is built mainly on top of the basin north of the early Cenozoic arc. We first summarize the principal features of these components of Java and then go on to discuss the stratigraphy of the Southern Mountains Arc.

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Southern Mountains Arc There are few exposures of basement rocks in Java, and in East Java they are known only in the western part of the study area. Based upon this limited evidence, the basement has been

interpreted to be arc and ophiolitic rocks of Cretaceous age. A volcanic arc was built upon basement rocks from the middle Eocene to the Miocene in southern Java (Smyth, 2005), which is known to extend from East Java into West Java (SoeriaAtmadja et al., 1994). The stratigraphic thickness of the arc

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products is >2500 m, and within this sequence andesites are well known (van Bemmelen, 1949; Soeria-Atmadja et al., 1994), but acidic volcanic rocks have not been previously reported. The zone containing the arc products is 50 km wide. The Southern Mountains Arc is now uplifted and partially eroded. The strata typically dip uniformly toward the south between 20° and 30°. Kendeng Basin The Kendeng Basin lies directly behind and to the north of the Southern Mountains Arc (Fig. 2). The deposits of the basin are poorly exposed. The depocenter is marked by a strong negative Bouguer gravity anomaly of more than −580 μm–2 (Fig. 5). The negative anomaly can be traced eastward into a negative marine free air anomaly (Sandwell and Smith, 1997), which extends from the Straits of Madura eastward to the north of Bali. At the west end of the Kendeng Basin a relatively abrupt change in the character of the anomaly is around the modern volcanoes of Merapi and

the Dieng Plateau (Fig. 5). To the west, in West Java, the anomaly is not well defined, as it becomes positive (40 μm–2). No basement rocks are exposed or known from drilling in this region. The basin fill has an age range of middle Eocene to Miocene, similar to the Southern Mountains Arc (de Genevraye and Samuel, 1972). The basin is east-west oriented, at least 400 km long, parallel to the Southern Mountains Arc, and filled with a succession of volcaniclastic turbidites and pelagic mudstones that are reported to be at least 6 km thick (Untung and Sato, 1978). Gravity calculations indicate that the basin may contain as much as 10 km of sediment (C.J. Ebinger, 2005, personal commun.). Based on field observations, seismic sections, and regional gravity interpretation (Smyth, 2005), the basin is estimated to have been ~100–120 km wide during the early Cenozoic. It is now partially exposed at the surface in the Kendeng Fold-Thrust Belt, where there is an estimated 10–30 km of shortening (de Genevraye and Samuel, 1972; Smyth, 2005); de Genevraye and Samuel (1972) report that deformation commenced in the Pliocene.

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Edge of Sunda Shelf To the north of the Kendeng Basin (Figs. 2 and 4) the hills of North Java and the offshore East Java Sea constitute an area interpreted to be the edge of the early Cenozoic Sunda Shelf (Hamilton, 1979). This region has been the focus of most hydrocarbon exploration. Pre-Cenozoic basement rocks sampled by drilling are known to be ophiolitic and arc rocks, which include chert and basic volcanic and metasedimentary rocks (e.g., Hamilton, 1979). Basin development began in the Eocene. There are between 2000 and 6000 m of Eocene to Pliocene shallow marine clastic and extensive carbonate sedimentary rocks within fault-controlled basins (e.g. Ardhana, 1993; Ebanks and Cook, 1993). The sedimentary sequences have been deformed since the late Miocene by numerous open, east-west–oriented folds; northward-verging, east-west–oriented thrusts interpreted to be Pliocene in age; and ENE-WSW normal faults (Chotin et al., 1984; Hoffmann-Rothe et al., 2001). Modern-Day Volcanic Arc The active volcanoes of the Sunda Arc are built mainly on the Kendeng Basin (Figs. 3 and 4) but locally overlap the edge

of the Southern Mountains Arc. Arc activity commenced in its present position at ca. 10 Ma in the late Miocene (SoeriaAtmadja et al., 1994). The volcanoes exceed 3000 m in elevation and form the central spine of the island (Fig. 2). The average composition of the present-day volcanic products is basaltic andesite (Nicholls et al., 1980), which is much more basic than the average composition of the Southern Mountains Arc products. Most of the active volcanoes are situated ~100 km above the subducting slab (England et al., 2004). There are also a number of unusual K-rich backarc volcanoes (Edwards et al., 1991), which occur to the north of the axis of the arc, including Muria (Fig. 2). The eruptive and epiclastic products of the Sunda Arc cover a significant proportion of the island, contributing to its fertile soils. STRATIGRAPHIC RECORD OF THE EARLY CENOZOIC SOUTHERN MOUNTAINS ARC The sedimentary rocks in the Southern Mountains Arc in East Java were deposited on the basement above a poorly dated regional unconformity (Fig. 6). The unconformity separates Upper Cretaceous basement rocks and the Cenozoic succession, suggesting a long period, potentially up to 30 m.y., during which uplift and erosion occurred. The stratigraphic succession above

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the unconformity in the Southern Mountains Arc and the Kendeng Basin has been subdivided into three “synthems” (Smyth et al., 2005), each representing a different period of arc activity (Fig. 6). The term synthem has been used because much of the work undertaken and published in East Java has been hydrocarbon oriented, and as a result the term sequence is often used with seismic and sequence stratigraphic implications (e.g., Catuneanu, 2006). We have therefore chosen to avoid the term sequence for the major stratigraphic subdivisions and in this study used the term synthem, defined as an unconformity-bounded stratigraphic package (Rawson et al., 2001). The three synthems are: • Synthem One: records the initiation of arc volcanism and the early stages of arc development during the middle Eocene (ca. 45 Ma) to early Oligocene (ca. 34–28 Ma). • Synthem Two: records the growth and termination of arc volcanism in the Southern Mountains Arc during the late Oligocene (ca. 28–23 Ma) to the early Miocene (ca. 20 Ma). • Synthem Three: records widespread carbonate growth, accompanied by the erosion and redeposition of rocks from earlier synthems during the middle Miocene (ca. 20– 10 Ma), with no significant volcanic activity. In the account of the stratigraphy that follows, the stratigraphic ages have been converted to numerical ages, using the Gradstein et al. (2004) time scale. Synthem One: Initiation of the Southern Mountains Arc Most parts of the Southern Mountains rocks of Synthem One are covered by younger deposits and are exposed only at

Karangsambung, Nanggulan, and the Jiwo Hills (Fig. 2). Synthem One is not exposed in the Kendeng Basin but was sampled in blocks brought to the surface by mud volcanoes. Southern Mountains Arc Basement exposures in East Java are rare and are found only in the western part of the area at Karangsambung and the Jiwo Hills (Fig. 2). The exposed basement rocks are Cretaceous in age (Parkinson et al., 1998; Wakita and Munasri, 1994) and include basaltic pillow lavas, radiolarian cherts, various metasedimentary lithologies, quartz-mica schist, and high-grade metamorphic rocks including jadeite-quartz-glaucophane–bearing rocks and eclogites. Quartz veins are commonly observed within the basement exposures and have not been identified in the overlying Cenozoic rocks. The high-grade metamorphic rocks have been interpreted to indicate subduction zone metamorphism (Miyazaki et al., 1998). The basement rocks are interpreted to represent fragments of arc and ophiolitic material accreted to the margin of Sundaland during the Late Cretaceous. Similar rocks have previously been assumed to extend beneath the rest of East Java, where there is no information from surface exposures or drilling (e.g., Hamilton, 1979; Parkinson et al., 1998). Synthem One is ~1000 m thick but is exposed only in the western part of the Southern Mountains Arc in East Java. The oldest sedimentary rocks rest directly on the basement above a regional angular unconformity. They are poorly dated fluvial conglomerates and interbedded sandstones (Figs. 7A and 8) and are at least 50 m thick, but their total thickness is difficult to assess owing to limited and patchy exposures. These are the only rocks

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exposed onshore that contain no fresh volcanic material (Fig. 9). They are dominated by quartz grains and metamorphic and igneous lithic clasts (Figs. 8 and 9), including vein and metamorphic quartz, chert, phyllite, schist, metasedimentary rock, and basalt. The lithologies identified as clasts are typical of those exposed in the Cretaceous basement (described above) and are interpreted to be the product of erosion and reworking of these basement rocks. The terrestrial conglomerates and sandstones lack palynomorphs and so cannot be directly dated, but they are overlain by a succession of well-dated middle Eocene strata (Lelono, 2000). The middle Eocene to lower Oligocene of the Southern Mountains Arc is represented by a transgressive succession, from base to top, of coals and conglomerates, sandstones, siltstones,

and mudstones (Fig. 8). The middle Eocene age of the lower part of the succession is based on palynomorphs in the coals and organic-rich mudstones (Lelono, 2000), and the occurrence of Nummulites and Discocyclina higher in the section. The lower part is at least 200 m thick and is a noncalcareous unit of well-bedded coals, conglomerates, quartz-rich sandstones, and organic-rich muds. The coals and conglomerates are restricted to the lower 50 m. The conglomerates are commonly channelized, are ~10–75 cm thick, and contain a range of lithic clasts similar to those identified in the basal section directly overlying the basement. The coals vary laterally in thickness from 5 to >50 cm and have an average vitrinite reflectance of 0.4% Ro (Smyth, 2005). For normal and high geothermal gradients typical

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of arc regions this would imply very shallow burial depths of <1 or 2 km (Madon et al., 1997; Watts, 1997). Above this transgressive succession lie thinly laminated quartz-rich sandstones and organic-rich mudstones that are at least 150 m thick. These sediments were deposited in a terrestrial setting with some marine incursions indicated by the presence of calcareous zones rich in Nummulites and Discocyclina. In addition to the organic-rich mudstones there are several gray mudstones, which are rich in smectite. The quartz-rich sandstones contain metamorphic lithic clasts, like those of the conglomerates, but also contain volcanic lithic clasts such as

pumice and andesite. In addition, a significant proportion of the quartz has a volcanic origin (the identification of volcanic quartz is discussed further below). At the base of the quartzrich sandstones the quartz is dominated by metamorphic grains, but this content gradually decreases in its relative contribution upsection, and at the top, volcanic quartz is dominant (Fig. 9B). The heavy mineral assemblage contains zircons that yielded spot sensitive high-resolution ion microprobe (SHRIMP) UPb ages of 41.8 ± 1.6 and 42.7 ± 1.5 Ma (Smyth, 2005). These ages are the same as the biostratigraphically determined ages for the host rocks, indicating that volcanic activity occurred

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Figure 9. Petrographic details of Synthem One. Scale bars are 1 mm. (A) Basement, volcanic, and other sources are depicted in a triangular diagram, showing the upsection trend identified in East Java, with a reduction in basement-derived material and an increase in volcanic contribution. (B) Pie charts showing the quartz types identified within quartz-rich sandstones of East Java; example from the Nanggulan Formation. Volcanic quartz increases upsection. (C) Photomicrograph of the basal sandstone, dominated by metamorphic quartz. (D) Volcaniclastic turbidite dominated by laths of plagioclase feldspar.

as the sediments were deposited. Higher in the section, in the sandstones, a number of Nummulites-rich zones mark the onset of fully marine conditions. Above the nummulitic units the sandstones become increasingly arkosic and rich in fresh laths of plagioclase feldspar (Fig. 9D), volcanic quartz, volcanic lithic fragments, elongate volcanic zircons, and volcanic clays such as smectite and

zeolites (Fig. 9A). These sandstones are interbedded with tuffaceous mudstones. Water depth is uncertain but is interpreted to be deeper than that for the nummulitic zones. There are no diagnostic sedimentary structures, but the sandstones contain pelagic foraminifers. Higher in the section the presence of volcaniclastic turbidites indicates an increase in water depth. Some turbidites are

Cenozoic volcanic arc history of East Java, Indonesia characterized by Bouma divisions A, C, and E, but more commonly C, D, and E. They are >150 m thick, with individual bed thicknesses ranging from 5 to 50 cm. The turbidites have a diverse planktonic foraminiferal assemblage (Fig. 8E), including Helicosphaera euphratis, H. reticulate, and H. wolcoxonii (P. Lunt, 2002, personal commun.), which provides an age of NP18 (36.8–36.2 Ma) for the lower parts of the succession. The turbidites extend into the lower Oligocene, based on biostratigraphy (P. Lunt, 2002, personal commun.). They are dominated by volcanic debris (Fig. 9) such as laths of plagioclase feldspar, volcanic lithic clasts, and volcanogenic clays, and there are no metamorphic lithic clasts within the coarser beds. The heavy mineral assemblage is dominated by volcanic zircons, which yielded a weighted mean SHRIMP U-Pb age from 17 grains of 41 ± 1.4 Ma, similar to the biostratigraphic age of the lower part of the section, suggesting some reworking. The upper boundary of Synthem One is an intra-Oligocene unconformity, which is interpreted to be the result of sea level change, as the sedimentary rocks directly above and below the gap have similar bedding orientations with no indication of deformation. This could be a local sea level change but could also be a global change. Unconformities of this age that record a global intra-Oligocene sea level fall are widely known, from the Haq et al. (1987) curve (30 Ma), the Marshall Paraconformity (32–29 Ma) in the Canterbury Basin, New Zealand (Fulthorpe et al., 1996), and the carbonates of Baldwin County, Alabama, USA (Baum et al., 1994). Within Synthem One the contribution of volcanic material increases upsection as the proportion of basement-derived material (metamorphic quartz, metamorphic lithic fragments, and illite, chlorite, and serpentinite clays) decreases, and the upper Eocene sedimentary rocks are almost entirely dominated by volcanic debris (Fig. 9A, B). The volcanic centers supplying this material cannot be separately mapped owing to their limited exposure, but they are interpreted to follow the same trend and occupy the same positions as the volcanoes of the Oligocene– Miocene arc (see below). Kendeng Basin In the Kendeng Basin there are limited exposures, and Synthem One is not seen at the surface, but blocks of the older lithologies have been brought to the surface by Pleistocene and modern mud volcanoes. These blocks are terrestrial and are composed of shallow marine sandstones and conglomerates (de Genevraye and Samuel, 1972) similar in character to the middle Eocene sedimentary rocks in the Southern Mountains Arc. Synthem Two: Growth and Catastrophic Termination of Arc Volcanism Southern Mountains Arc The upper Oligocene to lower Miocene deposits of Synthem Two are primary volcanic rocks and epiclastic rocks. These rocks are exposed extensively throughout the Southern Mountains Arc

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(Fig. 4) and within the fold-thrust belt of the Kendeng Basin. The oldest rocks of this synthem in the Southern Mountains Arc are reworked bioclastic tuffaceous mudstones dated biostratigraphically as Oligocene, NP24 (27.3–30 Ma; M. Fadel, 2002, personal commun.). Throughout the late Oligocene and early Miocene the volcanic activity in the Southern Mountains Arc was extensive, explosive, and intermediate to acidic in composition. The deposits range from andesite to rhyolite, with an average SiO2 content of 67 wt% (Smyth, 2005), and include thick mantling tuffs (Fig. 10A), crystal-rich tuffs, block and ash flows, pumice-lithic breccias, andesitic breccias (Fig. 10B), silicic lava domes, and lava flows. The Oligocene–Miocene volcanic centers can be mapped (Smyth, 2005) by the occurrence of vent proximal facies, and at least 13 centers are presently exposed (Fig. 2), which show a similar spacing to the volcanoes of the present-day arc. The thickness of the proximal volcanic deposits ranges from 250 to >2000 m. Thick (>700 m) successions of volcaniclastic sedimentary rocks surround the volcanic centers. These reworked deposits are commonly interbedded with beds of unreworked mantling tuffs and volcanic breccias >1 m in thickness. The reworked volcaniclastic beds vary in thickness from 5 to >100 cm and are crystaland volcanic lithic-rich sandstones and tuffaceous mudstones. They commonly contain abundant fragments of charcoal, indicating the presence of vegetated slopes on the terrestrial volcanic centers. Both terrestrial and shallow marine deposits have been identified; the latter do not contain abundant bioclastic material but are weakly calcareous, and their upper bedding surfaces are commonly intensely bioturbated with Cruziana-type facies traces. There are slump folds and mass wasting deposits on the flanks of the volcanic centers. In the Southern Mountains Arc is a record of a major eruption, the Semilir Eruption, toward the end of the period of arc activity. Extensive deposits of this eruption are widespread to the east of Yogyakarta (Fig. 4). The Semilir and Nglanggran Formations (Fig. 10A, B) are the products of this event, and they were deposited in a short period, possibly during one eruptive phase between 21 and 19 Ma (Smyth, 2005). Based on measured sections, the combined thickness of the two formations varies between 250 and 1100 m. They are well exposed over an area of 800 km2, and the total volume of volcanic material is estimated to be at least 480 km3. The Semilir Formation (Fig. 6) is a thick accumulation of dacitic air-fall, pyroclastic surge and flow deposits produced by an explosive eruption. Both terrestrial and shallow marine deposits have been identified, indicating that the erupted material entered the sea from the flanks of the volcanic center. The Nglanggran Formation (Fig. 6) is a series of monomict andesitic volcanic breccias. The individual beds are up to 10 m thick, have flat bases and tops, and can be mapped for tens of kilometers. The breccias contain some blocks >3 m across. They are interpreted as vent proximal facies such as flow breccias or block-and-ash–flow deposits. These deposits mark the end of volcanism in the

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Figure 10. Details of Synthem Two. (A) Field photograph of thick, well-bedded ashes (bar is 1m). (B) Field photograph of monomict andesitic breccia; the breccia clasts commonly exceed 1 m in size (bar is 1m). (C) Basement, volcanic, and other sources are depicted in a triangular diagram, showing a selection of formations from Synthem Two. These deposits lack material of basement origin. (D) Pie chart showing the quartz types identified petrographically in a quartz-rich sandstone in the upper part of Synthem Two. The quartz is dominated by volcanic quartz; the image is a bipyramidal grain from the Jaten Formation, Pacitan (bar is 1m).

Southern Mountains Arc. There is no significant break within or between the Semilir and Nglanggran Formations, based on zircon dating and biostratigraphic dating of the overlying formations (Smyth, 2005). A number of quartz-rich sandstones (Fig. 10D) are in the upper part of Synthem Two in the Southern Mountains Arc

(Smyth, 2005). Volcanic zircons from the sandstones form a single population with ages that are the same as that of the Semilir Formation (Smyth, 2005) and are contemporaneous with the Semilir Eruption. The quartz within these sandstones is entirely of volcanic origin (Fig. 10D). Volcanic quartz grains commonly appear clear and bright in thin section, have

Cenozoic volcanic arc history of East Java, Indonesia nonundulose extinction, and are monocrystalline (Leeder, 1982). These features and other characteristics, including well-developed crystal faces, bipyramidal shape, melt embayments, and melt inclusions, can be used to distinguish volcanic quartz from other types such as metamorphic, vein, plutonic, and sedimentary quartz. Volcanic crystal-rich sandstones can be produced by primary eruptive mechanisms and/or secondary epiclastic processes. The sandstones and other quartz-rich volcaniclastic rocks in East Java previously were interpreted as continental siliciclastic deposits (Harahap et al., 2003), and it is for this reason that many of the acidic products of the Southern Mountains Arc have been overlooked (as discussed below). Here they are interpreted to have a volcanic origin and to be largely the product of the Semilir Eruption with some subsequent reworking. Kendeng Basin Synthem Two exposures in the Kendeng Basin are very limited, occurring only in a small thrust-bound sliver. They comprise poorly lithified Globigerina-rich tuffaceous mudstones ranging in age from late Oligocene to early Miocene (de Genevraye and Samuel, 1972). The mudstones are at least 85 m thick (de Genevraye and Samuel, 1972) but have limited surface exposure. The mudstones are volcanogenic and apparently lack continental terrigenous material. The volcanic material is fine grained, reworked, and distal in character, very different from that identified within the arc at this time. The abundance of Globigerina and other planktonic foraminifers is indicative of pelagic sedimentation in an open marine setting at water depths of a few hundred meters or more. Thick volcanogenic turbidites are reported from wells in this area (P. Lunt, 2002, personal commun.), but no descriptions of the stratigraphy or well log interpretations have been published. Edge of the Sunda Shelf The offshore Eocene to Pliocene successions in the East Java Sea that have been investigated during hydrocarbon exploration are not reported to contain any volcanic material (e.g., Matthews and Bransden, 1995). Field investigations during this study in northeast Java have identified distal volcanic material on the edge of the Sunda Shelf within carbonates, including thin tuff layers, zones rich in smectite clays, and concentrations of volcanic quartz and zircons. This suggests that there may be more volcanic debris offshore than recognized up to now, and that some of the clay layers may be fine air-fall ash deposits. The stratigraphic record of Synthem Two within the Southern Mountains Arc and Kendeng Basin accounts for only volcanic and volcaniclastic rocks at this interval. There is no evidence of input of material from basement or other sources. This indicates that the volcanic arc was the only source of sediment at this time. There are no significant exposures of carbonates within Synthem Two in the Southern Mountains Arc or Kendeng Basin, but carbonates do occur farther to the north on the edge of the Sunda Shelf.

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Synthem Three: Reworking of the Southern Mountains Arc Southern Mountains Arc In the Southern Mountains, to the south of the Oligocene– Miocene volcanic centers, extensive calcareous volcanogenic turbidites (Fig. 11) occur with Bouma divisions A, C, D, and E. The turbidites are at least 500 m thick and have been dated within nannofossil subzones NN2–NN8 (Kadar, 1986) that correspond to dates between 19 and 10 Ma. Beds vary in thickness from 20 to 75 cm, and upper bedding surfaces are commonly bioturbated with traces of Cruziana facies. Flute casts indicate that the flow direction was toward the southeast, and there are slump folds with southeasterly vergence. Thin section examination shows that volcanic crystals and lithic fragments are well rounded, and there is no fresh or unreworked volcanic debris present. Dating of zircons (see below) supports the field and petrographic evidence that these are reworked volcaniclastic deposits, as the zircon ages are similar to the age range of rocks in Synthem Two. Synthem Three is therefore interpreted to be the product of reworking of older arc rocks of Synthem Two rather than the product of contemporaneous volcanism. Several tuff beds have been identified at the top of the turbidite sequence in the Southern Mountains but are not the result of volcanic activity in the Southern Mountains Arc. The tuffs are distal air-fall deposits and yield zircons with ages between 12 and 10 Ma (P.J. Hamilton, 2003, personal commun.). These tuffs are the product of the modern Sunda Arc (Soeria-Atmadja et al., 1994), and the ages mark the initiation of volcanic activity some 50 km to the north of the Southern Mountains Arc. In addition to the turbidites, there are the first widespread carbonates in the Southern Mountains (Fig. 11D). The carbonates range in age from late early Miocene to middle Miocene (Lokier, 2000; Smyth, 2005) and formed isolated reefs and extensive carbonate platforms, which can be observed overstepping the deposits of Synthem Two. The limestones are at least 200 m thick and are the source of the carbonate within the volcanogenic turbidites. Kendeng Basin Volcanogenic sedimentary rocks identified in the Kendeng Basin include channelized volcanic lithic conglomerates, crystalrich and volcanic lithic-rich sandstones, tuffaceous mudstones, and Globigerina mudstones. A thickness of up to 400 m of rocks is exposed (Smyth, 2005), and hydrocarbon exploration shows there is up to 3000 m in the basin (de Genevraye and Samuel, 1972). The sandstones contain Bouma divisions B, C, and E and are locally cut by the channelized conglomerates. These conglomerates can cut as much as 20 cm into the underlying beds and are up to 1 m thick. Slump folds are common, and most verge toward the northeast. Measurements of scours, grooves, channel structures, flutes, and slumping indicate that sediment was transported from the south toward the north. The Globigerina mudstone is similar in character to those of Synthem Two. A sample from the top of the exposed section yielded a biostratigraphic age

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Figure 11. Stratigraphic details of Synthem Three. (A) Logged section typical of the volcanogenic turbidites of Synthem Three, Sambipitu Formation. (B) Field photograph showing asymmetrical ripples on the upper bedding surfaces of the volcanogenic turbidites, Sambipitu Formation. (C) Field photograph of volcanogenic turbidites within the Kendeng Basin, Kerek Formation. (D) Field photograph of the reefal carbonates typical of Synthem Three in the Southern Mountains Arc, Campurdarat Formation.

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of foraminifer subzones N8–N9 (M. Fadel, 2003, personal commun.) corresponding to ages between 17.2 and 14.4 Ma.

in the East Java Sea (Bishop, 1980; van Bemmelen, 1949), and redeposited on the edge of the Sunda Shelf.

Edge of the Sunda Shelf On the edge of the Sunda Shelf are a number of lower to middle Miocene quartz-rich sandstones. They were previously interpreted to have a continental provenance and to have been derived from reworking of basement rocks of Sundaland (e.g., Ardhana, 1993; Sharaf et al., 2005), but new studies (Smyth et al., 2007) have shown that these sandstones contain a significant proportion of volcanic quartz, volcanic zircons, and reworked volcanogenic clays. These volcanic particles are interpreted to have fallen as ash onto the Sunda Shelf, to have been subsequently reworked and mixed with material derived from uplifted basement blocks

INTERPRETATION OF THE EOCENE TO MIOCENE STRATIGRAPHY Initiation of the Southern Mountains Arc The first evidence of arc volcanism in East Java is identified as middle Eocene at ca. 42 Ma with the first occurrence of tuff layers, volcanic lithic fragments, volcanic quartz, laths of plagioclase feldspar, and volcanic zircons. Prior to this, the sedimentary rocks, which are the oldest rocks exposed on East Java, lack evidence of contemporaneous arc volcanism and are the only rocks

Cenozoic volcanic arc history of East Java, Indonesia exposed on East Java that contain no volcanic debris. This suggests that there was no arc volcanism on Java prior to 42 Ma, which also implies no subduction of the Indian-Australian plate beneath Java at this time. Growth and Development of the Southern Mountains Arc Following the initiation of volcanism in the Southern Mountains the contribution of volcanic debris increased rapidly, and the contribution of basement-derived material decreased. The volcanic centers probably formed an east-west chain of volcanic islands separated by interarc basins, much like the present IzuBonin-Mariana and Aleutian Islands Arcs. A narrow volcaniclastic shelf built up around the isolated volcanic islands, where thick sequences of volcanic and epiclastic material were deposited. To the north and directly behind the arc, the Kendeng Basin began to subside, and material was transported from the volcaniclastic shelf into the basin. No reef carbonates are preserved within Synthem Two near the arc, suggesting that environmental conditions, such as the influx of volcanic detritus, prevented reef growth or, alternatively, that these carbonates were not preserved or exposed. However, some distance to the north, on the edge of the Sunda Shelf, a large area of carbonate platform was developing at this time; the shelf received volcanic debris as air fall, but much less than that deposited close to the arc, and the volcanic material did not substantially inhibit carbonate growth. Wilson and Lokier (2002) showed that volcanic input can influence carbonate development close to active arcs without killing carbonate-producing organisms and causing carbonate deposition to cease. During the Oligocene to early Miocene, volcanic activity along the Southern Mountains Arc was at its most voluminous and explosive. The volcaniclastic shelf close to the active volcanic centers increased in width and thickness. Material was fed into the Kendeng Basin, and only in the deepest part of the basin was there pelagic sedimentation without a volcanogenic component.

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reef growth. Volcanogenic turbidites built out as thick aprons northward from the inactive arc into the Kendeng Basin. Sedimentation was rapid and on an unstable northward-dipping slope, and the deposits began to fill the accommodation space within the basin. Material also traveled southward from the eroding arc into the Java forearc. On fault-bounded highs or in the shelter of the extinct volcanoes, reefs and carbonate platforms developed, marking the first period of extensive carbonate growth within the Southern Mountains during the middle Miocene (ca. 20–10 Ma). Arc volcanism resumed in the late Miocene, after a lull of ~8 m.y., 50 km to the north of the extinct Southern Mountains Arc at the position of the modern Sunda Arc. CHARACTER OF THE CRUST BENEATH THE SOUTHERN MOUNTAINS ARC Little is known of the crust beneath many young arcs, especially in Indonesia, and East Java is no exception. However, study of the stratigraphy of the Southern Mountains Arc has provided some new and surprising insights into the character of the deep crust beneath the arc. As discussed above, exposures of basement rocks in East Java are limited and are restricted to the west of the study area. These rocks have been interpreted to be fragments of arc and ophiolitic material accreted to the continental margin of Sundaland during the Late Cretaceous (Hamilton, 1979; Wakita, 2000). It has been generally considered that these rocks are typical of the basement beneath the whole of East Java. This is supported by oil company drilling on land in East Java and farther to the north in the East Java Sea, where deep wells penetrate a varied basement, including basic and acidic volcanic rocks, slaty metasedimentary rocks, quartzites, and cherts (e.g., Hamilton, 1979; Matthews and Bransden, 1995). However, the abundance of acidic volcanism in the Southern Mountains Arc and the dating of zircons indicate that the crust beneath much of the southern part of East Java is very different, and is much older.

Termination of Volcanism in the Southern Mountains Arc

Volcanism

Following a long period of arc activity between 42 and 18 Ma, volcanism in the Southern Mountains Arc ceased. The final stages of activity were marked by a phase of explosive volcanism, which included a major event, the Semilir Eruption. The age range, thicknesses, area, and estimated volumes of volcanic deposits in the vicinity of the Toba volcanic center (Chesner and Rose, 1991) are similar to those of the Semilir center, and the deposits of the Semilir Eruption may be distributed, like the Youngest Toba Tuffs (Song et al., 2000), over large parts of SE Asia. Work is in progress to assess their distribution.

The intermediate to acidic volcanic rocks of the Southern Mountains Arc range in composition from andesite to rhyolite (60–77 wt% SiO2), with an estimated average SiO2 content of 67 wt%. In addition, many of the high-level intrusive bodies exposed within the Southern Mountains Arc are granodiorites. These rocks are considerably more evolved than the basic to intermediate eruptive products of the present-day volcanoes. The chemistry of the modern arc volcanic rocks in East Java (Handley, 2006; Nicholls et al., 1980; Wheller et al., 1987) indicates that they are the products of relatively primitive subduction melts, which reveal no significant interaction with underlying crust. The difference in chemistry of the products of the Southern Mountains Arc and the Sunda Arc could be explained in a number of ways: fractional crystallization of more basic magma, interaction of more basic magma with felsic crust, and partial melting of

Lull and Resurgence of Volcanic Activity Following the termination of volcanism, there was widespread erosion of the deposits of Synthem Two and extensive

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felsic crust. Evidence from zircons (discussed below) suggests that involvement of continental basement in petrogenesis is the most likely explanation. Ancient Zircons At the beginning of this study there were very few good dates for volcanic rocks of the older arc on Java (50 analyses), and all published ages were K-Ar dates (Ben-Avraham and Emery, 1973; Soeria-Atmadja et al., 1994). No acidic rocks had been dated. During this study, zircons from 16 Cenozoic samples of igneous (9), volcaniclastic (3), and sedimentary rocks (4) were dated by the SHRIMP U-Pb method at Curtin University of Technology, Australia, using the methods described in Smyth et al. (2005, 2007). More than 453 spot ages were measured, and 270 of these were Cretaceous and older. Prior to SHRIMP dating, the expected age range of the zircon samples was Cenozoic to Cretaceous. This reflects the age of arc activity known from Java, and a contribution from the Cretaceous arc and ophiolitic terranes interpreted to form the basement of East Java. However, in addition to the expected ages, an unexpected range of Archean–Cambrian grains was identified in a large number of samples (Smyth, 2005, 2007). The samples containing Cretaceous ages and those containing Cambrian and older zircons occur in distinct areas of East Java. Lithology and Zircon Age Range The five intrusive and four extrusive igneous rocks analyzed contain a range of Cambrian and older zircons (n = 155) but lack Cretaceous zircons. In contrast, the three volcaniclastic rocks analyzed yielded a significant number of Cretaceous zircons (n = 42) but only one Cambrian or older zircon. The four quartzrich sandstones analyzed yielded varied zircon ages. All of these sandstones contain Cretaceous grains (n = 33). Three sandstones from the Southern Mountains and Kendeng Basin contain a range of Cambrian and older grains (n = 22). One sandstone from the edge of the Sunda Shelf contains a single Jurassic–Triassic grain but no older grains. Age Ranges and Spatial Distribution The samples containing different zircon ages do have a clear geographical pattern. The Archean grains are found only in rocks of the Southern Mountains (Fig. 12B), and the igneous rocks of this area contain no Cretaceous zircons (Fig. 12A). Cretaceous zircons are found only in the north and west of East Java, and these zircons have been identified only in sedimentary and volcaniclastic rocks. This is interpreted to indicate that the Southern Mountains are underlain by material providing Cambrian to Archean zircons but no Cretaceous zircons. To the north and west of the Southern Mountains the underlying rocks do not include Cambrian to Archean material but do include Cretaceous material. Figure 12C shows the distribution of basement rocks interpreted from the zircon ages.

If all the zircon ages are grouped together, five age peaks can be identified (Fig. 13): a Cenozoic peak (5–42 Ma) recording volcanism in the Southern Mountains and Sunda Arcs, a Cretaceous peak (65–135 Ma) consistent with the interpreted age of basement, a Cambrian to Neoproterozoic peak (500–1000 Ma), a Mesoproterozoic to Paleoproterozoic peak (1000–2500 Ma), and an Archean peak (2500–3200 Ma). The Cambrian and older ages were not expected, based on knowledge of the regional geology. Source of the Ancient Zircons Sundaland, the continental core of SE Asia, is the closest and most obvious source of the very old zircons (Figs. 1 and 13). Rocks that could have provided abundant old zircons include granites of SW Borneo (Hamilton, 1979), the Malay Peninsula (Liew and Page, 1985), the Thai-Malay Tin Belt (Cobbing et al., 1986), and Sumatra (Imtihanah, 2000; McCourt et al., 1996) (Fig. 13). However, in these areas there are no known Archean rocks, nor is there any indication that they are underlain by Archean crust. Geochemical and isotopic studies suggest a basement no older than Proterozoic in areas of Sundaland, such as the Thai-Malay peninsula, that have been studied (e.g., Liew and Page, 1985). The closest area with extensive granites is the Schwaner Mountains of SW Borneo. Paleogene sedimentary rocks of north Borneo contain debris eroded mainly from the Schwaner granites and the Malay Tin Belt, including detrital zircons (van Hattum et al., 2006). The zircon populations are dominated by Cretaceous zircons with age peaks different from those of East Java; there are abundant Permian–Triassic zircons, rare in East Java; and Precambrian zircons are mainly Paleoproterozoic with very rare Archean grains. The differences in zircon ages rule out a Schwaner or Thai-Malay provenance, and other granite sources in Sundaland are even more distant and paleogeographically unlikely. The largest and closest area of continental crust of Archean age is Australia, which formed part of Gondwana until the Cretaceous. It has been suggested that small Gondwana continental fragments collided with the east Sundaland margin in Sulawesi during the Cretaceous (e.g., Wakita et al., 1996; van Leeuwen and Muhardjo, 2005), although the source of the fragments was not identified. There is evidence of Late Jurassic–Early Cretaceous rifting of continental fragments from the northwest and west Australian margins (Müller et al., 2000) during continental breakup preceding the separation of India from Gondwana. One of these continental fragments could have collided with the Sundaland margin. Bergman et al. (1996) speculated that there could be old continental crust beneath west Sulawesi, on the basis of lead isotopic compositions of Neogene plutonic rocks, which were suggested to have had an Australian origin. Basement rocks in western Australia are dated as Proterozoic and Archean. Detrital zircons from young sediments of the Perth Basin, derived from the erosion of western Australian basement, have been well studied and dated (Cawood and Nemchin, 2000; Pell et al., 1997; Sircombe and Freeman, 1999). The zircon populations in Perth

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Basin sedimentary rocks are remarkably similar in age to those of the Southern Mountains in East Java. Both areas yield zircons with 4200–2500 Ma ages (predominantly 3200–2500 Ma) similar to the Archean Yilgarn and Pilbara Blocks, 2000–1600 Ma and 1300–1000 Ma ages similar to the Paleoproterozoic and

Mesoproterozoic Capricorn and Albany-Fraser orogenic belts, and 800–500 Ma ages similar to the Neoproterozoic Pinjara orogenic belt (Fig. 14). A west Australian origin is a geographically simple and plausible explanation for the Precambrian zircons of East Java, but how did they become incorporated in igneous rocks?

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Figure 13. Extent of area interpreted to be underlain by continental crust, and locations of granitic bodies that may have been available for erosion during the Cenozoic. The NE-SW–trending Karimunjawa Arch, which was a significant topographic barrier throughout the Cenozoic, is also denoted.

Incorporation of Zircons into East Java Igneous Rocks Zircons are highly refractory and are well known for surviving repeated episodes of melting (e.g., Hanchar and Hoskin, 2003). Ancient zircons could have been introduced into Paleogene magmas of the Southern Mountains Arc either by subduction of Gondwana-derived sediments or crust, or by passage of magmas through a previously emplaced Gondwana-origin continental fragment at depth beneath the Southern Mountains. Subduction of Gondwana Sediment Sediment from western Australia deposited on the Indo-Australian plate could have been carried north to the Java Trench as Australia moved north. However, even today a fan similar in size to the modern Bengal Fan would be required to bring material from Australia to the site of subduction at the Java Trench, and there is no evidence of a sediment supply and distribution system like this in the past. The huge Bengal Fan is fed by erosion from the Himalayas, but there is no evidence for a comparable orogenic belt in Australia during Cenozoic times. During the early Cenozoic, Australia lay much farther south of its current position (Hall, 2002), and there is little sedimentary cover today on the IndianAustralian plate south of Java (Masson et al., 1990; Kopp, 2002) in what would have been more proximal parts of such a fan. It is unlikely that material could have been introduced into the Java Trench by axial transport from the Himalayas

via the Bengal Fan, or from the Bird’s Head microcontinent, New Guinea. If sediment was produced by the erosion of the Himalayas, it would require transport via the Bengal Fan to the Java Trench and total transport distances >4000 km. Although mud-rich material is considered to have reached northern Sumatra in the distal parts of the Bengal Fan (Curray et al., 2002), there were topographic barriers at the Investigator Fracture Zone and the Ninety East Ridge, and even then a further 2000 km of along-trench transport. Today there is no material from the Bengal Fan reported in the Java Trench off the shore of East Java (e.g., Masson et al., 1990) and no significant sediment cover on the Indian plate south of Java (Kopp, 2002). It is equally unlikely that sediment from the Himalayan region could have traveled southeastward from Indochina across Sundaland, bypassing numerous basins and associated structural highs (Hall and Morley, 2004) to enter East Java. This would require transport distances >3500 km. The basement rocks of the Bird’s Head microcontinent were not available for erosion during most of the Cenozoic, and the Bird’s Head was the site of deposition of thick marine carbonates of the New Guinea limestones (Fraser et al., 1993; Pieters et al., 1983). Even had the Bird’s Head been supplying large volumes of sediment, along-trench transport distances >2000 km would have been required for material to enter the subduction zone beneath Java. This is equivalent to material from southern Mexico or northern Washington State entering the Allistos Arc

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Range of U-Pb SHRIMP zircon ages from the Southern Mountains Arc, East Java (Smyth, 2005). N = 453

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Figure 14. Comparison of the U-Pb SHRIMP ages of (A) East Java (Smyth, 2005), and (B) Western Australia (Bruguier et al., 1999; Sircombe and Freeman, 1999). Revised from Smyth et al. (2005).

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(Busby et al., 2006) of Baja California, or for northern Norway or Alabama to have been sources for the Lough Nafooey Arc (Draut and Clift, 2001) of the Irish Caledonides. Subduction of a Gondwana Fragment Subduction of continental crust is even less likely. Buoyancy forces make subduction of continental crust difficult, and

even if these were overcome, the fragment would have been required to supply material to the arc from the middle Eocene to the early Miocene, a period of ~20 m.y. The fragment of crust either remained stationary beneath the Southern Mountains or was 1500 km in length, based on the present subduction rate of 75 mm/year (McCaffrey, 1996), or on Cenozoic plate tectonic reconstructions (Hall, 2002).

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Accretion of a Gondwana Fragment The most likely explanation is that continental crust was already present beneath the Southern Mountains Arc when subduction began in the middle Eocene. We suggest that this was a western Australia–rifted continental fragment that collided with the margin of Sundaland during the Late Cretaceous and may have terminated the Cretaceous phase of subduction. Subduction was renewed in the middle Eocene, and Cenozoic melts were contaminated by interaction with continental crust beneath the Southern Mountains. The distribution of pre-Cenozoic zircons indicates that north of the Southern Mountains there is no continental basement, and the deep crust is arc or ophiolitic, perhaps representing material accreted during Cretaceous subduction or underplated during collision. The continental fragment could have been quite large and may be traceable into Sulawesi. Zircon dating and geochemical evidence (van Leeuwen and Muhardjo, 2005) show that the Malino Metamorphic Complex of NW Sulawesi (Fig. 13) contains zircons with ages up to 3500 Ma. DISCUSSION The stratigraphic record in East Java provides insights into the nature of the deep crust beneath the arc, the history of IndianAustralian plate subduction, and the contribution of the volcanic arc to basin formation. It also poses some questions about the continuity and location of volcanic activity, which have relevance for other arcs. Deep Character of the East Java Volcanic Arc Little is known about the deep crust beneath the young arcs of the Western Pacific and Indonesia. There have been few seismic refraction studies and few xenolith studies, and the deep crust is almost invariably covered by younger volcanic and sedimentary rocks. In East Java the old zircons in the Eocene to Miocene arc products provide evidence of the character of the deep crust. The insight gained was completely unexpected. When this study began, it was thought that the entire region was underlain by Cretaceous arc and ophiolitic fragments, as suggested by Hamilton (1979). The distribution of ages shows that north of the Southern Mountains there is indeed crust of this character, as indicated by the small exposures of basement, but beneath the Southern Mountains themselves there is a fragment of continental crust of Gondwana character and western Australian origin. The presence of a continental fragment beneath the arc accounts for the unusually acidic character of arc volcanism in the Eocene to early Miocene arc. The acidic products of this arc have been overlooked partly because the resistant “Old Andesites” are topographically so much more obvious, although acidic volcanic and minor intrusive rocks are well exposed in many parts of the Southern Mountains. However, the main reason why the acidic products of the Eocene to early Miocene arc were missed by earlier studies is because they were erupted explosively and dispersed as volcanic ash, the ash was reworked into sediments,

and the processes of eruption and tropical weathering combined to remove the unstable volcanic constituents, leaving well-sorted quartz-rich sandstones. These were previously interpreted as having been eroded from a continental Sundaland source, but careful examination of the sandstones reveals abundant evidence of their volcanic origin on the basis of textures, light mineral constituents, quartz character, clay mineralogy, and zircon character and ages (Smyth, 2005). Age of Subduction The character of the oldest parts of the East Java stratigraphic successions indicates that Cenozoic volcanic activity began in the middle Eocene (ca. 45 Ma). There is little evidence to support the common assumption (e.g., Hall, 2002; Heine et al., 2004; Metcalfe, 1996; Scotese et al., 1988) that subduction continued from the Mesozoic into the Cenozoic without significant interruption. The stratigraphy of East Java indicates that older subduction probably terminated in the Late Cretaceous with the collision of a continental fragment rifted from western Australia. We suggest that the continental fragment extended from East Java to North Sulawesi. Ophiolites were emplaced from Java to North Borneo during this collision. The oldest sedimentary rocks above the basement in East Java are of uncertain age but are middle Eocene or older and lack volcanic debris. There is little evidence for latest Cretaceous to early Eocene volcanic activity in most of the Sundaland margin between Sumatra and Sulawesi. Subduction resumed in the middle Eocene when Australia began to move northward rapidly after 45 Ma (Hall, 2002; Müller et al., 2000; Schellart et al., 2006), and has continued to the present day. Formation of the Kendeng Basin When volcanic activity resumed in the middle Eocene a basin began to form directly north of the arc. Most of the investigations of sedimentary basins on land in East Java and in the East Java Sea (Fig. 1) have been carried out as part of hydrocarbon exploration activity, and little attention has been given to areas close to the modern arc. The Sunda Shelf basins are typically >100 km from the arc and are characterized by many features typical of fault-controlled basins; explanations for their origin have therefore interpreted them as types of backarc basins resulting from subduction rollback, rift and thermal sag, or strike-slip faulting (e.g., Hamilton, 1979; Matthews and Bransden, 1995). The Kendeng Basin is not a typical backarc basin. Backarc basins (Taylor and Karner, 1983) may form by trench rollback, causing generation of oceanic crust (Dewey, 1980; Karig, 1971), rifting of continental crust (Kobayashi, 1985), gravitational collapse in the wake of arc-continent collision (Clift et al., 2003), or by the trapping of old oceanic crust behind an arc (Scholl et al., 1986). Retroarc basins (Dickinson, 1974) form as compressive foreland basins behind a volcanic arc. Busby and Ingersoll (1995) defined backarc basins as oceanic basins behind intraoceanic magmatic arcs, and continental basins behind continental

Cenozoic volcanic arc history of East Java, Indonesia margin arcs that lack foreland fold and thrust belts. Marsaglia (1995) identified similar categories of backarc basins with the addition of boundary basins resulting from extension along plate boundaries with translational components. Busby and Ingersoll (1995) claim that the Sunda Shelf basins of Indonesia are nonextensional backarc basins formed in a neutral strain regime, although the basis for this assertion is not clear, because there is abundant evidence for rifting of many of these basins (e.g., Cole and Crittenden, 1997; Hall and Morley, 2004). There is no evidence for either newly formed oceanic crust under the Kendeng Basin or for strike-slip faulting. The basin is much closer to the arc than other backarc basins and most of the “backarc basins” of the Sunda Shelf. The late Cenozoic deformation in the region led to ~10–30 km of contraction at the outer edge of the basin, so throughout the early Cenozoic this basin was very close to the arc. There was extension in the Sunda Shelf from the Eocene, but nowhere was oceanic crust formed. The Kendeng Basin is an asymmetrical depression, the deepest part of the basin lies directly behind the arc, and the Eocene to lower Miocene volcanic and sedimentary sequence thins toward the edge of the Sunda Shelf (Waltham et al., this volume). The basin fill was derived mainly from the Southern Mountains Arc on the south side of the basin. The basin began to form at the same time as arc activity began, and its subsidence history is closely linked to activity in the volcanic arc. There are no seismic lines crossing the basin, and it is not well exposed at the surface, so we cannot assess the role of faulting in the basin development, but it does not have the typical synrift-postrift stratigraphy of many other Sunda Shelf basins. There is no evidence of compressional loading having contributed to basin formation. Thrusting in East Java occurred in the late Neogene long after the Kendeng Basin had formed. The close relationship between volcanic activity and basin subsidence suggests that the two are linked, and we suggest that the load of the volcanic arc was the major cause of basin subsidence. The contribution of volcanic arc loading to basin formation in Indonesia is discussed in greater detail by Waltham et al. (this volume). Movement of the Volcanic Arc Activity in the Southern Mountains Arc terminated during the early Miocene (ca. 20 Ma). After a period with little magmatism, a new episode of arc volcanism began in East Java at ca. 12–10 Ma in a new location. The Miocene to Holocene arc is parallel to the Southern Mountains Arc but ~50 km north of it. A significant reduction in volcanic activity in the middle Miocene is a well-known feature of the Sunda Arc from Java eastward, although subduction was continuous during this period (Hall, 2002). Vigorous volcanic activity resumed at the end of the middle Miocene in the Sunda and Banda Arcs at ca. 10 Ma (Hall, 2002; Macpherson and Hall, 2002). Macpherson and Hall (1999, 2002) suggest that the decline in volcanic activity resulted from northward advance of the subduction hinge. Hinge advance prevented replenishment of the mantle wedge by fertile mantle, and consequently subduction-induced

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melting was inhibited as the wedge became depleted. The cause of hinge advance was the result of the collision of Australia with eastern Indonesia, and the counterclockwise rotation of the Borneo-Java region this collision induced (Hall, 2002). Arc volcanism resumed when hinge advance ceased and the mantle wedge was replenished by fertile mantle (Macpherson and Hall, 2002). Why the volcanic arc moved to its new position 50 km north of the Southern Mountains Arc is uncertain. There is evidence of thrusting and contraction in Java in the late Neogene, but its timing is not precisely known; this is part of our current research. It is possible that contraction was linked to large-scale arc dynamics, such as coupling of the subduction and overriding slab. There is evidence of an old consolidated accreted material in the Java forearc that acts as a backstop to the active accretionary prism (Kopp et al., 2001). This suggests two phases of forearc development, and these may be linked to the two phases of arc volcanism on land in East Java. It is also possible that when the subduction hinge advanced and the mantle wedge was not being replenished, one result was a rigid, stronger overriding lithospheric plate that later failed when hinge advance ceased and weaker, warmer mantle replenished the wedge. Another control could have been the nature of the crust beneath the arc. The zircon ages indicate that there was a change in basement type north of the Southern Mountains and that the boundary between continental and ophiolitic crust at depth may have been a preexisting structural weakness that influenced the position at which melts could rise when arc activity resumed. Other possibilities for explaining arc migration include subduction erosion and a change in dip of the subducting slab. Subduction erosion seems unlikely, as the width of the arc-trench gap today is >300 km, and because marine data support active accretion at that time (e.g., Kopp et al., 2001). A change in subduction angle cannot be ruled out, but it is notable that today the slab dip has increased to a very high angle after the slab descended to 100 km, as seen in seismicity data (England et al., 2004). A lower, not a higher, angle would be expected if the arc had moved north owing to the change in dip of the slab. The jump in position of the volcanic arc is a feature of other Indonesian arcs (e.g., Halmahera Arc; Nichols and Hall, 1991) and reflects the changing stresses at the plate boundary over time. We prefer an explanation that links the change in position of the arc to plate reorganization in the region. However, it is difficult to test different hypotheses. Comparison of arc development to the east and west of East Java, from Bali eastward, and toward West Java and Sumatra, might provide insights, but unfortunately even less is known of arc history in these regions than in East Java. This East Java study shows that our understanding of arc tectonics is still incomplete and that studies of arc stratigraphy are essential if our knowledge of arc dynamics is to be improved. CONCLUSIONS The case study of early Cenozoic arc volcanism in East Java shows the importance of examining arc stratigraphy. Insights

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have been gained into the eruptive history, formation of a deep basin behind the arc, and the character of the deep crust. The early Cenozoic stratigraphy of East Java provides a record of a cycle of arc activity from initiation in the middle Eocene (ca. 42 Ma) to termination in the early Miocene (ca. 20 Ma). The Kendeng Basin, directly behind the Southern Mountains Arc, contains >6 km of volcaniclastic and sedimentary rocks. The basin is not a typical backarc basin, and its subsidence history is linked to volcanic activity within the Southern Mountains Arc. The final stage of volcanic activity in the Southern Mountains Arc is marked by the Semilir Eruption (ca. 20 Ma), which distributed ash over a wide area and may be comparable to the Pleistocene eruption of Toba in Sumatra. Following this phase of major eruptions, there was a lull in volcanic activity during the middle Miocene, followed by resumption in arc activity to the north of the Southern Mountains Arc, along the axis of the modern Sunda Arc during the late Miocene. The mechanisms that resulted in the decline in volcanism, and northward movement in arc axis, are not yet understood and show that our understanding of arcs is still incomplete. The stratigraphic record of volcanic arcs can provide insights into the character of the deep crust. The entire East Java region was previously thought to be underlain by Cretaceous arc and ophiolitic fragments, but Archean to Cambrian zircons within acidic products of the Southern Mountains Arc point to the occurrence of a continental crust of Gondwanan character and western Australian origin beneath the old arc. This continental fragment is thought to have collided with Sundaland during the Cretaceous and is interpreted to have terminated the Cretaceous phase of subduction. The extent of this fragment is not known but may be traceable into Sulawesi. The use of inherited zircons to determine the character of the deep crust may be applicable in other arcs. ACKNOWLEDGMENTS This project was funded by the SE Asia Research Group at Royal Holloway University, supported by an oil company consortium. Zircons were analyzed using the SHRIMP facility at Curtin University of Technology, Perth, Australia. Financial assistance for SHRIMP dating was provided by the University of London Central Research Fund and CSIRO, Australia. The analyses were undertaken by Joseph Hamilton (CSIRO and University of the West Indies) and Pete Kinny (Curtin University of Technology). We are grateful to LIPI (Lembaga Ilmu Pengetahuan Indonesia) for field-work permissions. We wish to thank Eko Budi Lelono, Peter Lunt, Theo van Leeuwen, Moyra Wilson, Colin Macpherson, Heather Handley, Cindy Ebinger, and Dave Waltham. Marcelle BouDagher-Fadel of University College London provided biostratigraphic analyses. Peter Clift, Jason Ali, and an anonymous reviewer are thanked for their extensive reviews, which greatly improved this manuscript.

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MANUSCRIPT ACCEPTED BY THE SOCIETY 24 APRIL 2007

Printed in the USA

The Geological Society of America Special Paper 436 2008

New constraints on the sedimentation and uplift history of the Andaman-Nicobar accretionary prism, South Andaman Island R. Allen Department of Environmental Science, Lancaster University LA1 4YQ, UK A. Carter Research School of Earth Sciences, Birkbeck and University College London, Gower St., London WC1E 6BT, UK Y. Najman Department of Environmental Science, Lancaster University LA1 4YQ, UK P.C. Bandopadhyay Geological Survey of India, Geodata Division, Salt Lake, Kolkata, 91, India H.J. Chapman M.J. Bickle Department of Earth Sciences, Cambridge University, Downing St., Cambridge CB2 3EQ, UK E. Garzanti G. Vezzoli S. Andò Dipartimento di Scienze Geologiche e Geotecnologie, Universita Milano-Bicocca, Piazza della Scienza 4, 20126 Milano, Italy G.L. Foster Department of Earth Sciences, Bristol University, Queens Rd., Bristol BS8 1RJ, UK C. Gerring Department of Earth Sciences, The Open University, Walton Hall, Milton Keynes MK7 6AA, UK

ABSTRACT The Andaman Islands are part of the Andaman-Nicobar Ridge, an accretionary complex that forms part of the outer-arc ridge of the Sunda subduction zone. The Tertiary rocks exposed on the Andaman Islands preserve a record of the tectonic evolution of the surrounding region, including the evolution and closure of the Tethys Ocean. Some of the Paleogene sediments on Andaman may represent an offscraped part of the early Bengal Fan. Through field and petrographic observations, and use of a number of isotopic tracers, new age and provenance constraints are placed on the

Allen, R., Carter, A., Najman, Y., Bandopadhyay, P.C., Chapman, H.J., Bickle, M.J., Garzanti, E., Vezzoli, G., Andò, S., Foster, G.L., and Gerring, C., 2008, New constraints on the sedimentation and uplift history of the Andaman-Nicobar accretionary prism, South Andaman Island, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 223–255, doi: 10.1130/2008.2436(11). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Allen et al. key Paleogene formations exposed on South Andaman. A paucity of biostratigraphic data poorly define sediment depositional ages. Constraints on timing of deposition obtained by dating detrital minerals for the Mithakhari Group indicate sedimentation after 60 Ma, possibly younger than 40 Ma. A better constraint is obtained for the Andaman Flysch Formation, which was deposited between 30 and 20 Ma, based on Ar-Ar ages of the youngest detrital muscovites at ca. 30 Ma and thermal history modeling of apatite fission-track and U-Th/He data. The latter record sediment burial and inversion (uplift) at ca. 20 Ma. In terms of sediment sources the Mithakhari Group shows a predominantly arc-derived composition, with a very subordinate contribution from the continental margin to the east of the arc. The Oligocene Andaman Flysch at Corbyn’s Cove is dominated by recycled orogenic sources, but it also contains a subordinate arc-derived contribution. It is likely that the sources of the Andaman Flysch included rocks from Myanmar affected by India-Asia collision. Any contribution of material from the nascent Himalayas must have been minor. Nd isotope data discount any major input from cratonic Greater India sources. Keywords: Andaman, accretionary wedge, arc, subduction, thermochronology, provenance, uplift.

INTRODUCTION The Andaman-Nicobar Islands are part of an accretionary complex that forms the outer arc ridge of the northern Sunda subduction zone (Fig. 1). The Andaman Islands are in the southeastern part of the Bay of Bengal and make up part of a 3000– 5000 km chain that runs from the Myanmar Arakan-Yoma down to Sumatra and Java in the south. The Indian plate is subducting northward below the Euarasian plate and obliquely below the Sino-Burman plate along the Burma-Andaman-Java Trench. The structure of the Andaman Islands comprises an accretionary prism formed by an imbricate stack of east-dipping fault slices and folds that young to the west (Fig. 2), linked to a westward-shifting subduction zone (e.g., Roy, 1992; Pal et al., 2003). The geology of an accretionary wedge is complex, reflecting its dynamic environments and involving subduction, folding, and thrusting. Depositional ages and environments can change abruptly over relatively short distances, and uplift leads to recycling of sediment from the eroding wedge. Throughout subduction, new material introduced at the bottom of the accretionary wedge is accreted, uplifted, or subducted. Some of the accreted material may be uplifted and brought to the seafloor. Slope basins may develop behind folds in the accreted sedimentary rocks and trap sediment in a deep-water environment. Simultaneously, shallow-water sediments such as reefs can form on the prism top and be eroded and transported down the slope. Given this inherent complexity, many accretionary complex rocks are referred to as mélange; thus unraveling the sedimentation history in a subduction-accretionary setting is a major task. Interpretation of the geology of the Andaman Islands is hindered by the lack of isotopic age constraints, limited biostratigraphy, and poor outcrop exposure (Bandopadhyay and Ghosh, 1999, Bandopadhyay, 2005; Pal et al., 2003, 2005). The

aim of this paper is to build on previous field-based studies and apply petrologic, isotopic, and thermochronometric techniques to better understand the provenance, sediment deposition, and uplift history. The origin of the Andaman Flysch has been debated for >20 yr. It has been variously proposed that the Andaman Flysch was derived from the Irrawaddy Delta (Karunakaran et al., 1968; Pal et al., 2003) or, alternatively, from Bengal Fan material shed from the nascent Himalaya sourced either directly or by emplacement as an allochthon into the accretionary prism by oblique subduction (Curray et al., 1979; Curray, 2005). We are particularly interested to determine whether the Himalayan-Tibetan orogen contributed sediment to the Andaman Islands, because this might reveal information on the early evolution of the orogen not preserved elsewhere. LITHOLOGIES Previous Work The current stratigraphy (Table 1) of the Andaman Islands is based on lithological mapping and can be traced back to the pioneering work of Oldham (1885), who first divided the Andaman geology into an older Port Blair Series and a younger Archipelago Series, separated by volcanic rocks and serpentinites later recognized as an ophiolite. Over the past 50 yr the stratigraphy has been modified and formation names changed, but it was not until the 1960s that paleontological constraints were used to place the Paleogene–Neogene lithostratigraphic units within a temporal framework (Guha and Mohan, 1965; Karunakaran et al., 1968). The stratigraphy now comprises four units, which, in ascending order, are Cretaceous sedimentary rocks and ophiolite, the Eocene Baratang-Mithakhari Group,

Sedimentation and uplift history of the Andaman-Nicobar accretionary prism

Shillong

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drilled to the east and west of the islands (Fig. 2) that helped place the exposed geology within the context of the accretionary setting. Below we briefly describe the main rock units exposed on the Andaman Islands and report key observations from previous field and petrographic studies.

100 es g n Ra Mogok Belt

Myanmar Shan Plateau Central Basin

I nd

Bengal Basin

obu rm a

90 Himalaya

Bengal Fan

20

Thailand

10 lay Ma la su nin Pe

idge an Nicobar R

Andam

Andaman Sea Mergui basin

tra ma Su

0

95 E Myanmar Bengal Fan

Irrawaddy delta

Andaman Sea

South Andaman

v v vv vv vv

nic

Richie’s Archipelago & Havelock

arc

v

v vv v

Volc a

Andam an - Nic

obar Tr e

nch

15 N

North Andaman

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Figure 1. General location map of Andaman Islands and location of potential source regions. Lower figure corresponds to boxed area in top figure.

the Eocene to upper Oligocene Port Blair–Andaman Flysch Group, and the lower to upper Miocene Archipelago Group. In the mid-1970s the Indian state Oil and Natural Gas Commission (ONGC) carried out a detailed seismic reflection study across the Andaman Islands, calibrated against offshore boreholes

Ophiolite The Andaman ophiolite contains the main components of an ophiolite sequence that includes upper mantle–depleted harzburgites and dunites, lower crustal cumulate gabbros and peridotites, and upper crustal sheeted dikes, pillow lavas, and marine pelagic sediments (Halder, 1985; Ray et al., 1988; Roy, 1992). However, the sequence is tectonically disturbed, and much of the crustal section is deformed and difficult to identify in the field. Pillow lavas are abundant, but sheeted dikes have been identified only in disrupted small-scale faulting and folding. Both massive and layered gabbros are recognized in the preserved ophiolite. South Andaman has the best preserved and most complete sequence of ophiolite, which extends for ~30 km from Corbyn’s Cove in the north to Chiriyatapu in the south (Fig. 3). Pelagic Sedimentary Rocks The topmost part of the ophiolite complex contains thin and discontinuous lenses and streaks, and laterally continuous (at outcrop scale) bedded sequences of pelagic sedimentary rocks consisting of jasper, chert, cherty limestone, and shales (Bandopadhyay and Ghosh, 1999). Outcrops commonly show evidence of significant deformation and folding (Fig. 4). Bedded Chert Rhythmic alternations of centimeter-thick, milky white chert and millimeter- to centimeter-thick reddish-brown and purple shale-mudstone beds constitute the bedded chert facies. Chert and shale normally show uniform (0.5–4.0-cm-thick) beds that have sharp bases and tops and planar contacts. Some 10–15-cm-thick beds, and massive beds, of chert are present. Soft-sediment deformation is evident in some localities. Radiolarians are preserved to varying degrees in most cherts and indicate a Late Cretaceous to Paleocene depositional age, which constrains the underlying ophiolite sequence to a Late Cretaceous age. Shale At outcrop the shale facies form interbedded sequences of extremely variable thickness and lateral continuity. Basaltic volcanic rocks occur as thin intercalations, conformable lenses, and, at places, small crosscutting dikes. Thin beds of fine-grained sandstone, siltstone, and cherty limestone are also present. There is evidence of soft-sediment deformation and slump folds as well as cutting by normal and thrust faults. Some shales are clearly tuffaceous with plagioclase phenocrysts, vitric fragments, pumice clasts, and diagenetically altered volcanic lithic fragments. Chlorite is abundant in the matrix of the altered tuff and tuffaceous shale. Occasionally sharp edged, cuspate or platy, fresh glass shards can be found (Bandopadhyay and Ghosh, 1999).

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A

Andaman - Nicobar

West

A' East

50 km

123-

Andaman - Nicobar Ridge

B

Middle to South Andaman

B

B'

0-

N

depth (m)

4-

Sund a Tre nch

two-way time (sec)

0-

100020003000-

B'

25 Km A'

Melange/ deformed pelagic cover

Upper Cretaceous

A

Neogene Paleogene

Ophiolite

Archipelago Group Andaman Flysch Mithakhari Group

Figure 2. Cross section through Andaman Ridge, based on seismic sections from Roy (1992) and Curray (2005) that show the underlying structure of the accretionary wedge, comprising a series of folded east-dipping thrust slices.

Approximate depositional age range Miocene to Pliocene

Oligocene–late Eocene(?)

TABLE 1. SIMPLIFIED STRATIGRAPHY OF THE ANDAMAN ISLANDS Group Formation

Cross-stratified and graded sandstones, silty mudstones and limestones marls, and chalky limestones

Archipelago Group Andaman Flysch Group (formerly Port Blair Group)

Bouma sequences, sandstone-shale and mudstones

Namunagarh Grit Early to middle Eocene(?)

Mithakhari Group (formerly Baratang and Port Meadow Groups)

Hope Town Conglomerate Lipa Black Shale

Late Cretaceous to Paleocene(?)

Lithology

Ophiolite Group

Mithakhari Group The Mithakhari Group consists of immature gravels and coarse- to fine-grained sandstones, pebbly to fine-grained pyroclastic sandstones, and minor thin beds of mudstones and coal. Whereas the Mithakhari Group dominates the outcrop geology of the Andaman Islands, particularly in North and Middle Andaman,

Pebbly and coarse to fine-grained volcaniclastic sandstones and grits Interstratified massive and graded polymict conglomerates, massive cross-stratified and graded sandstones, shales, and thin coals Pyritiferous black shale Pillow lava, basalt, gabbro, pyroxinite, harzburgite, serpentinite, andesite, diorite, plagiogranite, rhyolite, serpentinized harzburgite, pyroxinite, and pelagic sediments; radiolarian chert and hematitic mudstones

a paucity of good exposures and poor access make it difficult to obtain continuous sections, and exposures are limited to isolated stone quarries, coastal areas, and road cuts. In South Andaman the Mithakhari Group occurs as a north-south–trending outcrop that extends for ~50 km, but the best sections, exposing the least weathered outcrops, are found only near Hope Town, Mungleton, Namunagarh, and Chiriyatapu (Fig. 3). Karunakaran et al. (1968)

Sedimentation and uplift history of the Andaman-Nicobar accretionary prism

Hope Town

Mithakhari Group

South Andaman

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Port Blair

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Figure 3. General location map of Andaman Islands and local geology (A, B) for the areas on South Andaman Island sampled for this study.

Ophiolite 3km

B) Andaman Andaman Flysch Flysch

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Ophiolite

Mithakhari Group Chiriyatapu

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Hope Town Conglomerates This unit is best seen near Hope Town on South Andaman Island, where ~6 m of conglomerates and pebbly sandstones are well exposed (Fig. 5), interbedded with thin beds of greenish gray coarse- and fine-grained sandstones. The sequence shows fining- and thinning-upward sequences with evidence of slumping and soft sediment deformation. Bed contacts are generally sharp and planar, and some evidence of fluvial channels can be found. Conglomerates are polymict, of mainly basic-ultrabasic sources, and with subordinate to minor amounts of andesite, sedimentary limestone, and cherts plus sporadic mudstone clasts and metamorphic quartz. Figure 4. Folded cherts and shales that represent the pelagic cover to the Cretaceous ophiolite exposed on the shore at Chiriyatapu, South Andaman Island.

first introduced the term Mithakhari Group, dividing the group into a lower Lipa Black Shale, a middle Hope Town Conglomerate, and an upper Namunagarh Grit Formation (Table 1). The Lipa Black Shale is a minor unit and not well exposed, and so will not be considered further.

Namunagarh Grit Formation This unit is characterized by coarse- to fine-grained sandstones and siltstone, with minor conglomerate at the base. On South Andaman the type section and best exposures are found in quarry sections near Namunagarh village (Fig. 3). These display 3–5-m-thick, green, matrix-supported sandstones. The sandstones are well bedded and laterally persistent along the quarry sections, and consist of coarse- and fine-grained beds. The coarse-grained beds, at the base of the section, are ≥1 m thick, with sharp nonerosive contacts. The finer grained beds consist of 4–8-cm-thick

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A)

1m

B)

10cm Figure 5. Massive polymict conglomerates exposed at Hope Town Quarry (A), South Andaman Island. These polymict conglomerates are dominated by basic and ophiolitic clasts (B) whereas the matrix contains abundant volcanic glass and felsitic volcanic grains, suggesting a dominant volcanic arc source, although zircon fission-track data also show evidence of continental Mesozoic sources.

beds of fine- to medium-grained volcaniclastic sandstones interbedded with thin mudstones. Some large fragments of volcanic rock fragments, including elongated pumice lapilli, are present, which resemble floating clasts in turbidites. For a long time the sandstones exposed at Namunagarh stone quarries had been described as graywacke formed from weathering and erosion of accreted ophiolite (Acharyya et al., 1989). Recently Bandopadhyay (2005) identified beds with abundant pyroclasts, including vesiculated glass fragments, pumice clasts and shards, euhedral feldspars, and angular lithic fragments diagnostic of tuff, indicating that some of the Namunagarh Grit beds were derived from direct volcanic arc sources (Fig. 6). The polymict conglomerates (Hope Town Conglomerate) and grits (Namunagarh Grit Formation) are interpreted as having

Figure 6. Namunagarh Quarry. Thin sections show that the sediments are largely tuffaceous with clear evidence of fresh arc volcanic material including devitrified glass.

been derived from the ophiolite and its pelagic–shallow marine cover. The succession, which includes thin coals and gypsum, was deposited in a delta-slope setting with facies associations ranging from subaerial alluvial plain to prodelta slope (Chakraborty et al., 1999). Its depositional age is not well defined owing to the lack of distinct biostratigraphic evidence, but shallow benthic foraminifers in the Hope Town Conglomerate, including Nummulites atacicus, constrain the age from late Ypresian to early Lutetian (Karunakaran et al., 1968). Many of the foraminifers, however, are broken and abraded (i.e., reworked). The relationship of the Namunagarh grits to the Hope Town conglomerates is not clear, although the Namunagarh grits are presumed to be younger. Andaman Flysch The Andaman Flysch is a siliciclastic turbidite sequence deposited on a submarine fan. It is bounded between the Mithakhari Group below and the Archipelago Group above. The (misleading) term flysch is derived from the resemblance of the turbidites to the classic Bouma turbidites described in the Swiss Alps. Similar looking beds are seen throughout the Andaman Islands; hence the Andaman Flysch is described as cropping out over a N-S strike length of 200 km from the southern part of South Andaman to the northern tip of North Andaman. The overall thickness is not well defined, with estimates varying from 750 m (Roy, 1983) to 3000 m (Pal et al., 2003). The best and most completely documented exposures are found on South Andaman at Corbyn’s Cove (Fig. 3), where outcrops of steep, westerly dipping beds are seen adjacent to the pillow basalt of the ophiolite sequence (Fig. 7), although the nature of the contact is uncertain. Individual sandstone beds can be traced along strike for distances of several kilometers, but the total thickness is only 250–300 m. Current directional structures in sandstone beds

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METHODOLOGICAL DETAILS

Figure 7. Outcrop of the Andaman Flysch, looking south across Corbyn’s Cove (see Fig. 3) toward outcrops of pillow basalts and ophiolite.

include flute casts, groove casts, and current bedding. The orientation of flute casts at the base of overturned sandstone beds near Corbyn’s Cove reveals southward-directed paleocurrents (Pal et al., 2003). The relationship between the turbidites and underlying lithostratigraphic units is unclear. Onlap of the Andaman Flysch with the Mithakhari Group has been reported (Chakraborty and Pal, 2001), but no supporting evidence was found in this study. However, there is a marked change in lithology and provenance with up to 50% quartz in the Andaman Flysch in contrast to the relatively quartz-free Mithakhari Group (Pal et al., 2003). Lithic fragments in the Andaman Flysch range from micaceous metamorphic clasts diagnostic of continental sources to cherts, basalts, and weathered volcanic glass consistent with derivation from volcanic arc and ophiolite sources. Biostratigraphic evidence is vague and spans the Oligocene to the early Miocene, ca. 36–21 Ma (Pal et al., 2003). Archipelago Group The Archipelago Group represents the topmost stratigraphic unit of the Tertiary succession. The lower units comprise basal conglomerates and sandstones, overlain by calcareous arenites of the Strait Formation. This is followed by chalk and limestone with some argillaceous limestones and shale, described as the Melville Limestone (Shell Limestone) Formation. Deposition was mostly in a slope environment (Pal et al., 2005; Roy, 1983). These sedimentary rocks most likely covered most of Andaman, but recent uplift and erosion means that today only small patches can be found on the main islands, with most exposures confined to Havelock Island and associated smaller islands to the east of South Andaman in Richie’s Archipelago (Fig. 1). Radiolarians, planktonic foraminifers, and calcareous nannofossils from John Lawrence Island (Singh et al., 2000) indicate a maximum depositional age for the calcareous chalk in the Archipelago Group of 18.3–15.6 Ma (Burdigalian to Serravallian), but elsewhere the Archipelago Group may span any age from Miocene to Pliocene (Pal et al., 2005).

The aim of this work is to provide improved understanding of the burial and uplift history of the accretionary wedge and provenance of the constituent sedimentary rocks, now exposed on the Andaman Islands, through the application of thermochronometric, petrographic, and geochemical analyses of these rocks. Multiple proxies are used to get the best insight to the source provenance and to avoid any potential bias that might arise by relying on a single type of mineral. For example, mica is generally not present in arc-derived volcanic rocks, whereas apatite and zircon are. Sampling is confined to South Andaman, where there are accessible exposures of the main type localities, studied by Bandopadhyay and Ghosh, (1999), Chakraborty and Pal (2001), Pal et al. (2003, 2005), and Bandopadhyay (2005). Both Middle and North Andaman are less developed and thus are less accessible, with extensive areas of jungle that include large tribal reserves that are restricted to aboriginal peoples. Samples from South Andaman were collected for detrital thermochronometric, heavy-mineral, biostratigraphic, and Sm-Nd whole rock and single grain analyses. Zircon and Apatite Fission-Track Analysis Apatite and zircon fission-track (FT) analyses are used to define provenance and low-temperature histories, and for places where sediments have been subjected to significant (>1.5 km) burial to determine their postdepositional burial and uplift history. Fission tracks in apatite are sensitive to relatively low temperatures (typically <60–110 °C) and are ideally suited to constrain levels of postdepositional burial and timing of subsequent rock uplift and exhumation (e.g., Green et al., 1995). Zircon FT data record higher temperature cooling histories (typically ~200–310 °C) and for sedimentary rocks are more suited to provenance studies that record either volcanic formation ages or postmetamorphic cooling-exhumation ages (Carter, 1999). Samples for FT analysis were irradiated in the well-thermalized (Cd ratio for Au >100) Hifar Reactor at Lucas Heights in Australia. Apatite grain compositions were monitored using etch pits and direct measurement on a JEOL microprobe using a defocused 15 KeV Beam to prevent F and Cl migration. Durango apatite and rock salt were used as standards. Samples with mixed ages, indicated by χ2 values <5% and large age-dispersion values (>20%), were deconvolved into their principal age components using the approach of Sambridge and Compston (1994) and incorporating the method of Galbraith and Green (1990). Apatite Helium Analysis Apatite helium dating was used to provide additional constraint on rock uplift histories. This method complements FT dating, as it is sensitive to closure at ~60 °C, the temperature at which FT begins to lose sensitivity to changes in cooling rate. Helium ages were based on replicate analyses of apatite grains

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handpicked to avoid mineral and fluid inclusions. Each selected grain was first photographed, and then its dimensions were recorded for later alpha-ejection correction. Samples were loaded into platinum microtubes for helium outgassing and U/Th determination. Outgassing was achieved using an induction furnace at a temperature of 950 °C. The abundance of 4He was measured relative to a 99.9% pure 3He spike in a Pfeiffer Prisma 200 quadrupole mass spectrometer. The quantification of U/Th was performed on an Agilent 7500 quadrupole mass spectrometer using spiked solutions of the dissolved apatite. Repeated analysis of the California Institute of Technology (CIT) laboratory Durango apatite standard gives an age of 31.3 ± 1.2 Ma (2σ), based on 39 analyses. This error of the mean (6.7%), combined with the U/Th and He analytical uncertainties, is used as a measure of the total uncertainty in sample age. Ar-Ar Age Dating of Detrital White Micas The 40Ar/39Ar age of a detrital muscovite records the timing of cooling and exhumation (or crystallization) of rocks in the source region, aiding the discrimination of potential source regions for clastic sequences (e.g., Sherlock and Kelley, 2002; Haines et al., 2004). For this study, single grains were totally fused using an infrared laser ablation microprobe (IRLAMP). Samples were monitored using the GA1550 biotite standard with an age of 98.8 ± 0.5 Ma (Renne et al., 1998). The calculated J value for the samples was 0.0138 ± 0.000069. Blanks were measured both before and after each pair of sample analyses, and the mean of the two blanks was used to correct the sample analyses for the measured isotopes. Overall mean blank levels for 40Ar, 39Ar, and 36Ar were (378, 6, and 11) × 10−12 cm3 at a standard temperature and pressure. The resulting analyses were also corrected for mass spectrometer discrimination, 37Ar decay, and neutron induced interferences. The correction factors used were (39Ar/37Ar)Ca = 0.00065, (36Ar/39Ar)Ca = 0.000264, and (40Ar/39Ar)K = 0.0085; these were based on analyses of Ca and K salts. Samples were irradiated for 33 h in the McMaster University reactor (Canada). U-Pb Dating of Detrital Zircon Zircon U-Pb dating reflects the time of zircon growth, which in most cases is the igneous rock’s crystallization age. The U-Pb system is mostly unaffected by high-grade metamorphism and is effectively stable up to ~750 °C (Cherniak and Watson, 2001; Carter and Bristow, 2000). Zircon U-Pb ages from detrital grains in a sedimentary rock are therefore expected to be representative of the range of crustal ages within the contributing drainage basin. Samples for this study were analyzed at University College London by laser ablation-inductively-coupled-plasma mass spectrometry (LA-ICP-MS) using a New Wave 213 aperture imaged frequency quintupled laser ablation system (213 nm) coupled to an Agilent 750 quadrupole ICP-MS. Real-time data were processed using GLITTER and repeated measurements of the external zircon standard PL (Svojtka et al., 2001; TIMS reference age

337.1 ± 0.7 Ma) to correct for instrumental mass bias. The results have not been corrected for common lead or ranked according to degree of discordance, as the latter involves choosing an arbitrary value and is therefore open to analyzer bias. Sm-Nd Isotope Analysis Whole rock Sm and Nd isotope data in sedimentary rocks are widely used to fingerprint sediment source. 143Nd/144Nd ratios are generally normalized and expressed in epsilon units as deviation from a chondritic uniform reservoir (CHUR), where εNd = 0. A single epsilon unit is equivalent to a difference in the 143 Nd/144Nd ratio at the 4th digit. For clastic sedimentary rocks, εNd will in part represent the weighted average of the time when the sediment sources were extracted from the mantle. When melt is extracted from the mantle, it has a lower Sm/Nd ratio than its parent and therefore evolves over time to have a lower εNd than CHUR; the residual has a higher Sm/Nd than CHUR, evolving to a higher εNd over time. For this study, sandstone and mudstone samples were collected from type localities from the Andaman Flysch and the Mithakhari Group. Whole rock samples were ignited overnight at 900 °C to remove any organic material. Dissolution and analytical methods follow Ahmad et al. (2000), with the exception that the samples were spiked with a mixed 150Nd-149Sm spike and the 143Nd-144Nd ratios were measured on the spiked fraction. εNd is calculated relative to the present day (i.e., at t = 0) using CHUR 143Nd/144Nd = 0.512638. Sm and Nd blanks were <10−3 of the sample, and the laboratory Johnson Mathey Nd internal standard gave 143Nd/144Nd = 0.511119 ± 5 (1σ = 24) over the period of the analyses. As whole rock εNd typically represents the weighted average of sediment sources, further insight into the origin of the Andaman Flysch was achieved by analyzing the Nd isotopic character of single apatite grains. This was achieved using a 193 nm homogenized ArF New Wave/Merchantek laser ablation system linked to a ThermoFinigan Neptune multicollector mass spectrometer at the University of Cambridge (UK). All ablation was carried out in a He environment and mixed with Ar and N after the ablation cell. Laser spot sizes were 65–90 µm. During the analytical period, standards reproduced to better than 0.5 epsilon units, while samples typically gave internal precisions of 1–2 epsilon units. The full methodology of this in situ approach is detailed in Foster and Vance (2006). Petrography and Heavy Minerals A total of 400 points were counted in six selected samples according to the Gazzi-Dickinson method (Dickinson, 1985). A classification scheme of grain types allowed for the collection of fully quantitative information on the sampled sandstones. Transparent dense minerals were counted on grain mounts according to the “ribbon counting” method, and 200 minerals were counted also to assess the percentage of etched and corroded grains. Dense minerals were concentrated with sodium metatungstate (density, 2.9 g/cm3) using the 63–250 µm fraction treated with hydrogen

Sedimentation and uplift history of the Andaman-Nicobar accretionary prism peroxide, oxalic acid, and sodium ditionite to eliminate organic matter, iron oxides, and carbonates, respectively. RESULTS AND INTERPRETATION Biostratigraphy Attempts to identify new, more robust biostratigraphic control for the Paleogene sedimentary rocks, based on nannofossils, failed owing to the barren nature of the mudstones. Samples of mudstone were taken from the Archipelago Group, Andaman Flysch Group, and the Mithakhari Group. Whereas nannofossils are present in the calcareous Neogene sedimentary rocks (Archipelago Group) (Singh et al., 2000), we can only conclude that nannofossils either were never present in the Paleogene sedimentary rocks or have since been dissolved by weathering or dissolution below the carbonate compensation depth. Similarly, the sampled rocks yield few (as yet undated) foraminifers. Those found were either broken or abraded, consistent with reworking. At Chiriyatapu, clasts of limestone were present within the coarse-grained Mithakhari sedimentary rocks (Fig. 8A). These were found to have shallow marine reef assemblages, including Nummulites spp., small miliolids and rare Morozovella spp., fragments of rhodophyte algae, and dasycladaceans (Belzungia spp.) of Thanetian–Ypresian age (ca. 58–49 Ma) (Fig. 8B). Fission-Track and (U/Th)-He Thermochronometry Samples from the Mithakhari Group contained relatively low concentrations of heavy minerals and yielded fewer apatites and zircons than those from the Andaman Flysch. For this

B)

A)

C)

Figure 8. A and B show clasts of limestone from Mithakhari sediments. These contain shallow-marine reef assemblages that in thin section (C) are seen to include Nummulites spp., small miliolids and rare Morozovella spp., fragments of rhodophyte algae, and dasycladaceans (Belzungia spp.) that indicate a Thanetian–Ypresian depositional age. No in situ outcrops of these beds have been found on South Andaman.

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reason, FT data sets from the Mithakhari Group were smaller in comparison with those from the Andaman Flysch. Nevertheless the results adequately provide a measure of the underlying detrital FT signatures for each key lithostratigraphic unit. Data are summarized as radial plots in Figures 9 and 10, and tabulated in Table A1 (Appendix). The interpretation of detrital apatite FT data requires comparing youngest ages with sediment depositional age. If all single grain ages are less than depositional age, total resetting took place (generally indicating burial heating to >100–120 °C; e.g., Green et al., 1995, their Fig. 12), whereas if the population of measured single grain ages ranges from younger to older than depositional age, partial resetting has taken place (generally indicating burial heating to <100 °C). The main issue with interpreting the Andaman FT data is that suitable depositional ages are missing, preventing robust use of FT analysis to determine exhumation rates. Samples from the Hope Town Conglomerate collected at the Hope Town quarry (Fig. 5) yield a single population of apatites with an age of 57 ± 9 Ma. The large uncertainty in age is due to low uranium concentrations. The zircon ages comprise three age modes, with the youngest (majority of analyzed grains) at 61 ± 2 Ma, within error of the apatite age. The other zircon ages indicate sources with Late Cretaceous and Permian cooling signatures. Given that the zircon grain ages from the Hope Town Conglomerate are at or older than the Eocene biostratigraphic age (maximum age owing to the reworked nature of the fossils), the FT ages must reflect different sources. Furthermore, given that the apatite grains have the same age as the youngest zircon ages, the apatite data must also reflect provenance. The youngest apatite and zircon ages at ca. 60 Ma constrain depositional age to being at or after this time for the Hopetown Conglomerate Formation at this location. The apatites are noticeably euhedral, contain variable chlorine, and are low in uranium (Fig. 11), typical of volcanic apatites, consistent with sample petrography that records a dominant flux of volcanic detritus (see below). Similar FT ages and volcanic affinities are seen in the apatite and zircon data from the Mungleton quarry (20 km inland from the Hope Town quarry; Fig. 5), exposing a 6-m-thick sequence of interbedded greenish-gray, fine- to medium-grained sandstones, siltstones, and mudstones, underlain by conglomerates. The succession here was described as the Namunagarh Grit Formation. In contrast, the apatite FT data from a quarry at Namunagarh village (also mapped as Namunagarh Grit Formation) gave a single population age of 40 ± 4 Ma, based on 24 grains. The zircon content of this sample was too low for FT analysis. This quarry, studied by Bandopadhyay (2005), comprises tuffaceous beds with well-preserved pumice fragments and glass shards. The apatites are euhedral and have variable chlorine and low uranium contents typical of volcanic sources. Detrital assemblages in these rocks contain pyroxenes, epidote, sphene, green-brown hornblende, and chrome spinel that are consistent with provenance from a volcanic arc. Major diagenetic dissolution is also evident, so it is questionable whether the apatite age reflects source or resetting. No track lengths were measured as

Andaman Flysch Apatite Partially reset by burial 100

AND-1a Central Age: 31±2 Ma P(X^2): 0.0% Relative Error: 29% No. of grains: 38 +2 0 -2

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Figure 9. Radial plots of fission-track data from the Andaman Flysch, showing the distribution of single grain ages. Where the sample data comprise mixed grain ages the deconvolved age modes are also shown.

Apatite Unreset by burial

Zircon Unreset by burial Hopetown Conglomerate

Central Age: 57±9 Ma P(X^2): 69% Relative Error: 5% No. of grains: 19 +2

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Namunagarh Grit Mangluton Central Age: 69±9 Ma P(X^2): 0.0% Relative Error: 65% No. of grains: 21

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Figure 10. Radial plots of fission-track data from the Mithakhari Group, showing the distribution of single grain ages. Where the sample data comprise mixed grain ages the deconvolved age modes are also shown.

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Apatite grain composition Hopetown Conglomerate

Chlorine (wt%)

Andaman Flysch 1

Figure 11. Plot comparing uranium and chlorine for apatite grains from the Andaman Flysch and Hopetown Conglomerate, Mithakhari Group. The plot clearly shows that the apatite came from different sources. The low uranium Hopetown Conglomerate apatites, which also have variable amounts of chlorine, are typical of volcanic apatites. The volcaniclastic petrography and euhedral form of the apatites support this and suggest that they came from arc related sources.

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a result of the low spontaneous track density. A sample of the Namunagarh Grit Formation from Chiriyatapu on the southern coast of South Andaman yielded only zircon, which produced three FT provenance age modes at 40 ± 3 Ma, 91 ± 7 Ma, and 329 ± Ma. The youngest zircon FT age mode, at 40 ± 3 Ma, indicates that the Thanetian–Ypresian limestone clasts found in these rocks were derived from erosion of significantly older material (by ca. 16 Ma). Thus we conclude that the age of the Mithakhari Group is constrained to be younger than 60 Ma at two locations, and younger than 40 Ma at a third location. Nevertheless, the volcanic origin of the material, with little sign of reworking, suggests that the rocks were deposited only shortly after the recorded mineral ages. This is consistent with diachronous deposition in local basins, as expected in a subduction zone setting (e.g., Draut and Clift, 2006). Apatite and zircon FT data from four samples collected through the section of Andaman Flysch at Corbyn’s Cove are closely similar in age (most are 35–40 Ma) to the youngest age modes in the Namunagarh Grit Formation beds. The youngest ages comprise most analyzed grains for zircon and all analyzed grains for the apatites. In addition, two of the samples show zircon age modes at 58–67 Ma, similar to the age modes detected in the Mithakhari Group. Three of the samples also show Mesozoic and Paleozoic FT ages. The Andaman Flysch is widely considered to be Oligocene in age, although biostratigraphic evidence

is not robust (Pal et al., 2003). While both zircon and apatite FT ages are older than the Oligocene, some partial resetting (burialrelated heating) may have taken place. Thermal Modeling To constrain postdepositional thermal history, apatites from one of the samples were analyzed by the (U/Th)-He method (Table A2, Appendix). Replicates yielded an FT corrected age of 16 ± 1 Ma, crudely representing the time at which the sample cooled to <50 ± 10 °C. To define more robustly the sample postdepositional thermal history, the FT and helium data can be jointly modeled. Ideally this requires incorporating a sample depositional age, but, as discussed, this is not well defined, and so we resorted to using the youngest detrital ages as an upper limit for the time of deposition. For the Andaman Flysch, deposition must have taken place at or after 35–30 Ma on the basis of the youngest argon mica ages. In addition the Andaman Flysch sedimentary rocks must have been at or near surface temperatures by the middle Miocene (ca. 16 Ma) because in the Hobdaypur area of South Andaman (Fig. 2) these rocks are seen to be conformably overlain by sedimentary rocks of the Archipelago Group, although in the eastern part of the island Archipelago rocks are juxtaposed with rocks of the Mithakhari Group along a faulted contact (Pal et al., 2005). A phase of reburial up to ~60 °C

Sedimentation and uplift history of the Andaman-Nicobar accretionary prism is also required on the basis of recent diagenetic evidence from the Archipelago Group sedimentary rocks (Pal et al., 2005). With these constraints the combined apatite FT and helium data were modeled, using the data-driven modeling approach of Gallagher (1995) that combines multicompositional FT annealing and helium diffusion models (Ketcham et al., 1999; Meesters and Dunai, 2002). The best-fit solution (Fig. 12) highlights three important stages: (1) a requirement for deposition and burial to peak temperatures of ~80–90 °C between ca. 30 and 25 Ma, (2) uplift to the surface between ca. 25 and 20 Ma, and (3) reburial in the Miocene (ca. 25–5 Ma) to temperatures of 50–60 °C (broadly equivalent to depths of ~1–1.5 km, assuming geothermal gradients of 30 °C/km).

Depositional age range

Temperature °C

0

Oldest depositional age of Archipelago Group Modeled results are consistent with Archipelago Group (Pal et al., 2005) diagenetic data that indicate maximum burial temperatures of 45-55°C

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Obs. Age : 36 Ma Pred. Age : 35 Ma Obs. Mean length : 13.24μm Pred. Mean length 13.66μm Obs. S.D. : 1.61μm Pred. S.D. : 1.61μm Oldest track 39 Ma

Track Length (microns) He/P 50

Helium diffusion model

40 He Obs : 13.0 He Pred: 14.7 Ma Grain R: 75.0 microns

30 20 10 0

40

30

20 Age (Ma)

Detrital Argon Data Four samples were analyzed from the Andaman Flysch, Corbyn’s Cove section, which yielded 114 grain ages. One additional sample from Mithakhari Group sedimentary rock at Chiriyatapu was also analyzed (33 grain ages). The results are displayed as probability plots in Figure 13 and tabulated in Table A3 (Appendix). Ages for the Mithakhari Group sample extend from 70 Ma to the Archean, with the bulk of ages spanning the Mesozoic and Paleozoic. The Andaman Flysch ages are in general younger, ranging from 30 Ma to the Proterozoic. Most of the ages (71 grains) are <200 Ma, with 20 grains falling between 30 and 60 Ma, and 32 grains between 60 and 80 Ma. Given that postdepositional heating was modest, as constrained by the FT data, which measure lower closure temperatures than argon, the young mica ages, some of which are younger than the FT data, point to an Oligocene depositional age: i.e., the Andaman Flysch was deposited at or after 30 Ma. Detrital Zircon U-Pb Data Preliminary zircon U-Pb data from the Andaman Flysch (Fig. 14; Table A4, Appendix) show a large Proterozoic population, with a few grains showing Archean and Cretaceous–Eocene ages. The youngest zircons (five grains) give an average age of 48 ± 5 Ma, which overlaps both zircon FT and mica ages. These grains are euhedral and have concentric zoning typical of magmatic zircon, suggesting a direct contribution from an early Eocene igneous source rather than reworking of older sedimentary deposits. Petrography and Heavy Mineral Data

1 2 3 4 5 6 7 8 9 10 1112 1314151617181920

50

235

10

0

Figure 12. Best-fit thermal-history model for apatite fission-track and (U-Th)/He data from the Andaman Flysch. Key features are a requirement for maximum burial temperatures between ca. 20 and 30 Ma and rapid cooling at ca. 20 Ma. Values given are for observed data (Obs.), model predicted values (Pred.), and standard deviation (S.D.). For the helium plot, R is the grain radius and He/P is the helium concentration profile.

The very fine to fine-grained turbidites of the Andaman Flysch have an intermediate quartz content (Q 49 ± 2%) (Fig. 15) and contain subequal amounts of feldspar (F 22 ± 6%; plagioclase feldspar 39 ± 9%) and lithic grains (L 29 ± 6%) as seen in Table A5A (Appendix). The latter chiefly consist of very low rank to medium rank and subordinately high rank metapelitemetapsammite grains (Garzanti and Vezzoli, 2003), indicating provenance from a collisional orogen (Dickinson, 1985). Volcanic grains are mainly felsitic and are significant (Lv 8 ± 1%), suggesting subordinate contributions from a volcanic arc. Transparent heavy-mineral assemblages in the Andaman Flysch (Table A5B, Appendix) have very low concentrations (0.3 ± 0.2%). Ultrastable species (zircon, tourmaline, rutile, chromian spinel) represent 21 ± 2% of the assemblage, and ferromagnesian minerals (pyroxenes, amphiboles) are invariably absent, clearly showing the strong influence of diagenetic dissolution. The very coarse grained sandstones of the Namunagarh Grit Formation chiefly consist of plagioclase and microlitic, lathwork, and felsitic volcanic grains. The quartz content is low (Q 14 ± 10%), and only a few K-feldspar and granitic lithic grains

236

Allen et al.

Andaman Flysch-1A (90 grains) Relative probability

Relative probability

Mithakhari Group Chiriyatapu (33 grains)

0

0

400

800

500

1000

1500

2000

2500

3000

Age (Ma)

1200 1600 2000 2400 2800

Age (Ma)

0

Relative probability

Relative probability

Andaman Flysch Corbyn’s Cove (114 grains)

200 400 600 800 1000 1200 1400 1600

0

Age (Ma)

50

100

150

200

250

300

350

400

Age (Ma)

Figure 14. Probability plots showing the distribution of concordant detrital zircon U-Pb ages analyzed in this study.

Relative probability

Younger grain ages (71 grains)

20

60

100

140

180

220

Age (Ma) Figure 13. Probability plots showing the distribution of detrital argon mica ages analyzed in this study.

occur, indicating provenance from a largely undissected volcanic arc (Marsaglia and Ingersoll, 1992). A few terrigenous (shalesandstone) and very low rank to medium rank metapelite and metapsammite grains also occur, suggesting minor reworking of sedimentary and low-grade metasedimentary sequences. Significant chert, along with traces of metabasite and serpentinite grains, point to minor contributions from oceanic rocks. In the Namunagarh Grit Formation, opaque grains represent 21 ± 16% of total heavy minerals. Transparent heavy-mineral assemblages have poor concentrations, pointing to significant diagenetic dissolution. These mainly include pyroxenes (mostly green augite; 35 ± 24%) and epidote (34 ± 19%), with subordinate sphene (11 ± 8%), green-brown hornblende (8 ± 8%), chromian spinel (8 ± 8%), minor apatite (2 ± 1%), garnet, rutile, and other

Sedimentation and uplift history of the Andaman-Nicobar accretionary prism

Q

Lmf

A) + cra ton

O

B)

Andaman Flysch Namunagarh Grit + Bengal Fan Irrawaddy 125 Irrawaddy 180 Irrawaddy 80

O O

rec

+

yc

led

oro

ge

O

n

O O

magmatic arc O

F

O

L

O

O

Lv + Lmv

titanium oxides, consistent with provenance from a volcanic arc. It must be kept in mind, however, that the original composition of the Namunagarh Grit Formation sandstones (abundant volcanic glass and ferromagnesian minerals) was markedly different with respect to orogen-derived turbidites, and therefore that dissolution reactions of ferromagnesian minerals may have progressed at different rates and in the presence of partially buffered or even saturated interstitial waters. Sm-Nd Whole Rock and Single Grain Analyses Sandstone and mudstone samples from the Mithakhari Group yielded whole rock εNd values of –7.2 and 3.1(+), respectively. The Andaman Flysch samples gave whole rock values of εNd –11.1 and –8.2 for the sandstone sample and mudstone sample, respectively. Such values are lower than those of Himalayan sources, which typically range between εNd –12 and –26 (Galy and France-Lanord., 2001) and are inconsistent with Indian Shield sources (Peucat et al., 1989). Very little data exist for the relevant Myanmar source regions. Analyses of modern river sediment from the Irrawaddy River give εNd values between −10.7 (Colin et al., 1999) and −8.3 (this paper: Table 2). The Western Indo-Burman ranges (Fig. 16), which may have been exhuming by the Oligocene (Mitchell, 1993), have more negative values (Colin et al., 2006; Singh and France-Lanord, 2002; this paper, Table 2, and Table A6A, Appendix). Figure 16, which compares the whole rock data from the Andaman Flysch with possible source areas, shows that the most probable source area lies within Myanmar. To understand the source area in more detail, and in particular the significance of apatite from the Andaman Flysch, which appears to come from more distant sources in comparison with local arc and ophiolite sources in the Mithakhari Group sedimentary rocks, Nd isotope composition was measured on single grains of apatite from the Andaman Flysch. Apatite typically has

Lms + Ls

237

Figure 15. Petrography of Tertiary sandstones of South Andaman. Whereas composition of the Namunagarh Grit Formation sandstones points to provenance from an undissected-transitional magmatic arc, a recycled orogenic provenance is clearly indicated for the Andaman Flysch (Dickinson, 1985). These very fine to fine-grained turbidites compare closely with modern sands of homologous grain size from the Irrawaddy Delta. Q—quartz; F—feldspars; L— lithic grains; Lv + Lmv—volcanic and low-rank metavolcanic; Ls + Lms—sedimentary and low-rank metasedimentary; Lmf—high-rank metamorphic.

Sm/Nd ratios of 0.2–0.5, therefore for non-age–corrected apatite the 143Nd/144Nd ratio tends to reflect the Nd isotopic composition of their parent whole rock, enabling a direct linkage between thermochronometric data and source in terms of Nd isotopic signature. This approach is particularly effective if the source rocks are young and/or their potential range of Sm-Nd has been quantified. It is also important to note that this approach yields information that is not the same as sediment whole rock values that record an average value (weighted by Nd concentration) of source rocks. It is also possible that the detrital apatite may be dominated by a single apatite-rich source. The majority (~60%) of the single grain analysis of apatite from the Andaman Flysch produced εNd values of –5 to +5, typical of juvenile volcanic whole rock values and contrasting significantly with the bulk sediment values of εNd –11.1 and –8.2 (Fig. 17; Table A6A, B, Appendix). The restricted range of εNd values for these apatites, which gave partially reset FT ages between 30 and 40 Ma (Table A1, Appendix), imply a single source region for the apatite grains. Little is known about the apatite Sm-Nd systematics for the source rocks of the Andaman Flysch. However, as Figure 17 shows, the Andaman Flysch apatite Nd isotope data contrast with apatite Nd data from a Holocene sand dominated by Himalayan sources collected from the Bengal Basin near Joypur, West Bengal. This plot clearly shows that there is little evidence for material eroded from Himalayan sources in this sample. DISCUSSION Constraints on Sedimentation and Uplift New thermochronometric evidence and field observations from type locations on South Andaman, combined with existing biostratigraphy and petrography, provide an improved chronology for deposition and uplift. The accretionary setting means that sedimentation history will be intimately tied to subduction

3,4

2

Values from south and east craton only; –30 to –3618

Dominantly Archean19

Precambrian: Archean and Proterozoic; Cretaceous and Paleogene (55–65 Ma) grains16

3 samples give values of –4.0; 1 sample (farther north) gives a value of –7.416

Dark-gray, very fine grained volcanic arenite of Eocene (Oligocene?) age. Only limited mica and heavy minerals. Many shale and siltstone lithics.16

Predominantly gneisses and granites; opaques, orthoproxene, and sillimanite found in rivers draining east craton.17

Irrawaddy: Proterozoic, Ordovician, Jurassic–Cretaceous, Paleogene, and Neogene grains14

Oligocene bedrock: Eocene and Miocene: Tertiary grains absent; 65–400 Ma grains very rare6

Signature today: As determined from Ganges sediments: mostly >400 Ma– Precambrian. Grains <30 Ma rare. Grains 30– 400 Ma very rare5

Irrawaddy: –10.7, –8.313

Oligocene bedrock: Eocene: av. –83 Miocene: –14 to –174

Signature today: Higher Himalayan as determined from Ganges: av. –17.52

Cretaceous arc rocks and Triassic forearc-backarc sediments on metamorphic basement. Mogok schists, gneisses, and intrusives, Shan-Thai Proterozoic–Cretaceous sedimentary rocks on schist basement. Irrawaddy River sediment plots in “recycled orogenic” province field of QFL plot (Dickinson, 1985).12

In the Oligocene, less metamorphosed “Higher Himalayan protolith” would likely have been exposed. Rocks were low-grade metamorphic, sub-garnet grade, nonmicaceous (as determined from Miocene foreland basin rocks).1

TABLE 2. TYPICAL SIGNATURES OF POTENTIAL SOURCE REGIONS U-Pb ages of zircons Rock description, heavy minerals, and petrography Sm–Nd Whole rock εNd(0)

Najman and Garzanti (2000). Galy and France-Lanord (2001). Robinson et al. (2001); Najman et al. (2000). 5 Campbell et al. (2005). 6 DeCelles et al. (2004). 7 Brewer et al. (2003); Najman and Pringle (unpublished data). 8 DeCelles et al. (1998); Najman and Garzanti (2000). 9 Campbell et al. (2005). 10 Najman et al. (2005). 11 Najman et al. (2004, 2005). 12 Mitchell (1993); Pivnik et al. (1998); Bertrand et al. (1999); R. Allen et al. (unpublished data). 13 Colin et al. (1999); R. Allen et al. (unpublished data). 14 Bodet and Schärer (2000). 15 R. Allen et al. (unpublished data). 16 R. Allen et al. (unpublished data). 17 Mallik (1976). 18 Peucat et al. (1989); Saha et al. (2004). 19 Mishra and Rajamani (1999); Auge et al. (2003).

1

Indian Shield Dominantly Arc hean craton. Subordinate Proterozoic mobile belts and Gondwanan sedimentary cover.

Western Indo-Burman Ranges Accretionary prism; may have been exhuming and thus a sediment source in late Eocene–Oligocene. Data determined from Arakan modern river sands (R. Allen et al., unpublished data) unless otherwise stated.

Burma Region drained by Irrawaddy Data from modern Irrawaddy River sediment. Shan-Thai block lies to east, forearc-backarc of Indo-Burman Ranges (IBR) to west. Paleocontinental margin.

Himalaya (Southern Flank) Drained by Ganges and tributaries Oligocene bedrock characteristics interpolated from Eocene and Miocene foreland basin rocks Bedrock signal today taken from modern river sediments; higher Himalaya dominates detritus

Source region

Ar-39Ar ages of white mica

No data available

No mica present in preMiocene sedimentary rocks16

Irrawaddy: Cretaceous to Miocene grains; peak 30–55 Ma15

Oligocene bedrock: Micas absent in Eocene, very rare in late Oligocene–Miocene8

Signature today: As determined from Ganges tributaries: Neogene peak; subordinate grains spanning to Precambrian7

40

No data available

Cretaceous and Paleocene to Eocene grains16

Irrawaddy: Neogene, Paleogene, Cretaceous grains15

Oliogocene bedrock: Eocene: peaks at 45 Ma, 119 Ma, and 343 Ma;10 Miocene: peaks at 30 Ma, 60–75 Ma, 117 Ma, 300–370 Ma11

Signature today: No FT data available for Ganges sediment; He data <55 Ma, Pliocene– Pleistocene peak9

Zircon fission-track ages

Andaman Flysch Namunagarh Grit IndoBurman Ranges

Himalayan Sources

Modern Irrawaddy River -30 to -36

Indian Shield present-day Brahmaputra / Ganges TSS

HHS LHS -28

-24

-20

-16

-12

-8

-4

0

4

8

Epsilon Nd Figure 16. Plot comparing Sm-Nd whole rock values from the Namunagarh Grit Formation and Andaman Flysch Formation with data from possible source regions (Colin et al., 2006; DeCelles et al., 2004; Peucat et al., 1989; Singh and France-Lanord, 2002). The data do not support the Himalayan region as the main source. TSS—Tethys Sedimentary Sequence; HSS—Higher Himalayan Sequence; LHS—Less Himalayan Sequence.

Holocene, Bengal delta Andaman Flysch

143 Nd / 144 Nd

0.5132 0.5130

+8

0.5128

+4

0.5126

0

Field of typical Himalayan apatites

0.5124

-4

0.5122

-8

0.5120

-12

0.5118

-16

0.5116

-20

0.5114

-24

0.5112

0

0.05

0.1

0.15

0.2

0.25

0.3

0.35

Epsilon Nd

0.5134

0.4

147Sm / 144Nd Figure 17. Single grain Nd measured on apatite from the Andaman Flysch Formation compared with a Holocene sand from the Bengal Basin known to have its source in the Himalayas. The clear difference between these two samples supports the Sm-Nd whole rock data plotted in Figure 16.

240

Allen et al. modeling, which indicates that burial to maximum temperatures of ~90 ± 5 °C was reached ca. 25–20 Ma, ending with a phase of rapid cooling at ca. 20 Ma. This implies a discrete episode of major uplift at ca. 20 Ma consistent with deposition of the Archipelago Group shallow marine volcanic-rich sedimentary rock from ca. 18 Ma onward (Pal et al., 2005; Singh et al., 2000), which was deposited on top of the Andaman Flysch. Uplift at ca. 20 Ma is also regionally significant, coinciding with major stratigraphic changes in the Irrawaddy Delta, Mergui Basin (Fig. 1), in sediments accreted onto the Andaman-Nicobar Ridge (Curray, 2005), and formation of an unconformity in the IndoBurman Ranges (Acharyya et al., 1989). Uplift of the Andaman Ridge at 20 Ma is also coincident with the uplift of the Owen and Murray Ridges in the Arabian Sea (Mountain and Prell, 1990), suggestive of a wider plate tectonic trigger. An outstanding question is, what was the original maximum thickness of the Andaman Flysch? Hydrocarbon-exploration

180150-

100-

Lee and Lawver, 1995 Guillot et al., 2003 5040302010-

Archipelago Group

uplift

Angle of convergence (from N)

Rate of Indian convergence (mm/yr)

behavior, offscraping, accretion, and uplift. While in such a setting uplift may be viewed as a more or less continuous process, we recognize some distinct uplift events that have affected the South Andaman geology. The earliest conglomerates and grits of the Mithakhari Group record subaerial exposure and erosion of the arc sequence and shallow marine limestone cover, as described above, constraining the age of ophiolite as pre–Mithakhari Formation (early to middle Eocene) and post–shallow marine sedimentation. The petrography of the Mithakhari Group is indicative of an undissected volcanic arc with minor reworking of sedimentary and low-grade metasedimentary sequences and minor contributions from oceanic igneous rocks. Our data provide maximum and minimum age constraints for deposition of the Andaman Flysch. Detrital Ar-Ar mica and FT data constrain maximum depositional age to ca. 30 Ma. The minimum depositional age is constrained by the thermal history

deposition

Figure 18. A comparison between Indian convergence history (after Lee and Lawver, 1995, and Guillot et al., 2003) and the uplift and sedimentation history of the rocks studied from South Andaman Island.

uplift

unconformity Andaman Flysch

deposition

Mithakhari Group

contact unclear Namunagarh Grit Hopetown Conglom. unconformity

continuous deposition ?????

Ophiolite pelagic cover sediments

uplift

deposition

? ? ? obduction

Ophiolite

? ? ? 0

20

40 Age (Ma)

60

80

100

Sedimentation and uplift history of the Andaman-Nicobar accretionary prism

QT QS

AL

NORTH CHINA

SPG

LHASA

SOUTH CHINA

India

WB

INA INDOCH

The Mithakhari Group shows clear evidence of a dominant contribution from an arc, in keeping with its interpreted forearc basin depositional environment and in line with previous provenance work on the basis of petrography (Bandopadhyay, 2005). The rocks plot in the magmatic arc field on the QFL petrographic plot (Fig. 15); volcanic fragments and glass shards are common, and mineral grains, e.g., apatite, are euhedral. A positive whole rock εNd signal (+3) is indicative of derivation from a maficjuvenile source. The source was likely that of the Cretaceous– Eocene arc that stretched from the Himalaya collision zone through Myanmar to Sumatra (Mitchell, 1993), as reflected in the FT ages of the euhedral apatites. A subordinate continental source is confirmed by petrography and Precambrian–Paleozoic Ar-Ar mineral grain ages. Such a source was probably the continental margin of the Shan-Thai Block (Sibumasu) extending

QD

SU

Sediment Provenance

TARIM KL

SIBUMA

wells drilled to the east and the west of Middle Andaman indicate thicknesses of <1000 m (Roy, 1992), whereas outcrop estimates range from ~750 to 3000 m (Pal et al., 2003). Results from thermal history modeling indicate burial of the sampled Andaman Flysch beds beneath at least 2.5 km of cover (assuming geothermal gradients in the range of 35–30 °C). In a growing accretionary wedge, burial is as likely to be due to thrust stacking as it is to sedimentation. Crucially ~2.5 km of burial is incompatible with the unannealed apatite data from the stratigraphically lower Mithakhari Group, highlighting the fact that these different units must have been deposited in different locations (sub-basins?) within the accretionary wedge. Thermal history modeling also shows a phase of accelerated cooling from ca. 10 to 5 Ma. This is probably tied to a phase of regional Miocene–Pliocene uplift linked to spreading in the Andaman Sea driven by increases in the subduction rate and dip of the subducted slab (Khan and Chakraborty, 2005). This last phase of uplift, the result of squeezing between subduction in the west and extension in the east, is likely responsible for the present topography of the Andaman Islands. The relationship of these episodes of uplift and erosion to regional events can be appreciated when these new constraints are compared against reconstructions of Indian plate convergence history. Figure 18 plots the India-Asia convergence rate and convergence angle from the studies of Lee and Lawver (1995) and more recently from that of Guillot et al. (2003). The graph shows that the studied outcrops of Andaman Flysch were deposited at a time of more northerly convergence and ended when subduction shifted to a less oblique angle. Seismic lines taken from across the Andaman Arc show how the angle of subduction can influence wedge development where the intensity of deformation, which increases from north to south, changes with obliquity of subduction (Curray, 2005). This relationship may explain the evidence for a discrete episode of uplift, involving the Corbyn’s Cove Andaman Flysch, at ca. 20 Ma.

241

QD = Quidam terrane QS = Qamdo-Simao terrane QT = Qiantang terrane KL = Kunlun terrane AL = Alao Shan terrane SPG = Song Pan Ganzi terrane WB = West Burma SWB = Southwest Burma

SWB

Figure 19. Map showing the principal terranes that have amalgamated to form SE Asia (after Metcalfe, 1996).

down to peninsular Thailand, adjacent to the forearc (Metcalfe, 1996) (Fig. 19). The data show a clear change in petrographic and isotopic characteristics between the Mithakhari Group and the overlying Andaman Flysch. Andaman Flysch composition represents the first sedimentary material to plot in the recycled orogenic province (Fig. 15). Initial results from the pilot study of Sm-Nd whole rock values differ markedly between the Mithakhari Group and the Andaman Flysch, particularly when the finer grained facies are compared. The positive εNd signal of the Mithakhari Group contrasts strongly with the more negative εNd signal (−11) of the Andaman Flysch mudstone, suggestive of a contribution from older continental crustal sources (Fig. 15). Detrital micas are of radically different ages, with the age peaks for the Mithakhari Group falling within the Precambrian–Paleozoic spectrum, whereas those from the Andaman Flysch are typically of late Mesozoic and Tertiary age (Fig. 13). A provenance contrast is not apparent from the zircon FT data (Figs. 9 and 10), as the detrital modes are similar. The apatite FT data cannot be directly compared, as the Andaman Flysch apatites have seen some postdepositional partial resetting. However, comparison of apatite

242

Allen et al. older than 50 Ma, present in both the Mithakhari Group and the Andaman Flysch, are consistent with known sources along the western margin of the Shan-Thai Block (Sibumasu) adjacent to the forearc. The presence in the Andaman Flysch of a small but recognizable population of mica ages between 30 and 40 Ma is not consistent with these Sibumasu sources, as the youngest granites on the Thai peninsula are ca. 50 Ma (Charusiri et al., 1993). Possible sources for the Tertiary mica ages might instead include a region affected by India-Asia collision at ca. 50 Ma, possible candidates being either the nascent Himalaya to the north, Transhimalaya, or a more northeastern source—namely, Myanmar locations of the Burman-Thai Block—where Tertiary metamorphism, magmatism, and grain isotopic characteristics and ages are ascribed to the effects of the India-Asia collision (Bertrand et al., 1999; 2001, Bodet and Schärer, 2000; Barley et al., 2003). Isotopic characteristics of these Myanmar rocks are incompletely documented at present. Those that are available are taken from bedrock, and analyses of sands from the Irrawaddy River, which drains the areas under discussion—i.e., the Central Myanmar Basin, to its east the Mogok Belt and Shan Plateau, and to its west the Indo-Burman Ranges (Bodet and Schärer, 2000; Colin et al., 2006; Singh and France-Lanord, 2002; Table 2, this study)—as verified by 30 Ma Ar-Ar ages from the Irrawaddy

chemistry (chlorine-uranium content, Fig. 11) does highlight a major difference in the apatite source. To some extent these differences between the Mithakhari Group and the Andaman Flysch may reflect different mixtures of the same sources. The continental-derived material that dominates the Andaman Flysch represents only a minor contribution in the arc-dominated Mithakhari Group. Variations may have as much to do with depositional setting than source variation. In this regard, sedimentary rocks of the Namunagarh Grit Formation may represent a perched forearc basin, whereas the Andaman Flysch represents deep-water turbidites. It was proposed by Karunakaran et al. (1968) and Pal et al. (2003) that these “recycled orogenic” Andaman Flysch rocks were sourced by the Irrawaddy River of Myanmar and deposited in a forearc basin. In contrast, Curray et al. (1979) considered the Andaman Flysch to be trench sediments offscraped from the Himalayan-derived Bengal Fan on the downgoing Indian slab. Significant contribution from the northward-drifting cratonic Greater India can be ruled out by the dissimilar petrography, Sm-Nd signature, and mineral cooling ages (Fig. 16; Table 2) as well as by the apatite Nd data (Fig. 17). No apatites were measured with significantly negative εNd values (e.g., −30) diagnostic of Indian Shield sources. The mica argon ages

Paleogene Eurasia

nas

cen tH fore land imalay a bas ins

Tsangpo River, flows along Himalayan suture draining east to South China Sea

rivers draining southern flanks

an

anic

-T ha

arc

Early Bengal Fan

Sh

volc

Greater India

Figure 20. Cartoon illustrating the Paleogene paleogeography and paleodrainage of SE Asia. Although this sketch simplifies the location of the subduction zone, the Asian margin and Andaman crust would have changed throughout the Paleogene as Greater India moved northward. The cartoon does serve to illustrate the general location of continental sources and drainages to the Andaman region.

continental drainage

iSu nd ala

Paleogene location of Andaman crust

nd

c ar nch nic re lca a T vo und S

Indian Ocean

Sedimentation and uplift history of the Andaman-Nicobar accretionary prism region (Table 2). Discrimination between these two possible sources, the Himalayan or Myanmar region, requires an understanding of the different paleogeography and paleo-drainage of the region in the Oligocene in comparison with the present day, as discussed below (Fig. 20). Prior to the India-Asia collision, the northward extension of the Sunda Arc as far as Pakistan permits correlation between the southern margin of Asia in the Himalayan region with equivalent rocks in Myanmar (Mitchell, 1993) and farther south. Subsequent to continental collision at ca. 50 Ma (Rowley, 1996), the Himalayan thrust belt started to develop in the north while the eastern (Myanmar) region remained an active continental margin with, for example, the Central Basin along which the Irrawaddy now drains, the site of the continental margin between the subduction zone to the west and the arc to the east (Pivnik et al., 1998). Movement along the Sagaing dextral strike-slip fault (initiation in pre-Miocene times: Mitchell, 1993) and the opening of the Andaman Sea in the late Miocene (Curray, 2005; Khan and Chakraborty, 2005) were responsible for changes in the coastal paleogeography of this region. In the Paleogene the Yarlung-Tsangpo River (draining the Himalayan Arc and suture zone) probably drained into the South China Sea prior to its capture by the Irrawaddy River, and finally the Brahmaputra River in the Neogene (Clark et al., 2004, 2005). Thus, Ganges-type rivers—those draining only the southern slopes of the Himalayan mountain belt, with no arc component—would have contributed the Himalayan signal to the Indian Ocean (Fig. 17). The southern slopes of the Himalayas consist of Indian crust. Characteristics of these rocks changed with time as Neogene Himalayan metamorphism subsequently increased the metamorphic grade and overprinted the metamorphic cooling ages of minerals and rocks exposed at the surface today. In contrast, the northeastern Myanmar region was devoid of Indian crustal rocks, and no thrust belt was present to bar the Asian and arc sources from draining south. Prior to the opening of the Andaman Sea, rivers draining the Myanmar region presumably would have drained into the trench-forearc system, which acted as a “sink,” preventing deposition farther west onto the Indian plate. The Andaman Flysch records mixed orogenic and arc provenance. The young (<100 Ma) U-Pb–dated detrital zircons from the Andaman Flysch are consistent with derivation from igneous sources, most likely an eastern (Myanmar) provenance based on similar aged grains found in modern Irrawaddy River sediment (Bodet and Schärer, 2000), as grains with young U-Pb zircon ages are extremely rare in the southern flanks of the Himalayas (Campbell et al., 2005; DeCelles et al., 1998, 2004). Myanmar sources also can adequately account for the majority of the “recycled orogenic” component of the detritus in the Andaman Flysch. Both petrography and εNd whole rock signatures, especially of the fine-grained material, are nearly identical to those of the modern Irrawaddy River (Figs. 15 and 16). 40Ar-39Ar and FT data from the Irrawaddy River also are

243

similar to data obtained from the Andaman Flysch, showing a mica age peak of ca. 30–60 Ma with a population of older Cretaceous grains and zircon FT ages of Miocene and Paleogene (20– 65 Ma) to Cretaceous (Table 2). Nevertheless, some disparities are evident. The lack of zircon grains with Jurassic U-Pb ages for the Andaman Flysch is surprising, considering their prevalence in the Mogok Belt (Barley et al., 2003) and their presence in modern Irrawaddy River sand (Bodet and Schärer, 2000). Recourse to a Himalayan contribution through the Bengal Fan is not required to explain the recycled orogenic component of the Andaman Flysch. However, a minor contribution from the Himalayan thrust belt cannot be ruled out: for example, grains aged 100–1500 Ma in the Andaman Flysch are common to both Himalayan and Myanmar rock types. In addition, the intermediate whole rock εNd values may represent a mixture between a more negative Himalayan source and more positive Myanmar sources, as implied by the single grain apatite analyses. The possibility of dual provenance, with Himalayan-derived Bengal Fan material and Myanmar-derived material meeting and mixing in the trench, is a model that would be consistent with the regional seismic data, which show folding and uplift of Bengal Fan sediments at the base of the slope (Curray, 2005). Definitive discrimination between Myanmar and Himalayan sources, if possible at all, awaits more information on the source rock geology of Myanmar, in particular that of the Shan Plateau and Mogok Belt, and analysis of as-yet-undrilled Oligocene sediments preserved in the Bengal Fan, and coeval sedimentary rocks of the Central Myanmar Basin. CONCLUSIONS The Mithakhari Group was deposited in the late Paleocene– Eocene, with a maximum age of ca. 60 Ma. This current study shows that the sedimentary rocks are predominantly arc derived from a proximal source in keeping with the interpreted forearc depositional environment. A subordinate contribution from an older continental source was most likely to have been the western Sibumasu margin, but a Transhimalayan arc unit source cannot be ruled out. The Oligocene Andaman Flysch was deposited between 30 and 20 Ma, and then uplifted by 20 Ma. It shows a clear change in petrographic and isotopic characteristics from the Mithakhari Group and is composed mainly of “recycled orogenic” sources with a subordinate arc provenance. It represents the earliest record in the region of major influx from a continental area. The recycled orogenic component is most simply explained by erosion from the northeastern (Myanmar) continental region. Isotopic and petrographic differences between formations may be explained to some extent by different mixtures and contributions from the same source regions. Although dual provenance is a favorable model for this region, detailed discrimination between Himalayan and Myanmar sources, at present, awaits more data from Myanmar source rocks.

Latitude Longitude

Mineral

N11°41.407′ E92°43.501′

N11°41.339′ E92°41.106′

N11°42.005′ E93°32.301′

Namunagarh Quarry

Chiriyatapu

N11°38.457′ E92°45.318′

N11°38.459′ E92°45.368′

N11°39.461′ E92°45.417′

AND-1B

AND-3

AND-4 42

15

Zircon

Zircon

50

Apatite

47

32

Zircon

Apatite

57

43

Zircon

Apatite

38

38

Apatite

Zircon

24

21

Apatite

Apatite

18

19

Apatite

Zircon

25

Zircon

No. of grains

0.505 (3450)

11.106 (6131)

0.108 (3450)

1.106 (6131)

0.515 (3450)

1.106 (6131)

0.516 (3450)

1.106 (6131)

0.525 (3450)

1.106 (6131)

1.106 (6131)

0.535 (3450)

1.106 (6131)

0.528 (3450)

10.63 (7575)

0.427 (923)

10.24 (1930)

0.497 (852)

6.115 (3085)

0.499 (738)

8.341 (6439)

0.264 (497)

12.53 (7468)

0.060 (108)

0.150 (322)

4.506 (3210)

2.652 (5733)

4.828 (910)

2.556 (4739)

0.0

0.04

0.0

0.03

0.0

7.3

2.499 (3696) 3.685 (1859)

0.0

0.02

0.0

99.6

0.0

0.0

0.03

0.0

(Pχ2)

70.1

20.2

73.0

22.6

41.5

16.3

44.5

28.8

92.6

0.0

64.5

42.3

69.1

56.2

RE%

Dispersion

4.07 (3147)

1.653 (3116)

2.255 (1344)

0.282 (506)

0.246 (946)

4.872 (1857)

0.209 (191)

0.064 (58)

10.972 (4182)

4.32 (1756)

ρι Ni

12.16 (4942)

ρs NS

67.95 ± 7.6

30.75 ± 1.5

58.55 ± 11.3

35.85 ± 1.9

53.75 ± 4.4

37.15 ± 1.8

62.15 ± 4.6

30.85 ± 2.2

143.0 ± 22

39.8 ± 4.2

68.9

75.15 ± 7.9

56.5 ± 8.5

91.8 ± 10.8

Central age ±1σ (Ma)

38 ± 1 (20)

31 ± 2 (47)

36 ± 2 (12)

36 ± 2 (50)

41 ± 2 (26)

37 ± 2 (57)

44 ± 2 (24)

31 ± 2 (38)

40 ± 3 (9)

40 ± 4 (24)

58 ± 4 (10)

101 ± 14 (6)

67 ± 4 (12)

91 ± 7 (8)

270 ± 16 (12)

616 ± 102 (3)

147 ± 10 (6)

329 ± 17 (22)

314 ± 51 (1)

55 ± 8 (19)

280 ± 32 (4)

226 ± 21(3) 151 ± 68 (1)

129 ± 10 (5)

59 ± 2 (15)

57 ± 9 (19)

61 ± 2 (16)

Age components

TABLE A1. FISSION-TRACK ANALYTICAL DATA FOR ANDAMAN ISLANDS–SOUTH ANDAMAN ρd Nd

13.42 ± 0.23

13.24 ± 0.22

13.06 ± 0.20

13.25 ± 0.17

Mean track length (μm)

1.80

1.61

1.75

1.51

S.D.

62

54

75

76

No. of tracks

6 –2 Note: (i) Track densities are (x10 tr cm ); numbers of tracks counted (N) shown in brackets. (ii) Analyses by external detector method using 0.5 for the 4π/2π geometric correction factor. (iii) Ages calculated using dosimeter glass CN-5; (apatite) ξCN5 = 338 ± 4; CN-2 (zircon) ξCN2 = 127 ± 4 calibrated by multiple analyses of IUGS apatite and zircon age standards (see Hurford, 1990). (iv) Pχ2 is probability for obtaining χ2 value for v degrees of freedom, where v = no. crystals –1. (v) Central age is a modal age, weighted for different precisions of individual crystals (see Galbraith and Green, 1990). (vi) Age modes deconvolved using approach of Sambridge and Compston (1994) and Galbraith and Laslett (1993). RE—% relative error about the central age; S.D.—standard deviation; bold type—apatite; bold italic type—zircon ages.

N11°38.430′ E92°45.292′

AND-1A

ANDAMAN FLYSCH

N11°34.223′ E92°39.541′

Mungleton Quarry

Namunagarh Grit Formation

Hopetown Quarry

Hopetown Conglomerate Formation

MITHAKHARI GROUP

Sample

0.356

N11°38.430′ E92°45.292′

N11°38.457′ E92°45.318′

N11°38.459′ E92°45.368′

AND-A

AND-B

AND-C

0.018

0.021

0.023

HB (ncc)

0.204

0.188

0.333

He-HB (ncc)

4

0.119

0.119

0.452

5.48E + 09

5.05E + 09

8.95E + 09

4

He (atoms)

6.52E + 06

6.01E + 06

4.04E + 07

Absolute ±

0.058

0.062

0.108

238

U (ng)

1.98

2.18

1.2

± (%)

TABLE A2. RAW APATITE-HELIUM DATA %S.D. in Q±

0.315

0.215

0.537

Th (ng)

232

1.4

1.8

1.33

± (%)

0.819 0.962 0.886

11.704 13.756 12.664

Error ± 7%

He age* (Ma)

82

66

77

Grain radius

0.81

0.76

0.81

FT corrected

15.6

18.1

14.4

FT corrected age (Ma)

Note: All samples are taken from the Andaman Flysch Formation at Corbyn’s Cove. Helium ages are based on replicate analyses of apatite grains. The total uncertainty in sample age is based on the reproducibility of 39 analyses at the California Institute of Technology laboratory. Durango apatite standard (this error of the mean = 6.7%) combined with the U/Th and He analytical uncertainties. HB—hot blank; ncc—nanno cc; S.D.—standard deviation; FT—fission track. *Uncorrected He age.

0.222

0.209

He (ncc)

4

Location

Sample

A1B Ms 1 Ms 2 Ms 3 Ms 4 Ms 5 Ms 6 Ms 8 Ms 9 Ms 10 Ms 11 Ms 12 Ms 13 Ms 15 Ms 16 Ms 17 Ms 18 Ms 19 Ms 20 Ms 21 Ms 22 Ms 23 Ms 24 Ms 25 Ms 26 Ms 27 Ms 28 Ms 29 Ms 30 Ms 31 Ms 32 Ms 33 A3 Ms 1 Ms 2 Ms 3 Ms 4 Ms 5 Ms 6 Ms 8 Ms 9 Ms 10 Ms 11 Ms 12 Ms 13 Ms 15 Ms 16 Ms 17 Ms 18 Ms 19 Ms 20 Ms 21 Ms 22 Ms 23 Ms 24 Ms 25 Ms 26 Ms 27 Ms 28 Ms 29 Ms 30 Ms 31 Ms 32

40

Ar 0.68068 0.203759 0.451179 0.329659 0.37799 0.269821 2.277688 2.143176 1.384602 0.253605 0.259847 3.391288 0.790418 0.531821 0.323132 0.832588 0.176399 0.458665 0.269899 1.721992 1.216759 0.085958 0.130443 2.371162 0.314057 0.336883 0.999109 0.124323 1.604755 0.176028 0.094113 40

Ar 0.274897 0.444523 0.297818 1.30819 0.346651 3.632999 0.311219 0.281996 0.180333 0.518299 0.604087 0.514859 0.190225 0.138431 0.560451 1.927513 0.549605 0.199519 0.792813 0.531614 0.290937 0.481724 1.163152 0.272653 0.422343 0.24212 0.358268 0.178945 0.273151 0.173772

39

Ar 0.213559 0.060257 0.185698 0.110714 0.046614 0.037161 0.11746 0.070945 0.074126 0.015673 0.082319 0.385463 0.252174 0.070046 0.105295 0.098415 0.061786 0.086807 0.070387 0.356929 0.459022 0.027352 0.038038 0.106391 0.064276 0.064142 0.106174 0.075656 0.163177 0.107915 0.056227 39

Ar 0.083084 0.143237 0.112368 0.466538 0.128071 0.500069 0.100244 0.054435 0.044863 0.095269 0.186293 0.086648 0.061068 0.029423 0.085584 0.246058 0.188385 0.062808 0.159513 0.155877 0.092139 0.087635 0.352631 0.094083 0.128727 0.086183 0.11507 0.044807 0.093756 0.089489

TABLE A3. RAW ARGON DATA Andaman Flysch (Corbyn’s Cove section) 38 37 36 40 39 Ar Ar Ar Ar*/ Ar 0.003536 0.005284 0.000779 2.109962 0.000705 0.000626 6.48E-05 3.063559 0.002571 0.002597 0.000354 1.865821 0.001661 0.002407 0.000334 2.085147 0.000787 0.000994 0.00013 7.286482 0.000531 0.000573 5.98E-05 6.784923 0.001906 0.002831 0.000274 18.7013 0.00093 0.000402 2.99E-05 30.08465 0.001043 0.001053 0.00018 17.96253 0.000291 5.75E-05 0 16.18154 0.00117 0.000824 7.98E-05 2.870187 0.004783 0.003164 0.000389 8.499624 0.003255 0.003477 0.000284 2.801527 0.000772 0.000291 5.49E-05 7.360776 0.001329 0.001905 0.000189 2.537016 0.001318 0.000894 4.98E-05 8.310543 0.00071 0.000428 7.99E-05 2.472932 0.001068 0.00179 0.00013 4.842794 0.000792 0.000117 0.0002 2.994993 0.004472 0.009619 0.000637 4.296717 0.005575 0.00187 0.00031 2.451514 0.000404 0.000195 0 3.142686 0.000618 0.002262 0.000284 1.219882 0.001354 0.001404 0.000115 21.96896 0.000828 0.000507 4.99E-05 4.656836 0.000767 0.000156 8E-05 4.883813 0.001329 0.00041 0.000125 9.062499 0.000961 0.000488 7.49E-05 1.350842 0.001993 0.000528 0.000145 9.572115 0.001318 0.000411 8.49E-05 1.398714 0.000787 0.000782 5.48E-05 1.385834 38

Ar 0.000981 0.00185 0.0014 0.005703 0.001671 0.006087 0.001252 0.00068 0.000593 0.001186 0.002397 0.001109 0.000797 0.000409 0.001022 0.003455 0.002402 0.000767 0.00209 0.00209 0.001272 0.000905 0.004461 0.001191 0.00161 0.001145 0.001441 0.000537 0.001232 0.00115

37

Ar 0.00039 0.001424 0.000746 0.002087 0.000476 0.001156 0.000476 0.000511 0.000392 0.000528 0.000648 0.000392 0.000672 0.000534 6.89E-05 0.002792 0.00138 0.000897 0.002278 0.001657 0.000881 0.00038 0.001418 0.005552 0.000571 0.000242 0.000277 0.000211 0.000316 1.76E-05

36

Ar 0.000145 0.000445 0.000115 0.000109 0.000155 0.000305 0.000185 0.000105 2.49E-05 0.000125 0.000155 2.49E-05 0.000145 0.000115 6.5E-05 0.000814 0.000265 0.000185 0.000214 0.000385 7.98E-05 9.99E-05 0.00038 8.85E-05 0.00011 9.94E-06 2.99E-05 3.49E-05 0.000115 5E-05

40

39

Ar*/ Ar 2.793324 2.186142 2.348489 2.734713 2.349362 7.084947 2.55964 4.611138 3.855628 5.053064 2.997079 5.857044 2.414197 3.551304 6.324168 6.855692 2.502356 2.307352 4.573025 2.681443 2.901768 5.160096 2.980373 2.619921 3.028746 2.775297 3.036643 3.763259 2.551233 1.77673

± 0.020067 0.006231 0.022963 0.038311 0.032957 0.089994 0.031643 0.047492 0.023198 0.035317 0.003993 0.011701 0.02637 0.036057 0.031682 0.035372 0.067916 0.034779 0.047513 0.010036 0.006712 0.011197 0.023433 0.014042 0.012754 0.04699 0.029797 0.039184 0.020485 0.006655 0.005724

Age (Ma) 51.8 74.7 45.9 51.2 172.9 161.5 414.2 626.5 399.5 363.6 70.1 200.1 68.4 174.5 62.1 195.9 60.5 116.7 73.1 103.9 60.0 76.6 30.1 477.7 112.4 117.7 212.6 33.3 223.8 34.5 34.2

± 0.5 0.4 0.6 1.0 1.1 2.2 2.0 2.8 1.9 1.8 0.4 1.0 0.7 1.2 0.8 1.2 1.7 1.0 1.2 0.6 0.3 0.5 0.6 2.1 0.6 1.2 1.2 1.0 1.1 0.2 0.2

± 0.037215 0.029916 0.038715 0.013397 0.023267 0.012826 0.004776 0.010622 0.00766 0.006582 0.016331 0.034464 0.048452 0.003826 0.019494 0.06817 0.018307 0.05385 0.019797 0.009091 0.033228 0.036978 0.009667 0.032203 0.023467 0.03496 0.005754 0.033264 0.016583 0.016944

Age (Ma) ± 68.2 1.0 53.6 0.8 57.5 1.0 66.8 0.5 57.6 0.6 168.3 0.9 62.6 0.3 111.3 0.6 93.5 0.5 121.6 0.6 73.1 0.5 140.2 1.0 59.1 1.2 86.3 0.4 151.0 0.9 163.1 1.7 61.2 0.5 56.5 1.3 110.4 0.7 65.6 0.4 70.8 0.9 124.1 1.0 72.7 0.4 64.1 0.8 73.9 0.7 67.8 0.9 74.1 0.4 91.3 0.9 62.4 0.5 43.7 0.5 (continued)

A4 Ms 1 Ms 2 Ms 3 Ms 4 Ms 5 Ms 6 Ms 8 Ms 9 Ms 10 Ms 11 Ms 12 Ms 13 Ms 15 Ms 16 Ms 17 Ms 18 Ms 19 Ms 20 Ms 21 Ms 22 Ms 23 Ms 24 Ms 25 Ms 26 Ms 27 Ms 28 Ms 29 Ms 30 Ms 31 Ms 32 Ms 33

40

Ar 1.757257 0.350481 0.226229 0.84056 0.387886 0.827359 0.124755 3.009747 1.209192 0.497906 3.00958 0.239454 1.727941 0.259758 0.102435 0.565702 0.461503 0.084844 0.051231 0.539888 0.224278 2.435008 0.28137 1.447467 0.131921 0.381247 0.400761 0.124287 0.077119 0.042558 0.057102 40

Ms 1 Ms 2 Ms 3 Ms 4 Ms 5 Ms 6 Ms 8 Ms 9 Ms 10 Ms 11 Ms 12 Ms 13 Ms 15 Ms 16 Ms 17 Ms 18 Ms 19 Ms 20 Ms 21 Ms 22 Ms 23 Ms 24 Ms 25 Ms 26 Ms 27 Ms 28 Ms 29 Ms 30 Ms 31 Ms 32 Ms 33

Ar 0.8284261 0.7446774 0.7987535 0.3026851 0.6288437 0.2056404 2.0267770 1.3058958 1.0895856 1.7582128 0.7200124 0.7633280 1.4321915 3.1189393 0.5782068 1.2848422 0.5290525 5.1561196 0.6433217 1.1030049 0.3852341 0.8201835 0.2651809 0.4454907 0.2823107 7.5439887 0.4338841 0.9228281 2.4734857 1.5381052 0.3426377

39

Ar 0.228175 0.117885 0.086674 0.203536 0.07517 0.256024 0.077014 0.428003 0.187151 0.161033 0.046496 0.097718 0.097393 0.122808 0.029899 0.067938 0.030601 0.03198 0.028354 0.175027 0.125809 0.11556 0.043934 0.066026 0.043514 0.061408 0.030962 0.055912 0.024908 0.024805 0.020714 39

Ar 0.0365259 0.0323472 0.0472551 0.0938337 0.0690488 0.0084045 0.0181880 0.0589862 0.0467023 0.0745294 0.0307043 0.0327602 0.0604223 0.0694722 0.0518422 0.0599626 0.0217523 0.0246141 0.0339743 0.0956311 0.0186428 0.0345528 0.0818851 0.0287515 0.0124546 0.1794691 0.0171655 0.0225529 0.0295677 0.0256109 0.0155174

TABLE A3. RAW ARGON DATA (continued) Andaman Flysch (Corbyn’s Cove section) 38 37 36 40 39 Ar Ar Ar Ar*/ Ar 0.003051 0.001952 0.000479 7.080388 0.001625 0.000598 0.000135 2.63508 0.001165 0.000669 0.000115 2.218651 0.00254 0.000722 0.000305 3.687254 0.001007 0.000476 3.49E-05 5.023023 0.00326 0.001164 0.00039 2.781793 0.001012 0.000247 0.00011 1.198084 0.005417 0.000953 0.000195 6.897607 0.00232 0.000706 0.000105 6.29557 0.00184 0.000212 7.49E-05 2.954419 0.000608 0.000395 7.99E-05 64.22046 0.001129 0.000897 0.00013 2.058057 0.001247 3.59E-05 7.5E-05 17.51444 0.001441 0.000395 3.49E-05 2.0312 0.000307 0 1.5E-05 3.277664 0.000869 0.000198 3.49E-05 8.174691 0.000363 0 5E-06 15.03286 0.000358 3.6E-05 9.99E-06 2.560674 0.000327 0.000144 9.96E-06 1.70302 0.002034 0.000648 0.000225 2.705021 0.001584 0.000793 0.000115 1.513072 0.001303 0.000343 8.99E-05 20.84143 0.000537 0.000162 3E-05 6.202931 0.00092 0.001368 0.000185 21.09624 0.000777 0.001825 0.000255 1.303279 0.000899 0.001902 0.000264 4.935659 0.000603 0.001484 0.000155 11.46811 0.000761 0.001294 0.000205 1.141272 0.000368 0.000686 8.98E-05 2.030571 0.000409 0.000952 4.97E-05 1.123077 0.000225 0.000305 5.99E-05 1.901885 38

Ar 0.0004242 0.0003628 0.0006388 0.0011243 0.0008483 0.0001022 0.0002760 0.0007666 0.0006133 0.0008279 0.0004191 0.0004344 0.0007308 0.0008943 0.0006746 0.0007052 0.0003015 0.0003628 0.0003577 0.0011039 0.0002044 0.0004804 0.0009965 0.0004242 0.0001073 0.0021310 0.0002402 0.0003373 0.0004191 0.0003475 0.0001942

Mithakhari Group, Chiriyatapu 37 36 Ar Ar 0.0004699 0.0000000 0.0000000 0.0000550 0.0002743 0.0000249 0.0001764 0.0000750 0.0003334 0.0000449 0.0000000 0.0000000 0.0003637 0.0000349 0.0002830 0.0000949 0.0004044 0.0000849 0.0005462 0.0000899 0.0003844 0.0001499 0.0003442 0.0000499 0.0002229 0.0000499 0.0007095 0.0000998 0.0002434 0.0000349 0.0001623 0.0000750 0.0002639 0.0000199 0.0002639 0.0000799 0.0001625 0.0000500 0.0005283 0.0001499 0.0000000 0.0000000 0.0002349 0.0000499 0.0008358 0.0001048 0.0007315 0.0001248 0.0000000 0.0000000 0.0008891 0.0001048 0.0000000 0.0000000 0.0004192 0.0000299 0.0005765 0.0000398 0.0004719 0.0000000 0.0002623 0.0000299

40

39

Ar*/ Ar 22.6805064 22.5188268 16.7471471 2.9897197 8.9150267 24.4679238 110.8679577 21.6634575 22.7933251 23.2345815 22.0072636 22.8502553 23.4587880 44.4702431 10.9540815 21.0580001 24.0509065 208.5188362 18.5010347 11.0708902 20.6639642 23.3100383 2.8603315 14.2118131 22.6671453 41.8625264 25.2765779 40.5267640 83.2566784 60.0566236 21.5108920

± 0.0316 0.030076 0.038875 0.018742 0.044232 0.00436 0.007833 0.010471 0.006128 0.018552 0.067476 0.005784 0.014541 0.024229 0.007241 0.007531 0.049761 0.046289 0.116858 0.007126 0.003637 0.030878 0.034616 0.0327 0.01007 0.026829 0.056235 0.059711 0.013455 0.119683 0.160137

Age (Ma) 168.2 64.4 54.4 89.5 120.9 68.0 29.6 164.0 150.3 72.1 1144.9 50.5 390.6 49.9 79.8 192.8 340.1 62.7 41.9 66.1 37.3 456.1 148.2 461.0 32.2 118.9 265.1 28.2 49.9 27.7 46.7

± 1.1 0.8 1.0 0.6 1.2 0.3 0.2 0.8 0.7 0.6 4.3 0.3 1.8 0.6 0.4 0.9 1.9 1.2 2.9 0.4 0.2 2.1 1.1 2.1 0.3 0.8 1.7 1.5 0.4 2.9 3.9

± 0.0750144 0.0957791 0.0197808 0.0315933 0.0157014 0.0604556 0.1272294 0.0479489 0.0380826 0.0322423 0.1088272 0.0157553 0.0323211 0.0356104 0.0577358 0.0540885 0.1038536 0.2883512 0.0472177 0.0175933 0.0133670 0.0867000 0.0027919 0.0282976 0.0457513 0.0295576 0.0434406 0.1467725 0.1423053 0.1661772 0.0746433

Age (Ma) 491.3 488.2 375.1 72.9 209.3 524.8 1674.6 471.9 493.4 501.7 478.5 494.5 506.0 863.3 254.0 460.2 517.1 2444.9 410.2 256.5 452.6 503.2 69.8 323.1 491.0 822.6 539.8 801.4 1380.1 1089.0 469.0

± 2.6 2.8 1.7 0.8 1.0 2.5 5.6 2.3 2.3 2.3 3.0 2.2 2.3 3.5 1.7 2.3 3.0 6.9 2.1 1.3 2.0 2.7 0.3 1.6 2.3 3.3 2.5 4.0 5.1 4.7 2.5

8 21 18 28 1 37 7 13 31 16 23 2 9 5 3 35 4 29 12 39 26 6 22 11 32 25 34 10 40 19 27 33 17 30 24 20 38 15 36 14

Grain no.

238

Pb/ U ratio 0.56820 0.52114 0.45548 0.30919 0.25134 2.68537 0.20471 0.19499 2.04475 0.19076 0.19007 0.18013 0.17942 0.17514 0.15851 1.54569 0.15069 0.15001 0.13848 1.43337 0.12184 0.11741 0.09140 0.08908 0.70554 0.08390 0.54771 0.03673 0.16376 0.01567 0.01427 0.09256 0.01183 0.09413 0.00927 0.00910 0.05721 0.00778 0.13338 0.00705

206

0.01491 0.01876 0.02377 0.01417 0.01476 0.14191 0.00761 0.00490 0.18301 0.00770 0.00612 0.00766 0.00506 0.00495 0.00357 0.07373 0.00383 0.00490 0.00336 0.20932 0.00484 0.00687 0.00258 0.00363 0.09115 0.00418 0.03128 0.00221 0.10117 0.00179 0.00191 0.01678 0.00088 0.02893 0.00087 0.00077 0.00743 0.00052 0.08419 0.00024

±

235

Pb/ U ratio 18.27837 17.80391 10.17671 5.34072 3.17496 0.23801 2.27006 2.12506 0.19077 2.00881 2.22350 2.41204 1.81366 1.78122 1.66696 0.15281 1.56797 1.70213 1.34274 0.12528 1.12034 0.91179 0.76164 0.96494 0.08743 0.73573 0.06830 0.25590 0.02074 0.10372 0.10618 0.01397 0.07776 0.01041 0.06214 0.11849 0.00799 0.05499 0.00723 0.04889

207

0.00211 0.00319 0.00390 0.00228 0.00192 0.00394 0.00089 0.00101 0.00536 0.00087 0.00114 0.00081 0.00083 0.00078 0.00100 0.00239 0.00082 0.00107 0.00107 0.00491 0.00076 0.00108 0.00058 0.00048 0.00262 0.00076 0.00115 0.00024 0.00132 0.00025 0.00022 0.00047 0.00012 0.00056 0.00010 0.00010 0.00020 0.00005 0.00081 0.00004

±

206

Pb/ Pb ratio 0.23368 0.24751 0.16214 0.12476 0.09164 0.08188 0.08055 0.07915 0.07779 0.07645 0.08469 0.09716 0.07343 0.07384 0.07632 0.07342 0.07553 0.08191 0.07044 0.08302 0.06648 0.05641 0.06035 0.07869 0.05857 0.06343 0.05821 0.05061 0.05728 0.04802 0.05378 0.04809 0.04771 0.06564 0.04851 0.09449 0.05193 0.05135 0.13385 0.05037

207

0.01172 0.01869 0.01534 0.01304 0.01199 0.00394 0.00685 0.00497 0.00536 0.00752 0.00712 0.00881 0.00503 0.00488 0.00404 0.00239 0.00458 0.00734 0.00431 0.00491 0.00719 0.00919 0.00473 0.00822 0.00262 0.00948 0.00115 0.01065 0.00132 0.02354 0.03759 0.00047 0.01416 0.00056 0.01680 0.02208 0.00020 0.01135 0.00081 0.00528

±

238

Pb/ U Age (Ma) 2900.4 2704.0 2419.6 1736.7 1445.4 1376.4 1200.5 1148.4 1125.6 1125.5 1121.8 1067.7 1063.8 1040.3 948.5 916.7 904.8 901.0 836.1 760.9 741.2 715.6 563.8 550.1 540.3 519.4 425.9 232.5 132.3 100.2 91.3 89.4 75.8 66.7 59.5 58.4 51.3 49.9 46.5 45.3

206

61.3 79.5 105.3 69.8 76.0 20.5 40.7 26.5 29.0 41.7 33.2 41.8 27.7 27.1 19.8 13.4 21.5 27.5 19.1 28.2 27.8 39.6 15.2 21.5 15.6 24.8 6.9 13.7 8.4 11.3 12.1 3.0 5.6 3.5 5.5 4.9 1.3 3.3 5.2 1.6

±1σ

235

Pb/ U Age (Ma) 3004.5 2979.2 2451.0 1875.4 1451.1 1324.4 1203.0 1157.0 1130.5 1118.5 1188.5 1246.2 1050.4 1038.6 996.0 948.8 957.6 1009.3 864.4 903.0 763.1 658.0 575.0 685.8 542.1 559.9 443.5 231.4 154.0 100.2 102.5 89.9 76.0 91.3 61.2 113.7 56.5 54.4 127.1 48.5

207

TABLE A4. ZIRCON U-Pb DATA (UNCORRECTED AGES) FOR ANDAMAN FLYSCH SAMPLE 1A

46.4 72.0 83.7 88.2 95.3 39.1 56.6 42.0 61.0 63.8 58.4 60.4 42.7 40.6 31.7 29.4 35.5 58.2 34.2 87.3 57.1 75.1 34.3 49.1 54.3 62.5 20.5 41.9 88.3 45.6 67.0 15.6 21.1 26.9 19.9 23.6 7.1 11.2 75.4 4.8

±1σ

206

Pb/ Pb Age (Ma) 3077.4 3168.9 2478.1 2025.4 1459.9 1242.7 1210.5 1175.9 1141.5 1106.9 1308.4 1570.4 1025.9 1037.0 1103.5 1025.7 1082.6 1243.4 941.1 1269.6 821.5 467.6 616.3 1164.5 551.3 722.7 537.1 223.1 501.8 99.0 361.9 103.5 84.0 795.0 124.2 1518.0 282.3 256.6 2149.0 212.2

207

77.9 114.9 151.3 174.4 230.3 99.7 158.8 119.3 172.8 184.9 154.9 160.9 132.6 128.1 102.3 94.7 117.0 166.2 120.5 266.2 210.9 325.9 160.8 194.1 264.1 288.8 122.9 425.1 982.4 874.6 1092.3 384.6 585.7 545.5 660.0 386.4 275.3 441.6 828.8 225.9

±1σ

Namunagarh Grit Modern river sand

Mungleton Quarry N11°34′22.3″ E92°39′54.1″

Irrawaddy River ~ Nyaungdoun

Myanmar

S Andaman

S Andaman

S Andaman

S Andaman

S Andaman

S Andaman

MY05 23A

NAM 26

NAM 25A

NAM 3D

FT4

FT3

FT1B

—

1800

850

1500

130

140

115

TABLE A5A. PETROGRAPHIC DATA Region Sample Grain size (μm)

47

19

21

2

47

48

50

Q

11

1

2

1

10

17

13

KF

11

30

27

41

9

11

5

P

6

45

40

52

5

3

4

Lv

0

0

0

0

0

0

0

Lc

6

1

2

0

3

2

1

Lp

1

1

2

0

1

0

0

Lch

15

3

5

1

24

16

25

Lm

2

0

0

0

1

2

1

M

1

0

1

2

0

1

1

HM

100.0

100.0

100.0

100.0

100.0

100.0

100.0

Total

Note: The MI (“Metamorphic Index”; Garzanti and Vezzoli, 2003) expresses the average rank of metamorphic rock fragments in the studied samples, and varies from 0 in detritus from sedimentary and volcanic cover rocks to 500 in detritus from high-grade basement rocks. Q—quartz; KF—potassium feldspar; P—plagioclase; Lv—volcanic; Lc—carbonate lithic fragments (including marble); Lp—terriginous lithic fragments (shale, siltstone); Lch—chert lithic fragments; Lm—metamorphic lithic fragments; M—micas; HM—heavy minerals.

Namunagarh Grit Namunagarh Grit

Hopetown Quarry N11°41′40.7″ E92°43′50.1″

Andaman Flysch

Corbyn’s Cove section below hotel

Namunagarh Quarry

Andaman Flysch Andaman Flysch

Corbyn’s Cove section ~100 m upsection from 1A

Formation

Corbyn’s Cove section N11°38′43.0″ E92°45′29.2″

S it e

112

11

15

2

150

160

175

MI

Formation

Namunagarh Grit Namunagarh Grit

Namunagarh Quarry

Hopetown Quarry N11°41′40.7″ E92°43′50.1″

Mungleton Quarry N11°34′22.3″ E92°39′54.1″

River sand

Namunagarh Grit

Corbyn’s Cove section below hotel

Irrawaddy River ~ Nyaungdoun

Andaman Flysch Andaman Flysch

Corbyn’s Cove section ~100 m upsection from 1A

Andaman Flysch

Corbyn’s Cove section N11°38′43.0″ E92°45′29.2″

Site

Namunagarh Grit River sand

NAM 25A

Namunagarh Grit

Hopetown Quarry N11°41′40.7″ E92°43′50.1″

Mungleton Quarry N11°34′22.3″ E92°39′54.1″

Irrawaddy River ~ Nyaungdoun

NAM 3D

Namunagarh Grit

Namunagarh Quarry

FT3

Corbyn’s Cove section below hotel

MY05 23A

NAM 26

NAM 25A

NAM 3D

FT4

FT3

FT1B

Sample

MY05 23A

NAM 26

FT4

Andaman Flysch Andaman Flysch

Corbyn’s Cove section ~100 m upsection from 1A

FT1B

Sample

35 37 42 53 79

1.5 0.6 0.5 4.0

5.0

HM% vfs–fs 0

0

0

0

0

0

0

1.0

1.3

4.0

0.6

0.4

0.9

0

0

0

0

0

0

0

1

3

0

24

30

24

44

0.2 0.2

Anatase

0

3

54

HM% transparent Brookite

0

100

17

0

3

1

3

11

17

10

0

0

0

0

0

0

0

0

0

0

0

0

1

0

0

0

0

0

0

0

0

0

0

0

0

0

1 100

40

42

100

7

16

0

0

0

0

0

0

0

0

0

0

0

44

0

0

1

0

0

0

0

0

0

2

0

0

2

0

0

0

1

1

0

3

16

15

1

0

2

3

0

0

0

0

0

0

2 1 (continued)

0

0

20

0

0

0

0

0

0

1

2

100

7

5

25

6

5

38

3

4

0 4

5 4

4

100

4

100

3

100

57

49

37

8

7

9

% transparent Ti aggregates

0.5

% opaque Apatite

Xenotime

Formation

% turbid Monazite

Total Barite

Andaman Flysch

Dravite Others

S it e

Schorlite Blue-green hornblende

Corbyn’s Cove section N11°38′43.0″ E92°45′29.2″

Rutile Green hornblende

Zircon

TABLE A5B. HEAVY-MINERAL DATA

Sagenite Green-brown hornblende

Sphene Brown hornblende

Formation

NAM 25A

Namunagarh Grit Namunagarh Grit

Hopetown Quarry N11°41′40.7″ E92°43′50.1″

Mungleton Quarry N11°34′22.3″ E92°39′54.1″

Namunagarh Grit Namunagarh Grit

Namunagarh Quarry

Hopetown Quarry N11°41′40.7″ E92°43′50.1″

Mungleton Quarry N11°34′22.3″ E92°39′54.1″

Note: HM—heavy minerals; vfs—very fine sand; fs—fine sand.

River sand

Namunagarh Grit

Corbyn’s Cove section below hotel

Irrawaddy River ~ Nyaungdoun

Andaman Flysch Andaman Flysch

Corbyn’s Cove section ~100 m upsection from 1A

MY05 23A

NAM 26

NAM 25A

NAM 3D

FT4

FT3

FT1B

Sample

Formation Andaman Flysch

MY05 23A

NAM 26

River sand

Corbyn’s Cove section N11°38′43.0″ E92°45′29.2″

Site

Irrawaddy River ~ Nyaungdoun

NAM 3D

Namunagarh Grit

Namunagarh Quarry

FT3

Corbyn’s Cove section below hotel

FT4

Andaman Flysch Andaman Flysch

Corbyn’s Cove section ~100 m upsection from 1A

FT1B

Sample

0 0

0 0

0

0

0

0

0

0

0

1

0

2

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0 0

0 0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

Zoisite

Andaman Flysch

Oxyhornblende Other epidotes

Corbyn’s Cove section N11°38′43.0″ E92°45′29.2″

Glaucophane Prehnite

Site

Sodic amphiboles

TABLE A5B. HEAVY-MINERAL DATA (continued)

Tremolite Pumpellyite

Chloritoid 0

0

0

0

25

12

12

Actinolite 0

0

0

0

0

0

0

Lawsonite 0

0

0

0

0

0

0

Antofillite 2

0

0

0

0

0

0

Carpholite 0

0

0

0

0

0

0

0

0

0

0

0

0

Garnet 6

1

0

0

15

13

14

Other amphiboles 0

0

0

0

0

0

0

0

0

0

0

0

1

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

0

100

100

100

100

100

100

100

0

0

0

0

0

0

0

33

44

43

11

0

2

18

0

7

16

0

2

3

2

Spinel Kyanite

Hypersthene Staurolite

Epidote Sillimanite

Olivine Andalusite

Clinozoisite Total transparent

Corbyn’s Cove

AF2

Thanlwe River

Thandwe River

Irrawaddy River

MY05 15A

MY05 17B

MY05 22B

MY05 23B

N18°48.391′, E95°12.218′

N18°27.466′, E94°23.563′

N18°59.097′, E94°15.244′

N17°57.139′, E94°33.087′

MR mud

MR mud

MR mud

MR sand

MR sand

Sandstone

Mudstone

Mudstone

Sandstone

Lithology

3.67

3.60

3.48

3.33

4.02

5.05

11.17

1.99

4.83

Sm

17.9317

18.0331

17.6156

16.4129

20.0489

25.7129

50.6900

7.18860

25.3134

Nd

TABLE A6A. WHOLE ROCK Sm-Nd DATA

N20°49.212′, E93°18.576′

Andaman Flysch

Andaman Flysch

Namunagarh Grit

Namunagarh Grit

Formation / GPS

Note: MR—modern river sample; GPS—Global Positioning System.

Lemyu River

Kyeintuli River

MY05 8A

Myanmar—Rivers draining Eocene rocks

Corbyn’s Cove

AF1A

Namunagarh Quarry

Namunagarh Quarry

AN05 3E

Location

AN05-31

Andaman Islands

Sample

0.2052

0.1997

0.1977

0.2035

0.2010

0.1966

0.2204

0.2769

0.1909

Sm/Nd

0.124

0.1207

0.1195

0.1250

0.1215

0.1188

0.1332

0.1674

Nd

144

0.1154

Sm/

147

Nd

144

0.512210

0.512424

0.512427

0.512434

0.512300

0.512220

0.512068

0.512799

0.512271

Nd/

143

18

14

18

14

18

12

10

19

16

1σ (ppm)

–8.3

–4.2

–4.1

–4.0

–7.4

–8.2

–11.1

3.1

–7.2

εNd

Andaman Flysch AF1A Corbyn’s Cove

8

Andaman Flysch AF1A Corbyn’s Cove

24

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

Apatite grain

0.172124

0.244791

0.131323

0.136207

0.082935

0.102799

0.085210

0.117803

0.105069

0.114968

0.067596

0.078972

0.120108

0.113837

0.099159

0.108342

0.093898

0.106736

0.100785

0.082561

0.073481

0.090724

0.070462

0.123272

0.000850

0.003438

0.000741

0.000550

0.000202

0.000424

0.000165

0.000238

0.001427

0.000317

0.000236

0.000114

0.000750

0.000346

0.000388

0.001112

0.000488

0.000552

0.002208

0.000180

0.000171

0.000581

0.000408

0.001140

TABLE A6B. SINGLE-GRAIN APATITE Sm-Nd DATA 147 144 Lithology 2 S.D. of NIST610 Sm/ Nd

0.000079

0.000076

0.000062

0.000118

0.000136

0.000051

0.000056

0.000055

0.000067

0.000043

0.000075

0.000067

0.000044

0.000051

0.000033

0.000047

0.000033

0.000046

0.000062

0.000031

0.000054

0.000058

0.000065

0.000069

2σ (internal)

Note: Global Positioning System (GPS) reference for AF1A Corbyn’s Cove: N11°38.430' E92°45.292'; S.D.—standard deviation.

Andaman Flysch AF1A Corbyn’s Cove

Andaman Flysch AF1A Corbyn’s Cove

22

Andaman Flysch AF1A Corbyn’s Cove

21

23

Andaman Flysch AF1A Corbyn’s Cove

Andaman Flysch AF1A Corbyn’s Cove

19

Andaman Flysch AF1A Corbyn’s Cove

18

20

Andaman Flysch AF1A Corbyn’s Cove

Andaman Flysch AF1A Corbyn’s Cove

16

17

Andaman Flysch AF1A Corbyn’s Cove

Andaman Flysch AF1A Corbyn’s Cove

14

15

Andaman Flysch AF1A Corbyn’s Cove

Andaman Flysch AF1A Corbyn’s Cove

12

13

Andaman Flysch AF1A Corbyn’s Cove

Andaman Flysch AF1A Corbyn’s Cove

7

11

Andaman Flysch AF1A Corbyn’s Cove

6

Andaman Flysch AF1A Corbyn’s Cove

Andaman Flysch AF1A Corbyn’s Cove

5

Andaman Flysch AF1A Corbyn’s Cove

Andaman Flysch AF1A Corbyn’s Cove

4

10

Andaman Flysch AF1A Corbyn’s Cove

3

9

Andaman Flysch AF1A Corbyn’s Cove

Andaman Flysch AF1A Corbyn’s Cove

1

Formation / Age

2

Sample Nd

144

0.512771

0.512607

0.512791

0.512192

0.512486

0.512782

0.512321

0.512607

0.512792

0.511822

0.512425

0.512291

0.512621

0.512794

0.512669

0.512545

0.511800

0.512211

0.512404

0.512485

0.512588

0.512237

0.512803

0.512855

Nd/

143

154

148

122

230

266

99

109

108

131

83

147

131

86

99

64

93

65

90

121

61

105

114

127

134

2σ ppm

2.59

–0.61

2.99

–8.71

–2.96

2.80

–6.19

–0.60

3.01

–15.92

–4.16

–6.77

–0.33

3.03

0.61

–1.81

–16.35

–8.33

–4.57

–2.98

–0.98

–7.83

3.21

4.24

εNd (0)

254

Allen et al.

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Printed in the USA

The Geological Society of America Special Paper 436 2008

Post-collisional collapse in the wake of migrating arc-continent collision in the Ilan Basin, Taiwan Peter D. Clift School of Geosciences, University of Aberdeen, Aberdeen, AB24 3UE, UK, and DFG-Research Centre Ocean Margins (RCOM), Geowissenschaften, Universität Bremen, Klagenfurter Strasse, 28359 Bremen, Germany Andrew T.S. Lin Department of Earth Sciences, National Central University, 300, Jungda Road, Jungli, Taiwan Andrew Carter School of Earth Sciences, University and Birkbeck College London, Gower Street, London, WC1E 6BT, UK Francis Wu State University of New York at Binghamton, Department of Geological Sciences and Environmental Studies, P.O. Box 600, Binghamton, New York 13902-6000, USA Amy E. Draut U.S. Geological Survey, Pacific Science Center, 400 Natural Bridges Drive, Santa Cruz, California 95060, USA T.-H. Lai L.-Y. Fei Central Geological Survey, 2, Lane 109, Hua-Hsin Street, Chung-Ho, Taipei, Taiwan 235 Hans Schouten Department of Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts 02543, USA Louis Teng Department of Geosciences, National Taiwan University, Roosevelt Road, Taipei, Taiwan

ABSTRACT The Ilan Basin of northern Taiwan forms the western limit of the Okinawa Trough, where the trough meets the compressional ranges of central Taiwan. Apatite fissiontrack ages of 1.2 ± 0.5 Ma and 3.5 ± 0.5 Ma, measured north and south of the basin, respectively, indicate faster exhumation rates in the Hsüehshan Range to the north (>1.6 mm/yr) than in the Backbone Range to the south (0.7 mm/yr). Reconstructed subsidence rates along the northern basin margin are also faster than in the south (6–7 compared with 3–5 mm/yr). Global positioning system (GPS) and active seismological data indicate motion of the southern basin margin to the east and southeast. Clift, P.D., Lin, A.T.S., Carter, A., Wu, F., Draut, A.E., Lai, T.-H., Fei, L.-Y., Schouten, H., and Teng, L., 2008, Post-collisional collapse in the wake of migrating arc-continent collision in the Ilan Basin, Taiwan, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 257–278, doi: 10.1130/2008.2436(12). For permission to copy, contact editing@geosociety. org. ©2008 The Geological Society of America. All rights reserved.

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Clift et al. We propose that the Ilan Basin is being formed as a result of extension of northern Taiwan, largely controlled by a major southeast-dipping fault, modeled at ~30° dip, and mapped as a continuation of the Lishan Fault, a major thrust structure in the Central Ranges. Flexural rigidity of the lithosphere under the basin is low, with elastic thickness ~3 km. A southwest-migrating collision between the Luzon Arc and southern China, accompanied by subduction polarity reversal in the Ryukyu Trench, has allowed crustal blocks that were previously held in compression between the Eurasian and Philippine Sea plates to move trenchward as they reach the northern end of the collision zone. Subduction polarity reversal permits rapid extension and formation of the Ilan Basin and presumably, at least, the western Okinawa Trough, as a direct consequence of arc-continent collision, not because of independent trench rollback forces. This conceptual model suggests that migrating arc-continent collision causes the rapid formation of deep marginal basins that are then filled by detritus from the adjacent orogen, and that these should be common features in the geologic record. Keywords: collision, extension, erosion, subduction, seismology.

INTRODUCTION The Earth’s continental crust is generally considered to be largely generated by magmatism along active plate margins, a process that balances the long-term loss of continental crust back into the mantle through subduction zones (e.g., von Huene and Scholl, 1991; Rudnick and Fountain, 1995; Clift and Vannucchi, 2004). This balance is maintained because arc crust generated in oceanic subduction settings is not entirely subducted when these features collide with other trench systems or passive continental margins. Instead, continental masses appear to be built up as a result of the progressive accretion of primitive arc blocks to older, preexisting continental blocks. Understanding the tectonism of arc-continent collision is important, not only because this process is fundamental to the conservation of the total volume of continental crust, but also because it is likely the most important method by which active continental margins are created (e.g., Casey and Dewey, 1984; Konstantinovskaya, 1999). Accreted oceanic island-arc terranes have been recognized in a number of orogenic suture zones (e.g., Kohistan in the Himalaya, Talkeetna in southern Alaska, South Mayo in the Irish Caledonides), and the process of arc-continent collision forms an inevitable part of the plate tectonic cycle. Arc-continent and continentcontinent collisions are unlikely to be identical because of the different mechanical properties of arc and cratonic lithospheres. In this paper we investigate the later stages of arc-continent collision, specifically the collapse of the collisional mountain belts to form sedimentary basins whose fill might be used to understand the collision process. In order to understand the collapse of arc collisional mountains, we examine the Ilan Basin of northern Taiwan. Taiwan, located off the coast of southeastern China, comprises some of the tallest and most rapidly uplifting and eroding mountains in East Asia. The orogen is formed by the collision between the oceanic Luzon Arc and the passive margin of southern China (Suppe,

1981; Chemenda et al., 1997), which is itself built on the remains of a Cretaceous continental arc (Davis et al., 1997). The collision between the Luzon oceanic arc and mainland Asia is oblique, so that the collision point must have migrated along the Asian continental margin through time. At any given time the north end of the island will represent arc units and Chinese passive-margin sequences that began to be involved in collision earlier than those in the south, where collision is just starting. Thus northern Taiwan is the optimal place to examine the processes that occur during and after the peak of arc-continent collision. These processes form the subject of this paper. Radiometric dating of the metamorphic rocks exposed in the Central Ranges, combined with biostratigraphic dating of orogenic sediment from the Coastal Ranges, indicates that collision of this part of the Luzon Arc with mainland Asia started ca. 6–9 Ma (e.g., Suppe, 1981; Teng, 1990; Sibuet et al., 2002; Malavieille et al., 2002; Huang et al., 2006). A wide consensus in the community agrees that Taiwan represents a simple collision between the Luzon Arc and mainland Asia, an assumption that underlies this present study. Nonetheless, we note that alternative models have been advanced to explain the mountain building. One theory involves collision of an exotic terrane with the Chinese passive margin, followed by later collision of the Luzon Arc (Lu and Hsü, 1992). Another model invokes compressional tectonism following collision of the Luzon Arc with the Ryukyu Arc, which would previously have extended farther southwest along the southern margin of China (Hsu and Sibuet, 1995; Sibuet and Hsu, 1997). It is unclear when the collision between Luzon and Eurasia began. Whereas some models favor collision to have initiated only ca. 6–9 Ma, effectively just to the east of the present collision zone (e.g., Suppe, 1984; Sibuet and Hsu, 1997; Huang et al., 2000, 2006), other models suggest a more steady-state collision that may have started earlier (e.g., Suppe, 1984; Teng, 1990, 1996; Clift et al., 2003). In the steady-state model the modern

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of the free edge provided by the new trench. Specifically, there is no need for slab break-off to allow polarity reversal or as a trigger of orogenic extension.

collision represents a snapshot of a continuous collision process that is migrating to the southwest along the Chinese margin, accreting at least part of the crust of the Luzon Arc to the edge of Eurasia, while generating a new Ryukyu Arc-Trench system and associated Okinawa Trough in its wake. Debate continues as to whether the Okinawa Trough is an active rift system formed by trench tectonic forces linked to the Ryukyu subduction zone and propagating to the southwest into the northern tip of the Taiwan orogen (e.g., Suppe, 1984; Sibuet et al., 1998; Wang et al., 1999), or whether it represents the product of gravitational collapse of the Taiwan orogen, following subduction polarity reversal (e.g., Teng, 1996; Clift et al., 2003). In the latter case, extension of the Ilan Basin has occurred because of the reversal of subduction polarity following arc-continent collision, which removes the stresses that support the high topography in central Taiwan. The earliest phases of extension form the Ilan Plain Basin. The Okinawa Trough represents the continuation of this extension to form a deep marine basin offshore. Without the tectonic push from the Philippine Sea plate the orogen is able to extend because

REGIONAL GEOLOGY The Ilan Basin forms the onshore western limit of the Okinawa Trough and lies in northeast Taiwan (Figs. 1 and 2). The basin is supplied with sediment from the southwest via the Lanyang River. This river terminates in a small delta, facing into the deeper waters of the Okinawa Trough. To the north the basin is bounded by exposures of shale and sandstone of the Meichi, Szeleng, and Chiayang Formations, which are dated as Oligocene, and typically interpreted to be part of the passive margin of China (Suppe, 1981; Teng, 1990; Lundberg et al., 1997), as well as minor amounts of the underlying Eocene Tachien Sandstone, which represents a synrift deposit, predating Oligocene seafloor spreading (Fig. 3). The whole sedimentary sequence was deformed and thrust to the northwest during the Taiwan orogeny, forming the

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Figure 2. Shaded topographic map of the Ilan Plain area, showing the location of the drill sites considered in this study, the position of apatite fission-track samples (with average age of the dominant population), and the orientation of topographic profiles shown in Figure 5. Map is from NASA Shuttle Radar data, made available through www.geomapapp.org.

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modern Hsüehshan Range. Peak metamorphic grades are low (e.g., Chen, 1984). Liu et al. (2001) used zircon fission-track data to show that these units had not been buried beyond the partial annealing temperature for that mineral (~200 °C; Tagami et al., 1998) and inferred a lower greenschist metamorphic facies. The southern edge of the basin is marked by outcrops of the shale-rich Miocene Lushan Formation, which Liu et al. (2001) considered to have reached prehnite-pumpellyite facies, and thus is of a slightly lower grade than the Hsüehshan Range. The Lushan Formation makes up much of the Backbone Range and unconformably overlies the Tananao Schist locally (Suppe et al., 1976), which is formed largely of Paleozoic and Mesozoic sedimentary rocks, metamorphosed to greenschist grade. The oldest sedimentary rocks overlying the Tananao Schist are Eocene conglomerates of the Pilushan Formation, but the onlap is time transgressive between basin and highs on the paleo-Chinese margin. Localized slices of higher grade Tananao metamorphic rocks, including mafic amphibolites, are well exposed on the coast at Tungao, south of Suao (Chen and Jahn, 1998). The Tananao Schist is dated to have been metamorphosed during the

Mesozoic and subsequently overprinted by greenschist conditions in the Pliocene (Lo and Yu, 1996; Wang et al., 1998). Southwest of the Ilan Basin the Lushan and Tatungshan Formations are juxtaposed across the Lishan Fault. In central Taiwan the Lishan Fault is manifested as a major, steep, NW-dipping thrust structure, having an ESE-vergent geometry, interpreted by Clark et al. (1993) as an antithetic backthrust relative to the dominant northwest-vergent thrust stack in the Central Ranges (Wu, 1978). Changes in stratigraphic thickness across the fault suggest that it is a reactivated structure that originally formed within an extensional rift structure on the South China Sea passive margin (Teng et al., 1991). NEOTECTONICS OF THE LISHAN FAULT Although the Lishan Fault is a NW-dipping thrust in central Taiwan, it can be readily traced northward to the SW apex of the Ilan Plain. In this region the fault is overturned and has an extensional sense of motion at the western limit of the Ilan Basin, meaning that the slip sense remains the same. The Lishan Fault is

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such a major tectonic division that it can be readily traced to pass from being a thrust in central Taiwan to an extensional fault in the Ilan Plain (Teng and Lee, 1996). In the western Ilan Basin the Lishan Fault forms a normal fault of relatively high angle (>60°) that separates the Hsüehshan Range from the Pleistocene basin fill (Teng and Lee, 1996). The topographic break along the northern edge of the basin is commonly sharp and suggestive of an active neotectonic structure (Fig. 4A; Shyu et al., 2005). Exposures of the sedimentary rocks that form the Hsüehshan Range

are relatively fresh and prominent at the northern edge of the basin along much of its length. Figure 4B shows an example of well-bedded, relatively unweathered sedimentary rocks, exposed close to the trace of the fault near the southwestern end of the basin. In this area the southern boundary of the basin is also sharp and consistent with neotectonic faulting, as shown in map view (Fig. 2) and in topographic profiles drawn across the strike of the basin (Fig. 5). However, the shaded topographic map (Fig. 2) also shows that the southern basin margin is less linear and poorly

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Figure 4. Field photographs. (A) Steep escarpment along the northeast edge of the Ilan Plain, testifying to the faultbounded nature of the boundary. (B) A relatively fresh exposure of sandstone and shale exhumed by the Lishan Fault at the western end of the Ilan Plain, testifying to the recent activity of the fault along the northern edge of the basin.

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defined in comparison with the northern boundary. Exposures along the southern basin margin are generally not fresh but show strong deep weathering of the basement, inconsistent with recent tectonic exhumation. The eastern part of the plain in particular shows a strong topographic contrast between a gently sloping southern margin and a steep northern margin (Fig. 5). This general contrast between the northern and southern expressions of the Lishan Fault suggests that the northern boundary is mostly fault controlled. However, the relationship is less clear in the south and is more consistent with erosion of a tilted hanging block rather than an uplifted footwall block. Although Shyu et al. (2005) proposed a neotectonic fault along the easternmost part of the southern boundary, it is noteworthy that this feature has little topographic expression, suggesting that motion on this structure is rather less than on the Lishan Fault. Sharp increases in topographic gradient within the Backbone Range south of the basin boundary nonetheless indicate that faulting is active in this region. Although the eastern and western ends of the northern margin of the Ilan Basin are well defined, and apparently fault-controlled, these trends do not align with one another but instead are separated by a section ~10 km wide with less topographic relief (Fig. 2). We infer a relay section in the Lishan Fault extension in this area,

where motion is transferred laterally. As is often the case in fault relay zones in rifts, this area is marked by rivers that enter the basin laterally, adding to the sediment provided by the axial Lanyang River (e.g., Leeder and Gawthorpe, 1987). DRILLING CONSTRAINTS ON NEOTECTONISM Drilling by the Central Geological Survey of Taiwan has allowed the basement contact along the faulted northern edge of the Ilan Basin to be determined at depth as well as at outcrop. Cores from three sites were examined in this project, wells TC, TL, and CH, trending east to west close to the northern edge of the basin (Fig. 2). These wells intersect the basement at depths of 17, 84, and 38 m, respectively, suggesting an average dip of the basement-cover interface under the northern edge of only ~10°, although this is over a short distance and may not be representative of the dip over the whole basin width. The overlying sediment comprises sequences of alluvial, coastal, and some shallow marine sediment that has been dated as late Pleistocene to Holocene by AMS 14C methods (Lai and Hsieh, 2003). The sediment recovered at well CH includes fine-grained muds, silts, and fine sands but is dominated by thick-bedded debris-flow gravels, with clasts up to 10 cm across (Fig. 6). Well TL recovered the finest

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Figure 6. Sedimentary logs showing the cored sequences found along the northern edge of the Ilan Plain, where Pleistocene sediment is deposited over low-grade metamorphosed sedimentary rock, commonly below a zone of fault breccia that marks a detachment surface.

Migrating arc-continent collision in the Ilan Basin, Taiwan grained sand of the three drilled sequences, though the sediment is still characterized by coarse sands and gravels. Well TC recovered a sequence of medium- to thick-bedded sands and muds. In each location the sediment is underlain by well-lithified, lowgrade metamorphosed shale, typical of that seen in outcrop in the Hsüehshan Range. All drilled basement samples are distinguished by the presence of discrete brecciated layers, generally 2–5 m thick at the basement contact but also present within the basement below that level. These breccias are monomict and comprise only fragments of metamorphosed shale (Fig. 7). The lack of material in the breccias from the shallower basin fill sediment probably reflects the fact that they are loose, unconsolidated materials and unable to support the confining stresses needed to be incorporated into the gouge. Texturally the breccias are generally very angular, unsorted, and matrix supported. The breccias are relatively well lithified in comparison with the overlying soft Pleistocene sediment. We interpret these rocks to represent fault breccias formed by cataclasis in the shallow subsurface of the Lishan Fault. The brittle, angular, broken character of the fault-breccia clasts is apparent in thin section. Critically, the fault breccias postdate the development of cleavage in the metapelites that form the footwall (Fig. 8). As a result, motion on the fault is constrained to being post–peak metamorphism and relatively shallow (<8 km) in the crust. The breccias are preferentially formed close to the basement contact, because this is the fault plane exposed to the surface by the exhumation of the footwall block. Not all the breccia and fault layers lie exactly along the basement-sediment contact, indicating that the Lishan Fault has been active along several fault splays within the footwall block during uplift of the Hsüehshan Range. The drill data confirm the presence of a major southeast-dipping structure that controls the basin’s northern margin (Fig. 9).

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Because the sediment cored in the Ilan Basin is all close to sea level and preserves evidence of the environment of its sedimentation, these materials can be used to determine rates of tectonic subsidence across the plain. In the Ilan Basin the fast rate of the sediment supply, largely from the Lanyang River, appears generally to match rates of basement subsidence, keeping the basin full while allowing shoreline progradation toward the Okinawa Trough. Lai and Hsieh (2003) demonstrated that subsidence reaches a maximum rate of ~19 mm/yr around the mouth of the Lanyang River and decreases toward the edges of the plain, both north and south, as well as toward the west. However, what is important toward understanding the tectonics of the Ilan Basin is the recognition that the northern part of the plain subsides faster than the southern sector, reaching average Holocene rates of 6–7 mm/yr versus 3–5 mm/yr. This disparity is consistent with the observations presented above that indicate a dominant asymmetric character to the basin structure. We thus favor basin formation above a dominant south-dipping extensional detachment (Fig. 9). MODELING BASIN GEOMETRY The large-scale structure of the basin can be considered and modeled using a theoretical forward of an asymmetric basin, controlled by a dominant detachment. We employed the flexural cantilever model of Kusznir et al. (1991) to forward model the deformation and subsidence that would be expected to result from the extension measured across the normal fault that bounds the northern edge of the basin, and to test the idea that the regional dip is 10°, as inferred from the drilling. Although the flexural cantilever model is not the only, or universally accepted, theoretical model for extensional basins, it does have a number of

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Figure 7. Photographs of cored prePleistocene rocks from well CH in the western Ilan Basin. (A) Fault-brecciated zone at 75.78 m depth. (B) Coherent block of metapelite from 84.50 m depth.

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Figure 8. Photomicrographs of thin sections cut from basement rocks in well CH. (A) Typical fine-grained character of the basement metapelites, with a well-developed cleavage. (B) Fault rock from within one of the breccia zones, demonstrating low-temperature shattering textures.

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important features that allow the basin tectonics to be understood at a first-order level, and it has been used to effect in understanding basins in several parts of the world (Kusznir et al., 1995; Roberts et al., 1993). The model is complicated by our lack of subsurface data and constraint on the degree of extension across the fault, although this is also limited by the total size of the basin. For modeling purposes we chose a default of 5 km horizontal extension and a crustal thickness of 40 km. In this approach we made no attempt to replicate the basin morphology but simply extended a model

continental lithosphere using the major detachment fault and predicted what sort of basin this would form. Flexural rigidity is expected to be low in an arc orogenic setting. Effective elastic thickness (Te) is only 13 km in the Taiwan foreland (Lin and Watts, 2002) and is expected to be rather less in the orogenic core that we consider here. The deformation in the flexural cantilever model assumes brittle faulting in the upper crust, set at 10 km thickness here, and ductile deformation distributed in a sinusoidal fashion over a wavelength of 100 km. The results of the models are shown in Figure 10.

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Figure 10. Proposed basin geometries and comparisons with possible model basins in order to limit the possible faulting geometries and strength of the lithosphere under the Ilan Plain. (A) Proposed basin cross section inferred from outcrop, coring, and gravity constraints. (B–D) Forward models, made using the flexural cantilever model of Kusznir et al. (1991) and a 30° detachment fault, with variations in the flexural rigidity of the lithosphere. (E–G) Forward models using a 10° detachment fault, showing the effects of varying amounts of extension and flexural rigidity. (H) Forward models using a 45° detachment fault, showing a poor match to observed basin shape. Te—elastic thickness.

An important result of our forward models is that basins generated above a 10° detachment surface do not produce realistic basin geometries (Fig. 10A, E–G), mostly because they do not predict significant footwall uplift of the size seen in the Hsüehshan. Also, 10° detachments fail to produce footwall uplift even when Te is increased to 10 km (Fig. 10A, G), or if the extension across the fault is increased to 10 km. We conclude that the 10° dip inferred from coring cannot be representative of the basin. The best first-order fit to basin geometry was acquired

using a 30° detachment, with a 5 km extension and a Te of 3 km (Fig. 10A, B). Although this is not a unique solution, it does give some guidance as to the type of fault and mechanical state of the crust under the Ilan Plain. Changing flexural rigidity to either 1 km or 5 km changed the basin width to less close fits (Fig. 10A, C, D). The modeling is also effective at eliminating very high angle faults as possible mechanisms, because even a 45° fault produces too much footwall uplift, hanging-wall subsidence, and a wide basin (Fig. 10A, H).

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FISSION-TRACK ANALYSIS The exhumation history of the ranges around the Ilan Basin can be examined using low-temperature thermochronometry. Although previous studies have used fission-track methods in the central parts of Taiwan (e.g., Liu et al., 2000, 2001; Willett et al., 2003), no similar studies around the Ilan Basin have been undertaken. Liu et al. (2001) used zircon fission-track methods to date the cooling history and included a transect north of the basin, across the Hsüehshan Range. Fission-track studies in zircon grains record cooling through a partial annealing zone of 200–320 °C (Tagami et al., 1998), and so this method is sensitive to exhumation driven by erosion and has been widely used in orogenic exhumation studies. However, Liu et al. (2001) recorded zircon fission tracks north of the Ilan Basin that are mostly older than the Taiwan orogeny. Central ages are as low as 22 Ma close to the topographic front of the Hsüehshan Range but become older farther north, consistent with more exhumation at the margin of the basin (Liu et al., 2001). That the zircon fission tracks are this old indicates that they have been partially but not fully reset by late Miocene–Holocene collision-related burial. A lower temperature thermochronometer is thus required to reconstruct the cooling of rock units in shallower parts of the crust. In this study we employed the fission-track method applied to apatite, which records cooling through ~125–60 °C over time scales of 1–10 m.y. (Green et al., 1989). Apatite fission-track analysis was performed at University College, London, UK. Polished grain mounts were etched with 5N HNO3 at 20 °C for 20 s to reveal the spontaneous fission tracks. Subsequently, the uranium content of each crystal was determined by irradiation, which induced fission of 235U. The induced tracks were registered in external mica detectors. The samples for this study were irradiated in the thermal facility of the Hifar Reactor at Lucas Heights, Australia. The neutron flux was monitored by including Corning glass dosimeter CN-5, with a known uranium content of 11 ppm, at either end of the sample stack. After irradiation, sample and dosimeter mica detectors were etched in 48% HF at 20 °C for 45 min. Only crystals with sections parallel to the c-axis were counted, as these crystals have the lowest bulk etch rate. To avoid biasing results through preferred selection of apatite crystals, the samples were systematically scanned, and each crystal encountered with the correct orientation was analyzed, irrespective of track density. The results of the fission-track analysis are presented in Table 1.

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Of 12 samples analyzed, only 2 yielded sufficient apatite to produce meaningful results. Fortunately these samples came from opposite sides of the basin (Fig. 2) and allow the exhumation histories of the margins to be compared. Central ages of 1.2 ± 0.5 Ma and 3.5 ± 0.5 Ma were recorded at the coast on the north and south basin margins, respectively (Fig. 11). The southern sample contains a minor number of older grains (older than 150 Ma), reflecting an earlier phase of cooling. However, for this study we focus on most of the population, which records cooling linked only to Pliocene–Holocene tectonism. The single, young grain population seen on the northern margin indicates total resetting and recent, rapid cooling. Exactly why the southern margin sample contains a few grains that are not reset at 3.5 Ma is not clear. However, the vast majority of grains in that sample show a well-defined 3.5 Ma cooling trend (Fig. 11). The spread of ages would be greater if only partial annealing had occurred. We conclude that the sample underwent rapid cooling through the AFTA annealing temperature zone ca. 3.5 Ma. Our result indicates more recent cooling of the north margin in comparison with the south, as might be expected for an asymmetric basin controlled by a south-dipping detachment. What is surprising is that this result is the opposite of Liu et al.’s (2001) zircon fission-track result that has central ages of only 1.5 ± 0.3 Ma in the south close to Suao and 22.2 ± 2.3 Ma along the northern margin, which would imply the opposite sense of motion. Such motion is not consistent with the structural asymmetry presented above. Although the apatite and zircon results are consistent in the north, it is impossible to have younger fission-track ages for zircon than for apatite, calling into question whether the zircon or apatite data from Suao are representative of the region. Assuming a geothermal gradient of 30 °C/km (Barr and Dahlen, 1988; Willett et al., 2003), the apatite data imply average exhumation rates of 1.6 ± 0.3 km/m.y. in the north and 0.7 ± 0.1 km/m.y. in the south. Because we demonstrate the importance of extensional motion on the Lishan Fault in generating uplift of the Hsüehshan Range, much of the exhumation on the north margin may be linked to tectonic unroofing by detachment faulting rather than by erosion. SEISMIC EVIDENCE Observation of recent earthquakes is important to understanding the neotectonic evolution of the Ilan Basin, as they provide a measure of where faulting is active and a snapshot of the current stress regime. The seismicity used here is a subset of

TABLE 1. FISSION-TRACK ANALYTICAL DATA Spontaneous Induced Age dispersion 2 Ns Ni RE% ρs ρi Pχ 0.185

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Age components mode 1 mode 2 mode 3 (number of grains) 3.5 ± 0.5 22 ± 12 294 ± 78 (26) (2) (2) Single pop

Apatite 15 1.387 3845 0.009 13 1.666 2116 19.8 97.0 1.2 ± 0.5 6 –2 Note: (i) Track densities are (×10 tr cm ) numbers of tracks counted (N) shown in parentheses; (ii) analyses by external detector method using 0.5 for the 4π/2π geometric correction factor; (iii) ages calculated using dosimeter glass CN-5; (apatite) ξCN5 = 338 ± 4; CN-2; (zircon) ξCN2 = 127 ± 5 calibrated by multiple analyses of IUGS apatite and zircon 2 2 age standards (see Hurford, 1990); (iv) Pχ is probability for obtaining χ value for v degrees of freedom, where v = no. crystals – 1; (v) central age is a modal age, weighted for different precisions of individual crystals (see Galbraith, 1990). TW116-3

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TW114-1 Central age: 14.4 ± 7 Ma P(χ2): 0.0% Relative error: 256% Number of crystals: 30

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events initially reported by the Central Weather Bureau of Taiwan. The phase data for these events are used for relocation by means of the double-difference method of Waldhauser and Ellsworth (2000). The relocation procedure minimizes the errors in hypocentral determination because of lateral velocity variations. The resulting events tend to cluster around known structures, as shown by Waldhauser (2001). The Benioff zone in this area is below 50 km and strikes nearly E-W and dips at ~30° to the north (Fig. 12; Wu et al., 1997). In the Ilan area the deeper seismicity (~50 km) is concentrated along the southwestern margin of the basin, as well as more generally across the northern half of the basin. The seismicity in the upper 50 km is shown in Figure 13. Cross sections through the basin show that the deeper earthquakes form a NW-dipping array (Fig. 14A–C).

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However, a coherent array of deep earthquakes is not seen under the western Okinawa Trough in the 0–50 km depth range considered here. Our deeper profile (Fig. 12) indicates that the Philippine oceanic slab is deeper in this location and that therefore the array of seismic events seen in Figure 14A–C reflects a separate lithospheric structure. The NE-trending seismic zone in the central part of the Ilan Plain is populated by shallow normal-faulting events with NW-SE directed T-axes (extension). In the southeast basin, shallow seismicity appears to be arranged into two shallow E-W–trending zones dominated by sinistral strike-slip events and N-S–oriented normal faulting, but with T-axes in the same general direction as in the basin center. The two areas of intense seismicity are not that different in that both show a shallow layer of seismicity between a narrow range of 8–13 km. Analysis of the map and sections in Figures 13 and 14 reveals the subtle differences in seismicity associated with different belts. In the first place, the shallow events are concentrated in a fairly narrow depth range. This is curious; one possible explanation is that the seismic layer is limited on its upper surface by the presence of relatively incompetent and therefore nonbrittle sediment. In contrast, the lower limit of seismic activity may be bounded by hot middle crust at ~10 km depth, where ductile deformation precludes seismogenesis. Sections B and E (Fig. 14) show the depth distribution of the E-W–trending strike-slip events around 24°34′N (southern end of the profiles). The NE-SW–trending belt does show well in the cross section, with only a few events >30 km in Figure 14B, C. No evidence exists for a shallow-dipping active slip plane bounding the basin to the north on the basis of seismicity. Whatever fault motion we can confirm is either diffused or concentrated in a narrow depth range, probably between the bottom of the sedimentary layer and the top of the basement, judging from the rheological properties of these materials. The events >50 km deep under the basin axis are interpreted as thrust earthquakes, consistent with their association with the Ryukyu subduction zone, which is now starting under northern Taiwan by lateral motion of the Philippine Sea plate into a tear in the passive margin lithosphere of southern China (Lallemand et al., 2001). However, analysis of earthquakes shallower than 20 km from the nearby Okinawa Trough shows a dominant extensional character, with extension perpendicular to the strike of the trough, i.e., NW-SE. These shallow earthquakes form an offshore continuation of the shallow events seen under the Ilan Basin, which are similarly interpreted as being extensional. No strong evidence exists from earthquakes to support the presence of a currently active major south-dipping detachment (e.g., the Lishan Fault). However, if the major extensional fault is shallow dipping, then it is too shallow to store elastic stresses and to be seismogenic under the northern part of the Ilan Plain. Deeper faulting under the southern edge of the basin is consistent with these asymmetric structural models. Strike-slip faulting in the southeast basin may be interpreted as part of the strain accommodation related to the bend in the tectonic fabric from a N-S orientation in central Taiwan to an E-W orientation in the Ryukyu Arc. The strain observed indicates motion of the southern margin

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of the Ilan Basin to the ESE. The sense of motion is the opposite of that expected from the bending of the structural fabric, but it is consistent with the lateral motion of orogenic crust away from the Taiwan mountains toward the Ryukyu Trench, as might be predicted for orogenic collapse driven by a change in stresses triggered by generation of that trench. GEODYNAMIC EVOLUTION Additional constraints on the nature of current tectonic strain are provided by global positioning system (GPS) monitoring of the region. Here we consider the motions reconstructed by Chang et al. (2003), as shown in Figure 15. Relative to stable southern China the ranges north of the Ilan Basin are moving slowly to the northwest (~10–15 mm/yr), effectively a continuation of the west-directed thrusting that characterizes most of Taiwan, albeit slightly rotated clockwise in this particular area. In contrast, GPS locations in the Ilan Basin itself are moving almost due east, while the basement around Suao on the southern boundary is being displaced toward the southeast at 15–20 mm/yr. The net result indicates motion of the Hsüehshan and Backbone Ranges away from one another, consistent with the seismic evidence for extension across the basin. The eastward motion of the Ilan Basin shows the same sense of motion as the strike-slip fault data derived from the fault plane solutions of the seismic data. It should be noted that the major strike-slip structures (Fig. 11) are south of the GPS stations on the Ilan Plain. The GPS data thus indicate a broader zone of crustal extrusion than that based

on the seismology alone, implying extrusion of this crustal block toward the Okinawa Trough. We do recognize the importance of the strike-slip zone resolved by the seismic data and note that GPS motions are more south-oriented, directly toward the trench, south of that lineament. The GPS motion data can be understood in the context of the tectonic setting that is characterized by E-W compression in central Taiwan, contrasting with N-S subduction to the east along the Ryukyu Trench. In practice the Ryukyu Trench provides a free edge, allowing the southward motion of crustal blocks in the Ryukyu forearc, compared with the compression by the Luzon-China collision (Fig. 16). If Taiwan is considered an arc-continent collision orogen that is progressively migrating along the Chinese margin toward the southwest, then the passive-margin units that are compressed and uplifted by collision may then collapse and extend, once the restraining buttress of the Philippine Sea plate has been removed and subduction with the opposite polarity is initiated. The load of the Taiwan orogen causes flexure of the underlying Chinese continental crust (Yu and Chou, 2001; Lin and Watts, 2002; Lin et al., 2003). Where that load slides southeastward away from the margin’s edge in northern Taiwan, the continental crust quickly rebounds and regains much of its normal thickness (Rau and Wu, 1995). Thus SE-directed motion of the basement south of the Ilan Basin reflects collapse of the flank of the Taiwan orogen and motion of the Backbone Range toward the Ryukyu Trench. Extension in the Okinawa Trough, driven by southward motion of the Ryukyu forearc toward the trench also provides space for

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Figure 14. Cross sections down to 50 km depth across the basin show concentrated seismicity ~13–8 km depth. Those in the 13–50 km range are evidently associated with collision of the Philippine Sea and Eurasian plates, and the subduction events are at depths >50 km in this region. Section locations are shown in Figure 13.

the eastward motion of material in the Ilan Basin, driven by the gravitational potential of the thickened crust in central Taiwan. This motion is partially accommodated by strike-slip tectonism as seen in the seismic data, as well as by extension, especially that focused on the major fault bounding the Ilan Plain to the north, inferred to be the Lishan Fault, reversed in direction in comparison with its thrust sense in the Central Ranges. DISCUSSION The fact that the Lishan Fault can be traced along strike from a major thrust fault into an extensional detachment is consistent

with the extension in the Ilan Basin being the product of postorogenic collapse, triggered by subduction polarity reversal, as argued by Teng (1996), and not the result of extension propagating from the Okinawa Trough, driven by trench forces. The extensional fault that controls the Ilan Basin does not represent a westward propagating rift of oceanic origin cutting across the older orogenic fabric, but rather a reversal of motion on thrusts as compressional stresses are released toward the northern end of the orogen. The major bounding fault can be clearly traced as an extension of the Lishan Fault, which is a major thrust structure in the Central Ranges but is overturned and moving as an extensional structure where it reaches the Ilan Plain. The initiation of

Figure 15. Map showing the motions of crustal blocks around the Ilan Basin relative to stable Eurasia, recorded by Chang et al. (2003) using GPS methods over the period 1990–1995.

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Figure 16. Schematic depiction of the origin of the Ilan Basin as a result of gravitational collapse of the Taiwan Central Ranges during a SW-migrating collision of the Luzon Arc and mainland China.

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the Ryukyu Trench removes the compressive force of the Philippine Sea plate from the orogen and allows the edifice to extend as the trench provides a free edge toward which material can be displaced. In turn this suggests that at least the southwestern end of the Okinawa Trough has also formed as a result of migrating arc-continent collision and subsequent orogenic collapse and is not generated by subduction slab rollback (cf. Suppe, 1984). This is an important revision of the generally held belief that the Okinawa Trough is formed by trench forces, principally rollback of the trench, much as suggested for the Mariana Trough or the Lau Basin (e.g., Hawkins, 1974). However, in the rollback model the presence of the westernmost propagating end of the rift at the northern tip of Taiwan would be coincidental, whereas rifting related to collision and subduction polarity reversal would predict this position. In practice, little evidence exists for active rift propagation. The Okinawa Trough is not V-shaped as seen in other western Pacific basins, where a propagating rift culminates in seafloor spreading. The Okinawa Trough appears to be opening to the west, yet the older eastern parts of the basin are no wider than that closest to Taiwan. In the Okinawa Trough, rifting is followed by a cessation of extension, whereas in contrast rollback and arc rifting are normally followed by further extension in the form of seafloor spreading. Evidence that the westward migration of the Okinawa Trough is matched by arc migration was provided by Shinjo (1999), who noted that middle Miocene volcanic rocks from the southern Ryukyu Arc are not subduction related but instead are similar to intraplate volcanism seen in China. Such an observation implies that the Ryukyu Arc is new and that the subductionrelated Ryukyu volcanic front has propagated into the area since that time. This hypothesis is also supported by the geochemistry of recent volcanic rocks from the southernmost Okinawa Trough (Chung et al., 2000). In a collapse model the Okinawa Trough becomes increasingly younger to the south, the opposite of the rollback model in which the Okinawa Trough spreading centers are propagating into the basin away from Taiwan. It is noteworthy that active magmatism and faulting that cuts right to the seafloor in the Okinawa Trough is restricted to that part of the basin closest to Taiwan (Sibuet et al., 1998). In contrast, middle to late Miocene (6–9 Ma) extension ages are recorded in the northern Okinawa Trough (Letouzey and Kimura, 1985). The northern Okinawa Trough thus is either the remnant of an earlier arc-continent collision, as favored by Clift et al. (2003), or has a separate origin from that part of the basin closer to Taiwan. The basement of the southern Okinawa Trough is inferred to be the extended remnants of the Taiwan orogen, equivalent to the Backbone and Hsüehshan Ranges. The recognition of a continuous Taiwan-Sinzi folded zone under the SE edge of the East China Sea (Hsiao et al., 1999) would suggest a continuous migration of the orogen from Sinzi at ca. 12 Ma to present-day Taiwan. Volcanism is the manifestation of the new arc volcanic front to the Ryukyu subduction zone, which overlies the nonvolcanic forearc ridge only in the central and northern Ryukyu Arc, where active extension has ceased.

Here we propose that the Ilan Basin can best be understood in the context of a migrating arc collision, and especially as the culmination of gravitational collapse of the resultant orogen, having taken place over relatively short periods of geologic time (Fig. 17). Collision of the modern Taiwan section of the Luzon Arc with China is thought to have begun ca. 6–9 Ma (e.g., Dorsey, 1988; Teng, 1990; Sibuet et al., 2002; Malavieille et al., 2002; Huang et al., 2006) and is already finished and in a state of rapid exhumation around the Ilan Basin. Byrne and Crespi (1997) reported extension throughout the Backbone and Central Ranges. The stronger extension seen around the Ilan Basin can be understood as an extension of this, made possible as the new Ryukyu Trench form, removing the compressive stresses of collision and allowing the flank of the Taiwan orogen to move laterally into that space. Figure 18 shows the relationship between tectonically driven rock uplift rates and exhumation rates in Taiwan, assuming that “hard collision” in northern Taiwan initiated ca. 5 Ma, whereas in the south hard collision between the Luzon Arc and China is just starting at the modern coast. Initial collision between the Luzon forearc and the Chinese margin begins farther south, with the development of an accretionary prism, which progressively overthrusts the North Luzon Trough (forearc basin; Lundberg et al., 1997). Regional trends in rock uplift rates can be determined from the current elevations and the age of the collision, together with estimates for the amount of exhumation derived from the metamorphicgrade and fission-track data (e.g., Dadson et al., 2003). Although in some areas modern rates of uplift have been determined by dating terraces (e.g., Lin, 1969; Peng et al., 1977; Vita-Finzi and Lin, 1998), these terraces are necessarily limited to the coastal regions, mostly in the Coastal Ranges of eastern Taiwan. Exhumation rates driven by erosion reach a peak in the south of the island, because rates of rock uplift are highest during the most intense period of collisional compression between Luzon and China; these are partly balanced by erosion driven largely by precipitation but also by tectonic extension (Crespi et al., 1996; Teng et al., 2000). Exhumation and vertical uplift rates decrease abruptly toward the northern end of the Central Ranges, especially around the Ilan Basin, although active motion along a detachment reversing the Lishan Fault causes increased exhumation in the Hsüehshan Range. The rates of burial and exhumation are some of the highest known on Earth. Exhumation rates at Nanga Parbat in the Pakistan Himalaya reach 5 mm/yr (Zeitler et al., 1993), slightly less than the rates seen in the Central Ranges (Dadson et al., 2003). Although exhumation rates are lower in the Hsüehshan Range, reaching >1.6 mm/yr adjacent to the Lishan Fault, these levels are still relatively high, far exceeding rates of 0.33 mm/yr in the Cascade Range of the northwestern United States (Reiners et al., 2003), somewhat more than the ~1 mm/yr found in the Alaska Range (Fitzgerald et al., 1993), and comparable to the 2.5–0.8 mm/yr measured by Tippett and Kamp (1993) in the southern Alps of New Zealand. Clearly arccontinent collision forms some of the most dynamic geology on the planet.

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CONCLUSIONS A variety of geological and geophysical data demonstrates active extension along a NNW-SSE axis across the Ilan Basin of northern Taiwan. The structure of the basin appears to be largely controlled by a SE-dipping detachment fault, likely dipping at ~30°, that causes uplift of the Hsüehshan Range to the north of the basin and preferential fast subsidence of the northern Ilan Basin. Apatite fission-track ages confirm geomorphic evidence for faster exhumation of the northern margin of the basin, reaching rates of at least 1.6 ± 0.3 km/m.y. The main basin-bounding extensional fault is mapped as a continuation of the Lishan Fault, a major thrust structure from central Taiwan. The southern Backbone Range is shown by seismic and GPS data to be moving southeast toward the newly formed western Ryukyu Trench. Formation of a free edge in the trench and release of the E-W compressive stresses allow the Taiwan orogen to collapse in the Ilan Basin–Okinawa Trough region. Because the Taiwan orogen is migrating to the southwest along the passive margin of southern China, we suggest that the extension and associated basin formation must also move in this direction. Although the Ilan Basin was formed by the

start of collapse onshore, the process reached its culmination offshore in the Okinawa Trough, indicating that at least some of this basin has formed as a result of collision and not by trench rollback forces as previously believed (cf. Suppe, 1984). We propose that mountain building and rifting of the Ilan Plain and the Okinawa Trough are all results of a single common process, arc-continent collision, rather than being the unique interaction of a collisional orogen and a propagating backarc rift. Like the Alboran Sea in the western Mediterranean the Ilan-Okinawa Trough shows how far postorogenic collapse may drive extension and basin formation (e.g., Platt and Vissers, 1989), although in this case without the need to invoke delamination of the mantle lithosphere. Because arc-continent collision is a common process in the plate-tectonic cycle we anticipate that collapse basins similar to the Ilan Basin should be a common feature and should be recognizable in ancient collision zones. Although the Ordovician South Mayo Trough of the Irish Caledonides has been interpreted to be such a basin (e.g., Clift et al., 2003), the scarcity of other examples likely reflects a lack of understanding of existing sequences rather than the absence of gravitational collapse in ancient arc collisional events.

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ACKNOWLEDGMENTS P.C. thanks the University of Aberdeen, College of Physical Sciences; the Alexander von Humboldt Foundation; and the Joint Oceanographic Institutions for contributions to the cost of this research. A.T.L. is grateful for financial support for conducting

fieldwork from an NSC (Taiwan) grant of 93–2116-M-008–001. F.T.W. acknowledges support from the U.S. National Science Foundation’s Continental Dynamics Program. We thank Don Fisher, Neil Lundberg, Dave Topping, and Patrick Barnard for their helpful reviews, as well as editor Dave Scholl for his help in improving this manuscript.

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MANUSCRIPT ACCEPTED BY THE SOCIETY 24 APRIL 2007

Printed in the USA

The Geological Society of America Special Paper 436 2008

The Guerrero Composite Terrane of western Mexico: Collision and subsequent rifting in a supra-subduction zone E. Centeno-García* Instituto de Geología, Universidad Nacional Autónoma de México, Ciudad Universitaria, México D.F. 04510, México M. Guerrero-Suastegui O. Talavera-Mendoza Unidad Académica de Ciencias de la Tierra, Universidad Autónoma de Guerrero, AP 197, Taxco el Viejo, Guerrero, México

ABSTRACT The Guerrero Composite Terrane of western Mexico is the second largest terrane in North America. Mostly characterized by submarine volcanism and formed by five terranes, the Guerrero records vast and complex subduction-related processes influenced by major translation and rifting. It is composed of the Teloloapan, Guanajuato, Arcelia, Tahue, and Zihuatanejo Terranes. The Teloloapan Terrane is made up of Lower Cretaceous island-arc (IA) andesitic to basaltic submarine lava flows, interbedded with limestone and shallow-marine volcaniclastic rocks. The Guanajuato and Arcelia Terranes are characterized by Lower Cretaceous supra-subduction ophiolite successions formed by deep-marine volcanic and sedimentary rocks with mid-oceanic-ridge basalt (MORB), oceanic-island basalt (OIB), and island-arc basalt (IAB) signatures. These two terranes are placed between the continent and the more evolved arc assemblages of the Zihuatanejo Terrane. The Tahue Terrane is composed of Paleozoic accreted arc and eugeoclinal sedimentary rocks, Triassic rift-related metaigneous rocks, and overlain unconformably by pillow basalts, limestone, and volcaniclastic rocks. The Zihuatanejo Terrane was formed by Triassic ocean-flank to ocean-floor assemblages accreted in Early Jurassic time (subduction complexes). The subduction complexes are overlain by Middle Jurassic–evolved volcanic arc rocks, which are in turn unconformably overlain by Early and Late Cretaceous subaerial and marine arc-related volcano-sedimentary assemblages. Mesozoic stratigraphy at the paleocontinental margin of Mexico (Oaxaquia and Mixteca Terranes) is formed by Triassic submarine fan turbidites accreted during Early Jurassic time; Middle Jurassic–evolved volcanic arc rocks are unconformably covered by a Late Jurassic to Cretaceous calcareous platform. Six stages in the tectonic evolution are proposed on the basis of the stratigraphic and deformational events recorded in western Mexico: (1) A passive or rifting margin developed along the western margin of continental Mexico throughout the Triassic. A

*[email protected] Centeno-García, E., Guerrero-Suastegui, M., and Talavera-Mendoza, O., 2008, The Guerrero Composite Terrane of western Mexico: Collision and subsequent rifting in a supra-subduction zone, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 279–308, doi: 10.1130/2008.2436(13). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Centeno-García et al. thick siliciclastic turbiditic succession of the Potosi Submarine Fan was accumulated on the paleo-continental shelf-slope and extended to the west in a marginal oceanic basin. (2) Subduction began in the Early Jurassic, and the turbidites of the Potosi Fan with slivers of the oceanic crust were accreted, forming a wide subduction prism. (3) Exhumation of the accretionary prism and development of a Middle Jurassic continental arc onto the paleo-continental margin (Oaxaquia and Mixteca Terrane) took place, and also in the Zihuatanejo Terrane. (4) Intra-arc strike-slip faulting and rifting of the Middle Jurassic continental arc took place along with migration of the subduction toward the west and development of a calcareous platform in Oaxaquia and the Mixteca Terrane (continental Mexico). (5) Drifting of the previously accreted Tahue and Zihuatanejo Terranes formed a series of marginal arc-backarc systems, or one continuously drifting arc with intra-arc and backarc basins during Early to middle Cretaceous time. (6) Deformation of the arc assemblages, and development of Santonian to Maastrichtian foreland and other basins, date the final amalgamation of the Guerrero Composite Terrane with the continental margin. Keywords: Guerrero Terrane, Mexico, tectonics, Triassic subduction complex, Cretaceous arc volcanism.

INTRODUCTION The present configuration of continental Mexico was built after accretion of basement remnants and oceanic terranes. During most of their Mesozoic history, Proterozoic to Paleozoic accreted terranes formed a relatively narrow neck of land adjacent to the North American craton. This was bordered on its eastern side by rifting and on its western side by active subduction. Thus Mexico is probably one of the most suitable regions in North America for studying the interaction between these two differing tectonic scenarios. We suggest in this paper, based on evidence recorded in the stratigraphy of the Guerrero Composite Terrane and surrounding terranes, that the almost continuously subducting Pacific margin of Mexico was directly influenced by extensional tectonics associated with the breakup of Pangea and the formation of the Gulf of Mexico. The Guerrero Composite Terrane (Campa and Coney, 1983) constitutes approximately one-third of Mexico. As originally described, it is the largest of all the Mexican terranes and probably the second largest of the North America Cordillera after Wrangellia (Campa and Coney, 1983; Centeno-García et al., 1993a). The Guerrero Composite Terrane is characterized mostly by submarine and locally subaerial volcanic and sedimentary successions that range in age from Jurassic (Tithonian) to middle–Late Cretaceous (Cenomanian), and scarce exposures of older rocks. A wide variety of models has been proposed for the origin of the Guerrero Composite Terrane. Like other terranes of the North America Cordillera, it was first interpreted as an exotic terrane formed by a fartraveled Cretaceous oceanic arc. Some authors have suggested that it was an oceanic arc terrane that was accreted to nuclear Mexico in Late Cretaceous time via a westward-dipping subduction zone that closed a major ocean basin (Lapierre et. al., 1992; Tardy et al., 1994; Dickinson and Lawton, 2001, etc.). Other authors have

suggested that the Guerrero Composite Terrane might represent one or more complex systems of two or three peripheral arcs that developed relatively close to the continent (Campa and Ramírez, 1979; Ramírez-Espinosa et al., 1991; Mendoza and Suastegui, 2000; Centeno-García et al., 2003; Centeno-García, 2005). Some models even proposed that the arc was autochthonous and was built upon Proterozoic continental crust of nuclear Mexico (de Cserna, 1978; Elías-Herrera and Sánchez-Zavala, 1990). In other words, there is a model for each likely possibility, but each lacks strong supporting evidence. New findings on the stratigraphy, discussed in this paper, suggest a more complex evolution, implying a series of accretions to the continent followed by rifting, and later by collision. In this paper we attempt to present our insights into the evolution of western Mexico gained from examining the stratigraphy and structure, and the geochemical and geochronological data, of such a vast area. However, we discuss in this paper only stratigraphic units and localities that are keys for reconstructing the tectonic evolution. This paper synthesizes the work done by many authors. Although there is the need for more geochronological and detailed field work, we consider that the preliminary tectonic model presented in this paper is consistent with the evidence collected to date. OVERVIEW OF THE GUERRERO COMPOSITE AND NEIGHBORING TERRANES The stratigraphy of western Mexico is synthesized in this paper under the framework of tectono-stratigraphic terranes, which are regions that share the same geological history and are bounded by major faults. As mentioned before, by the early Mesozoic, the Paleozoic and Proterozoic terranes were already accreted to the southern part of the North American craton.

Guerrero Composite Terrane of western Mexico Those that already formed part of the continental margin during the Mesozoic were Oaxaquia and the Mixteca, Parral, and Cortes Terranes (Fig. 1). Terranes accreted or displaced during the Mesozoic were those of the Guerrero Composite, the Central, as well as terranes of the western Baja California Peninsula. The latter will not be reviewed in this paper. A brief summary of the stratigraphy is described as follows; more detailed descriptions of key areas and events are discussed later.

al., 1995; Ramírez-Ramírez, 1992; Lawlor et al., 1999; Solari et al., 2003; Keppie et al., 2003). It is covered by Paleozoic sedimentary rocks (Fig. 2) that are capped by Permian volcanic and volcaniclastic rocks (McKee et al., 1999; Stewart et al., 1999; Rosales-Lagarde et al., 2005). Triassic (Carnian–Norian) sedimentary rocks (La Ballena Formation) are exposed at the western margin of Oaxaquia (Labarthe et al., 1982; SilvaRomo, 1993; Tristán-Gonzalez and Torres-Hernández, 1994; Centeno-García and Silva-Romo, 1997; Barboza-Gudiño et al., 1998, 1999, 2004; Bartolini et al., 2002). These rocks are made up of a thick succession of turbidites (Fig. 2) deposited in a submarine fan environment named the Potosi Fan (CentenoGarcía, 2005). Triassic rocks of the Potosi Fan were deformed prior to deposition of Jurassic volcanic-volcaniclastic rocks (Centeno-García and Silva-Romo, 1997). They are interpreted as a Jurassic continental arc and rest unconformably on the Triassic Potosi Fan. Jurassic arc strata are made up of subaerial andesitic-rhyolitic lava flows, interbedded with volcaniclastic rocks (Silva-Romo, 1993). The arc sequence changes transitionally upsection to shallow-marine

OAXAQUIA At the end of the Paleozoic, Proterozoic basement terranes of Gondwanan affinity were already accreted to the southern part of the North American craton. The largest of these is the Oaxaquia block (Fig. 1), a crustal fragment, subcontinent in size, of Grenville affinity (Ortega-Gutiérrez et al., 1995). This crustal block forms the backbone of eastern Mexico and is referred to herein as continental Mexico for the Mesozoic. Oaxaquia has a Precambrian (1157–900 Ma) crystalline basement (gneisses and anorthosites; Patchett and Ruíz, 1987; Ortega-Gutiérrez et

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volcaniclastic rocks, limestone, and some evaporites (Fig. 2; Silva-Romo, 1993; Tristán-González and Torres-Hernández, 1994; Barboza-Gudiño et al., 2004). Calcareous sedimentation in Oaxaquia ranges in age from late Oxfordian–Kimmeridgian to Turonian and is interpreted as the southern extension of the North

American seaway. A major change upsection from calcareous to clastic sedimentation occurred at the uppermost part of the Cretaceous, forming a thick succession of sandstone, shale, and conglomerate (Caracol Formation; Silva-Romo, 1993). Oaxaquia is overthrust by the Guerrero Composite Terrane (Fig. 1).

Guerrero Composite Terrane of western Mexico MIXTECA TERRANE The basement of the eastern Mixteca Terrane is made up of pre-Mississippian polydeformed metamorphic rocks of the Acatlán Complex (Ortega-Gutiérrez, 1981; Ruíz et al., 1988; Yañez et al., 1991). This complex is considered to be the result of complex interactions between Gondwana and Laurentia previous and during the assembling of Pangea (Ortega-Gutiérrez et al., 1999). It is unconformably overlain by Permian sedimentary rocks, which are in turn overlain unconformably by Middle Jurassic volcanic and sedimentary rocks (Fig. 2; García-Díaz et al., 2004). At the western part of the terrane, near the limit with the Guerrero Composite Terrane, partly metamorphosed volcanic and volcaniclastic rocks are exposed (Taxco Schist and Chapolapa Formation; de Cserna and Fries, 1981; Talavera-Mendoza, 1993; Campa and Iriondo, 2004). The Taxco Schist is made up of andesitic to rhyolitic lavas and volcaniclastic rocks of Early Cretaceous age (TalaveraMendoza, 1993; Campa and Iriondo, 2004). The Taxco Schist is unconformably overlain by a thick limestone succession of Albian to Cenomanian age and by Turonian–Maastrichtian clastic rocks (Mexcala Formation; Campa and Ramírez, 1979; TalaveraMendoza et al., 1995). Contacts between the Mixteca Terrane and Oaxaquia, as well as between the Mixteca and Guerrero Composite Terranes, are partially exposed. The Mixteca Terrane is on strike-slip fault contact with Oaxaquia, and rocks of the Guerrero Composite Terrane are thrust over the Mixteca Terrane. PARRAL TERRANE The Parral Terrane (Figs. 1 and 2) was first defined by Pacheco et al. (1984) and Coney and Campa (1987) and was redefined by Centeno-García (2005). The basement of the Parral Terrane is formed by Devonian to Carboniferous metamorphic rocks (Pescadito Schist; Eguiluz and Campa, 1982; Araujo and Arenas, 1986; Zaldivar and Garduño, 1984). These Paleozoic metamorphic rocks are unconformably overlain by red beds and volcanic successions (Nazas Formation; Pantoja-Alor, 1963), which change transitionally to Tithonian limestone (Araujo and Arenas, 1986; Contreras-Montero et al., 1988). Cretaceous calcareous and clastic sedimentation of the Parral Terrane is laterally continuous with the calcareous-clastic deposits that cover Oaxaquia and the Central Terrane. Relationships among the Parral, Central, and Cortes Terranes and the Parral Terrane and Oaxaquia are unknown, because the contacts are covered by Cretaceous limestone or by Cenozoic volcanic successions. Therefore, the exact locations of their boundaries are unknown but are inferred by the difference in styles of deformation of the Cretaceous rocks.

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Paleozoic “miogeosyncline” of Nevada. It was transferred toward the south by Middle to Late Jurassic time via the Mohave-Sonora megashear (Anderson and Silver, 1979, 2005; Stewart et al., 1990). The Cortes Terrane is interpreted as an autochthonous terrane to North America, which probably evolved at the margin of the Caborca Terrane (Stewart et al., 1990). It is made up of a thick succession of Paleozoic deep-marine turbidites that were thrust over platform limestone of the Caborca Terrane (Figs. 1 and 2). The Cortes Terrane is interpreted as continental-slope deposits, and it is considered the southern extension of the Paleozoic Cordilleran “eugeoclinal” deposits from Nevada and California (Poole and Madrid, 1988; Coney and Campa, 1987; Stewart et al., 1990). The previously deformed Paleozoic deep-marine rocks of the Cortes Terrane are overlain by Triassic (Carnian–Norian) terrestrial and marine sedimentary rocks (Stewart et al., 1990; Stewart and Roldán-Quintana, 1991). The Triassic rocks are overlain by Cretaceous red beds and volcanic rocks (Stewart and RoldánQuintana, 1991). Contact relationships between the Cortes and Guerrero Composite Terranes have not been well constrained, but the contact is inferred to be a Late Cretaceous thrust fault. CENTRAL TERRANE The nature of the basement of the Central Terrane is unknown, but it is assumed to be different from the Proterozoic basement of Oaxaquia because its oldest exposed rocks near its contact are a subduction-related accretionary complex (Taray Formation; Anderson et al., 1990; Diaz-Salgado et al., 2003; Anderson et al., 2005; Centeno-García, 2005). The subduction zone on which the Taray Formation was deformed was probably constructed along the Oaxaquia continental margin between Late Permian and Early Jurassic time (Diaz-Salgado et al., 2003; Anderson et al., 2005). The complex is unconformably overlain by Oxfordian subaerial rhyolitic to andesitic volcanic rocks and red beds (Jones et al., 1995). These rocks change transitionally to shallow-marine limestone that ranges in age from Late Jurassic to Late Cretaceous (Córdoba-Méndez, 1964). The location of the northern and eastern contact between the Central Terrane and Oaxaquia is inferred on the basis of the location of the last exposures of Paleozoic– early Mesozoic rocks, and a contrast in deformation styles of Cretaceous rocks in both (Fig. 1). The contact between the Central and Guerrero Composite Terranes has not been studied in detail but is inferred on the basis of the distribution of the northernmost exposures of Cretaceous marine volcanic rocks that belong to the Guerrero Composite Terrane. Structural trends on both sides of the contact suggest that the Central Terrane is overthrust by the Guerrero Composite Terrane to the south (Fig. 1). The thrusting is inferred to have occurred about Late Cretaceous time.

CABORCA AND CORTES TERRANES GUERRERO COMPOSITE TERRANE The Caborca Terrane has a Proterozoic basement older than 1.7 Ga (Anderson and Silver, 1981), covered by a thick Paleozoic sedimentary succession. It has been interpreted to be a displaced block of continental North America, originally located along the

Areas with large volumes of Lower Cretaceous volcanic and volcaniclastic rocks, located toward the west of Oaxaquia and the Mixteca Terrane, were originally grouped as the

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Guerrero Terrane by Campa and Coney (1983) and were, 10 years later, divided into the Tahue, Nahuatl, and Tepehuano Terranes by Sedlock et al. (1993). Subsequent regional mapping has shown that the divisions proposed by Campa and Coney (1983) are closer to the field locations of faults delimiting the terranes than those of Sedlock et al. (1993). Therefore, more recent reviews of the terrane distribution of Mexico (e.g., Centeno-García, 2005) have been based on Campa and Coney (1983). The Guerrero is a composite terrane, formed by at least five terranes: Tahue, Zihuatanejo, Guanajuato, Arcelia, and Teloloapan (Figs. 1 and 2; Talavera-Mendoza et al., 1995; Mendoza and Suastegui, 2000; Centeno-García et al., 2003; Centeno-García, 2005). Their stratigraphy is briefly described from NNW to ESE (Fig. 1): Tahue Terrane The Tahue Terrane contains the oldest rocks found so far within the Guerrero Composite Terrane (Fig. 2; Centeno-García, 2005). These rocks comprise Ordovician marine rhyoliticandesitic lavas and clastic and calcareous rocks, all deformed and metamorphosed to low-greenschist facies (El Fuerte Complex; Mullan, 1978; Roldán-Quintana et al., 1993; Poole and Perry, 1998). These rocks may have originated as an oceanic arc that apparently was accreted previous to the deposition of Pennsylvanian–Permian deep-marine sedimentary rocks (San José de Gracia Formation; Carrillo-Martínez, 1971; Gastil et al., 1991; Arredondo-Guerrero and Centeno-García, 2003; Centeno-García, 2005). These deep-marine turbidites are strongly deformed but do not show the metamorphism of the El Fuerte Complex; thus an unconformable contact relationship between these two units is inferred. Paleozoic rocks of the Tahue Terrane are unconformably overlain by Cretaceous marine arc volcanic rocks and are interpreted as part of the Guerrero Arc (OrtegaGutiérrez et al., 1979; Henry and Fredrikson, 1987; RoldánQuintana et al., 1993; Freydier et al., 1995). These rocks are also cut by mafic and ultramafic intrusions that are part of the same Cretaceous arc magmatism (Henry and Fredrikson, 1987; Gastil et al., 1999; Arredondo-Guerrero and Centeno-García, 2003). Therefore, the Paleozoic units form the basement upon which the arc was built. The Tahue Terrane also contains metamorphic rocks of Triassic age (Keppie et al., 2006). The contact relationship between the Cortes and Tahue Terranes has not been studied in detail, but it is inferred to be a thrust (Fig. 1; Roldán-Quintana et al., 1993). The contact between the Tahue and Zihuatanejo Terranes is not exposed. Zihuatanejo Terrane The Zihuatanejo Terrane is the largest of all terranes that form the Guerrero Composite Terrane (Fig. 1). It extends north of the Mexican Volcanic Belt and along the Pacific Coast of Mexico (Centeno-García et al., 1993a, 1993b; Talavera-Mendoza et al., 1995; Mendoza and Suastegui, 2000). Its basement is made

up of large volumes of Triassic (Norian) quartz-rich turbidites (sandstone and shale) that are tectonically imbricated (Campa et al., 1982; Centeno-García et al., 1993a, 1993b). The turbidites form a matrix within which are blocks and slabs of pillow basalts, diabase, banded gabbros, chert, and limestone (Fig. 2). These rocks have received different names at different outcrops: Zacatecas Formation, Arteaga Complex, and Las Ollas Complex (Burckhardt and Scalia, 1906; Ranson et al., 1982; Cuevas-Pérez, 1983; Monod and Calvet, 1991; Centeno-García and SilvaRomo, 1997; Talavera-Mendoza, 2000; Centeno-García et al., 2003). The deformation of these rocks varies from gently folded strata to highly sheared block-in-matrix textures, and their metamorphism ranges from none to high-greenschist–amphibolite facies (Centeno-García et al., 2003). Blueschist facies have been reported only in one locality (Las Ollas Complex; TalaveraMendoza, 2000). These lithologies are interpreted to constitute an Upper Triassic(?)–Lower Jurassic subduction-related accretionary complex. Scattered exposures of rocks of Middle to Late Jurassic– evolved arc volcanism lie along the Pacific Coast of the Zihuatanejo Terrane. These rocks are made up of submarine rhyolitic lavas and volcaniclastic rocks, and granitoids that were emplaced in rocks of the accretionary complex (Bissig et al., 2003; Centeno-García et al., 2003). The Middle to Upper Jurassic arc rocks were in turn deformed and exhumed previous to the deposition of uppermost Jurassic–Cretaceous arc-related strata (Centeno-García et al., 2003). The Cretaceous arc succession ranges from Berriasian to Cenomanian in age, and it includes andesitic, basaltic, and some rhyolitic volcanic and volcaniclastic rocks, interbedded with limestone, evaporites, and some red beds (Grajales and López, 1984). The arc succession contains abundant fossils such as rudists, gastropods, microfossils, fossil logs, and vertebrates. This arc succession was deformed prior to the intrusion of large granitoids of latest Cretaceous to Paleogene age (Schaaf et al., 2000). Also, uppermost Cretaceous (Santonian to Maastrichtian) red beds and volcanic rocks rest unconformably on all previous units (Altamira Areyán, 2002; Benammi et al., 2005). The contact between the Zihuatanejo Terrane and Oaxaquia is exposed at its northern limit, where Cretaceous arc rocks of the Zihuatanejo Terrane are thrust over shallow-marine limestone of Oaxaquia. Its contact with the Arcelia and Guanajuato Terranes is inferred to be an east-verging thrust, but it is covered by uppermost Cretaceous and Cenozoic red beds and volcanic rocks. Guanajuato Terrane The Guanajuato Terrane has been interpreted as a complete crustal section through a primitive island arc that appears to lack an older basement (Ortiz-Hernandez et al., 1991; OrtizHernandez, 1992). It has also been interpreted as the remains of an oceanic basin that lay between the Guerrero arc and the continental margin (Freydier et al., 2000). This terrane was formed by a series of tectonic slivers that placed lower crust rocks (gabbro,

Guerrero Composite Terrane of western Mexico tonalite, serpentinite, wehrlite, and dike swarms) on pillow basalts, rhyolitic tuffs, volcanic turbidites, chert, and black detrital limestone (Quintero-Legorreta, 1992; Ortiz-Hernandez et al., 1992; Lapierre et al., 1992; Monod et al., 1990; Martínez-Reyes, 1992; Ortiz-Hernandez et al., 2003). These rocks were poorly dated as Tithonian–Hauterivian in age (Ortiz-Hernandez et al., 2003; Hall and Mortensen, 2003). Previously deformed volcanic turbidites are unconformably overlain by Aptian–Albian limestone (Ortiz-Hernandez et al., 2003). This suggests that sedimentation and at least one phase of deformation occurred previous to the Aptian–Albian (Ortiz-Hernandez et al., 2003). At present the Guanajuato Terrane is thrust over the calcareous platform of Oaxaquia (Ortiz-Hernández et al., 2002). Contact relationships between the Guanajuato and Zihuatanejo Terranes have not been constrained. Arcelia Terrane The Arcelia Terrane is made up of basaltic pillow lavas and ultramafic bodies, black shale and chert, and volcanic turbidites, all intensively deformed and partly metamorphosed (RamírezEspinosa et al., 1991; Talavera-Mendoza et al., 1995). It is characterized by Early Cretaceous deep-marine primitive arc or arcrelated oceanic facies and shows the least evolved magmatism of all the arc successions of the Guerrero Composite Terrane (Talavera-Mendoza et al., 1995; Mendoza and Suastegui, 2000). The Arcelia Terrane appears to lack an older basement (TalaveraMendoza et al., 1995; Mendoza and Suastegui, 2000). Rocks of the Arcelia Terrane apparently were thrust over the assemblages of the Teloloapan Terrane, and were in turn overthrust by rocks of the Zihuatanejo Terrane. However, these contacts are inferred because they are covered by younger red beds. Teloloapan Terrane The Teloloapan Terrane consists of two distinct regions: the eastern region is characterized by shallow-marine volcanic and sedimentary deposits (Fig. 2), and the western region by deeper volcanic and sedimentary facies (Guerrero-Suastegui et al., 1991; Ramírez-Espinoza et al., 1991; Talavera-Mendoza et al., 1995; Mendoza and Suastegui, 2000; Guerrero-Suastegui, 2004). Both are marine arc assemblages, which vary in composition from basalt-andesite to scarce dacite-rhyolite (Talavera-Mendoza et al., 1995). This unit contains microfossils (radiolarians and coccoliths), gastropods, and bivalves that range in age from Hauterivian to Aptian; these rocks change transitionally upsection to Aptian–Albian island-arc carbonates (Guerrero-Suastegui et al., 1991; Ramírez-Espinoza et al., 1991; Talavera-Mendoza et al., 1995). The Teloloapan Terrane (Fig. 1) is exposed in the easternmost parts of the Guerrero Composite Terrane. It is characterized structurally by a complex thrust-fault system that verges eastward. Its Lower Cretaceous rocks are severely deformed and metamorphosed in low-grade greenschist facies. The Teloloapan Terrane overrides either Lower to Middle Cretaceous platform

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carbonates or Upper Cretaceous clastic sediments that belong to the Mixteca Terrane (Fig. 1; Campa and Ramirez, 1979). The nature of its basement remains unknown. Metamorphic rocks that are exposed near the northwestern boundary of the Teloloapan Terrane with the Arcelia Terrane have been interpreted as a possible basement for the former (Elías-Herrera and SánchezZavala, 1990; Sanchez-Zavala, 1993). The rocks in this area are of uncertain age and origin. TECTONIC MODEL The most abundant rocks of the Guerrero Composite Terrane are marine, and rarely subaerial, arc volcanic and sedimentary successions that range in age from latest Jurassic (Tithonian) to middle Late Cretaceous (Cenomanian). The composition of the few scattered exposures of older units suggests a complex earlier tectonic evolution. These older rocks were not taken into consideration for the tectonic models proposed by previous authors (de Cserna, 1978; Campa and Ramirez, 1979; Elías-Herrera and Sánchez-Zavala, 1990; Tardy et al., 1994; Lapierre, et. al., 1992; Dickinson and Lawton, 2001, etc.). Based on the available information, we identified six main tectonic stages in the evolution of the Guerrero Composite Terrane. These stages are represented in Figure 3 and are briefly described in this section. Detailed discussion of the data that support the reconstruction of each stage is presented in the following section. Stage I: Collision of a Paleozoic Oceanic Arc?—Basement of the Tahue Terrane The basement of the Tahue Terrane (Fig. 3) is composed of the early Paleozoic accreted volcanic-sedimentary rocks of the El Fuerte Metamorphic Complex. There are not enough data available to constrain the origin of this complex. Preliminary interpretations considered these rocks as remnants of Gondwanan crust accreted during the formation of Pangea (Poole et al., 2005). In this model, metamorphic rocks of El Fuerte could be the western continuation of basement rocks of the Parral Terrane (Figs. 1 and 2). An alternative interpretation is that the El Fuerte Complex may be a displaced fragment of the early Paleozoic arc (Antler Arc) that collided with the western continental margin of North America during late Paleozoic time (Burchfiel et al., 1992; Sánchez-Zavala et al., 1999; Dickinson, 2004; Centeno-García, 2005). Carboniferous deep-marine turbidites (San José de Gracia Formation) that apparently cover the lower Paleozoic arc rocks unconformably may be correlative with deep-marine sedimentary rocks exposed in the eastern peninsular ranges of Baja California and the southwestern Cordillera of North America (Gastil et al., 1991; Centeno-García, 2005). In either of the two scenarios, deformed Paleozoic rocks of the Tahue Terrane are the basement upon which Cretaceous volcanism was built, indicating an earlier history of accretion of the Guerrero Composite Terrane than was previously interpreted by other authors.

W

Tahue (Guerrero Composite Terrane) Passive margin sediments

E

Cortes

Stage I Arc collision and development of a passive margin

Caborca

Paleozoic Vizcaíno?

Zihuatanejo (Guerrero Composite) and Central terranes

Oaxaquia

Continent margin Siliciclastic turbidites share the same provenance

Stage II Marginal oceanic basin with active rift volcanism

n Potosí Submarine Fan

Vizcaino?

Oaxaquia and Mixteca

Arteaga Basin

Triassic Carnian-Norian Vizcaíno?

Guerrero

Central

Oaxaquia Potosi Fan (deformed)

Taray Complex Arteaga Complex

Stage III Accretion via subduction wide accretionary prism

?

?

Early Jurassic

Oaxaquia and Mixteca

pre 180-163 Ma Guerrero

Central

Oaxaquia

transtension? Tumbiscatío Granitoid 163-158 Ma

158 Ma

?

+Oaxaquia and Mixteca Sub

Guerrero Zihuatanejo/Tahue

Composite Terrane Arcelia Guanajuato

duc

+

Central and Mixteca

Oaxaquia (north)

Teloloapan Calcareous K platform

?

Middle to Late Jurassic

tion

transtension?

?

Stage IV Continental arc and contemporaneous strike-slip and extension (roll-back of the subducting plate?)

?

Oaxaquia and Mixteca

Stage V Formation of a multiple arcs system or a single arc with intra-arc/back arc rifting Cretaceous Berriasian-Cenomanian

Sub

duc tion

Guerrero Zihuatanejo/Tahue

Composite Terrane Arcelia Teloloapan Guanajuato

Central / Oaxaquia and Mixteca (south)

compression

granitoids 105 Ma

Oaxaquia and Mixteca

Stage VI a Deformation and locally metamorphism, cut by 105 Ma granitoids in the Zihuatanejo terrane Cretaceous Cenomanian?-pre Santonian

Red Beds

Marine turbidites

Stage VI b Active continental arc in the west and syntectonic marine turbidites (foreland) in the east and deformation Cretaceous Santonian?-Maastrichtian

Figure 3. Tectonic models for the evolution of western Mexico, showing the alternating stages of subduction-collision and rifting.

Guerrero Composite Terrane of western Mexico Stage II: Late Triassic Passive Margin—Deposition of the Potosi Fan The paleo-continental edge of Mexico lay approximately at the western boundary of the Oaxaquia and Mixteca Terranes in the early Mesozoic (Fig. 1; Centeno-García, 2005). Thus the Central and Guerrero Composite Terranes (Fig. 1) were accreted or displaced to their present position during the Mesozoic. Sedimentation along the western continental margin of Oaxaquia was dominated by large volumes of siliciclastic turbidites (quartz-rich sandstone and shale) that were deposited in the distal continental shelf or at the continental slope at least during Carnian–Norian time (Fig. 3). Accretionary complexes that form the basement of the Central Terrane and parts of the Guerrero Composite Terrane (Zihuatanejo Terrane) are formed largely (up to 60% of the total area of exposures) by similar quartz-rich sandstone and shale turbidites that made up the matrix within which blocks of variable composition are embedded. These turbidites in the accreted terranes contain fossils of the same age as those from turbidites deposited at the continental slope of Oaxaquia. Detrital zircon ages obtained from turbidites from all the localities of the Carnian–Norian turbidites, from Oaxaquia to basal accretionary complexes of the Central and Zihuatanejo Terranes, show the same populations, which suggest that the fan turbidites spread into a marginal oceanic basin that was later accreted to the continental margin. These siliciclastic rocks are grouped as the Potosi Fan (Centeno-García, 2005) and are important because they can be traced from Oaxaquia to the present Pacific Coast of Mexico, and they tie together the Central Terrane, the westernmost part of the Guerrero Composite Terrane (Zihuatanejo Terrane), and the continental margin of southern North America (Oaxaquia) during Late Triassic time. Thus, the Potosi Submarine Fan may have been a large sedimentary feature, probably close to the dimensions of the present Bengal Fan. There is no evidence of Triassic magmatism in continental Mexico, and detrital zircon geochronology of the fan turbidites show that the youngest age populations are much older than depositional ages in all the studied localities of the Potosi Fan (Fig. 3; Centeno-García et al., 2005; Centeno-García, 2005). Therefore, the Potosi Fan probably was deposited across a passive margin, or at least a margin that had no active subduction along the length of the fan at the time of deposition. Stage III: Accretion of the Potosi Fan to the Continental Margin via Subduction—Basement of the Central and Zihuatanejo Terranes All the Triassic units of central and western Mexico are strongly deformed and partially metamorphosed, indicating that a major compressional event occurred during latest Triassic–Early Jurassic time. This event is characterized by tight folding, shearing, and axial cleavage in the continent-slope deposits of the Potosi Fan in Oaxaquia (La Ballena Formation), and blockin-matrix texture in the Taray Formation (Central Terrane),

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in the Zacatecas Formation, and in the Arteaga and Las Ollas Complexes (Zihuatanejo Terrane). These last three units formed in the distal ocean-floor zone of the Potosi Fan. The presence of mélanges (Arteaga Complex and Taray Formation) as well as blueschist in the Las Ollas Complex (Zihuatanejo Terrane) indicates that deformation occurred in a subduction zone. During this deformational event the turbidites of the Potosi Submarine Fan, with slivers of the oceanic crust and its sedimentary cover, were accreted to the continent. This accretionary prism apparently was very wide, as suggested by the large areas that are floored by it. Whether the subducting slab was dipping toward the west (under an oceanic arc) or the east (under continental Mexico) has not been constrained. There are two isolated reports of dated Early Jurassic volcanic rocks in Oaxaquia (Barboza-Gudiño et al., 2004; Fastovsky et al., 2005), but whether they are part of a continental arc or not is not known. Evidence of contemporaneous oceanic-arc magmatism is exposed in the Vizcaíno Peninsula of Baja California (Kimbrough and Moore, 2003), where Triassic–Jurassic volcanic rocks have geochemical signatures of primitive arc affinity. It is possible that the rocks in the Vizcaíno Peninsula represent a displaced fragment of an oceanic arc that accreted to the Arteaga and Las Ollas Complexes of the western Guerrero Composite Terrane, which in turn accreted to the Taray, Zacatecas, and La Ballena Formations, but this model needs to be supported by more evidence. Stage IV: Late Jurassic Continental Arc—Overlapping Assemblage for Guerrero Composite Terrane, Central Terrane, Oaxaquia, and Mixteca Terrane Subaerial volcanic and sedimentary rocks, as well as shallow porphyritic intrusives, dikes, and sills, overlie or cut previously deformed Triassic sedimentary rocks in Oaxaquia and rocks of the accretionary prism in the Central Terrane. These rocks range in age from 174 to 158 Ma (Jones et al., 1995; Barboza-Gudiño et al., 2004). A common attribute of all the outcrops of these rocks is that they are mostly rhyolitic in composition, with minor dacitic-andesitic lava flows and tuffs, and show evolved-arc geochemical signatures (Centeno-García and Silva-Romo, 1997; Centeno-García, 2002; Centeno-García and Díaz-Salgado, 2002). Coeval volcanic rocks have been reported in the Mixteca Terrane as well, suggesting that arc volcanism was widespread in continental Mexico at that time (García-Díaz et al., 2004). Rocks of similar age range and similar evolved-arc geochemical signatures are exposed in the western Zihuatanejo Terrane of the Guerrero Composite Terrane (Bissig et al., 2003; Centeno-García et al., 2003). This suggests that the Guerrero Composite Terrane may have been incorporated into the continental margin by that time. Summarizing the data described above: (1) Triassic basement rocks of the Zihuatanejo Terrane (Guerrero Composite Terrane) share a provenance linkage with rocks of the same age in Oaxaquia and the Central Terrane; (2) all Triassic rocks, from those deposited on the paleo-continent’s margin of Mexico to those within the accreted terranes, were deformed previous to the

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development of a Late Jurassic continental arc; and (3) evolved Upper Jurassic continental-arc volcanism was widespread among continental Mexico and accreted terranes (Central and Zihuatanejo Terranes). On the basis of these facts, we propose in this paper that the first accretion of the Guerrero Terrane occurred during latest Triassic–Early Jurassic time instead of near the end of the Cretaceous, as previously proposed by other authors. Therefore, the Late Jurassic magmatic event represents an overlapping assemblage that stitches all the terranes of central and western Mexico for that period. Stage V: Late Jurassic–Early Cretaceous Intra-Arc Strike Slip(?)–Rifting of the Continental Arc—Drifting of the Guerrero Composite Terrane It has been proposed that major lateral displacements occurred during the activity of the Jurassic continental arc of stage IV (Anderson and Silver, 2005). Therefore, the arc was originally in a more northerly position, and it was displaced, via the Mojave-Sonora Megashear, to its present position in central Mexico prior to, or at, the early stage of development of the calcareous platform (Anderson and Silver, 2005). Whether this major strike-slip system existed or not has been widely discussed (see GSA Special Paper 393). We consider that extensive geological evidence of major tectonism during and after arc volcanism exists (see following discussion). The cessation of magmatism in the Central Terrane and Oaxaquia suggests a change in the location of the subduction zone. Then, a major regional calcareous platform developed over the arc and other older rocks. This major transgression initiated the deposition of limestone on Oaxaquia, and on the Mixteca and Central Terranes. Calcareous sedimentation in central and eastern Mexico was characterized by high subsidence rates (Goldhammer, 1999). Arc magmatism continued only in a small area in the western Mixteca Terrane and became widespread in the Guerrero Composite Terrane. Although there is some overlap in age ranges of arc volcanism among the terranes that form the Guerrero Composite Terrane, there is a general trend from older ages in eastern Oaxaquia and the Central Terrane to younger ages in the western Guerrero Composite Terrane (Fig. 3). This suggests a possible W-SW migration of the subduction zone. We propose that during and after the continental arc activity (Late Jurassic–Early Cretaceous time), large amounts of extension and lateral translations may have occurred (see inferred faults in Fig. 1). This extensional-transtensional(?) event split the continental arc, initiating the drifting of parts of previously accreted oceanic rocks (basements of the Tahue and Zihuatanejo Terranes) and the generation of new oceanic crust (Guanajuato and Arcelia Terranes). With the data available, it seems that volcanic activity at the northern Zihuatanejo Terrane and at the Guanajuato and Teloloapan Terranes was restricted to latest Jurassic–Early Cretaceous time (Fig. 3). In contrast, in the Arcelia, Tahue, and southern Zihuatanejo Terranes, arc volcanism apparently continued up to Albian–Cenomanian time (Fig. 3). Geochemical and isotopic

compositions of most of the Upper Jurassic–Cretaceous igneous rocks of the different arc assemblages of the Guerrero Composite Terrane suggest primitive sources, with little or no influence on an evolved continental crust (e.g., Ortiz-Hernandez et al., 1991; Lapierre et al., 1992; Centeno-García et al., 1993a; Freydier et al., 1995; Gastil et al., 1999; Talavera-Mendoza et al., 1995; Mendoza and Suastegui, 2000, among others). Basalts with oceanisland (OI) and mid-oceanic-ridge basalt (MORB) signatures of the Arcelia and Guanajuato terranes (Lapierre et al., 1992; OrtizHernandez et al., 2003; Mendoza and Suastegui, 2000) suggest the influence of a mantle source for the magmatism. Regional differences in the strata suggest abrupt lateral changes in the depositional environments from shallow marine to deep marine. Also, lateral differences in thickness of the successions suggest that they may have been deposited in alternate subsiding basins and basement highs where the deposits draped thinly or were absent. These major geological differences suggest that intra-arc rifting was considerable and was probably associated with a complex paleogeography of marginal arc and backarc systems in western Mexico. Whether or not the different terranes of the Guerrero Composite Terrane were formed in a single arc has not been constrained. Some authors proposed that the Guerrero Terrane formed from a complex system of two or three arcs (Ramírez-Espinosa et al., 1991; Mendoza and Suastegui, 2000). However, no Cretaceous subduction-related accretionary prisms have been identified within any of the terranes of the Guerrero Composite Terrane. Stage VI: Final Accretion of the Guerrero Composite Terrane, and Development of a New Continental Arc A major Late Cretaceous–early Paleogene orogenic phase is recorded throughout Mexico, coeval to the Sevier and Laramide orogenies in western North America. This event is associated with the Mexican Fold and Thrust Belt of the Sierra Madre Oriental. Apparently, final amalgamation of the Guerrero Composite Terrane occurred during this orogenic event, and volcanic and sedimentary rocks of the Teloloapan, Guanajuato, Zihuatanejo, and Tahue Terranes were thrust over the calcareous platform rocks of Oaxaquia and the Central, Cortes, and Mixteca Terranes. The amount of tectonic transport apparently is significant, as xenoliths of Precambrian continental crust were found in Cenozoic volcanic rocks that erupted onto accreted rocks of the Guanajuato Terrane (Urrutia-Fucugauchi and Uribe-Cifuentes, 1999). Significant tectonic transport is also suggested by the amount of shortening that produced tight folding and major thrusting within the northern Zihuatanejo Terrane and the Arcelia and Teloloapan Terranes (Salinas-Prieto et al., 2000). In contrast, deformation of Cretaceous rocks in the southern parts of the Zihuatanejo Terrane formed wide regional anticlines, and some overturned folds and minor thrust faults locally. The structures generally trend NWSE, although locally some structures trend N-S and E-W. Santonian terrestrial sedimentation covers unconformably the previously deformed arc assemblages of the Zihuatanejo

Guerrero Composite Terrane of western Mexico Terrane (Benammi et al., 2005). Synorogenic sedimentary basins (Caracol Formation in Oaxaquia, and Mexcala Formation in the Mixteca Terrane) containing clasts derived from the Guerrero Composite Terrane suggest that these terranes were deformed and exhumed by that time. In addition, synorogenic sedimentation overlaps the Arcelia and Teloloapan terranes (Miahuatepec Formation), which suggests that these two terranes were also amalgamated during the same orogenic event (Mendoza and Suastegui, 2000; Guerrero-Suastegui, 2004). All these synorogenic basins range in age from Turonian to Maastrichtian. In addition, Paleocene granitoids along the coast cut the previously folded units of the Zihuatanejo Terrane and suggest a Late Cretaceous– early Paleogene deformation. Therefore, final amalgamation of the Guerrero Composite Terrane occurred between Santonian and Turonian–Maastrichtian time. DISCUSSION This section summarizes the stratigraphic, structural, and geochemical data that support the proposed stages for the tectonic evolution of western Mexico. Stage I: Origin of the Basement of the Tahue Terrane Exposures of pre-Cretaceous rocks in northwest Mexico are scattered; thus contact relationships among them can only be indirectly inferred (Figs. 1 and 4). Approximate distribution of the contacts among the terranes of western Mexico (Caborca, Cortes, and Tahue; Figs. 1 and 4) was outlined on the basis of the geographic distribution of pre-Cretaceous outcrops and lateral changes in the isotopic signatures of Cretaceous–Paleogene granitoids (Valencia-Moreno et al., 2001). Thus the nature of the contacts and the amount of displacement among different basements are unknown. In this section the main stratigraphic units that define the terranes are described following a NW to SE transect throughout the Paleozoic rocks of the Caborca, Cortes, and Tahue Terranes (Guerrero Composite Terrane). At the southern margin of the Caborca Terrane a thick shelfal limestone succession is exposed that contains Carboniferous– Permian fusulinids and other shallow-marine fossil fauna (Stewart et al., 1990). These rocks are overridden by a north-verging major thrust fault that places deeper marine sedimentary rocks of the Cortes Terrane on the shelfal rocks of the Caborca Terrane (Fig. 1; Coney and Campa, 1987; Poole and Madrid, 1988; Stewart et al., 1990). Basal metamorphic rocks are not exposed in the Cortes Terrane, but its basement has been interpreted as thinned Proterozoic rocks, perhaps the same as in the Caborca Terrane, or else Proterozoic metamorphic rocks different from those of the Caborca Terrane (McDowell et al., 1999; Valencia-Moreno et al., 1999; Valencia-Moreno et al., 2001). The deep-marine sedimentary rocks of the Cortes Terrane are sandstone and shale turbidites, graptolitic shale, chert, and layered barite that range in age from

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Ordovician to Devonian–Early Mississippian and were deformed during the Mississippian (Poole and Madrid, 1988; Stewart et al., 1990). These rocks are in turn overlain by Upper Carboniferous and Permian turbidites (Fig. 5; Poole and Madrid, 1988; Stewart et al., 1990; Poole et al., 2005). They all were deposited in a deep-marine environment and are interpreted to be part of the Paleozoic continental slope-rise deposits of western North America (Poole and Madrid, 1988; Stewart et al., 1990). All these units of the Cortes Terrane were deformed and thrust over the Caborca Terrane by Late Permian to Early Triassic time, and they are unconformably covered by Upper Triassic terrestrial and marine sedimentary rocks (Stewart et al., 1990). Therefore, the Caborca and Cortes Terranes were assembled by early Mesozoic time. The nature of the contact between the Cortes and Tahue Terranes (the latter belonging to the Guerrero Composite Terrane) has not been mapped in detail. It is inferred to be a thrust that verges toward the north, and it is probably north of El Fuerte town in Sinaloa State (Fig. 5), based on the northernmost exposures of Cretaceous marine volcanic rocks of the Guerrero Terrane (Servais et al., 1982; Henry and Fredrikson, 1987; RoldánQuintana et al., 1993; Freydier et al., 1995). The oldest Paleozoic rocks of the Tahue Terrane (Guerrero Composite Terrane) are exposed in the area of El Fuerte (El Fuerte Complex; Figs. 4 and 5). The El Fuerte Complex is formed by marine rhyolitic to andesitic lava flows and volcaniclastic rocks, interbedded with quartz-rich sandstone, shale, and thin-bedded limestone (Mullan, 1978; Roldán-Quintana et al., 1993; Poole and Perry, 1998). All these various components are deformed and metamorphosed to greenschist facies (Mullan, 1978; Roldán-Quintana et al., 1993). Sedimentary rocks of the El Fuerte Complex contain Ordovician conodonts (Poole and Perry, 1998). Preliminary geochemical analyses indicate a calc-alkaline island-arc affinity for the volcanic rocks of the El Fuerte Complex, and are similar to those from coeval Paleozoic arc rocks in the Klamath Mountains of Northern California (Lapierre et al., 1987). However, more detailed geochemical and geochronological work needs to be done to constrain their origin and relationships. Upper Paleozoic deep-marine sedimentary rocks are exposed south of El Fuerte, in San José de Gracia town, Mazatlán City, and other scattered localities in Sinaloa State (Figs. 4 and 5). These rocks belong to the San José de Gracia Formation (Carrillo-Martínez, 1971; Gastil et al., 1991; Arredondo-Guerrero and Centeno-García, 2003) and are made up of quartz-rich sandstone and shale turbidites, thin-bedded calcareous debris flows, black shale, and chert. The turbidites contain olistoliths of limestone with chert nodules, which in turn contain Middle Pennsylvanian to Early Permian fossils at the San José de Gracia locality (Carrillo-Martínez, 1971; Gastil et al., 1991). The San José de Gracia Formation has been interpreted as deposits in a deepmarine environment (Gastil et al., 1991). The contact between the El Fuerte Complex and the San José de Gracia Formation is not exposed. However, major differences in deformation and metamorphism (turbidites of the San José de Gracia Formation

Hermosillo

e a Terran Caborc e n a rr e Cortes T

Ciudad Obregón

S O N O R A

Navojoa 27°

Caborca Terrane Limestone (Lower Cretaceous) Shallow marine limestone and shale (Paleozoic) Cortes Terrane Terrestrial and marine sedimentary rocks (Triassic) Deep Marine turbidites, chert and barite (Paleozoic)

ne rra e e s T rran rte Te Co ue h a T

Sonobari El Fuerte C H

N

I

San Jose de Gracia

H

26°

U A H

Sinaloa de Leyva-Porohui

U A

D U R A

25°

N

Overlapping units

G

CULIACÁN

Recent sediments

O

Recent basalts Miocene-Pliocene Volcanic and sedimentary (conglomerate-sandstone) Granitoids (Paleocene-Oligocene) Upper Cretaceous-Oligocene Tarahumara Formation and Lower and Upper 24° Volcanic Supergroup

Tahue Terrane (Guerrero Composite Terrane) Granitoids (Upper Jurassic-Lower Cretaceous) Cretaceous Guerrero Arc Marine volcanic and sedimentary rocks San Francisco Gneiss (Triassic) 23°

MAZATLÁN

San Jose de Gracia Fm. (Carboniferous) El Fuerte Complex (Ordovician) km 0

109°

108°

107°

25

50

100

106°

Figure 4. Geologic map of Sinaloa and the southern Sonora states, showing the geology of Caborca, Cortes, and Tahue Terranes (after Carrillo-Martínez, 1971; Mullan, 1978; Gastil et al., 1978; Henry and Fredrikson, 1987; Stewart and Roldán-Quintana, 1991; Ortega et al., 1992; and our own work).

metamorphosed volcanic and quartz-rich turbidites, limestone, chert, basalts and rhyolites

El Fuerte Complex Ordovician

migmatized gneisses and amphibolites

Francisco gneiss Upper Triassic

unconformity

dike swarm

linked by provenance (Potosí Fan)

?

matrix is quartz-rich turbidites blocks of pillow lavas, chert, serpentinite and limestone accretionary complex

Taray Fm. Paleozoic-Triassic?

unconformity

felsic lavas and volcaniclastics continental Oxfordian

Transitional from continental to marine shale, sandstone, evaporites, limestone Kimmeridgian-Tithonian

Limestone Titonian to Aptian

Central Terrane

Hauterivian-Tithonian? deep marine volcaniclastic turbidites,chert, limestone

Arperos Fm.

unconformity

Albian-Aptian massive limestone, black limestone shales and sanstone

La Luz basaltic flows La Perlita and tuff limestone

limestone, chert, volcaniclastic turbidites felsic lavas 146 Ma

Esperanza Fm.

Tuna Mansa diorite Cerro Pelon tonalite Santa Ana

Guanajuato Terrane

matrix is quartz-rich turbidites blocks of pillow lavas accretionary complex

Zacatecas Fm. Norian-Carnian

unconformity

150-147 Ma

felsic and basic lavas

volcanic turbidites calcareous debris flows, tuff and chert pillowed lavas and dikes deformed and in part metamorphosed

La Borda, El Saucito and ChilitosFms. Lower Cretaceous Aptian?

Zihuatanejo Terrane

volcanic turbidites limestone, tuff and chert, pillowed lavas and dikes deformed and in part metamorphosed

Tahue Terrane

GUERRERO COMPOSITE TERRANE

deformed quartz-rich turbidites continent-slope facies

La Ballena Fm. Carnian-Norian

unconformity

Transitional from continental to marine shale, sandstone, evaporites, limestone Oxfordian-Tithonian felsic lavas and volcaniclastic rocks continental unknown age Felsic lavas and volcaniclastic rocks 189-172 Ma

Limestone Tithonian to Aptian

shale and sandstone foreland-basin fill

western Oaxaquia Caracol Fm.

EAST

Figure 5. Simplified stratigraphic columns for Oaxaquia and terranes north of the Transmexican Volcanic Belt. The columns are in an east-west order. They include the Tahue Terrane, which is part of the Guerrero Composite Terrane. Vertical scale shows the age range (in Ma).

Paleozoic

Ladinian Early Triassic

Carnian

Norian

Rhaetian

Hettangian

Sinemurian

Pliensbachian

Toarcian

Aalenian

Bathonian Bajocian

Callovian

Oxfordian

Kimmeridgian

Tithonian

Berriasian

Valanginian

Hauterivian

Barremian

Aptian

Albian

Cenomanian

Turonian

Coniacian

Santonian

Campanian

Maastrichtian

WEST

292

Centeno-García et al.

are strongly deformed but not metamorphosed) indicate that the contact is probably an angular unconformity. Both units of the Tahue Terrane (El Fuerte Complex and San José de Gracia Formation) are important because they can constrain the paleogeography of the northern Guerrero Terrane. Preliminary single-grain, detrital zircon geochronology from the quartz-rich sandstone from the turbidites of the San José de Gracia Formation shows populations that have a North American affinity (Centeno-García et al., unpublished data) and are similar to those from Paleozoic rocks in Baja California, in the Cortes Terrane, and in Nevada (Gehrels et al., 2002). The stratigraphy, geochemistry, and provenance of the Paleozoic rocks suggest that the Tahue Terrane (Guerrero Composite Terrane) was linked to the tectonic evolution of the western continental margin of North America, probably up to Permian–Triassic time. After that, there were major differences in the composition of the Mesozoic sedimentary cover of the Caborca-Cortes Terranes with respect to that of the Tahue Terrane. Therefore, it is likely that a fragment of previously accreted island-arc and continent-margin assemblages drifted from the continental margin sometime in the early Mesozoic. Contact relationships between the Paleozoic sedimentary rocks of the San José de Gracia Formation (Tahue Terrane) and the Triassic subduction-related complex of the Zihuatanejo Terrane are unknown because the contact is covered by younger rocks. However, the Tahue and Zihuatanejo Terranes share similar Cretaceous volcanic and sedimentary cover. Stages II and III: Triassic Potosi Fan and Its Accretion to the Continental Margin There are few exposures of Triassic rocks in Mexico, and they are limited to the Caborca and Cortes Terranes, western Oaxaquia, the Central and Zihuatanejo Terranes, and a small outcrop in the Vizcaíno Peninsula in Baja California. Triassic rocks have not been found in the Mixteca Terrane or in other terranes of Mexico. In this section we briefly describe Triassic rocks of the Cortes and Tahue Terranes (Barranca Group and Francisco Gneiss) and focus on the marine Triassic rocks of Oaxaquia (La Ballena Formation), the Central Terrane (Taray Formation), and the Guerrero Composite Terrane (Zacatecas Formation, and the Arteaga and Las Ollas Complexes). Barranca Group and Francisco Gneiss Triassic (Carnian–Norian) sedimentary rocks of the Cortes Terrane are made up of fluvial sandstone and shale that contain abundant coal layers (Barranca Group; Stewart and RoldánQuintana, 1991). These sediments were deposited unconformably on previously deformed Paleozoic deep-marine rocks. The Triassic fluvial deposits change transitionally up the column to shallow-marine siliciclastic deposits. These rocks have no evidence of contemporaneous volcanism. In contrast, Triassic rocks of the Tahue Terrane (Guerrero Composite Terrane) are made up of metamorphosed igneous rocks of the Francisco Gneiss near

Sonobari (Figs. 4 and 5; Mullan, 1978; Keppie et al., 2006). The Francisco Gneiss is made up of migmatized gneisses and amphibolites that have within-plate and continental tholeiite geochemical signatures (Keppie et al., 2006). This suggests that the Tahue and Cortes terranes may have been geographically separated by that time. La Ballena Formation Triassic rocks of Oaxaquia crop out on its western margin, near its boundary with the Guerrero Composite Terrane (Figs. 1 and 6). They are grouped as the La Ballena Formation (SilvaRomo, 1993; Silva-Romo et al., 2000), and their largest exposures are in the Peñón Blanco, Charcas, and Real de Catorce areas (Fig. 6; Silva-Romo, 1993; Tristán-González and TorresHernández, 1994; Barboza-Gudiño et al., 2004). The La Ballena Formation is made up of quartz-rich sandstone and shale, and scarce conglomerates deposited as small channel-fill lenses. The sedimentary structures of these Triassic rocks indicate deposition mostly by turbidity currents, although some debris flows and large slumps are present. This sequence contains abundant trace fossils and ammonites and bivalves of Late Triassic (Carnian) age at the Peñón Blanco and Charcas areas (Cantu-Chapa, 1969; Silva-Romo et al., 2000; Bartolini et al., 2002). Sedimentary structures and fossil fauna suggest that the deposition of this unit occurred in a submarine fan that developed on an external platform or continental slope setting. These rocks form part of the Potosi Submarine Fan (Centeno-García, 2005). The original thickness is unknown, but up to 4640 m was penetrated by exploration drilling without reaching the base of the succession (López-Infanzón, 1986). Taray Formation Similar marine siliciclastic rocks crop out at the Pico de Teyra region in the Central Terrane (Figs. 5 and 6). They belong to the Taray Formation, made up of highly deformed quartz-rich turbidites (sandstone and shale) interbedded with some black chert and scarce detrital limestone that contains fragments of crinoids, gastropods, corals, bivalves, and bryozoans (Diaz-Salgado et al., 2003). The Taray siliciclastic turbidites form a matrix within which blocks of black and green chert, pillow basalts, serpentinite, and crystallized limestone can be found (Figs. 5 and 6; Diaz-Salgado et al., 2003). The age of this unit remains undetermined; however, there are reports of fusulinids from one of the limestone blocks (Anderson et al., 1990). The youngest detrital zircons collected from the sedimentary matrix are Late Permian in age (Diaz-Salgado et al., 2003). There is also a report of molds of bivalves of possible Carnian age (Barboza-Gudiño et al., 1999; Bartolini et al., 2002). Thus deposition of the sedimentary matrix should have occurred between the Late Permian to the Late Triassic. The Taray Formation has a block-in-matrix structural style, formed by centimeter-size blocks to blocks of hundreds of meters in size, all in a highly sheared sedimentary matrix. This characteristic is typical of a subduction accretionary complex (Anderson et al., 1990, 2005; Diaz-Salgado et al., 2003).

Guerrero Composite Terrane of western Mexico 100°

101°

102°

293

Cenozoic cover

Caopas

Oaxaquia and Central Terrane

Concepción del Oro

U/Pb 158 Ma Pico de Teyra

Ju-K Calcareous Platform Ju felsic and intermediate lava flows, epiclastic rocks and redbeds

24°

Huiznopala Granjeno

Real de Catorce

Central terrane

Matehuala

Huizachal

U/Pb190 Ma

Tr Quartz-rich shale and sandstone turbidites (Potosí Fan) Tr(?) Accretionary complex blocks of pillow basalt, ultramafic, chert, and marble in a sedimentary quartz-rich matrix Proterozoic and Paleozoic metamorphic and sedimentary rocks Guanajuato Terrane Ju-K pillow basalts, deep marine volcanic turbidites, chert, and mafic and ultramafic plutons

Charcas

23°

Zacatecas U/Pb 147-150 Ma

Oaxaquia

Zihuatanejo Terrane Ju-K pillow basalts, volcanic turbidites, detrital limestone

Peñón Blanco San Luis Potosí

Tr Accretionary complex blocks of pillow basalt, in a sedimentary quartz-rich matrix

22°

o at aju e an ran Gu ter

Zihuatanejo terrane

100 km

U/Pb146 Ma Guanajuato Figure 6. Geologic map of central Mexico, showing main stratigraphic units of Oaxaquia and the Central, Zihuatanejo, and Guanajuato Terranes (modified from Ortega et al., 1992). Tr—Triassic; Ju—Jurassic; K—Cretaceous.

Zacatecas Formation The oldest rocks of the Zihuatanejo Terrane in its northernmost exposure are Triassic in age as well (Fig. 6). They make up the Zacatecas Formation, which crops out in a small tectonic window at the western margin of Zacatecas City (Fig. 6; Burckhardt and Scalia, 1906; Ranson et al., 1982; Cuevas-Pérez, 1983; Monod and Calvet, 1991). This formation is made up of quartzrich turbidites (sandstone and shale) that contain blocks of pillow basalts that have MORB geochemical signatures (Fig. 7; Centeno-García and Silva-Romo, 1997). The Zacatecas Formation contains fossil ammonites and bivalves of Late Triassic (Carnian) age (Burckhardt and Scalia, 1906; Bartolini et al., 2002). Its contact with the La Borda Formation of Late Jurassic(?)–Cretaceous age is inferred to have been originally an unconformity, but it was sheared and detached during Late Cretaceous thrusting and folding (Fig. 3). Rocks of the Zacatecas Formation show

structures associated with two distinct deformational events, one of them prior to the deformation that is recorded in the Cretaceous rocks as well. The small size of the outcrop prohibits constraints on the tectonic origin of the Zacatecas Formation, but its lava flows and siliciclastic turbidites are similar to those from the Arteaga Accretionary Complex, which is exposed in the southern part of the Zihuatanejo Terrane. Arteaga Complex More exposures of Triassic(?) rocks are found in the southern part of the Zihuatanejo Terrane (Fig. 7). Their largest outcrops are located in the Arteaga, Placeres del Oro, and Tiquicheo areas (Arteaga Complex) and near Zihuatanejo City (Las Ollas Complex) (Fig. 8; Centeno-García et al., 1993a, 1993b; Talavera-Mendoza et al., 1995; Mendoza and Suastegui, 2000). The Arteaga Complex is made up of quartz-rich turbidites (sandstone and shale),

unconformity

163-155 Ma

Arcelia Terrane

matrix is quartz-rich turbidites blocks of pillow lavas, chert, serpentinite and limestone accretionary complex

Arteaga Complex Norian-Carnian

poly deformed sedimentary and volcanic rocks continent-continent collision

felsic lavas and volcaniclastics continental affinity 127-133 Ma Las Lluvias Ignimbrite 168-179 Ma

unconformity Taxco Schist, Zicapa and Chapolapa

Limestone Aptian-Albian

Cuautla and Morelos Fms.

shale and sandstone foreland-basin fill

Mexcala Fm.

Acatlan Complex mid-Paleozoic

felsic and andesitic lavas and volcaniclastics shallow marine in the east, deep marine in the west 145-137 Ma

Limestone Aptian-Albian

TeloloapanTerrane

shallow marine limestone,volcaniclastic rocks, few andesitic-basaltic lava flows Villa de Ayala Fm. Aptian Albian MORB basaltic pillow lavas, IA lavas and Angao and San Lucas Fms. dikes, chert, deep volcanic turbidites, chert marine volcaniclastic felsic and basic lavas turbidites shallowing upward Early Cretaceous Neocomian

Comburindio and Mal Paso Fms.

Huetamo region

Cuale felsic lavas and Tumbscatío granitoids

volcaniclastic rocks and limestone age not well constrained unconformity

Albian-Cenomanian subaerial few marine volcaniclastic rocks, felsic lavas Aptian

subaerial to marine shale, sandstone, evaporites, limestone, basaltic-andesitic lava flows

Coastal region

Zihuatanejo Terrane

GUERRERO COMPOSITE TERRANE

EAST

Mixteca Terrane (western Continent-margin)

Figure 7. Simplified stratigraphic columns for the terranes described in the text that are south of the Transmexican Volcanic Belt, except for the Guanajuato Terrane, and include the Mixteca Terrane and the Teloloapan, Arcelia, and Zihuatanejo Terranes (Guerrero Composite Terrane). Vertical scale shows the age range (in Ma). MORB—mid-oceanic-ridge basalt; IA—island arc.

Paleozoic

Ladinian Early Triassic

Carnian

Norian

Rhaetian

Hettangian

Sinemurian

Pliensbachian

Toarcian

Aalenian

Bathonian Bajocian

Callovian

Oxfordian

Kimmeridgian

Tithonian

Berriasian

Valanginian

Hauterivian

Barremian

Aptian

Albian

Cenomanian

Turonian

Coniacian

Santonian

Campanian

Maastrichtian

WEST

10

0

Tecoman

20

PA

Aquila

km

Colima

Nevado de Colima

Jilotlán

40

CIF IC OC EA

Coalcoman

Pijuamo Tepalcatepec

Tecalitlán

103°

N

Infiernillo Dam

Zihuatanejo

iver as R Bals

Lázaro Cárdenas

Arteaga

Ar/Ar 152-158 Ma

Tumbiscatío

U/Pb 163+3 Ma

Playa Azul

Aguililla

101°

Balsas

River

Huetamo

Arcelia

Zitácuaro

100°

XOLAPA TERRANE

Iguala

Roads

Chapolapa

Chilpancingo

Zicapa

MIXTECA TERRANE

Teloloapan

Taxco

State limit

Rivers

inferred terrane boundary (strike-slip fault)

inferred terrane boundary (thrust fault)

Observed terrane boundary (thrust fault)

Cenozoic volcanic and sedimentary rocks

Upper Cretaceous to Paleogene granitoids

CutzamalaFormation Santonian-Maastrichtian red beds (overlaps Zihuatanejo and Arcelia terranes)

Miahuatepec Formation Upper Cretaceous clastic rocks foreland basin-fill (overlaps Arcelia and Teloloapan terranes) Mexcala Formation Turonian-Maastrichtian clastic rocks foreland basin-fill (overlaps Mixteco and Teloloapan terranes)

Overlapping assemblages

Tejupilco 19°

Jurassic to Cretaceous migmatites, gneisses, and plutons

Xolapa Terrane

Paleozoic Acatlán Complex

Z I H U ATA N E J O T E R R A N E

Nueva Italia

102°

Cretaceous MORB pillow lavas, fine-grained volcanic turbidites and chert (deep marine)

Cretaceous IAB pillow lavas, fine-grained volcanic turbidites and chert (deep marine)

Arcelia Terrane

Lower Cretaceous continental arc assemblages, marine and terrestrial rhyolite to andesite lava flows and epiclastic rocks, quartz-rich clastic rocks

Aptian-Albian Calcareous Platform

Mixteca Terrane

18°

Figure 8. Geologic map of southwestern Mexico, showing the simplified geology of the Mixteca, Teloloapan, Arcelia, and Zihuatanejo Terranes (after Campa and Ramirez, 1979; Ortega et al., 1992; Talavera-Mendoza et al., 1995; Corona-Chávez and Israde-Alcántara, 1999; Mendoza and Suastegui, 2000; Centeno-García et al., 2003). IAB—island-arc basalt; MORB—mid-oceanic-ridge basalt.

N

Manzanillo

Minatitlán

Zapotitlán

104°

Arteaga Subduction Complex (Triassic)

Coastal Cretaceous arc assemblage shallow marine and terrestrial rhyolite, andesite and some basalt, limestone, volcaniclastic and basement-derived clastic rocks Middle to Upper Jurassic plutons Las Ollas Subduction Complex (age unknown)

Lower Cretaceous arc assemblages shallow marine andesitic to basaltic lava flows and volcaniclastic rocks, massive and reefal limestone (east) deep marine lava flows, calcareous debris flows and volcanic turbidites (west)

Teloloapan Terrane

deep marin

Huetamo area, Cretaceous arc assemblage, mostly marine volcaniclastic rocks, limestone and some pryroclastic and lava flows

TERRANE

ARCELIA

e

TELOLOAPAN TE RRANE

Zihuatanejo Terrane

marine shallow

296

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black and green chert, and mafic tuff that form a matrix that contains blocks and slabs of pillow basalts, diabase, banded gabbros, chert, and limestone, all deformed in a block-in-matrix structural style (Centeno-García et al., 2003). Chert layers contain radiolarians of Triassic (Ladinian–Carnian) age (Campa et al., 1982). Pillow basalts and gabbros have oceanic geochemical signatures (MORB; Centeno-García et al., 1993a; Centeno-García et al., 2003). Sedimentary structures preserved in some exposures of unmetamorphosed turbidites, along with the affinity of the few fossils found in the sedimentary rocks of the matrix, suggest that the sequence was deposited in a deep-ocean environment. Apparently the quartz-rich turbidites were contemporaneous with oceanic magmatic activity, as they are interbedded with volcaniclastic rocks (Centeno-García et al., 2003). The block-in-matrix style of deformation of the Arteaga Complex, as well as its lithological associations, indicate that it was formed in a subduction accretionary prism. Metamorphism ranges from none to amphibolite facies; blueschist facies has not been found in the area. Las Ollas Complex The Las Ollas Complex forms part of the Zihuatanejo Terrane and is exposed near Zihuatanejo City (Figs. 7 and 8; Talavera-Mendoza, 2000). This complex is a tectonic mélange formed by highly sheared blocks of metabasalt, banded and massive gabbro, metadolerite, ultramafic rocks, and shale and quartzrich sandstone (Talavera-Mendoza, 2000). These blocks are enveloped in a highly sheared clastic (quartz-rich sandstone) or serpentinitic matrix (Talavera-Mendoza, 2000). Blueschist facies were reported by Talavera-Mendoza (1993, 2000). Geochemical compositions of the basalts are typical of MORB and primitive oceanic-arc magmas (Talavera-Mendoza, 2000). 40Ar/39Ar and K/Ar ages obtained from amphibole from several metagabbro blocks range from 223 Ma to 96 Ma (Permian to early Cenomanian) (Delgado, 1982; A. Iriondo, 2003, personal commun.). This has been interpreted to be the subduction complex of the Cretaceous arc (Vidal-Serratos, 1991; Talavera-Mendoza, 1993); however, its contact relationships with Cretaceous arc-related rocks, and similarities with the Arteaga Complex, suggest an earlier origin. Quartz-rich turbidites from the La Ballena Formation of Oaxaquia (continental Mexico), the matrix of the Taray Formation of the Central Terrane, and the Arteaga and Las Ollas Complexes and the Zacatecas Formation of the Zihuatanejo Terrane (Guerrero Composite Terrane) have similar and distinctive compositions and detrital-zircon age populations (Centeno-García et al., 2005; Talavera-Mendoza et al., 2007). Therefore, Triassic sedimentation of the central and western terranes of Mexico is linked by provenance. The youngest zircon age populations from all the samples (latest Permian) are much older than the depositional ages of the turbidites (Carnian–Norian), which means that there was no active volcanism at that time. In other words, there is no evidence of Late Triassic continental arc volcanism in Mexico. Zircon age populations of the Potosi Fan are different from those of Triassic quartz-rich sandstone from the Caborca

and Cortes Terranes (González-León et al., 2005) but are similar to those from Triassic fluvial sedimentary rocks of Arizona (Anderson, 2006). This suggests that at the end of the Triassic the terranes of central and western Mexico may have been to the north of their present locations. Based on this evidence, we propose that the margin of the western paleo-continent of Mexico was passive or rifting at the end of the Triassic. This passive margin received abundant clastic sedimentation, forming the large Potosi Fan. Sediments of this fan were deposited on oceanic crust (Arteaga Basin in Fig. 3). When subduction started, slivers of the ocean floor were tectonically mixed with the already existing passive-margin quartz-rich turbidites that were forming the Taray and Zacatecas Formations as well as the Arteaga and Las Ollas Complexes. Whether the ocean basin that was covered by sediments of the Potosi Fan was an active marginal oceanic basin, a marginal backarc basin, or an open ocean–continent flank is still uncertain. The only potential evidence of association of the Potosi Fan sediments with tholeiitic oceanic volcanism is the volcaniclastic rocks interbedded with the siliciclastic turbidites in the Arteaga Complex, as the volcaniclastic rocks have geochemical signatures between primitive island arc and MORB (Centeno-García et al., 2003). At least two phases of deformation are found in all the Triassic rocks of Oaxaquia and the Central and Zihuatanejo Terranes. The first event comprised strong shearing and tight folding, and the block-in-matrix structures. A second event was recorded only in the Arteaga Complex. This event deformed the Jurassic granitoids as well, and it is characterized by a mylonitic fabric. The third event was common to all the Triassic units and is also recorded in the Jurassic and Cretaceous cover sediments, and it is characterized by axial cleavage, open to tight folding, reverse faulting, and thrusting. The time of accretion of the Central Terrane with Oaxaquia is assumed to have been prior to the Middle Jurassic, because the La Ballena Formation of Oaxaquia and the Taray Formation of the Central Terrane were deformed and locally metamorphosed prior to deposition of Upper Jurassic terrestrial volcanic and clastic formations (Tristán-González and Torres-Hernández, 1994; Jones et al., 1995; Silva-Romo et al., 2000). The Zihuatanejo Terrane (Guerrero Composite Terrane) was also accreted at that time, because the Arteaga Complex is cut by granitoids of Middle Jurassic age as well (Centeno-García et al., 2003). The subduction zone that formed the Taray and Zacatecas Formations, and the Arteaga and Las Ollas Complexes, was probably constructed along the continental margin of Oaxaquia in Early Jurassic time. Whether the subducting slab was dipping to the east or to the west has not been determined. Stage IV: Jurassic Continental Arc of Western Mexico Erosion and exhumation of the accreted continental slope sediments and the accretionary complexes occurred prior to the initiation of Middle to Late Jurassic magmatism. This is indicated by the major angular unconformity that separates the Jurassic arc

Guerrero Composite Terrane of western Mexico succession from the deformed Triassic rocks of Oaxaquia and the Central and Zihuatanejo Terranes. Jurassic arc magmatism has also been identified in the Mixteca Terrane. The Jurassic arc rocks have different names at different locations; they are hereby described by their occurrence in different terranes: Nazas, Huizachal, and La Joya Formations in Oaxaquia The La Ballena Formation (Oaxaquia) is unconformably overlain by the volcanic rocks and red beds of the Nazas Formation in the Peñón Blanco, Charcas, and Real de Catorce areas (Figs. 5 and 6; Silva-Romo, 1993; Tristán-González and TorresHernández, 1994; Barboza-Gudiño et al., 2004). The Nazas Formation is made up of dacitic and minor rhyolitic and andesitic lava flows and pyroclastic flows, dikes, and porphyritic shallow intrusives. The volcanic rocks are interbedded with conglomerate, sandstone, and scarce paleosols. Conglomerate is formed mostly by volcanic clasts and a few clasts of sandstone and shale derived from the underlying La Ballena Formation. The volcaniclastic conglomerate and sandstone form lens-shaped bedding with low-angle cross-bedding, interbedded with some debris flows, suggesting that they were deposited in a terrestrial (fluvial and alluvial fan) environment. Although their age has not been well constrained at all the exposures, there is a report of U/Pb ages as old as 189 Ma at a subaerial volcanic-sedimentary succession in Huizachal (Huizachal Formation; Fastovsky et al., 2005), which might not belong to the same volcanic arc event (Figs. 5 and 6). Rocks of the Nazas Formation at Real de Catorce yielded U/Pb ages of 172 ± 5 Ma (Barboza-Gudiño et al., 2004). The Nazas Formation changes transitionally upward to shallow-marine volcaniclastic rocks, some evaporites, and thin-bedded limestone, which in turn become a thick limestone succession in the Peñón Blanco and Charcas areas (Fig. 6). The basal part of this limestone succession contains late Oxfordian–Kimmeridgian fossil faunas (Centeno-García and Silva-Romo, 1997). In contrast, there is an internal angular unconformity in the Real de Catorce locality (Fig. 6), which separates in two units, the terrestrial volcanic and sedimentary successions (Nazas and La Joya Formations; Barboza-Gudiño et al., 2004). The upper La Joya Formation changes transitionally upward to shallow-marine volcanic sandstone and shale interbedded with thin limestone strata. The oldest fossils reported from the base of the limestone succession in Real de Catorce are Oxfordian in age (Barboza-Gudiño et al., 2004). Caopas, Rodeo, and Nazas Formations of the Central Terrane The volcanic cover of the Taray Formation (Central Terrane; Figs. 5 and 6) belongs to the Caopas, Rodeo, and Nazas Formations (Córdoba-Méndez, 1964; López-Infanzón, 1986; Jones et al., 1995). The Rodeo and Nazas Formations are lateral equivalents of the same rocks but named differently in separate outcrops (Díaz-Salgado, 2004). Both units are made up of rhyolitic to andesitic lava flows and dikes, and pyroclastic deposits that are interbedded with fluvial sedimentary rocks, mostly sandstone and conglomerate (Anderson et al., 1990, 1991; Jones et al., 1995;

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Díaz-Salgado, 2004). The Caopas Formation was formed by shallow porphyritic intrusives. Felsic volcanic rocks of the Rodeo Formation yielded a K-Ar age of 183 Ma (López-Infanzón, 1986), and the Caopas Formation a U/Pb age of 158 Ma (Jones et al., 1995). Terrestrial volcaniclastic rocks of the Rodeo Formation are interpreted to have been deformed previous to the deposition of late Oxfordian limestone (Anderson et al., 1991; Bartolini et al., 2002). However, in another locality nearby the Nazas Formation changes transitionally upward to shallow-marine calcareous rocks that range in age from Late Jurassic to Late Cretaceous (Córdoba-Méndez, 1964; Díaz-Salgado, 2004). All these volcanic-sedimentary units are interpreted in this work as the first overlapping succession that stitches the Central Terrane with Oaxaquia. The Caopas and Rodeo Formations, as well as the Nazas Formation, are interpreted as continental intraarc assemblages (Jones et al., 1995). Las Lluvias Ignimbrite of the Mixteca Terrane Jurassic arc volcanism was also recorded in the Mixteca Terrane in which ignimbrites, interbedded with fluvial and shallow-marine siliciclastic deposits, yielded U/Pb ages of 168.2 ± 1.2 Ma, 177.3 ± 1.5 Ma, and 179.1 ± 1.5 Ma (Campa and Iriondo, 2003). Cuale Assemblage and Tumbiscatio Granitoids of the Zihuatanejo Terrane Evidence of coeval Jurassic magmatism has been found in two localities in the Zihuatanejo Terrane (Guerrero Composite Terrane). One of the exposures is NE of Puerto Vallarta City, in the Cuale mining district, and the other locality is in the Tumbiscatio region, both along the Pacific Coast (Figs. 7 and 8). Rocks at Cuale contain volcanogenic massive sulfide (VMS) deposits and are composed of submarine rhyolitic lavas and tuffs, volcanic sandstone with evolved-arc geochemical affinity (Bissig et al., 2003), and shale that yielded U/Pb ages of 162.4 and 155.9 Ma (Bissig et al., 2003). These rocks are strongly deformed and partially metamorphosed, and their contact with Cretaceous unmetamorphosed marine volcanic and sedimentary successions has not been determined. Two Jurassic granitoids crop out in the Tumbiscatio region. They were emplaced in previously deformed sedimentary rocks of the Arteaga Complex, and vary in composition from granodiorite to granite to quartz monzonite. Their geochemical compositions are typical of calc-alkaline subduction-related granites, which are more evolved than granitoids of Cretaceous and Cenozoic ages from the same area. Both granitoids show intense shearing and internal deformation. Grajales and López (1984) obtained one K/Ar date of Late Jurassic age (158 Ma). U/Pb isotopic analysis yielded a 163 Ma age, and Ar/Ar ages are 158 and 152.4 Ma (Centeno-García et al., 2003). The igneous rocks of the Cuale and Tumbiscatio regions have strong similarities in geochemical composition and age with volcanic rocks of the Central Terrane (Caopas, Rodeo, and Nazas); thus we suggest that they probably originated in the same volcanic arc. Therefore,

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the Arteaga Complex was probably accreted to the continental margin, either near or along the strike from central Mexico. The Jurassic volcanic event did not produce a thick stratigraphic column and apparently did not have large volumes of volcanic products. The column changes transitionally upward to shallow-marine calcareous rocks. Therefore, the lithologic associations and vertical facies changes of this volcanic-sedimentary event are similar to those of a continental rift. However, the scarce geochemical analyses from its volcanic rocks suggest an arc setting (Jones et al., 1995). These rocks have been interpreted as the southern continuation of the Jurassic continental arc that developed along the southwestern margin of North America (Jones et al., 1995). It has been proposed that major strike-slip faults were probably active during the arc activity (Mojave-Sonora Megashear; Jones et al., 1995). This could explain the fact that the Potosi Fan is south of its possible continental fluvial correlative in Arizona, as well as the southward displacement of the Tahue Terrane. Whether or not the Jurassic volcanic event was coeval with a major transform fault has not been well documented. The evidence in favor of an important synsedimentary deformation involving major extension is as follows: (1) Minor synsedimentary normal faults and local angular unconformities are present within the Jurassic arc volcanic and sedimentary successions, and pre-Cretaceous mylonitic shearing is recorded in the Jurassic granitoids of the Tumbiscatio region (Zihuatanejo Terrane). (2) Arc magmatism suddenly ceased in the Central Terrane and Oaxaquia, followed by a rapid transgression recorded in a few meters of transitional sedimentation. (3) Subsidence rates apparently were significant during the early stages of the Oxfordian–Kimmeridgian marine sedimentation, because the calcareous rocks show evidence of deeper sedimentation at higher stratigraphic levels as well as overall rapid sedimentation. (4) Although fault planes have been obliterated by younger deformational events, they have been inferred by the rapid lateral changes in thickness and facies of the calcareous succession through the interval from the end of the Jurassic to the Early Cretaceous. (5) In addition, major regional lineaments have been identified in central and eastern Oaxaquia, including the San Marcos and La Babia Faults (Fig. 1) (Goldhammer, 1999; Chávez-Cabello et al., 2005). Stage V: Rifting of the Guerrero Terranes and Formation of a Complex Arc System In this section we list the main stratigraphic features of the volcanic-sedimentary successions of the Guerrero Composite Terrane and the Mixteca Terrane. Arc volcanism was absent in Oaxaquia and the Central Terrane through the end of the Jurassic and the Cretaceous. During this period, oceanic crust was emplaced toward the east of Oaxaquia in the Gulf of Mexico, and continuous subsidence prevailed throughout the Early Cretaceous, resulting in a thick calcareous platform that covered all the Central Terrane and Oaxaquia.

Although much detailed work needs to be done in order to reconstruct the paleogeography of western Mexico during the Cretaceous, the available evidence indicates three important features: 1. Magmatism prograded generally east to west through time, from the oldest ages in the Oaxaquia and Mixteca Terranes to the youngest ages in the coastal Zihuatanejo Terrane. There is some overlap of age ranges for the volcanism among the different terranes, e.g., volcanism of the Mixteca Terrane overlaps in age with part of the volcanism of the Teloloapan Terrane (Guerrero Composite Terrane). However, on a large scale, Albian–Cenomanian volcanism is absent in the Mixteca and the Teloloapan Terranes, and it is widespread in the coastal region of the Zihuatanejo and Arcelia Terranes. 2. Magma chemistry changed through time toward a more primitive melt. The Middle Jurassic volcanic and intrusive rocks in all the terranes show mostly felsic-evolved continental-arc geochemical signatures, including the Mixteca Terrane and Oaxaquia. In contrast the Cretaceous volcanic rocks of the Guerrero Composite Terrane range from tholeiitic basalts to andesites, with few rhyolites. They show more primitive island-arc (IA) geochemical signatures overall, and some even have MORB to oceanicisland basalt (OIB) signatures. The Mixteca Terrane is the exception to this trend; its magmatism remained evolved, with continental arc signatures, into the Cretaceous. 3. Within different assemblages of the Guerrero Composite Terrane there are major differences in the stratigraphy, sediment composition, and depositional environments. And the Guerrero Composite Terrane overall is different from the volcanic-sedimentary rocks of the Mixteca Terrane to the east. In their present distribution, areas with shallow-marine and terrestrial volcanic-sedimentary successions alternate with areas with deep-marine volcanicsedimentary successions, and suggest a complex paleogeography for that time. These three features are hereby interpreted as evidence of intra-arc rifting-translation. We propose, as a hypothesis to be tested, that the subduction zone might have migrated to the west. This would have produced thinning of the crust, which in turn would have originated more primitive IA geochemical signatures of the magmas and promoted the development of deep basins. Whether the amount of extension was large enough to develop oceanic basins and several parallel subduction zones has not been determined. The stratigraphy, depositional environments, age, and geochemical affinities of the main units are summarized by terrane. First, those of southern Mexico are described, following a section from east to west. Next, Cretaceous rocks of the northern terranes are described from east to west as well. Mixteca Terrane Three localities with Early Cretaceous volcanism have been identified in the western Mixteca Terrane near the contact with

Guerrero Composite Terrane of western Mexico the Guerrero Composite Terrane: the Taxco Schist and the Chapolapa and Zicapa Formations (Fries, 1960; de Cserna and Fries, 1981; Talavera-Mendoza, 1993; Campa and Iriondo, 2003; Fitz et al., 2002). The Taxco Schist is made up of submarine andesitic to rhyolitic lava flows and tuffs interbedded with epiclastic rocks and quartz-rich sandstone and shale (de Cserna and Fries, 1981; Talavera-Mendoza, 1993). Its volcanic rocks have a continentalarc geochemical affinity, more evolved than contemporaneous magmatism from the Guerrero Composite Terrane (TalaveraMendoza; 1993; Centeno-García et al., 1993a). The Zicapa Formation is made up of dacitic to rhyolitic lava flows interbedded with fluvial deposits (Fitz et al., 2002). The Chapolapa Formation is composed mostly of marine lava flows and epiclastic rocks. The abundance of quartzites within the volcanic-sedimentary successions of the Taxco Schist and Zicapa Formation suggests that a crystalline basement was exposed during the arc activity. U/Pb dating of lavas from the Taxco Schist by sensitive high-resolution ion microprobe (SHRIMP) methods yielded 130–131 Ma ages (Campa and Iriondo, 2004), and from the volcanic-volcaniclastic rocks of the Zicapa Formation, 127 Ma (Fitz et al., 2002). Lava flows from the Chapolapa Formation have 129–133 Ma SHRIMP U/Pb ages. The Taxco Schist shows one phase of deformation and metamorphism prior to the deposition of Aptian–Albian carbonates. Thus, Early Cretaceous volcanic rocks of the Mixteca Terrane are unconformably covered by a carbonate platform that ranges in age from Early to middle Cretaceous (Fries, 1960). The limestone succession in the western Mixteca Terrane changes upward to a thick clastic succession (Mexcala Formation) of Turonian to Maastrichtian age (Guerrero-Suastegui, 2004). The Mexcala Formation is made up of alternating sandstone, shale, and conglomerate, deposited in deltaic and submarine-fan environments (Figs. 7 and 8). It is a synorogenic deposit (foreland basin-fill) associated with regional thrusting and folding of both the Guerrero Composite and Mixteca Terranes at the end of the Cretaceous. Therefore, the Mexcala Formation is the first overlapping assemblage that stitches the Guerrero Composite Terrane and the Mixteca Terrane, and marks the final amalgamation of the Guerrero Composite Terrane to continental Mexico. Teloloapan Terrane The Teloloapan Terrane (Figs. 1 and 8) is exposed in the easternmost parts of the Guerrero Composite Terrane. This terrane is characterized structurally by a complex east-vergent thrust-fault system. Its rocks are severely deformed and metamorphosed to low-grade greenschist facies. The Teloloapan Terrane overrides either Cretaceous platform carbonates or Upper Cretaceous siliciclastic rocks that belong to the Mixteca Terrane (Figs. 7 and 8; Talavera-Mendoza et al., 1995). The nature of the basement of the Teloloapan Terrane remains unknown. Metamorphic rocks of the Tejupilco area (Fig. 8) were interpreted as a possible basement for the Teloloapan Terrane by Elías-Herrera and Sánchez-Zavala (1990), and Sanchez-Zavala (1993). These authors suggested that the

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Tejupilco volcanic-sedimentary sequence might represent an arc assemblage older than the rest of the Guerrero Terrane magmatism. They based this conclusion on U-Pb dates from associated sulfide deposits. The ages they obtained vary broadly from Carnian (227 Ma) to Oxfordian (156 Ma). However, the same volcanic-sedimentary rocks were considered a part of the Cretaceous arc assemblage by other authors (Campa and Ramirez, 1979; Talavera-Mendoza et al., 1995). The arc assemblage of the Teloloapan Terrane consists of two distinct regions with different volcanic and sedimentary rocks. The eastern region is characterized by shallow-marine deposits, and the western region is composed of deeper facies (Guerrero-Suastegui et al., 1991; Ramírez-Espinoza et al., 1991; Talavera-Mendoza et al., 1995; Mendoza and Suastegui, 2000; Guerrero-Suastegui, 2004).The stratigraphy of the eastern region, from base to top, is made up of basaltic to andesitic pillow and massive lava flows, volcanic breccias, and pyroclastic flow deposits (Villa de Ayala Formation; Talavera-Mendoza et al., 1995). These deposits are interbedded with epiclastic sandstone and conglomerate. Primary structures in the volcaniclastic rocks suggest a marine depositional environment (Guerrero-Suastegui et al., 1991; Guerrero-Suastegui, 2004). Storm deposits, coral fragments, and other fossils suggest shallow and warm waters. This unit contains fossil gastropods and bivalves that range in age from Hauterivian to Aptian (Guerrero-Suastegui et al., 1991; RamírezEspinoza et al., 1991; Talavera-Mendoza et al., 1995). Geochemical analyses of volcanic rocks of the Villa de Ayala Formation of the Teloloapan Terrane indicate that the magmatism is calc-alkaline and similar to that of active intraoceanic arcs (Talavera-Mendoza, 1993; Talavera-Mendoza et al., 1995; Lapierre et al., 1992; Mendoza and Suastegui, 2000; Centeno-García et al., 1993a). The base of the Villa de Ayala Formation is not exposed. The maximum thickness is considered to be ~3000 m (Guerrero-Suastegui, 2004). The volcanic succession of this formation changes transitionally upward to thick, massive reefal limestone of the Teloloapan Formation. At the base the Teloloapan Formation is composed of intertidal limestone interbedded with volcaniclastic rocks containing rudists and nerineas of late Aptian–early Albian age (Guerrero-Suastegui et al., 1991, 1993; Guerrero-Suastegui, 2004). Thus magmatism ceased prior to the late Aptian (Guerrero-Suastegui et al., 1991; Mendoza and Suastegui, 2000; Guerrero-Suastegui, 2004). The Teloloapan Formation grades upward into the Pachivia Formation of Turonian age, which is made up of shale and fine-grained sandstone and shale. The Pachivia Formation is the western equivalent of the Mexcala Formation of the Mixteca Terrane and indicates that the Teloloapan and Mixteca Terranes were already in close proximity (Guerrero-Suastegui et al., 1991; Talavera-Mendoza et al., 1995; Guerrero-Suastegui, 2004). The stratigraphy of the western part of the Teloloapan Terrane comprises submarine basaltic, andesitic, and felsic lava flows and volcaniclastic rocks (Villa de Ayala Formation) deposited in deeper water conditions than the sediments of the eastern Teloloapan Terrane. It is in transitional contact upsection with

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the Acapetlahuaya Formation, composed of thin-bedded volcanic shale and sandstone at the base, at some localities interbedded with dark, thinly laminated limestone. It changes transitionally upward to shale, with little or no volcanic material at the top (Campa and Ramirez, 1979; Guerrero-Suastegui, 2004). This unit has been highly tectonized, making it difficult to calculate its original thickness and contact relationships. Apparently, the Acapetlahuaya Formation changes laterally toward the west and overlies transitionally the volcaniclastic deposits of the Villa de Ayala Formation. Its upper contact with the Amatepec Formation is highly tectonized. The Acapetlahuaya Formation contains ammonoids, and radiolarians that are late Aptian in age (Campa et al., 1974; Guerrero-Suastegui et al., 1993; Talavera-Mendoza et al., 1995; Guerrero-Suastegui, 2004). The Amatepec Formation is made up of thin-bedded black detrital limestone and is devoid of volcanic material. It is interpreted as deep-basin–slope deposits. This formation is tightly folded and overlies either the Villa de Ayala or the Acapetlahuaya Formation. It is late Albian to early Cenomanian in age, based on calcispherulids, planktonic foraminifers, and radiolarians (Campa and Ramirez, 1979; Guerrero-Suastegui et al., 1991, 1993; Talavera-Mendoza, 1993; Talavera-Mendoza et al., 1995). The deep-marine limestone is overlain by turbiditic sandstone-shale successions of the Miahuatepec Formation (Talavera-Mendoza et al., 1995). Fossils have not been found, but it is at least post–early Cenomanian because of its stratigraphic position. The Miahuatepec Formation was deposited, during the amalgamation of the Zihuatanejo, Arcelia, and Teloloapan Terranes, in a thrust-related basin (Guerrero-Suastegui et al., 1991; Ramírez-Espinoza et al., 1991; Talavera-Mendoza et al., 1995; Mendoza and Suastegui, 2000; Guerrero-Suastegui, 2004). The ages of magmatism of the Teloloapan Terrane have been poorly constrained by the limited fossils found in the volcaniclastic levels. A few U/Pb isotopic ages from felsic lavas at the base of the succession range in age from 137.4 to 145.9 Ma (Tithonian–Hauterivian; Mortensen et al., 2003). Thus, magmatism of the Teloloapan Terrane is in part contemporaneous with that of the Mixteca, Guanajuato, and Zihuatanejo Terranes. There are three distinctive differences in the Cretaceous stratigraphy between the Mixteca and Teloloapan Terranes: (1) Volcanism of the Mixteca Terrane is more evolved, and its isotopic signatures show influence of old continental crust in the magma generation. In contrast, volcanism of the Teloloapan Terrane is more primitive and has no traces of contamination by old continental crust (Centeno-García et al., 1993a; Talavera-Mendoza et al., 1995; Mendoza and Suastegui, 2000). (2) Metamorphic and quartz clasts are abundant (up to 70%) in the sandstones that are interbedded with volcanic rocks in the Mixteca Terrane but are absent throughout the stratigraphic column of the Teloloapan Terrane. (3) Magmatism ceased in the Mixteca Terrane before the Aptian, and part of the volcanic-sedimentary succession was deformed and metamorphosed (Taxco Schist). In contrast, volcanism continued in the Teloloapan Terrane until Aptian–Albian time, and no internal deformation has been identified.

The arc volcanism of the Mixteca and Teloloapan Terranes has been interpreted as part of a single arc-backarc system in which volcanism of the Mixteca Terrane would be the backarc basin (Cabral-Cano et al., 2000; Monod et al., 1994). An alternative interpretation is that these two terranes belong to different arcs, separated by a double-dipping subduction of an oceanic basin (Guerrero-Suastegui, 2004). Arcelia Terrane Thrust over the Teloloapan Terrane is the Arcelia Terrane (Guerrero Composite Terrane), which shows deeper marine facies and less evolved magmatism than the rest of the arc successions of the Guerrero Composite Terrane (Talavera-Mendoza et al., 1995; Mendoza and Suastegui, 2000). This terrane is made up of basaltic pillow lavas and ultramafic bodies, black shale and chert, and fine-grained volcanic turbidites (Fig. 7), all intensively deformed and partly metamorphosed (Ramírez-Espinoza et al., 1991; Talavera-Mendoza et al., 1995). The chert layers contain radiolarians reported as Albian–Cenomanian in age (Dávila and Guerrero, 1990; Ramírez-Espinoza et al., 1991). Ar/Ar and K/Ar ages (93.4–105 Ma; Delgado et al., 1990; Ortiz and Lapierre, 1991; Elías-Herrera, 1993) are compatible with biochronology, but detrital zircon ages from volcanic turbidites are older (mean age, 130 Ma; Talavera-Mendoza et al., 2007). Geochemical signatures of the Arcelia magmas are similar to those in recent primitive IAs and oceanic basins (MORB) (Talavera-Mendoza, 1993; Talavera-Mendoza et al., 1995; Mendoza and Suastegui, 2000). There are no exposures of older rocks in the Arcelia Terrane, and no clasts of older metamorphic or sedimentary rocks have been found in its sedimentary strata. Mendoza and Suastegui (2000) suggest that this terrane is entirely oceanic, that it may have originated as an independent oceanic arc and backarc basin, and that it represents partly developed oceanic crust. An alternative interpretation is that the Arcelia Terrane could also be a backarc basin of the Zihuatanejo Terrane (Centeno-García et al., 2003a). Southern Part of the Zihuatanejo Terrane Uppermost Jurassic–Cretaceous volcanic-sedimentary assemblages of the Zihuatanejo Terrane can be grouped in three main regions: northern Zihuatanejo Terrane (Zacatecas area), Huetamo area, and coastal Zihuatanejo-Colima region (Figs. 7 and 8). The uppermost Jurassic to Cretaceous strata of the southern Zihuatanejo Terrane are not as strongly deformed as those in other terranes, and original contact relationships and complete stratigraphic columns are well preserved. The strata are characterized by numerous lateral facies changes and internal erosional and angular unconformities. The geographic distribution of the facies is highly irregular, and it has not yet been determined in detail. Therefore, compiling, correlating, and synthesizing the stratigraphy of the area is difficult because it varies considerably from one locality to another. The stratigraphy of the northern Zihuatanejo Terrane (Zacatecas area) is described later. The stratigraphic column of the southern Zihuatanejo Terrane in the Huetamo area is made up of Triassic basement rocks

Guerrero Composite Terrane of western Mexico of the Arteaga Complex, overlain by uppermost Jurassic to Cretaceous volcanic and sedimentary cover. These rocks are thrust over the Arcelia Terrane (Figs. 7 and 8). Arc-related rocks of the Huetamo region (Figs. 7 and 8) overall have been formed by a thick succession of alternating shale, sandstone, and conglomerate, with scattered basaltic pillow lavas, submarine ignimbrite flows, and other intermediate pyroclastic and epiclastic flows in the lower parts of the succession (Angao and San Lucas Formations; Pantoja, 1959; Guerrero-Suastegui, 1997). These arcrelated rocks lie unconformably on the Arteaga Complex (Figs. 7 and 8). Fossils of Late Jurassic age have been reported from the Angao Formation (Pantoja, 1959), although the major exposures of volcaniclastic rhythmic sedimentary rocks are Berriasian to upper Aptian (Guerrero-Suastegui, 1997). Their depositional environment changes upsection from deep to shallow marine. The volcaniclastic rocks of the San Lucas Formation change toward the top to thick limestone zones with fossil ammonites, orbitolinids, and rudists of late Aptian–early Albian age (El Cajon and Mal Paso Formations; Guerrero-Suastegui, 1997; Pantoja-Alor and Caballero, 2003). This sequence alternates with or changes laterally into marine and terrestrial volcanic sandstone and conglomerate (Comburindio Formation; Guerrero-Suastegui, 1997; Pantoja-Alor and Caballero, 2003). The conglomerate is covered by massive, thick packets of limestone (Huetamo Formation) that contain fossils of late Albian–Cenomanian age. This unit is found only in the central parts of the Huetamo region (Pantoja, 1990). The arc succession of the Zihuatanejo Terrane in the Huetamo area was deformed prior to the deposition of a thick, subaerial red-bed succession that is interbedded with volcanic rocks (Cutzamala Formation of Campa and Ramirez, 1979) and is related to a continental arc of Santonian–Maastrichtian age (AltamiraAreyán, 2002; Benammi et al., 2005). The oldest Cretaceous rocks of the Zihuatanejo Terrane in the Zihuatanejo-Colima region of coastal Mexico that have been penetrated by drilling are Berriasian–Hauterivian in age (Alberca Formation; Cuevas, 1981). The lower member of the Alberca Formation is made up of interbedded black shale, sandstone, and limestone, and some tuff. The upper member is composed mostly of andesitic-basaltic lava flows interbedded with limestone and shale. The Alberca Formation changes transitionally upward to andesitic and basaltic lava flows, with some rhyolitic flows, interbedded with pyroclastic (intermediate tuffs and ignimbrites) and epiclastic deposits. It contains limestone packets interbedded with subaerial conglomerate and sandstone, red siltstone, and some evaporites, and continues into limestone with scarce basaltic pillow lavas at the top (Tecalitlán, Tepalcatepec, and Madrid Formations). The age range of these units, based on their fossil content, is Barremian to Cenomanian (Grajales and López, 1984). Along the west coast between the cities of Colima and Zihuatanejo are exposures of an important succession of red beds, alternating with lesser amounts of limestone in comparison with other areas of the Guerrero Composite Terrane. The assemblage is made up of rhyolitic lavas (lava flows, breccias,

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and ignimbrites) and minor andesitic and dacitic lavas (Tecalitlán Formation, Titzupa-La Unión assemblage, Playitas Formation, etc.; Ferrusquía et al., 1978; Grajales and López, 1984; Pantoja and Estrada, 1986; Centeno-García et al., 2003). These units are interbedded with epiclastic deposits such as tuff, volcanic shale, and sandstone, and some conglomerate. The assemblage also contains thin beds of limestone containing orbitolinids, gastropods, and some pelecypods of late Albian–Cenomanian age (Ferrusquía et al., 1978; Grajales and López, 1984). Raindrop marks, desiccation polygons, and dinosaur footprints can be found in this succession (Ferrusquía et al., 1978). The lower parts of the Cretaceous succession are missing in the Arteaga region, where nonmarine and shallow-marine volcanic and volcaniclastic rocks of Aptian–Albian age rest unconformably on the Arteaga Complex. Overall, Cretaceous volcanic rocks of the southern Zihuatanejo Terrane show geochemical and isotopic signatures that suggest a transitional composition between oceanic island arcs and active continental margins (Centeno-García, 1994; Freydier et al., 1997; Mendoza and Suastegui, 2000). The high potassium content, abundance of felsic lavas, and trace element abundances of these volcanic rocks are similar to those observed in IAs where the crust is thick (>~20 km), allowing magmatic differentiation (Centeno-García, 1994). Rocks of the southern Zihuatanejo Terrane are distinctive from the rest of the terranes because they were deposited in shallow-marine and fluvial environments, contain fossil vertebrates, and show calc-alkaline volcanism more evolved than that of the Teloloapan and Arcelia Terranes. Sedimentary rocks interbedded with the volcanic flows contain clasts of their basement rocks, made up of sandstone, quartz, and mylonitic granite. Thus its stratigraphy is similar to that of arcs constructed on intermediate crust with a previous history of accretions. The presence of fossil vertebrates suggests proximity to the continent. Northern Guerrero Terrane Following a section from east to west in the northern part of the Guerrero Terrane, the main stratigraphic characteristic is an absence of rocks similar to those of the Mixteca or Teloloapan Terrane. Instead, deep-marine volcanic-sedimentary successions of the Guanajuato Terrane were thrust directly over limestone of the calcareous platform of Oaxaquia. Contact relationships between the Guanajuato Terrane and the northern Zihuatanejo Terrane are unconstrained because the contact is covered by younger units. It is inferred that the Guanajuato Terrane is overthrust by the Zihuatanejo Terrane on the basis of regional vergence of the structures. Contact relationships between the Tahue and Zihuatanejo Terranes are unknown because the contact is covered by overlapping Cenozoic assemblages. Guanajuato Terrane The succession at the Guanajuato Terrane has been described as a complete stratigraphic column of an accreted volcanic arc, as its assemblages vary from the roots of the arc (gabbros and

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diabases, and dike swarms) to pillow basalts, interbedded with thin-bedded siltstone, shale, chert, and fine-grained volcanic sandstone (Figs. 6 and 7; Ortiz-Hernandez et al., 1991; OrtizHernandez, 1992). However, all the different stratigraphic levels are in the form of tectonic slivers (Fig. 7), with the deepest mafic levels (gabbro, tonalite, serpentinite, wehrlite, dike swarms) thrust over the upper stratigraphic levels (pillow basalt and volcanic turbidites). The uppermost thrust sheet is made up of ultramafic-mafic rocks of the Cerro Pelón tonalite and the Tuna Mansa diorite. These ultramafic rocks are thrust over a succession incorporating a diabasic feeder dike swarm, basaltic pillow lavas (La Luz basalts), rhyolitic tuffs (Cubilete Tuff), and a deep-marine volcaniclastic succession made up of sandstone and shale turbidites, chert, and black detrital limestone (Esperanza Formation; Quintero-Legorreta, 1992; Ortiz-Hernandez et al., 1992; OrtizHernandez et al., 2003). Basalts of this assemblage show geochemical signatures similar to present primitive volcanic island arcs (Ortiz-Hernandez, 1992). The third and lowermost structural level (Fig. 7) is composed of a thick turbidite succession of volcanic graywackes, quartzites, micritic limestone, radiolarian chert, black shale, and rare conglomerate resting on basaltic pillow lavas (Arperos Formation; Ortiz-Hernandez et al., 1992; Lapierre et al., 1992; Quintero-Legorreta, 1992; Monod et al., 1990; Martínez-Reyes, 1992; Ortiz-Hernandez et al., 2003). Pillow basalts at the base of the Arperos Formation are more alkaline than the La Luz basalts and show OI geochemical signatures (Ortiz-Hernandez et al., 2003). The Arperos Formation is unconformably overlain by the Aptian–Albian La Perlita Limestone (Ortiz-Hernandez et al., 2003). It is difficult to reconstruct the role of the Guanajuato Terrane in the tectonic evolution of western Mexico because of the lack of enough geochronological data. The only U/Pb zircon age reported from the area comes from the El Gordo volcanogenic massive sulfide ore deposit (Hall and Mortensen, 2003), which is considered part of the lowermost succession by Hall and Mortensen, (2003), but it is at the stratigraphic level of the second thrust sheet (Cubilete tuff?) in the stratigraphy proposed by Ortiz-Hernandez et al. (1992). The age of a rhyolite from El Gordo volcanogenic massive sulfide ore deposit reported by Hall and Mortensen (2003) yielded a 146.1 Ma U/Pb age. There are also reports of badly preserved radiolarians from the Arperos Formation that are not in good enough condition to be age indicators (possibly Valanginian–Turonian in age), but a report of nannofossils suggests a Tithonian–Hauterivian age (Ortiz-Hernandez et al., 2003). Other ages reported from the Guanajuato area are from K/Ar analyses and seem to have been reset by later thermal events (Ortiz-Hernandez et al., 1992, 2003). The sedimentary rocks of the La Luz and Arperos Formations seem to be distal volcanic turbidite deposits, but the abundance of limestone associated with the pillow lavas suggests that deposition occurred above the carbonate compensation depth (Ortiz-Hernandez et al., 2003). Aptian–Albian limestone of the La Perlita Formation rests

unconformably on the Arperos Formation and suggests that sedimentation and at least one phase of deformation occurred prior to the Aptian–Albian (Ortiz-Hernandez et al., 2003). Whether or not this deformation is related to the accretion of the Guanajuato Terrane to the continental margin has not been determined. At present the Guanajuato Terrane is thrust over the calcareous platform of Oaxaquia in the San Miguel de Allende area (OrtizHernández et al., 2002). Rocks of the Guanajuato Terrane have been correlated with the Arcelia Terrane, and both were interpreted as having formed part of an oceanic arc independent of the Zihuatanejo and other arc terranes (Ortiz-Hernandez et al., 1992). Also, these rocks are considered relicts of an oceanic basin consumed by subduction related to the arc of the Zihuatanejo Terrane (Lapierre et al., 1992; Tardy et al., 1994). An alternative preliminary interpretation, based on provenance and stratigraphy, is that the Guanajuato Terrane may have been the backarc basin of the Zihuatanejo Terrane (Centeno-García et al., 2003). Zihuatanejo Terrane The Upper Jurassic–Cretaceous stratigraphy of the Zacatecas area in the northern Zihuatanejo Terrane is very different than the stratigraphy of the neighboring Central Terrane and Oaxaquia (Figs. 5–7). Whereas the strata in the northern Zihuatanejo Terrane are mostly composed of volcanic and volcaniclastic rocks, northern Oaxaquia and the Central Terrane were covered by a thick, shallow-marine calcareous platform during the Late Jurassic–Cretaceous (Centeno-García and Silva-Romo, 1997). This suggests that the Zihuatanejo Terrane was probably undergoing dislocation from the continental margin during that time. The arc stratigraphy of the Zacatecas area is formed by the La Borda, Chilitos, and El Saucito Formations (de Cserna, 1976; Yta et al., 1990; Olvera-Carranza et al., 2001; Olvera-Carranza, 2002). These three formations are made up of pillow basalts and volcanic breccias, interbedded with thin-bedded siltstone, shale, chert, and volcanic sandstone and conglomerate, with scarce felsic tuff beds and detrital limestone (Centeno-García and SilvaRomo, 1997; Olvera-Carranza, 2002). The chert layers contain radiolarian fossils of Neocomian(?) to Aptian–Albian(?) age (Yta et al., 1990; Olvera-Carranza, 2002). However, older U/Pb ages have been reported (150–148 Ma) from the base of the succession (Danielson, 2000; Mortensen et al., 2003). Lapierre et al. (1992) and Freydier et al. (1995) characterized this magmatism as primitive IA and OI basalts. Sedimentary structures and fossil content suggest that the La Borda, El Saucito, and Chilitos Formations were deposited as distal turbidites and grain flows in a volcaniclastic submarine apron (Centeno-García et al., 2003). These Jurassic–Cretaceous arc successions contain important volcanogenic, massive sulfide ore deposits (Yta et al., 1990; Danielson, 2000; Mortensen et al., 2003). Tahue Terane Cretaceous successions of the Tahue Terrane are exposed mostly in the Sinaloa de Leyva–Porohui region (Fig. 4). They

Guerrero Composite Terrane of western Mexico were formed by submarine pillow lavas, volcaniclastic rocks, shale, and limestone. They contain Albian ammonites (OrtegaGutiérrez et al., 1979; Freydier et al., 1995; Gastil et al., 1999), but Ar/Ar ages from the lavas are younger (86 Ma; Gastil et al., 1999), suggesting resetting. The basaltic lavas show MORB and OIB geochemical affinities, but the volcaniclastic rocks are more evolved and show IA geochemical signatures (Freydier et al., 1995; Gastil et al., 1999). Although this volcanic succession has been interpreted as the northern continuation of the Arperos Formation of the Guanajuato Terrane, and part of a major oceanic basin that originally lay between the Guerrero arc and the continent (Tardy et al., 1994; Lapierre, et. al., 1992; Dickinson and Lawton, 2001), the stratigraphy does not support such a tectonic scenario because (1) the Cretaceous volcanic rocks rest unconformably on a Paleozoic basement, (2) the stratigraphy and facies associations are not indicative of deep-pelagic sedimentation and oceanic-ridge volcanism, and (3) the Guanajuato successions apparently are older than the arc assemblages of the Tahue and other parts of the Guerrero Composite Terrane. SUMMARY • The stratigraphy of the Guerrero Composite Terrane of western Mexico is characterized by a series of terranes whose basements were formed by Paleozoic to Triassic fragments of oceanic arcs, continental slope sediments, and ocean floor assemblages that were accreted to the continent and consecutively rifted and translated. • Metamorphosed Ordovician volcanic and marine sedimentary rocks and a thick succession of deep-marine turbidites of the NW Guerrero Composite Terrane (Tahue Terrane) make up the record of a middle Paleozoic collision and development of a Carboniferous to Permian passive margin. These rocks might be equivalent to the early Paleozoic Antler Arc and eugeoclinal sedimentation in the SW Cordillera of North America. • The continental margin during the early Mesozoic was located in the middle of Mexico, approximately along the boundary between Oaxaquia and the Central–Guerrero Composite Terranes. This continental margin was active during the Permian–Carboniferous, when a continental arc developed in Oaxaquia. • Permian–Carboniferous arc-related magmatism ceased, and a passive or rifted margin developed along the western continental margin of Mexico, extending throughout the Triassic. This development is suggested by the thick submarine siliciclastic turbidite succession that accumulated on the western paleo-continental shelf–slope region (Potosi Submarine Fan). The siliciclastic fan turbidites are mostly continent-derived, quartz-rich sandstone, siltstone, and shale, containing fossils of Carnian–Norian age. • The Potosi Fan is interpreted as passive-margin deposits, as there is no evidence of contemporaneous magmatism either in the stratigraphy or in the provenance.

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• The siliciclastic rocks of the Potosi Fan extended to the west in a marginal oceanic basin (Arteaga Basin) that at present forms the basement of the Zihuatanejo Terrane of the Guerrero Composite Terrane. • The first compressional event that deformed the Triassic rocks originated tight folding, shearing, and axial cleavage in the La Ballena Formation, and block-in-matrix texture in the Taray and Zacatecas Formations and the Arteaga Complex. This deformation was related to subduction along the early Mesozoic continental margin. It may have started sometime between the Late Triassic and Early Jurassic, accreting the turbidites of the Potosi Submarine Fan, with slivers of the oceanic crust, to the continent. • Whether the subducting slab was dipping toward the west or the east is not well constrained, but the accretionary prism apparently was very wide. Evidence of contemporaneous oceanic arc magmatism is found in the Vizcaíno Peninsula, where a volcanic sequence of primitive arc affinity is exposed. It is possible that the rocks in the Vizcaino Peninsula represent a displaced fragment of an oceanic arc that accreted to the Arteaga Complex of the Guerrero Composite Terrane, but this model needs more evidence. • Arc-related volcanic and sedimentary rocks unconformably overlie the deformed Triassic rocks of Oaxaquia and the Central and Guerrero Composite Terranes. They are characterized by continental rhyolitic to andesitic lava flows, interbedded with fluvial and alluvial deposits. The succession shows minor angular unconformities, probably related to tilting. These rocks have been interpreted as the southern continuation of the Jurassic continental arc that developed along the southwestern margin of the United States. Magmatism was active from ca. 163 to 155 Ma (Callovian–Oxfordian), although older volcanic rocks have been reported for eastern Mexico (189 Ma). The Jurassic arc shows more evolved geochemical signatures than the subsequent volcanic events. • During and after the continental arc activity (Late Jurassic–Early Cretaceous), large amounts of extension and lateral translations probably occurred, as suggested by the changes in the stratigraphy. It has been proposed that major strike-slip faults were probably active during the arc activity (Mojave-Sonora Megashear). Arc magmatism ceased in central Mexico, and considerable subsidence and extension is evidenced by the fast deepening of the calcareous platform that developed over the arc rocks. • Major stratigraphic, geochemical, and isotopic differences are evident in the different Cretaceous stratigraphic assemblages among the Guerrero terranes. They are, from east to west: Andesitic-basaltic submarine lava flows and tuff (IA geochemical signatures), interbedded with limestone and shallow-marine volcaniclastics (Teloloapan Terrane) that were thrust over contemporaneous but more evolved arc successions and the calcareous platform of southern continental Mexico (Mixteca Terrane). Ophiolite successions,

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Centeno-García et al. with deep-marine volcanic and sedimentary rocks with MORB, OIB, and IA signatures (Guanajuato and Arcelia Terranes), are placed between the continent and the more evolved arc in the north (Zihuatanejo Terrane) and between the two shallow-marine arcs (Teloloapan and Zihuatanejo Terranes) in the south. • These major geological differences suggest that intra-arc rifting was considerable and originated a series of marginal arc-backarc systems in western Mexico, with complex paleogeography. Two possible scenarios can be proposed for the Cretaceous paleogeography of western Mexico: (1) that there was one single rifting arc, with westward migration of the magmatism and development of deep-marine intraarc and backarc basins (Guanajuato and Arcelia Terranes); and (2) that rifting during the end of the Jurassic was large enough to allow the formation of multiple marginal island arcs, separated by oceanic backarc basins. • The proposed timing of the final amalgamation of the Guerrero terranes to the margin of older terranes that form the eastern part of Mexico is Turonian to Maastrichtian, as suggested by the age span of foreland basins associated with the deformation of the arc. Overlapping the previously deformed Arcelia and Zihuatanejo Terranes, a new arc developed along the coast by Santonian time.

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l’évolution géodynamique des cordillères mexicaines [Ph.D. thesis]: Université Joseph Fourier-Grenoble I, France, 462 p. Talavera-Mendoza, M.O., 2000, Mélanges in southern México: Geochemistry and metamorphism of Las Ollas complex (Guerrero Terrane): Canadian Journal of Earth Sciences, v. 37, p. 1309–1320. Talavera-Mendoza, O., Ramirez-Espinosa, J., and Guerrero-Suástegui, M., 1995, Petrology and geochemistry of the Teloloapan subterrane, a Lower Cretaceous evolved intra-oceanic island-arc: Geofísica Internacional, v. 34, p. 3–22. Talavera-Mendoza, O., Ruiz, J., Gehrels, G., Valencia, V., and Centeno-García, E., 2007, Detrital zircon U/Pb geochronology of southern Guerrero and western Mixteca arc successions (southern Mexico): new insights for the tectonic evolution of southwestern North America during Late Mesozoic: Geolgical Society of America Bulletin, v. 119, no. 9/10, p. 1052 1065, doi: 10.1130/B26016.1. Tardy, M., Lapierre, H., Freydier, C., Coulon, C., Gill, J.B., Mercier de Lepinay, B., Beck, C., Martinez, J., Talavera, M., Ortiz, E., Stein, G., Bourdier, J.L., and Yta, M., 1994, The Guerrero suspect terrane (western Mexico) and coeval arc terranes (the Greater Antilles and the Western Cordillera of Colombia): A late Mesozoic intra-oceanic arc accreted to cratonal America during the Cretaceous: Tectonophysics, v. 234, p. 49–73. Tristán-González, M., and Torres-Hernández, J.R., 1994, Geología del área de Charcas, Estado de San Luis Potosí, 1994: Universidad Nacional Autónoma de México, Instituto de Geología: Revista Mexicana de Ciencias Geológicas, v. 11, p. 117–138. Urrutia-Fucugauchi, J., and Uribe-Cifuentes, R.M., 1999, Lower crustal xenoliths from the Valle de Santiago Maar Field, Michoacan-Guanajuato Volcanic Field, Central Mexico: International Geology Review, v. 41, p. 1067–1081. Valencia-Moreno, M., Ruíz, J., and Roldán-Quintana, J., 1999, Geochemistry of Laramide granitic rocks across the southern margin of the Paleozoic North American continent, Central Sonora, Mexico: International Geology Review, v. 41, p. 845–857. Valencia-Moreno, M., Ruíz, J., Barton, M.D., Patchett, P.J., Zürcher, L., Hodkinson, D., and Roldán-Quintana, J., 2001, A chemical and isotopic study of the Laramide granitic belt of northwestern Mexico: Identification of the southern edge of the North American Precambrian basement: Geological Society of America Bulletin, v. 113, p. 1409–1422, doi: 10.1130/0016– 7606(2001)113<1409:ACAISO>2.0.CO;2. Vidal-Serratos, R., 1991, Estratigrafía y tectónica de la región de Zihuatanejo, Estado de Guerrero, Sierra Madre del Sur: Convención sobre la evolución Geológica Mexicana, 1er Congreso Mexicano de Mineralogía, Pachuca, Memoir, p. 231–233. Yañez, P., Ruiz, J., Patchett, P.J., Ortega-Gutiérrez, F., and Gehrels, G., 1991, Isotopic studies of the Acatlan Complex, southern Mexico: Implications for Paleozoic North American tectonics: Geological Society of America Bulletin, v. 103, p. 817–828. Yta, M., Lapierre, H., Monod, O., and Wever, P., 1990, Magmatic and structural characteristics of the Lower Cretaceous arc-volcano-sedimentary sequence of Saucito-Zacatecas-Fresnillo (central Mexico), geodynamic implications: Munich, Geowisenschaftliches Lateinamerika, Kolloquium, Memoir, p. 21.11–23.11. Zaldivar, R.J., and Garduño, M.V.H., 1984, Estudio estratigráfico y estructural de las rocas del Paleozoico Superior de Santa Maria del Oro, Durango, y sus implicaciones tectónicas [abstract]: Reunión Anual, Sociedad Geológica de México, p. 37–38.

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The Geological Society of America Special Paper 436 2008

Tectonic architecture of an arc-arc collision zone, Newfoundland Appalachians Alexandre Zagorevski* Cees R. van Staal* Vicki McNicoll Neil Rogers Geological Survey of Canada, 601 Booth Street, Ottawa, Ontario, K1A 0E8, Canada Pablo Valverde-Vaquero Instituto Geológico y Minero de España (IGME), La Calera 1, Tres Cantos (Madrid), 28760, Spain

ABSTRACT The Appalachian-Caledonian orogen records a complex history of the closure of the Cambrian−Ordovician Iapetus Ocean. The Dunnage Zone of Newfoundland preserves evidence of an Ordovician arc-arc collision between the Red Indian Lake Arc, which forms part of the peri-Laurentian Annieopsquotch accretionary tract (ca. 480– 460 Ma), and the peri-Gondwanan Victoria Arc (ca. 473–453 Ma). Despite the similarity in age, the coeval arc systems can be differentiated on the basis of the contrasts that are apparent across the suture zone, the Red Indian Line. These contrasts include structural and tectonic history, stratigraphy, basement characteristics, radiogenic lead in mineral deposits, and fauna. The arc-arc collision is considered in terms of modern analogues (Molucca and Solomon Seas) in the southwest Pacific, and the timing is constrained by stratigraphic relations in the two arc systems. The Victoria Arc occupied a lower-plate setting during the collision and underwent subsidence during the collision, similar to the Australian active margin and Halmahera arcs in the southwest Pacific. The timing of the subsidence is constrained by three new ages of volcanic rocks in the Victoria Arc (457 ± 2; 456.8 ± 3.1; 457 ± 3.6 Ma) that immediately predate or are coeval with deposition of the Caradoc black shale. In contrast the Red Indian Lake Arc contains a sub-Silurian unconformity and a distinct lack of Caradoc black shale, suggesting uplift during the collision. The emergent peri-Laurentian terranes provided detritus into the newly created basin above the Victoria Arc. The evidence of this basin is preserved in the Badger Group, which stratigraphically overlies the periGondwanan Victoria Arc but incorporated peri-Laurentian detritus. Thus the Badger Group forms a successor basin(s) over the Red Indian Line. Following the collision, subduction stepped back into an outboard basin, the Exploits-Tetagouche backarc, closing the Iapetus Ocean along the Dog Bay Line in the Silurian. Correlative tracts in

*Email, Zagorevski: [email protected]; Present address, van Staal: Geological Survey of Canada, 625 Robson Street, Vancouver, British Columbia V6B 5J3, Canada. Zagorevski, A., van Staal, C.R., McNicoll, V., Rogers, N., and Valverde-Vaquero, P., 2008, Tectonic architecture of an arc-arc collision zone, Newfoundland Appalachians, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 309–333, doi: 10.1130/2008.2436(14). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Zagorevski et al. the Northern Appalachians and British Caledonides support the Ordovician arc-arc collision; however, the evidence is less obvious than in Newfoundland. Keywords: arc-arc collision, Dunnage Zone, U/Pb geochronology, Sandbian and Katian black shale, Red Indian Line.

INTRODUCTION The Cambrian–Ordovician closure of the main tract of the Iapetus Ocean generated a diverse set of arc-backarc terranes and microcontinents that were accreted sequentially to the Laurentian margin, leading to significant growth of the eastern margin (present coordinates) of North America and the formation of the Appalachian-Caledonian orogen. The reconstruction of the paleogeography of the Iapetus Ocean requires identification of the various terranes and their tectonic setting and has important implications for a wide variety of topics ranging from the distribution of mineral deposits to the tectonic development of convergent margins and continental growth. Newfoundland lies in a critical position in the Appalachians because it forms the geological link between the Northern Appalachian and the Caledonide segments of this once-continuous orogen. As such the development of ideas in Newfoundland has greatly influenced the evolution of thought on the tectonic development of the Appalachian-Caledonian orogen as a whole. In addition to occupying a key position in the orogen, Newfoundland is also perhaps the best constrained portion of the Appalachians, featuring excellent coastal exposures, detailed mapping, high-density geochemical, isotopic, geochronological, and geophysical data that include two crustal-scale seismic-reflection transects in central Newfoundland (van der Velden et al., 2004). The central mobile belt of Newfoundland comprises a complex tectonic collage of arc-backarc terranes that formed within the Iapetan realm outboard of Laurentia and Gondwana. The juxtaposition of these terranes was piecemeal, involving several arc-continent and arc-arc collisions, with paleogeographic complexity similar to the modern arcs in the southwest Pacific (e.g., van Staal et al., 1998). Similar to the southwest Pacific (Hall, 2002), the accretionary processes in the Iapetus operated on very short time scales, with the accretion of many arc-backarc terranes that occurred within 5–10 m.y. following their formation (Lissenberg et al., 2005a; Zagorevski et al., 2006, 2007a). Consequently, the resolution of these processes requires highly detailed geochronology, which is available in central Newfoundland. This paper reviews the evidence for an Ordovician arc-arc collision recorded in the central mobile belt of the Newfoundland Appalachians. Three new U-Pb zircon geochronological ages constrain the timing of volcanism and sedimentation in the Exploits Subzone and the age of the arc-arc collision. Select structural data for central Newfoundland is presented and integrated into a tectonic model, which also incorporates

the seismic-reflection data of van der Velden et al. (2004). This tectonic model is considered in the context of the tectonics and sedimentation in recent active arc-arc collisions in the southwest Pacific, specifically the Huon-Finisterre (Solomon Sea) and Halmahera-Sangihe (Molucca Sea) arc-continent and arcarc collisions (e.g., Abers and McCaffrey, 1994; Pubellier et al., 1999; Whitmore et al., 1999). Modern Arc-Arc Collisions Prior to reviewing the evidence for the arc-arc collision in central Newfoundland we briefly introduce two modern analogues and review the key evidence that would enable the recognition of arc-arc collisions in ancient orogens. We have chosen the Huon-Finisterre arc-continent collision in the Solomon Sea (Abbott et al., 1994; Abers and McCaffrey, 1994) and the Halmahera-Sangihe arc-arc collision in the Molucca Sea (Lallemand et al., 1998; Pubellier et al., 1999) because they are active and preserve different stages of collision along the strike of the orogen. In addition, we believe that they most closely resemble the relationships in Newfoundland. Solomon Sea The Solomon Sea plate is currently subducting under the Bismarck Arc to the north (New Britain Trench) and under the active Australian continental margin to the south (Trobriand Trough: Fig. 1; e.g., Abers and McCaffrey, 1994). In the western Solomon Sea the Australian-Bismarck-Solomon-plate triple junction forms the site of a modern arc-arc collision where the Bismarck Arc is colliding with the active Australian margin. The collision is diachronous from the northwest to the southeast and has occurred northeast of the triple junction, forming the Huon and Finisterre Ranges on the Huon Peninsula. The deformation is accommodated by southwest-directed thrusting of the Bismarck Arc and its accretionary prism over the Australian active margin along the Ramu-Markham Fault, a continuation of the New Britain Trench on land. An important out-of-sequence thrust, the Gain Thrust, cuts across the Ramu-Markham Fault and emplaces the volcanic rocks of the Bismarck Arc over its accretionary prism (Fig. 2). Continuation of movement on the Gain thrust may in time structurally mask the accretionary prism, leading to the juxtaposition of the Bismarck Arc and active Australian margin. As a result of the collision and loading by the Bismarck Arc, parts of the Australian continental shelf underwent rapid subsidence (2 km over the last 348 m.y.: Galewsky et al., 1996) indicated by the presence of the drowned carbonate platform above

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the Morobe Shelf (Fig. 1; Galewsky and Silver, 1997). The degree of subsidence of the Australian margin decreases to the east of the Australia-Bismarck-Solomon-plate triple junction, where the collision has not yet occurred (Galewsky and Silver, 1997). The emergent orogen to the west of the triple junction supplies abundant sediment into the Solomon Sea. The detritus is dominantly transported as turbidity currents parallel to the trench along the Markham Canyon and is deposited on the subducting Molucca Sea plate and Australian margin (Galewsky and Silver, 1997). Molucca Sea The Molucca Sea is the site of the collision between the facing Sangihe and Halmahera Arcs, which formed as a result of the double subduction of the Molucca Sea plate to the west and east, respectively (Fig. 1; e.g., Hall, 2002). The collision initiated in the northern Molucca Sea during the Pliocene and migrated to the south, where the collision is still in its infancy (Pubellier et al., 1999; Hall, 2000). The Molucca Sea plate has been entirely overridden by the collisional complex derived primarily from the Sangihe forearc. In the southern Molucca Sea, where the collision is the least advanced, the Halmahera forearc is in part overridden by the Sangihe forearc. To the north, the Halmahera forearc is completely overridden by the Sangihe forearc, and the Snellius Plateau, the continuation of the Halmahera Arc to the north, is subducting at present (Lallemand et al., 1998; Pubellier et al., 1999; Hall, 2000). As a result of subduction, the Snellius Plateau underwent rapid subsidence, as evidenced by the presence of a drowned carbonate shelf, whereas parts of the overriding collisional complex have been uplifted above sea level (Talaud Island and Pujada Ridge: Figs. 1, 2). The Molucca Sea plate is entirely overridden by the collisional complex but is still actively subducting in the south Molucca Sea. To the north, evidence from deep earthquakes suggests that the Molucca Sea plate has delaminated from the overriding plate beneath Mindanao (Lallemand et al., 1998). Mindanao is characterized by widespread Pliocene and younger syn- to postcollisional basaltic to dacitic magmatism. The compositions of the magmas range from calc-alkaline to shoshonitic. The diversity of the magmas and the geochemistry has been related to the delamination of the Molucca Sea plate (Sajona et al., 2000). Recognition of Ancient Arc-Arc Collisions The juxtaposition of coeval arc terranes, such as those described for the Molucca and Solomon Seas, presents a difficulty with respect to their identification and interpretation in ancient orogens. In the case of both collision zones described above, the facing arc terranes may be differentiated on the basis of the (1) tectonic history and basement characteristics, (2) geochemistry, and (3) stratigraphic record. (1) Prior to the collision the Halmahera Arc-backarc underwent a period of shortening that resulted in thrusting of the backarc rocks onto the arc and a gap in volcanism, whereas the Sangihe Arc was not shortened during this time. The Australian margin in the Solomon Sea area was not active until recently, whereas the basement to the Bismarck Arc

should have undergone several tectono-thermal events relating to protracted arc volcanism, based on the reconstructions of Hall (2002). (2) The geochemistry of the volcanic rocks of the Halmahera and Sangihe Arcs shows distinct characteristics (Macpherson et al., 2003), which may be utilized in their separation. (3) The pre- and syncollisional stratigraphy may prove extremely useful in differentiation of the arc sequences. Both the Australian margin and the Halmahera Arc–Snellius Plateau were drowned prior to and during the collision, while the overriding arcs and forearc basins were undergoing uplift and erosion. Emergence of the overriding arc and collisional orogen supplies syntectonic sediment into the basin created above the subducting arc. The appearance of sediment with provenance consistent with derivation from the overriding arc thus would provide information on the timing of the emergence of the orogen. Finally, the delamination of the subducting oceanic plate may lead to extensive syn- to postcollisional magmatism, marking the end of the collision, such as in Mindanao (Sajona et al., 2000). The recognition of ancient arc-arc collision zones similar to those in the southwest Pacific thus would require an approach that integrates stratigraphy, geochemistry, geochronology, structural geology, and geophysics, combined with a good understanding of basement-cover relationships and affinity of the various terranes. We believe that the quality, quantity, and distribution of data make the Newfoundland Appalachians unique among ancient orogens in that they permit detailed reconstruction of the tectonic architecture of an Ordovician arc-arc collision. GEOLOGICAL SETTING The complexity of the closure of the Iapetus Ocean is reflected in the zonal division of the northern Appalachians, where five zones have been defined: Humber, Dunnage, Gander, Avalon, and Meguma (Fig. 3; Williams, 1995b). The Humber Zone represents the Cambrian–Ordovician Laurentian margin. The Dunnage Zone (the central mobile belt) lies outboard of the Humber Zone and contains allochthonous ensialic and ensimatic arc-backarc complexes that formed within the Iapetus Ocean. Farther outboard, the Gander, Avalon, and Meguma Zones represent microcontinents derived from Gondwana (see van Staal, 2005, for a review). The margin of Laurentia experienced three orogenic episodes from Ordovician to Early Devonian time, namely the Taconic, Salinic, and Acadian. These episodes occurred as a result of the successive arrival of the Dashwoods, Gander, and Avalon microcontinents (van Staal, 2005). The Dunnage Zone is subdivided into the peri-Laurentian Notre Dame and the peri-Gondwanan Exploits Subzones (Fig. 3; Williams et al., 1988). Other subsequent subdivisions (Williams, 1995a) reflect minor regional differences internal to these two subzones that are beyond the scope of this paper. The Notre Dame Subzone comprises predominantly Early to Middle Ordovician Notre Dame Arc plutons that intruded a ribbon microcontinent, referred to as the Dashwoods microcontinent (Waldron and van Staal, 2001), as well as accreted peri-Laurentian terranes. The

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Exploits Subzone contains peri-Gondwanan ensialic and ensimatic arc-backarc complexes. The two subzones are differentiated on the basis of faunal, paleomagnetic, isotopic, and structural contrasts (Williams et al., 1988). Notre Dame–Exploits Relationships in Newfoundland The identification of the marked geological contrasts between the Notre Dame and Exploits Subzones has led to recognition of the Red Indian Line (Fig. 3; Williams et al., 1988), which is now accepted as the main suture zone in the northern Appalachians, along which several thousand kilometers of Iapetus oceanic lithosphere has been consumed (e.g., van Staal, 2005). Following the formation of the Red Indian Line the convergence continued by means of subduction in Iapetan marginal

basins and seaways underlain by oceanic crust (e.g., ExploitsTetagouche backarc basin: van Staal et al., 1998). In Notre Dame Bay the Red Indian Line is marked by late brittle faults and locally by Ordovician mélange (McConnell et al., 2002; Williams et al., 1988), although the exact nature and location of this boundary is debated (e.g., Arnott et al., 1985; Wasowski et al., 1986; van der Voo et al., 1991). In central Newfoundland the surface trace of the Red Indian Line is marked by a mélangephyllonite belt (Fig. 3; Rogers and van Staal, 2002; Zagorevski et al., 2006, 2007b), which coincides with the surface projection of a major crustal-scale reflector, interpreted to be a fault, in seismic-reflection profiles (van der Velden et al., 2004). The surface trace of the Red Indian Line is cut out in central Newfoundland by the Devonian Victoria Lake shear zone (Fig. 3; Valverde-Vaquero and van Staal, 2002), which is a continuation

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of the Noel-Paul’s Line (Williams et al., 1988). At depth the Red Indian Line is truncated by a seismic reflector interpreted to be a Devonian wedging-related fault (van der Velden et al., 2004). Notre Dame Subzone The Notre Dame Subzone in Newfoundland comprises several composite terranes formed outboard of the Laurentian margin. The western margin and central part of the Notre Dame Subzone is dominated by the Dashwoods microcontinent (Waldron and van Staal, 2001) and its Notre Dame Arc suprastructure. The eastern margin of the Notre Dame Subzone contains a tectonic collage of continental and intraoceanic supra-subduction zone complexes collectively referred to as the Annieopsquotch accretionary tract (Figs. 4–6; van Staal et al., 1998; Lissenberg et al., 2005b; Zagorevski et al., 2006). The Annieopsquotch accretionary tract formed above a westdipping subduction zone and was facing the main Iapetan basin. As such, it contains the most outboard arc-backarc complexes of the Notre Dame Subzone and provides constraints on the development of the Notre Dame Subzone and its juxtaposition with the Exploits Subzone.

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Subduction in the Annieopsquotch accretionary tract initiated at ca. 480 Ma, leading to the formation of the Annieopsquotch ophiolite belt (Dunning and Krogh, 1985; Dunning, 1987; Lissenberg et al., 2005b). This may have occurred in a reentrant in the Dashwoods margin (Waldron and van Staal, 2001; Zagorevski et al., 2006). Subsequently, the subduction developed under the Dashwoods margin, forming a continental arc and backarc by ca. 473 Ma (Buchans Group and Lloyds River ophiolite complex: Dunning et al., 1987; Swinden et al., 1997; Zagorevski et al., 2006). This arc rifted, forming an oceanic basin (Pickett, 1987) and a younger arc, represented by the Red Indian Lake Group in central Newfoundland (Zagorevski et al., 2006). Red Indian Lake Group The Red Indian Lake Group comprises mafic, andesitic, and felsic volcanic rocks, epiclastic sedimentary rocks, red shale, and ferruginous chert. Tholeiitic mafic volcanic rocks are predominant at the stratigraphic base, while calc-alkaline mafic, andesitic, and felsic volcanic rocks become more common at the stratigraphic top, suggesting a maturing arc sequence (Zagorevski et al., 2006). Negative εNd values of felsic volcanic rocks (−6.4 to −7.7) and

57°

L

55°

RI 12

11 10

13

49°

G

R

U

B

49°

5 6 4

8

9

7 1

3

48° 48°

2 57°

Annieopsquotch accretionary tract

Badger Group

55°

Victoria-Exploits forearc, arc and backarc volcanic and sedimentary rocks

Penobscot Arc

Figure 4. Distribution of Annieopsquotch accretionary tract, Victoria Arc, Penobscot basement, and Badger Group in Newfoundland (after Colman-Sadd et al., 1990). 1—Dashwoods Subzone; 2—Windsor Point Group; 3—Annieopsquotch ophiolite belt; 4—Lloyds River ophiolite complex; 5—Buchans Group; 6—Red Indian Lake Group; 7—Pats Pond Group and Wigwam Brook Group; 8—Sutherlands Pond group; 9—Tally Pond Group and Noel Paul’s Brook group; 10—Exploits Group; 11—Wild Bight Group; 12—Summerford Group; 13—Dunnage mélange; GRUB—Gander River Ultrabasic Belt; RIL—Red Indian Line.

430 440 Silurian Silurian Llandovery Wenlock

Notre Dame Subzone

Red Indian Line

Exploits Subzone Botwood Group

Springdale Group and equivalents

Badger Group

VICTORIA ARC

Sutherlands Pond group (8)

Hgs

Exploits Group (10)

Tally Pond Group (9) Victoria Lake Supergroup Penobscot Black shale arc

Ophiolite

Peri-Gondwanan volcanic rocks

Peri-Laurentian volcanic rocks

Limestone

Sedimentary rocks

U/Pb zircon age

Hiatus

Map Location

Red Indian Line Lithoprobe seismic reflection profiles

Unclassified contact Thrust fault

Wild Bight Group (11)

Penobscot orogeny

Pats Pond Group (7)

Annieopsquotch accretionary tract

Summerford Group (accreted seamount; 12)

Dunnage melange (13)

Red Indian Lake Group (6)

Buchans Group (5)

Lloyds River Ophiolite Complex (4)

Annieopsquotch ophiolite belt (3)

*

NPBg

Figure 5. Relationships between selected tectono-stratigraphic units in Newfoundland. Subdivision of the Ordovician Period into epochs follows McKerrow and van Staal (2000). Numbers in parentheses refer to the locations shown in Figure 4. Hgs—Harpoon gabbro suite; NPBg—Noel Paul’s Brook group; WwBG—Wigwam Brook Group. See text for discussion.

PENOBSCOT ARC

Widsor Point Group (2)

WBG WwBG

Dashwoods Subzone(1)

Age (Ma) 500 490 510 480 470 460 450 Middle Upper Middle Upper Lower Cambrian Cambrian Ordovician Ordovician Ordovician Tremadoc Arenig Llanvirn Caradoc Ash.

*

Plutonic rocks

Buchans

Red Indian Lake

48o45’

56o30’

Notre Dame Subzone Annieopsquotch accretionary tract

o

57 15'

S1 n=162

Exploits Subzone

Victoria Lake Supergroup o

Victoria arc

48 30’ o

57 00'

Penobscot arc e

ak

sL

d loy

10 km

S1 n=289

L 57o45'

ria Lake

to Vic

S2 n=190

o

48 15' 57o30’

S1 n=284

Figure 6. Simplified geology of central Newfoundland (after Zagorevski et al., 2006). Stereonets are lower hemisphere, equal area, contoured using the Kamb method, with the number of points indicated. Stereonets of S1 are for the indicated portions of the mapped area, south and southwest of Red Indian Lake. The S2 stereonet is for the entire mapped area south and southwest of Red Indian Lake.

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abundant Mesoproterozoic zircon inheritance suggest that this arc formed on continental crust. The range of U/Pb zircon ages of the volcanic rocks of the Red Indian Lake Group (462 +2/–9, 463 ± 3, 464.8 ± 3.5, 465 ± 2 Ma: Zagorevski et al., 2006) indicates that arc volcanism is predominantly Llanvirn (i.e., 459–465 Ma: McKerrow and van Staal, 2000). The Red Indian Lake Group is the most outboard and the youngest sequence identified within the Annieopsquotch accretionary tract. Correlatives of the Red Indian Lake Group likely occur throughout the Annieopsquotch accretionary tract and may include the undated Crescent Lake Formation (Fig. 3; Bostock, 1988) and parts of the Cottrell’s Cove Group (Arenig: in O’Brien, 2003; i.e., 465–480 Ma: McKerrow and van Staal, 2000) in the Notre Dame Bay region. The Red Indian Lake Group and its correlatives are henceforth referred to as the Red Indian Lake Arc, which comprises the youngest arc sequence in the Annieopsquotch accretionary tract. Notre Dame Subzone Cover The Notre Dame Subzone is unconformably overlain by the volcaniclastic and sedimentary rocks of the spatially restricted Windsor Point Group in southwestern Newfoundland (Fig. 4; 453 +5/–4 Ma: Dube et al., 1996); however, the most widespread unconformity is sub-Silurian. The Notre Dame Subzone is regionally overlain unconformably by the rocks of the Silurian Springdale Group or its equivalents (Chandler et al., 1987; Dunning et al., 1990). In central Newfoundland these rocks comprise felsic to andesitic volcanic rocks, breccia, polymictic conglomerate, and red sandstone (e.g., van Staal et al., 2005a, b). The provenance of the polymictic conglomerate suggests derivation predominantly from the Notre Dame Subzone. Some clasts bear a predepositional foliation, indicating deformation and exhumation of the Ordovician tectonites by Silurian time. Exploits Subzone To the east of the Red Indian Line the Exploits Subzone in Newfoundland can be subdivided into two distinct arc-related sequences (Fig. 4; Jenner and Swinden, 1993; O’Brien et al., 1997; MacLachlan and Dunning, 1998a, b; van Staal et al., 1998; Rogers et al., 2006; Zagorevski et al., 2007a). The older sequence, the Penobscot arc and backarc, contains ca. 513– 485 Ma volcaniclastic and sedimentary rocks that formed above an east-dipping subduction zone along the margin of Ganderia (Rogers et al., 2006; Zagorevski et al., 2007a), a postulated periGondwanan microcontinent (van Staal et al., 1996). A gap in magmatism in the Exploits Subzone (485–478 Ma) is correlated with the collision of the Penobscot arc with Ganderia at ca. 480– 485 Ma and obduction of the Penobscot backarc ophiolites (Jenner and Swinden, 1993) onto the Newfoundland Gander Zone (Fig. 3; e.g., Colman-Sadd et al., 1992; Tucker et al., 1994; van Staal et al., 1998). The Penobscot arc is stratigraphically overlain by ca. 473–453 Ma volcano-sedimentary rocks (Fig. 4; e.g., MacLachlan and Dunning, 1998b; O’Brien et al., 1997; Evans and Kean, 2002; Zagorevski et al., 2007a), which are referred to

as the Victoria Arc in central Newfoundland. The Victoria Arc is a correlative of the Popelogan arc of New Brunswick, Canada (Fig. 3: Popelogan inlier; van Staal et al., 1998). Victoria Arc The tectonic setting of the Victoria Arc, and its correlative Popelogan arc in New Brunswick, are well constrained in the northern Appalachians (e.g., van Staal et al., 1998). The earliest Victoria Arc volcanism in Newfoundland occurred at ca. 473 Ma (Fig. 4; Wild Bight Group: MacLachlan and Dunning, 1998b). Coeval eruption of arc- and backarc-related volcanic rocks (Exploits Group: O’Brien et al., 1997; Wild Bight Group: MacLachlan and Dunning, 1998b) indicates an extensional phase of arc magmatism, which led to the rifting of the Victoria and Popelogan arcs from Ganderia and the opening of the Exploits-Tetagouche backarc basin in Newfoundland and New Brunswick (Fig. 3; O’Brien et al., 1997; MacLachlan and Dunning, 1998b; van Staal et al., 1998; Valverde-Vaquero et al., 2006; Bathurst Supergroup: van Staal et al., 2003). Abundant fossil and some age-dating evidence indicate that the Victoria Arc was active until at least middle Llanvirn time (e.g., O’Brien et al., 1997). For example, Victoria Arc–related basalts of the Sops Head Complex in Notre Dame Bay form peperitic contacts with Llanvirn limestone (McConnell et al., 2002). In central Newfoundland the Victoria Arc is represented by the Sutherlands Pond group (ca. 457–462 Ma: Dunning et al., 1987; Rogers et al., 2005; see following), Noel Paul’s Brook group (ca. 457–465 Ma: van Staal et al., 2005c), and Wigwam Brook Group (453 ± 4 Ma: Zagorevski et al., 2007a). The youngest dated Victoria Arc–related volcanic rocks in the lower Wigwam Brook Group (van Staal et al., 2005c; Zagorevski et al., 2007a) comprise a coarsening-up sequence of felsic tuff and breccia that are intercalated with turbiditic sandstone, conglomerate, black shale, and basalt with arclike chemistry. The upper Wigwam Brook Group comprises turbiditic sandstone and black shale, suggesting cessation of arc magmatism (Zagorevski et al., 2007a). Exploits Backarc Similarly to the Victoria Arc, the Exploits-Tetagouche backarc was magmatically active until the late Llanvirn (e.g., Lawrence Head Formation: O’Brien et al., 1997). The coeval sedimentation in the Exploits backarc was continuous from the early Arenig until the Caradoc (Caradoc: 449–459 Ma: McKerrow and van Staal, 2000) and mostly consisted of deep-marine sediments interrupted by occasional pulses in magmatism, such as the eruption of backarc-basin basalts of the Laurence Head Formation (O’Brien et al., 1997). The end of the Llanvirn was marked by cessation of magmatism, deposition of chert and regionally discontinuous limestone, followed by the black shale of the Lawrence Harbour Formation (Williams, 1995a). In central Newfoundland, backarc volcanism is preserved in the Red Cross Group, which comprises felsic volcanic rocks (466 ± 3 Ma: Valverde-Vaquero et al., 2006), basalt, turbiditic sandstone, black shale, and calcareous shale locally intruded by gabbro (457

Tectonic architecture of an arc-arc collision zone ± 3 Ma: Valverde-Vaquero et al., 2006). Similar to the rest of the Exploits Subzone, the upper parts of the Red Cross Group commonly contain abundant black shale and calcareous shale. Exploits Subzone Cover One of the key characteristics of the Exploits Subzone is the presence of the regionally transgressive Caradoc black shale cover, accompanied by the general cessation of volcanism (Fig. 5; van der Pluijm et al., 1987; Williams, 1995a). In the Notre Dame Bay area (Fig. 3), this time was marked by deposition of the black shale of the Lawrence Harbour and Shoal Arm Formations (e.g., Williams et al., 1995; O’Brien et al., 1997). In central Newfoundland the Wigwam Brook Group (Zagorevski et al., 2007a) and Stanley Waters formation (see following discussion; van Staal et al., 2005c) mirror the relationships in Notre Dame Bay and indicate cessation of volcanism accompanied-followed by deposition of the Caradoc black shale. The cessation of arc magmatism and deposition of black shale are commonly assumed to have occurred at the Llanvirn–Caradoc boundary (e.g., O’Brien, 2003). However, in the Wigwam Brook Group this transition occurred in the middle Caradoc (453 ± 4 Ma: Zagorevski et al., 2007a). Volcanism of similar age may also have been the source of felsic volcanic boudins, which are common in the mélange belts of central Newfoundland (Rogers and van Staal, 2002; Valverde-Vaquero et al., 2006). The cessation of volcanism and the subsidence of the Victoria Arc thus appear to have been somewhat diachronous along strike, and may have started earlier in Notre Dame Bay. Alternatively this may be an artifact of sparse dating in Notre Dame Bay or of local basin architecture such that parts of the basin were effectively isolated from the volcaniclastic input and received only background sedimentation. Regionally the Ashgill to Wenlock (424–449 Ma: McKerrow and van Staal, 2000) Badger Group (Williams et al., 1993) stratigraphically overlies the Caradoc black shale and comprises an upward-coarsening sedimentary sequence of deep-marine turbidites to shallow-marine conglomerate. Structural and sedimentological investigations of the Badger Group indicate deposition in a syntectonic setting (e.g., Kusky et al., 1987; Williams et al., 1995). The detrital provenance of the sedimentary rocks suggests derivation from the Notre Dame Subzone (see Williams et al., 1995), confirmed by the U/Pb detrital zircon data (McNicoll et al., 2001). The syntectonic setting and provenance have led to the interpretation of the Badger Group as a successor basin or basins over the Red Indian Line in the forearc region (arc-trench gap) of the Late Ordovician to Early Silurian subduction complex formed during the closure of the Tetagouche-Exploits backarc basin (Kusky et al., 1987; Pickering, 1987; van Staal et al., 1998; Valverde-Vaquero et al., 2006). STRUCTURAL GEOLOGY Central Newfoundland displays a relatively simple macroscopic structure with internally folded structural panels bounded by curviplanar shear zones (Fig. 6). In the Annieopsquotch

317

accretionary tract, for example, the thickness of the individual tectonic panels varies between 0 and 5 km, indicating partial to total excision of tectono-stratigraphic units along strike. Detailed investigations of the structural geology indicate that the macroscopic simplicity is deceptive and that the rocks have undergone multiple phases of deformation. Seven deformation episodes (D1 through D7) have been recognized in central Newfoundland, based on overprinting relationships as well as available age and stratigraphic constraints (Zagorevski et al., 2007b). The regional geometry of the rocks is primarily the result of D1, D2, and D4 episodes, although only D1 structures formed during the arc-arc collision. Hence, only these are discussed below. D3, 5–7 so far have been shown to be present only on a mesoscopic scale (Zagorevski et al., 2007b) and therefore are not imperative for this chapter. The oldest phase of deformation identified in central Newfoundland (D1) was heterogeneous and led to strain localization into northwest-dipping brittle to ductile shear zones ten to several hundred meters in thickness. These deformation zones are characterized by mylonite, phyllonite, and/or black shale mélange that mark the boundaries between the various tectono-stratigraphic units (Figs. 6, 7). In proximity to the Red Indian Line in central Newfoundland, D1 shear zones accommodated SSE-directed motion (Lissenberg, 2005; Zagorevski et al., 2007b). D1 shear zones are overprinted by upright moderate- to shallow-plunging asymmetric F2 folds and axial planar S2 foliation. S2 is commonly a composite foliation formed by transposition of S1 by tight to isoclinal F2 folds (Figs. 6, 7). The composite S1–2 is the regionally dominant foliation. The enveloping surface of the F2 folds dips northwest, suggesting that the D1 shear zones were originally shallow northwest-dipping structures prior to the D2 structures. In addition, the emplacement of “old over young” rocks suggests that the D1 shear zones represent thrust faults (Fig. 8; Zagorevski et al., 2007b). This interpretation is supported by areas of weak D2 strain, such as Buchans (Fig. 6), where a southeast-directed antiformal thrust stack is developed (Calon and Green, 1987). D2 strain is commonly concentrated along D1 shear zones where the D1 tectonites provided a strong planar anisotropy. This resulted in transposition of S1 and reactivation of D2-steepened D1 shear zones as southeast-directed reverse faults. Similar to D1 shear zones, D2 reverse faults also accommodated SSE-directed translation. The differentiation of D1 and D2 structures in areas of high D2 strain thus can be tenuous in the absence of overprinting relationships. D4 structures include S4 foliation, F4 folds, and D4 shear zones that overprint D1 and D2 structures. Although the D4 structures timewise are not related to the arc-arc collision discussed in this paper, they are locally paramount to the kinematic interpretation of the geometry of the deformed rocks. D4 structures were associated with formation of a northwest-directed thrust and fold belt in central Newfoundland. An excellent example of a D4 structure is the Devonian Victoria Lake shear zone (Fig. 6; Valverde-Vaquero et al., 2006). This shear zone emplaces the metamorphic tectonites of the Gander Zone over the Exploits

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Zagorevski et al.

A

B

C

D

S1

S1

S5 F1 S2 S2

S4 (composite)

Figure 7. Representative relationships between D1 and D2 structures. (A) Horizontal outcrop of polydeformed black shale mélange with a basaltic knocker marking the boundary between the Sutherlands Pond group and the Penobscot arc. (B) Horizontal outcrop of highly deformed Ordovician limestone along the same boundary, with sinistral boudinage of a vein. (C) Vertical photomicrograph of F2-folded D1 metabasic tectonite along the boundary between Lloyds River ophiolite complex and Annieopsquotch ophiolite belt (see Fig. 5). (D) Vertical outcrop demonstrating the relationships between F1 folded quartz veins and upright F2 folds in felsic tuff of the Red Indian Lake Group.

A

B Notre

RIL

N

Dame Arc

AAT

RIL

Victoria Lake Sg. Ganderia

20

Laurentia

B

RIL

30 Moho

C

10

C

Notre Dame Arc

40 50 km

RIL 10 Ganderia

Laurentia Laurentia

20 30 40 50 km

Figure 8. Interpretation of D1-D2 relationships in central Newfoundland. (A) Interpretation of the three-dimensional geometry of the Red Indian Line, based on the asymmetry of F2 folds, which indicate a northwest-dipping enveloping surface. Inset shows a close-up of the central area with folded shear zones (light gray) and folded unconformity between the Pats Pond and Wigwam Brook groups (dark gray). Location of seismicreflection profiles is indicated by panels B and C. Interpretation of the D1-D2 relationship in panel A is consistent with the seismic-reflection profiles shown in panels B and C (simplified from van der Velden et al., 2004). AAT—Annieopsquotch accretionary tract; RIL—Red Indian Line; Sg—supergroup.

Subzone and in part buries the D1 and D2 shear zones, including the Red Indian Line. Ages of D1 and D2 Deformation Zones The ages of the D1 zone have been obtained from the syntectonic intrusions in the westernmost units of the Annieopsquotch accretionary tract. These syntectonic intrusions span from Arenig to early Caradoc (ca. 470 Ma Ar/Ar hornblende;

468 ± 2 Ma U/Pb zircon; ca. 464–458 Ma U/Pb zircon: Lissenberg et al., 2005a). The oldest ages of D1 deformation predate the youngest peri-Laurentian arc volcanism in the Annieopsquotch accretionary tract (Fig. 5; i.e., Red Indian Lake Group: Zagorevski et al., 2006). Therefore, Arenig to early Llanvirn deformation is strictly peri-Laurentian and is not related to the accretion of any peri-Gondwanan terranes. This deformation is probably a result of back thrusting that initiated following the Taconic collision of the Dashwoods microcontinent with the

Tectonic architecture of an arc-arc collision zone Laurentian margin at ca. 470 Ma (Waldron and van Staal, 2001; van Staal et al., 2006). The earliest constraints on the start of deformation that affected both the Notre Dame and Exploits Subzones are provided by the stratigraphic relationships. The dismemberment of the Caradoc black shale unit into mélange and broken formation along the Red Indian Line, without incorporation of Ashgill-age rocks, suggests that the D1 deformation that affected both subzones was under way in the late Caradoc (also observed in the Sops Head–Boones Point Complex: McConnell et al., 2002). Hence, although the D1 deformation was protracted, lasting from at least 470 Ma until at least 450 Ma, only the youngest part of the D1 deformation is related to the accretion of peri-Gondwanan terranes. The upper limit of D1 deformation and the span of D2 deformation are constrained by the Silurian plutonic rocks (435–427 Ma: Dunning et al., 1990; Whalen et al., 2006; Zagorevski et al., 2007b). Some of these plutons stitch major D1 shear zones and contain rare enclaves of D1 tectonites. Despite displaying cross-cutting relationships, the Silurian plutonic rocks are also deformed along the shear zones at amphibolite facies within error of the age of intrusion and as such are interpreted as synD2 plutons (Zagorevski et al., 2007b). Hence, D2 is constrained by these plutons to have started by at least 435 Ma and continued to at least 426 Ma (Zagorevski et al., 2007b). The age of the D2 deformation is also constrained by the unconformable Silurian cover of the Notre Dame Subzone. The age of the red beds

Penny Brook Formation 0.082 Felsic Tuff

480

Pb/238U

206

Penny Brook Formation A sample (RAX05–900, Z8777; UTM zone 21, 594708 E, 5474803 N; NAD 83: Fig. 10) from a thin seam of felsic tuff,

Stanley Waters formation Fine grained felsic tuff

465 Concordia age = 0.074 457.5 ± 2.7 Ma

0.5

0.6

Pb/238U

0.7

0.8

520 500

455

Z1

0.63

Z4A

485 475

460

0.074

Z2

0.54 0.57 0.60 0.082 505 Sutherlands Pond group Felsic volcanic 0.078

480

0.078

Z3

Z1 LI = 457 ± 3.6 Ma (MSWD=0.1) 100 µ m TIMS data

SHRIMP II data

400

Sutherlands Pond group 0.082 Rhyolite

835 Ma

495

485

440

0.062 0.4

505

475

420

206

U-Pb thermal ionization mass spectrometry (TIMS) and sensitive high-resolution ion microprobe (SHRIMP II) analyses were conducted at the Geological Survey of Canada (GSC) in order to constrain the timing of the arc volcanism and black shale deposition in the Exploits Subzone. U-Pb TIMS analytical methods are outlined in Parrish et al. (1987), with treatment of analytical errors after Roddick et al. (1987). SHRIMP II analyses were conducted using analytical procedures described by Stern (1997), with standards and U-Pb calibration methods following Stern and Amelin (2003). U-Pb TIMS and SHRIMP analyses are presented in Tables 1 and 2, respectively, and are plotted in concordia diagrams with errors at the 2σ level (Fig. 9). Further details on the U-Pb analytical techniques are presented in the Appendix. Geochronology sample locations are indicated in Figure 10.

0.078

0.070

465

440

0.070

Concordia age = 456.8 ± 3.1 Ma SHRIMP II data

420

0.062 0.4

U/Pb GEOCHRONOLOGY

460

0.074

0.066

(ca. 429 Ma: Dunning et al., 1990), their intimate association with D2 reverse faults, and the presence of clasts containing predepositional S1 foliation indicate that they were deposited during the D2 deformational episode, following the exhumation of the D1 tectonites.

500

0.078

0.066

0.082

520

400

0.5

0.6 207

319

235

Pb/ U

0.7

0.8

0.074

455

0.070 0.54

Z5A

0.57

457 ± 2 Ma TIMS data

0.60 207

0.63 235

0.66

Pb/ U

0.69

Figure 9. U/Pb SHRIMP II and TIMS concordia diagrams (2σ, decay constants included). MSWD—mean square of weighted deviates.

320

Zagorevski et al. Badger Bay Victoria Lake Supergroup

5477000 Badger Group

Shoal Arm Fm.

l’s au p el P rou No ok g Bro

5358000 Stanley Waters 457±4 Ma

5475000 Wild Bight Group

ck te ba i le Su p ip ive Cr rus t In

457±3 Ma

593000

480000

Red Indian Lake Group

Red Indian Lake Wigwam Brook Group

462+4/-2 Ma1

5399000 457±3 Ma

457±1 Ma

Victoria Lake Supergroup 5396000

5355000 596000

Sutherlands Pond gr. Mafic volcanic Felsic volcanic

Victoria Lake

483000

518000

524000

Figure 10. Simplified geology of the geochronology sample locations (after McConnell et al., 2002; Rogers et al., 2005; van Staal et al., 2005c). For explanation of the fills, see Figure 5. From Dunning et al. (1987).

contained within a section of sandstone, was collected from the Penny Brook Formation at the top of the Wild Bight Group ~30 m below the stratigraphic contact with the Shoal Arm Formation (B. O’Brien, 2005, personal commun.). The sample contained a small number of euhedral prismatic zircons, many of which contain abundant inclusions. Backscatter SEM images reveal faint oscillatory zoning in many of the grains. As the sample did not yield much zircon, and the presence of xenocrystic zircon in this thin tuff unit was a possibility, the sample was analyzed by SHRIMP. A concordia age, utilizing all of the SHRIMP analyses, is calculated to be 457.5 ± 2.7 Ma (mean square of weighted deviates [MSWD] of concordance and equivalence = 0.82; n = 18) (Fig. 9; Table 1). This date is interpreted to be the crystallization age of the felsic tuff. Noel Paul’s Brook Group Crystal tuff from the Stanley Waters formation of the Noel Paul’s Brook group (van Staal et al., 2005c) was sampled along the eastern shore of the Stanley Waters arm of Victoria Lake (VL01–9104; Z7334; UTM zone 21, 418768 E, 5356382 N; Fig. 10). The Stanley Waters formation forms part of a narrow but regionally extensive belt of coarse volcanogenic graywacke, fine-grained crystal tuff, and minor intercalations of green volcanogenic siltstone and shale (van Staal et al., 2005c). This sequence is directly adjacent to the ca. 565 Ma Valentine Lake monzonite (Evans and Kean, 2002). Although the contact is not exposed, it is presumed to be a reverse fault. The volcanogenic siltstone grades into Caradoc black shale, which caps the sequence. Therefore, the tuff is interpreted to represent the upper part of a volcanogenic sequence at the base of the Caradoc black shale, which forms a regional cover for the Victoria Lake Supergroup. This sample yielded abundant stubby to elongated prismatic zircon grains with melt inclusions through their cores (Fig. 9). The three zircon fractions analyzed define a mixing line (MSWD = 0.114) with an upper intercept of 835 ± 170 Ma and a lower intercept of 456 ± 12 Ma (Table 1; Fig. 9). This line is

anchored by fraction Z1 with a concordant age of 457 ± 3.6 Ma. This concordant age provides the best estimate for the extrusion age of the tuff. Sutherlands Pond Group The first Sutherlands Pond group sample (RAX01–904, z7157; UTM zone 21, 517905 E, 5398510 N: Fig. 10) was collected from a quartz-feldspar phyric rhyolite unit that contains small black quartz phenocrysts. This rhyolite occurs as small beds within a rhyolitic breccia sequence that typifies the Tims Creek formation (Rogers et al., 2005). The sample contains abundant euhedral, prismatic zircons. Three-multigrain zircon TIMS analyses from this sample are highly discordant (15%–57%) owing to the presence of inherited zircon (Table 2, not plotted). Backscatter and cathodoluminescence scanning electron microscope (SEM) images of the zircons reveal the presence of inherited cores rimmed by thick, oscillatory-zoned magmatic rims. A concordia age, calculated from the SHRIMP analyses of magmatic grains and rims (n = 21), is 456.8 ± 3.1 Ma (MSWD of concordance and equivalence = 1.6) (Table 1; Fig. 9). Inherited cores and entirely xenocrystic grains analyzed from this rhyolite range in age from ca. 970 to 1300 Ma (analyses in Table 2 marked by an asterisk, not plotted). The date of 456.8 ± 3.1 Ma is interpreted as the crystallization age of the rhyolite. A second sample from the Sutherlands Pond group (MRB01–06, z7098; UTM zone 21, 521960 E, 5398445 N: Fig. 10) was collected from an aphanitic, altered, brecciated felsic volcanic rock. A very small number of fair quality, small euhedral zircons with abundant inclusions and fractures were retrieved from this sample. There was only enough material to analyze two multigrain fractions. One TIMS analysis is 34% discordant and contains a large inherited component (Z4A; Table 2; Fig. 9). The other analysis (Z5A) is concordant, with an age of 456.9 ± 1.3 Ma. Although there is only one concordant analysis, the age of 457 ± 2 Ma is consistent with the crystallization age of the RAX01–904 sample from the Sutherlands Pond group.

Th (ppm)

Th U Pb* (ppm)

Pb (ppb)

204

771

453

30

73

70

170

169

102

114

237

214

115

72

131

94

144

223

101

86

33

0.606

0.289

0.378

0.435

0.598

0.470

0.398

0.418

0.555

0.599

0.331

0.390

0.488

0.454

0.512

0.498

0.363

0.282

0.297

8

63

7

15

13

23

29

19

21

34

30

27

14

21

16

22

35

21

23

2

7

3

4

1

1

3

4

6

3

3

3

3

5

3

7

0

1

4

47

432

559

495

152

385

123

126

364

274

269

7157-46.1

7157-87.1

7157-43.1

7157-39.1

7157-84.1

7157-55.1

7157-78.1

7157-86.1

7157-99.1

7157-83.1

7157-92.1

171

177

232

35

66

251

104

228

369

277

35

0.656

0.665

0.658

0.285

0.559

0.673

0.710

0.475

0.682

0.663

0.773

21

22

29

9

9

30

12

37

44

34

4

3

2

5

2

5

2

4

1

2

3

4

RAX01-904 (Z7157): Rhyolite, Sutherlands Pond group

8777-39.1

166

8777-45.1

200

293

8777-28.1

108

372

8777-29.1

8777-44.1

266

8777-24.1

8777-41.1

441

282

8777-27.1

369

8777-20.1

359

8777-19.1

214

8777-15.1

8777-18.1

290

8777-14.1

277

462

8777-13.1

191

287

8777-11.1

8777-17.1

316

8777-10.1

8777-16.1

116

8777-7.1

RAX05-900 (Z8777): Felsic tuff, Penny Brook Formation

Spot name

U (ppm)

0.000163

0.000117

0.000227

0.000296

0.000630

0.000094

0.000420

0.000049

0.000059

0.000114

0.001226

0.000037

0.001031

0.000272

0.000368

0.000055

0.000041

0.000164

0.000226

0.000223

0.000134

0.000148

0.000213

0.000179

0.000353

0.000162

0.000229

0.000010

0.000058

0.000539

Pb 206 Pb

204

0.000059

0.000054

0.000067

0.000151

0.000139

0.000046

0.000106

0.000033

0.000049

0.000036

0.000525

0.000054

0.000365

0.000138

0.000217

0.000099

0.000069

0.000138

0.000193

0.000076

0.000073

0.000092

0.000181

0.000163

0.000137

0.000142

0.000063

0.000010

0.000128

0.000380

± 204Pb 206 Pb

0.0028

0.0020

0.0039

0.0051

0.0109

0.0016

0.0073

0.0009

0.0010

0.0020

0.0213

0.0007

0.0179

0.0047

0.0064

0.0010

0.0007

0.0028

0.0039

0.0039

0.0023

0.0026

0.0037

0.0031

0.0061

0.0028

0.0040

0.0002

0.0010

0.0093

f(206)204

0.2079

0.2081

0.2040

0.0992

0.1678

0.2140

0.2295

0.1524

0.2135

0.2109

0.2544

0.1904

0.0703

0.1129

0.1333

0.1866

0.1518

0.1173

0.1370

0.1702

0.1807

0.1035

0.1235

0.1546

0.1280

0.1613

0.1553

0.1197

0.0910

0.0830

Pb 206 Pb

208

0.0049

0.0033

0.0035

0.0062

0.0062

0.0028

0.0053

0.0020

0.0030

0.0033

0.0211

0.0031

0.0143

0.0063

0.0094

0.0052

0.0040

0.0072

0.0098

0.0058

0.0080

0.0046

0.0094

0.0097

0.0064

0.0063

0.0048

0.0028

0.0053

0.0166

± 208Pb 206 Pb

0.5714

0.5723

0.5530

0.5832

0.5335

0.5592

0.5736

0.5660

0.5584

0.5508

0.5945

0.5951

0.4621

0.5684

0.5549

0.5672

0.5749

0.5629

0.5627

0.5481

0.5711

0.5682

0.5641

0.5521

0.5520

0.5744

0.5493

0.5618

0.5737

0.5217

Pb 235 U

207

0.0165

0.0129

0.0152

0.0283

0.0258

0.0113

0.0204

0.0099

0.0140

0.0109

0.0848

0.0153

0.0601

0.0263

0.0380

0.0200

0.0166

0.0260

0.0327

0.0162

0.0227

0.0202

0.0322

0.0304

0.0268

0.0264

0.0141

0.0111

0.0283

0.0643

±207Pb 235 U

TABLE 1. U/Pb SHRIMP ANALYTICAL DATA

0.0729

0.0728

0.0727

0.0723

0.0723

0.0722

0.0719

0.0718

0.0714

0.0712

0.0711

0.0759

0.0715

0.0745

0.0741

0.0738

0.0748

0.0725

0.0721

0.0735

0.0755

0.0742

0.0734

0.0728

0.0723

0.0732

0.0732

0.0730

0.0738

0.0730

Pb 238 U

206

0.0010

0.0009

0.0010

0.0012

0.0011

0.0008

0.0009

0.0008

0.0009

0.0009

0.0015

0.0010

0.0013

0.0010

0.0010

0.0009

0.0009

0.0010

0.0009

0.0009

0.0012

0.0009

0.0010

0.0011

0.0010

0.0010

0.0009

0.0009

0.0009

0.0011

± 206Pb 238 U

0.586

0.627

0.609

0.458

0.435

0.663

0.478

0.754

0.619

0.725

0.267

0.597

0.258

0.408

0.323

0.460

0.535

0.409

0.336

0.523

0.511

0.465

0.346

0.385

0.408

0.422

0.561

0.692

0.362

0.244

Corr Coeff

0.0569

0.0571

0.0552

0.0585

0.0535

0.0561

0.0579

0.0572

0.0568

0.0561

0.0607

0.0569

0.0469

0.0554

0.0543

0.0557

0.0557

0.0563

0.0566

0.0541

0.0548

0.0556

0.0557

0.0550

0.0554

0.0569

0.0544

0.0558

0.0564

0.0518

Pb 206 Pb

207

0.0013

0.0010

0.0012

0.0025

0.0024

0.0009

0.0018

0.0007

0.0011

0.0008

0.0084

0.0012

0.0059

0.0024

0.0036

0.0018

0.0014

0.0024

0.0031

0.0014

0.0019

0.0018

0.0030

0.0028

0.0025

0.0024

0.0012

0.0008

0.0026

0.0062

± 207Pb 206 Pb

454

453

452

450

450

450

448

447

444

443

443

472

445

463

461

459

465

451

449

457

469

461

457

453

450

455

456

454

459

454

Pb 238 U

206

6

5

6

7

7

5

6

5

6

5

9

6

8

6

6

5

6

6

5

5

7

6

6

6

6

6

5

5

5

7

Pb Pb

486

494

420

548

351

458

524

498

482

456

627

486

43

427

385

441

441

464

476

375

406

435

442

412

426

489

388

444

467

278

206

207

Ages (Ma) ± 206Pb 238 U

(continued)

53

39

50

98

102

34

71

26

45

31

330

47

278

98

154

72

56

97

127

58

79

72

125

119

103

95

49

32

106

278

± 207Pb 206 Pb

Th (ppm)

Th U Pb* (ppm)

Pb (ppb)

204 206

204

Pb Pb

124

95

117

75

340

129

81

7157-18.1*

7157-36.1*

7157-101.2*

7157-37.2*

7157-87.2*

7157-80.1*

7157-45.1*

25

66

165

13

38

28

44

27

49

92

51

100

63

346

227

95

118

3

135

166

68

152

0.320

0.526

0.501

0.182

0.339

0.302

0.366

0.544

0.757

1.375

0.718

0.723

0.586

0.717

0.563

0.431

0.426

0.040

0.329

0.550

0.339

0.557

19

28

70

13

21

17

22

9

13

15

13

12

9

41

33

17

22

5

31

25

15

22

2

3

1

6

2

5

3

5

1

7

1

1

2

0

2

4

4

3

5

1

5

5

0.000148

0.000121

0.000021

0.000474

0.000128

0.000380

0.000162

0.000678

0.000123

0.000717

0.000090

0.000110

0.000230

0.000010

0.000081

0.000248

0.000228

0.000650

0.000199

0.000044

0.000386

0.000274

0.000069

0.000066

0.000030

0.000118

0.000098

0.000086

0.000049

0.000211

0.000095

0.000162

0.000095

0.000100

0.000138

0.000010

0.000047

0.000083

0.000086

0.000306

0.000048

0.000093

0.000102

0.000081

± 204Pb 206 Pb

0.0026

0.0021

0.0004

0.0082

0.0022

0.0066

0.0028

0.0118

0.0021

0.0124

0.0016

0.0019

0.0040

0.0002

0.0014

0.0043

0.0040

0.0113

0.0035

0.0008

0.0067

0.0048

f(206)

204

208

Pb Pb

0.0975

0.1605

0.1527

0.0551

0.1050

0.0891

0.1089

0.1539

0.2431

0.4230

0.2278

0.2293

0.1869

0.2272

0.1776

0.1351

0.1332

0.0112

0.1039

0.1796

0.1034

0.1739

206

0.0044

0.0032

0.0016

0.0047

0.0040

0.0039

0.0032

0.0090

0.0052

0.0079

0.0047

0.0050

0.0063

0.0019

0.0025

0.0037

0.0037

0.0115

0.0024

0.0055

0.0043

0.0037

± 208Pb 206 Pb Pb U

2.8578

2.2415

2.0648

1.8560

1.8462

1.7502

1.7502

1.6470

1.7060

1.5824

1.5942

0.6135

0.5954

0.5791

0.5741

0.5666

0.5638

0.5582

0.5726

0.5716

0.5683

0.5568

235

207

0.0708

0.0476

0.0307

0.0606

0.0515

0.0451

0.0324

0.0895

0.0463

0.0687

0.0493

0.0208

0.0274

0.0079

0.0111

0.0177

0.0171

0.0520

0.0131

0.0176

0.0195

0.0215

±207Pb 235 U Pb U

0.2226

0.2008

0.1932

0.1850

0.1791

0.1748

0.1729

0.1686

0.1680

0.1624

0.1614

0.0767

0.0761

0.0752

0.0745

0.0743

0.0741

0.0740

0.0740

0.0739

0.0737

0.0729

238

206

0.0038

0.0027

0.0023

0.0027

0.0026

0.0023

0.0020

0.0030

0.0021

0.0027

0.0021

0.0009

0.0010

0.0008

0.0008

0.0009

0.0009

0.0012

0.0009

0.0008

0.0009

0.0008

± 206Pb 238 U

0.762

0.722

0.849

0.554

0.618

0.601

0.701

0.442

0.555

0.498

0.527

0.464

0.391

0.844

0.648

0.480

0.509

0.291

0.631

0.466

0.461

0.408

Corr Coeff

Pb Pb

0.0931

0.0810

0.0775

0.0728

0.0748

0.0726

0.0734

0.0708

0.0737

0.0707

0.0717

0.0580

0.0567

0.0558

0.0559

0.0553

0.0552

0.0547

0.0561

0.0561

0.0559

0.0554

206

207

0.0015

0.0012

0.0006

0.0020

0.0017

0.0015

0.0010

0.0035

0.0017

0.0027

0.0019

0.0018

0.0024

0.0004

0.0008

0.0015

0.0015

0.0049

0.0010

0.0015

0.0017

0.0020

± 207Pb 206 Pb

1296

1180

1139

1094

1062

1039

1028

1004

1001

970

964

476

473

468

463

462

461

460

460

460

458

454

Pb 238 U

206

20

15

12

15

14

12

11

17

11

15

12

5

6

5

5

5

5

7

5

5

5

5

1490

1220

1134

1007

1062

1003

1026

953

1032

947

976

531

481

446

449

423

420

400

457

457

450

427

Ages (Ma) 207 ± 206Pb Pb 238 206 U Pb

31

29

16

57

45

43

27

104

47

80

55

68

97

17

33

63

60

215

40

62

70

81

± 207Pb 206 Pb

Note: See Stern (1997). Uncertainties reported at 1σ (absolute) and are calculated by numerical propagation of all known sources of error; f(206)204 refers to mole fraction of total 206Pb that is due to common Pb, calculated using the 204Pb method; common Pb composition used is the surface blank.

67

50

69

7157-100.1*

7157-79.1*

73

7157-5.1*

7157-88.1*

112

142

7157-7.1

498

7157-4.1

7157-6.1

416

7157-14.1

69

7157-91.1

287

425

7157-45.3

229

312

7157-20.1

7157-101.1

208

7157-69.1

7157-37.1

281

7157-77.1

RAX01-904 (Z7157): Rhyolite, Sutherlands Pond group (continued)

Spot name

U (ppm)

TABLE 1. U/Pb SHRIMP ANALYTICAL DATA (CONTINUED)

Pb§ (ppm)

Pb# 204 Pb

206

Pb** (pg)

Pb 206 Pb

208

3 9 6

1.10375 0.74132 2.14224

0.70183 0.56821

Pb 235 U

207

0.00233 0.00191 0.00323

0.00134 0.00102

± 1SE Abs

0.11263 0.08806 0.13993

0.07973 0.07344

Pb 238 U

0.00025 0.00019 0.00019

0.00008 0.00011

± 1SE Abs

Isotopic ratios†† 206

0.859 0.793 0.953

0.769 0.696

Corr.§§ Coeff.

0.07107 0.06106 0.11103

0.06385 0.05611

Pb 206 Pb

207

0.00008 0.00010 0.00005

0.00008 0.00007

± 1SE Abs

688.0 544.1 844.3

494.5 456.9

Pb 238 U

206

2.9 2.2 2.1

1.0 1.3

± 2SE

755.2 563.2 1162.5

539.9 456.9

235

Pb U

2.3 2.2 2.1

1.6 1.3

± 2SE

Ages (Ma)## 207

Pb Pb

959.5 641.2 1816.4

736.5 456.8

206

207

4.7 6.7 1.7

5.5 5.8

± 2SE

VL01-9104 (Z7334): Felsic Tuff, Noel Paul’s Brook group Z1 (18) Co,Clr,Pr,Eu,aIn,NM5° 6 219 17 1262 5 0.12 0.58395 0.00231 0.07490 0.00011 0.500 0.05654 0.00020 465.6 1.3 467.0 3.0 473.8 15.3 Z2 (19) Co,Clr,Pr,Eu,aIn,NM5° 10 579 45 2367 11 0.15 0.58595 0.01076 0.07446 0.00124 0.572 0.05708 0.00093 463.0 14.9 468.3 13.8 494.4 70.1 Z3 (20) Co,Clr,Pr,Eu,aIn,NM5° 20 327 26 3228 10 0.13 0.63209 0.00456 0.07926 0.00056 0.967 0.05784 0.00011 491.7 6.6 497.4 5.7 523.7 8.1 *All fractions are zircon and have been abraded following the method of Krogh (1982). Number in parentheses refers to the number of grains in the analysis. † Zircon descriptions: Co—colorless; Clr—clear; fFr—few fractures; aIn—abundant inclusions; rIn—rare inclusions; Eu—euhedral; Pr—prismatic; St—stubby prism; Dia—diamagnetic; M1°—magnetic @ 1.8A, 1°SS; M3°— magnetic @ 1.8A, 3°SS; NM5°—nonmagnetic @ 1.8A, 5°SS. § Radiogenic Pb. # Measured ratio, corrected for spike and fractionation. **Total common Pb in analysis corrected for fractionation and spike. †† Corrected for blank Pb and U and common Pb; errors quoted are 1σ absolute; procedural blank values for this study ranged from 0.1 to 0.3 pg for U and 2 to 10 pg for Pb; Pb blank isotopic composition is based on the analysis of procedural blanks; corrections for common Pb were made using Stacey and Kramers (1975) compositions. §§ Correlation coefficient. ## Corrected for blank and common Pb; errors quoted are 2σ in Ma.

7290 2313 9106

0.18 0.14 0.22

U (ppm)

RAX01-904 (Z7157): Rhyolite, Sutherlands Pond group Z1 (10) Co,Clr,Eu,Pr,rIn,Dia 26 128 15 Z2 (7) Co,Clr,Eu,Pr,rIn,Dia 30 129 12 Z3A (4) Co,Clr,Eu,St,rIn,Dia 29 206 33

Wt. (ug) 0.19 0.21

Description†

MRB01-06 (Z7098): Brecciated felsic volcanic rock, Sutherlands Pond group Z4A (42) Co,Clr,Eu,St,aIn,fFr,M1° 15 251 22 1983 10 Z5A (48) Co,Clr,Eu,St,aIn,fFr,M3° 18 239 19 2503 8

Fract.*

TABLE 2. U-Pb TIMS ANALYTICAL DATA

1.8 6.6 6.3

29.8 15.8 57.0

34.1 0.0

Disc (%)

Zagorevski et al.

Hence we interpret the 457 ± 2 Ma as the crystallization age of the Sutherlands Pond group felsic volcanic rock. DISCUSSION Identification of four new ages of Caradoc peri-Gondwanan volcanism in the Victoria Arc, in addition to a previously obtained age (453 ± 3 Ma: Zagorevski et al., 2007a), demonstrates that Caradoc volcanism was much more widespread than previously postulated (e.g., O’Brien et al., 1997; MacLachlan et al., 2001), allowing more precise timing constraints to be placed on the deposition of the cover and the cessation of magmatism in the Victoria Arc. The age obtained for the Penny Brook Formation extends the age range of volcanism in Notre Dame Bay until 457.5 ± 2.7 Ma and constrains the maximum age of sedimentation in the overlying manganiferrous chert, shale, and black shale of the Caradoc Shoal Arm Formation, which forms part of the Victoria Arc cover sequence. The volcanic rocks from the Sutherlands Pond and Noel Paul’s Brook groups are the same age (457 ± 2; 456.8 ± 3.1; 457 ± 3.6 Ma). Because they predate deposition of most of the black shale, these ages constrain the maximum age of deposition of the cover of the Victoria Arc independently from sparse fossil evidence. Contrasting Ordovician Arc Systems: Red Indian Lake and Victoria Arcs The ages obtained in this study and by Zagorevski (2006) indicate that the Victoria Arc was magmatically active from ca. 473 Ma until 453 Ma. The Victoria Arc magmatism was coeval with the Annieopsquotch accretionary tract, where magmatism initiated at ca. 480 Ma (Annieopsquotch ophiolite belt: Dunning and Krogh, 1985; Lissenberg et al., 2005b) and continued until ca. 460 Ma (Red Indian Lake Arc: Zagorevski et al., 2006). The subsequent juxtaposition of these coeval arc systems along the RIL can make their separation very difficult in the absence of a multidisciplinary data set. Williams et al. (1988) noted the broad contrasts in stratigraphy, structure, fauna, plutonic rocks, radiogenic lead in mineral deposits, magnetic anomalies, and gravity anomalies as the basis for separation of the Dunnage Zone into the Notre Dame and Exploits Subzones. Detailed mapping in central Newfoundland has revealed additional criteria that can be utilized for the separation of the arc systems, namely the zircon provenance and the Sm/Nd isotopic characteristics of volcanic rocks. In the following sections we discuss the differences in the basement characteristics, tectonic history, and cover sequences between the Red Indian Lake and Victoria Arcs. These will then be utilized to reconstruct the Ordovician arc-arc collision in Newfoundland based, in part, on modern analogues in the southwest Pacific. Basement Characteristics The Red Indian Lake and Victoria Arcs can be effectively differentiated on the basis of the zircon inheritance and the Sm/Nd

isotopic characteristics of felsic volcanic rocks. The lowest εNd values recorded in the volcanic rocks of the Exploits Subzone occur in the youngest rocks (Wigwam Brook Group: εNd −4), whereas most of the rocks have generally maintained positive to slightly negative εNd values (εNd +8 to −1) with TDM ages generally in the 1.30–0.6 Ga range (Fig. 11; Zagorevski et al., 2006). In contrast to the Exploits Subzone, volcanic rocks of the Notre Dame Subzone display consistently lower εNd values, with εNd in felsic volcanic rocks commonly in the −3 to −10 range (Whalen et al., 1997; Lissenberg et al., 2005a; Zagorevski et al., 2006), indicating a greater contribution of continental crust and/or more mature continental crust (Fig. 11). Zircon inheritance supports the Sm/Nd isotope data and indicates that the volcanic rocks of the Exploits Subzone have been deposited on continental basement that underwent tectonomagmatic events at ca. 560 Ma and ca. 900–1200 Ma (Zagorevski et al., 2007a). This is supported by the zircon inheritance in the Sutherlands Pond group rhyolite (ca. 970–1200 Ma: Table 2). This basement is similar to the Proterozoic Crippleback Igneous Suite and Sandy Lake Group (ca. 560 Ma, TDM ~1300: Kerr et al., 1995; Rogers et al., 2006), which stratigraphically underlie the oldest parts of the Penobscot arc in Newfoundland (Rogers et al., 2006). Similar relationships occur in the broadly correlative New River Belt in New Brunswick (Johnson and McLeod, 1996). This basement likely represents a fragment of Ganderia (van Staal et al., 2004), a Gondwana-derived microcontinent, which is characterized by zircon provenance in the 0.54–0.55, 0.6–0.8, 1.0–1.55, and 2.5–2.7 Ga age ranges (van Staal et al., 1996), εNd values greater than −4, and TDM ages of late Precambrian igneous rocks spanning 0.9 to 1.35 Ga (Kerr et al., 1995; Rogers et al., 2006).

Notre Dame/ Exploits Subzone Dashwoods sz.

Gander Zone

8

3

εNd

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-2

-7

-12

Age (Ma)

480

460

460

500

540

Legend Mafic Felsic Gneiss

Figure 11. Sm/Nd isotopic evolution of the Notre Dame and Exploits Subzones. Compiled from Jenner and Swinden (1993), Kerr et al. (1995), Swinden et al. (1997), Whalen et al. (1997), MacLachlan and Dunning (1998a, b), Rogers (2004), Lissenberg et al. (2005b), Rogers et al. (2006), and Zagorevski et al. (2006, 2007a).

Tectonic architecture of an arc-arc collision zone The ensialic plutonic and volcanic rocks of the Annieopsquotch accretionary tract and the Red Indian Lake Arc have abundant inheritance in the 935–1845 Ma age range (Dec et al., 1997; Zagorevski et al., 2006), consistent with the presence of Laurentian basement at depth (e.g., Cawood et al., 2001; Cawood and Nemchin, 2001; Cawood et al., 1995). The presence of inheritance in the 1.0–1.55 Ga age ranges is common to both subzones and thus is not a reliable differentiation tool. However, zircon inheritance in the 1.7–1.9 Ga age range is generally associated with the Laurentian margin (Cawood and Nemchin, 2001). In addition, zircon inheritance in the 520–565 Ma age range appears to be restricted to the peri-Gondwanan terranes in the Dunnage Zone, owing to the extensive magmatism of this age in the Ganderia-derived basement (Dunning and O’Brien, 1989; Evans et al., 1990; Rogers et al., 2006). Although ca. 550 Ma magmatism (e.g., Cawood et al., 2001, and references therein) related to the opening of Iapetus is present in Laurentia (Cawood et al., 2001), it is generally restricted to the Humber Zone. Sm/Nd isotopic signatures and the presence of locally abundant zircon inheritance in felsic volcanic rocks indicate the contribution of continental basement to both the Red Indian Lake and Victoria Arcs. The contrasts in the Sm/Nd isotopic characteristics and the presence of specific age ranges of zircon indicate that the nature of the arc basement is fundamentally different and reflects the peri-Laurentian and peri-Gondwanan derivation of the Red Indian Lake and Victoria Arcs, respectively (Fig. 11). Tectonic History The Annieopsquotch accretionary tract and the Victoria Arc underwent distinctly different tectonic histories in the Early to Middle Ordovician. The Annieopsquotch accretionary tract underwent deformation throughout much of its history, starting at ca. 470 Ma and lasting until the collision with the Victoria Arc (Fig. 5; Lissenberg et al., 2005a). Some of this deformation may be related to the simultaneous collision between the Dashwoods microcontinent and the Laurentian margin (Waldron and van Staal, 2001). In contrast, the initiation of magmatism in the Victoria Arc at ca. 473 Ma followed a short orogenic episode (Fig. 5; Penobscot orogeny: Colman-Sadd et al., 1992) in the Exploits Subzone and Gander Zone. From the onset of magmatism, the continental Victoria Arc was extensional (e.g., MacLachlan and Dunning, 1998b; O’Brien et al., 1997; Rogers et al., 2003), opening a wide Tetagouche-Exploits oceanic backarc basin in Newfoundland and New Brunswick (van Staal et al., 1998). The earliest deformation that is common to both the Red Indian Lake and Victoria Arcs started during Caradoc time. This is constrained by the involvement of the Caradoc black shale and volcanic rocks in the mélange zones that mark the D1 thrusts, including the Red Indian Line (e.g., McConnell et al., 2002; Rogers and van Staal, 2002). The D1 mélange zones most likely represent the syncollisional terrane boundaries formed during the accretion of the Victoria Arc to the Laurentian margin. Shear sense indicators along the D1 shear zones suggest sinistral oblique

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SSE-directed emplacement of the Red Indian Lake Arc over the Victoria Arc. This observation is confirmed by seismic-reflection studies of Thurlow et al. (1992) and van der Velden et al. (2004), which clearly show the Red Indian Line to emplace the Annieopsquotch accretionary tract and the Red Indian Lake Arc over the Exploits Subzone and the Victoria Arc (Fig. 8). Interpretation of the migrated seismic reflection data by van der Velden et al. (2004) suggests that the Victoria Arc and its basement were subducted to a depth of at least 18 km and underthrust the Annieopsquotch accretionary tract by at least 60 km (Fig. 8). Cover Sequences of the Notre Dame and Exploits Subzones The contrast between the cover sequences of the Notre Dame and Exploits Subzones is highly distinctive in the Newfoundland Appalachians and can serve as a reliable tool to differentiate the two arc systems. Adhering to the original observations of Williams et al. (1988) the Notre Dame Subzone, and therefore the Annieopsquotch accretionary tract, contain a sub-Silurian unconformity that is overlain by Silurian continental red sandstones and volcanic rocks. In contrast, the rocks of the Exploits Subzone, and thus the Victoria Arc, are overlain by Caradoc black shale and display a generally continuous sedimentation through the Ordovician-Silurian boundary (i.e., Badger Group: Williams et al., 1995). Although several models for the depositional setting of the Badger Group have been proposed (Wasowski et al., 1986), the potential tectonic causes for the contrasts between the subzones have not yet been addressed. We now briefly discuss several key observations that allow a comparison with modern arc-arc collisions. Parts of the Victoria Arc were covered by late Llanvirn limestone (Figs. 5, 7; e.g., Williams, 1995b; O’Brien et al., 1997; McConnell et al., 2002), suggesting relatively shallow water prior to the collision. The deposition of limestone was locally accompanied by formation of mélange, suggesting initiation of deformation in the Caradoc (e.g., Sops Head Complex: McConnell et al., 2002). This was followed by widespread deposition of Caradoc black shale (Fig. 5; e.g., Williams, 1995b). In central Newfoundland, Caradoc black shale of the Laurence Harbour Formation unconformably overlies the volcanic rocks of the Tally Pond Group (ca. 513 Ma: Rogers et al., 2005, 2006), suggesting that the Penobscot basement of the Victoria Arc was locally emergent prior to the Caradoc. The deposition of black shale above the Victoria Arc and its basement, followed by deposition of deepmarine turbidites of the Badger Group (Williams et al., 1993), support rapid submergence of the Victoria Arc (e.g., Williams, 1995b). The zircon provenance of the turbidites in the lower Badger Group suggests derivation from the Laurentian margin or the Notre Dame Subzone (McNicoll et al., 2001), indicating that the collision of the Victoria Arc and the Annieopsquotch accretionary tract had already started by Early Ashgill time. In modern arc-arc collisions, the downgoing plate (i.e., Snellius Plateau and Morobe Shelf) undergoes rapid subsidence, as indicated by drowned carbonate platforms (e.g., Abers and McCaffrey, 1994; Galewsky and Silver, 1997; Pubellier et al.,

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1999). At the same time the overriding plate is rapidly uplifted and forms an emergent orogen (e.g., Talaud Island, Huon-Finisterre Ranges: Fig. 1). The detritus derived from the emergent orogen is transported parallel to the trench and deposited as thick turbidite sequences such as the Markam Canyon turbidites (Galewsky and Silver, 1997; Whitmore et al., 1999). This detritus would have its source in the overriding plate. The concurrence of the subsidence of the Victoria Arc and the influx of the peri-Laurentian detritus suggest that this was caused by the collision with the Red Indian Lake Arc (Figs. 12,

A 465–460 Ma Red Indian

Iapetus Ocean

Lake group (AAT)

Dunnage melange

13). Consistent with this interpretation, structural and seismicreflection studies indicate that the Victoria Arc occupies a lower plate setting with respect to the Red Indian Lake Arc (Fig. 8; e.g., van der Velden et al., 2004). Hence it was partially subducted under the Red Indian Lake Arc, resulting in loading of the Victoria Arc crust, rapid subsidence, and widespread deposition of black shale above the arc and backarc. This is consistent with the absence of the Upper Wild Bight Group equivalents in central Newfoundland, which were probably overridden by the Annieopsquotch accretionary tract during the arc-arc collision (Fig. 8)

Victoria arc

ExploitsTetagouche backarc

Accreted terranes Iapetus Ocean Plate

B

460–455 Ma

Uplift of peri-Laurentian terranes

ExploitsTetagouche backarc

Loading of Victoria arc

RIL SHC

forearc subduction?

C 455–450 Ma

Molucca Iapetus Sea Plate Ocean Plate

Emergent peri-Laurentian terranes WPG

Caradoc black shale and RIL volcanic rocks Future DBL

initiation of subduction

Slab detachment

Iapetus Ocean Plate

D

450–430 Ma

Sub-Silurian unconformity RIL

Badger basin

Future DBL

Figure 12. Tectonic model for the Caradoc arc-arc collision in central Newfoundland (based on van Staal et al., 1998). AAT—Annieopsquotch accretionary tract; DBL—Dog Bay Line; RIL—Red Indian Line; SHC—Sops Head Complex; WPG—Windsor Point Group.

Tectonic architecture of an arc-arc collision zone

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c. 450 Ma

Laurentia Baltica

RILA

VPA

AB ETB

Gan

Gondwana

deri

a

Figure 13. Caradoc paleogeography of arc-arc collision in the AppalachianCaledonian orogen, slightly modified from van Staal et al. (1998). Original figure courtesy of Conall MacNiocaill. ETBAB—Exploits-Tetagouche backarc basin; RILA—Red Indian Lake Arc; VPA—Victoria-Popelogan Arc.

nia

lo Ava

similar to the Snellius Plateau (Fig. 2). The loading and subsidence of the Victoria Arc would have occurred contemporaneously with the Caradoc arc-arc collision, as indicated by the formation of the Caradoc black shale mélange (Fig. 12; e.g., Sops Head Complex: McConnell et al., 2002). Any subduction of the buoyant Victoria Arc crust and its continental basement would have most likely resulted in the emergence of the Annieopsquotch accretionary tract and the Notre Dame Subzone above sea level (e.g., Huon-Finisterre Range: see previous discussion). Although no unconformities of the right age have been yet described within the Annieopsquotch accretionary tract in central Newfoundland, an upper Caradoc unconformity has been defined in southwestern Newfoundland below the Windsor Point Group (453 +5/–4 Ma: Dube et al., 1996), which unconformably overlies the Notre Dame Subzone plutons (ca. 469–488 Ma Cape Ray Igneous Complex: Dube et al., 1996). We interpret this unconformity and the lack of late Caradoc rocks in the Annieopsquotch accretionary tract to imply that the Annieopsquotch accretionary tract had indeed breached sea level, leading to nondeposition and/or erosion throughout the Caradoc (Fig. 12). Hence, the influx of turbidites with

peri-Laurentian provenance and their deposition on the subsided Victoria Arc (McNicoll et al., 2001) signaled the emergence of the Notre Dame Subzone and Laurentia during the arc-arc collision. Syncollisional Magmatism The presence of abundant Caradoc volcanism in the Exploits Subzone indicates that volcanism and deformation were coeval. Caradoc felsic volcanic rocks in the Victoria Arc achieve distinctly low εNd values (εNd −4 in the Wigwam Brook Group: Zagorevski et al., 2006; εNd −4 in the Sutherlands Pond group: Rogers, 2004) that are atypical for the Exploits Subzone but common in the Notre Dame Subzone (Fig. 11). The presence of Caradoc black shale and moderately low εNd values distinguishes the Caradoc volcanic rocks of the Exploits Subzone from the peri-Laurentian felsic volcanic rocks of the Red Indian Lake Group (εNd −6.4 to −7.7; Fig. 11; Zagorevski et al., 2006); however, the distinctly low εNd values in the Caradoc volcanic rocks are rather atypical for the Exploits Subzone, which generally contained positive to slightly negative εNd values throughout its history (Fig. 11). This may have resulted from tapping of a larger proportion or a different part of

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the Ganderian basement during magma generation, leading to a rapid change in εNd values in the Caradoc. An alternate solution is more appealing in terms of an arcarc collision model, although it would be difficult to prove. The low εNd in the Sutherlands Pond group rhyolite may be a result of subduction of peri-Laurentian sediment below the Victoria Arc. Comparison with the geochemical evolution of the HalmaheraSangihe arc collision zone (Macpherson et al., 2003) indicates that this is a feasible, albeit not a unique, solution. In both the Halmahera and Sangihe Arcs, the geochemistry of volcanic rocks that erupted prior to and during the collision indicates increased sediment flux through the subduction zone. The increased sediment flux could be a result of the collision of the facing accretionary prisms, as is probably the case in the Molucca Sea, or deposition of emergent orogen-derived sediment along active trenches (e.g., Markham Canyon turbidites in the Trobriand Trough). A similar explanation can be utilized for the Wigwam Brook Group felsic tuff; however, since the tuff may be epiclastic in nature, the low εNd values could also have been achieved by mixing with sediment of low εNd values. Hence, the most likely source would still be the emergent peri-Laurentian terranes. Although the discussion of the chemical characteristics of the volcanic rocks is beyond the scope of this paper the (pre-) Caradocian volcanic rocks have a range of compositions from arclike to nonarc, which suggests eruption in an extensional suprasubduction zone setting (e.g., Penny Brook Formation: McConnell and O’Brien, 2000; Sutherlands Pond group: Evans and Kean, 2002; Rogers, 2004; Rogers et al., 2005; Diversion Lake Group: Swinden et al., 1989; Evans and Kean, 2002; Rogers et al., 2005; Sops Head Complex: McConnell et al., 2002). An alternate setting for the syncollisional extensionlike magmatism may be above a delaminated Iapetus slab (Fig. 12), similar to the Pliocene and younger volcanic rocks in Mindanao (Sajona et al., 2000). This is an appealing model because it would also explain the change of isotopic characteristics of the younger volcanic rocks in the Exploits Subzone, as both peri-Gondwanan and periLaurentian sources can be involved. Along-Strike Variations Peri-Laurentian magmatism in the Annieopsquotch accretionary tract (ca. 460–480 Ma: Dunning and Krogh, 1985; Zagorevski et al., 2006) was in part coeval with the Victoria Arc (ca. 453–473 Ma: e.g., MacLachlan and Dunning, 1998b; Zagorevski et al., 2007a), requiring subduction on both sides of the Iapetus Ocean, which culminated in a Molucca Sea–type arc-arc collision (Fig. 13; van Staal et al., 1998). The key evidence that can be used to infer the presence of distinct and separated arc sequences in Newfoundland include structural and tectonic history, stratigraphy, basement characteristics, radiogenic lead in mineral deposits, fauna, zircon provenance, and Sm/Nd isotopic characteristics (see previous discussion; Williams et al., 1988). Extrapolation of these interpretations is difficult outside of Newfoundland, where the equivalents of the Dunnage Zone tend to

be poorly preserved owing to metamorphism and deformation as well as the presence of extensive cover sequences. In addition to these obstacles, the age of deformation, tectonic relationships, nature of the basement, and preservation of tectonic elements can vary significantly along the strike of the collision zone. First, both the Molucca and Solomon Sea arc-arc collisions are diachronous along strike of the orogen; hence the duration of arc volcanism, the age of the cover sequences, and the timing of deformation will also be diachronous (Fig. 1). Second, the polarity of thrusting can change along strike and may result in stratigraphic differences in the cover sequences. In the central Molucca collision zone, for example, the Halmahera Arc is subducted under the Sangihe Arc. At the same time, back thrusts emplace the Halmahera Arc over the Sangihe Arc in the northern Molucca Sea (Fig. 2). Hence the polarity of thrusting changes from north to south. Third, long-range correlation of the Appalachian terranes may require adjustments to the local basement characteristics, as the nature of the basement is unlikely to remain exactly the same along the whole strike of the orogen. Finally, the absence of tectonic elements can result from either strike-slip excision (e.g., Elders, 1987) or the local peculiarity of the colliding arc systems. Analogue experiments suggest that the behavior of the overriding arc during the collision will depend on the strength of the arc lithosphere and the retroarc region; i.e., the arc and forearc may be deformed but preserved, the forearc may be subducted, the forearc and part of the arc may be subducted, or the whole arc may be subducted (Boutelier et al., 2003). The lack of preservation of the forearc basin to the Red Indian Lake Arc, for example, indicates that the forearc from the central Newfoundland section was removed by either strike-slip translation or subduction. Northern Appalachians Despite the potential problems with preservation and longrange correlations, equivalents of the Annieopsquotch accretionary tract and Victoria Arc are recognized elsewhere in the northern Appalachians. The remnants of the Dashwoods microcontinent and its Notre Dame Arc suprastructure in New England are preserved in the Shelburne Falls Arc (ca. 470–485 Ma: Karabinos et al., 1998) and the Chain Lakes Massif (>473 Ma: Gerbi et al., 2006) both of which have clear peri-Laurentian affinities. The Annieopsquotch accretionary tract is largely obscured in Atlantic Canada by cover sequences; however, correlatives of the Annieopsquotch accretionary tract include the Boil Mountain Complex and the Jim Pond Formation (ca. 477–485 Ma: Coish and Rogers, 1987; Gerbi et al., 2006), which are adjacent to the Chain Lakes Massif (Fig. 3) as well as to the Caucomgomoc inlier (van Staal et al., 1998). The equivalents of the Victoria Arc (Fig. 3: Popelogan Inlier: Wilson, 2003; Munsungun inlier: Ayuso and Schulz, 2003; Ayuso et al., 2003) and backarc (Fig. 3: Bathurst Supergroup: van Staal et al., 2003) are well preserved in New Brunswick and Maine; however, the equivalents of the Victoria Arc in western Maine, New Hampshire, and Vermont are less clear. The Middle to Late Ordovician Bronson Hill Arc (Moench and Aleinikoff, 2003),

Tectonic architecture of an arc-arc collision zone which lies to the east of the peri-Laurentian Shelburne Falls Arc, is a probable correlative of the Munsungun inlier, which is an equivalent of the Popelogan Arc in Maine (Ayuso and Schulz, 2003; Ayuso et al., 2003). The Bronson Hill Arc contains two distinct sequences: the Ammonoosuc (457–470 Ma) and Quimby (435–456 Ma) sequences, which are thought to have formed above subduction zones of opposite polarity (Moench and Aleinikoff, 2003). Although the Ammonoosuc sequence has been tectonically correlated with the peri-Laurentian Shelburne Falls Arc (e.g., Moench and Aleinikoff, 2003), the contact relationships are obscured by the Connecticut Valley Trough and younger deposits, and the age span of the Ammonoosuc sequence overlaps with the age of the Victoria Arc, although relationships with the Penobscot Arc basement have not been identified. In addition, the Ammonoosuc sequence shares several characteristics with the periGondwanan Victoria Arc (van Staal et al., 1998; Hibbard et al., 2004), including an important period of extension as exemplified by the Chickwolnepy sheeted intrusions (ca. 467 Ma; Fitz, 2002; Moench and Aleinikoff, 2003). The cover sequences are also remarkably similar to the Victoria Arc. The type Ammonoosuc volcanic rocks are conformably overlain by the Caradoc sulfidic black shale of the Partridge Formation, followed by an unconformity and deposition of marine graywacke of the Ashgill Quimby Formation (Moench and Aleinikoff, 2003). The presence of a Caradoc-Ashgill unconformity between the Ammonoosuc and Quimby sequences (Moench and Aleinikoff, 2003) differentiates the Bronson Hill Arc from the Victoria Arc; however, this can be easily reconciled by considering the along-strike variations in an arc-arc collision zone. The Quimby sequence, overlying the Ammonoosuc sequence, comprises marine shales, graywackes, and conglomerates intruded by Upper Ordovician to Lower Silurian plutons (Moench and Aleinikoff, 2003). The Quimby sequence corresponds to the Bronson Hill Arc in southern New Hampshire and central Massachusetts (454–442 Ma: Tucker and Robinson, 1990; Karabinos et al., 1998). This phase of the Bronson Hill Arc development has been interpreted to have occurred following an arc collision of the Shelburne Falls Arc and the Ammonoosuc sequence of the Bronson Hill Arc with the Laurentian margin (Karabinos et al., 1998; Moench and Aleinikoff, 2003). The collision was followed by establishment of a west-dipping subduction zone along the composite Laurentian margin following a polarity flip (Karabinos et al., 1998; Moench and Aleinikoff, 2003). The tectonic evolution of New England broadly mirrors the Ordovician arc-arc collision in Newfoundland. The accretion of the Ammonoosuc sequence of the Bronson Hill Arc to the composite Laurentian margin, comprising the peri-Laurentian Shelburne Falls Arc and Jim Pond Formation, occurred at ca. 457 Ma, followed by a magmatic gap lasting several million years, deepmarine sedimentation, and probable subduction reversal (Moench and Aleinikoff, 2003). This closely agrees with the accretion of the Victoria Arc to the Red Indian Lake Arc, where the Victoria arc was loaded, leading to subsidence and deep-marine sedimentation (Fig. 12B, C). Subsequent convergence was accommodated

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by a change of polarity and a step-back of the subduction into the Tetagouche-Exploits backarc (Fig. 12C, D) and formation of the Quimby sequence of the Bronson Hill Arc. British Caledonides A direct correlation between Newfoundland and the British Caledonides is difficult, because the well-preserved sedimentary sequences of the late Llanvirn to the Late Silurian Southern Uplands accretionary complex (e.g., Legget, 1987) are largely absent in the northern Appalachians, a feature that may reflect strike-slip excision or subduction. Van Staal et al. (1998) proposed a correlation of the Annieopsquotch accretionary tract with the South Connemara Group in Ireland (e.g., Ryan and Dewey, 2004) and the correlative Northern Belt of the Southern Uplands terrane in Scotland. This correlation was in part supported by Armstrong and Owen (2001), who proposed a new terrane, Novantia, to be a correlative of the Annieopsquotch accretionary tract. Novantia has been subducted underneath the Southern Uplands; however, its presence was deduced on the basis of geophysical modeling. Armstrong and Owen (2001) also recognized correlatives of the Victoria Arc in the Grangegeeth Terrane of eastern Ireland and proposed that these have also been in part subducted under the Southern Uplands. The juxtaposition of the Novantia and Grangegeeth terranes in the Ordovician was subsequently supported by the presence of distinctly peri-Gondwanan (ca. 560 Ma) zircons in the Caradoc Portpatrick Formation of the Southern Uplands (Phillips et al., 2003). Comparison of the Newfoundland arc-arc collision with the Southern Uplands illustrates the importance of along-strike variation in a collisional setting. The Southern Uplands accretionary complex contains both the peri-Laurentian and the periGondwanan provenance, indicating emergence of both Laurentia and the Grangegeeth Terrane during the collision. The Badger Group in Newfoundland, on the other hand, contains a distinctly peri-Laurentian provenance (McNicoll et al., 2001), supporting the emergence of the peri-Laurentian terranes and subsidence of the Victoria Arc during the collision. CONCLUSIONS Comparison with the modern analogues indicates that the recognition of ancient arc-arc collision zones requires an integrated and multidisciplinary approach. In Newfoundland, two coeval but distinct and separated arc systems are preserved in the Dunnage Zone: the peri-Laurentian Red Indian Lake Arc (ca. 473–460 Ma) and the peri-Gondwanan Victoria Arc (ca. 473–453 Ma). The two arc systems collided above a doubly dipping subduction zone during early to middle Caradoc time along the Red Indian Line (Figs. 12, 13), the main Iapetus suture zone, marking the end of the Taconic orogeny and the start of the Salinic orogeny. Following the collision the subduction stepped back into the ExploitsTetagouche backarc basin, leading to the closure of that tract along the Silurian Dog Bay Line (Williams et al., 1993; van Staal et al., 1998).

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The separation of the Annieopsquotch accretionary tract and the Victoria Arc is supported by contrasts in structural and tectonic history, stratigraphy, basement characteristics, radiogenic lead in mineral deposits, and fauna that are apparent across the Red Indian Line (Williams et al., 1988). A brief review of the northern Appalachians and the British Caledonides reveals great similarity in the correlative tracts that also preserve evidence of an Ordovician arc-arc collision, although the interpretation of the polarity of subduction zones (e.g., Armstrong and Owen, 2001; Moench and Aleinikoff, 2003) differs from the models based on Newfoundland (e.g., van Staal et al., 1998). Future research should aim to reconcile these differences. The timing of the arcarc collision appears to have been restricted to the Caradoc in all of the correlative tracts. This similarity may reflect either coeval collision along the entire length of the tract or the postcollisional dispersal of the accreted terranes along the Laurentian margin (e.g., Elders, 1987).

heavy liquid techniques. Mineral separates were sorted by magnetic susceptibility using a FrantzTM isodynamic separator. Multigrain zircon fractions analyzed were very strongly air abraded following the method of Krogh (1982). Treatment of analytical errors follows Roddick et al. (1987) with errors for the ages reported at the 2σ level (Table 2). U-Pb concordia diagrams are presented in Figure 9. The data are plotted in concordia diagrams with errors at the 2σ level (Fig. 9), using Isoplot v. 2.49 (Ludwig, 2001) to generate the plots. A concordia age (Ludwig, 1998) is calculated for some of the samples presented in this paper. A concordia age incorporates errors in the decay constants and includes both an evaluation of concordance and an evaluation of equivalence of the data (how well the data fit the assumption that they are repeated measurements of the same point). The calculated concordia ages and errors quoted in the text are at 2σ, with decay constant errors included. ACKNOWLEDGMENTS

APPENDIX: U-Pb GEOCHRONOLOGY ANALYTICAL TECHNIQUES SHRIMP II analyses were conducted at the Geological Survey of Canada (GSC) using analytical procedures described by Stern (1997), with standards and U-Pb calibration methods following Stern and Amelin (2003). Zircons from the samples were cast in 2.5 cm diameter epoxy mounts (GSC mount 373 for sample RAX05–900 and GSC mount 254 for sample RAX01– 904 along with fragments of the GSC laboratory standard zircon (z6266, with 206Pb/238U age = 559 Ma). The midsections of the zircons were exposed using 9, 6, and 1 µm diamond compound, and the internal features of the zircons were characterized with backscatter electrons (BSE) and cathodoluminescence (CL) utilizing a Cambridge Instruments scanning electron microscope (SEM). Mount surfaces were evaporatively coated with 10 nm of high-purity Au. Analyses were conducted using a 16O– primary beam, projected onto the zircons at 10 kV. The sputtered area used for analysis was ca. 25 µm in diameter with a beam current of ~5 nA and 13 nA for RAX05–900 and RAX01–904, respectively. The count rates of 10 isotopes of Zr+, U+, Th+, and Pb+ in zircon were sequentially measured over 6 scans with a single electron multiplier and a pulse counting system with deadtime of 22 ns. Off-line data processing was accomplished using customized in-house software. The 1σ external errors of 206Pb/238U ratios reported in Table 1 incorporate a ±1.0% error in calibrating the standard zircon (see Stern and Amelin, 2003). No fractionation correction was applied to the Pb isotope data; common Pb correction utilized the measured 204Pb/206Pb and compositions modeled after Cumming and Richards (1975). The 206Pb/238U ages for the analyses have been corrected for common Pb using both the 204 and 207 methods (Stern, 1997), but there is generally no significant difference in the results. U-Pb TIMS analytical methods utilized in this study are outlined in Parrish et al. (1987). Heavy mineral concentrates were prepared by standard crushing, grinding, WilfleyTM table, and

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Printed in the USA

The Geological Society of America Special Paper 436 2008

The Catalina Schist: Evidence for middle Cretaceous subduction erosion of southwestern North America M. Grove* Department of Earth & Space Sciences, University of California, Los Angeles, 3806 Geology, Los Angeles, California 90095, USA G.E. Bebout Department of Earth & Environmental Sciences, Lehigh University, 31 Williams Drive, Bethlehem, Pennsylvania 18015, USA C.E. Jacobson Department of Geological and Atmospheric Sciences, 253 Science I, Iowa State University, Ames, Iowa 50011-3212, USA A.P. Barth Department of Earth Sciences, Indiana–Purdue University, 723 West Michigan Street, Indianapolis, Indiana 46202, USA D.L. Kimbrough Department of Geological Sciences, San Diego State University, 5500 Campanile Drive, San Diego, California 92182-1020, USA R.L. King School of Earth and Environmental Sciences, Washington State University, Pullman, Washington 99164-2812, USA Haibo Zou O.M. Lovera Department of Earth & Space Sciences, University of California, Los Angeles, 3806 Geology, Los Angeles, California 90095, USA B.J. Mahoney Department of Geology, University of Wisconsin–Eau Claire, Phillips 157, Eau Claire, Wisconsin 54702-4004, USA G.E. Gehrels Department of Geosciences, University of Arizona, Tucson, Gould-Simpson Building 529, Tucson, Arizona 85721, USA

ABSTRACT The Catalina Schist underlies the inner southern California borderland of southwestern North America. On Santa Catalina Island, amphibolite facies rocks that recrystallized and partially melted at ca. 115 Ma and at 40 km depth occur atop an inverted metamorphic stack that juxtaposes progressively lower grade, high-pressure/ temperature (PT) rocks across low-angle faults. This inverted metamorphic sequence

*[email protected] Grove, M., Bebout, G.E., Jacobson, C.E., Barth, A.P., Kimbrough, D.L., King, R.L., Zou, H., Lovera, O.M., Mahoney, B.J., and Gehrels, G.E., 2008, The Catalina Schist: Evidence for middle Cretaceous subduction erosion of southwestern North America, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 335–361, doi: 10.1130/2008.2436(15). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Grove et al. has been regarded as having formed within a newly initiated subduction zone. However, subduction initiation at ca. 115 Ma has been difficult to reconcile with regional geologic relationships, because the Catalina Schist formed well after emplacement of the adjacent Peninsular Ranges batholith had begun in earnest. New detrital zircon U-Pb age results indicate that the Catalina Schist accreted over a ~20 m.y. interval. The amphibolite unit metasediments formed from latest Neocomian to early Aptian (122–115 Ma) craton-enriched detritus derived mainly from the pre-Cretaceous wall rocks and Early Cretaceous volcanic cover of the Peninsular Ranges batholith. In contrast, lawsonite-blueschist and lower grade rocks derived from Cenomanian sediments dominated by this batholith’s plutonic and volcanic detritus were accreted between 97 and 95 Ma. Seismic data and geologic relationships indicate that the Catalina Schist structurally underlies the western margin of the northern Peninsular Ranges batholith. We propose that construction of the Catalina Schist complex involved underthrusting of the Early Cretaceous forearc rocks to a subcrustal position beneath the western Peninsular Ranges batholith. The heat for amphibolite facies metamorphism and anatexis observed within the Catalina Schist was supplied by the western part of the batholith while subduction was continuous along the margin. Progressive subduction erosion ultimately juxtaposed the high-grade Catalina Schist with lower grade blueschists accreted above the subduction zone by 95 Ma. This coincided with an eastern relocation of arc magmatism and emplacement of the ca. 95 Ma La Posta tonalite-trondjhemite-granodiorite suite of the eastern Peninsular Ranges batholith. Final assembly of the Catalina Schist marked the initial stage of the Late Cretaceous–early Tertiary craton-ward shift of arc magmatism and deformation of southwestern North America that culminated in the Laramide orogeny. Keywords: Catalina Schist, U-Pb, zircon, subduction erosion.

INTRODUCTION Subduction erosion is a fundamental convergent margin process that tectonically shortens the forearc region by displacing rocks from the overriding plate to subcrustal positions and/or into the asthenosphere (Scholl et al., 1980; von Huene and Scholl, 1991, 1993). Roughly half of all present-day convergent margins exhibit geologic and/or geophysical evidence for subductionrelated removal of rocks from the forearc regions of the overriding plate to positions beneath the arc and/or into the mantle (e.g., Clift and Vannucchi, 2004). The globally average rate at which continental debris is transported toward the mantle is significant (~2.5 km3/year; Scholl and von Huene, 2007). Unfortunately, its consumptive nature means that direct evidence for subduction erosion in ancient convergent margins is generally ambiguous or lacking. Indirect evidence for subduction erosion such as “missing” or “telescoped” forearc crust is difficult to distinguish from alternative tectonic processes such as margin-parallel strike-slip faulting (e.g., Karig, 1980). A classic illustration of the difficulty in deciding between subduction erosion and margin-parallel strike-slip dispersal of forearc rocks is provided by the Nacimiento fault in west-central California (Page, 1970). There, plutonic and metamorphic rocks of the Salinian block are juxtaposed against a high-pressure, low-temperature mélange of the Sur-Obispo terrane. This anomalous juxtaposition of lithotectonic belts has been explained both by underthrusting (subduction erosion) of

intervening rocks (Page, 1981; Hall, 1991) or by lateral transport of hundreds (Hill and Dibblee, 1953; Suppe, 1970; Dickinson, 1983; Dickinson et al., 2005) to thousands of kilometers (Page, 1982; McWilliams and Howell, 1982; Vedder et al., 1983; Debiche et al., 1987). Like the Sur-Obispo terrane, seismic data and geologic relationships indicate that the high-pressure/temperature (PT) Catalina Schist of southern California is also in direct contact with its coeval magmatic arc. This relationship was revealed by middle Miocene extension that exhumed the Catalina Schist from beneath the northwestern margin of the Peninsular Ranges batholith across an east-dipping Miocene detachment fault (Crouch and Suppe, 1993) (Fig. 1). The dramatic extension of the borderland occurred during a middle Miocene microplate capture event (e.g., Nicholson et al., 1994) triggered by Pacific–North American shearing along the margin (e.g., Atwater, 1970). During this event, Late Cretaceous forearc strata and underlying Jurassic and earliest Cretaceous basement rocks were ripped from the western margin of the Peninsular Ranges batholith, rotated clockwise, and translated northwestward into the outer California borderland along a detachment system underlain by the Catalina Schist (Wright, 1991; Crouch and Suppe, 1993; Bohannon and Geist, 1998; Ingersoll and Rumelhart, 1999). Although the Catalina Schist contains abundant blueschist facies rocks typical of subduction zone environments, it also has higher temperature lithologies that define an inverted

Figure 1. Geologic setting of the Catalina Schist. Location map in the upper right shows the distribution of highpressure/temperature (PT) subduction complexes and Cretaceous–Tertiary batholiths in southwestern North America. The offshore distribution of subduction complexes is estimated primarily from Crouch and Suppe (1993), Bohannon and Geist (1998), Sedlock (1988a, b), Bonini and Baldwin (1998), and Fletcher et al. (2007). The box outlines the location of the southern California continental borderland. The simplified map of the borderland shown in the lower left is based upon Bohannon and Geist (1998), whereas cross section Y–Y′ is after Crouch and Suppe (1993). Symbols P and P* denote formerly contiguous rocks that dextrally sheared and rotated clockwise during middle Miocene rifting. The box in the lower left shows the location of Catalina Island. The geologic map and cross section X–X′ of Catalina Island in the upper left are after Platt (1976). The distribution of lawsonite-albite and actinolite-albite rocks is from Altheim et al. (1997). Note that epidote-blueschist and epidote-amphibolite rocks are distinguished from Platt’s (1976) greenschist unit. A coherent kilometer-scale mass of epidote-amphibolite facies metagabbro that directly underlies the amphibolite unit was included by Platt (1976) within his amphibolite unit. K–T—Cretaceous–Tertiary.

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metamorphic sequence that includes amphibolite facies rocks at the highest structural levels (see inset in Fig. 1). These partly melted rocks underwent peak grade conditions at ca. 115 Ma (Suppe and Armstrong, 1972; Mattinson, 1986; Sorensen and Barton, 1987; Grove and Bebout, 1995; Anczkiewicz et al., 2004). A longstanding explanation for the unusually high-temperature, inverted metamorphism of the Catalina Schist is that it reflects progressive accretion of early subducted rocks beneath unrefrigerated mantle lithosphere during nascent subduction (e.g., Platt, 1975). Although Platt’s (1975) nascent subduction hypothesis was based upon the Catalina Schist, he extended it to explain the origin of meter- to kilometer-scale high-grade blocks within the Franciscan Complex and other accretionary complexes along the western margin of North America (e.g., Platt, 1975; Cloos, 1985; Wakabayashi, 1990, 1992, 1999; Anczkiewicz et al., 2004; Wakabayashi and Dumitru, 2007). The nascent subduction hypothesis makes sense for the Franciscan Complex because the blocks are coeval with a major pulse of Middle Jurassic magmatic activity along the western North American continental margin (Evernden and Kistler, 1970; Saleeby and Sharp, 1980; Stern et al., 1981; Chen and Moore, 1982; Wright and Fahan, 1988; Barton et al., 1988; Staude and Barton, 2001; Irwin, 2002). The Middle Jurassic orogenesis included formation of the Coast Range ophiolite (Hopson et al., 1981; Shervais et al., 2005), which is considered by some (e.g., Stern and Bloomer, 1992) to be a product of the initiation of Middle Jurassic subduction (see also Dickinson et al., 1996). The equivalence in age of the Coast Range ophiolite and the high-grade blocks of the Franciscan Complex corresponds well with a common relationship exhibited by many of the world’s ophiolites: namely, age equivalence of the ophiolite and its metamorphic sole (Jamieson, 1986; Hacker, 1990a, b; Wakabayashi and Dilek, 2000). Whereas establishment of a new subduction regime during the Middle Jurassic may generally explain the occurrence of high-grade rocks of this age within the Franciscan Complex of central and northern California, an equivalent event at ca. 115 Ma off the southern California margin makes little sense with respect to the convergent margin evolution of southwest North America. Emplacement of the 750-km-long Peninsular Ranges batholith had begun by Middle Jurassic time (Shaw et al., 2003) and was well under way prior to the proposed subduction-initiation event at 115 Ma (Silver and Chappell, 1988; Kistler et al., 2003). Furthermore, no middle Cretaceous ophiolite exists within the southern California region. Only fragments of older forearc basement are preserved. The latter is best represented by Triassic and Middle Jurassic ophiolitic rocks that border the southern part of the Peninsular Ranges batholith (Kimbrough and Moore, 2003). Just as in the case of the Franciscan Complex, the Middle Jurassic ophiolite of west-central Baja California is associated with high-PT blocks of equivalent age on the Vizcaino Peninsula and Cedros Island (Fig. 2; Baldwin and Harrison, 1989, 1992). This paper presents new detrital zircon U-Pb results from metagraywackes of the major metamorphic units of the Catalina Schist that clarify its accretion history. We confirm a genetic

relationship to forearc and batholith rocks of the adjacent Peninsular Ranges batholith. These new results indicate that the anomalously high-T, high-PT amphibolite facies rocks of the Catalina Schist had been deposited, accreted, and metamorphosed at peak grade conditions ~15–20 m.y. before the lawsonite-blueschist and lower-temperature, high-PT rocks of the complex were accreted. We conclude that amphibolite facies, epidote-amphibolite, and possibly epidote blueschist units of the Catalina Schist formed from forearc strata and basement rocks that were underthrust and sheared together with lithospheric mantle beneath the western margin of the Peninsular Ranges batholith in a subduction erosion process. Our model is consistent with recent deep seismic-reflection imaging of subparallel megathrusts within modern subduction zones (e.g., Calvert, 2004). Continued subduction erosion ultimately juxtaposed the high-grade rocks of the Catalina Schist with lawsonite-blueschists and lower grade rocks that had formed within a subduction zone setting. BACKGROUND Catalina Schist High-Pressure/Temperature Complex High-PT rocks distributed along the western margin of southern and Baja California appear to represent the southern continuation of the better exposed Franciscan Complex of central and northern California (Fig. 1; Woodford, 1924; Suppe and Armstrong, 1972; Kilmer, 1979; Blake et al., 1984). Within the southern California region, on-land exposures of a high-PT complex referred to as the Catalina Schist occur on Santa Catalina Island (Bailey, 1941; Platt, 1976) and the Palos Verdes Peninsula (Woodring et al., 1946; Dibblee, 1999). More widespread submarine exposures of the Catalina Schist occur throughout the inner southern California borderland, and equivalent basement is detected in boreholes within the western and southwestern Los Angeles Basin (Schoellhamer and Woodford, 1951; Yerkes et al., 1965; Yeats, 1968, 1973; Sorensen, 1985, 1988a, 1988b; Wright, 1991; Crouch and Suppe, 1993; Bohannon and Geist, 1998; ten Brink et al., 2000). The Catalina Schist on Santa Catalina Island has been described by Woodford (1924), Bailey (1941), Platt (1975, 1976), Sorensen and Barton (1987), Sorensen (1986, 1988a, b), Sorensen and Grossman (1989), Bebout and Barton (1989, 1993, 2002), and Grove and Bebout (1995) and is only briefly summarized here. It consists of metasedimentary, metavolcanic, and ultramafic protoliths that were metamorphosed and sheared together under amphibolite facies to lawsonite-blueschist facies and lower grade conditions during the middle Cretaceous (Suppe and Armstrong, 1972; Platt, 1976; Mattinson, 1986; Sorensen and Barton, 1987; Sorensen, 1988a, 1988b; Grove and Bebout, 1995). Individual tectonic slices (Fig. 1) contain rocks of broadly equivalent metamorphic grade. These are juxtaposed across low-angle faults in an apparently inverted metamorphic sequence (Platt, 1976; Fig. 1). The structurally highest unit is an amphibolite facies shear zone composed primarily of intercalated and metasomatically altered

Figure 2. Geologic map of the Peninsular Ranges batholith, modified after Gastil et al. (1975) and Kimbrough et al. (2001). Generalized stratigraphic relationships of the forearc region in the northern Vizcaino Peninsula modified after Kimbrough et al. (2001), Kimbrough and Moore (2003), and references cited within these papers. Simplified stratigraphy of the northern Santa Ana Mountains is based upon references provided within Lovera et al. (1999). Western limit of the forearc is the restored pre–middle Miocene position based upon Bohannon and Parsons’s (1995) reconstruction and our interpretation of aeromagnetic data presented by Langenheim and Jachens (2003). PRB—Peninsular Ranges batholith; TTG—tonalite, trondjhemite, granodiorite.

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mafic and former harzburgite/dunite protoliths (Sorensen and Barton, 1987; Sorensen, 1988b; Bebout and Barton, 1993, 2002). The proportion of sediment, particularly immature graywacke, increases structurally downward. At the structurally lowest levels the lawsonite-blueschist and lower grade units are sediment dominated. Each of the major metamorphic units is characterized by meter- to kilometer-scale, compositionally heterogeneous shear zones that appear to have facilitated metasomatic fluid infiltration at near peak-grade conditions (Bebout and Barton, 1989, 1993, 2002; King et al., 2006, 2007). Primary geologic mapping of the Catalina Schist on Santa Catalina Island was performed by Bailey (1941) and Platt (1976). Although we rely heavily upon Platt’s (1976) mapping of the central part of the island, we recognize different tectonometamorphic units than he did. Platt’s (1976) Catalina greenschist unit has been subdivided into epidote-amphibolite and epidoteblueschist units (Grove and Bebout, 1995). Results presented in Grove and Bebout (1995) and in this paper indicate that the epidote-amphibolite and epidote-blueschist rocks had important differences in accretion and metamorphic history. In addition, we have delineated lawsonite-albite and actinolite-albite facies units on the west end of the island. These lower grade rocks appear to structurally underlie a lawsonite-blueschist unit (Altheim et al., 1997). Because the Catalina Schist terrane was significantly extended as it was exhumed during the early middle Miocene formation of the inner continental borderland (Wright, 1991; Crouch and Suppe, 1993; Nicholson et al., 1994; Bohannon and Geist, 1998; Ingersoll and Rumelhart, 1999), original contacts bounding the major units within the schist are likely to have been strongly modified or excised altogether by middle Tertiary structures. In addition, the extent to which the Santa Catalina Island exposures of Catalina Schist represent the much larger area of high-PT rocks that underlie the inner continental borderland is difficult to assess. Study of widely distributed deposits of the San Onofre Breccia indicates that the Santa Catalina Island exposures are probably broadly representative of the overall terrane (Stuart, 1979). The San Onofre Breccia is an early middle Miocene deposit that accumulated as the Catalina Schist was exhumed (Wright, 1991; Crouch and Suppe, 1993). Clast counts performed by Stuart (1979) appear to be broadly consistent with an unroofing sequence in which amphibolite and epidote-amphibolite grade Catalina Schist and low-PT mafic basement from the hanging wall were much more abundant at lower stratigraphic levels within the San Onofre Breccia. Our own observations confirm this relationship for the thick deposits of this breccia that accumulated along the western margin of the Peninsular Ranges batholith (Fig. 1). In contrast, the stratigraphically younger San Onofre Breccia, including deposits laid down atop extended borderland crust underlain by the Catalina Schist, tends to be dominated by lawsonite-blueschist and lower-grade detritus. Based upon these relationships, we are reasonably confident that our sampling of the Catalina Schist on Santa Catalina Island is broadly representative of the Catalina Schist terrane as a whole.

Forearc Basement and Strata The relatively complete forearc sequence of the Vizcaino Peninsula at the southern end of the Peninsular Ranges batholith provides the basis to interpret the forearc of the highly disrupted southern California borderland (Figs. 1, 2). On the Vizcaino Peninsula and on Cedros Island, both Late Triassic (221 Ma) and Middle Jurassic (173 Ma) ophiolite sequences (Moore, 1985; Kimbrough, 1985; Kimbrough and Moore, 2003) are the basement for Upper Jurassic, Lower Cretaceous, Upper Cretaceous, and lower Cenozoic arc volcanic and forearc sedimentary rocks (Boles, 1978; Kilmer, 1979; Kienast and Rangin, 1982; Boles and Landis, 1984; Patterson, 1984; Smith and Busby, 1993; Busby et al., 1998; Kimbrough et al., 2001; Critelli et al., 2002; Busby, 2004). High-PT Cretaceous rocks of the Western Baja terrane constitute an accretionary complex that is juxtaposed beneath the Triassic and Jurassic ophiolitic rocks well outboard of the batholithic margin (Suppe and Armstrong, 1972; Moore, 1986; Sedlock, 1988a, b; Baldwin and Harrison, 1989; Baldwin and Harrison, 1992). Only fragmentary evidence exists for the early Mesozoic forearc basement within the southern California area. Within the outer California continental borderland, offshore drilling and seismic exploration detect mafic basement beneath forearc strata (Bohannon and Geist, 1998; ten Brink et al., 2000). Surface exposures of lower Mesozoic basement rocks that occur on Santa Cruz Island (the Willows Complex of Weaver and Nolf, 1969; Hill, 1976; Mattinson and Hill, 1976) are correlated with the Coast Range ophiolite (Jones et al., 1976). The associated Santa Cruz Island Schist and the correlative Santa Monica slate of the western Transverse Ranges exhibit distinctly arclike compositions and could be equivalent to accreted Jurassic terranes of the Sierran Foothills (Sorensen, 1985, 1988a). “Saussurite” gabbro similar to altered zones of the Willows Complex gabbros is in fault contact with the Catalina Schist on Santa Catalina Island (Platt, 1976) and within the subsurface of the southwestern Los Angeles Basin (Schoellhamer and Woodford, 1951; Yeats, 1973; Sorensen, 1985, 1988b). Saussurite gabbro clasts are a major component within the San Onofre Breccia (Stuart, 1979) and confirm that low-PT mafic basement was exhumed along with the Catalina Schist during middle Miocene borderland rifting. Cenomanian and younger forearc strata onlap the western margin of the northern Peninsular Ranges batholith (Fig. 2; Woodring and Popenoe, 1942; Yerkes et al., 1965; Flynn, 1970; Nordstrom, 1970; Peterson and Nordstrom, 1970; Kennedy and Moore, 1971; Sundberg and Cooper, 1978; Schoellhamer et al., 1981; Nilsen and Abbott, 1981; Bottjer et al., 1982; Bottjer and Link, 1984; Fry et al., 1985; Girty, 1987; Bannon et al., 1989). Thick sections of Upper Cretaceous strata also occur throughout the outer borderland (Howell and Vedder, 1981; Vedder, 1987; Bohannon and Geist, 1998). The existence of Lower Cretaceous forearc strata is far less certain, however. Sedimentary rocks of this age have not been described anywhere along the western margin of the northwestern Peninsular Ranges, the Santa Monica Mountains, or on the offshore islands. Their existence in the subsurface

The Catalina Schist is mostly inferred in all areas that have been drilled (Howell and Vedder, 1981; Vedder, 1987; Bohannon and Geist, 1998; ten Brink et al., 2000). Our model for the formation of the Catalina Schist explicitly accounts for the scarcity of these Lower Cretaceous forearc rocks in the southern California region (see Discussion). Peninsular Ranges Batholith The Peninsular Ranges batholith of southern and Baja California constitutes a classic Cordilleran continental margin batholith (Fig. 2; Larsen, 1948; Jahns, 1954; Gastil et al., 1975). The better studied northern segment of this batholith consists of longitudinal western and eastern zones based on age, petrology, prebatholithic wall rock, geophysical parameters, and depth and style of emplacement (Gastil et al., 1981; Baird and Miesch, 1984; Taylor, 1986; Gromet and Silver, 1987; Silver and Chappell, 1988; Ague and Brimhall, 1988; Todd et al., 1988; Hill and Silver, 1988; Gastil, 1993; Johnson et al., 1999; Lovera et al., 1999; Todd et al., 2003; Kistler et al., 2003; Langenheim and Jachens, 2003). The oldest recognized intrusive rocks are gneissic S-type plutons that occupy the medial zone of the Peninsular Ranges batholith and yield Middle Jurassic emplacement ages (Todd and Shaw, 1985; Thomson and Girty, 1994; Shaw et al., 2003; Kistler et al., 2003). Cretaceous plutons as old as 140 Ma also occur within this batholith (Silver and Chappell, 1988; Alsleben et al., 2005; D.L. Kimbrough, personal observ.). These are cut by a regionally extensive dike swarm emplaced at ca. 130–120 Ma (Böhnel et al., 2002; D.L. Kimbrough, personal observ.). The well-developed western zone of the Peninsular Ranges batholith is composed mainly of 125–100 Ma gabbro to monzogranite plutons with primitive island arc geochemical affinities. The eastern zone of the batholith is defined by a belt of large-volume, 95 ± 3 Ma tonalite, trondjhemite, and low-K granodiorite (TTG) plutons (Gastil et al., 1975; Silver and Chappell, 1988; Walawender et al., 1990) that constitute the La Posta TTG suite (Tulloch and Kimbrough, 2003; Kimbrough and Grove, 2006). These Na- and Al-rich plutons have a deep, garnet-present, melt-source signature, as seen in high Sr, Ba, Sr/Y, and La/Yb. Field and thermobarometric data indicate emplacement at ~2–6 kbar pressures (Rothstein and Manning, 2003) followed by rapid Cenomanian–Turonian uplift and denudation at rates of ~1–2 mm/yr (Krummenacher et al., 1975; Lovera et al., 1999; Johnson et al., 1999; Kimbrough et al., 2001; Ortega Rivera, 2003; Grove et al., 2003b). A delayed late Campanian–Maastrichtian phase of uplift (e.g., Krummenacher et al., 1975) appears to be related to Laramide shallow subduction and removal of the deep crustal and lithospheric roots of the La Posta belt (Lovera et al., 1999; Grove et al., 2003a, b). SAMPLING AND METHODS Detrital Zircon U-Pb Age Measurements

of the Catalina Schist collected from Santa Catalina Island. Many of the samples examined were previously studied by Grove and Bebout (1995). We focused upon metasedimentary rocks from each of the major metamorphic units in order to obtain an upper bound upon the depositional age of their sedimentary protolith and to determine sediment provenance. In our U-Pb analysis of zircons from the Catalina Schist, we employed secondary ionization mass spectrometry (SIMS) methods using the UCLA Cameca ims 1270 ion microprobe. The extensive metamorphic recrystallization that affected zircons from the amphibolite and epidote-amphibolite units was significantly ameliorated by the high spatial resolution of the ion microprobe in conjunction with cathodoluminescence (CL) imagery (e.g., only ~1 nanogram of sputtered zircon required to yield a U-Pb age). The techniques are described in Grove et al. (2003b) with additional details included in the GSA Data Repository.1 We analyzed conventionally sectioned and polished zircons in epoxy mounts. The zircons were hand selected from heavy mineral concentrates produced from standard crushing, density, and magnetic methods. Further details are included in the Data Repository. Because the number of zircons measured from individual samples was typically small owing to the low yields realized during mineral separation, we pooled data from multiple samples to obtain statistically meaningful results for each of the major tectonic units of the Catalina Schist. As noted above, metamorphic zircon growth was a significant issue for some of the grains we examined, particularly for the metasediments from the amphibolite unit. Anczkiewicz et al. (2004) also described evidence for metamorphic zircon growth in reconnaissance U-Pb zircon measurements they reported from the Catalina Schist amphibolite unit. Although we selected our analysis sites on the basis of morphologic and optical criteria in CL imagery, a significant number of the spot analyses from zircons of the amphibolite and epidote-amphibolite units overlapped regions affected by metamorphic recrystallization. Metamorphic overgrowths on igneous zircons generally have Th/U <0.1; Kröner et al., 1994; Rubatto, 2002; Williams and Claesson, 1987). We have found a cutoff of Th/U = 0.1 to be empirically well supported by detrital zircon results from the Cenomanian and younger, pluton-derived strata that overlie the Peninsular Ranges batholith. In our sampling of these units (Mahoney et al., 2005), 1506 out of 1527, or 98.6% of the zircon U-Pb analyses, yielded Th/U values equal to or greater than 0.1 (Fig. 3A). The average Th/U value measured was 0.48 ± 0.41, with 3.3% of the analyses falling within the Th/U = 0.1–0.2 range. Thus a cutoff of Th/U = 0.1 is useful for identifying detrital igneous zircons adversely affected by metamorphic recrystallization. Accordingly, we have excluded all analyses with Th/U <0.1 from the summary plots or calculations presented below (Fig. 3B). Complete data tables of U-Pb age measurements, lithologic descriptions, and sample locations are available from the 1

A total of 645 U-Pb zircon ages were measured from 33 samples of amphibolite through sub-blueschist facies metagraywacke

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GSA Data Repository item 2008105 is available online at www.geosociety. org/pubs/ft2008.htm, or on request from [email protected] or Documents Secretary, GSA, P.O. Box 9140, Boulder, CO 80301, USA.

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Grove et al. GSA Data Repository. In this paper, quoted uncertainties are ±1σ errors unless otherwise specified. The U-Pb ages are generally 206Pb/238U values for <1 Ga zircons and 207Pb/206Pb ages for older grains. Analyses with age uncertainties >10% have been discarded. The U-Pb ages of zircons that exhibited resolvable 206 Pb/238U versus 207Pb/235U discordance are 207Pb/206Pb values. Rb-Sr Measurements To clarify the significance of previous 40Ar/39Ar phengite step-heating results (Grove and Bebout, 1995) that had indicated a possible Middle Jurassic age for garnet-bearing blueschist blocks within the lawsonite-blueschist mélange, we carried out Rb-Sr measurements with one of the blocks (330–4B) that had been analyzed in this previous study. All column chemistry and mass spectrometer measurements were carried out at UCLA. There, Rb and Sr were separated in cation exchange columns containing AG50W-X8 resin, using 2.5N HCl. The Sr isotopic compositions were measured with a VG54–30 multicollector thermal ionization mass spectrometer. 87Sr/86Sr ratios were normalized to 86Sr/88Sr = 0.1194. In this analysis session, the measured 87Sr/86Sr value for our Sr standard (NBS 987) value was 87Sr/86Sr = 0.710239 ± 16 (2σ, n = 13). The Rb and Sr concentrations were measured by isotope dilution using VG54–30. Model Rb-Sr isochron ages were calculated using ISOPLOT/Ex Version 3 (Ludwig, 2003). RESULTS Amphibolite Facies

Figure 3. (A) Measured zircon U-Pb age versus Th/U of Cretaceous forearc sedimentary rocks of the Peninsular Ranges batholith. Data, from Mahoney et al. (2005), demonstrate that virtually all detrital igneous zircons derived from the Peninsular Ranges have Th/U >0.1 (see also Williams and Claesson, 1987). (B) Equivalent plot for zircon results from the Catalina Schist amphibolite unit. Based upon cathodoluminescence imaging and other criteria, we regard a Th/U = 0.1 as a meaningful cutoff to distinguish analyses in which the sputter pit overlapped metamorphic zircon growth (open symbols). The filled symbols represent detrital grains that we consider to be largely unaffected by metamorphic zircon growth (see also Kröner et al., 1994; Rubatto, 2002). Zircon results with Th/U <0.1 have been excluded from all subsequent plots and calculations.

Five biotite + muscovite + garnet ± kyanite-bearing metagraywackes from the amphibolite unit yielded abundant subhedral to well-rounded zircon. Alternatively, both metachert samples examined contained only metamorphic zircon. Although external overgrowths <1–5 µm or more were common, zircons also exhibited internal areas of patchy diffuse recrystallization in CL imagery. Of the 169 analyses, 48 (28%) overlapped metamorphic zircon based upon Th/U (Fig. 3B). About half of the affected analyses yielded obviously mixed ages. U-Pb ages calculated for the remaining analyses with Th/U <0.1 fell between 107 and 126 Ma, with a peak at 116 ± 6 Ma. This result is in good agreement with independent estimates for the timing of amphibolite unit recrystallization (ca. 115 Ma; Suppe and Armstrong, 1972; Mattinson, 1986; Grove and Bebout, 1995; Anczkiewicz et al., 2004). The data for zircons with Th/U >0.1 clearly indicate that the sedimentary protolith of the amphibolite facies metagraywackes contained a large proportion of craton-derived detritus (Fig. 4A). Roughly 50% of the analyses were from Middle Proterozoic zircon (Fig. 4G). Unfortunately, many of these grains exhibited variable Pb loss and high degrees of discordance. This condition complicates characterization of the Middle Proterozoic age distribution. Although 207Pb/206Pb ages more accurately approximate the crystallization age than the highly discordant U-Pb ages, they

Figure 4. Relative probability plots of detrital zircon U-Pb age distributions from major units of the Catalina Schist on Catalina Island. (A) Amphibolite. (B) Epidote-amphibolite. (C) Epidote-blueschist. (D) Lawsonite-blueschist. (E) Actinolite-albite. (F) Lawsonite-albite. Note that we use a split horizontal axis at 300 Ma and that relative probability plots between 300 and 3000 Ma have a 2× scaling factor to improve resolution of the overall age distribution. (G) True scale cumulative probability spectra for all units.

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are insufficiently precise to reveal strong clustering of ages. Nevertheless, the subdued maxima that occur at 1.15, 1.40, and 1.65– 1.80 Ga are compatible with a southwestern North American provenance. Age peaks also occur at 126, 145, and 162 Ma. The five youngest zircons measured from the amphibolite unit metagraywackes yielded a weighted mean age of 122 ± 3 Ma. These analyses yielded a mean Th/U value of 0.30 ± 0.20 and were from oscillatory zoned regions in Cl− imagery that likely reflect igneous crystallization. Accordingly, we regard 122 ± 3 Ma as a geologically meaningful upper bound upon the depositional age of the sedimentary protolith. Epidote-Amphibolite Facies Five muscovite ± biotite ± garnet epidote-amphibolite facies metagraywackes yielded zircon that was generally more subhedral to euhedral grains than the well-rounded grains common in the amphibolite unit metagraywackes. CL imaging and Th/U indicate that metamorphic recrystallization of detrital grains was less common than in the amphibolite unit. Only 13 of 135 or 10% of analyses of the zircons from epidote-amphibolite facies metagraywackes had Th/U <0.1. The U-Pb age distribution indicates a diminished cratonal provenance relative to amphibolite facies metagraywackes; only ~20% of the zircons yielded Middle Proterozoic ages (Fig. 4G). Although fewer analyses are available, the overall age distribution of Middle Proterozoic zircons in epidote-amphibolite metagraywackes is broadly similar to that of the amphibolite facies rocks (Fig. 4B). The relative lack of Proterozoic grains is accompanied by a proportionate increase of <200 Ma grains. Distinct peaks are present at 115 and 126 Ma, with a strong peak at 150 Ma (Fig. 4B). The five youngest detrital zircons from the epidote-amphibolite metagraywackes yielded an age of 113 ± 3 Ma, with a mean Th/U value of 0.42 ± 0.05. Epidote-Blueschist Facies Thirty analyzable zircons were recovered from a single metagraywacke sample intercalated with Na-amphibole + clinozoisite + albite-bearing mafic rocks (glaucophanic greenschists of Platt, 1975; Sorensen, 1986). Six similar-appearing samples from equivalent field settings failed to yield zircon. This poor zircon yield probably reflects abundant volcanic detritus in the protolith. Nearly all of the zircons recovered consisted of clear subhedral to euhedral grains. There was no evidence for metamorphic zircon growth either petrographically or in terms of measured Th/U. Two distinct maxima occur at 103 and 142 Ma (Fig. 4C). Although results from the epidote-blueschist facies unit are transitional between those obtained from the amphibolite and epidote-amphibolite facies units and lower grade parts of the Catalina Schist (Fig. 4G), the epidote-blueschist facies unit appears more closely allied with the lower grade units, based upon the age of the youngest zircons detected in each population. The five youngest zircons measured yielded an average U-Pb age of 101 ± 3 Ma with a mean Th/U of 0.63 ± 0.23.

Lawsonite-Blueschist Facies and Lower Grade Rocks Data bearing upon the sedimentary protoliths of the lawsonite-blueschist and lawsonite-albite, and albite-actinolite facies units are discussed together because of the highly similar results produced (Fig. 4D–G). Overall, we obtained 164 analyses from 11 lawsonite-blueschist facies rocks, 59 analyses from 3 albite-actinolite facies rocks, and 82 analyses from 7 lawsonitealbite facies metagraywackes. No metamorphic zircon growth was detected; Th/U values were all >0.1. The Ca-rich metagraywackes prevalent in low-grade parts of the Catalina Schist overwhelmingly contain clear subhedral to euhedral grains. Most of these grains yield U-Pb ages of 95–130 Ma (Fig. 4D–F). Late Jurassic–Early Cretaceous zircon was greatly subordinate, and only trace quantities of Proterozoic zircon were present. Because the youngest zircon U-Pb ages measured from the three lowest grade units all agreed within error (96 ± 3 Ma, 97 ± 3 Ma, and 98 ± 3 Ma for lawsonite-blueschist, actinolite-albite, and lawsonitealbite, respectively), they define a 97 ± 3 Ma upper bound for the depositional age of the sedimentary protolith for all of the lowgrade Catalina Schist units. Garnet-Bearing Blueschist Block in Lawsonite-Blueschist Mélange Garnet-bearing blueschist blocks occur within mélange zones in the lawsonite-blueschist unit. Phengite 40Ar/39Ar results for two such blocks indicate that they formed prior to peak-grade recrystallization in the amphibolite unit (Grove and Bebout, 1995). To confirm this relationship we undertook Rb-Sr measurements of phengite, Na-amphibole, and whole rock for one of the samples (330–4B; Fig. 5A). Results are shown in Figure 5B. Phengite, Na-amphibole, and whole-rock results do not define a statistically meaningful isochron. However, the model Rb-Sr age defined by phengite + Na-amphibole, 135 Ma, is identical to the total gas age calculated from 40Ar/39Ar step-heating results obtained from 330 to 4B phengite. DISCUSSION Evaluation of the Nascent Subduction Model for the Catalina Schist The origin of the amphibolite facies rocks of the Catalina Schist is central to assessing the tectonic significance of this terrane. Amphibolite facies metamorphism and anatexis are atypical features of subduction complexes and are thought to reflect unusual processes along a convergent margin. High-temperature conditions accompany the earliest developmental stages of subduction before the hanging wall is refrigerated (Platt, 1975; Cloos, 1985; Peacock, 1987). Although ridge subduction and/or slow underflow of very young oceanic crust (e.g., Peacock, 1987, 1992; Hacker, 1990a, 1990b; Peacock et al., 1994) can also produce high-temperature metamorphism in subduction zones, there

The Catalina Schist

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Figure 5. (A) Garnet-bearing blueschist block in mélange matrix of the lawsonite blueschist unit. (B) Rb-Sr results for phengite, Na-amphibole, and whole rock. A statistically meaningful isochron is not defined. Phengite and Na-amphibole define a model isochron (135 Ma) that is identical to the total gas 40Ar/39Ar age yielded by this sample.

is no evidence that either process affected the middle Cretaceous southwestern margin of North America. The detrital zircon results (Fig. 6) require substantial modification of Platt’s (1975) hypothesis that amphibolite through blueschist facies rocks formed in an inverted metamorphic aureole in response to transient heating during nascent subduction at ca. 115 Ma. Sequential accretion of the major tectonic units of the

Catalina Schist over a ~20 m.y. interval is required. As indicated in Figure 6A, the protolith of the amphibolite unit metagraywackes was deposited between 122 Ma (youngest detrital zircon ages) and 115 Ma (the time of peak-grade recrystallization). In contrast, the sedimentary protolith of the lawsonite-blueschist and lower grade units was deposited after 97 Ma or at least 18 m.y. after peak-grade recrystallization of the amphibolite unit (Fig. 6C, D).

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Figure 6. Estimated temperature-time histories for major units of the Catalina Schist. (A) Amphibolite. (B) Epidoteamphibolite. (C) Epidote-blueschist. (D) Lawsonite-blueschist and lower grade units of the Catalina Schist. Yellow circles represent measured detrital zircon U-Pb ages. Maximum bounds upon the depositional age of the sedimentary protolith represent average and standard deviation of U-Pb ages of the five youngest zircons with Th/U >0.1 (15 youngest zircons for composite of the three lowest grade units). Data sources for other thermochronometry include Suppe and Armstrong (1972), Mattinson (1986), Grove and Bebout (1995), and Anczkiewicz et al. (2004). Note that we have varied the bulk closure temperatures assigned to micas from 400 °C for coarse-grained muscovites within the high-grade units to 350 °C for finer grained, less retentive white mica within lawsonite-albite rocks.

Available data for the Catalina Schist amphibolite unit indicate that it cooled slowly (ca. 25 °C/m.y.) from peak-grade conditions (650–700 °C) at 115 Ma to K-Ar muscovite closure temperatures (350–400 °C) at 105–100 Ma (Fig. 6A; Grove and Bebout, 1995). The Catalina epidote-amphibolite unit also cooled slowly (Fig. 6B): a 7–12 m.y. interval separated peakgrade recrystallization of epidote-amphibolite mafic gneisses and cooling through phengite K-Ar closure. Slow cooling exhibited by the amphibolite unit is compatible with very slow subduction or subduction of very young oceanic lithosphere. However, environments such as these are far too warm to permit formation of lawsonite-blueschist and lower temperature metamorphic rocks.

Within the Catalina Schist, only the lawsonite-blueschist and lower grade rocks record metamorphic conditions that require formation within a low-T, high-PT subduction zone. The amphibolite and epidote-amphibolite units were retrograded under greenschist facies conditions, as was the epidote-blueschist unit (Platt, 1976; Sorensen, 1986; Bebout and Barton, 1989; Grove and Bebout, 1995). Given their present association with lawsonite-blueschist and lower grade rocks, we find it remarkable that the amphibolite and epidote-amphibolite units exhibit such little evidence for equilibration at blueschist facies conditions (Platt, 1976; Sorensen, 1986; Bebout and Barton, 1989; Grove and Bebout, 1995). Blueschist facies overprinting of higher grade

The Catalina Schist rocks is the rule for subduction complexes of the North American Cordillera (Ernst, 1988). For example, tectonic blocks of garnet amphibolite in the Franciscan and Shuksan subduction complexes are heavily overprinted by blueschist facies assemblages (e.g., Brown et al., 1982; Cloos, 1985; Wakabayashi, 1990). In contrast, evidence for blueschist facies overprinting within the Catalina Schist amphibolite unit is obscure and limited to sparse veining by pumpellyite and trace amounts of lawsonite in white mica + zoisite + albite assemblages that replace oligoclase in amphibolite facies rocks (Grove and Bebout, 1995). We interpret the available thermochronology and greenschist facies overprinting of the high-grade units of the Catalina Schist to indicate that they occupied a subcrustal position characterized by 400–500 °C temperatures prior to 100–95 Ma. The lack of blueschist facies overprinting within the amphibolite and epidote-amphibolite units indicates to us that these rocks were sufficiently distant from an active subduction zone to be largely unaffected by the low geothermal gradient (<5–10 °C/km) recorded by the Franciscan and other Cordilleran subduction complexes during the middle Cretaceous. Provenance Ties Linking the Catalina Schist to the Peninsular Ranges Batholith King et al. (2007) measured whole-rock Pb isotopes from mélange matrix sampled from the amphibolite, lawsoniteblueschist, and lawsonite-albite units of the Catalina Schist. Although details of the U-Th-Pb systematics indicate the possible elemental redistributions during the devolatization history of the schist (see King et al., 2007), it is remarkable how well the measured Pb isotopic compositions of the sediment-dominated lawsonite-blueschist and lawsonite-albite mélange matrix agree with those measured from granitoids of the adjacent Peninsular Ranges batholith (Fig. 7A, B; Kistler et al., 2003). The wholerock Pb isotopic compositions from the mafic- and ultramaficdominated mélange of the Catalina Schist amphibolite unit also overlap strongly with the Peninsular Ranges batholith (Fig. 7C) but tend to more radiogenic values than the lower grade units (Fig. 7A, B). The amphibolite unit mélange results are more difficult to interpret than those from the lower grade units, because the sources of Pb are far less uncertain owing to the scarcity of sedimentary material within the mélange (see Bebout and Barton, 1993, 2002). Nevertheless, these results are consistent with a cratonal provenance similar to that exhibited by southeastern

Figure 7. Whole-rock Pb isotopic data of King et al. (2007) for: (A) Lawsonite-albite. (B) Lawsonite-blueschist. (C) Amphibolite unit mélange matrix of the Catalina Schist. Also shown are wholerock Pb isotopic data for northern Peninsular Ranges batholith plutons (red circles; Kistler et al., 2003) and equivalent data for southeastern Arizona (orange circles; Wooden and Miller, 1990). The yellow field represents the generalized distribution of whole-rock Pb yielded by the Mojave–Transverse Ranges province (based upon Barth et al., 1995, and personal observ. from J. Wooden, D. Coleman, and A. Barth).

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Arizona (Wooden and Miller, 1990) and can be differentiated from the Mojave–Transverse Ranges province (Fig. 7C; Barth et al., 1995; J. Wooden, D. Coleman, and A. Barth, personal observ.). Because the early Mesozoic wall rocks of the Peninsular Ranges batholith are dominated by craton-derived Proterozoic zircon that indicates a similar source (see below), we speculate that the Proterozoic zircon contributed from the cratonally derived sediment eroded from the batholith’s wall rocks recrystallized during mélange formation, liberating the highly radiogenic Pb that was dispersed throughout the amphibolite unit mélange as it was recrystallized and metasomatically altered (Bebout and Barton, 1993, 2002; King et al., 2007). In the same manner as the Pb isotope data discussed above, the detrital zircon results from the Catalina Schist also indicate a close genetic relationship between the Catalina Schist and the Peninsular Ranges batholith: (1) early Mesozoic, Paleozoic, and Proterozoic detrital zircon from the two terranes are similar; and (2) middle Cretaceous detrital zircon age distributions from the lawsonite-blueschist and lower grade Catalina Schist are virtually identical to the distribution of crystallization ages from the northern Peninsular Ranges batholith and the detrital age distribution from the adjacent Upper Cretaceous forearc rocks. Lower Mesozoic flysch wall rocks of the Peninsular Ranges batholith crop out within the deeply denuded axial zone of the northern batholith (Fig. 2) and represent a major potential source of early Mesozoic, Paleozoic, and Proterozoic zircon. A composite detrital zircon distribution based upon results obtained by Morgan et al. (2005) from three different localities in southern California (Bedford Canyon Formation, French Valley area, and Julian Schist; localities 1–3 in Fig. 2) and northern Baja California (Vallecitos area; locality 4 in Fig. 2) is shown in Figure 8. As indicated, the age distribution defined by >200 Ma detrital zircon from the Catalina Schist (Fig. 8A) is consistent with a source region similar to the Peninsular Ranges batholith wall rocks (Fig. 8B). The expected age distribution for detritus that originated from the southwest North American craton is represented by late Miocene to Holocene sands from the Colorado River system (Fig. 8C). Well-defined peaks at 1.0–1.2, 1.45, and 1.65–1.75 Ga are distinctive of the basement assemblage (Gehrels and Stewart, 1998). Early Paleozoic and latest Neoproterozoic (400–650 Ma) zircon was contributed by supracrustal cover rocks that had their source in the Appalachian and Ouachita orogenic belts (Dickinson and Gehrels, 2003). Finally, Permian–Triassic zircon is derived from the earliest Phanerozoic magmatic arcs established along the southwestern North American margin (Barth and Wooden, 2006; González-León et al., 2006). The lawsonite-blueschist and lower grade rocks of the Catalina Schist that contain Upper Cretaceous forearc strata of the Peninsular Ranges are very similar (Fig. 8E), as is the distribution of U-Pb zircon crystallization ages from the adjacent batholith (Fig. 8F; Silver and Chappell, 1988; Walawender et al., 1990; Kistler et al., 2003; D.L. Kimbrough, personal observ.). The fact that the Catalina Schist results are skewed to slightly older ages is

readily understood in terms of the Cenomanian age of the protolith. Cenomanian-age forearc strata from the Peninsular Ranges batholith are also skewed to older ages because they were deposited prior to massive exhumation of the ca. 95 Ma La Posta plutonic suite within the eastern batholith (Mahoney et al., 2005). Most of the results from the Upper Cretaceous forearc rocks are Campanian or Maastrichtian and hence are enriched in detritus from the eastern plutonic zone. In order to further evaluate the strength of the provenance tie between the Catalina Schist and the Peninsular Ranges, we have carried out ternary mixing calculations. We have identified three distinctive components: (1) lower Mesozoic wall rocks, represented by the Bedford Canyon, French Valley, Julian Schist, and Vallecitos flysch localities (Fig. 9A; Morgan et al., 2005); (2) Early Cretaceous volcanics, represented by results from Lower Cretaceous sandstones and volcanics within the volcanic arc (Fig. 9B; Alsleben et al., 2005; D.L. Kimbrough, personal observ.); and (3) Upper Cretaceous forearc, represented by Cenomanian–Maastrichtian strata distributed along the western margin of the Peninsular Ranges batholith (Fig. 9C; Mahoney et al., 2005). Cumulative probability plots for these three components are shown in Figure 9D, along with equivalent curves representing the major units of the Catalina Schist. We have linearly mixed these three components to obtain the best fit to the detrital zircon age distributions from the Catalina Schist. Results for the amphibolite unit, epidote-amphibolite unit, epidote-blueschist unit, and lawsonite-blueschist and lowergrade units are shown in Figure 9E through H, respectively. The relative proportions of the three end members required to produce these fits are shown in Figure 9I. As indicated in Figure 9I, the detrital zircon provenance signature of the earliest accreted material within the amphibolite unit of the Catalina Schist is best approximated by a 68:32 mixture of sediment derived predominantly from lower Mesozoic wall rocks and the Early Cretaceous volcanic arc, respectively. Such a high proportion of cratonally derived material within the amphibolite unit metasediments makes sense, given the abundance of Middle Proterozoic zircon, the aluminous and quartz-rich nature of the protolith, and the fact that pegmatites, metasediments, and mélange sampled from the amphibolite unit tend to yield a relatively radiogenic Sr, Nd, and Pb isotopic signature (i.e., similar to evolved arc crust; Bebout and Barton, 1993, 2002; King et al., 2006, 2007). The provenance signature of sediment shed from the Peninsular Ranges batholith shifted dramatically between early Aptian to early Cenomanian time (Mahoney et al., 2005). A parallel shift is recorded by the Catalina Schist (Fig. 4). In the case of the epidote-amphibolite unit, the relative proportions of lower Mesozoic wall rock to Early Cretaceous volcanic arc sediment diminishes to 44:56 (Fig. 9I). Compositionally, there is still a resolvable cratonal contribution in the epidote-amphibolite metagraywackes. A further shift is exhibited by the epidote-blueschist metagraywackes. Their age distribution is well modeled by a 20:58:22 mixture of lower Mesozoic wall rock, Lower Cretaceous volcanic arc, and Upper Cretaceous forearc sediment,

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Figure 8. Relative probability plots (200–2200 Ma) of detrital zircon U-Pb ages. (A) All results from the Catalina Schist. (B) Representative early Mesozoic flysch wall rocks of the northern Peninsular Ranges batholith (Morgan et al., 2005); data originate from localities 1–4 in Figure 2. (C) Representation of the southwestern North American detrital zircon provenance signature as represented by late Miocene–Holocene sediments of the Colorado River system (D. Kimbrough and M. Grove, personal observ.). Yellow bands denote the primary Middle Proterozoic crystallization age maxima contributed by cratonal basement and supracrustal rocks of southwestern North America. (D) All detrital zircon U-Pb age results from the lawsonite-blueschist and lower grade units of the Catalina Schist. (E) Representative Late Cretaceous (K) detrital zircon data from the forearc of the northern Peninsular Ranges batholith (PRB) (Mahoney et al., 2005). (F) Pluton U-Pb zircon crystallization ages from the northern Peninsular Ranges batholith (Silver and Chappell, 1988; Walawender et al., 1990; Kistler et al., 2003; D. Kimbrough, personal observ.).

respectively (Fig. 9I). The provenance shift was completed by the time the protolith of the lawsonite-blueschist and lower grade metagraywackes was deposited. The detrital zircon age distribution of the latter is well described by a 49:51 mixture of Lower Cretaceous volcanic arc and Upper Cretaceous forearc sediment, respectively (Fig. 9I). Implications of Pre-115 Ma High-PT Tectonic Blocks within Low-Grade Catalina Schist Grove and Bebout (1995) reported 40Ar/39Ar results for two garnet-bearing, blueschist blocks within the lawsonite-blueschist unit of the Catalina Schist that indicated that the blocks predated

115 Ma peak-grade metamorphism within the amphibolite unit. Because the pronounced age gradients yielded by both samples reached 150–160 Ma at the highest temperatures of gas release, Grove and Bebout (1995) speculated that the garnet-bearing blueschist blocks formed during the Middle Jurassic and had the same tectonic significance as similar high-grade blocks present within the Franciscan Complex along the margin to the north (Wakabayashi, and Dumitru, 2007, and references cited therein) and within the western Baja terrane along the margin to the south (Baldwin and Harrison, 1989, 1992). The concordance of the 135 Ma Rb-Sr and K-Ar model ages from sample 330–4B (Fig. 5) indicate that 40Ar/39Ar results from it cannot be explained by excess 40Ar contamination. Petrographic observations of

Figure 9. Relative probability plots (0–2000 Ma) of detrital zircon U-Pb age results used as end members in ternary mixing calculations: (A) Early Mesozoic flysch wall rocks of the northern Peninsular Ranges batholith (PRB) from Morgan et al. (2005). (B) Early Cretaceous volcanic sandstones (Alsleben et al., 2005) and volcanic rocks (D. Kimbrough, personal observ.) of the northern Peninsular Ranges batholith. (C) Late Cretaceous forearc strata of the northern Peninsular Ranges batholith (Mahoney et al., 2005). (D) Cumulative probability plots of the three end members. Best-fit ternary-mixing results: (E) Amphibolite unit. (F) Epidote-amphibolite unit. (G) Epidote-blueschist unit. (H) Lawsoniteblueschist (and lower grade) units of the Catalina Schist. (I) Ternary diagram showing best-fit solutions. Note the general trend away from Peninsular Ranges wall rock and volcanic zircon signature exhibited by the oldest accreted units toward a pluton-dominated zircon provenance that characterizes the youngest rocks within the Catalina Schist. Mz—Mesozoic; K—Cretaceous.

The Catalina Schist sample 330–4B indicate that the K-Ar and Rb-Sr model ages are most sensibly interpreted as mixed ages that resulted from partial replacement of 0.1–1 mm diameter Middle Jurassic phengite by much finer grained (<50 µm) middle Cretaceous phengite. This evidence for prior (Middle Jurassic?) subduction metamorphism along the margin is in good agreement with the geologic history of the Peninsular Ranges batholith. Conversely, pre–115 Ma subduction along the segment of the margin underplated by the Catalina Schist makes it difficult to explain how mantle lithosphere composing the overriding hanging wall could have retained sufficient heat to produce amphibolite facies metamorphism within a newly formed subduction zone at 115 Ma. Subduction Erosion Model for the Formation of the Catalina Schist The results and geologic relationships described above lead us to conclude that the amphibolite, epidote-amphibolite, and possibly the epidote-blueschist units of the Catalina Schist are unlikely to have formed within the same subduction zone in which the lawsonite-blueschist and lower grade rocks of the Catalina Schist were also accreted and metamorphosed. Whereas studies of plate kinematics involving the Farallon, Kula, and Pacific plates are not definitive, they indicate that continuous subduction of comparatively old oceanic crust likely prevailed along the southwestern North American margin throughout the middle Cretaceous (Engebretsen et al., 1986; Stock and Molnar, 1988). Accordingly, we favor a model in which the Catalina amphibolite and epidote-amphibolite units represent Early Cretaceous forearc strata and basement (Fig. 10A) that were underthrust and metamorphosed beneath a forearc thrust that was well separated from the deeper subduction megathrust (Fig. 10B). A modern analogue for paired megathrusts depicted in Figure 10B is imaged by deep seismic-reflection data from the northern Cascadia subduction zone (Calvert, 2004). Forearc thrusting on this scale is capable of displacing forearc rocks to subcrustal positions and is consequently an important manifestation of subduction erosion. In the Catalina Schist, high-PT amphibolite facies metamorphism and partial melting may have occurred when a part of the Early Cretaceous forearc (Fig. 10A) was underthrust to a position beneath the western margin of the northern Peninsular Ranges batholith between 122 and 115 Ma (Fig. 10B). This event coincided broadly with previously documented intra-arc thrusting within the adjacent Peninsular Ranges batholith (Gastil et al., 1981; Todd et al., 1988; Silver and Chappell, 1988; Thomson and Girty, 1994; Busby et al., 1998; Johnson et al., 1999: Schmidt et al., 2002; Wetmore et al., 2002; Schmidt and Paterson, 2002; Wetmore et al., 2003; Busby, 2004). Shortening of the forearc could have been triggered by lateral expansion related to batholith emplacement coupled with collapse of thin, hot backarc crust that had separated the Peninsular Ranges batholith from the southwestern craton margin in the Early Cretaceous. Underthrusting of the Early Cretaceous forearc rocks to a stalled position beneath the western margin of the magmatic arc

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(Fig. 10B) explains the heat source required to produce amphibolite facies metamorphism and anatexis within the Catalina amphibolite unit. It also explains why the amphibolite unit underwent protracted greenschist facies metamorphic conditions after peak metamorphism occurred at 115 Ma. Greenschist facies, rather than blueschist facies overprinting, is expected if the amphibolite unit was metamorphosed close to the magmatic arc and far removed from the subduction zone. Based upon detrital zircon results, the epidote-amphibolite unit was accreted after 113 ± 3 Ma or several million years after peak grade recrystallization in the overlying amphibolite unit (Fig. 6B). Accretion of the epidote-amphibolite unit created an imbricate thrust stack beneath the western margin of the Peninsular Ranges batholith (Fig. 10C). Peak grade recrystallization of the epidote-amphibolite unit occurred at 110–107 Ma on the basis of 40Ar/39Ar hornblende results from the mafic gneiss directly underlying the amphibolite unit (Fig. 6B). Just as in the case of the amphibolite unit, greenschist facies metamorphic conditions likely persisted within the epidote-amphibolite unit for up to 10 m.y. on the basis of 40Ar/30Ar phengite results from epidote-amphibolite unit metagraywackes that indicate that cooling below 400–350 °C was delayed until ca. 97 Ma (Fig. 6B). Only fragmentary constraints are available for the metamorphic evolution of the epidote-blueschist unit. Previous work had indicated that the rocks of the epidote-amphibolite unit and epidote-blueschist unit were contiguous (the Catalina greenschist unit of Platt, 1976) and shared a common greenschist facies overprint (Platt, 1976; Sorensen, 1986). Our detrital zircon results from a single sample of epidote-blueschist metasediment intercalated with clinozoisite-albite–bearing blueschists cast doubt upon the likelihood that the epidote-amphibolite and epidoteblueschist units were closely related prior to ca. 101 Ma. Results from this single sample indicate that the epidote-blueschist rocks were accreted after 101 ± 3 Ma or roughly 6 m.y. after peak-grade recrystallization of the epidote-amphibolite unit had occurred. The epidote-blueschist assemblages could reflect subcrustal accretion beneath the collapsing forearc at an intermediate position between the subduction zone and the arc. Greenschist facies overprinting of the epidote-blueschist assemblages resulted as the rocks were underthrust beneath the epidote-amphibolite and higher grade rocks of the Catalina Schist between 101 and 97 Ma. The lawsonite-blueschist and lower grade rocks of the Catalina Schist were accreted after 97 ± 3 Ma during a major Cenomanian pulse of synbatholith erosion (Fig. 10D; Kimbrough et al., 2001). Evidence that accretion took place within the subduction zone is provided by the lawsonite-blueschist and lower grade mineralogy and the fact that phengite 40Ar/39Ar ages and the youngest detrital zircon ages coincide, indicating that accretion, metamorphism, and cooling took place rapidly (Fig. 6D). Merging of the higher grade units of the Catalina Schist with the subduction complex likely occurred by ca. 95 Ma and triggered tectonic denudation via low-angle extensional faulting that attenuated the crust and established the presently observed contact relationships within the Catalina Schist (e.g., Platt, 1986).

(A)

(B)

(C)

(D)

(E)

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Figure 10. Subduction erosion model for the formation of the Catalina Schist. (A) Early Cretaceous geometry of the convergent margin of the northern Peninsular Ranges batholith. The backarc extension and sedimentation shown are inferred from geologic relationships within eastcentral Peninsular Ranges batholith and formerly adjacent mainland Mexico. Detrital zircon results from Early Cretaceous volcanic sediments (Alsleben et al., 2005) require close proximity to southwestern North America. (B) Initial underthrusting of Early Cretaceous forearc beneath northwestern Peninsular Ranges and amphibolite facies metamorphism and anatexis at 115 Ma. Amphibolite unit stalls beneath western batholith at a position well separated from active subduction zone. Evidence for intra-arc shortening at this time is detailed within the text. (C) Underthrusting of epidote-amphibolite unit during progressive subduction erosion of the Early Cretaceous forearc beneath northwestern Peninsular Ranges batholith. Shortening in the backarc region preconditions the crust for deep melting (>40 km) to generate La Posta magmatism at 95 Ma. (D) Accretion of the lawsonite-blueschist and lower grade units of the Catalina Schist at 95 Ma occurs concomitantly with a major pulse of denudation and erosion within the Peninsular Ranges batholith (e.g., Kimbrough et al., 2001). Deep, garnet-involved melting to generate La Posta magmatism is facilitated by overthickened arc-ophiolite basement in the former backarc, focused asthenospheric counterflow caused by initiation of flat subduction and eastern relocation of magmatic arc, and massive devolatilization of deep underplated Catalina Schist within the subduction channel. (E) Laramide flat subduction tectonically erodes deep crustal root beneath eastern Peninsular Ranges batholith between 80 and 65 Ma, causing renewed deep exhumation of the eastern part of this batholith (Grove et al., 2003b). Tr—Triassic; J—Jurassic; K—Cretaceous; Pz—Paleozoic; TTG—tonalite, trondjhemite, granodiorite.

Devolatilization and Fluid Flow within the Catalina Schist Recent studies (Bebout et al., 1999, 2007; Bebout, 2007) have discussed varying extents of devolatilization in the tectonometamorphic units of the Catalina Schist within the context of varying prograde P-T paths that took place in a subduction zone setting (also see discussion in Grove and Bebout, 1995). Whereas the subduction erosion model (Fig. 10) does not affect the general conclusions of this previous work as they pertain to devolatilization histories, the sources and pathways for infiltrating fluids that metasomatized the higher grade rocks of the Catalina Schist do require further consideration within the context of the subduction erosional model presented here. Bebout and Barton (1989) and Bebout (1991a, b) proposed that, in the amphibolite unit, the O isotope compositions in mélange and other more permeable zones 18 reflect infiltration by externally derived aqueous fluids with δ O inherited from prior equilibration with similar metasedimentary rocks but at lower temperatures (~300–600 °C). The source of the fluids that ascended into the underplated (but still hot) amphibolite unit was primarily related to 400–550 °C chlorite breakdown reactions in rocks equivalent to those within the lower grade units of the Catalina Schist (Bebout, 1991a, b; also see discussion in Grove and Bebout, 1995). In the subduction erosion model (Fig. 10), fluids with appropriate lower T metasedimentary signatures derived from the subducting slab are required to traverse a significant thickness (tens of kilometers) of mantle lithosphere in order to infiltrate the amphibolite unit at near peak-grade metamorphic conditions. We suggest that this could have occurred without substantial 18 modification of the δ O of the fluids. Devolatilization of subducting oceanic crust and trench fill is considered capable of hydrating mantle lithosphere over broad regions (Peacock, 1993; Humphreys et al., 2003). Infiltration of aqueous fluids derived from the subduction zone likely began during, or prior to, the Middle Jurassic. This previous history was likely sufficient to thoroughly hydrate and isotopically re-equilibrate the intervening lithospheric mantle with lower temperature, slab-derived fluids prior to the onset of Early Cretaceous subduction erosion of

the forearc. Pervasive hydration of mantle lithosphere underlying the forearc region is consistent with observations, in modern forearcs, of extensive hydration in the hanging wall leading to distinctive seismic-velocity signatures (Bostock et al., 2002; Brocher et al., 2003) and forearc serpentinite diapirs (Fryer et al., 1995). As continued subduction erosion brought the loci of forearc thrusting nearer to the active subduction thrust system (see Fig. 10C, D, E), progressively lower T rocks, such as those in the epidote-amphibolite and epidote-blueschist rocks, would similarly have been infiltrated by these fluids emanating from the still lower T (lawsonite-blueschist) domains produced in the active subduction zone thrust. Relationship to Emplacement of the La Posta TonaliteTrondjhemite-Granodiorite Suite At ca. 100 Ma the locus of Peninsular Ranges batholith magmatism shifted abruptly eastward in conjunction with intrusion of the voluminous La Posta tonalite-trondjhemite-granodiorite plutonic suite (Fig. 2; Gastil et al., 1975; Silver and Chappell, 1988; Walawender et al., 1990; Tulloch and Kimbrough, 2003; Kimbrough and Grove, 2006). The La Posta suite emplacement involved a sustained (ca. 98–92 Ma) magmatic flux of >100 km3/ m.y./km strike length over >1200 km (Kimbrough and Grove, 2006). The high Sr, Ba, Sr/Y, Na2O, Al2O3, and highly fractionated rare-earth-element patterns exhibited by these plutons indicate deep, garnet-involved melting of a fundamentally mafic source region (Gromet and Silver, 1987). Oxygen and Rb-Sr isotopic measurements reported by Taylor (1986), Gromet and Silver (1987), Silver and Chappell (1988), Hill and Silver (1988), and Kistler et al. (2003) have revealed that the northern La Posta belt features elevated δ18O (9–12‰ whole-rock equivalent) at intermediate initial 87Sr/86Sr values (typically 0.704–0.708). The supracrustal contribution implied by these combined isotopic attributes cannot be accounted for by high-level assimilation of highly radiogenic, early Mesozoic, cratonally derived flysch host rocks (Shaw et al., 2003). These characteristics are more readily explained by partial melting and/or devolatilization of isotopically

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primitive low-grade metasedimentary and metavolcanic rocks and altered oceanic crust within the deep source region (Taylor, 1986; Gromet and Silver, 1987). Emplacement of the La Posta belt plutons involved an eastern relocation of the locus of magmatism within the Peninsular Ranges over a strike length of 1200 km (Fig. 10D; Kimbrough and Grove, 2006). This event was the initial stage of an eastern sweep of magmatism into northern Mexico that occurred during the Late Cretaceous–early Tertiary (e.g., Silver and Chappell, 1988; McDowell et al., 2001; Staude and Barton, 2001; Henry et al., 2003). Subduction erosion (Fig. 10D) could have delivered the Catalina Schist into close proximity to the La Posta source region between 100 and 95 Ma. The coincidence of crustal thickening of the former backarc extensional basin (compare Fig. 10A with 10D) focused asthenospheric corner flow (Fig. 10D), and massive devolatilization of the Catalina Schist (Fig. 10D) may have set up the optimal conditions required to trigger the La Posta TTG flare-up (Kimbrough and Grove, 2006). Recent δ18O measurements performed by Kimbrough and Grove (2006) indicate that the La Posta plutons have high δ18O only in the northern segment of the batholith, where it might have been underplated by the Catalina Schist (Fig. 2). This is a good indication that Catalina Schist subduction erosion was sufficiently important to modify the source region characteristics of the La Posta belt magmatism. An important implication is that substantial fractions of subducted forearc supracrustal material can be recycled via subduction erosion processes and incorporated into arc batholith magmas over a short (<10 Ma) time interval. Relationship to Laramide Underthrusting The accretion of the lowest grade units of the Catalina Schist beneath the western Peninsular Ranges and the eastern relocation of the La Posta plutonic belt to the eastern Peninsular Ranges at 95 Ma can be viewed as important precursors to the Laramide craton-ward shift of arc magmatism and contractional deformation (Coney and Reynolds, 1977; Dickinson and Snyder, 1978). During the Laramide episode of Late Cretaceous–early Tertiary deformation, large tracts of southern California and southwestern Arizona were underplated by the high-PT Pelona Schist and related schists, with no intervening lithospheric mantle preserved (Fig. 10E; Ehlig, 1968, 1981; Crowell, 1968, 1981; Yeats, 1968; Haxel and Dillon, 1978; Haxel et al., 2002; Burchfiel and Davis, 1981; Jacobson, 1983, 1990; Jacobson et al., 1988, 2002, 2007; Dillon et al., 1990; Malin et al., 1995; Wood and Saleeby, 1997; Saleeby, 2003; Grove et al., 2003a). Late Cretaceous eclogitic xenoliths have been recovered from kimberlite pipes as far east as northeastern Arizona (Usui et al., 2003). The earliest recognized accretion of the distinctive Pelona and related schists occurred at 91 ± 1 Ma along the southwesternmost tip of the Sierra Nevada batholith (Rand Schist within the San Emigdio Mountains; Saleeby, 2003; Grove et al., 2003a). Lawsonite-blueschist and the lower grade Catalina Schist represent nearly equivalent material underplated at a slightly older

time (i.e., between 97 and 95 Ma) beneath the northern Peninsular Ranges (Fig. 6B). Thermochronology and detrital zircon results obtained from the NW-SE–trending belt of schist exposures have revealed that schist underplating associated with the Laramide event was widespread by 80–70 Ma beneath the middle Cretaceous arc (Jacobson, 1990; Jacobson et al., 2000; Barth et al., 2003; Grove et al., 2003a). By 70–60 Ma, cratonal rocks were being underplated by schist at positions well east of the medial Cretaceous arc (Grove et al., 2003a; Usui et al., 2003; Jacobson et al., 2007). Laramide shallow subduction processes tectonically removed the deep lithospheric mantle roots of the La Posta belt within the northern Peninsular Ranges batholith between 80 and 65 Ma (Fig. 10E). Receiver function seismic studies indicate that the deep crust and lithospheric mantle roots no longer exist beneath or within the northeastern Peninsular Ranges batholith (Ichinose et al., 1996; Lewis et al., 2001). Whereas this batholith exhibited predominantly syn- to late-batholithic cooling in its southern extent (>85 Ma; e.g., Ortega Rivera, 2003), the northeastern segment of the batholith was further characterized by a delayed and very significant pulse of rapid cooling between 80 and 65 Ma (Krummenacher et al., 1975; George and Dokka, 1994; Lovera et al., 1999; Grove et al., 2003b). Lovera et al. (1999) and Grove et al. (2003b) attributed this cooling to denudation related to the removal of lithospheric mantle and the underplating of schist during Laramide shallow subduction. Holk et al. (2006) reported that Laramide-age deformation of the northern Peninsular Ranges was associated with infiltration of high δD and δ18O fluids that are most readily explained by devolatilization of subducted oceanic crust or underplated volcanogenic sediments. The impact of the Laramide orogeny on the northern Peninsular Ranges appears to have contrasted significantly with its effect upon the southern Sierra Nevada batholith, where an inflection in the Laramide subduction zone (Pickett and Saleeby, 1993; Malin et al., 1995; Saleeby, 2003) apparently allowed the deep crust and upper mantle to be preserved until much more recently (Ducea and Saleeby, 1996, 1998; Zandt et al., 2004). CONCLUSIONS 1. Major tectonometamorphic units of the Catalina Schist were successively accreted over a 15–20 m.y. interval, beginning with the amphibolite unit at ca. 120–115 Ma and concluding with the lawsonite-blueschist and lower grade lithologies by 97–95 Ma. 2. The amphibolite unit resided at high temperatures for a protracted period of time (10–15 m.y.), whereas the lawsonite-blueschist and lower grade units that most likely formed near or within the subduction zone were deposited, accreted, metamorphosed, and cooled over a time interval too brief to be resolved by the methods employed in this study (<3 m.y.). 3. The provenance of the Catalina Schist metasedimentary rocks shifted as a function of time of accretion (and now,

The Catalina Schist metamorphic grade). Metagraywackes from the earliest accreted amphibolite unit were derived from an early Aptian sediment that apparently originated from erosion of Late Triassic–Jurassic flysch wall rocks and Early Cretaceous volcanics of the Peninsular Ranges batholith. Successively younger accreted materials became enriched with Early Cretaceous plutonic zircon from the Peninsular Ranges. The last accreted lawsonite-blueschist and lower grade rocks were derived from Turonian sediment with a detrital zircon provenance virtually identical to similarly aged sediment within the Peninsular Ranges forearc. 4. The Catalina Schist likely does not represent a synchronously formed, inverted metamorphic aureole arising from nascent subduction. The highest grade portions of the complex appear to have formed by a subduction erosion process in which parts of the forearc were underthrust beneath the western margin of the Peninsular Ranges batholith at ca. 122–115 Ma. Progressive subduction erosion of the forearc by continued underthrusting in the forearc region ultimately juxtaposed the higher grade units of the Catalina Schist with the subduction complex by 97–95 Ma. Only the lawsonite-blueschist and lower grade rocks are considered to have originated along the dominant subduction zone thrust within a high-PT thermal regime characteristic of such settings. 5. The accretion of the Catalina Schist marked the initial stage of shallowing subduction and inboard migration of magmatism and sediment underplating that culminated in the Late Cretaceous–early Tertiary Laramide orogeny. The sedimentary protolith of the youngest Catalina schist units are nearly identical in age and provenance to the oldest representative of the eugeoclinal schists emplaced beneath the southwesternmost Sierra Nevada batholith at ca. 92 Ma. ACKNOWLEDGMENTS We gratefully acknowledge discussions with Pat Abbott, Suzanne Baldwin, Andy Barth, Mark Barton, Ned Brown, Cathy Busby, Mark Cloos, Mihai Ducea, Ray Ingersoll, Peter Lonsdale, John Platt, Jason Saleeby, Dave Scholl, Sorena Sorensen, John Wakabayashi, and Jim Wright that helped shape ideas presented in this paper. We particularly thank Bob Bohannon, Mark Cloos, John Platt, Sorena Sorensen, and John Wakabayashi, whose reviews of previous versions of the manuscript led to significant improvements. Work carried out by Brian Altheim and Cathy Christoffel, as parts of their M.S. research at Lehigh University, provided important information on lowgrade metamorphism within the Catalina Schist that helped guide our sampling. Helge Alsleben, Brian Mahoney, and J.R. Morgan provided permission to use detrital zircon U-Pb age data from the Peninsular Ranges batholith to compare with our results from the Catalina Schist. Joe Wooden and Drew Coleman shared unpublished whole-rock Pb isotopic data from the

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Printed in the USA

The Geological Society of America Special Paper 436 2008

Sedimentary response to arc-continent collision, Permian, southern Mongolia C.L. Johnson* Geology and Geophysics, WBB 609, University of Utah, 135 South 1460 East, Salt Lake City, Utah 84112, USA J.A. Amory Salym Petroleum Development, Griboedova Street #2, 625000 Tyumen, Russia D. Zinniker Geology and Geophysics, Yale University, P.O. Box 2080109, New Haven, Connecticut 06520, USA M.A. Lamb Geology Department, OWS 153, University of St. Thomas, 2115 Summit Avenue, St. Paul, Minnesota 55105, USA S.A. Graham Geological and Environmental Sciences Department, Stanford University, 450 Serra Mall (Building 320) #118, Stanford, California 94305, USA M. Affolter Geology and Geophysics, WBB 609, University of Utah, 135 South 1460 East, Salt Lake City, Utah 84112, USA G. Badarch† Institute of Geology and Mineral Resources, Mongolian Academy of Sciences, Ulaanbaatar 210351, Mongolia

ABSTRACT The Eurasian Tien Shan–Yin Shan suture is a ~3000-km-long boundary between Paleozoic arc and accretionary complexes (the Altaids) and Precambrian microcontinental blocks (Tarim and North China block). Stratigraphic data are presented from localities in southern Mongolia spanning >800 km along the northern margin of the suture. Facies descriptions, climatic indicators, sandstone provenance, and paleocurrent data help reconstruct Permian basin evolution during and following arc-continent collision, and results are integrated with previously published data to create a preliminary regional synthesis. Upper Permian strata of southern Mongolia comprise fluvial successions in the southwest and marine turbidite deposits in the southeast. Floral assemblages show mixing of Siberian craton and North China block communities, indicating their close proximity to Mongolia by Permian time. There was a rapid transition from humid environments in the Late Permian to more arid

*[email protected] † Deceased. Johnson, C.L., Amory, J.A., Zinniker, D., Lamb, M.A., Graham, S.A., Affolter, M., and Badarch, G., 2008, Sedimentary response to arc-continent collision, Permian, southern Mongolia, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 363–390, doi: 10.1130/2008.2436(16). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Johnson et al. conditions in the Early Triassic, corresponding to the global Permian–Triassic boundary event, but which may also reflect more local driving mechanisms such as rain shadow effects. Permian sandstones from Mongolia have undissected to dissected arc provenance, with little contribution from continental or recycled orogen sources. The timing of the nonmarine-marine facies transition and cessation of arc magmatism broadly supports earlier collision along the western part of the suture zone than along the eastern part (e.g., Late Carboniferous–Late Permian). However, when regional geologic constraints are integrated, a more complex model involving differential rotation of Tarim and the North China block is preferred. Late Paleozoic rocks of southern Mongolia have been subsequently dismembered along Mesozoic–Cenozoic strike-slip faults, and thus also represent the long-term record of intracontinental deformation within accreted, heterogeneous crust. Keywords: collisional tectonics, arc accretion, palynology, sedimentary basins, Mongolia.

INTRODUCTION Continental growth via accretion is well illustrated by the Cenozoic India-Eurasia collision (Harrison et al., 1992; Rowley, 1996). This phenomenon broadly represents the entire Phanerozoic history of the region, as demonstrated by a series of generally southward-younging suture zones extending from the Siberian craton through Mongolia and China (Zhang et al., 1984; Fig. 1). Such rapid continental growth not only highlights an exceptional protracted record of terrane accretion but also illustrates subsequent intracontinental deformation along inherited zones of crustal weakness, even far inboard of the active margins (Tapponnier et al., 1982; Burchfiel et al., 1989). Multiple studies highlight

the timing, mechanisms, and sedimentary record of Mesozoic and younger collisional events in this region (Graham et al., 1993; Graham, 1996; Yin and Nie, 1996). However, the suture zone that formed the main nucleation point for Mesozoic accretion remains poorly understood; specifically, the boundary between Paleozoic arc terranes that lie south of the Siberian craton and tectonic blocks in China that are floored mainly by Precambrian continental crust (e.g., North China block and Tarim–Central Tien Shan blocks; Watson et al., 1987). This accretionary complex (the Altaid tectonic collage or “Altaids” after S¸engör and Natal’in, 1996, Fig. 1; cf. the Central Asian Orogenic Belt of Jahn, 1999) is well exposed in the study area of southern Mongolia. As such, this region constitutes a significant portion of one of the largest zones

Figure 1. Map of major tectonic units, boundaries, and fault zones in eastcentral Asia. ATF—Altyn Tagh fault; BS—Bogda Shan; CTS—Central Tien Shan; EGFZ—East Gobi fault zone; JH—Junggar-Hegen suture; NTS— North Tien Shan; QFS—Qilian fault system; SIB—Siberian craton; THB—Turpan-Hami basin; TLF—Tan Lu fault; UB—Ulaanbaatar (capital of Mongolia).

Sedimentary response to arc-continent collision of arc accretion on earth (S¸engör et al., 1993), so there is much to learn about arc-arc collision as well as terminal arc-continent collision, the subject of this paper. Few integrated studies of Permian rocks in southern Mongolia are available, so this study is a first step toward documenting that record. The boundary between the Altaids and Precambrian microcontinental blocks of north China has been variously named the Suolon, Solonker, Junggar-Hegen, South Mongolia–Hinggan, and Tien Shan–Yin Shan sutures (S¸engör et al., 1993; Wang and Mo, 1995; Badarch et al., 2002; Cope et al., 2005; Shi, 2006). The term Tien Shan–Yin Shan suture is favored because it implies a regional continuation of the collision zone from northwest to northeast China (Fig. 1; Yin and Nie, 1996; Xiao et al., 2004), a correlation that is supported by this study. The geology of this region is poorly defined on many levels beyond these problematic naming conventions. The latitude and extent of the suture zone relative to the China-Mongolia political border is debated (Xiao et al., 2003). Assuming that collision involved a combined Tarim–North China block moving north relative to the Altaids (this assumption is discussed more fully below; Enkin et al., 1992; Ziegler et al., 1996), the Tien Shan–Yin Shan suture spans >3000 km from northwest to northeast China, making it one of the longest and oldest tectonic boundaries within the amalgamated crustal material south of the Siberian craton. Formation of this suture during closure of the Paleo-Asian ocean (Dobretsov et al., 1994) is generally thought to have been complete by Late Permian to Early Triassic time, but alternate interpretations range from Devonian to Cretaceous (reviewed in Xiao et al., 2003). Timing relationships across the region have also been interpreted as either broadly synchronous or highly diachronous collision from west to east (Amory, 1996; Yin and Nie, 1996). Although many studies provide important constraints for parts of the suture zone (particularly in northwest China; Coleman, 1989; Windley et al., 1990; Laurent-Charvet et al., 2003), in general the paleogeography of the entire closing ocean basin is poorly understood in terms of polarity of subduction, and the number and nature of magmatic arcs that may have been active (Lamb and Badarch, 2001; Zhou et al., 2001; Xiao et al., 2004). Without a better understanding of the location, timing, and mechanisms associated with this event, it is difficult to evaluate implications such as whether postorogenic collapse could have been active in the Permian (Tang, 1990) or later into the Triassic (Davis et al., 2004), whether later intracontinental deformation has been concentrated along major crustal boundaries, and whether regional climate change in the Permian of China can be attributed to a rain shadow effect created by collisional orogens (Cope et al., 2005). This study reviews known, inferred, and newly constrained relationships relevant to the tectonic and stratigraphic history on either side of the nominal collision zone in order to extract a new model regarding its formation. Reconstruction of the Tien Shan– Yin Shan suture is aided by decades of research in northern China. We first review these studies, with an emphasis on available basin analysis constraints, to provide a regional tectonic context and to

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highlight outstanding questions related to regional tectonic evolution. Complementary data are then presented from seven main localities spanning >800 km from southwestern to southeastern Mongolia (Table 1; Fig. 2). New data and observations are derived primarily from Upper Permian accretionary and postaccretionary sedimentary complexes. Considered in the regional context, these data help constrain the timing, mechanisms, and implications of this early phase of continental growth in Asia. Similarly, the arc setting of southern Mongolia in the middle to late Paleozoic has been compared to the modern southwest Pacific Ocean (Lamb and Badarch, 2001), and additional field constraints permit evaluating this analogy and various models that describe the stratigraphic record of interarc and arc-continent collisions (Ingersoll et al., 1995; Huang et al., 2006). LATE PALEOZOIC SEDIMENTATION AND TECTONICS OF NORTH CHINA The southern margin of the Altaids is inferred to follow either a strongly arcuate shape roughly parallel to the modern political border, or alternatively a more continuous east-west– trending boundary across northern China (Fig. 1, JunggarHegen versus Tien Shan–Yin Shan sutures; Zhang et al., 1984; Yin and Nie, 1996). Tarim and the North China block represent continental margins with Precambrian crust, forming distinctive southern boundaries to the accretionary complex. In contrast, the Altaid tectonic collage primarily comprises late Paleozoic subduction-accretion complexes with probably only small and isolated microcontinent blocks (S¸engör et al., 1993). Confusion over the location of the suture that marks final closure of the Paleo-Asian ocean is exacerbated by this rather diffuse northern margin, which consists of several magmatic arc and ophiolite zones, and thus potentially multiple terrane boundaries. Northwest China Tectonostratigraphy Tarim and the North China block bound the Altaids south of Mongolia (Fig. 1). These two microcontinents are inferred by many to have amalgamated prior to the Ordovician (Yin and Nie, 1996; Heubeck, 2001). Similar middle to upper Paleozoic

TABLE 1. LATITUDE AND LONGITUDE COORDINATES FOR KEY STUDY AREAS* Shin Jinst Noyon Uul Tavan Tolgoy Naftgar Uul Tsaagan Tolgoy Nomgon Hovsgol *See Figure 2.

N 44° 30', E 99° 18' N 43° 12', E 101° 54' N 43° 09', E 102° 04' N 43° 33', E 105° 31' N 43° 14', E 105° 12' N 42° 53', E 105° 27' N 42° 57', E 108° 20' N 42° 46', E 108° 29' N 43° 35', E 109° 31' N 43° 37', E 109° 47'

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Johnson et al. Faults associated with the East Gobi fault zone Mesozoic-Cenozoic rocks, undifferentiated

Shin Jinst

Permian sedimentary and volcanic rocks Devonian-Carboniferous arc sequences, undifferentiated

Triassic sedimentary rocks

44º N

Ordovician-Silurian sequences, undifferentiated

Permian granite

Rocks mapped as Precambrian, including carbonate klippen and possible Mesozoic tectonites

Tavan Tolgoy Noyon Uul

s Ea

Naftgar Uul

tG

o

b bi

as

in

Hovsgol

Nomgon Tsaagan Tolgoy

China border-Mongolia

Onch Hayrhan

Southwest localities

105º E

South-central localities

108º E

Bulgan Uul

102º E

99º E

42º N

N

0

50

100 km

Southeast localities

Figure 2. Geologic map of southern Mongolia, showing main study areas. This “timeslice” map is generalized and modified from Tomurtogoo (1999) and Yanshin (1989) and highlights Triassic and older rocks most relevant to this study. See Table 1 for latitude and longitude coordinate locations of the study areas.

stratigraphy on either side of the Qilian Shan supports this early assembly (Yang et al., 2001; Ritts et al., 2004), and no clear Phanerozoic suture has been found between Tarim and the North China block. On the contrary, paleomagnetic data indicate separate apparent polar wander paths throughout the late Paleozoic, and such reconstructions imply that the North Tien Shan and Yin Shan segments of the suture zone formed at about the same time (Carboniferous–Permian) but at very different latitudes (McElhinny et al., 1981; Li et al., 1988; Enkin et al., 1992; Zhao et al., 1996). Floral distributions generally contradict this interpretation, instead indicating that Tarim and the North China block were at similar latitudes in the Late Permian (Ziegler et al., 1996). Paleozoic paleomagnetic poles from Tarim may be overprinted, not fully reconciled with Mesozoic–Cenozoic shortening and vertical axis rotations, and/or compared with problematic Eurasian reference poles (Gilder et al., 1996; Roger et al., 2003). With respect to late Paleozoic collision with the Altaids, it seems that Tarim and the North China block can be treated as constituting a northward-migrating Andean-style margin, but it is important to note that the paleogeography is poorly known and controversial. Amalgamation of northwest China occurred in two main phases, beginning with collision of Tarim and the Central Tien Shan by Early Carboniferous time (Fig. 1; Carroll et al., 1990; Windley et al., 1990; Allen et al., 1993; Zhou et al., 2001; Wang et al., 2005). The second phase of amalgamation in northwest China is inferred to represent the southern boundary of the Altaids in northwest China and thus is of primary interest for this study. Alternate interpretations place the boundary to the north of the Junggar basin (e.g., Junggar-Hegen suture, Fig. 1; Zhang et al.,

1984; Watson et al., 1987). Although crustal thickness throughout northwest China ranges up to 50 km, this is mainly attributed to post-Paleozoic sedimentary basin deposits and Cenozoic crustal thickening owing to India-Asia collision (Wang et al., 2003). No Precambrian rocks have been found surrounding the Junggar basin, in contrast to widespread Paleozoic ophiolite complexes and mafic igneous rocks that are present (Ren et al., 1987; Coleman, 1989; Wang et al., 2003). Thus the Junggar basin does not seem to be underpinned by Precambrian continental crust on the basis of multiple lines of evidence, and the area is better grouped into Altaid arc affinities (S¸ engör et al., 1993). Collision between the combined Tarim–Central Tien Shan and the North Tien Shan–Bogda Shan Arcs occurred between Late Carboniferous and Early Permian time (ca. 300–280 Ma), based on radiometric dating of arc-related igneous rocks (Fig. 1; Yin and Nie, 1996; Dumitru et al., 2001; Zhou et al., 2001). Sedimentary basin analyses provide further constraints on formation of the North Tien Shan. The youngest marine rocks directly north of Tarim (in the Junggar basin) are Upper Carboniferous–Lower Permian turbidites and shallow-marine sandstone, mudstone, and carbonates (Carroll et al., 1995). Overlying Permian deposits are nonmarine, with at least 3000 m of upward-coarsening successions ranging from lacustrine mudstone to coarse alluvial-fluvial conglomerate (Carroll et al., 1990). Apparent climatic shifts accompany and postdate collision in the Junggar and Turpan basins (connected in the Permian– Triassic, Hendrix et al., 1992; Carroll et al., 1995). There is a Late Permian transition from organic-rich, wood- and coal-bearing fluvial and lacustrine rocks associated with relatively humid climates, to braided fluvial red beds with common desiccation

Sedimentary response to arc-continent collision features (Carroll, 1991; Wartes et al., 2002). Arid conditions persisted from the latest Permian to the Early Triassic, whereas coal-bearing, wet-climate rocks were once again widespread in the Middle to Late Triassic (Hendrix et al., 1992; Greene et al., 2001). Along with paleocurrent and provenance data, these lithologic and climatic changes have been attributed to syndepositional deformation during uplift of the North Tien Shan in the Early Permian and continued deformation and partitioning of the flexural Junggar-Turpan-Hami basin from the Late Permian into the Mesozoic (Hendrix, 2000; Greene et al., 2001; Vincent and Allen, 2001). Carroll et al. (1992) also linked evidence for saline versus fresh-water depositional conditions in Permian lacustrine oil shale to an orographic effect, driven by increased precipitation runoff adjacent to the North Tien Shan. The Junggar basin likely originated in an arc setting (back-, intra- or interarc), and by the late Paleozoic it was a trapped ocean basin, evolving into a flexural foreland with ongoing collisional deformation in the Mesozoic (Allen et al., 1991; Carroll et al., 1990; Vincent and Allen, 2001). Sediment transport in the Permian Junggar trapped ocean basin is inferred to have been toward the west-northwest, as indicated by regional paleocurrent data (Carroll et al., 1995). This may reflect closure of a wedge-shaped ocean that opened toward the west, perhaps accommodated by clockwise rotation of Tarim and the Central Tien Shan during this time. This interpretation is consistent with published remnant ocean-basin models, where sediment is dispersed outboard of the migrating suture zone and mainly parallel to colliding margins (Ingersoll et al., 1995). However, it may contradict the broader interpretation of the entire Tien Shan–Yin Shan suture having closed progressively from west to east (Nie et al., 1990; Mueller et al., 1991; Amory et al., 1994). Other geologic evidence for Permian block rotation in northwest China includes dextral strike-slip shear zones in the North Tien Shan with synkinematic biotite 40Ar/39Ar ages of 290–245 Ma (Laurent-Charvet et al., 2003). There is also paleomagnetic evidence for continued movement of Tarim with respect to Siberia since the Late Permian (Li et al., 1988; Gilder et al., 1996), so it seems permissible that rotation could have begun during late Paleozoic collision, coincident with dextral shear in the North Tien Shan. Northeast China Tectonostratigraphy The broader interpretation of west-to-east closure across the entire Tien Shan–Yin Shan suture zone is based mainly on evidence for younger collision along the Yin Shan belt (northeast China) ~2000 km east of the North Tien Shan (Fig. 1; Amory et al., 1994). Cope et al. (2005) interpreted active arc magmatism from ca. 400–275 Ma via detrital zircon ages from Permian sandstones. This is further constrained by a geochemical shift to inferred postcollisional granites younger than ca. 257 Ma (Chen et al., 2000). Dissected continental arcs are found south of the Mongolian border in the Yin Shan highlands, indicating south-directed subduction in the region through Late Permian time. Alternatively, Xiao et al. (2003) define both a northern and southern boundary to the

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Solonker (e.g., Yin Shan) suture in northeast China, with dualpolarity subduction beneath two continental margins (the North China block to the south and a Devonian continental margin to the north), followed by Himalayan-style, final-stage continental collision of these margins at the end of the Permian. Although few paleomagnetic data are available for Carboniferous–Permian rocks of the North China block (cf. Huang et al., 2001), Mesozoic paleogeographic reconstructions suggest that Mongolian arc terranes had accreted to the North China block by the Late Permian, and as a combined block they rotated counterclockwise relative to the Siberian craton from the Triassic to the Cretaceous during closure of the Mongol-Okhotsk ocean (Fig. 1; Enkin et al., 1992; Heubeck, 2001). As noted for paleomagnetic data from northwest China, it seems possible that this rotation began during or prior to Permian collision. Along the North China block boundary in the Yin Shan Range (Fig. 1), >1000 m of Upper Carboniferous–Upper(?) Permian sedimentary rocks were deposited in braided fluvial to meandering fluvial and floodplain environments (Mueller et al., 1991; Cope et al., 2005). The youngest marine rocks in this region are Ordovician shallow-marine carbonates that unconformably underlie the Carboniferous and a younger nonmarine sequence. Significant changes occur within the upper Paleozoic stratigraphy in northeast China. Coeval with a shift from braided fluvial to meandering fluvial deposition, paleocurrents reversed from north-directed dispersal in Carboniferous–Lower Permian strata to south-directed dispersal in Upper Permian–Lower Triassic(?) rocks (Cope et al., 2005). Sandstone petrography indicates a recycled orogen to dissected arc provenance, representing unroofing of a thin sedimentary cover into continental arc plutons during the Late Permian (Cope et al., 2005). The late Paleozoic succession in northeast China is also marked by a climatic shift from humid-environment, coal-bearing deposits in the Carboniferous–Early Permian to widespread aridclimate, red bed deposition by Late Permian time (Mueller et al., 1991). Cope et al. (2005) suggest that the climatic shift represents a rain shadow effect created during uplift of the continental margin owing to collision, and/or northward movement of northern China to arid subtropical latitudes during this time. Uplift and significant crustal thickening along this margin during the Permian may have been further supported by Late Triassic–Early Jurassic (ca. 220–190 Ma) collapse of the orogen on the basis of radiometric ages of rocks involved in metamorphic core complex formation in northeast China (Davis et al., 2004). Summary Multiple lines of evidence indicate that collision between the southern Altaids and microcontinent blocks of north China (Tarim and the North China block) occurred during Carboniferous to Permian time. More specifically, suturing appears to have occurred earlier in the Tien Shan than in the Yin Shan Ranges, suggesting diachronous west-to-east closure. However, available radiometric data indicate that arc activity was at least partly overlapping in

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northeast and northwest China, and the transition from nonmarine to marine deposition is not well preserved across the length of the suture zone. Paleomagnetic data generally support the possibility of differential rotation of Tarim and the North China block during the late Paleozoic, and stratigraphic studies suggest that the Junggar trapped ocean basin may have opened toward the west rather than having been contiguous to the east. Thus the prospect of a more complex collisional event, with Tarim and the North China block acting somewhat independently and perhaps separated by an indenter (Lamb and Badarch, 2001), is intriguing and perhaps supported by our new data from Mongolia. Another important but poorly constrained aspect of this collision is what, if any, climate response occurred during and after arc-continent collision, and whether this might have been driven by orographic relief along the suture zone. Rain shadow and other topographic influences driven by collisional and postcollisional tectonics have been cited to explain observed stratigraphic changes in Permian basins of northwest and northeast China, but this is only one of multiple possible driving factors. Outstanding questions that we now seek to address include whether the Permian of Mongolia supports or refutes the timing and location of this suture zone, as suggested by data from the continental margin to the south. The potential climatic effects of this event are also investigated. Additional data from the southern Altaids help to evaluate how well various modern analogues, including the southwest Pacific Ocean, may apply to this setting. This ancient record addresses existing models of basin evolution, evolving sedimentary systems and source areas, and issues of preservation in comparison with modern analogues. Finally, interpreting subsequent Mesozoic–Cenozoic deformation requires understanding the pre- and syn-assembly record of southern Mongolia provided by this study. LATE PALEOZOIC SEDIMENTATION AND TECTONICS OF SOUTHERN MONGOLIA Overview The broad swath of land from southwestern to southeastern Mongolia encompasses some 1000 km along the northern edge of the Tien Shan–Yin Shan suture zone. Arc and accretionary complexes that define the Altaids are well exposed along basement uplifts in southern Mongolia (Fig. 2). Nevertheless, the number and nature of these intraoceanic arc systems is debated; the geologic complexity of the region is often represented by dozens of individual tectonic “terranes” and intricate models for their collision and evolution (Zonenshain et al., 1990; S¸engör et al., 1993; S¸engör and Natal’in, 1996; Badarch et al., 2002). Based on geochemical data and sandstone provenance data from >20 sites in southern Mongolia, Lamb and Badarch (2001) proposed a simpler interpretation of two main evolving arc systems, Devonian and Carboniferous, followed by Permian closure of the Paleo-Asian ocean. Devonian units are generally assigned to an island arc setting above a north-dipping subduction zone,

likely close to a continental margin or highly dissected sediment source to the north (Badarch et al., 2002). The Carboniferous arc shifted slightly south, closer to the present-day China-Mongolia border, and is at least partly built on Devonian strata. Carboniferous sandstones suggest a possible granitic source, although this may represent uplift and dissection of the Devonian arcs rather than proximity to a microcontinent (Lamb, 1998). Both systems have analogous components across the border and likely extended west into the Bogda Shan–North Tien Shan arc system (Fig. 1; Lamb and Badarch, 2001), albeit with some changes toward more oceanic arc affinities to the southwest. Similarly, the system is tentatively linked through southeast Mongolia and into northeast China (Xiao et al., 2003). These arcs were likely more complicated regionally, but the simpler interpretation reflects the resolution afforded by available data (Lamb and Badarch, 2001). This interpretation is also more representative of continuous arc systems, such as the modern Philippine Arc or other parts of the southwestern Pacific Ocean, where arc type, geometry, and subduction polarity are variable along strike but nevertheless are part of the same geodynamic setting. Regional geologic maps of southern Mongolia show multiple Precambrian units, implying the existence of a microcontinent along this part of the Tien Shan–Yin Shan suture, e.g., the South Gobi microcontinent (Fig. 2; S¸engör et al., 1993; Xiao et al., 2003). This interpretation is supported by a single U-Pb zircon age (916 ± 156 Ma) from gneiss on the Chinese side of the Yagan-Onch Hayrhan metamorphic core complex (Fig. 2; Wang et al., 2001); however, protolith ages have not been tested rigorously across the region. High-grade gneiss-schist-mylonite complexes in southeastern Mongolia are mainly Mesozoic tectonites (Webb et al., 1999; Webb and Johnson, 2006), and in many areas inferred protolith lithologies are more consistent with Paleozoic arc successions than with continental crust (Johnson et al., 2001; Webb and Johnson, 2006). In addition to metamorphic rocks, Proterozoic marble-quartzite units are mapped throughout the region. These are typically in flat, thrust-fault contact over Permian and younger rocks and are inferred to be klippen of allochthonous lower Mesozoic thrust sheets (Webb et al., 1999). Thus the existence of the South Gobi microcontinent is uncertain, and in contrast to the Tarim and North China blocks the Precambrian crust likely forms a very minor component of the southern Altaids in Mongolia, if at all. Although these rocks have not been described extensively in the English-language literature, Permian sequences in southern Mongolia are inferred to represent final closure of the PaleoAsian ocean between Mongolian arc terranes and continental block(s) to the south (Amory, 1996). The timing of this terminal collision is perhaps best constrained by Permian floral assemblages that indicate mixing between polar-temperate, Siberian (Angaran) communities and tropical, north China (Cathaysian) communities (Durante, 1971, 1983; Shen et al., 2006). Similarly, interpreted syn- or postcollisional granites in northeast China have ca. 250–230 Ma U-Pb and Rb-Sr ages, supporting Late Permian collision (Chen et al., 2000).

Sedimentary response to arc-continent collision The ages of Permian rocks in Mongolia are generally only poorly known. Volcanic lithologies are uncommon except for the Early Permian (Fig. 3), and even for this interval have not been extensively dated by modern radiometric techniques. Invertebrate biostratigraphy for marine sections in southeastern Mongolia relies mainly on brachiopod zonations, some with identification of bryozoans, crinoids, and foraminifers (Manankov, 1988, 2004). Nonmarine strata are constrained by plant fossils, invertebrates, and spore-palynology where available. These forms tend to be low-resolution age constraints (>1 m.y.), and the correlation from Russian “horizons” to the most recent global absolute time scales (Gradstein et al., 2004) is just currently under way (Shen et al., 2006; Shi, 2006; Manankov et al., 2006). Thus it is difficult to provide stage or even epoch assignments for many of the studied intervals. Improving age resolution and correlations to a global time scale are a focus of our ongoing work in Mongolia. For the

Standard Chronostratigraphy

Southwest Mongolia

purposes of this study we retain the use of Upper-Late Permian (Cisuralian, 299–270.6 Ma) versus Lower-Early Permian (Guadalupian–Lopingian, 270.6–253 Ma) terminology (Fig. 3; cf. Gradstein et al., 2004), as this is still widely used in regional literature and provides the best means for suggesting regional correlations. Generalized stratigraphic correlations (Fig. 3) show broad subdivisions for the Permian of southern Mongolia. Lower Permian red bed and rhyolite-andesite volcanic sequences are observed across the region, typically in fault contact with Carboniferous– Devonian arc sequences. Overlying middle Permian strata include shallow-marine carbonate and clastic successions, which conformably underlie turbidite deposits of the Lugyn Gol Formation (Fig. 3). A regional unconformity divides these flysch deposits from the nonmarine overlap succession, which includes coal-bearing fluvial strata that graded into braided fluvial successions in the Early Triassic (Fig. 3).

South-central Mongolia

Triassic 250 and younger

Early Permian

Unconformity

Unconformity

Tavan Har and Har Erdene Formations, >300 m thick. Shallow marine clastics, minor carbonate and volcanic units (Figs. 11, 12).

Agui Uul and Bulgan Uul Formations, >3000 m thick. Shallow marine carbonate and clastic deposits. Ulaan Nur Formation, ~300 m thick. Rhyolitic-andesitic volcanic rocks and nonmarine clastic deposits (Figs. 4-5).

?

? contact not exposed

290

Unconformity

Argalant Formation, >100 m thick, red beds and volcanic rocks.

295

Carboniferous and older

Nonmarine Upper Permian units not exposed.

Lugyn Gol Formation, >500 m thick. Distal turbidite deposits (Figs. 11, 12).

Lugyn Gol Formation, >100 m thick (partially preserved in south-central Mongolia). Distal turbidite deposits.

?

275

285

?

? Unconformity

?

280

Southeast Mongolia

Tavan Tolgoy and Tsaagan Tolgoy Formations, ~100-300 m thick. Coal bearing fluvial deposits (Figs. 6-10).

~1600 m meandering fluvial deposits, Noyon Uul (Figs. 4-5).

Lopingian

270

Cisuralian

Age (Ma)

265

Late Permian

260

Guadalupian

255

369

Carboniferous ? sedimentary and volcanic C2 rocks thrust over Permian (Noyon Uul, Fig. 5E) or P1 angular unconformity with Carboniferous and older.

?

Unconformable and fault contact with Carboniferous and older.

Argalant Formation, >100 m thick, red beds and volcanic rocks. ?

Unconformable and fault contact with Devonian.

Figure 3. General chronostratigraphic correlation of Permian sedimentary and volcanic rocks from southwest to southeast Mongolia (cf. localities shown in Fig. 2). Lithologies are indicated by patterns as shown in Figure 4 (key). Formation names, approximate thicknesses, lithologies, and contact relationships are compiled from our own work and published interpretations (Berkey and Morris, 1927; Grabau, 1931; Durante, 1971; Mossakovsky and Tomurtogoo, 1976; Zaitsev, 1974; Nikolov et al., 1981; Manankov et al., 2006). Dashed lines between sections represent inferred lithostratigraphic correlations, which in some cases are interpreted to be time-transgressive. Time scale is from Gradstein et al. (2004), although we have retained the use of Early versus Late Permian epochs, as discussed in the text. Age assignments are only approximate, given the low resolution of available age constraints from Mongolia.

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To fill in such broad correlations with more detailed fieldbased constraints, we now present new data from a regional southwest to southeast transect, highlighting localities where Permian strata are best preserved. The sections were visited from 1992 to 2005, and so this is a long-term compendium of our studies. We present these observations at a reconnaissance scale in order to emphasize regional relations; much of this interpretation is derived from more detailed descriptions in Amory (1996) and Lamb et al. (1999). This work focuses on the identification of sedimentary facies to highlight basin setting and paleogeography, to constrain marine to nonmarine transitions, and to identify shifting climatic and tectonic controls on sedimentation. Stratigraphic relations were identified on the basis of measuring and describing sections and, where possible, collecting paleocurrent data. Climatic indicators are mainly inferred from study of nonmarine sedimentary facies, particularly the distinction between highly organic and coal-bearing strata that formed in wet to seasonally humid environments versus red beds, with common exposure and desiccation features indicative of aridity (Parrish et al., 1982; Wright, 1992; Dubiel and Smoot, 1994). Spores and other palynomorphs from the Tsaagan Tolgoy locality (Fig. 2) support these interpretations. Sedimentary provenance, mainly derived from sandstone petrofacies, is also discussed as an additional constraint on basin tectonic setting and evolving source areas. Southwest Localities Permian strata in southwest Mongolia (Fig. 2) lie unconformably on Lower Carboniferous and older marine sedimentary and volcanic units. Although Devonian–Carboniferous successions have distinct arc-related geochemistry, Permian strata are nonmarine and are interbedded with basalts that have intraplate geochemical signatures (Lamb and Badarch, 2001). These relations support a postcollisional setting by Permian time. At Shin Jinst (Fig. 4), Zonenshain et al. (1974) reported Late(?) Permian plant fossils in a section of ~100 m of reddish argillite, sandstone, and conglomerate (Fig. 5A). Although beds are cleaved and deformed in places, sedimentary structures are relatively well preserved. Organic material is abundant throughout, including in-place petrified wood stumps. Sandstone beds are typically lenticular over several meters, have sharp and erosional bases, and are trough to low-angle cross-bedded and current-rippled (Fig. 5B–D). These features are interpreted as fluvial channels, and given the abundance of red and gray mudstone, representing floodplain deposits, we suggest mainly meandering fluvial environments. The presence of soft sedimentary deformation features, climbing ripples, and syndepositional clastic dikes may also be consistent with a delta-plain environment (Battacharya and Walker, 1992), even though distal marine or lacustrine successions are not preserved. There is an overall upward-coarsening trend, possibly representing a shift from meandering to braided fluvial deposition from Late Permian to Early Triassic time.

Approximately 250 km southeast of Shin Jinst at Noyon Uul (Fig. 2), ~1600 m of Upper Permian strata include interbedded sandstone and mudstone intervals that are also interpreted to represent meandering fluvial environments (Fig. 4; full section shown in Hendrix et al., 1996). Sandstone beds are lenticular over ~10 m (Fig. 5F) with minor erosional scouring at their bases and are upward fining, with common trough and planar cross-bedding. Reworked terrestrial organic matter is present throughout. The section is dominated by floodplain deposits of variegated silt and mudstone; carbonate nodules and minor paleosols (argillosols) are also present (Fig. 5G, H). Age control in this part of the section is based on well-preserved Late Permian flora and fauna (Zaitsev et al., 1973; Durante, 1976; other references in Hendrix et al., 1996). However, the uppermost portion of the meandering fluvial succession has been assigned an Early Triassic age, based upon tetrapod fossils (Gubin and Sinitza, 1993). This generally fine-grained sequence is sharply overlain by thick Middle(?) Triassic conglomerates (Zaitsev et al., 1973; Hendrix et al., 2001). The coarse-grained Triassic–Jurassic units at Noyon Uul were deposited in an intraplate setting, likely within a flexural foreland basin (Hendrix et al., 1996). They clearly postdate the main collisional event along the Tien Shan–Yin Shan suture but likely record ongoing intraplate contractile deformation. The Late Permian floral assemblage at Noyon Uul is also notable because it contains a mixture of Cathaysian (north China affinity) and Angaran (Siberian affinity) flora (Durante, 1971; Amory, 1996). This is one of the early phases of population mixing prior to complete homogenization in the Triassic, and is detected at localities in south-central Mongolia as well. In particular, an unusual cycad genus (Guramsania) has been found at both Noyon Uul and Tsaagan Tolgoy (Figs. 2, 6), providing further evidence for a connection between these localities (Vakrameyev et al., 1986). Such floral mixing in the transition zone along the China-Mongolia border provides compelling evidence of close approach and/or collision with the North China block (Li, 2006; Shen et al., 2006). Although Upper Permian fluvial strata measured within the Noyon Uul syncline are generally consistent with at least periodically wet-humid conditions (Hendrix et al., 1996), Permian strata unconformably underlying this section are quite different. Anatoleva (1974) describes ~315 m of distinctively red extrusive volcanic units interbedded with sandstone and mudstone exposed southeast of the Noyon Uul syncline (Figs. 2, 4). Our observations of this section are generally consistent with Anatoleva’s (1974) descriptions. The lowermost unit is in thrust-fault contact with middle–Upper Carboniferous arc sequences (Fig. 5E). It consists of ~55 m of poorly sorted cobble-boulder conglomerate interbedded with andesitic-rhyolitic crystalline tuffs ranging from 0.5 to 2.5 m thick. This sequence fines upward into a distinctive red and reddish-brown succession (~140 m total) of interbedded siltstone, sandstone, and conglomerate interbedded with tuff, tuff breccias, and rhyolitic ash-fall deposits and rare trachytebasalt flows. The uppermost ~120 m of section grades to mainly

Key for all stratigraphic sections siderite horizon root marks

current ripples low angle cross beds low angle / wavy bedding planar laminae trough cross beds

mud chips carbonate nodules soft sediment deformation cover coal mudstone-siltstone fine-medium grained sandstone volcanic and volcaniclastic sandstone gravel-pebble conglomerate limestone or dolomite cobble-boulder conglomerate

Upper Permian, Shin Jinst

(m)

Upper Permian (?) - lowermost 100 Triassic, Noyon Uul syncline (Western Transect, trough Hendrix et al., 1996) cross beds n=58

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(m) angular unconformity 300

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0m r t. t. l. ve lts ss cg co d/si d. ble e b u m m pe

0m r t. t. l. ve lts ss cg co d/si d. ble e b u m m pe

r t. t. l. ve lts ss cg co d/si d. ble e b u m m pe

Figure 4. Measured sections from southwest localities (see Fig. 2 for locations). Key applies to all measured sections (Figs. 4, 6, 9, 11). The left-hand section (Lower Permian, southeast of Noyon Uul syncline) is a schematic representation based on Anatoleva (1974) and our own field observations. The Noyon Uul syncline detailed section was measured in 1992 by M. Hendrix and E. Sobel (Western Transect of Hendrix et al., 1996; ~1500 m from base), and is redrafted with permission. This section was measured close to the Permian–Triassic boundary, and its exact age is uncertain; however, the depositional style is representative of the underlying, continuous Permian section. Paleocurrent measurements are from the same approximate stratigraphic interval, measured ~10 km along strike (Sain Sar Bulag transect; Hendrix et al., 2001). The Shin Jinst section was measured as part of this study, as described in the text.

A

B

~ 30 meters

C

D

E

F

~10 m

G

H

Figure 5. Photographs from Permian strata at the southwest localities (F–H from M. Hendrix). (A) Overview of Shin Jinst section. (B, C) Channel cut-and-fill geometries at Shin Jinst. (D) Soft sedimentary deformation within a channel sandstone body, Shin Jinst. (E) Lower Permian rocks southeast of Noyon Uul syncline, showing red beds and volcaniclastic rocks in the foreground. High ridges in the background are Carboniferous sedimentary and volcanic rocks thrust over Permian (teeth on overriding plate). (F) Overview of detailed Upper Permian section at Noyon Uul. (G) Carbonate nodules within interdistributary mudstone-siltstone deposits, Upper Permian, Noyon Uul. Stratigraphic up is to the upper left. (H) Variegated red-green argillosol and resistant carbonate zone, Upper Permian, Noyon Uul (hammer for scale).

U. Permian-L.Triassic Tsaagan Tolgoy, detailed section U. Permian-L.Triassic Tsaagan Tolgoy, long section

(m)

70

last petrified wood fragment (at about 80 meters)

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last siderite horizon

major shift in pollen assemblage

detailed section at right (measured along strike)

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color change to dark gray to greenish-light gray mudstones

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l r oa t. t. l. ve c ilts . ss cg o c d/s d ble e b u m m pe

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al er coltst. sst. cgl. v i co d/s d. ble e b u m m pe

Figure 6. Long (at left) and detailed (right) measured sections across the Permian–Triassic boundary at south-central locality Tsaagan Tolgoy. Xs mark palynological samples used in Figure 7. See Figure 4 for key to lithologies and symbols.

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cross-bedded sandstone and siltstone, which are variegated with yellow and greenish gray more commonly than in the underlying sections (Fig. 4). In general the section is coarse grained and poorly sorted, with a trend to better sorting and rounding starting at ~100 m from the base. Organic matter is rare, probably owing to volcanic contributions and poor preservation under oxidizing conditions. These observations are consistent with an alluvial fan setting, close to an active volcanic source area, possibly grading to braided fluvial deposition through time. A Late Permian age for this volcanic and clastic succession southeast of the Noyon Uul syncline has been suggested (Anatoleva, 1974); however, this lower section appears to have an angular relationship with the Upper Permian strata described by Hendrix et al. (1996) and is also quite distinctive lithologically (with its red color, abundant volcanic rocks, and lack of coal measures). The age of the lower section is bracketed between Late Carboniferous, which is in fault contact with this section, and Late Permian; thus we favor an Early Permian age for these clastic and volcanic units. South-Central Localities One of the better known nonmarine Permian successions in south-central Mongolia is found at Tsaagan Tolgoy (also known as Ih Uvgon, Zinniker and Badarch, 1997; Figs. 2, 6). Age control here is excellent, based on our own and previous paleontological analyses (Durante, 1976; Vakrameyev et al., 1986; Zinniker and Badarch, 1997; Uranbileg, 2003; Aristov, 2005), which document Upper Permian through Lower Triassic strata (Fig. 6). Characteristic Permian pollen genera found at Tsaagan Tolgoy (identified by D. Zinniker, Fig. 7) include Cordaitina, Florintes, Potonieisporites, and Neoraistrickia. Triassic and transition zone species include Lueckisporites virkiae, Lundbladispora sp., and Taenisporites sp. The Permian–Triassic palynological shift is similar to that documented along the southern margin of the Junggar basin (Ouyang and Norris, 1999). The Upper Permian section at Tsaagan Tolgoy (Fig. 6) is ~300 m thick and consists primarily of dark, organic-rich mudstone, coaly zones, and thin sandstone beds (Figs. 7, 8A). Petrified wood is common, including paleohorizontal and growthposition logs encased in tabular sandstone beds (Fig. 8B). The section is generally fine grained, with multiple concentrated zones of centimeter-scale siderite nodules, indicating swampy, reducing conditions (Landuydt, 1990). Sandstone units range from thin beds to 1–4-m-scale individual and amalgamated units and commonly exhibit convolute bedding and slumping and upward-fining trends (Fig. 8C). Trough cross-beds and asymmetric current ripples are also found within the discontinuous sandstone beds. Overall, this section is consistent with poorly drained interdistributary or backswamp deposition, with sandier units representing mainly crevasse splay deposition or meandering of minor channels across the floodplain. Above the Permian–Triassic boundary (~280 m above the base of the section; Figs. 7, 8D), the section at Tsaagan Tolgoy

coarsens upward abruptly, with the first appearance of pebblecobble conglomerate at ~400 m. Conglomeratic and sandy intervals contain large-scale trough cross-beds and sharp to erosive bases (Fig. 8E). Above 280 m, mudstone units grade abruptly from organic-rich, dark-gray units to lighter colored tan-green and organic-poor beds. Carbonate nodules also appear in the upper section, and siderite is comparatively rare, indicating a shift to relatively well-drained soils. Overall, this vertical trend is interpreted as a change to higher energy fluvial environments, possibly including braided-stream deposition, with more oxidized and less organic-rich floodplain deposition. A distinctive shift in floral assemblages also occurs at Tsaagan Tolgoy (Fig. 7). Late Permian flora are dominated by Cordaitales (pre-gymnosperms), ferns, and bottryococcus algae, whereas the Triassic assemblage includes taeniate-striate bisaccate pollen typically associated with evaporites and more arid conditions. Cavate microspores (Lundbladispora sp.) and megaspores characteristic of Early Triassic heterosporous lycopods (e.g., the desert xerophyte Pleuromeia) were also found in the upper part of the section. An observed 4–6‰ decrease in the δ13C ratio of organic carbon is coincident with changes in lithology and pollen assemblage (Fig. 7A). Similar shifts across the Permian–Triassic boundary in marine and nonmarine sections around the world are believed to reflect a large, globally synchronous contribution of isotopically light carbon into the atmosphere (Baud et al., 1989; Krull and Retallack, 2000). Along with the lithologic changes noted previously, the rapid floral and isotopic shift is thought to represent the demise of Late Permian coal-bearing swamp facies and replacement with primarily herbaceous flora tolerant of seasonal droughts. The change may reflect the global record of warming across the Permian–Triassic boundary (Faucett et al., 1994; Retallack et al., 1996; Berner, 2002). However, a similar aridification trend observed within Permian rocks of northeast China has also been attributed to a possible rain shadow effect, reflecting ongoing collision and uplift along the Tien Shan–Yin Shan suture (Cope et al., 2005). Supporting evidence for this climate shift is found at the Tavan Tolgoy and Naftgar Uul localities, described below. Although neither of these sections records the actual Permian–Triassic transition, as at Tsaagan Tolgoy, individually they provide supporting evidence of distinctive and shifting climate regimes. Additional Upper Permian coal-bearing fluvial deposits are preserved within coal quarries at Tavan Tolgoy (Durante, 1971; Fig. 9). Coal seams are individually up to 14–33 m thick and interbedded with coarser units, including rare pebble conglomerate and medium- to fine-grained sandstone. Sandstones are trough cross-bedded with sharp, erosional bases, and are lenticular, although amalgamated units are laterally continuous over tens of meters (Fig. 10). Sandstone beds exhibit normal grading, and fine upward to rippled siltstone and mudstone. Paleocurrent measurements at Tavan Tolgoy are primarily south-southwest directed (Fig. 9). Overall, outcrops at Tavan Tolgoy are broadly analogous to the Upper Permian at Tsaagan Tolgoy and are consistent with a meandering fluvial system in a relatively humid environment.

A

Relative abundance of palynomorph groups: present

common

abundant

characteristic Triassic taxa

smooth or small sculptured cavate spores (7E) megaspore wall fragments spore tetrad taeniate bisaccate (7D)

transitional

striate bisaccate unsculptured acavate spores normal bisaccate

characteristic Permian taxa

monosaccate sculptured acavate spores (7B and 7C) large sculptured cavate spores bottryococcus cycadopites

d 13 C (‰)

-23 -25 -27 Permian

Triassic

-29 0

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10 µ µm m

40

80

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120 160 200 240 280 320 distance in meters from base of section

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400

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10 µm

Figure 7. (A) Spore and pollen abundance and δ13C isotope data from Tsaagan Tolgoy, plotted against distance above the base of the long section shown in Figure 6. (B–E) Photomicrographs of select spores including Neoraistrickia (B), Lycopodiumsporites (C), Leuckisporites (D), and Lundbladispora (E), keyed to the palynomorph groups shown above. Carbon isotopic data were collected on acid-resistant kerogen isolated for palynological analysis, conducted at Petrobras stable isotope laboratory in 1997 following industry-standard methods.

A

B

C

D

E

10 cm Figure 8. Photographs from Tsaagan Tolgoy. (A) Coaly intervals at the base of the section. (B) Petrified log in growth position. (C) Convolute bedding in overbank deposits. (D) Transition from dark, organic-rich mudstone to tan, organic-poor mudstone at the Permian–Triassic boundary (~290 m). (E) Trough cross-beds in a lenticular, channelized sandstone.

Figure 9. Measured sections from south-central localities Naftgar Uul and Tavan Tolgoy (see Fig. 2 for locations). Paleocurrent data from Naftgar Uul include data from possible Triassic strata overlying this section. See Figure 4 for key to symbols and lithologies.

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A

The lower part of the section is more consistent with sedimentary facies found in the Triassic across the region, and thus it is likely that these strata are younger than mapped. The section at Naftgar Uul includes more than 300 m of interbedded sandstone, siltstone, and rare granular to pebble conglomerate. Sandstone beds contain planar laminations, low-angle cross-stratification, and trough cross-beds with >1 m of relief. These beds typically have sharp, slightly erosive bases, and amalgamated units are laterally continuous over >50 m. Paleoflow directions measured in trough cross-beds are variable but are generally west-directed (Fig. 9). Coal is absent in the section, and only rare, poorly preserved and oxidized plant fossils are present. Lower Triassic strata here and at Tsaagan Tolgoy are consistent with sand-rich meandering rivers to braided-stream environments, with oxidizing environments and likely periodic arid-drought conditions.

B

Southeast Localities

Figure 10. Photographs of the Tavan Tolgoy coal mine section shown in Figure 9. (A) Lenticular sandstone bodies and interbedded coals. Person (circled) for scale. (B) Detailed photo of finely laminated siltcoal beds in interdistributary deposits.

As at Noyon Uul, Upper Permian strata at Tavan Tolgoy and Tsaagan Tolgoy contain uniquely mixed Angaran-Cathaysian assemblages (Vakrameyev et al., 1986) and thus indicate that south Mongolia and north China were in close proximity at that time. Widespread terrestrial deposition in this area likely occurred in syn- and postcollisional flexural basins (Hendrix et al., 2001). Another similarity to the Noyon Uul section is that Upper Permian strata near Tavan Tolgoy overlie extrusive intermediate-felsic volcanic rocks and interbedded reddish conglomerate, sandstone, and siltstone that are thought to be Lower Permian (Ulaan Nur Formation, Fig. 3; Amory, 1996). Thus there is apparently an older climatic shift from generally arid to more humid climates in the middle Permian, followed by a return to aridification at or near the Permian–Triassic boundary. In comparison with Tsaagan Tolgoy, Triassic strata are better exposed nearby at Naftgar Uul (Figs. 2, 9). Although shown as Permian on regional maps (Tomurtogoo, 1999), one of us (D. Zinniker) found Late Triassic to Middle Jurassic pollen assemblages (Classopollis) in lacustrine facies near the top of the section.

The final phase of marine deposition in southern Mongolia is preserved in outcrops of the Permian Lugyn Gol Formation, exposed at Nomgon and Hovsgol (Figs. 2, 11). Devonian turbidite deposits are also reported in the area, which has led to some confusion in mapping, where these widespread outcrops are alternately assigned either to the Early–Middle Devonian or the Late Permian (cf. Yanshin, 1989, and Tomurtogoo, 1999). However, the Late Permian age of these deposits is well established in multiple areas (Pavlova et al., 1991) and is confirmed by our own findings in 1995 and 2005 of Late Permian crinoids and bryozoans (J. Undariya, 2005, personal commun.). It is possible that Devonian crinoids and corals found in the section are reworked and transported. The Lugyn Gol Formation crops out along northeast-trending ridges at Nomgon (Fig. 2). Beds are highly deformed and cleaved with large-scale isoclinal folding and overturned units. Thus true thickness is difficult to determine, but we estimate a minimum thickness of ~800 m, based on individually continuous, non-repeated sections (Fig. 11). Permian flysch units mainly lie unconformably or in fault contact with Carboniferous and older metasedimentary, volcanic, and intrusive rocks, including ophiolites, carbonate klippen, and metamorphic rocks mapped as Precambrian (Fig. 3; Tomurtogoo, 1999). The lowest part of the section includes a possible shallowmarine succession that is marked by 10–50-m-scale upwardcoarsening and -thickening sandstone packages (Figs. 11, 12A). Dolomitic units include crinoid, brachiopod, and bivalve coquina and grainstone. Low-angle cross-bedding includes trough and possible hummocky cross-stratification, and ripple marks are also present (Figs. 12B, C). Beds are generally tabular rather than strongly lenticular, but they do pinch out over hundreds of meters along strike. Sandstone beds have sharp bases but generally lack channel-like scour features or bar forms, and no subaerial exposure surfaces were identified. This lowest section is interpreted as a shallow-marine succession, possibly along a fluvial shoreline with upward-shallowing prodeltaic parasequence packages (Fig. 11; Posamentier and Allen, 1999). Similar units

Figure 11. Measured sections through the Upper Permian Lugyn Gol Formation at Nomgon, southeast Mongolia (see Fig. 2 for location). Sections are noncontinuous but are shown in stratigraphic order from oldest to youngest (left to right). Paleocurrent data are all from uppermost stratigraphic section. See Figure 4 for key to lithologies and symbols.

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A

B

C

D

~25 m

E

F

Figure 12. Photographs of the Lugyn Gol Formation exposed at Nomgon. (A) Upward-coarsening and thickening packages (lowermost stratigraphic section, Fig. 11). (B) Trough cross-beds (lowermost section). (C) Ripple marks at top of bed, middle section. (D) Overview of turbidite sequence. (E) Single upward-fining and -thinning unit (overturned beds, from uppermost section, Fig. 11). (F) Flute casts at the base of sandy turbidite bed.

are found to the east at Hovsgol (Fig. 2), where crinoid and brachiopod grainstones are present with sporadic plant fossils, bedding-plane parallel and perpendicular bioturbation, and ripples and convolute bedding, which may also support shallow-marine and/or prodeltaic deposition. There is a rapid transition from the shallow-marine packages into regular turbidite sequences (Fig. 11, uppermost section) at Nomgon. The classic turbidite flysch architecture of the Lugyn

Gol Formation includes hundreds of meters of tabular, continuous sandstone and shale beds (Fig. 12D). Sandstone units are centimeter to decimeter scale in thickness and are typically upward fining (Fig. 12E). Basal contacts are sharp to slightly erosive and commonly preserve tool marks and flute casts, which generally show south-southeast transport directions (Figs. 11, 12F). Structures include normal grading, sporadic load structures, planar laminae, and some current ripples at the tops of beds, corresponding to Ta,

Sedimentary response to arc-continent collision

mixed

oro ge n

co nti ne nta lb loc

k

led

Sandstone samples were collected for petrographic analysis and point counting (Fig. 13; Table 2) to complement facies analyses. Sandstone provenance studies permit evaluation of the tectonic setting in which sedimentary sequences are deposited and can demonstrate evolution of source areas through time (Dickinson and Suczek, 1979; Dickinson et al., 1983). In other arc collisional settings, sandstone compositional trends have been attributed to erosion of magmatic arcs and unroofing of collisional orogens (Dorsey, 1988; Graham et al., 1993) and thus record timing and evolution of suturing and postcollisional uplift. Sandstone provenance studies are best conducted on mediumgrained sandstone samples that are relatively unaltered. Given the level of deformation and burial in many Permian outcrops of the region, not all localities are represented. Nevertheless, 82 Permian samples from 6 localities across southwestern to southeastern Mongolia were point-counted using a modified Gazzi-Dickinson method (Ingersoll et al., 1984), which compensates for the potential adverse effect of grain-size variations. Principal framework grains were counted by standard categories (Dickinson, 1970) and then normalized as detrital modes on standard ternary diagrams for comparison with global modal distribution fields (Table 2; Dickinson et al., 1983; Marsaglia and Ingersoll, 1992). In some cases, sandstones were extensively altered, such that distinguishing polycrystalline quartz and chert from altered and devitrified zeolites or volcanic lithic fragments was impossible. Thus the QmFLt ternary diagram (Fig. 13; Qm = monocrystalline quartz,

yc

Sandstone Provenance Data

Qm

rec

Tb, and Tc/Te petrofacies (respectively) of Mutti and Ricci Lucchi (1978). Local zones of intense cleavage preclude identification of small-scale sedimentary structures, but the rhythmic sand-shale interbedding and tabular, laterally continuous units (Fig. 11) are consistent with turbidite deposition, likely within Mutti and Ricci Lucchi’s (1978) mixed sand-mud facies (facies C). These turbidite deposits are tentatively assigned to an outer submarine fan setting that is intermediate between proximal and distal fan deposition (Walker, 1992). Coarse-grained facies, highdensity turbidite flows, slumping and channel-levee systems were not observed, which argues against a more proximal inner fan setting. Similarly, the system is quite sandy, and bioturbation is rare, arguing against a distal basin-floor fan environment. However, overall packaging is variable between more shaly and more sandy intervals, and this likely indicates fluctuations in relative sea level through time or shifting feeder channels within the fan (Amory, 1996). The distal submarine environments observed at Nomgon and Hovsgol are consistent with depositional settings observed in analogous remnant ocean-basin fan systems, where basin geometry favors trough-like axial deposition, and turbidite pulses are funneled outboard over long distances (Graham et al., 1975). We suggest that the underlying shallow-marine successions formed along the margin of an extinct Carboniferous–Devonian arc, immediately prior to final collision stage marked by rapid subsidence and deposition of overlying turbidite successions.

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Dissected arc

Transitional arc F

Lt Undissected arc Shin Jinst (n=11) Noyon Uul (n=21)

Southwest localities (Upper Permian)

South-central locality (Triassic?) Nomgon (n=21) Southeast localities (Upper Permian) Hovsgol (n=10) Effect of dynamic metamorphism (Nomgon and Shin Jinst samples) Naftgar Uul (n=18)

Figure 13. Ternary diagram of sandstone modal compositions, using monocrystalline quartz (Qm), potassium plus plagioclase feldspar (F), and total lithic fragments (Lt). Mean values are shown with a field of one standard mean deviation (Table 2). Provenance fields are from Dickinson et al. (1983), following methods described in the text.

F = plagioclase + potassium feldspar, Lt = sedimentary, metamorphic, and volcanic lithic fragments + carbonate and chert) is the most reliable comparison between all localities. Recent studies of modern arc-derived sands also use the QtFL ternary plot (Marsaglia and Ingersoll, 1992), which combines monocrystalline, polycrystalline, and chert into the Qt (total quartz) field. However, Qp (polycrystalline quartz) is a minor component in most modern arc–derived sands (Marsaglia, 1992), so use of QmFLt modal compositions provides a reasonable comparison for the southern Mongolia Permian sample suite. Upper Permian sandstones from southern Mongolia range from undissected to transitional to dissected arc provenance fields (Fig. 13), indicating recycling or inheritance of late Paleozoic arc sources (Graham et al., 1993). None of the samples has a recycled orogen or continental block signature. There does not appear to be a consistent spatial trend in samples from across southern Mongolia: i.e., dissected versus undissected arc provenance fields are not

TABLE 2. POINT COUNT DATA FROM PERMIAN SANDSTONES OF SOUTHERN MONGOLIA Shin Jinst 94-SJ-181 94-SJ-182 94-SJ-230 94-SJ-231 94-SJ-240 94-SJ-241 94-SJ-243 94-SJ-245 94-SJ-246 94-SJ-247 94-SJ-251 SJ Mean SJ St Dev

Qm 17.66 16.36 41.11 26.98 24.24 25.87 30.70 32.75 34.67 31.99 23.48 27.80 7.03

F 29.58 43.18 42.78 43.25 34.34 11.89 28.73 21.62 21.07 18.01 30.19 29.51 10.24

Noyon Uul (TU = Tost Uul, near Noyon Uul) 95-NU-2 0.00 4.55 95-NU-3 0.76 15.78 95-NU-4 3.57 13.74 95-NU-8 5.79 3.22 92-NU-8 1.43 9.07 92-NU-10 1.20 19.60 92-NU-13 6.20 19.90 92-NU-14 6.80 22.20 92-NU-15 10.59 9.85 92-NU-39 5.52 4.94 92-NU-40 6.10 18.30 92-NU-41 8.25 18.45 97-TU-1 6.40 18.50 97-TU-4 8.80 12.50 97-TU-7 5.40 16.40 97-TU-8 5.90 10.80 97-TU-13 4.80 18.30 97-TU-14 3.70 15.50 97-TU-15 5.90 11.80 97-TU-16 10.20 18.50 97-TU-19 2.00 8.90 NU Mean 5.21 13.85 NU St Dev 2.89 5.41

Lt 52.76 40.45 16.11 29.76 41.41 62.24 40.57 45.63 44.27 50.00 46.33 42.69 11.47

95.45 83.46 82.69 91.00 89.50 79.20 73.90 71.00 79.56 89.53 73.30 73.30 75.10 78.70 78.30 83.30 76.80 80.80 82.20 71.30 89.10 80.83 6.79

Naftgar Uul (continued) 93-NG-14a 93-NG-15 93-NG-15a 93-NG-24 93-NG-25 93-NG-26 93-NG-27 93-NG-27a NG Mean NG St Dev

Qm 18.37 10.83 14.46 20.60 17.49 16.84 25.27 17.36 16.52 4.24

F 38.86 35.04 28.19 36.54 28.72 31.02 22.87 38.14 33.52 5.08

Lt 42.77 54.13 57.35 42.86 53.79 52.14 51.86 44.50 49.96 5.71

Nomgon (Lugyn Gol Fm.) 93-LG-201 30.80 93-LG-202 38.70 93-LG-204 26.10 93-LG-206 39.10 93-LG-207 19.30 93-LG-208 24.60 93-LG-209 23.40 93-LG-210 26.40 93-LG-211 25.60 93-LG-301 26.30 93-LG-302 31.00 93-LG-303 32.90 93-LG-304 30.90 93-LG-305 20.60 93-LG-306 27.40 93-LG-307 30.90 93-LG-308 26.60 93-LG-309 27.50 93-LG-401 32.00 93-LG-402 43.20 93-LG-403 26.60 LG Mean 29.04 LG St Dev 5.78

22.90 22.10 25.00 19.00 28.90 23.50 27.10 23.40 25.90 29.30 27.00 25.20 18.00 29.40 31.80 21.10 29.60 30.50 24.70 24.80 15.70 25.00 4.17

46.20 39.10 48.90 41.90 51.90 51.80 49.50 50.20 48.50 44.40 42.00 41.90 51.10 50.00 40.80 48.00 43.80 41.90 43.30 32.00 57.60 45.94 5.54

Hovsgol (Lugyn Gol Fm.) XO-1A XO-1B 5.08 XO-1C 1.52 XO-1D 6.57 93-XO-3A 1.63 93-XO-3B 1.70 93-XO-3C 0.59 93-XO-3E 5.23 93-XO-3F 5.85 93-XO-3G 1.56 XO Mean 3.39 XO St Dev 2.09

4.14 28.43 33.59 28.71 12.23 14.16 24.56 18.90 22.15 17.19 22.52 6.52

25.30 66.50 64.90 64.72 86.14 84.14 74.85 75.87 72.00 81.25 74.09 7.40

Naftgar Uul 93-NG-1 15.20 41.52 43.27 93-NG-2 15.01 29.01 55.98 93-NG-3 18.25 31.22 50.53 93-NG-5 18.69 26.11 55.19 93-NG-6 12.00 39.50 48.50 93-NG-7 5.43 33.33 61.23 93-NG-9 15.79 34.59 49.62 93-NG-12 20.59 32.94 46.47 93-NG-13 20.17 40.34 39.49 93-NG-14 14.91 35.50 49.59 Note: QmFLt normalized to 100%. Qm—monocrystalline quartz; F—total feldspar, Lt—total lithic fragments (see text for explanation of methods); St Dev—standard deviation.

Sedimentary response to arc-continent collision linked to the southwestern, south-central, or southeastern study areas, respectively. The dissected arc provenance indicated by sandstone samples from Shin Jinst and Nomgon may in fact be an artifact of alteration level. Dynamic, greenschist-grade metamorphism causes notable recrystallization and breakdown of labile grains, including chert, calcite, and fine-grained lithic volcanic fragments, into a slaty fabric matrix. This causes a relative compositional increase in the quartzofeldspathic fraction, which might indicate a more transitional-arc origin for the original composition of these sandstones (gray arrow in Fig. 13). This may explain the difference between Hovsgol (undissected arc) and Lugyn Gol sand compositions, as the Lugyn Gol turbidite sand samples have undergone a comparatively high degree of alteration. By comparison, Naftgar Uul sandstone provenance samples lack any metamorphic overprint. These samples are more compositionally mature, but they are also interpreted as possibly younger (Triassic) deposits than the other localities plotted here. They are distinctly richer in monocrystalline quartz. At Naftgar Uul there is no statistical change in sandstone composition over a 2000-m-thick sampling interval (Amory, 1996). In contrast, sandstone composition at Noyon Uul changes drastically near the Permian–Triassic boundary from lithic-rich to more quartzofeldspathic (Hendrix et al., 1996; note that only the Permian sample compositions from Noyon Uul are presented here). DISCUSSION Tectonic Setting and Stratigraphic Evolution Permian sedimentary basins of southern Mongolia formed along the northern margin of an evolving collision zone between

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amalgamated middle to late Paleozoic oceanic arc systems and microcontinental blocks of north China. Upper Permian sedimentary successions in southwestern Mongolia (e.g., Shin Jinst and Noyon Uul localities) comprise entirely nonmarine deposits, including meandering to braided fluvial environments. Similar successions are found in south-central Mongolia (e.g., Naftgar Uul, Tavan Tolgoy, and Tsaagan Tolgoy), albeit with more extensive coal measures. These deposits likely formed in separate intramontane, flexural basins formed in response to collision and continued crustal thickening along the western end of the Tien Shan–Yin Shan suture (cf., Hendrix et al., 2001). In contrast, Upper Permian strata in southeast Mongolia (e.g., Nomgon and Hovsgol) include distal turbidite sequences that record the final phase of marine deposition in a closing, trapped ocean basin immediately prior to final collision by Early Triassic time. These regional facies relationships (summarized in Fig. 14) broadly support diachronous collision along the Tien Shan–Yin Shan suture zone. This interpretation is based mainly on the progression of youngest marine rocks from the middle Carboniferous in southwestern Mongolia to the Late Permian in the southeast, as well as evidence for earlier cessation of arc-related magmatism in northwest China (Zhou et al., 2001; Cope et al., 2005). However, several caveats accompany this general interpretation. The marine-to-nonmarine facies transition is not smoothly continuous from west to east. The only classic flysch deposits preserved in southeastern Mongolia are part of the Lugyn Gol Formation (Nomgon and Hovsgol localities, Fig. 14). Carboniferous or Lower Permian flysch deposits are not present in the study areas; indeed, distal marine deposits predating the Upper Permian are sparse throughout southern Mongolia (Lamb and Badarch, 2001). Furthermore, Upper Carboniferous strata across the border in the

Mongol-Okhotsk ocean: T1 marine sedimentation in NE Mongolia SW localities (NU, SJ)

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Junggar: W-facing P1 trapped ocean basin? T1 P2 P1 C2

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0 250 500 km nonmarine, arid climate facies nonmarine, humid climate facies marine strata (ocean basin) and/or oceanic arc sequences

sediment dispersal direction Inferred Late Paleozoic rotation direction.

Figure 14. Schematic summary of regional tectonic and stratigraphic relationships discussed in this study. EGFZ—East Gobi fault zone. Southwest (SW) localities: NU—Noyon Uul; SJ—Shin Jinst. South-central (SC) localities: NA—Naftgar Uul; TT—Tavan Tolgoy; TsT—Tsaagan Tolgoy. Southeast (SE) localities: NO—Nomgon; HO—Hovsgol. Age divisions are only approximate, based on available resolution as discussed in the text.

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Junggar basin may indicate west-directed closure and clockwise rotation of Tarim, which suggests that perhaps Tarim and the North China block were not colliding as a single block (Fig. 14). Thus the sedimentary record appears to support Late Carboniferous–Early Permian collision between the Altaids and Tarim, versus Late Permian amalgamation between the Altaids and the North China block; however, this does not seem to have occurred in a contiguous east-west basin with simple progressive suturing proceeding toward the east (Fig. 15). Provenance data underscore that the progressive closure model should not be oversimplified. A transition from undissected-dissected arc to recycled orogen provenance is observed as an unroofing trend in many arc-collisional settings (Dorsey, 1988; Clift et al., 2003). If closure of the Paleo-Asian ocean was accommodated by progressive counterclockwise rotation and northward movement of Tarim and North China blocks combined, we might expect to see a west to east trend from dissected to undissected arc provenance in our study areas indicating earlier arc dissection in the west. No such trend was observed, nor is

there evidence for systematic cessation of arc activity from west to east in Mongolia. Sand compositions represent recycling of Carboniferous and Devonian arc systems (Graham et al., 1993; Lamb and Badarch, 2001) in contrast to Permian sandstones in northeast China, which range from dissected arc to recycled orogen provenance (Fig. 13; Mueller et al., 1991; Cope et al., 2005). This indicates a much stronger continental (quartzofeldspathic) contribution to Permian basins along the southern margin of the Tien Shan–Yin Shan suture in China and argues against exposure of significant Precambrian crust in Mongolia (i.e., the hypothesized South Gobi microcontinent). Finally, the lack of direct evidence for Permian arc activity in southern Mongolia supports south-dipping subduction beneath the North China block, at least during this terminal collision phase. One way of reconciling the simple regional interpretation (of progressive west-to-east suturing) with the complexities indicated by field data is to depart from some aspects of traditional remnant ocean-basin models. According to Ingersoll et al. (1995, p. 363), “A remnant ocean basin is a shrinking ocean

turbidite fan system Altaids, southern Mongolia SE SC Yin Shan

B

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A

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North Tien Shan

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subducting Altaid lithosphere microcontinent crust

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Junggar basin and North Tien Shan, northwest China - Late Carboniferous to Early Permian collision and closure of Junggar trapped ocean basin from east to west during clockwise (?) rotation of Tarim.

B

North Tien Shan--Yinshan suture zone - Collisional zone formed during late Paleozoic closure of the Paleo-Asian ocean. The orogen grows through bidirectional suturing, first westerly with closure of the Junggar ocean basin (Late Carboniferous to Early Permian), then easterly with closure of the Solonker ocean (Early to Late Permian). The orogenic belt comprises mainly uplifted Carboniferous-Devonian arc sequences, which are the main source for sediment delivered by turbidite fan deposits in the closing ocean basins to the west and east.

C

Solonker ocean basin - Early to Late Permian collision between the North China block and the Altaids of southeast Mongolia. Suturing progresses eastward over time due to counterclockwise rotation of the North China block relative to Mongolia. Sediment is mainly derived from the growing orogenic belt to the west (including uplifted Carboniferous-Devonian arc sequences), and distributed in the form of turbidite fan deposits within the trapped ocean basin.

Figure 15. Paleogeographic block diagram for the late Paleozoic of the China-Mongolia border zone, summarizing observed tectonic and stratigraphic relationships discussed in this study. Cartoon is highly schematic, and no specific scale is implied by the features shown. SW, SC, and SE refer to approximate locations of the southwest, south-central, and southeast study areas in Mongolia, discussed in the text (cf. Fig. 14).

Sedimentary response to arc-continent collision basin, which is flanked by at least one convergent margin and whose floor is typically covered by turbidites derived predominantly from associated suture zone(s).” Although this definition broadly applies to the Permian setting in the China-Mongolia border region, many aspects of the model are derived mainly from end-member cases in which significant volumes of sediment are deposited and preserved as turbidite fans—i.e., the Cenozoic Himalayan-Bengal system and the late Paleozoic AppalachianOuachita orogen (Graham et al., 1975). In these case studies, sequential suturing of an irregular continental margin results in transition from fine-grained marine flysch to coarse-grained terrestrial molasse in the direction of progressive collision (Dewey and Burke, 1974). This predictive aspect of the remnant ocean basin model may be broadly supported by the fact that turbidite facies in southeastern Mongolia are mainly distal, outer fan deposits, favoring long-distance funneling of submarine flows outboard of a growing orogen. However, the flysch-molasse transition is not well preserved across the entire suture zone, and this part of the Altaids may be better represented by ongoing intra-arc and microcontinent collisions with small trapped ocean basins, as are found in parts of the modern southwest Pacific Ocean (cf. Fig. 15). Collision and Climate Regional sedimentary facies relationships demonstrate changing climate regimes that accompany collision. In Mongolia, two possible transitions from relatively arid climate to humid climate conditions occur (Fig. 14). The older shift is poorly constrained but seems to have occurred in the middle Permian at Noyon Uul and Tavan Tolgoy, where red shale and sandstone beds generally representing well-drained soils and oxidizing conditions are distinct from overlying Upper Permian coal-bearing units. The younger climatic shift is better constrained in southwest and south-central Mongolia, where coal-bearing Upper Permian strata are overlain by Lower Triassic rocks that lack extensive coal deposits. This is illustrated by a dramatic record of floral transition to droughttolerant organisms at Tsaagan Tolgoy. As noted previously, this sudden shift suggests a correlation with the global warming event at the Permian–Triassic boundary (Retallack et al., 1996; Berner, 2002). However, orographic relief generated by local tectonic uplift may also have influenced such climatic changes. A similar aridification trend, though inferred to be somewhat older, is reported from northeast China, where Carboniferous and Lower Permian humid-climate facies were replaced by arid-climate indicators that started in the Late Permian (Mueller et al., 1991). The relationship is also present in northwest China, based on a Late Permian transition from humid to arid environments (Carroll, 1991; Hendrix et al., 1992; Wartes et al., 2002; Fig. 14). In the Yin Shan belt, the drying event is attributed either to northward latitudinal migration of north China into subtropical realms or to a rain shadow effect following development of high topography during collision with the Altaids (Carroll et al., 1992; Cope et al., 2005).

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Data from Mongolia help to address these hypothesized climatic forcing factors. Regarding the impact of northward plate movement, southern Mongolia already had mixed CathaysianAngaran floral assemblages by Late Permian time, underscoring the influence of flora with Siberian and North China craton affinities. However, paleomagnetic data still indicate a wide latitudinal gap between Siberia and the North China block (and by inference the amalgamated Mongolian arcs) extending into the Cretaceous (Enkin et al., 1992). The sedimentary record in northeast Mongolia does not support such a late collision with the Siberian craton, although Lower Triassic marine rocks are present in northeast Mongolia (Zonenshain et al., 1990; Badarch et al., 2002), indicating that final amalgamation with Siberia likely extended into the early Mesozoic along this remnant ocean (the Mongol-Okhotsk ocean, Fig. 14) to the north of the Tien Shan–Yin Shan suture. As such, continued northward movement of both the North China block and southern Mongolian arcs is still possible, at least into the Triassic. However, the Permian–Triassic climatic shift in southern Mongolia apparently postdates the middle Permian event in China, whereas we would expect an earlier response to the north if caused by such plate movement. Although it is currently unclear how much crustal thickening and orographic relief accompanied collision in the southern Altaids, Davis et al. (2004) have documented metamorphic core complex formation in northeast China at ca. 220–190 Ma. This may indicate collapse of the Permian–Early Triassic collisional orogen and would support the presence of high topography prior to collapse. Cope et al. (2005) suggested creation of a rain shadow, resulting in red bed deposition, starting in the middle to Late Permian. Interestingly, this coincides with Upper Permian coal-bearing deposits in south-central Mongolia (Fig. 14), which could also indicate trapping of moisture on the north side of an inferred mountain range along the Tien Shan–Yin Shan suture. The rain shadow interpretation seems permissible within the range of available age constraints, although we note that turbidite deposition continued in the Late Permian in southeast Mongolia (directly north of the Yin Shan belt), suggesting slightly later collision. In any case, the Early Triassic appears to have been a time of arid climate across the region, which likely reflected the global warming event at the Permian–Triassic boundary, perhaps modified by local topography. IMPLICATIONS If, as is argued here, the Tien Shan–Yin Shan suture links the Tarim-Altaids and the North China block–Altaids collision zones, the entire suture is >3000 km long and thus rivals some of the longest accretionary zones known (e.g., Indian-Eurasia collision, Ouachita-Appalachian orogen). The diffuse northern boundary represented by the Altaids creates a particular challenge in delineating the final suture zone: perhaps the best marker is the southern boundary against known Precambrian-floored blocks, with the understanding that multiple arc systems mark the collision zone to the north. This record seems to be well represented by

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modern accretionary processes in the southwestern Pacific ocean, including the “triple-junction” region of convergence between the Pacific-Philippine, Australia, and Southeast Asia plates (Hinschberger et al., 2005), with several magmatic arcs and microplates occupying the Java, Molucca, and Banda Seas. Although the geometry of this collision zone may not directly apply to the Altaids, it is notable that the oceanic collision zone is marked by multiple arc and backarc basins between major plates that are moving independently, which may be analogous to the Siberian craton, the Tarim–North Tien Shan, and the North China block in the late Paleozoic. Multiple trapped ocean basins are present along the Tien Shan–Yin Shan suture, but evidence is lacking for continuous zipperlike regional closure (Fig. 15). This broad comparison is supported by other modern Pacific arc-collision analogues where climatic effects of orogeny are documented (e.g., a present-day rain shadow effect in southern Papua New Guinea) as well as the evolution to syn- or postcollisional collapse and creation of metamorphic core complexes, as in Taiwan (Clift et al., 2003; Huang et al., 2006). Perhaps the main strength of Permian studies near the China-Mongolia border is that this region provides a good ancient analogue for such modern systems and predicts what the southwest Pacific might look like following terminal collision and closure over the next tens of millions of years. Strike-slip deformation is also an important aspect of the long-term results of these accretionary complexes. The role of syncollisional strike-slip faults in accommodating rotational closure and oblique collisions is noted in modern analogs (Moore and Silver, 1983; Hinschberger et al., 2005; Huang et al., 2006). Late Paleozoic strike-slip deformation has been documented in parts of the Altaids (S¸engör and Natal’in, 1996; Laurent-Charvet et al., 2003; Zhou et al., 2001) but has not yet been well documented in the Permian of Mongolia. Reconstruction of syncollisional paleogeography is complicated by later strike-slip deformation that has only recently been recognized in southern Mongolia. At least two phases of left-lateral movement are documented along the East Gobi fault zone (Figs. 1, 2). The first phase occurred in the latest Triassic (ca. 207–209 Ma) associated with a northeasttrending mylonite zone that is continuous for >250 km (Lamb et al., 1999). This was followed by Late Cretaceous–Cenozoic reactivation via brittle sinistral faulting (Johnson, 2004). The amount of offset associated with these sinistral movements is poorly constrained but likely represents ~200–400 km of total slip (Webb and Johnson, 2006). One important implication of this postcollisional deformation is that attempts to reconstruct the southern margin of the Altaids by conducting mainly west-to-east correlations may incorrectly assume that these units were originally formed at the same latitude. For example, turbidite deposits in the Nomgon area may have equivalents along the southernmost border of Mongolia at Bulgan Uul (Figs. 2, 14; Lamb et al., 1999), and thus the record of continued marine deposition may extend farther westward than currently realized. It is common for convergent margins to be reactivated as transform or strike-slip settings from shifting convergence directions along a plate margin (e.g., Cenozoic of California; Graham,

1978) or from intracontinental deformation within an accretionary complex (Tapponnier et al., 1982). The latter seems to have been the case in Mongolia, where these two phases of sinistral motion have been tentatively linked to post-Paleozoic collisions along the southward-growing margin of Asia, including the Cenozoic India-Asia collision >2000 km from the East Gobi fault zone. Cenozoic strike-slip is also demonstrated in southwestern Mongolia (Cunningham et al., 2003). Rather than following a single terrane boundary, this intracontinental fault system seems to exploit multiple zones of weakness within heterogeneous crust of the Altaids. Thus consideration of syn-and postcollisional strike-slip deformation is critical to reconstruction of these collisional zones, and, conversely, better constraints on the location and nature of arc-collision zones help to predict modes and locations of subsequent deformation. CONCLUSIONS Stratigraphic data presented here are a first step toward unraveling the history of arc-arc and ultimately arc-continent collision that is dramatically represented in the southern Altaids, particularly from north of the China-Mongolia border. Future studies that attempt to further constrain the timing and modes of continental growth by accretion, both in this part of Asia and in analogous settings worldwide, should take into account several key constraints afforded by this study. Regional relations that extend some 3000 km from northwest China, through south Mongolia, to northeast China, argue for broad correlation of the Tien Shan– Yin Shan sutures. The main suture zone is best drawn along the northern margin of known continental blocks; tectonic boundaries that extend into the Altaids represent only extensions of this broad collision zone into a heterogeneous accretionary margin. Generally south-dipping subduction during the Permian is supported by the lack of active Permian arcs in southern Mongolia, and by complementary studies of active Permian arc systems in north China. In very general terms, west-to-east diachronous closure may be argued mainly on the basis of the age of the youngest marine strata from the Tien Shan to the Yin Shan (from Upper Carboniferous to Upper Permian). However, paleogeographic reconstructions based on sedimentary facies and paleocurrent data, sandstone petrofacies, and dating of magmatic activity indicate that the collisional event was not a simple progressive suturing process. There is no record of progressive arc dissection in the Permian from southwest to southeast Mongolia, nor is there a continuously eastward-younging flysch to molasse transition. Sandstone provenance studies argue for very little input from continental crust or recycled orogen sources into Permian basins of Mongolia, in contrast to analogous deposits in north China. Final arc-continent accretion may have included both west- and east-facing trapped oceans and differential rotation of the North China block and Tarim, possibly driven by uneven colliding margins with indenters that are currently poorly constrained. Thus Permian strata of southern Mongolia are best considered an endmember representation of published models for remnant ocean

Sedimentary response to arc-continent collision basin closure. This record also provides a predictive glimpse into future results of arc-arc and arc-continent collisions in the present southwest Pacific Ocean. The possible implications of this collisional event are significant. Regional climatic shifts are linked to evolution of the accretionary belt, including the possibility of a rain shadow effect south of the collisional orogen, which may have subsequently collapsed during the Triassic. Similar impacts are reported in modern arc-continent collisions such as Taiwan and Papua New Guinea. Finally, strike-slip deformation is an important syncollision mechanism to consider as a way of accommodating rotational closure. The recognition of subsequent strike-slip deformation is also critical to reconstruction of the suture zone, which in southern Mongolia is now partly offset and dismembered along the East Gobi fault zone. It is likely that multiple phases of post-Permian movement along this fault system were driven by continued collisions along the growing margin of Asia, and thus the long-term record of intracontinental deformation is also tied to this region as an inherited zone of crustal weakness. ACKNOWLEDGMENTS This work was supported by U.S. National Science Foundation grants to Graham (EAR-9614555 and EAR-9708207) and Johnson (EAR-0537318, Tectonics and OISE, L.E. Webb, coPI), and funding from the American Chemical Society (Petroleum Research Fund 40193-G8 to Johnson). Funds from Stanford University also supported this work, including the Graduate Fellowship program, the McGee Fund, and the Stanford Undergraduate Research Opportunity Grant. The American Association of Petroleum Geologists and the Geological Society of America supported field seasons in Mongolia through several student grants to the authors. We thank Ch. Minjin and colleagues at the Mongolian University of Science and Technology for their support, and M.S. Hendrix and L.E. Webb for many years of collaboration. D. Zinniker thanks S. Fowell and G. Norris for assistance in paleopalynology. Countless field assistants and colleagues assisted in field work, including A. Chimitsuren, J. Crider, M. Heumann, A. Keller, T. Hickson, N. Manchuk, G. Sersmaa, E. Sobel, and J. Undariya. We appreciate thoughtful and thorough reviews by A. Carroll, R. Dorsey, M. Faure, and P. Clift, which improved the final manuscript. This study is dedicated to our colleague, advisor, and dear friend Gombosuren Badarch, who died during preparation of this manuscript. This work would not have been possible without Badarch’s guidance, expertise, and patience. We miss him deeply, but his love for Gobi field work and the geology of Asia continues to inspire us. REFERENCES CITED Allen, M.B., Windley, B.F., Zhang, C., Zhao, Z., and Wang, R.-G., 1991, Basin evolution within and adjacent to the Tien Shan Range, NW China: Geological Society [London] Journal, v. 148, p. 369–378.

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Allen, M.B., Windley, B.F., and Zhang, C., 1993, Palaeozoic collisional tectonics and magmatism of the Chinese Tien Shan, central Asia: Tectonophysics, v. 220, p. 89–115, doi: 10.1016/0040-1951(93)90225-9. Amory, J., 1996, Permian sedimentation and tectonics of southern Mongolia [M.S. thesis]: Stanford, California, Stanford University, 183 p. Amory, J., Hendrix, M., Lamb, M., Keller, A., Badarch, G., and Tomurtogoo, O., 1994, Permian sedimentation and tectonics of southern Mongolia; implications for a time-transgressive collision with North China: Geological Society of America Abstracts with Programs, v. 26, no. 7, p. 242. Anatoleva, A.I., 1974, The structure and composition of the Permian red volcanogenic-sedimentary complex in southern Mongolia: Geologiya i Geofizika, v. 15, p. 32–43. Aristov, D.S., 2005, New Grylloblattids (Insecta: Gryllobattida) from the Triassic of Eastern Europe, Eastern Kazakstan, and Mongolia: Paleontological Journal, v. 39, p. 173–177. Badarch, G., Cunningham, W.D., and Windley, B.F., 2002, A new terrane subdivision for Mongolia; implications for the Phanerozoic crustal growth of Central Asia: Journal of Asian Earth Sciences, v. 21, p. 87–110, doi: 10.1016/S1367-9120(02)00017-2. Battacharya, J.P., and Walker, R.G., 1992, Deltas, in Walker, R.G., and James, N.P., eds., Facies Models: St. Johns, Newfoundland, Geological Association of Canada, p. 157–178. Baud, A., Magaritz, M., and Holser, W.T., 1989, Permian–Triassic of the Tethys; carbon isotope studies: Geologische Rundschau, v. 78, p. 649–677, doi: 10.1007/BF01776196. Berkey, C.P., and Morris, F.K., 1927, Geology of Mongolia, in Natural History of Central Asia, Volume 2: New York, American Museum of Natural History, 475 p. Berner, R.A., 2002, Examination of hypothesis for the Permo–Triassic boundary extinction by carbon cycle modeling: Proceedings of the National Academy of Sciences of the United States of America, v. 99, p. 4172– 4177, doi: 10.1073/pnas.032095199. Burchfiel, B.C., Deng, Q., Molnar, P., Royden, L., Wang, Y., Zhang, P., and Zhang, W., 1989, Intracrustal detachment within zones of continental deformation: Geology, v. 17, p. 748–752, doi: 10.1130/00917613(1989)017<0448:IDWZOC>2.3.CO;2. Carroll, A.R., 1991, Late Paleozoic tectonics, sedimentation, and petroleum potential of the Junggar and Tarim basins, northwest China [Ph.D. thesis]: Stanford, California, Stanford University, 405 p. Carroll, A.R., Liang, Y., Graham, S.A., Xiao, X., Hendrix, M.S., Chu, J., and McKnight, C., 1990, Junggar basin, northwest China: Trapped Late Paleozoic ocean: Tectonophysics, v. 181, p. 1–14, doi: 10.1016/00401951(90)90004-R. Carroll, A.R., Brassell, S.C., and Graham, S.A., 1992, Upper Permian lacustrine oil shales, southern Junggar basin, northwest China: American Association of Petroleum Geologists Bulletin, v. 76, p. 1874–1902. Carroll, A.R., Graham, S.A., Hendrix, M.S., Ying, D., and Zhou, D., 1995, Late Paleozoic tectonic amalgamation of northwestern China: Sedimentary record of the northern Tarim, northwestern Turpan, and southern Junggar basins: Geological Society of America Bulletin, v. 107, p. 571–594, doi: 10.1130/0016-7606(1995)107<0571:LPTAON>2.3.CO;2. Chen, B., Jahn, B.-M., Wilde, S., and Xu, B., 2000, Two contrasting Paleozoic magmatic belts in northern Inner Mongolia, China: Petrogenesis and tectonic implications: Tectonophysics, v. 328, p. 157–182, doi: 10.1016/ S0040-1951(00)00182-7. Clift, P.D., Schouten, H., and Draut, A.E., 2003, A general model of arc-continent collision and subduction polarity reversal from Taiwan and the Irish Caledonides, in Larter, R.D. and Leat, P.T., eds., Intra-oceanic Subduction Systems: Tectonic and Magmatic Processes: Geological Society [London] Special Publication 219, p. 81–98. Coleman, R.G., 1989, Continental growth of northwest China: Tectonics, v. 8, p. 621–635. Cope, T., Ritts, B.G., Darby, B.J., Fildani, A., and Graham, S.A., 2005, Late Paleozoic sedimentation on the northern margin of the North China Block; implications for regional tectonics and climate change: International Geology Review, v. 47, p. 270–296. Cunningham, D., Dijkstra, A.H., Howard, J., Quarles, A., and Badarch, G., 2003, Active intraplate strike-slip faulting and transpressional uplift in the Mongolian Altai in Storti, F., Holdsworth, R.E., and Salvini, F., Intraplate Strike-slip Deformation Belts: Geological Society [London] Special Publication 210, p. 65–87.

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The Geological Society of America Special Paper 436 2008

Links among mountain building, surface erosion, and growth of an accretionary prism in a subduction zone—An example from southwest Japan Gaku Kimura* Department of Earth and Planetary Science, University of Tokyo, Japan, and Institute for Frontier Research on Earth Evolution, Japan Agency for Marine-Earth Science and Technology, Yokosuka, Japan Yujin Kitamura Department of Earth and Planetary Science, University of Tokyo, Japan, now at IFM-GEOMAR, Kiel, Germany Asuka Yamaguchi Hugues Raimbourg Department of Earth and Planetary Science, University of Tokyo, Japan

ABSTRACT The relationships between mountain building, surface erosion, sediment supply to the trench, and growth of the accretionary prism are examined in southwest Japan and the Nankai Trough. Mountain building caused by the subduction of the Philippine Sea plate in the Nankai Trough and collision in central Japan has resulted in a rock uplift rate of ~4 mm/yr. Surface denudation rates in the mountain regions are on the order of 3–4 mm/yr, resulting from the heavy rainfall of the Asian monsoon. This fact suggests that mountain building is almost in an equilibrium stage in which surface erosion and rock uplift balance each other, resulting in a constant altitude of ~2000 m. Several drainage systems on land and in offshore submarine canyons enable the transport of eroded sediments directly into the Nankai Trough. Most of the terrigenous sediments supplied to the Nankai Trough are accreted in the subduction zone of the Philippine Sea plate. The accretion rate of the sediments in the eastern Nankai Trough is ~1.68 × 107 m3/yr, which is consistent with the denudation rate of the Akaishi Mountains, contributing to the supply of 1.72 × 107 m3/yr of sediment in central Japan. The growth of the accretionary prism is an important controlling factor for the onset of large earthquakes in the Nankai Trough, because the hanging wall of the rupture area of the seismogenic zone is composed entirely of the accretionary prism. Repeated large earthquakes with a recurrence time of ~100–200 yr, which are well recorded in the Nankai Trough, in turn promote surface erosion through consecutive landslides and tsunamis.

*[email protected] Kimura, G., Kitamura, Y., Yamaguchi, A., and Raimbourg, H., 2008, Links among mountain building, surface erosion, and growth of an accretionary prism in a subduction zone—An example from southwest Japan, in Draut, A.E, Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 391–403, doi: 10.1130/2008.2436(17). For permission to copy, contact editing@ geosociety.org. ©2008 The Geological Society of America. All rights reserved.

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Kimura et al. Southwest Japan, with its extensive record of both erosional processes and seismic events, shows the intimate long-term relationships between tectonically driven mountain building, surface erosion under the Asian monsoon climate, growth of the accretionary prism in the trench, and the generation of large earthquakes. Keywords: Nakai Trough, Asian monsoon, denudation, accretionary prism, large earthquake.

INTRODUCTION The growth of accretionary prisms is controlled mainly by the sediment supply into the trench, which depends on surface erosion on land and on the drainage system down to the trench. The development of the drainage system from land to trench and active surface erosion are necessary for the development of an accretionary prism. Surface erosion on land is especially intense in mountain regions under climatic conditions that include heavy rainfall, such as tropical to subtropical rainy mountains (e.g., in southeastern Asia) or high-latitude glacier-covered mountains (e.g., Alaska, Aleutian Trench, or southernmost regions of South America). Mountain building in these regions results from plate convergence associated with continent-continent, continent-arc, or arc-arc collisions, or intracontinental shortening. Southwest Japan is a particularly appropriate context for analysis of the complex interplay among active mountain building on land, active surface erosion from the Asian monsoon climate, and the resulting development of a huge accretionary prism in the Nankai Trough (Taira et al., 1980, 1988). Taira et al. (1988) and Taira and Niitsuma (1986) emphasized that effective drainage systems from the central mountain region to the Nankai Trough in Japan have contributed to the buildup of the Nankai accretionary prism. The mountain building in central Japan is the tectonic result of the collision between the Izu-Bonin Arc and the main Honshu Arc, and the collision between the southwest Japan and northeast Japan Arcs (Sugimura, 1972; Huzita, 1980). The Asian monsoon brings heavy rain to Japan every year during the spring rainy season, called Tsuyu, which means continuous rain. In addition, typhoons attack the Japanese islands, usually at the end of summer. This climatic condition enhances the surface erosion of the mountains, and eroded debris and sediments drain into the rivers down to the ocean. Frequent earthquakes also might promote surface erosion. Earthquakes in the Nankai Trough have repeatedly taken place during historic time (Ando, 1975). Most of the earthquakes in the Nankai Trough are larger than Mw = 8. Ruff and Kanamori (1980) suggested that the occurrence of large earthquakes in the subduction zone might be related to the development of an accretionary prism in the trench. Thus, mountain building, heavy surface erosion, development of the accretionary prism, and large earthquakes in the Nankai Trough might ultimately be linked to one another. In this paper, we examine the linkage in southwest Japan and discuss the relationships among these geologic processes.

TECTONIC SETTING OF SOUTHWEST JAPAN AND MOUNTAIN BUILDING IN CENTRAL JAPAN The tectonic setting of central to southwest Japan has resulted from the convergence of four plates or blocks, and it is commonly decomposed into three distinct layers (Fig. 1): the top layer is a conjunction of the Amurian plate of southwest Japan and the northeast Japan block, which have collided with each other in central Japan; the second layer is the Philippine Sea plate; and the lowest layer is the Pacific plate, which is subducting beneath the above two layers. The upper plate of the Japanese islands is composed mainly of volcanic arcs and ancient accretionary complexes, whose rigid parts consist only of upper brittle crust without rigid mantle lithosphere. As the island arcs are strongly deformed, they are separated into several microblocks by large faults or tectonic discontinuities. In such cases the distinction between plate boundaries and intraplate faults is subtle, and the definition of plates differs from one author to another, from North America (Nakamura, 1983), to Okhotsk (Savostin and Karasik, 1981; Kimura and Tamaki, 1986), to the northeast Japan plate or block (Seno, 1985), to the northeast Japan Arc, to Eurasia (Nakamura, 1983), to the Amurian (Savostin and Karasik, 1981; Kimura and Tamaki, 1986) plate, to the southwest Japan Arc (Fig. 1). Whether the northeast and southwest Japan Arcs are plates or microblocks, they are separated by a large fault in central Japan (Fig. 1; Nakamura, 1983), and ongoing convergence between the two is responsible for mountain uplift of the Japan Alpine Mountains in central Japan. The convergence and collision of these two arcs began in Pliocene time (Tsunakawa and Takeuchi, 1986) and accelerated at ca. 2 Ma (Kaizuka, 1975; Huzita, 1980; Nakamura and Uyeda, 1980; Tsunakawa and Takeuchi, 1986). The southern part of central Japan is also the collision zone between the Izu-Ogasawara Arc on the Philippine Sea plate and the northeast Japan Arc (Sugimura, 1972; Fig. 1). Since the onset of collision at ca. 15 Ma the arc crust has continuously flaked and accreted to the upper plate of central Japan (Takahashi, 2006). The Izu Peninsula is a colliding part of the Philippine Sea plate, and the intracrustal breakage of the peninsula has already started because of the collision. The Pacific side of the southwest Japan Arc (Fig. 1) is composed mainly of Jurassic, Cretaceous, and Tertiary accretionary complexes, which are exposed parallel to the Nankai Trough. The complexes in central Japan are bent in a northward-convex shape. This bending is the result of the continuous collision and accretion of the Izu-Ogasawara Arc (Matsuda

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and Uyeda, 1971). The main part of the Philippine Sea plate is now subducting beneath southwest Japan along the Nankai and Sagami Troughs at a rate of 4–6 cm/yr (Seno, 1989). The Pacific plate is subducting beneath the northeast Japan Arc at ~10 cm/yr along the Japan Trench. A trench-trench-trench type of triple junction lies to the southeast of central Japan at the southeastern end of the Sagami Trough. UPLIFT RATE OF THE MOUNTAINS IN CENTRAL AND SOUTHWEST JAPAN

system (GPS), put in place after the 1995 Hanshin-Awaji earthquake, has enabled highly detailed measurements over the past 10 yr (Sagiya, 2004). The century-scale record shows that uplift in central Japan is concentrated in mountain regions (Figs. 2, 3), especially in the southern part—the Akaishi Mountains—which lie to the west of the Itoigawa-Shizuoka Tectonic Line and the Izu Peninsula. The mountains of the Kii Peninsula, and the Shikoku and Kyushu island regions in southwest Japan, also have undergone some uplift, especially in their southern parts (Figs. 2, 3). The most uplifted regions are parallel to the Nankai Trough but are not continuous. They actually alternate with subsiding regions characterized by flat plains, channels, and bays, resulting in the following pattern from east to west: the Akaishi Mountains (uplift), Ise Bay (subsidence), Kii Mountain in the Kii Peninsula

The present crustal movement of the Japanese islands has been precisely recorded with modern geodetic observations within eight distinct periods from 1883 to 1999 (Kunimi et al., 2001). In addition to these observations, the global positioning

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(uplift), Kii channel (subsidence), eastern Shikoku Mountains (uplift), Tosa Bay (stable), and the western Shikoku Mountains in Kyushu (uplift). This geometry is characteristic of 10-km-scale crustal folds in the along-arc orientation and is explained as a superimposed compression that resulted from two different convergence directions owing to the subduction of the Philippine Sea plate along the Nankai Trough and the collision in central Japan (Huzita, 1980). The Seto Inland Sea to the north of the island of Shikoku (Fig. 2) is a subsiding region. This aspect was explained by the same tectonic origin of crustal-scale folding, superposition caused by the subduction of the Philippine Sea plate, and the collision in central Japan similar to that in the outer arc of southwest Japan (Kaizuka, 1975; Huzita, 1980). The recent compilation of geodetic observations by GPS (Sagiya, 2004) shows almost the same vertical crustal movement as was recorded earlier by long-term geodetic observations

(Fig. 3B) except for the southern tip of the outer arc. Mountain regions in the outer arc are generally uplifted, but their southern parts in Capes Omaezaki in Tokai, Shionomisaki in Kii, and Muroto in Shikoku, from east to west, are undergoing subsidence (Fig. 2B). This subsidence has been explained as an interseismic phenomenon between the large earthquakes in 1944 in Tonankai and in 1946 in Nankai, and a near-future one predicted for the twenty-first century (Sagiya, 2004). The uplift rate in the mountain region has reached a maximum of 4 mm/yr (Fig. 2) in the Akaishi Mountains and other nearby areas. SURFACE DENUDATION RATE OF THE MOUNTAINS IN CENTRAL JAPAN The relationship between sediment production and tectonic uplift in the Japanese islands, and how the climate contributes to sediment production, were thoroughly studied by Yoshikawa

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(1974) and Ohmori (1978). Their contribution is a milestone in this research field. We briefly introduce their results in this section and discuss them in light of data acquired from geographic information system (GIS) analyses in central Japan. The denudation rate of the mountain region was estimated from the amount of accumulated sediments in the water reservoirs, which were constructed by complete damming of the primary river system (the mainstream system important to the nation as defined by the Ministry of Land, Infrastructure and Transport of the Japanese Government) that drains the mountains. As the volume of small clayey particles and solution chemicals from erosion is likely to be smaller by several orders of magnitude than that of solid accumulation in the reservoirs, Ohmori (1978) ignored such components for a first-order approximation. The period studied by Ohmori (1978) is also important. Many small dams were constructed in upstream regions in the tributaries of the primary rivers since the 1970s to prevent landslides and debris flows in the mountains. This means that sediment accumulation in the large reservoirs can no longer be easily related to the denudation rate of the mountains. Thus, the data set collected in the 1960s is the only available quantitative measurement of the natural erosion rate in the mountains in the Japanese islands. Ohmori (1978) clarified that an empirical relationship between the mean altitude of mountains and sediment production is expressed by

SDR = W × H a , where SDR = the mean annual sediment delivery rate (m3/km2/ yr), H = the mean altitude of mountains (m), W = 0.4642−3 (/yr), and a = 2.1894. Thus SDR is almost a quadratic function of the mean altitude. The denudation rate (DR) is obtained from

DR = SDR ×1.75 /2.68 ×10−6 , where 1.75 and 2.68 are the mean bulk densities of sediments and bedrocks, respectively. Then, the relationship between the denudation rate and the mean altitude is represented by

DR = 0.3031×10−9 × H 2.1894 . As reviewed in the previous section, most of the mountains in Japan are undergoing tectonic uplift. The uplift started in the Pliocene and accelerated at ca. 2 Ma as described. Geologically long-term uplift is a result of tectonic uplift (U) minus surface erosion. A geological long-term surface uplift rate dH/dt might be expressed by

dH /dt = U − DR = U − 0.3031×10−9 × H 2.1894 . Integrating the function from a mean altitude h1 to another mean altitude h2,

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This is a law of diminishing returns, which implies that the mountains are equilibrated at some altitude at some time t. In the case of 4 mm/yr of U, ~1 m.y. is enough to reach the equilibrium stage at which surface erosion and uplift balance each other, resulting in a constant altitude of ~2000 m. Figure 4B shows that most of the Akaishi Mountains are steeper than 30°, which is the critical slope angle consistent with internal friction of crustal rocks. Slopes steeper than 30° are not stable and collapse easily by gravitational sliding. By multiplying the denudation rate (DR) obtained from the local stream system (Ohmori, 1978) by the surface of the land whose altitude is higher than 1000 m, a total solid erosion rate is obtained. We reanalyzed the surface areas of the Akaishi Mountains higher than 1000 m, which amounted to 4.631 × 106 m2, obtained from GIS analysis. Applying the DR from the river systems in the Akaishi Mountains in central Japan, an erosional solid volume of ~1.72 × 107 m3/yr is estimated (Table 1). Ohmori (1978) pointed out that the lithology of the mountains and location-controlled climatic differences (such as snowy mountains in the north or rainy mountains in the south within the Japanese islands) are not strong factors in denudation rates. This is important, because climatic conditions during the Last Glacial Maximum were different from those of the present. Sea level was much lower, and the land area was much greater, than in the present, but low-altitude mountains might have not contributed much to surface erosion, as suggested by the SDR-H function described above. The present Japanese islands widely cover the climatic zones of the Asian monsoon from south to north, but Ohmori (1978) points out the absence of serious differences in the denudation rate in northern and southern Japan. DRAINAGE SYSTEMS TO THE TRENCH FLOOR OF THE NANKAI TROUGH The sediments from the uplifting mountains have not been directly transported into the trench in many places but are instead trapped in the forearc, e.g., on continental shelves, in forearc basins, and in piggyback slope basins on the accretionary prism. In the eastern part of the Nankai forearc off the Tokai district, the main submarine canyons—Suruga Trough and Tenryu Canyon— directly connect the on-land river systems to the trench floor of the Nankai Trough (Fig. 5). The Sionomisaki Submarine Canyon from the Kii Peninsula is the main drainage canyon in the middle part of the Nankai Trough. Other, rather small canyons connect the Shikoku and Kyushu islands with the Nankai Trough (Fig. 5). These canyons are topographic manifestations of faults that cut the accretionary prism (Tokuyama et al., 1999). Detailed topography of the mouth of these canyons shows that submarine fans (Tokuyama et al., 1999) composed of turbidites (Ike et al., 2005) pass through the canyon channels directly from the land (Fig. 6).

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TABLE 1. DENUDATION RATE* AND SURFACE AREA HIGHER THAN 1000 m ALTITUDE IN THE AKAISHI MOUNTAINS IN CENTRAL JAPAN River system

Erosion of solids 6 3 (×10 m /yr) 6.04 4.92 2.08 0.19 4.01

Total erosion of solids 17.24 Note: Denudation rate and each river system are shown in Figure 4. *(Ohmori, 1978)

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Surface area 6 2 (10 m ) 1,678 2,191 636 55 2,125

Denudation rate –3 (10 m/yr) 3.598 2.247 3.267 3.435 1.887

The eastern parts of the Nankai Trough are filled by turbidites, which are apparent on seismic profiles (Park et al., 2002a; Ike et al., 2005) most of which originate from the canyons mentioned above (see Fig. 7). The trench floor of the Nankai Trough gently deepens from east to west owing to subduction of the topographic highs in the eastern part (Tokuyama et al., 1999). Therefore, the turbidites flow down to the west. The total volume of solid sediments draining directly into the trench per year, passing through these channels, is ~1.44–1.91 × 107 m3/yr, as estimated from the solid volume of denudation in central Japan, given in the previous section.

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Figure 5. Map showing mean sediment delivery rate (SDR) to the reservoirs in central to southwest Japan. Black and gray solid circles indicate the river systems draining into the Pacific Ocean and Japan Sea, respectively. Numbers in circles represent the volume (m3/km2/yr), modified from Ohmori (1978).

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ACCRETION RATE OF THE SOLID VOLUME IN THE NANKAI ACCRETIONARY PRISM The mode of accretion changes at the trench-slope break from an in-sequence and out-of-sequence thrusting mode to an underplating-dominant mode (Kimura et al., 2007). Terrigenous sediments supplied from the mountainous region on land dominantly return back to the land owing to frontal accretion in the Nankai Trough (e.g., Moore et al., 2001; Kimura et al., 2007). In the case of the Nankai Trough the critical taper condition (Davis et al., 1983) of the accretionary prism is operated trenchward from the trench-slope break (Saffer and Bekins, 2002; Wang and Hu, 2006; Kimura et al., 2007), and the accretionary-wedge

taper angle varies between 2° and 11°, depending on hydrological conditions within the wedge and basal décollement (Saffer and Bekins, 2002; Wang and Hu, 2006; Kimura et al., 2007) (Fig. 7). The frontal accretion rate Vt per unit length along the trench is estimated from

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and zd = 1200 m is the averaged thickness of trench-filling sediments involved in frontal accretion in the eastern half of the Nankai Trough (Moore et al., 2001; Park et al., 2002b). We obtain Vt = 47.95 m2/yr. This value is the growth rate of the area of the profile of the accretionary prism normal to the trench axis. The distance of the accretionary prism along the trench is ~500 km from the mouth of the Suruga Trough to the deformation front off Shikoku. The results of ocean drilling off Shikoku show that the trench-filling sediments are dominated by turbidites directly derived from the land, as mentioned above (Taira and Niitsuma, 1986; Taira et al., 1992). However, large amounts of terrigenous sediments are also supplied to the trench through Shionomisaki Canyon and other canyons to the west, diluting the sediment transported from the northeast. Because it is difficult to quantify the contributions from various sources, we consider only turbidites accreted in the trench to the east of the mouth of Shionomisaki submarine canyon, along ~300 km of the trench (Fig. 5). Using this value, the total volumetric accretion rate is ~1.68 × 107 m3/yr. Such a rapid rate of accretion suggests that the accretionary prism outboard of the trench-slope break at ~30 km from the deformation front is composed dominantly of Quaternary turbidites. This is verified by direct drilling into the accretionary prism (Moore et al., 2001) and by samples using submersible dives (J. Ashi, 2006, personal commun.).

DISCUSSION We quantitatively compiled data on the present mountain building, surface erosion, and growth of the accretionary prism in southwest Japan and the Nankai Trough. Our treatment is just a first-order approximation and needs further development, but these relationships are important toward understanding the relationship and feedback among these processes, which have been studied separately until now. We discuss in the following section how present relationships can be extrapolated to longterm evolution. Relationship between Long-Term and Short-Term Linkages The denudation rate in the mountains reviewed by Ohmori (1978) is averaged over ~10 yr, and the uplift rate in terms of geodetic observation is averaged over 10–100 yr. It is questionable whether these rates are applicable to geologic time scales of millions of years. Was the denudation rate accelerated or decelerated during the most recent ice age? It is well known that high mountains in Japan were capped by glaciers during the last ice age (Minato et al., 1965). Such climatic conditions appear to have enhanced surface erosion and thus have produced many more sediments

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than in the present day, as a wide consensus concludes that average denudation rates have been higher than normal since the start of northern hemispheric glaciation ca. 2.7 Ma. During the last ice age the average annual temperature in the Japanese islands was ~7–8 °C lower than that at present. Ohmori (1978) discussed the difference in denudation rate between the mountains in northern Japan from the example of Hokkaido, the northernmost island of Japan, and those in southwest Japan. Their locations are separated in latitude by >10°, and the difference in average annual temperature is ~10 °C at present. Ohmori’s conclusion is that the denudation rates are almost the same in northern, central, and southern Japan, as reviewed. Climatic conditions in central and southwest Japan are different from those in northern Japan. A heavy rainy season is common in central and southwest Japan but not in northern Japan. In contrast, northern Japan has a heavy snowy winter, whereas southwest Japan does not have such a severe winter. Such climatic factors may compensate each other, resulting in similar denudation rates. Such present analysis of the denudation rate in the Japanese islands suggests that a much cooler climate during the last ice age might not have resulted in higher denudation rates than at present. The East Asian summer monsoon rains were much weaker in glacial times, and because rainfall is a major control on erosion rates (Reiners et al., 2003; Dadson et al., 2003) this means that denudation rates during the last glacial maximum likely were much lower than in the present interglacial. Thus, the wide consensus is that average denudation rates higher than normal during an ice age should be reexamined on a global scale. Another factor controlling denudation rate is sea level, which was much lower during the last ice age (Fairbanks, 1989). Ohmori (1978) pointed out that the altitude of the mountains is the primary factor controlling the erosion rate and the lowest part of land does not seriously contribute to sediment production but just affects transportation and accumulation. All these observations by Ohmori (1978) suggest that the denudation rate during the last ice age was not very different from the present interglacial. The intensification of the Asian monsoon controlled by global-scale atmospheric currents can play back to the beginning of the Quaternary Period. In addition, the present tectonic framework of the Japanese islands also dates back to the beginning of the Quaternary (Huzita, 1980; Honza and Fujioka, 2004). Thus the linkage among mountain building, denudation rates, and sediment supply into the trench might be applicable at least to the Quaternary Period. This hypothesis is to be tested by studying the sedimentation history in the Nankai Trough and the northern Shikoku Basin, where drilling in the framework of the Integrated Ocean Drilling Program will soon operate.

plate smooth the surface and make a widely coupled asperity in the seismogenic zone. The rupture areas of the 1944 Tonankai and 1944 Nankai earthquakes lie beneath and within the accretionary prism (e.g., Park et al., 2002b), which developed from accretion of sediment transported into the Nankai Trough since the late Miocene Epoch. What controls the size of the rupture area of large earthquakes in subduction zones is a complicated question. One of the hypotheses involves sediment subduction: Subducted sediments not only smooth the surface as suggested by Ruff and Kanamori (1980), but water is released from sediments, especially those that are clayey, dynamically lubricating the plate-boundary fault for rupture propagation (e.g., Kimura et al., 2007). Most of the terrigenous turbidites are accreted at the toe of the accretionary prism, but old pelagic to hemipelagic sediments underthrust, and their lithification works for the onset of the seismogenic zone (Moore and Saffer, 2001; Kimura et al., 2007). The dynamic weakening resulting from lubrication of the fault might be the most likely mechanism to propagate the rupture area, and the size of the fluid-rich area might control the magnitude of earthquakes. In turn, the shaking of the land and tsunamis caused by large earthquakes enhance landslides in the mountainous areas and promote surface erosion.

Onset of Large Earthquakes, the Seismogenic Zone, and Growth of the Accretionary Prism

Long-Term Feedback System among Tectonic, Sedimentary, and Climatic Processes

The growth of the accretionary prism is important for the onset of the rupture of large earthquakes in subduction zones (Ruff and Kanamori, 1980). Sediments covering the subducting oceanic

A generic relationship among tectonic, sedimentary, and climatic processes is shown in Figure 8. An important link between the tectonic and sedimentary processes is the enhancement of

Accretion versus Subduction Erosion One of the unsolved but important questions in the research of subduction zones concerns the balance between accretion and subduction erosion (e.g., von Huene and Scholl, 1991; Clift and Vannucchi, 2004). The review in this paper suggests that the amount of underthrust sediments, even in the frontal part of the prism, might be one or two orders of magnitude smaller than accreted sediments. Only the distal parts of turbidites transported far from the trench and deposited as hemipelagic sediments will come back to the trench and will underthrust, together with biogenic and chemically precipitated pelagic sediments overlying oceanic basement of the subducting plate. Subducting sediments are much smaller in volume because underplating takes place beneath the frontally accreted sediments. Thus the amount of sediment subduction at the toe of the deformation front in the eastern Nankai Trough might be 1–10 × 105 m3/yr along a unit kilometer along the trench, because the difference between the sediment supply rate of 1.72 × 107 m3/yr and the accretion rate of 1.68 × 107 m3/yr is the amount that should subduct along 300 km of the trench axis. This value is one or two orders of magnitude smaller than the value estimated for the erosional margins (von Huene and Scholl, 1991; Clift and Vannucchi, 2004), although the value is estimated without taking into account the amount of basal erosion or underplating.

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Tectonic process

Sedimentary process

Climatic process

Surface erosion

Glacier Heavy rainfall

Collision Mountain building

Sedimentary transport Accretionary prism Large earthquake >Mw=8.0

Trench filling

Figure 8. Diagram showing linkage among tectonic, sedimentary, and climatic systems.

Shaking & Tsunami Smooth the plate interface? Fluid induced rupture propagation/lubrication?

Change the climate/ocean current system

surface erosion owing to mountain uplift and the resulting growth of the accretionary prism owing to sediment transport into the trench. The growth of the accretionary prism is a key for largeearthquake generation in subduction zones, as discussed above, and then strong shaking and tsunamis further enhance surface erosion in the mountains on land. Thus, tectonic events and sediment production are interrelated at long-term geological scales. ACKNOWLEDGMENTS This study was supported by the Plate Dynamics Program of the Japan Agency for Marine-Earth Science & Technology, and the 21st Century Center of Excellence Program of the University of Tokyo. Discussions with G.F. Moore, Y. Kaneda, T. Shibata, S. Okamoto, and T. Ike were quite useful. Reviews by A. Draut, C. Wobus, P. Clift, and G.F. Moore improved the early draft. We appreciate all these persons. REFERENCES CITED Ando, M., 1975, Source mechanisms and tectonic significance of historical earthquakes along Nankai Trough, Japan: Tectonophysics, v. 27, p. 119– 140, doi: 10.1016/0040-1951(75)90102-X. Bray, C.J., and Karig, D.E., 1985, Porosity of sediments in accretionary prisms and some implications for dewatering processes: Journal of Geophysical Research, v. 90, p. 768–778. Clift, P., and Vannucchi, P., 2004, Controls on tectonic accretion versus erosion in subduction zones: Implications for the origin and recycling of the continental crust: Reviews in Geophysics, v. 42, no. RG2001. Dadson, S.J., Hovius, N., Chen, H.G., Dade, W.B., Hsieh, M.L., Willett, S.D., Hu, J.C., Horng, M.J., Chen, M.C., Stark, C.P., Lague, D., and Lin, J.C., 2003, Links between erosion, runoff variability and seismicity in the Taiwan orogen: Nature, v. 426, p. 648–651, doi: 10.1038/nature02150. Davis, J.D., Suppe, F.A., and Dahlen, Q., 1983, Mechanics of fold-and-thrust belts and accretionary wedges: Journal of Geophysical Research, v. 88, p. 1153–1172. Fairbanks, R.G., 1989, A 17,000-year glacio-eustatic sea level record: Influence of glacial melting rates on Younger Dryas events and deep ocean circulation: Nature, v. 342, p. 637–642, doi: 10.1038/342637a0. Honza, E., and Fujioka, K., 2004, Formation of arcs and backarc basins inferred from the tectonic evolution of Southeast Asia since the Late Cretaceous: Tectonophysics, v. 384, p. 23–53, doi: 10.1016/j.tecto.2004.02.006.

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