Special Paper 466
THE GEOLOGICAL SOCIETY OF AMERICA®
THE ORDOVICIAN EARTH SYSTEM
EDITED BY STANLEY C. FINNEY AND WILLIAM B.N. BERRY
The Ordovician Earth System
edited by
Stanley C. Finney Department of Geological Sciences California State University at Long Beach Long Beach, California 90840 USA William B.N. Berry Department of Earth and Planetary Science University of California, Berkeley Berkeley, California 94720 USA
Special Paper 466 3300 Penrose Place, P.O. Box 9140
Boulder, Colorado 80301-9140, USA
2010
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Contents
Preface Stanley C. Finney and William B.N. Berry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v 1. Global setting of Ordovician orogenesis Cees R. van Staal and Robert D. Hatcher, Jr. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1 2. Ordovician explosive volcanism Warren D. Huff, Stig M. Bergström, and Dennis R. Kolata . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13 3. Toward identifying potential causes for stratigraphic change in subtropical to tropical Laurentia during the Mohawkian (early Late Ordovician) Achim D. Herrmann and Bernd J. Haupt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29 4. The Upper Ordovician Guttenberg δ13C excursion (GICE) in North America and Baltoscandia: Occurrence, chronostratigraphic significance, and paleoenvironmental relationships Stig M. Bergström, Birger Schmitz, Matthew R. Saltzman, and Warren D. Huff . . . . . . . . . . . . . . . . 37 5. The Ordovician brachiopod radiation: Roles of alpha, beta, and gamma diversity David A.T. Harper . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69 6. Ordovician paleogeography and tectonics of the major paleoplates of China Chen Xu, Zhou Zhi-yi, and Fan Jun-xuan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 85 7. Ordovician of the Siberian Platform Alexander V. Kanygin, Tatiana N. Koren, Anastasia G. Yadrenkina, Alexander V. Timokhin, Oleg V. Sychev, and Tatiana Yu. Tolmacheva . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 8. Early–Middle Ordovician conodont paleobiogeography with special regard to the geographic origin of the Argentine Precordillera: A multivariate data analysis Guillermo L. Albanesi and Stig M. Bergström . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 119 9. Black shales: An Ordovician perspective William B.N. Berry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 141 10. Paleogeographic, paleoceanographic, and tectonic controls on early Late Ordovician graptolite diversity patterns Daniel Goldman and Wu Shuang-Ye . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 149
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Contents 11. Origin of Late Ordovician (mid-Mohawkian) temperate-water conditions on southeastern Laurentia: Glacial or tectonic? Frank R. Ettensohn . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 12. Correlations across a facies mosaic within the Lexington Limestone of central Kentucky, USA, using whole-rock stable isotope compositions John W. Coates, Frank R. Ettensohn, and Harold D. Rowe. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 177
Preface
Stanley C. Finney† Department of Geological Sciences, California State University at Long Beach, Long Beach, California 90840, USA William B.N. Berry† Department of Earth and Planetary Science, University of California, Berkeley, Berkeley, California 94720, USA With a duration of 44 million years (488–444 Ma), the Ordovician Period is a most significant chapter in Earth’s history. It includes the great Middle Ordovician biodiversification event, the Hirnantian glaciation, and long-term greenhouse conditions. During the period, many continents were dispersed widely along the equator and in the southern temperate belt (Laurentia, Baltica, Siberia, South China, and Tarim), and large Gondwana (Africa, South America, India, Antarctica, and Australia) occupied much of the Southern Hemisphere from the pole to the equator. Richly fossiliferous carbonate sediments, along with occasional shale facies and quartz sands, were deposited in shallow seas that flooded tropical continents, while siliciclastic sediments, including Hirnantian glacial and glacio-marine sediments, were deposited over large parts of Gondwana and peri-Gondwana terranes. After nearly complete submergence in the Early Ordovician, Laurentia underwent a dramatic, extensive regression and subsequent transgression that produced the continentwide, post-Sauk, pre-Tippecanoe unconformity, which may be expressed to various degrees in stratigraphic successions on most other paleocontinents that lay at low to middle paleolatitudes. The margins of the Ordovician oceanic realm, now preserved in orogenic belts (Antler, Taconic, Ouachita, and Ellesmerian of North America, Caledonides of northwest Europe, Central Asia, and Lachlan in eastern Australia), were sites of deep-marine sedimentation of thick turbiditic sandstone packages, shale, and chert. Magmatic-volcanic arcs along paleo-plate boundaries produced impressive belts of Ordovician granites in northwestern Argentina, and numerous, often thick, geographically widespread K-bentonites that are found today in eastern North America, northwest Europe, China, and the Argentine Precordillera and which represent some of the largest known fallout ash deposits of the Phanerozoic Era. Subduction along the margins of Laurentia, Gondwana, Baltica, and Siberia formed extensive linear orogenic belts. Accordingly, rocks of the Ordovician System are widespread on all continents except Antarctica (where the present ice cover may hide them), and capping Mount Everest they occupy the highest elevation on Earth (Ross, 1984). A 30 yr project of formally defining the Cambrian-Ordovician and Ordovician-Silurian boundaries, and a single set of global series-epochs and stages-ages for the Ordovician System-Period was recently completed by the International Subcommission on Ordovician System when the GSSP (global stratotype section and point) for the base of the Dapingian Stage and the Middle Ordovician Series was approved by the International Commission on Stratigraphy and ratified by the International Union of Geological Sciences in 2007 (Fig. 1). This scientific process was especially difficult and complex for the Ordovician System, because the high degree of biogeographic provincialism and ecologic differentiation of Ordovician faunas greatly limited global correlations (Finney, 2005). British series-epochs were often used as de facto nomenclature on stratigraphic correlation charts and geologic time scales. However, they were not widely adopted outside of the British Isles because they could not be correlated biostratigraphically with precision and high †
E-mails:
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Preface
GLOBAL SERIES
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443.7 Ma 445.6 Ma
GLOBAL STAGES
BRITISH SERIES
NORTH AMERICAN SERIES
HIRNANTIAN ASHGILL
UPPER
CINCINNATIAN
471.8 Ma
478.6 Ma
MIDDLE
468.1 Ma
CARADOC
MOHAWKIAN
SANDBIAN
DARRIWILIAN
LLANVIRN WHITEROCKIAN
DAPINGIAN ARENIG FLOIAN
LOWER
460.9 Ma
ORDOVICIAN
455.8 Ma
KATIAN
IBEXIAN TREMADOCIAN
TREMADOC
488.3 Ma
Figure 1. Chart showing correlation between global series and stages and regional series for Britain and North America. Calibrated ages from Ogg et al. (2008).
resolution. As a result, several independent and very different regional sets of series and constituent stages were established for the Ordovician System, with each generally applicable to a different paleoplate or modern continent, but this greatly confused Ordovician chronostratigraphy-geochronology and often resulted in imprecise correlations of Ordovician stratigraphic successions. The new set of global series and stages, and the precise definitions of their boundaries as well as those of the system, now provide a common language for correlation of Ordovician strata and for expressing geologic time (Fig. 1). During the many years required to establish the complete set of new global units, many Ordovician stratigraphers used the regional classification that they preferred, and some lamented the replacement of the British Series as the standard units of Ordovician stratigraphy (Cope, 2007). Nevertheless, now that the global set of units is complete, they are universally accepted and regularly used in papers on Ordovician stratigraphy. Regional chronostratigraphic classifications can still be used for those regional stratigraphic successions on which they were defined and may offer the most precise local correlation, and the correlation of the varied regional units to the new global series and stages is fully elucidated in Bergström et al. (2009). The process of establishing global chronostratigraphic units for the Ordovician System depended on, and resulted in, extensive, focused, often multidisciplinary studies of stratigraphic successions worldwide by a multitude of specialists. It generated a huge body of stratigraphic information, primarily on biostratigraphy and paleontology, but also on chemostratigraphy, sequence stratigraphy, and magnetostratigraphy as well as sedimentology and geochemistry. In addition it fostered the acquisition of radiometric dates that could be used to calibrate the ages of the chronostratigrapic units, particularly their boundaries, thus refining a significant part of the Geologic Time Scale (Cooper and Sadler, 2004). The primary goal of this long-term project was establishment of a single set of global chronostratigraphic units for precise, high-resolution correlations worldwide. However, it also fostered integrated, multidisciplinary studies that documented the state of, and important changes in, the Earth system during the Ordovician Period (e.g., Huff et al., 1992; Finney et al., 1999). As the chronostratigraphic project neared completion, members of the Ordovician Subcommission redirected their attention toward using the greatly expanded knowledge of Ordovician stratigraphy to address major events of the Ordovician Period. Former subcommission chair Barry Webby organized and led the most
Preface successful International Geological Correlation Programme (IGCP) 410, the Great Ordovician Biodiversification Event, which produced an impressive examination and summary of one of the major biotic events of not only the Ordovician but also the entire Phanerozoic (Webby et al., 2004). Along the same lines, a symposium titled “Global Ordovician Earth System” was organized at the 32nd International Geological Congress in Florence, Italy, in 2004. It was composed largely of contributions that used an Earth systems approach to examine aspects of the Earth’s lithosphere, biosphere, hydrosphere, and atmosphere during the Ordovician and to integrate them where possible. Several of those contributions, further elaborated since the Congress in Florence, form the core of this volume with additional contributions later solicited by the editors. The contributions in this volume range from global to regional in scope. Their focus varies from tectonic to paleogeographic and from sedimentological to paleontological. Yet all provide views of the Ordovician world. They are only a small sample of the diverse current research on the Ordovician System, yet they provide entry into the Ordovician world and in varying degree to interactions among Earth’s systems. Thus, the title of the volume is The Ordovician Earth System. In the lead article, Cees van Staal and Robert Hatcher conclude from their study of early Paleozoic orogenic belts that Ordovician orogenesis was due principally to accretion of arc terranes and ribbon microcontinents. Warren Huff and co-authors review the distribution of Ordovician K-bentonites in eastern North America, northwest Europe, Argentina, and China. Achim Herrmann and Bernd Haupt employ numerical models of the Ordovician ocean-climate system to propose that upwelling cold-water masses that penetrated the epicontinental sea of Laurentia during the early Late Ordovician indicate the onset of cooling conditions that plunged the Earth system toward icehouse conditions that later led to the Hirnantian glaciation as well as produced a regional extinction during the early Late Ordovician (Mohawkian) of eastern Laurentia. Stig Bergström and co-authors demonstrate the powerful correlation potential of the early Katian (early Late Ordovician) Guttenberg carbon isotope excursion, which occurred within the time of massive fallout ash deposits and the Mohawkian extinction, but they consider its cause to be enigmatic. Turning to the biosphere, David Harper’s detailed analysis of brachiopod diversity illustrates well the great Middle Ordovician biodiversification event, the Hirnantian mass extinction, and less dramatic changes and relates them to paleogeographic changes and tectonic events. Contributions by Chen Xu and colleagues on the Ordovician paleoplates of modern China and by Alexander Kanygin and co-authors on the Ordovician of the Siberian platform provide valuable summaries of stratigraphic information that are essential for reconstructing Ordovician paleogeography for significant continents of the Ordovician world. Through statistical analysis of Ordovician conodonts from globally widespread areas, Guillermo Albanesi and Stig Bergström examine paleobiogeographic relationships between Laurentia, Baltica, the Argentine Precordillera, and the proto-Andean margin of Gondwana and find that the evidence does not support a Laurentian origin of the Argentine Precordillera (the Cuyania terrane), a hypothesis that has received much recent attention. Organic-rich black shales are prominent, widespread Ordovician lithofacies and commonly contain abundant graptolites, which are essential for chronostratigraphic correlations. Using modern analogues, William Berry concludes that they accumulated in settings similar to modern hypoxic-anoxic environments. The final three papers in this volume are complementary examinations of the Earth system interactions in eastern Laurentia during the Katian (early Late Ordovician or Mohawkian). Daniel Goldman and Wu Shuang-Ye conclude that a significant episode of graptolite extinction was regional and not global and likely due to cold-water masses that penetrated the epicontinental sea of Laurentia, as proposed by the ocean–climate system model of Herrmann and Haupt (this volume). Frank Ettensohn associates the influx of cold-water masses to regional tectonics that altered regional paleogeography. John Coates and co-authors demonstrate that carbon-isotope stratigraphy provides high-resolution correlation between varied lithofacies within this complex depositional setting. Thorough reviews of manuscripts submitted for this volume were provided by Leho Ainsaar, Peter Cawood, Robin Cocks, Roger Cooper, Olda Fatka, Don Gorsline, Jack Green, Christopher Holmden, Dmitri Kaljo, Thomas Kelty, Dennis Kolata, Ed Landing, Stephen Leslie, Greg Ludvigson, Michael Melchin, Godfrey Nowlan, Mark Patzkowsky, Michael Pope, Leonid Popov, John Repetski, Rong Jiayu, Matthew Saltzman, Hans Peter Schonlaub, Brian Witzke, and Jan Zalaciewicz. Their efforts are very much appreciated.
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Preface REFERENCES CITED Bergström, S.M., Chen Xu, Gutiérrez-Marco, J.C., and Dronov, A., 2009, The new chronostratigraphic classification of the Ordovician System and its relations to major regional series and stages and to δ13C chemostratigraphy: Lethaia, v. 42, p. 97–107. Cooper, R.A., and Sadler, P.M., 2004, The Ordovician Period, in Gradstein, F.M., Ogg, J.G., and Smith, A.G., eds., A Geologic Time Scale 2004: Cambridge, UK, Cambridge University Press, p. 165–187. Cope, J.C.W., 2007, What have they done to the Ordovician?: Geoscientist, v. 17, p. 19–21. Finney, S., 2005, Global series and stages for the Ordovician System: A progress report: Geologica Acta, v. 3, p. 309–316. Finney, S.C., Berry, W.B.N., Cooper, J.D., Ripperdan, R.L., Sweet, W.C., Jacobson, S.R., Soufiane, A., Achab, A., and
Noble, P.J., 1999, Late Ordovician mass extinction: New perspective from stratigraphic sections in central Nevada: Geology, v. 27, p. 215–218, doi: 10.1130/0091-7613(1999) 027<0215:LOMEAN>2.3.CO;2. Huff, W.D., Bergström, S.M., and Kolata, D.R., 1992, Gigantic Ordovician volcanic ash fall in North America and Europe: Biological, tectonomagmatic, and event-stratigraphic significance: Geology, v. 20, p. 875–878, doi: 10.1130/0091 -7613(1992)020<0875:GOVAFI>2.3.CO;2. Ogg, J.G., Ogg, G., and Gradstein, F.M., 2008, The Concise Geologic Time Scale: Cambridge, UK, Cambridge University Press, 177 p. Ross, R.J., Jr., 1984, The Ordovician System, progress and problems: Annual Review of Earth and Planetary Sciences, v. 12, p. 307–335, doi: 10.1146/annurev.ea.12.050184.001515. Webby, B.D., Paris, F., Droser, M.L., and Percival, I.G., 2004, The Great Ordovician Biodiversification Event: New York, Columbia University Press, 484 p.
The Geological Society of America Special Paper 466 2010
Global setting of Ordovician orogenesis Cees R. van Staal† Geological Survey of Canada, 625 Robson Street, Vancouver, British Columbia V6B 5J3, Canada Robert D. Hatcher, Jr. Department of Earth and Planetary Sciences and Science Alliance Center of Excellence, University of Tennessee, Knoxville, Tennessee 37996-1410, USA
ABSTRACT The global distribution, setting, and dynamic implications of Ordovician orogenesis are reviewed. Evidence for true Ordovician continent-continent collision is absent. Orogenesis is principally due to accretion of arc terranes and/or ribbon microcontinents. Most arc terranes are ensialic and separated from the adjacent continents by backarc or marginal basins, the episodic closure of which commonly was responsible for orogenesis. Little evidence is preserved for true intra-oceanic juvenile arcs during the Early to Middle Ordovician. Instead, subduction appears to have been localized near the margins of Laurentia, Gondwana, Baltica, and Siberia, forming extensive linear orogenic belts during relatively short periods when the upper plate switched from extension to compression. Such tectonic switching appears to have taken place along the entire length of the Pacific and Iapetan margins of Gondwana (>10,000 km) from Middle–Late Cambrian to Early Ordovician time. The onset of orogenesis along Gondwana’s Pacific margin during the end of the Early Cambrian (ca. 513 Ma) coincided with subduction initiation along both margins of the Iapetus Ocean. Orogenesis and subduction initiation are causally related to a global-scale plate reorganization, probably induced by terminal amalgamation of Gondwana. During the Paleozoic, Laurentia’s Iapetan margin steadily grew in size and expanded southward owing to continuous accretion of suprasubduction zone oceanic crust, peri-Gondwanan arc terranes, and ribbon microcontinents. In contrast, the Pacific, Iapetan, and Rheic margins of Gondwana saw little addition of new, allochthonous crust. Accretion mainly involves reattachment of previously rifted-off arc terranes and small slivers of the intervening marginal basin crust. INTRODUCTION
and Eurasia. There is also minor Ordovician deformation in Africa, but this mainly relates to transpression and basin inversion induced by far-field stresses and is not relevant to this paper. What makes Ordovician orogenesis stand out from all other periods in the Paleozoic is that nowhere it appears to have involved true continent-continent collision. Some workers have invoked a Middle Ordovician collision between Gondwana and Laurentia (e.g., Dalziel et al., 1994), mainly based on orogenesis taking
Orogenic events (orogenesis) affected many parts of the world during the Ordovician, and the remnants of the mountain belts created can be found principally in the present-day continents of North America, South America, Australia, Antarctica, †
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Van Staal, C.R., and Hatcher, R.D., Jr., 2010, Global setting of Ordovician orogenesis, in Finney, S.C., and Berry, W.B.N., eds., The Ordovician Earth System: Geological Society of America Special Paper 466, p. 1–11, doi: 10.1130/2010.2466(01). For permission to copy, contact
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place nearly coevally on both margins of Iapetus (Fig. 1) and involving accretion of Laurentian-derived crust (Precordillera) to Gondwana. However, this model was abandoned by Dalziel (1997) and replaced by a short-lived collision between a large promontory on Laurentia that bulged into the Iapetus Ocean (Texas Plateau) and Gondwana. Most other workers favor collision between a Laurentian-derived microcontinent (Cuyania/ Precordillera) and Gondwana instead (e.g., Ramos et al., 1986; Astini et al., 1995; Thomas and Astini, 1996). Irrespective of whether or not the fleeting promontory-Gondwana collision ever happened, we do not consider it a true continent-continent collision, because the area involved in this putative collision was very small compared with the length of the margins. Instead the four principal continental masses present on Earth at that time (Laurentia, Gondwana, Baltica, and Siberia; see Fig. 1 and also Cocks and Torsvik, 2002; Torsvik and Cocks, 2005) were involved in accretion of peri-cratonic arc terranes and/or ribbon-shaped microcontinents and remained separated by oceans (principally the Iapetus, Aegir, Tornquist-Ran, and Pacific Oceans). These accretionary events were precursors to major continent-continent collisions that occurred during the middle to late Paleozoic. Formation and accretion of ribbon-shaped microcontinents, principally derived from Gondwana and to a lesser extent Laurentia, and the closure of marginal oceanic basins on a global scale
during the Ordovician are important tectonic phenomena that pose important geodynamic questions. The purpose of this paper is to provide a brief review and summary of the data and ideas related to Ordovician orogenic events, mainly concentrating on their tectonic evolution and whether they are kinematically and/or dynamically linked, and if so, how and why, and finally to discuss the implications of these orogenic events. The latter is done to further our understanding of the global tectonic climate and the driving tectonic forces during the Ordovician. In this paper we have adopted the usage of the new Ordovician stage names (Finney, 2005; Bergström et al., 2006). OROGENESIS AND TIME In general, classification and correlation of structures and other evidence of tectonic events were historically (pre–plate tectonics) done on the basis of time alone, which is a hazardous process (although one sometimes has no other choice), because orogenic events are invariably diachronous, and kinematically unrelated tectonic events may take place coevally in different parts of the world (e.g., van Staal, 1994). These, by chance, may have been juxtaposed later as a result of subsequent plate motion(s) and closure of intervening oceans (van Staal et al., 1998).
Figure 1. Distribution of the main continents, oceans, subduction zones and arc-backarc systems, and micro-continents in the Early Ordovician (Tremadocian). The South Pole is situated in North Africa. The Pacific margin of Gondwana with the Ross-Delamerian belt of Australia and Antarctica that is continuous with the active proto-Andean margin, together forming a very long circum-Gondwanan subduction system, is just out of view. The Corel Draw figure originally made by Conall MacNiocaill for van Staal et al. (1998) was modified for this paper. Narrow volcanic arc terranes are indicated by ellipses. Dashed lines are used where the existence of subduction zones and/or transform faults is uncertain or contentious. Stippled structures may not have been active at this time. Arm—Armorica terrane; Aval—Avalonia; Boh—Bohemia terrane; Cadom—Paleozoic Cadomian arc terrane of Stampfli and Borel (2002); Car—Carolinia; Dashw & ND arc— Dashwoods microcontinent and Notre Dame arc; Famat arc—Famatinian arc-backarc system; Finnm arc—Finnmark arc-backarc system; Gan—Ganderia; Høl.—Hølanda terrane; Ib—Iberia terrane; Kip arc—Kipchak arc; M—Meguma terrane; Mu arc—Mugodzhar arc-microcontinent; Pc—Precordillera terrane; Penobs. arc—Penobscot arc-backarc system; S-M—Sakmara-Magnitogorsk marginal basin; Svalb—estimated position of the Svalbard and Pearya terranes.
Global setting of Ordovician orogenesis ORDOVICIAN OROGENESIS Orogenic events discussed in this paper include the Finnmarkian, Penobscottian, Taconian, Grampian, Shelveian, and the early stages of the Salinic and Scandian in the Appalachians and Caledonides of North America, British Isles, and Scandinavia; M’Clintock in Pearya of Arctic Canada and its unnamed equivalent in neigboring Svalbard; Famatinian and/or Ocloyic in South America; Sardic or Eo-Variscan in the Eurasian Variscan orogen; events in the Central Asian orogenic belt (Altaids); and parts of the Ross-Delamerian and Lachlan orogenies in eastern Australia and Antarctica (Fig. 1). Most of these orogenies are restricted to the Ordovician, but some started or mainly took place during the Middle to Late Cambrian (Finnmarkian and Ross-Delamerian), whereas others started during the latest Ordovician and climaxed during the Silurian (Salinic and Scandian). Their inclusion here is either due to diachroneity, such that their orogenesis locally extends into the Ordovician, and/or to their dynamics having had an impact on understanding Ordovician orogenesis in general. Following the definition of McKerrow et al. (2000), we restrict usage of the term Caledonian orogeny to describe all orogenic events involved in closing the Iapetus Ocean. Ordovician Paleogeography Comprehensive reviews of Earth’s geography, position of tectonic elements, and migration of terranes relevant to this paper are presented in van Staal et al. (1998), Cocks and Torsvik (2002), Stampfli and Borel (2002), Torsvik and Cocks (2005), and Cawood (2005). Although these authors do not necessarily agree on the evolution of each element, they broadly agree on the disposition of the main continental landmasses summarized in Figure 1. We adhere to these views, which are significantly different from the one proposed by Dalziel (1997). North American Appalachians and British Caledonides Ordovician orogenesis affected the Appalachians differently in different parts of the mountain chain. These orogenic events generally are diachronous along the length of the mountain chain, and to make matters more confusing, they occurred on both sides of the Iapetus ocean while the ocean was still very wide (Fig. 1), which makes the classification of Ordovician structures difficult without knowing the provenance and tectonic setting of the deformed rocks involved. The oldest orogenic events that affected the Appalachian orogen are represented by the latest Cambrian–Early Ordovician phases of the Taconian on the Laurentian side and the Penobscottian on the Gondwanan side (e.g., Colman Sadd et al., 1992; van Staal et al., 2007). Early phases of both orogenies partly overlap in time, and both involve ophiolite accretion-obduction and arc-continent collision (Fig. 1). The onset of ocean closure along opposite margins of the Iapetus Ocean (Fig. 1) signals a major plate reorganization at the end of the Early Cambrian (van Staal et al., 1998; also see below).
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The main phase of the Taconian (Figs. 1 and 2) in the northern Appalachians is the collision between magmatic arc(s) and Laurentia, which started during the Early Ordovician (ca. 475 Ma; Humberian in Newfoundland) and lasted until the Late Ordovician. Deformation and accompanying tectonothermal activity are principally related to collision of a magmatic arc (Notre Dame–Shelburne Falls arc) built upon a ribbon-shaped peri-Laurentian microcontinent (Karabinos et al., 1998) named Dashwoods in Newfoundland (Taconic 2 suture in Fig. 1; Waldron and van Staal, 2001; van Staal et al., 2007), rapidly followed by Late Ordovician (ca. 450 Ma) softdocking of the peri-Gondwanan Popelogan–Victoria–Bronson Hill arc (Taconic 3, Fig. 2) built on a sliver of Ganderian crust (van Staal, 1994; Rogers et al., 2006 ; Zagorevski et al., 2007). This phase represents the terminal phase of the Taconian and closed the main oceanic tract of the Iapetus between Laurentia and Ganderia (Fig. 2), although a wide oceanic backarc basin (Tetagouche-Exploits backarc) and seaway (Acadian seaway) remained open, to be closed during the Silurian and Early Devonian, respectively (van Staal et al., 1998, 2007; van Staal, 2005; Valverde-Vaquero et al., 2006). In the southern Appalachians the Taconian is termed Blountian (Kay, 1942; Rodgers, 1971; Drake et al., 1989). Orogenesis here is particularly evident from the accumulation of the substantial Blount-Sevier foredeep clastic wedge on Laurentia’s continental margin near the beginning of the Middle Ordovician (ca. 468 Ma, Darriwilian), but it is also recorded by abundant metamorphic ages, formation of eclogite, and generation of arclike plutons in the internal parts of the mountain chain (central and eastern Blue Ridge and western Inner Piedmont). The most attractive model, which is virtually identical to the northern Appalachians, is one involving formation of a west-facing magmatic arc on a peri-Laurentian microcontinent (Hatcher, 1989; Miller et al., 2006; Hibbard et al., 2007). Metamorphism as high as granulite facies affected the rocks in the internal parts of the southern Appalachians between 455 and 465 Ma (Moecher et al., 2004) as a result of the arc-continent collision and associated obduction. Hibbard (2000) proposed that immediately subsequent to this collision the leading edge of the peri-Gondwanan Carolina terrane docked in the latest Ordovician–Early Silurian. If Hibbard’s model is correct, it indicates a remarkably similar kinematic/dynamic evolution for the entire Appalachian margin of Laurentia (Fig. 2) during the Ordovician (Hibbard et al., 2007), virtually eliminating the possibility of a fleeting Middle– Late Ordovician Gondwana-Laurentia collision as proposed by Dalziel et al. (1994) and Dalziel (1997). Other compelling sedimentologic, geochronologic, and field data from the southern and central Appalachians also support closure of the bulk of the Iapetus Ocean in the Late Ordovician, but these data also indicate that the trailing Rheic Ocean remained open until its late Paleozoic destruction (Bream, 2003; Merschat et al., 2005; Hatcher and Merschat, 2006; Merschat and Hatcher, 2007). Possibly several other, yet to be identified terranes outboard of Carolina were accreted to Laurentia prior to final closure of the Rheic Ocean,
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Figure 2. Paleogeography of the tectonic elements shown in Figure 1 in the Late Ordovician (450 Ma). Figure modified from van Staal et al. (1998). Abbreviations are the same as in Figure 1. Tac 2 and 3 refer to suture zones formed during the Taconian orogeny (van Staal et al., 2007). ACSW—Acadian seaway between Ganderia and Avalonia, the closure of which caused the Acadian orogeny; Hun—Hun superterrane of Stampfli et al. (2002); PVA— Popelogan-Victoria arc; RILA—Red Indian Lake arc; TEB—Tetagouche-Exploits backarc basin; Thor—Thor suture zone, forming part of the TESZ (Trans-European suture zone).
which caused major underthrusting of the “westerly terranes,” situated near the leading edge of Laurentia beneath Carolina. The Taconian also affected the central Appalachians, but its effects were generally recorded slightly later (Late Ordovician) on the foreland. Here, the Martinsburg (Martinsburg-Juniata) clastic wedge is thicker (~5 km) than the Blount-Sevier wedge (~3 km) farther to the south. Deposition of the Martinsburg in the type area started approximately at the same time as the BlountSevier wedge, but deposition continued into the Late Ordovician. These relationships attest to the diachronous and composite nature of the Taconian orogeny in the Appalachians (Shanmugam and Lash, 1982; Drake et al., 1989; Hatcher, 1989). Several exposures of the angular unconformity from Pennsylvania and New Jersey to eastern New York reveal Early Silurian (Llandovery) molasse overlying tilted Upper Ordovician clastic wedge rocks (Rodgers, 1971). Extensive peri-cratonic island arc systems fringing both the Iapetan margins of Laurentia and Gondwana until the Late Ordovician (van Staal et al., 1998), combined with paleomagnetic and faunal evidence (e.g., Cocks and Torsvik, 2002), require that both continents remained widely separated until at least the Late Devonian. Hence, we conclude that there is no evidence to support collision of Laurentia with any large continental landmass other than ribbon-shaped continental arc terranes during the Ordovician.
The Taconian is kinematically and temporally equivalent to the Grampian in the British Isles (Dewey and Shackleton, 1984). Like the Taconian, the Grampian starts with ophiolite obduction immediately prior to arc-continent collision in the Early Ordovician (Dewey and Mange, 1999). The main Grampian orogenesis is a short-lived ca. 15 Ma event (ca. 475–460 Ma), although it comprises a complex tectonometamorphic cycle involving early blueschist and eclogite facies metamorphism of the subducting Laurentian (Dalradian) margin (Friend et al., 2000; Chew et al., 2003), polyphase deformation, and medium to high grade Barrovian metamorphism of the collision complex concurrent with syntectonic intrusion of plutons (Friedrich et al., 1999; Soper et al., 1999; Dewey and Mange, 1999). In this respect, the Grampian is nearly coeval and kinematically remarkably similar to the main phase of the Taconian in central and western Newfoundland (van Staal et al., 1998, 2007). Grampian orogenesis was rapidly followed by the Late Ordovician–Early Silurian Scandian orogenesis similar in nature to the Salinic in the northern Appalachians. While Grampian orogenesis occurred on the Laurentian side, Early to Middle Ordovician orogenesis also took place on the Gondwanan side in the British Isles (e.g., Max et al., 1990; Todd et al., 2000). This orogenic event, which has no name attached to it as yet in the British Isles, is the equivalent of the Penobscottian in the Appalachians (van Staal et al., 1998).
Global setting of Ordovician orogenesis Scandinavian, Greenland, and Svalbard Caledonides Ordovician orogenesis in the Scandinavian Caledonides was, until recently, mainly thought to be represented by an early event that principally took place in the Late Cambrian along the Baltic margin (Finnmarkian; Andréasson et al., 1998; Sturt and Ramsay, 1999; Grenne et al., 1999; Roberts, 2003). However, the structurally highest allochthons, in the Scandinavian Caledonides, which have a Laurentian provenance, have preserved ample evidence for a major Ordovician tectonic event that took place prior to their Silurian (Scandian) docking with the Baltic margin (Yoshinobu et al., 2002; Roberts, 2003). The Ordovician event is kinematically equivalent to the Taconian–Grampian orogeny in the Appalachians and British Caledonides. Surprisingly, no evidence for this event has been preserved in the East Greenland Caledonides, the area from which these rocks presumably were derived. However, the presence of Middle Ordovician blueschist, eclogite, and other evidence of major tectonism in western Svalbard (Dallmeyer et al., 1990; Gee and Page, 1994) and Pearya (M’Clintock orogeny; Trettin, 1987) leave little doubt that the Taconian–Grampian was probably continuous along the entire length of Laurentia’s Iapetan margin north of the Appalachians. The lack of evidence of this event in East Greenland suggests that most Ordovician tectonism was restricted to the peri-Laurentian arc terrane preserved in the Caledonides upper allochthon (Hølanda terrane of Grenne et al., 1999), which is thought to have been present directly outboard of Laurentia at the latitude of Greenland (Fig. 1) during the Ordovician (Grenne et al., 1999). A similar relationship has been observed in Newfoundland, where most Taconian tectonothermal activity was restricted to the peri-Laurentian Notre Dame arc and its basement (the Dashwoods microcontinent of Waldron and van Staal, 2001) and largely absent in the exposed parts of the adjacent passive margin, because most of the underthrusted margin rocks were never fully exhumed (van Staal et al., 2007). The Finnmarkian event is generally thought to be mainly restricted to the latest Cambrian, but age dating revealed that orogenesis continued into the Early–Middle Ordovician in Finnmark in northern Norway (e.g., Rice and Frank, 2003). The Finnmarkian involved arc development, ophiolite obduction, arc accretion, and metamorphism to eclogite facies conditions that principally affected rocks that now reside in the middle and upper allochthons in Norway and Sweden (Gee et al., 1985). The main Finnmarkian event is thought to have been the result of collision of a peri-cratonic arc terrane with the adjacent Baltic margin (Sturt and Roberts, 1991; Grenne et al., 1999; Roberts, 2003). Holtedahl (1920) recognized a slightly younger Early Ordovician event in central Norway and called it the Trondheim event, which probably is related tectonically to Finnmarkian orogenesis. It includes ophiolite emplacement and blueschist and eclogite facies metamorphism (Eide and Lardeaux, 2002; Roberts, 2003). Ophiolites have yielded uranium lead ages of 497–482 Ma (Dunning and Pedersen, 1988; Roberts, 2003), which are similar in age to the Penobscot ophiolites preserved in Newfoundland.
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In general, the close similarities in age and style of orogenesis between the Penobscottian and Finnmarkian, both positioned on the southern side of Iapetus, invite correlation. If correct, Baltica and adjacent Gondwana (Ganderia and parts of Avalonia; Rogers et al., 2006) may have been characterized by a more or less continuous (excluding transform faults) peri-cratonic arc terrane with a length on the order of thousands of kilometers. Thus consistent to what has been observed in the British Caledonides and the northern Appalachians, Ordovician orogenesis also appears to have taken place on both sides of the Iapetus Ocean in the Scandinavian Caledonides. The Baltic–East European craton (Baltica) also converged with Ganderia and Avalonia during closure of the Tornquist Ocean, which separated Baltica from Gondwana, along the former’s southern margin (Valverde-Vaquero et al., 2006) and which led to formation of the Thor suture. The Tornquist closure was completed by the end of the Ordovician (454–443 Ma) and caused the Shelveian orogeny, whose effects on the exposed parts of Baltica (lower plate) appear to be minor (Torsvik and Rehnstrom, 2003). In general, this collision appears to have been relatively soft, with deformation and metamorphism (low grade) mainly restricted to the collsion zone (Trans-European suture zone [TESZ]; see below). Scandian orogenesis in the Scandinavian Caledonides mainly took place during Silurian times. The Scandian was principally caused by the main collision between Laurentia (Greenland) and the Baltic craton. It is coeval and is correlated with the Salinic orogeny (Laurentia-Ganderia collision) in the northern Appalachians (van Staal, 2005; van Staal et al., 2008). Since Baltica and Ganderia were already together by Late Ordovician times as a result of the Shelveian orogeny (see above and Valverde-Vaquero et al., 2006), the Salinic and Scandian orogenies are both dynamically related to collision of Laurentia (upper plate) and the composite craton comprising Baltica and Ganderia (lower plate). Early Paleozoic Orogenic Events in Eurasian Gondwana The orogenic events that involved the assembly of Paleozoic Europe comprised kinematically and temporally diverse, and in part, unrelated events that were scattered throughout the Paleozoic, culminating in the collision that formed Pangea at the end of the Paleozoic. Orogenic events that took place during the Ordovician include the Shelveian orogeny, which formed as a product of convergence among Baltica, Ganderia, and Avalonia (McKerrow et al., 1991; Pharaoh et al., 1993; van Staal et al., 1998; Torsvik and Rehnstrom, 2003; Valverde-Vaquero et al., 2006). The structures related to the Shelveian orogeny are mainly present in the unexposed basement beneath the North Sea and Denmark. Shelveian structures near the Thor suture were incorporated in a poorly understood but complex, long-lived deformation zone that characterizes the TESZ. The TESZ separates the Baltic–East European craton from younger accreted lithosphere of western Europe over a strike length exceeding 2000 km. Accretion-related deformation in the TESZ possibly had already started during the
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Late Cambrian–Early Ordovician (Sandomierz orogenic phase of Belka et al., 2000) owing to accretion of a small peri-Gondwanan crustal block to Baltica. Evidence for Early to Late Ordovician orogenic events (referred to as the Sardic or Eo-Variscan) is also preserved farther west in Europe (e.g., Armorican terrane assemblage) from the Alps to Spain and Portugal (von Raumer, 1998; Handy et al., 1999; Martínez Catalán et al., 2002; Matte, 2002; Robardet, 2002; and references therein). This evidence includes arc magmatism, ophiolite generation, polyphase deformation, and eclogite and high-grade metamorphism, which are generally thought to be related to formation of a short-lived volcanicmagmatic arc terrane(s) and backarc basin(s) fringing the Eurasian margin (Fig. 1) of the Gondwanan supercontinent (Paleozoic Cadomia of Stampfli et al., 2002) and an accretionary event. Ordovician orogenesis here does not represent a major collisional event but more of an amalgamation and/or reattachment of the peri-Gondwanan, ribbon-like arc terrane(s) to the Eurasian margin of Gondwana. According to Stampfli and Borel (2002), the reattached arc terranes were assembled into the long Hun superterrane (Fig. 2), which became the leading edge of the Eurasian margin until it rifted off Gondwana as a coherent ribbon continent during the latest Silurian and started to move north, opening part of the Paleotethys in its wake. Reattachment of the Paleozoic Cadomian arc terrane and closure of the intervening marginal basin during the Early to Middle Ordovician took place after departure of Avalonia and concomitant eastward migration of the trailing Rheic Ocean spreading center to the north of the Cadomian arc terrane. The eastward migration of the Rheic spreading center may have put the arc terrane into compression (Stampfli and Borel, 2002) and caused it to re-accrete. Final accretion of the Hun superterrane to Laurussia took place later during the Variscan (Martínez Catalán et al., 1997, 2007; Matte, 2002). The ribbon terranes fringing the Eurasian margin extended at least as far east as the Himalayan margin of India (Stampfli and Borel, 2002; Cocks and Torsvik, 2002). The latter also preserved evidence for extensive Cambrian–Ordovician orogenesis (e.g., Gehrels et al., 2006). Accretion-reattachment of a peri-Gondwanan ribbon-like terrane to the Himalayan margin of India, possibly kinematically related to coeval events along the European segment of this Gondwanan margin, is also an attractive mechanism for explaining Ordovician orogenesis here. Central Asia Orogenic Belt (Altaids) and the Urals The notion that the Ordovician globally was a time of development of a series of rapidly evolving and amalgamating pericratonic arc terranes, which, when built on continental crust commonly formed ribbon continents, may also be valid for the Central Asia orogenic belt of central Asia and the Uralides. The central Asia orogenic belt and narrow Uralides formed mainly during closing of the Aegir Ocean (Fig. 1) and separated the Siberian (Angara) craton from the Baltic–East European craton to the east and the accreted terranes of the Tethysides to the south. Ordovician orogenesis in the Central Asia orogenic belt
mainly involved closure of marginal basins and assembly of separated fragments of the ribbon-shaped Kipchak arc, according to S¸engör and Natal’in (1996). In addition, collision of part of the Kipchak arc with the Mugodzhar arc-microcontinent, which was positioned outboard off the Baltic–East European craton’s eastern margin and separated from it by the Sakmara-Magnitogorsk marginal basin (Fig. 1), may have induced oroclinal bending and strike-slip imbrication of the Kipchak arc. The Mugodzhar arcmicrocontinent was later re-accreted to the East European craton during the middle–late Paleozoic Variscan orogeny (Puchkov, 1997). Kröner et al. (2007), on the other hand, suggested that the central Asia orogenic belt is an accretionary orogen that formed by across-strike rather than by highly oblique accretionary processes, and the precursor of the Kipchak arc was not a ribbon continent. In addition, detailed structural, petrologic, geochemical, and geochronologic studies in the Ol’khon (Baikal) region in southern Siberia by Fedorovsky et al. (2005) suggest that the central Asia orogenic belt includes an Early Ordovician orogenic event. This event is represented by a polydeformed tectonic collage of ophiolite, arc volcanic, plutonic, and sedimentary rocks, produced during its accretion to the southern margin of the Siberian craton. Accretion was oblique, with strike-slip dominating over dip-slip tectonics. Metamorphic grade ranges from granulite facies against the Siberian craton, with little or no metamorphic overprint on the craton itself, to amphibolite facies to the south. Shortening is indicated by craton-vergent, ductile-brittle Paleozoic folds and thrusts in the adjacent Siberian craton, which formed during accretion of the tectonic collage to the craton. An intriguing element of this part of the orogen consists of linear to blocky bodies of marble associated with metasiliciclastic and volcanic rocks to the northwest that give way to dominance of less deformed and metamorphosed volcanic and plutonic rocks to the south. Famatinian Belt in South America Ordovician orogenesis was mainly localized along the protoAndean margin, and is particularly well preserved in north-central Argentina, adjacent Chile and southern Peru, and Bolivia. This belt probably extends through Ecuador into Colombia. Orogenesis comprises subduction-related magmatism, metamorphism, deformation, and foreland basin development, which are variably referred to as the Famatinian and/or Ocloyic orogenic cycle (Ramos, 1988). During the early Paleozoic the Famatinian belt was an active margin with development of a magmatic arc, which appears to have switched on several occasions from a compressional to an extensional state (e.g., Rapela et al., 1998). Transient compression of the magmatic arc caused localized deformation of arc plutons, which started during the latest Cambrian and/or earliest Ordovician (Pankhurst et al., 1998). Middle Ordovician compression of the Puna’s Famatinian arc-backarc basin system closed the backarc basin and led to concomitant deformation, retroarc–foreland basin development (Bahlburg and Hervé, 1997), and metamorphism in the underlying Mesoproterozoic
Global setting of Ordovician orogenesis crust of the Arequipa-Antefolla terrane (Loewy et al., 2004). Farther south, orogenesis mainly involves accretion of the ribbonlike Precordillera-Cuyania microcontinent, whose provenance remains conjectural, although it is generally regarded as having been derived from Laurentia (Ramos et al., 1986; Astini et al., 1995; Thomas and Astini, 1996). Terra Australis Orogen in Australia and Antarctica Early Paleozoic orogenesis is also evident along the eastern margin of Australia, Tasmania, and the Transantarctic Mountains of Antarctica, which were positioned directly along strike of the proto-Andean margin of South America during that time. All three continents formed part of a very extensive convergentaccretionary margin that nearly circumnavigated the entire length of the Pacific margin of the Gondwana supercontinent (Terra Australis orogen of Cawood, 2005). The first phase of Ordovician orogenesis was represented by the waning stages of the predominantly Cambrian Ross-Delamerian orogen (Foden et al., 2006). However, age dating of the slaty cleavage of the most outboard Robertson Bay terrane (northern Victoria Land) suggests that deformation, at least locally, continued well into the Early Ordovician (Dallmeyer and Wright, 1992). This event was followed by the first phase of the Lachlan orogeny during the Late Ordovician (Foster et al., 1999; Foster and Gray, 2000; Collins, 2002). Both the Ross-Delamerian and Lachlan orogenies involve accretion of peri-cratonic volcanic arcs to the AustralianAntarctic cratons related to closure of the intervening marginal basins. The Ross-Delamerian orogeny may have involved more than one subduction zone (see reviews by Boger and Miller, 2004, and Cawood, 2005, and references therein) and was a tectonically complex event. Ophiolite obduction locally accompanied closure of the marginal basins. The passive margin of the Lachlan backarc basin was built partly on the foundered Delamerian Mountains and hence was filled by a large volume of clastic detritus from this orogen. This clastic wedge was incorporated into a thick accretionary wedge during Late Ordovician–Silurian backarc basin closure. The latter was achieved by subduction of marginal basin crust beneath both the passive and active sides of the basin. The Lachlan orogenic events continued into the middle Paleozoic, around the Devonian-Carboniferous boundary, which led to full development of the Lachlan segment (Foster and Gray, 2000) of the Terra Australis orogen. DISCUSSION AND CONCLUSIONS The evidence for worldwide Ordovician orogenesis clearly indicates that it was a global phenomenon, but in contrast to the Proterozoic Grenvillian, late Paleozoic Appalachian, or Tertiary Alpine-Himalayan orogenic cycles, it did not involve true continent-continent collision(s), and the major continents remained separated by wide oceans. Instead the Ordovician is marked by formation of peri-cratonic arcs of great strike-length along the margins of Laurentia, Baltica, Siberia, and Gondwana;
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linear orogenic belts, several characterized by low-grade metamorphism; marginal basins; suprasubduction ophiolites; ribbon microcontinents; global high-water stands; and black shale deposition. Evidence for formation of true intra-oceanic juvenile arcs starved of continental-derived sediment (such as found in the Pacific today) is rare or absent. The latter is significant because it indicates that subduction was localized along continental margins. In addition, prior to the start of the Ordovician, during the Middle and Late Cambrian (513–490 Ma), the margins of virtually all continental cratons underwent either the onset of orogenesis or initiation of subduction. The latter commonly led to generation of large tracts of suprasubduction zone ophiolite and associated boninite. For example, subduction had initiated between 515 and 510 Ma directly outboard of the Ganderian, Baltic, and Laurentian margins (van Staal et al., 1998, 2007) at a time when the Pacific margins of Antarctica and eastern Australia saw the onset of Ross-Delamerian orogenesis. Early Cambrian Pampean orogenesis along the West Gondwanan margin had started slightly earlier (ca. 530 Ma; e.g., Rapela et al., 1998), but may also have been causally related. Combined, these phenomena suggest a causal link to a global-scale plate reorganization. An attractive mechanism was the Early Cambrian terminal suturing between East and West Gondwana (Boger and Miller, 2004; Cawood, 2005) and final assembly of Gondwana, which undoubtedly must have had a major impact on the Early–Middle Cambrian global plate motion budget. The rapid volume increase of oceanic spreading ridges that probably accompanied the global-scale plate reorganization may also be responsible for the eustatic rise in sea level during the Cambrian. Surprisingly, subduction initiation in the Iapetus appears to have taken place near the passive margins of Laurentia, Ganderia, and Baltica. In the Phanerozoic geological record there are no obvious examples in which passive margins were directly transformed into long-lived, active margins. Margins are generally characterized by old, strong oceanic lithosphere and hence are not favorable sites for subduction initiation (Cloetingh et al., 1982). The locations of nascent subduction zones are generally thought to be controlled by the presence of a zone of weakness, buoyancy, and compressive forces, the latter related to convergence (Mueller and Phillips, 1991, and references therein). The presence of compressive forces is critical, because no combination of a lithosphere-penetrating fault and negative buoyancy forces alone would produce a subduction zone (Hall et al., 2003). For this reason, van Staal et al. (2007) proposed that subduction had nucleated on old transform faults and/or abandoned spreading centers near the Laurentian margin of the Iapetus Ocean, created by an Early Cambrian inboard ridge jump(s), whereas on the opposite side of the Iapetus, near Avalonia and Ganderia, an extensive transform fault system had been created by late ridge-trench collisions during the late Neoproterozoic and Early Cambrian, respectively (Rogers et al., 2006). However, in light of the evidence discussed above and elsewhere (Cawood and Buchan, 2007), the possibility that subduction can initiate along a passive margin should not be ruled out. Erickson (1993)
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suggested that subduction could start near a passive margin, provided that the continent-oceanic interface becomes decoupled as a result of reactivation of old, rifting-related faults during strikeslip or extension. Rollback of downgoing slabs probably was mainly responsible for the great tract of Early Ordovician peri-cratonic arcs and associated marginal basins that were present in the Iapetus, Aegir, and Pacific Oceans. In particular, the propagation of the peri-cratonic arc-backarc systems that were present along the entire length of the Pacific and Iapetan margins of Gondwana (Australia, Antarctica, proto-Andean margin of South America, Ganderia, and Avalonia) and probably also Baltica was very impressive and comparable in scale to the active margins that surrounded the Pacific Ocean during the Mesozoic and Cenozoic. Tensional forces induced by slab rollback probably were also responsible for the terrane dispersal (Avalonia and Ganderia) along Gondwana’s Iapetan margin (e.g. Schultz et al., 2008) and their subsequent motion toward Laurentia (van Staal et al., 1998). Terrane dispersal along the Laurentian margin—e.g., Precordillera (Astini et al., 1995; Thomas and Astini, 1996), several internal southern Appalachian terranes (Merschat et al., 2005; Hatcher and Merschat, 2006), and Dashwoods (Waldron and van Staal, 2001; Cawood et al., 2001)—on the other hand was probably due to an inboard ridge jump from a spreading center positioned close to the Laurentian margin. It calved off ribbonshaped slivers of continental crust from adjacent promontories (van Staal et al., 2007). During the Ordovician, peri-cratonic terranes along the Gondwanan margin underwent several short-lived periods (10– 20 Ma) during which large arc segments switched from extension to compression, commonly leading to closure of the trailing marginal basins or narrow seaways (Zagorevski et al., 2007). The latter appears to be the main cause of Ordovician orogenesis along large segments of the Gondwanan margin (Penobscottian, Famatinian-Ocloyic, Sardic, late Ross-Delamerian, and early Lachlan) and probably also the margin of Baltica (Finnmarkian). Why the stresses along these margins switched from extension to transient compression is poorly constrained and not well understood at present. Collins (2002) proposed that compression was due to short-lived periods of flattening of the downgoing oceanic slab as a result of the entrance of relatively buoyant lithosphere, such as oceanic plateaus, into the subduction zone. Although this model is at first sight appealing and may be applicable to some events, it is difficult to accept that flattening of the circum-Gondwanan subduction zones more or less happened coevally (within a period of 20 m.y.) over a strike length exceeding 10,000 km simply because of arrival of buoyant oceanic crust; neither is there much evidence preserved for accretion of oceanic plateaus during the Early Ordovician along the Pacific and Iapetan margins of Gondwana. Perhaps the switch from extension to compression along the Pacific and Iapetan margins was due to far-field stresses generated by Gondwana’s drift and/ or associated rotation during the Ordovician.
While the Ordovician mobile belts along the Gondwanan margins were dominantly non-accretionary orogens sensu stricto (i.e., little or no accretion of truly exotic terranes, but rather re-accretion of rifted-off arc terranes: the extensional accretionary orogens of Collins, 2002), the Iapetan margin of Laurentia was the recipient of intra-Iapetan and/or Gondwana-derived arc terranes and/or microcontinents (van Staal et al., 1998, 2007; Hibbard, 2000; Hibbard et al., 2007; Merschat et al., 2005; Zagorevski et al., 2006, 2007) and hence was steadily growing in size. This process continued throughout the Paleozoic and basically ended with the formation of Pangea during the Permian. ACKNOWLEDGMENTS We thank Jim Hibbard and Monica Escayola for discussions on aspects of Cambrian–Ordovician orogenesis in the Appalachians and the Famatinian and Pampean mountain belts, respectively. Conall MacNiocaill provided us with the original paleomagnetic diagrams published in van Staal et al. (1998), which became the starting points for Figures 1 and 2. Jim Monger, Peter Cawood, and an anonymous reviewer provided insightful comments, all of which improved the manuscript. RDH acknowledges U.S. National Science Foundation grants EAR-8305832, EAR8417894, EAR-8816343, and EAR-9814800. This is Geological Survey of Canada publication 2006371. REFERENCES CITED Andréasson, P.G., Svenningsen, O.M., and Albrecht, L., 1998, Dawn of Phanerozoic orogeny in the North Atlantic tract: Evidence from the Seve-Kalak superterrane, Scandinavian Caledonides: GFF, v. 120, p. 159–172, doi: 10.1080/11035899801202159. Astini, R.A., Benedetto, J.L., and Vaccari, N.E., 1995, The Early Paleozoic evolution of the Argentine Precordillera as a Laurentian rifted, drifted and collided terrane: A geodynamic model: Geological Society of America Bulletin, v. 107, p. 253–273, doi: 10.1130/0016-7606(1995)107<0253: TEPEOT>2.3.CO;2. Bahlburg, H., and Hervé, F., 1997, Geodynamic evolution and tectonostratigraphic terranes of northwestern Argentina and northern Chile: Geological Society of America Bulletin, v. 109, p. 869–884, doi: 10.1130/0016 -7606(1997)109<0869:GEATTO>2.3.CO;2. Belka, Z., Ahrendt, H., Franke, W., and Wemmer, K., 2000, The BalticaGondwana suture in central Europe: Evidence from K-Ar ages of detrital muscovites and biogeographical data, in Franke, W., Haak, V., Oncken, O., and Tanner, D., eds., Orogenic Processes: Quantification and Modelling in the Variscan Belt: Geological Society [London] Special Publication 179, p. 87–102. Bergström, S.M., Finney, S.C., Chen, X., Goldman, D., and Leslie, S.A., 2006, Three new Ordovician global stage names: Lethaia, v. 39, p. 287–288, doi: 10.1080/00241160600847439. Boger, S.D., and Miller, J.McL., 2004, Terminal suturing of Gondwana and the onset of the Ross-Delamerian orogeny: The cause and effect of an Early Cambrian reconfiguration of plate motions: Earth and Planetary Science Letters, v. 219, p. 35–48, doi: 10.1016/S0012-821X(03)00692-7. Bream, B.R., 2003, Tectonic implications of geochronology and geochemistry of para- and orthogneisses from the southern Appalachians crystalline core [Ph.D. thesis]: Knoxville, University of Tennessee, 296 p. Cawood, P., 2005, Terra Australis Orogen: Rodinia breakup and development of the Pacific and Iapetus margins of Gondwana during the Neoproterozoic and Paleozoic: Earth-Science Reviews, v. 69, p. 249–279, doi: 10.1016/ j.earscirev.2004.09.001.
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Global setting of Ordovician orogenesis Torsvik, T.H., and Cocks, R.M., 2005, Norway in space and time: A centennial cavalcade: Norwegian Journal of Geology, v. 85, p. 91–104. Torsvik, T.H., and Rehnstrom, E.F., 2003, The Tornquist Sea and BalticaAvalonia docking: Tectonophysics, v. 362, p. 67–82, doi: 10.1016/S0040 -1951(02)00631-5. Trettin, H.P., 1987, Pearya: A composite terrane with Caledonian affinities in northern Ellesmere Island: Canadian Journal of Earth Sciences, v. 24, p. 224–245. Valverde-Vaquero, P., van Staal, C.R., McNicoll, V., and Dunning, G., 2006, Middle Ordovician magmatism and metamorphism along the Gander margin in Central Newfoundland: Journal of the Geological Society [London], v. 163, p. 347–362, doi: 10.1144/0016-764904-130. van Staal, C.R., 1994, The Brunswick subduction complex in the Canadian Appalachians: Record of the Late Ordovician to Late Silurian collision between Laurentia and the Gander margin of Avalon: Tectonics, v. 13, p. 946–962, doi: 10.1029/93TC03604. van Staal, C.R., 2005, The northern Appalachians, in Selley, R.C., Cocks, R.L.M., and Plimer, I.R., eds., Encyclopedia of Geology: Oxford, UK, Elsevier, v. 4, p. 81–91. van Staal, C.R., Dewey, J.F., MacNiocaill, C., and McKerrow, W.S., 1998, The Cambrian–Silurian tectonic evolution of the northern Appalachians and British Caledonides: History of a complex, west and southwest Pacifictype segment of Iapetus, in Blundell, D.J., and Scott, A.C., eds., Lyell: The Past Is the Key to the Present: Geological Society [London] Special Publication 143, p. 199–242. van Staal, C.R., Whalen, J.B., McNicoll, V.J., Pehrsson, S., Lissenberg, C.J., Zagorevski, A., van Breemen, O., and Jenner, G.A., 2007, The Notre Dame Arc and the Taconic orogeny in Newfoundland, in Hatcher, R.D., Jr., Carlson, M.P., McBride, J.H., and Martínez Catalán, J.R., eds., The 4-D Framework of Continental Crust: Geological Society of America Memoir 200, p. 511–552, doi: 10.1130/2007.1200(26).
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van Staal, C.R., Currie, K.L., Rowbotham, G., Goodfellow, W., and Rogers, N., 2008, Pressure-temperature paths and exhumation of Late Ordovician-Early Silurian blueschist and associated metamorphic nappes of the Salinic Brunswick subduction complex, northern Appalachians: Geological Society of America Bulletin, v. 120, p. 1455–1477, doi: 10.1130/B26324.1. von Raumer, J.F., 1998, The Palaeozoic evolution in the Alps: From Gondwana to Pangea: Geologische Rundschau, v. 87, p. 407–435, doi: 10.1007/ s005310050219. Waldron, J.W.F., and van Staal, C.R., 2001, Taconian orogeny and the accretion of the Dashwoods Block; a peri-Laurentian microcontinent in the Iapetus ocean: Geology, v. 29, p. 811–814, doi: 10.1130/0091-7613(2001)029 <0811:TOATAO>2.0.CO;2. Yoshinobu, A.S., Barnes, C.G., Nordgulen, O., Prestvik, T., Fanning, M., and Pedersen, R.B., 2002, Ordovician magmatism, deformation and exhumation in the Caledonides of central Norway: An orphan of the Taconic orogeny: Geology, v. 30, p. 883–886, doi: 10.1130/0091-7613(2002)030 <0883:OMDAEI>2.0.CO;2. Zagorevski, A., Rogers, N., van Staal, C.R., McNicoll, V., Lissenberg, C.J., and Valverde-Vaquero, P., 2006, Lower to Middle Ordovician evolution of peri-Laurentian arc and back-arc complexes in the Iapetus: Constraints from the Annieopsquotch accretionary tract, Central Newfoundland: Geological Society of America Bulletin, v. 118, p. 324–342, doi: 10.1130/ B25775.1. Zagorevski, A., van Staal, C.R., McNicoll, V.C., and Rogers, N., 2007, Upper Cambrian to Upper Ordovician peri-Gondwanan island arc activity in the Victoria Lake Supergroup, Central Newfoundland: Tectonic development of the northern Ganderian margin: American Journal of Science, v. 307, p. 339–370, doi: 10.2475/02.2007.02. MANUSCRIPT ACCEPTED BY THE SOCIETY 6 OCTOBER 2009
Printed in the USA
The Geological Society of America Special Paper 466 2010
Ordovician explosive volcanism Warren D. Huff† Department of Geology, University of Cincinnati, Cincinnati, Ohio 45221, USA Stig M. Bergström School of Earth Sciences, Ohio State University, 155 South Oval Mall, Columbus, Ohio 43210, USA Dennis R. Kolata Illinois State Geological Survey, 615 East Peabody Drive, Champaign, Illinois 61820, USA
ABSTRACT Explosive eruptions from volcanoes are recorded in the stratigraphic record throughout the Phanerozoic, but evidence of these eruptions in the form of preserved tephra layers appears to be concentrated at times of active plate collision and concomitant high stands of sea level. The products of volcanic eruptions are lavas, tephra, and gases. Basaltic magmas (low-silica content) are usually erupted in the form of lava flows, whereas rhyolitic magmas (high-silica content) are commonly explosively erupted as plinian and ultraplinian plumes, and associated pyroclastic flows. Fallout tephras are preserved in ancient sedimentary sequences as tonsteins, bentonites, and K-bentonites. Middle Ordovician K-bentonites represent some of the largest known fallout ash deposits in the Phanerozoic Era. They cover minimally 2.2 × 106 km2 in eastern North America and 6.9 × 105 km2 in central and northwestern Europe as a result of explosive volcanism, which affected both Laurentia and Baltica during the closure of the Iapetus Ocean. The three most widespread beds are the Deicke and Millbrig K-bentonites in North America and the Kinnekulle K-bentonite in northwestern Europe. Similar successions are well known in South America and China. Sedimentation rates of volcanic ejecta range from meters per year locally to ~1 mm/1000 yr in the deep sea. Volcanogenic sediments react with seawater to produce secondary phases such as zeolites and clay minerals. Studies of recent ashfall behavior suggest that the preservation potential in the stratigraphic record can be viewed as somewhat remarkable in that such sudden events are preserved at all, much less produce such a wealth of valuable geologic information. OVERVIEW
perturb the Earth’s climate on time scales of one to five years, generally resulting in net global surface cooling. Historically documented eruptions do not, however, represent the full range of intensity and magnitude of all explosive eruptions in geologic history. Deposits in the geologic record provide compelling evidence for eruptions that have been orders of magnitude larger than ones witnessed by mankind. Modern studies have shown
Explosive volcanism plays a fundamental role in the exchange of material and energy from the Earth’s interior to the hydrosphere and atmosphere, and major explosive events can †
E-mail:
[email protected].
Huff, W.D., Bergström, S.M., and Kolata, D.R., 2010, Ordovician explosive volcanism, in Finney, S.C., and Berry, W.B.N., eds., The Ordovician Earth System: Geological Society of America Special Paper 466, p. 13–28, doi: 10.1130/2010.2466(02). For permission to copy, contact
[email protected]. ©2010 The Geological Society of America. All rights reserved.
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that lithofacies associations of near-vent subaerial phenomena typically include pyroclastic surge deposits, thin welded tuff beds, various lava flow morphologies, abundant erosional unconformities, and fluvial and laharic facies. While volcano-tectonic subsidence might aid in the preservation of such deposits, they generally are not well known in the geologic record. Ash layers in marine sediments, on the other hand, are the best geologic record of explosive volcanism, and many of the Phanerozoic volcanic ash layers represent enormous explosive eruptions that dispersed fine ash and aerosols through the atmosphere over tens of thousands of square kilometers. Most of the world’s southern landmasses were assembled in the supercontinent of Gondwana at the beginning of the Ordovician (Scotese and McKerrow, 1990). As Gondwana drifted gradually toward the South Pole, fragments separated and moved toward the equator (Cocks and Torsvik, 2005). These included Baltica, Siberia, Avalonia, and Laurentia, and much smaller terranes such as the Precordillera of western Argentina and several blocks in south China (Fig. 1). Laurentia and Baltica are estimated to have moved at rates up to 23 cm/yr (Gurnis and Torsvik, 1994). Much of Ordovician explosive volcanism during this time was associated with closure between these various tectonic units. In the Northern Appalachians the timing of the emplacement of the Taconian allochthons and related low-grade metamorphism is constrained by faunal control and isotopic dating, which suggests emplacement ages of latest Middle Ordovician to Late Ordovician (Sasseville et al., 2003). In the southern and central Appalachians the Taconic terrane collided in a non-orthogonal fashion, first in southwest Virginia and then later in southeast Pennsylvania, and consisted of several discrete terranes (Bronson
Hill, etc.) that collided at different times in the Middle Ordovician (Aleinikoff et al., 2003). The northward drift of the Baltoscandian plate and similar activity along the Asian margin of Gondwana produced similar collisions and resulted in dramatic changes in sea level and related climatic and biological changes (Chen and John, 1998; Nordgulen et al., 2003). Each of these tectonic events included subduction-related explosive volcanism, and each has left a record of these eruptions in the form of K-bentonite beds. The Ordovician Period was not necessarily unusual in this respect. Collision-related volcanic ash beds can be found in all periods of the Phanerozoic and well back into Precambrian time. But some of the Ordovician eruptions were exceptionally large and produced massive amounts of fallout ash that, in a contemporary setting, would have been considered disastrous. A major episode of glaciation was centered in Africa during the Late Ordovician that contributed to ecological disruption and mass extinction. Although explosive volcanism was probably not a major contributor to these events, as the extinction events occurred nearly 10 Ma following the largest eruptive events, it no doubt exacerbated already fragile atmospheric conditions and might well have been a minor factor in driving environmental shifts (Rampino and Self, 1992). Miller (1997) examined faunal diversity patterns for the Ordovician and concluded that, while it is tempting to correlate global patterns of biodiversity with global-scale environmental processes, a careful examination of the data argues more strongly for a signal that is, in reality, “an aggregation of patterns and processes that are unique to particular regions or scales” (p. 100). The closing of the Iapetus Ocean, separating Baltica, Avalonia, and Laurentia, occurred by means of the subduction of
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Figure 1. Paleogeographic map for the Late Ordovician (Tychsen and Harper, 2004).
160°E
Ordovician explosive volcanism oceanic crust beneath, and the consequent collision of, volcanically active island arcs or microplates against the southeastern margin of Laurentia (Scotese and McKerrow, 1991). These collisions were associated with the Taconic orogeny, which began during the Middle Ordovician and produced a complex deformational and sedimentological record that has been extensively documented (Rowley and Kidd, 1981; Stanley and Ratcliffe, 1985; Tucker and Robinson, 1990) and which includes numerous K-bentonite beds in both the eastern North American, British, and Baltoscandian successions. Baltica was surrounded by a passive margin during the Middle Ordovician, but it apparently was in close proximity to Laurentia (Cocks and Torsvik, 2005; Huff et al., 1992; McKerrow et al., 1991). Consequently we attribute the origin of the ~150 Middle Ordovician ash beds in southern Sweden, including the Kinnekulle K-bentonite, to be the explosive volcanic activity in the magmatic arcs associated with the Taconic orogeny. The Ordovician record of explosive volcanism consists of examples of both near-vent pyroclastic flows and ignimbrites and distal sequences of altered fallout tephras known as K-bentonites. Middle Ordovician K-bentonites represent some of the largest known fallout ash deposits in the Phanerozoic Era (Huff et al., 1996). They cover minimally 2.2 × 106 km2 in eastern North America and 6.9 × 10 5 km2 in northwestern Europe as a result of explosive volcanism that affected both Laurentia and Baltica during the closure of the Iapetus Ocean. The three most widespread beds are the Deicke and Millbrig K-bentonites in North America and the Kinnekulle K-bentonite in northwestern Europe. The source vents are thought to have been near the Laurentian margin of Iapetus on an arc or microplate undergoing collision with Laurentia. The volume of ash preserved in the stratigraphic record, converted to dense rock equivalent (DRE) of silicic magma, is minimally estimated to be 943 km3 for the Deicke, 1509 km3 for the Millbrig, and 972 km3 for the Kinnekulle (Huff et al., 1996). One means of expressing the magnitude of an eruption is by use of the volcanic explosivity index (VEI). The VEI was proposed in 1982 as a way to describe the relative size or magnitude of explosive volcanic eruptions (Newhall and Self, 1982). It is a 0-to-8 index of increasing explosivity, and each successive increase in number represents an increase of about a factor of ten. In terms of the VEI these eruptions would probably be ranked between 7 and 8. The Millbrig and Kinnekulle beds are coeval and possibly equivalent, yielding a combined DRE volume of nearly 2500 km3. Some unknown but probably large amount of additional ash fell into oceanic regions of the Iapetus, but these areas became subducted, and the ash is not preserved in the geologic record. The symmetry of the thickness contours (Huff et al., 1992) is suggestive that one or more ash clouds interacting with equatorial stratospheric and tropospheric wind patterns dispersed pyroclastic material to both the northwest and southeast, in terms of Ordovician paleogeography. Well-documented examples of near-vent facies are relatively few, however, and are briefly summarized below. Therefore, the primary focus here will be on the
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stratigraphic record of distal ashes and their application to stratigraphic, tectonic, and paleogeographic reconstructions. PROXIMAL FACIES Some of the most carefully documented examples of Ordovician near-vent facies have been described in North Wales (Branney and Kokelaar, 1992; Kokelaar et al., 1984; Kokelaar and Königer, 2000), where units such as the Caradocian Pitts Head Tuff represent episodes of felsic arc magmatism. An Early Ordovician episode of subaerial arc magmatism was, in turn, succeeded by extension-related marine-marginal-basin magmatism that produced thick intercalations of volcanic rocks and coeval intrusions (Kokelaar, 1986). Two main subsequent cycles of Ordovician volcanism are distinguished within the Late Ordovician successions, the Llewelyn Volcanic Group and the Snowdon Volcanic Group. In the English Lake District the onset of volcanism occurred in the Late Ordovician and produced the Eycott and Borrowdale Volcanic Groups (Fitton et al., 1982). The Borrowdale Volcanic Group of northwest England represents part of a Late Ordovician ensialic arc related to subduction beneath the Avalonian margin of the Iapetus Ocean during late stages of closure between the continents of Laurentia and Avalonia (Branney, 1991; Millward and Evans, 2003; Millward et al., 1978). The volcanic succession includes subaerially erupted calc-alkaline basaltic, andesitic, and rhyolitic lavas, and pyroclastics with continental margin, subduction-related geochemical affinities. A similar succession known as the Komarov Volcanic Complex (Darriwilian) occurs in the Prague Basin (Chlupácˇ et al., 1998). In the eastern Lachlan Fold Belt in New South Wales, mantlederived shoshonitic magmatism commenced in the Early Ordovician and was replaced by tholeiitic magmatism by Late Ordovician time (Wyborn, 1992). The favored model for the development of the shoshonites is from melting of subcontinental lithosphere that was previously enriched in incompatible elements (notably Ba, K, Sr, and P). The prior enrichment is thought to have taken place by subduction processes in the Cambrian, when basement blocks defined by granite studies were being assembled (Gray and Foster, 2004; Keay et al., 1997). In this scenario, melting in the Ordovician subcontinental lithosphere was triggered by overturning associated with asthenospheric upwelling. Subsequent Silurian and Early Devonian volcanism became strongly felsic as a consequence of the increasing involvement of crustal melting in magma generation. In North America the Late Ordovician Ammonoosuc Volcanics and overlying Partridge Formation have been interpreted to be Ordovician island arc volcanics formed during the Taconic orogeny and emplaced as part of the Bronson Hill Anticlinorium (Hollocher, 1993; Karabinos et al., 1998; Leo, 1985; Schumacher, 1988). These units contain substantial quantities of metamorphosed bimodal mafic and felsic volcanics, including an array of pyroclastics, that have been dated at 453–449 Ma (Tucker and Robinson, 1990). They are thought to have been built on a
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margin of Avalonia that collided with the Laurentian margin in Late Ordovician or Early Silurian time (McKerrow et al., 1991). Although these rocks are of the same general age as several widespread distal airfall tephras, now altered to K-bentonites, in the eastern Midcontinent, no correlations have been made that would suggest that the Bronson Hill was a source area for these ashes. DISTAL FACIES Explosive eruptions, particularly those involving magmawater interaction and/or production of large ignimbrites, can generate huge volumes of mostly ash-grade tephra that is widely transported in the atmosphere (Walker, 1981). This material eventually settles out to form distinctive layers that can be correlated across thousands to millions of square kilometers and spanning different depositional environments to form chronostratigraphic markers (Bergström et al., 1995; Drexler et al., 1980; Huff et al., 1996; Rose and Chesner, 1987). Such tephra layers provide information on source magma composition, eruption dynamics, and chronology as deduced from mineral chemistry, granulometric properties, and thickness and distribution patterns (Fredlund et al., 2000; Ninkovich et al., 1978). Finer material (ash <2 mm and lapilli 2–64 mm) is convected upward in an eruption column (Self and Walker, 1994) before settling out downwind to form pyroclastic fall deposits. These deposits are composed of various proportions of vitric, crystal, or lithic particles. Vitric particles are glass shards or pumice derived from magma, whereas crystals are minerals derived from phenocrysts or microlites developed in the magma. Different minerals reflect the composition of different magmas. The most common minerals are shown in Table 1. Thickness and median grain size of ash deposits generally decrease exponentially with distance from a volcano (Walker, 1971). The distribution of ash will depend on the initial grainsize distribution of the ejecta (reflecting fragmentation during the eruption), dynamics of the eruption column, and the column’s interaction with wind (Carey and Sparks, 1986; Sparks et al., 1992). The glass component of the deposited tephra is unstable in most sedimentary environments, and within a relatively short
time following fallout it will begin to devitrify. For example, Hodder et al. (1993) and Naish et al. (1993) reported on studies from North Island, New Zealand, that showed that a downcore decrease in the abundance of volcanic glass is matched by a downcore increase in the abundance of smectite, supporting an early diagenetic origin for the smectite from glass. The transformation is described by a two-stage model involving a combination of parabolic and linear kinetics, reflecting both the hydration of glass and the formation of smectite. Their model indicates that the half-life for volcanic glass is only 1500 yr. Alteration products will depend upon the geochemical nature of the depositional environment, but generally they are zeolites and clay minerals (Hay and Guldman, 1987). Altered ash falls are known in the stratigraphic record as bentonites and K-bentonites. The latter most commonly occur in pre-Mesozoic strata and reflect a more advanced state of diagenetic alteration in which smectite has been metasomatized partially or completely to illite. Paleozoic K-bentonites consist predominantly of clay minerals, dominantly illite and smectite, derived from the breakdown of vitric pumice fragments. However, a small proportion of the beds consists of unaltered primary phenocrystic mineral grains whose morphology, composition, and mineral character retain information generated by the magmatic and tectonic origins of the source volcanoes. FIELD IDENTIFICATION OF K-BENTONITES K-bentonites in unmetamorphosed stratigraphic sections are characterized by a soapy to waxy texture when wet and display a range of colors including light to dark shades of gray, buff, orange, red, and green, depending upon the composition, the activity of circulating groundwater and the extent of local weathering reactions. In cores and subsurface exposures in which there has been less oxidation, K-bentonites are more commonly green to greenish gray to bluish gray (Fig. 2A). Some K-bentonites contain conspicuous and abundant euhedral to anhedral biotite flakes along with quartz, feldspar, amphibole, zircon, and apatite. The typical appearance of a K-bentonite bed in outcrop is that of a fine-grained clay-rich band ranging between 1 mm and 2m in
TABLE 1. COMPOSITION OF MAJOR PHENOCRYST PHASES IN MAGMA Basalt Basaltic Andesite Dacite Rhyolite andesite Plagioclase ** *** *** *** ** Olivine ** ** * Pyroxene ** ** ** * Hornblende * * ** ** * Biotite * ** ** Alkali feldspar * ** *** Quartz ** *** Zircon * ** *** Apatite * ** *** Fe-Ti oxide ** ** * Note: This table is adapted from Schmincke (2004). Mineral occurrences: abundant—***; common—**; rare—*; absent to rare— -.
Ordovician explosive volcanism
17
M D
B
A
M
D C
D
Figure 2. (A) Core from northern Estonia, showing a K-bentonite in an Ordovician carbonate section. (B) Road cut at Gladeville, Tennessee, with Deicke (D) and Millbrig (M) K-bentonites. (C) Deicke and Millbrig at Minke Hollow, Missouri. (D) Deicke at Carthage, Tennessee.
thickness that has been deformed by static load from the enclosing siliciclastic or carbonate sequence (Fig. 2B–2D). Accelerated weathering of K-bentonites causes them to be recessed into the outcrop face. The clay minerals tend to be primarily mixed-layer illite-smectite with lesser amounts of kaolinite (Fig. 3). With increasing burial metamorphism, corrensite (chlorite-smectite)
becomes common. For thicker K-bentonites there is often a zone of nodular or bedded chert in the adjacent strata at both the base and the top of the bed. Although cherts associated with K-bentonites are generally considered to reflect the destiny of silica released during ash devitrification, this relationship has not been thoroughly documented.
18
Huff et al.
33.3Å 9.9Å
7.14Å
3.33Å 3.58Å
5.25Å
12.9Å
5.02Å ARG_22
ARG_23
ARG_47
ARG_50
Figure 3. Sequence of X-ray-diffraction tracings of oriented, glycol-saturated K-bentonite clays from Argentina. Typical of many K-bentonites, these samples range from R0 ordered (ARG_6) to R3 ordered (ARG_22) illite-smectite. Kaolinite is also present in some samples, with peaks at 7.14Å and 3.58Å.
ARG_52 ARG_32
ARG_5B
ARG_6
5
10
15
°2Θ
20
25
30
ORDOVICIAN K-BENTONITES
North America
As with every Phanerozoic System, many Ordovician successions contain a number of K-bentonites representing episodes of explosive volcanism, most commonly associated with collisional tectonic events. Figure 4 shows the global stratigraphic and geographic distribution of K-bentonite beds that have been reported in the literature. Numerous beds have been reported from North and South America, Asia, and Europe. A brief review will summarize some of the pertinent literature for each region.
The Ordovician successions of North America are known to contain nearly 100 K-bentonite beds, one or more of which are distributed over 1.5 × 106 km2 (Kolata et al., 1996). The first report of an Ordovician K-bentonite in North America was made by Ulrich (1888), who described a thick bed of clay in the upper part of what is now known as the Tyrone Limestone, near High Bridge, Kentucky. The Tyrone Limestone, along with the Camp Nelson Limestone and the Oregon Formation, constitute the
UK GLOB. STAGE SER. SER.
UPPER ORDOVICIAN
MIDDLE ORDOVICIAN
LOWER ORDOVICIAN
HIRNANTIAN
KATIAN
SANDBIAN
DAPINGIAN DARRIWILIAN
FLOIAN
TREMADOCIAN
ASHGILL
CARADOC
LLANVIRN
ARENIG
North America
England Wales Norway
Sweden
East
Baltic
Poland
Figure 4. Global stratigraphic distribution of Ordovician K-bentonites.
N. Ireland Scotland China Argentina
Carnic Alps
20
Huff et al.
High Bridge Group of Late Ordovician age. Subsequent work by Nelson (1921, 1922) showed that the bed was volcanogenic in origin and that it could be correlated into Tennessee and Alabama. From the 1930s on, K-bentonite beds of Ordovician age began to be reported from localities throughout eastern North America (Brun and Chagnon, 1979; Huff and Kolata, 1990; Kay, 1935; Kolata et al., 1996; Weaver, 1953). Their clay mineralogy is typically dominated by a regularly interstratified illite-smectite in which the swelling component accounts for 20–40% of the total structure. Two prominent K-bentonites occur within the Tyrone Limestone of central Kentucky (Fig. 5). The Millbrig or “Mud Cave” K-bentonite of local drillers is found at or near the contact between the Tyrone Limestone and the overlying Curdsville Limestone Member of the Lexington Limestone. In parts of the region the Millbrig has been removed along the preLexington disconformity (Cressman, 1973) and hence has somewhat restricted usefulness in regional correlation. The equivalent bed in the Carters Limestone of central Tennessee is the T-4 bed of Wilson (1949). The Deicke or “Pencil Cave” K-bentonite of local usage occurs ~4–6 m below the top of the Tyrone Limestone and varies in thickness from a few centimeters to 1.5 m. Some reworking of the original ash by waves and bottom currents undoubtedly occurred. However, the influx of terrestrial clastics was so minimal as to preclude contamination of the K-bentonite by anything other than carbonate mud. The Deicke is the most persistent K-bentonite marker in the area. Its equivalent in central Tennessee is the T-3 bed (Wilson, 1949). The chemical characteristics of the K-bentonite beds along the Cincinnati Arch were reported by Huff and Türkmenoglu (1981). Using immobile trace elements, the Deicke and Millbrig have been correlated by chemical fingerprinting from southeastern Minnesota to southeastern Missouri (Kolata et al., 1987) and by wireline logs from Missouri to the southern Appalachians and into the Michigan Basin and southern Ontario (Huff and Kolata, 1990; Kolata et al., 1996). Both beds range from 1.5 m or more in thickness in the southern Appalachians (Haynes, 1994) to 3 cm or less in western Iowa. Unpublished data from wireline logs and recent studies of the Bromide Formation in southern Oklahoma suggest that both beds extend farther to the southwest than has previously been mapped. K-bentonites include a thin (5–6 cm) widespread but locally absent bed ~24 m below the top of the Tyrone Limestone, another thin bed between the Deicke and Millbrig, and a bed in the Curdsville Member of the Lexington Limestone that Conkin and Conkin (1992) labeled the Capitol Metabentonite. Huff et al. (1996) calculated the DRE values to be 943 km3 for the Deicke and 1509 km3 for the Millbrig. Kolata et al. (1996) documented the stratigraphic distribution of at least seven named K-bentonite beds traceable throughout the mid-Mohawkian of the upper Mississippi Valley region and subsequently named them the Hagan K-bentonite complex (Kolata et al., 1998). The youngest of the series, the House Springs K-bentonite, can be correlated as far as the southern Appalachians on the basis of both distinctive outcrop and wireline log characteristics. The Hagan complex is a sequence of altered fall-
out ashes that originated from one or more volcanic centers in the southern Appalachian Mountains. The Deicke and Millbrig beds are part of the complex and, to date, have the widest known distribution. Recent discoveries in the Bromide Formation of Oklahoma suggest that both the Deicke and Millbrig are continuous into that area (Leslie et al., 2006). A unique feature of the Hagan complex is that it straddles the Black River–Trenton unconformity, generally believed to be a significant stratigraphic sequence boundary in the eastern United States (Holland and Patzkowsky, 1996) and is traceable throughout much of the southern and central Appalachians. South America Ordovician K-bentonites have been recognized since 1994 in the upper San Juan Limestone and overlying clastic strata of the Gualcamayo and Los Azules Formations in the Argentine Precordillera (Huff et al., 1998). They have been recorded from more than 20 localities, mainly in the eastern thrust belts, in the San Juan Limestone and the overlying Gualcamayo Formation (Fig. 6), but a few ash beds are known also from the central thrust belts. The oldest occur in the upper Lower Ordovician Oepikodus evae conodont Zone and the youngest in the middle Darriwilian Pterograptus elegans Zone. The older beds are in the topmost part of the Lower Ordovician San Juan Limestone in outcrops along the Gualcamayo River in the Guandacol region. In the overlying, ~200 m thick Gualcamayo Formation are numerous ash beds, especially along a small tributary to the Guandacol River where more than 170 separate ash beds have been observed (Huff et al., 1998). The K-bentonite interval contains graptolites of the Isograptus victoriae maximus Zone (Ortega et al., 1993) and is of early Middle Ordovician age (Dapingian and Darriwilian/Global Stages). The top part of the underlying San Juan Limestone contains conodonts of the Oepikodus evae Zone (Hünicken and Pensa, 1989). Because the latter zone is elsewhere no younger than the Isograptus victoriae lunatus Zone, the contact between the San Juan Limestone and the Gualcamayo Formation at this locality may be unconformable, and the succession may have a stratigraphic gap corresponding to the Isograptus victoriae victoriae Zone (Bergström et al., 1996). The widespread occurrence of K-bentonite beds in the Argentine Precordillera constitutes one of the most extensive suites of such beds known anywhere in the Ordovician System of the world and serves as testimony to the high intensity of explosive volcanism along this margin of Gondwana during the early and middle parts of that period. Previous and ongoing studies of the sedimentology, mineralogy, and geochemistry of these beds provide both insight and constraints concerning the magmatic, tectonic, and paleogeographical settings under which the explosive volcanism was generated, and also permit comparisons with lower Paleozoic K-bentonites on other continents. Although recent field work has revealed an extensive succession of K-bentonite beds in the exposures along the Gualcamayo River and its tributaries in western La Rioja Province, most of
Oregon Limestone
40 m
K-b 60 m
20 m
K-b
40 m
K-b
Oregon Limestone
GROU P BRI D GE
50 m
10 m
HI GH
Tyrone Limes tone
Deicke
Camp Nelson Limestone
Millbrig
BRIDGE
GROU P
30 m
H IGH
70 m
LEXINGTON LIMESTONE
Logana Mbr. Capitol
C urd s v ille L im es ton e M em b er
80 m
0m
Figure 5. Stratigraphic column of a portion of the upper Ordovician High Bridge Group and Lexington Limestone in central Kentucky, showing K-bentonite zones, either named or indicated by the K-b symbol. The sequence consists mainly of micritic and heavily burrowed limestone and dolomite.
22
Huff et al.
Figure 6. Stratigraphic sections from the eastern thrust belts of the Precordillera, showing the distribution of K-bentonite (K-b) beds.
our detailed particle-size, mineralogical and geochemical studies to date have been of samples from the extensive sections at Cerro Viejo, near Jáchal, and at Talacasto, north of San Juan, in San Juan Province. Further, whereas most known evidence for pre-Andean explosive volcanism on the Gondwana margin is preserved in the Ordovician sections of the Argentine Precordillera, additional beds of pyroclastic origin have also been reported from the Balcarce Formation of the Tandilia region, south of Buenos Aires (Dristas and Frisicale, 1987). Trace fossils have traditionally been used to assign the Balcarce Formation to the
Lower Ordovician, owing to the presence of Cruziana furcifera (Poiré et al., 2003). At least one, and perhaps as many as four, altered pyroclastic beds occur in the white quartzite sequence that ranges from 18 to 500 m in thickness and unconformably overlies Precambrian basement (Dalla Salda et al., 1988). In contrast to the illite-smectite-rich beds of the Precordillera, the Balcarce beds consist mainly of well-crystallized kaolinite with scattered crystals of altered ilmenite, and are considered to be tonsteins that are the product of altered mafic ashes (Dristas and Frisicale, 1987).
Ordovician explosive volcanism The Argentine sequence is nearly unique in both the number and lateral distribution of K-bentonite beds, and geochemical and grain-size data are consistent with a source area along the Gondwana margin, such as the Puna-Famatina arc (Huff et al., 1998). They provide no supporting evidence of a close association between the Precordillera and Laurentian sedimentary basins at that time, as has been proposed by Thomas and Astini (1996).
23
suggesting a possible link with exotic units of the Scandinavian Caledonides. In the Late Ordovician to Early Silurian (ca. 450–430 Ma) gabbroic to granitic plutons were emplaced into the earlier assembled oceanic and continental rock units. The plutons show evidence of mixed crust and mantle sources and probably represent continued magmatism along the Laurentian margin. These collisions were associated with the Taconic orogeny, which began during the Middle Ordovician and produced a complex deformational and sedimentological record that has been extensively documented (Rowley and Kidd, 1981; Stanley and Ratcliffe, 1985; Tucker and Robinson, 1990) and which includes numerous K-bentonite beds in both the eastern North American, British, and Baltoscandian successions. Baltica was surrounded by a passive margin during the Middle Ordovician, but it apparently was in close proximity to Laurentia (Cocks and Torsvik, 2005; Huff et al., 1992; McKerrow et al., 1991). Consequently we attribute the origin of the ~150 Middle-Upper Ordovician ash beds in southern Sweden, including the Kinnekulle K-bentonite, to be the explosive volcanic activity in the magmatic arcs associated with the Taconian orogeny. Figure 8 shows the comparative stratigraphic distribution of K-bentonites in Baltoscandia and in North America. Many beds are coeval, most notably the Viruan Kinnekulle bed in Baltoscandia and the Mohawkian Millbrig bed in North America. The possibility of a common source for the Millbrig and Kinnekulle giant ash beds was suggested by Huff et al. (1992), but subsequently
Northern and Central Europe The closing of the Iapetus Ocean separating Baltica, Avalonia, and Laurentia (Fig. 7) occurred by means of the subduction of oceanic crust beneath, and the consequent collision of, volcanically active island arcs or microplates against the southeastern margin of Laurentia (Scotese and McKerrow, 1991). An Early to Middle Ordovician (ca. 480–465 Ma) magmatic and tectonometamorphic event is well documented in the Karmøy-Bergen area (southern Norway) and in the Helgeland Nappe Complex (Uppermost Allochthon, north-central Norway) (Nordgulen et al., 2003). Granitoids related to this event are peraluminous and were essentially produced during high-grade metamorphism and anatexis of crustal protoliths during collision and amalgamation of continental fragments above an east-dipping (present coordinates) subducting slab near the Laurentian margin. The northwestern parts of the Scottish and Irish Caledonides, which were derived from the Laurentian margin, underwent time-equivalent deformation, metamorphism, and magmatism,
o
0
Equator
Baltica
Figure 7. Closing of the Iapetus Ocean during the Late Ordovician initiated a line of volcanic arcs parallel to the margins of Laurentia, Avalonia, and Baltica. Yellow areas represent the sites of the most active explosive volcanism, and arrows indicate the direction of ash transport (Bergström et al., 1997a).
Scotland
Laurentia
N. Ireland nt n na cea UK m Re tus O e S. Ireland n Iap ea ia Oc n c ei lo Rh va
A
Nova Scotia
0
o
30 S 1000 km
H. teretiusculus
N. gracilis
C. bicornis
C. americanus
ab. 10 beds
ab. 9 beds
ab. 13 beds K
Darriwilian
Sandbian
D
M
Katian
H. teretiusculus
N. gracilis
Many thin beds
D. foliaceus
D. clingani
P. linearis
P. serra
P. anserinus
A. tvaerensis
A. superbus
Figure 8. Stratigraphic distribution of Ordovician K-bentonites in North America and Europe. Symbols indicate Millbrig (M), Deicke (D), and Kinnekulle (K) K-bentonites. Pertinent graptolite and conodont zones are also shown.
friendvillensis
“sweeti”
aculeata
quadridactylus
compressa
undatus
tenuis
C. spiniferus
confluens O. ruedemanni
G. pygmaeus
velicuspis
robustus
grandis
complanatus ordovicicus
HARJUAN VIRUAN
A. manitoulinensis
BALTOSCANDIC NORTH AMERICAN NORTH MIDCONTINENT ATLANTIC BALTIC BRITISH GLOBAL AMERICAN CONODONT GRAPTOLITE CONODONT GRAPTOLITE STAGES K-BENTONITE SERIES SERIES K-BENTONITE SERIES ZONES ZONES ZONES ZONES Hirnantian shatzeri D. D. D. complanatus divergens
C I N C I N N AT I A N
MOHAWKIAN
WHITEROCKIAN
ASHGILL
CARADOC LLANVIRN
Ordovician explosive volcanism questioned by Haynes et al. (1995). Close examination in the field shows that both beds consist of several internally graded units, suggesting that each bed represents the cumulative deposition of multiple ashfalls in environments characterized by low background sedimentation rates. This aspect was examined in some detail by Kolata et al. (1998) and Haynes (1994) for the Millbrig and by Huff et al. (1999) for the Kinnekulle, all of whom showed systematic mineralogical and grain size variation within individual subunits. Given higher rates of sediment accumulation, it is conceivable that these units could be preserved as a series of closely spaced coeval beds (Bergström et al., 1997b; Huff et al., 1999). The stratigraphic position of these beds in England and Wales is essentially occupied by the massive Snowdon and Borrowdale volcanics of north Wales and the English Lake District, as described above, although a possible occurrence of the Kinnekulle K-bentonite in central Wales was reported by Bergström et al. (1995). The Middle Ordovician section at Röstånga in Scania (southern Sweden) contains 18 K-bentonite beds ranging from 1 to 67 cm in thickness, and all occur within the D. foliaceus (formerly multidens) graptolite Biozone. At Kinnekulle, 290 km to the north, this interval includes the type section of the Kinnekulle K-bentonite, which is widespread and has been correlated throughout northern Europe (Bergström et al., 1995). In most sections the Kinnekulle K-bentonite can be recognized by distinctive geochemical fingerprints, its prominent thickness, and its biostratigraphic and lithostratigraphic position (Bergström et al., 1995). However, at Röstånga, whole rock chemistry is inconclusive at identifying which of the 18 beds is the Kinnekulle K-bentonite. Several beds at Röstånga correlate equally well with the Kinnekulle bed (Bergström et al., 1997b) and thus argue strongly for the composite nature of what is called the Kinnekulle K-bentonite. The Deicke, on the other hand, appears to be a single event deposit, but it has not been recognized in Europe. The regional aspects of ash accumulation on submarine surfaces was discussed by Kolata et al. (1998) and Ver Straeten (2004). The Millbrig in eastern North America and the Kinnekulle in northern Europe both display macroscopic and microscopic evidence of multiple event histories, a characteristic that is explainable only by invoking a history of closely spaced episodic ash accumulations in areas with essentially no background sedimentation (Kolata et al., 1998; Ver Straeten, 2004). Parts of the Millbrig and Kinnekulle beds contain biotite grains that are compositionally indistinguishable from one another, although the majority of samples analyzed show a clear distinction between the two beds on the basis of Fe, Mg, and Ti ratios. Tectonomagmatic discrimination diagrams, combined with Mg number data, indicate that the Deicke-Millbrig-Kinnekulle sequence represents the transformation from calc-alkaline to peraluminous magmatic sources, consistent with a model of progressively evolving magmatism during the closure of the Iapetus Ocean (Huff et al., 2004). Published isotopic age dates are inconclusive as to the precise
25
age of each bed. Thus, it appears that the Millbrig and Kinnekulle beds are coeval and represent separate but simultaneous episodes of explosive volcanism, although it cannot be excluded that parts of these beds were derived from the same eruption(s). Similar intercontinental correlations elsewhere in Europe or China have not yet been attempted. Although most Ordovician K-bentonites reported in Europe are from the British Isles (Fortey et al., 1996; Huff et al., 1993) and Baltoscandia (Bergström et al., 1995; Bergström et al., 1997b), there are also occurrences in Poland (Tomczyk, 1970), the Carnic Alps (Histon et al., 2003), and Lithuania (Sliaupa, 2000). The Alpine orogen represents a collage of Alpine and preAlpine crustal fragments. Schönlaub (1992, 1993) showed that some fragments reflect a true odyssey of near global wandering. These segments have been dated as ranging from Late Ordovician to the Permian on the basis of various rather well-known climate-sensitive biofacies and lithofacies markers, thus adding further to the controversy with regard to the paleogeography and the relationship of the Paleozoic proto-Alps and the coeval neighboring areas such as Baltica, the British Isles, the Prague Basin (Barrandian), Sardinia, Southern France and Spain, and North Africa. Ninety-five K-bentonite levels have been recorded to date from the Upper Ordovician (Ashgill) to Lower Devonian (Lochkov) sequences of the Carnic Alps (Histon et al., 2003). These bentonites occur in shallow- to deep-water fossiliferous marine sediments, suggesting a constant movement from a moderately cold climate of ~50° south latitude in the Late Ordovician to the Devonian reef belt of some 30° south. China The first Ordovician K-bentonite recognized in China (Ross and Naeser, 1984) was a single bed in the Upper Ordovician Wufeng Shale. Subsequently, Huff and Bergström (1995) reported two beds in the Lower Ordovician Ningkuo Formation at Hentang in the Jiangshan Province, southeastern China. More recently, a number of K-bentonite beds have been recognized in the Ordovician–Silurian transition (Ashgill–early Llandovery) in the Yangtze Block, south China (Su et al., 2004). A preliminary analysis of the geochemical composition of the K-bentonites has suggested a parental magma origin of trachyandesite to rhyodacite with some rhyolite in general, which came from volcanic-arc and syn-collision to intra-plate tectonic settings. Regional correlation of these K-bentonite beds indicates that they clearly have the potential of increasing southeastward, both in thickness and grain size. These features suggest that the original volcanic ash may have come from the southeastern part of present-day south China. In addition, along the southeast margin of the Yangtze Block, typical flysch successions have also been identified both from the Zhoujiaxi Group (early Llandovery) and Tianmashan Formation (Ashgillian) in the southern part of Hunan Province, south China. Geochemical analysis of the silicate minerals has
26
Huff et al.
suggested that the flysch successions were deposited in the basin on a passive continental margin (Fletcher et al., 2004). Field observations of the paleocurrents, cross-bedding, ripple marks, and flute marks all suggest that the detrital components must have been transferred from the southeast part of present south China, in good agreement with the conclusion drawn from the analysis of the K-bentonites as mentioned above. Furthermore, the flysch successions both in the Tianmashan Formation and the Zhoujiaxi Group clearly show a northwestward progradation in space and time during the Ordovician–Silurian transition. Both the K-bentonites and flysch successions could be regarded as the products from responding distantly of the area to the continuous northwestward collision and accretion process of the Cathaysia Block to the Yangtze Block (Su et al., 2004). DISCUSSION AND SUMMARY One of the important findings made by the Deep Sea Drilling Project and the Ocean Drilling Program is the discovery that the frequency of ash layers in the sedimentary record at sites close to volcanic arcs is not uniform but highly episodic (Cawood and Leitch, 1997; Pletsch and Reicherter, 2001). Volcanic ash in deepsea sediments may be in discrete layers or dispersed throughout other sediments. Size sorting by the wind may occur with distance from the source. Eruptions into the troposphere (5–12 km altitude) are the most common, with residence times of the ash measured in hours or days. Global ashfalls occur after extremely explosive eruptions inject ash into the stratosphere. Extremely fine-grained ash may remain in the atmosphere for a year or more, and often, prevailing wind patterns will distribute the ash over glacial and arid regions or in the deep ocean. Sedimentation rates of volcanic ejecta range from meters per year locally to ~1 mm/1000 yr in the deep sea. Volcanogenic sediments react with seawater to produce clay minerals. Against this backdrop of the studies of recent ashfall behavior the preservation potential in the stratigraphic record can be viewed as somewhat remarkable in that such sudden events are preserved at all, much less produce such a wealth of valuable geologic information. Burial compaction and devitrification of airfall tephra will typically reduce the thickness of an ash bed by a factor of 3–4× (Huff et al., 1996). Thus, many K-bentonites whose stratigraphic thicknesses are measured in millimeters and yet can be traced for tens or even hundreds of kilometers clearly represent the remains of what, by contemporary standards, would be considered large-magnitude eruptions. What is most remarkable about many Phanerozoic K-bentonites in general, and Ordovician beds in particular, is that their preservation could only occur with relative highstands of sea level and widespread epicontinental flooding. Kolata et al. (1998), and more recently Ver Straeten (2004), have provided compelling arguments for the selective preservation of fallout ash in marine sequences during flooding and highstand intervals. Thus, many of the stratigraphic characteristics of K-bentonites that we see in Ordovician sequences likely represent more general features of ash beds throughout the Phanerozoic.
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MANUSCRIPT ACCEPTED BY THE SOCIETY 6 OCTOBER 2009
Printed in the USA
The Geological Society of America Special Paper 466 2010
Toward identifying potential causes for stratigraphic change in subtropical to tropical Laurentia during the Mohawkian (early Late Ordovician) Achim D. Herrmann† Barrett Honors College, Arizona State University, Tempe, Arizona 85287, USA Bernd J. Haupt Earth and Environmental Systems Institute, The Pennsylvania State University, University Park, Pennsylvania 16802, USA
ABSTRACT Numerical models of the ocean-climate system indicate that during the early Late Ordovician, water from the higher southern latitudes flowed north toward the equator. The cold-water masses welled up into and penetrated the epicontinental sea of Laurentia. The “cold-water conditions” existed despite high levels of pCO2 (~15× preindustrial atmospheric levels) and did not necessarily indicate the onset of glaciation during the early Late Ordovician; rather the cold-water conditions may indicate the onset of a cooling event that plunged the Ordovician Earth system toward icehouse conditions that would lead later to the end-Ordovician (Hirnantian) glaciation. Furthermore, the observed distribution of cold-water masses across the southeastern margin of Laurentia is consistent with the interpretation that a cold-water event caused a regional extinction in the Mohawkian of eastern Laurentia.
margin of Laurentia indicate upwelling and initiation of glaciation prior to the Hirnantian (ca. 445–443 Ma; Webby et al., 2004). Furthermore, a prolonged glacial episode was suggested based on glaciogenic sediments (Hamoumi, 1999), and early Late Ordovician carbonates from the Baltic platform suggest a temperature drop of as much as 15 °C (Tobin et al., 2005). Identifying the timing of the onset of the Late Ordovician glaciation and the dynamics of the Ordovician climate system leading up to the glaciation is important, because the glaciation coincided with the second largest mass extinction during the Phanerozoic (Sepkoski, 1996). Sedimentological and faunal changes indicate a transition from warm-water to cool-water conditions within the shallow epicontinental sea of Laurentia during the early Late Ordovician (Brookfield and Brett, 1988; Brookfield and Elgadi, 1998;
INTRODUCTION It is generally thought that the Ordovician was characterized by high pCO2 levels (14–16 times preindustrial levels: e.g., Berner and Kothavala, 2001; Crowley and Baum, 1995; Yapp and Poths, 1992) and a warm, greenhouse climate (Barnes, 2004). However, a very short-lived glaciation (~1 m.y.) was proposed based on oxygen and carbon isotope stratigraphy and eustatic sea-level changes during the latest stage of the Ordovician (Brenchley et al., 1995, 2003). The proposed short duration of the glaciation was challenged based on stratigraphic evidence from Late Ordovician deposits. For example, Pope and Steffen (2003) suggested that deposits along the southern †
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Herrmann, A.D., and Haupt, B.J., 2010, Toward identifying potential causes for stratigraphic change in subtropical to tropical Laurentia during the Mohawkian (early Late Ordovician), in Finney, S.C., and Berry, W.B.N., eds., The Ordovician Earth System: Geological Society of America Special Paper 466, p. 29–35, doi: 10.1130/2010.2466(03). For permission to copy, contact
[email protected]. ©2010 The Geological Society of America. All rights reserved.
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Holland and Patzkowsky, 1996; Lavoie and Asselin, 1998; Patzkowsky and Holland, 1996; Pope and Read, 1997). This transition occurs across the M4-M5 sequence boundary (Fig. 1), which is part of 14 sequences that are identified in the Mohawkian strata of eastern North America (Holland and Patzkowsky, 1996). In particular the change to a skeletal association that is reminiscent of brynoderm and bryomol associations of modern and ancient temperate water settings supports the notion that cool-water conditions existed in the epicontinental sea. Changes in lithofacies were accompanied by a regional extinction event that occurred over the eastern United States (Patzkowsky and Holland, 1999) and affected articulate brachiopods (Patzkowsky and Holland, 1993, 1996; Patzkowsky et al., 1997), cephalopods (Frey, 1995), corals (Patzkowsky and Holland, 1996), and crinoids (Eckert, 1988). However, this transition from tropical to cold-water carbonates is problematic because paleogeographic reconstructions and paleoclimatic indicators place eastern North America in a tropical to subtropical region during this time (Scotese and McKerrow, 1990, 1991). Nevertheless, based on sequence stratigraphic evidence and carbon isotope data, Saltzman and Young (2005) suggested that the transition to Late Ordovician icehouse conditions occurred during this time. Therefore, global cooling could have led to cool-water conditions in equatorial regions, similar to other icehouse conditions in the Paleozoic (Beauchamp and Desrochers, 1997; Samankassou, 2002). Alternatively, it has also been proposed by several workers that these linked paleontologic and lithologic changes were caused by changing basin bathymetry owing to the Taconic orogeny, changing ocean circulation, and increased siliciclastic detritus from the Taconic highlands that enhanced water turbidity, thereby potentially suppressing and hampering tropical carbonate production (Holland and Patzkowsky, 1996; Patzkowsky and Holland, 1996). The purpose of this paper is to discuss results of numerical model simulations of
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the global ocean-climate systems in the early Late Ordovician and to address the hypothesis of whether global cooling could have led to a substantial decrease in surface temperature in a shallow ocean that was located in tropical to subtropical regions during this period. Our results indicate lower surface temperatures along the eastern and southern margins of Laurentia, places where cool-water conditions were postulated (e.g., Pope and Steffen, 2003). This supports the idea that global cooling contributed to the observed turnover event in the Mohawkian of eastern North America. METHODS The numerical simulations were performed with the Geophysical Fluid Dynamics Laboratory Modular Ocean Model (MOM) version 2.2 with a horizontal global resolution of 4° (zonal) by 4° (meridional) and 16 vertical, unevenly spaced levels that gradually widen with depth (50 m for the surface layer, 850 m for the bottommost layer). This grid is fine enough to represent the main ocean circulation features, including vertical thermohaline structure (Haupt and Seidov, 2001), necessary for use of a realistic paleogeography and paleobathymetry (Fig. 2). The Gent and McWilliams parameterization (Gent and McWilliams, 1990) for improved mixing along isopycnal surfaces (rather than across isopycnals to fixed-depth averages) associated with mesoscale eddies (50–100 km scales) was used (Pacanowski, 1996). The use of the Gent-McWilliams mixing improves the ability of coarse-resolution ocean circulation models to simulate large-scale aspects of general circulation. The ocean model is driven at the sea surface by annual mean atmospheric boundary conditions (momentum [wind stress], sea surface temperature, and freshwater flux) derived from the output of the atmospheric general circulation model GENESIS, a three-dimensional
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Figure 1. Summary of long-term lithologic changes and biotic turnover for Middle and Upper Ordovician strata of North America. Abrupt shifts in lithofacies occur in the lower Chatfieldian (Chat.) across the M4-M5 sequence boundary. Biotic turnover is shown for articulate brachiopod genera in the eastern United States (Patzkowsky and Holland, 1997). Similar patterns of turnover are known among other benthic invertebrates. Pulses of extinction and migration are coincident with the lithologic changes. Global stages and series are from Finney (2005) and Bergström et al. (2006).
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Figure 2. Paleogeography of the early Late Ordovician (after Scotese and McKerrow, 1990, 1991). Light gray indicates permanently exposed land areas, and flooded areas are outlined with black lines. LA—Laurentia; BA—Baltica; SA—South America; SCE—South Central Europe; SI—Siberia; CH—China; AU—Australia; I—Iapetus; Gond—Gondwana. Note incursion of colder water masses from higher southern latitudes deflected northward toward the southern margin of Laurentia. Simulation was performed with pCO2 level of 15× preindustrial atmospheric levels (PAL) of 280 ppm. Note that 15× PAL is within the proposed range of atmospheric pCO2 during the Late Ordovician (Berner, 1994; Berner and Kothavala, 2001; Yapp and Poths, 1992). Surface ocean circulation patterns of Herrmann et al. (2004a) are indicated by arrows.
atmospheric general circulation model coupled to a 50 m thick slab ocean (Herrmann et al., 2004b). The GENESIS output is averaged (weighted average for monthly output) over the last ten years of model simulation. These surface boundary conditions, which are a function of longitude and latitude, but independent of time, are interpolated from the GENESIS model grid to the MOM grid. The model is run without flux corrections—there are no direct (paleo) observations available—from a state of rest, and no separate spin-up of the ocean is required. Using given prescribed boundary conditions, the three-dimensional (3-D) model ocean adjusts to this particular forcing and produces a unique distribution of salinity, temperature, and velocity. The model is run for an integration period of 10,000 yr. We use built-in model diagnostics, which are large-scale integral quantities like the global energy cycle, gyre components, meridional overturning, and tracer budgets and fluxes to ensure that a steady-state solution is reached. Table 1 shows main surface boundary conditions for the numerical model. Additional details of both model setups and experiments are described in greater detail in Herrmann et al.
(2004a, 2004b). The model results presented here focus on the regional conditions across Laurentia. RESULTS Our numerical model experiments indicate that the Late Ordovician surface ocean was characterized predominantly by zonal circulation patterns in the Northern Hemisphere as a result of zonal wind patterns and a lack of continental barriers. The location of most continents in the Southern Hemisphere obstructed zonal currents and led to the development of large gyre systems in the southern ocean basins between the continents. The gyre system between the eastern margin of Gondwana and Laurentia was characterized by a strong western boundary current along the eastern margin of Gondwana, which transported warm equatorial water masses southward toward higher southern latitudes. Furthermore, surface water masses with lower temperatures from the high latitudes reached subtropical latitudes in the Iapetus region along the southeastern margin of Laurentia (Fig. 2).
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Herrmann and Haupt TABLE 1. SURFACE BOUNDARY CONDITIONS FOR NUMERICAL MODEL Boundar y conditions Atmospheric CO2 4200 ppm (15× preindustrial level) Other trace gas concentrations Methane: 0.65 ppm; N2O: 0.275 ppm; CFCs: 0 ppm Land-sea distribution Paleogeography of Scotese and McKerrow (1991) Orbital parameters Eccentricity: 0.06; obliquity: 22.0; precession 270, perihelion to N.H. vernal equinox Solar luminosity 95.5% of present-day value Note: For additional details on model setup and boundary conditions, see Herrmann et al. (2004a, 2004b).
The global ocean of the early Late Ordovician is typified by warm surface waters and cold bottom waters (Fig. 3A and 3B). The thermal structure is caused by strong vertical downwelling of water masses in the higher southern latitudes (Fig. 4). A less significant downwelling occurs in the high latitudes of the Northern Hemisphere because of fresher (less saline) and thus less dense waters compared with the high latitudes in the Southern Hemisphere. The strong deepwater formation in the Southern Hemisphere led to a predominantly zonal distribution of temperature and salinity, especially in the Northern Hemisphere. Deepwater masses below 2000 m had a temperature range of ~4–5 °C (Fig. 3A). On the shallow epicontinental sea of Laurentia, temperatures were above 20 °C (Fig. 3B). The water temperatures in the deeper ocean basins surrounding Laurentia decreased with depth to ~10 °C at a depth of ~1000 m. DISCUSSION The lithological and geochemical changes in the Mohawkian of Laurentia include (1) a shift from tropical-type to temperatetype carbonates, (2) an increase in fine-grained siliciclastics, (3) an increase in the amount and distribution of phosphatic sediments, and (4) a positive shift in the δ13C of carbonates and organics (Lavoie and Asselin, 1998; Patzkowsky and Holland, 1996; Patzkowsky et al., 1997; Pope and Read, 1997; Saylor et al., 1997). Global interpretations for this transition suggest that the lithologic and oceanographic changes in North America are the result of the onset of icehouse conditions that eventually led to a glaciation in the latest Ordovician (Lavoie and Asselin, 1998; Pope and Read, 1997). Kolata et al. (2001) identified the Sebree Trough as a potential passageway through which cold water from the southern (margin) opening could penetrate deeply into the shallow epicontinental sea of Laurentia. The presence of cold, open-ocean water masses close to the vicinity of the Sebree Trough was subsequently demonstrated by Pope and Steffen (2003). Pope and Read (1997) had speculated previously about the presence of cold-water masses along the southern margin of Laurentia that were carried into lower latitudes. Our simulations of the global ocean circulation support the hypothesis that cold ocean-water masses must have reached the southern margin of Laurentia during the Late Ordovician (Fig. 2). Using pCO2 values as high as 15× the preindustrial level, our modeled climate scenario indicates that the southern part of
the epicontinental sea, where the cooling event of the TurinianChatfieldian has been identified, had surface waters with temperatures below 28 °C. Temperatures above 28 °C occur only much farther north, closer to the equator. The average marine tropical seawater temperature required for precipitation of most carbonate particles ranges from 28 °C to 30 °C (James, 1997; Lees, 1975; Lees and Buller, 1972). At these temperatures, abiotic precipitation of carbonate is possible, and both pellets and ooid/aggregate grains can form. At the lower temperatures predicted by our modeling, carbonate precipitation rates would have been greatly reduced, consistent with the observed lithological and faunal changes in the Late Ordovician. These changes therefore could have been the result of the onset of the global cooling trend. Our numerical model results show that the global ocean was characterized by cold, deep ocean-water masses of 4 °C, with warmer water near the surface. This is in contrast to geochemical data from the Taconic foreland basin, which suggest that the epicontinental ocean was inversely stratified (Railsback et al., 1990). Based on paleoecological studies (Cisne et al., 1982; Cisne and Rabe, 1978), Railsback et al. (1990) suggested that along a depth transect across the eastern margin of Laurentia, warmer, saline bottom waters underlay colder, brackish surface waters. Two alternative explanations can be used to resolve this disagreement. First, the observed geochemical signal from the Taconic foredeep might be related to the outflow of saline waters from the extensive epicontinental shelf of Laurentia, similar to present-day outflow of warm saline waters from the Mediterranean Sea into the Atlantic Ocean (e.g., Johnson, 1997) and from the Persian Gulf into the Indian Ocean (e.g., Pratt et al., 1999). Second, studies by Railsback et al. (1990) and Ackerly et al. (1993) focused on the brachiopod genus Paucicrura. Ackerly et al. (1993) correlate the number of punctae in Paucicrura rogata with water depth and temperature, and showed that water temperature increased with depth. However, dysoxic conditions within the foreland basin may have contributed to reduced growth rates of brachiopods. Dysoxic conditions would have increased the packing density of punctae per unit area, and therefore packing densities would not have been a good predictor for any depthdependent temperature variations. The cold-water conditions would have existed in the epicontinental sea of Laurentia despite high pCO2 values of 15× preindustrial levels. Using an atmospheric general circulation model, Herrmann et al. (2003) showed that at those pCO2 levels,
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accompanied by a high sea level, no ice sheets will form. This indicates that despite the predominance of cold-water masses within the low-latitude shallow epicontinental sea and global cooling, high-latitudinal ice sheets are not necessary (Tobin et al., 2005). Until unambiguous glacial deposits are found, the early Late Ordovician should be referred to as a transition from hothouse to icehouse, rather than the onset of glaciation. CONCLUSIONS Although Laurentia was situated in tropical to subtropical latitudes during the Late Ordovician, lithological evidence indicates that part of this time period was characterized by deposition under cool-water conditions. Whereas Turinian strata of eastern North America contain abundant warm-water carbonates, the Lower Chatfieldian marks a shift to cool-water carbonates, an increase in siliciclastic influx, and phosphatic sediments. Numerical modeling results show that global cooling had a major impact on sea surface temperatures in the shallow epicontinental sea of Laurentia during the Late Ordovician. The presence of cold-water masses that originated in the higher latitudes and were brought toward the southern and eastern margins of Laurentia could have led to the observed stratigraphic changes. Our current global climate model does not include a turbidity function, and this alternative hypothesis cannot be directly tested with our model setup. Future modeling studies therefore need to extend the modeling approach by using high-resolution regional sediment-transport models that include turbidity functions. ACKNOWLEDGMENTS This paper was much improved thanks to the careful reviews by M. Saltzman and M. Pope, and to the editorial work of S. Finney. L. Wasylenki provided useful comments that greatly improved this manuscript. ADH expresses his gratitude for financial support from The Sol and Esther Drescher Faculty Development Grant. REFERENCES CITED Ackerly, S.C., Cisne, J.L., Railsback, L.B., and Anderson, T.L., 1993, Punctal density in the Ordovician orthide brachiopod Paucicrura rogata: Anatomical and paleoenvironmental variation: Lethaia, v. 26, p. 17–24, doi: 10.1111/j.1502-3931.1993.tb01506.x. Barnes, C.R., 2004, Ordovician oceans and climate, in Webby, B.D., Paris, F., Droser, M.L., and Percival, I.G., eds., The Great Ordovician Biodiversification Event: New York, Columbia University Press, p. 72–76. Beauchamp, B., and Desrochers, A., 1997, Permian warm- to very cold-water carbonates and cherts in northwest Pangea, in James, N.P., and Clarke, J.A.D., eds., Cool-Water Carbonates, v. 56: Tulsa, SEPM (Society for Sedimentary Geology), p. 327–347. Bergström, S.M., Finney, S.C., Xu, C., Goldman, D., and Leslie, S.A., 2006, Three new Ordovician global stage names: Lethaia, v. 39, p. 287–288, doi: 10.1080/00241160600847439. Berner, R.A., 1994, GEOCARB II: A revised model of atmospheric CO2 over Phanerozoic time: American Journal of Science, v. 294, p. 56–91. Berner, R.A., and Kothavala, Z., 2001, Geocarb III: A revised model of atmospheric CO2 over Phanerozoic time: American Journal of Science, v. 301, p. 182–204, doi: 10.2475/ajs.301.2.182.
Brenchley, P.J., Carden, G.A.F., and Marshall, J.D., 1995, Environmental changes associated with the ‘first strike’ of the Late Ordovician mass extinction: Modern Geology, v. 22, p. 69–82. Brenchley, P.J., Carden, G.A., Hints, L., Kaljo, D., Marshall, J.D., Martma, T., Meidla, T., and Nõlvak, J., 2003, High-resolution isotope stratigraphy of Late Ordovician sequences: Constraints on the timing of bio-events and environmental changes associated with mass extinction and glaciation: Geological Society of America Bulletin, v. 115, p. 89–104, doi: 10.1130/ 0016-7606(2003)115<0089:HRSISO>2.0.CO;2. Brookfield, M.E., and Brett, C.E., 1988, Paleoenvironments of the MidOrdovician (Upper Caradocian) Trenton limestones of southern Ontario, Canada: Storm sedimentation on a shoal-basin shelf model: Sedimentary Geology, v. 57, p. 75–105, doi: 10.1016/0037-0738(88)90019-X. Brookfield, M.E., and Elgadi, M., 1998, Sedimentology and paleocommunities of the Black River and Trenton Limestone groups (Ordovician), Lake Simcoe area, Ontario: Geological Society of America Annual Meeting, Field Trip Guidebook 6, 35 p. Cisne, J.L., and Rabe, B.D., 1978, Coenocorrelation: Gradient analysis of fossil communities and its applications in stratigraphy: Lethaia, v. 11, p. 341– 364, doi: 10.1111/j.1502-3931.1978.tb01893.x. Cisne, J.L., Karig, D.E., Rabe, B.D., and Hay, B.J., 1982, Topography and tectonics of the Taconic outer trench slope as revealed through gradient analysis of fossil assemblages: Lethaia, v. 15, p. 229–246, doi: 10.1111/ j.1502-3931.1982.tb00647.x. Crowley, T.J., and Baum, S.K., 1995, Reconciling Late Ordovician (440 Ma) glaciation with very high CO2 levels: Journal of Geophysical Research, v. 100, p. 1093–1101, doi: 10.1029/94JD02521. Eckert, J.D., 1988, Late Ordovician extinction of North American and British crinoids: Lethaia, v. 21, p. 147–167, doi: 10.1111/j.1502-3931.1988 .tb00805.x. Finney, S.C., 2005, Global series and stages for the Ordovician System: A progress report: Geologica Acta, v. 3, p. 309–316. Frey, R.C., 1995, Middle and Upper Ordovician Nautiloid Cephalopods of the Cincinnati Arch Region of Kentucky, Indiana, and Ohio, USA: U.S. Geological Survey Professional Paper 1066, 126 p. Gent, P.R., and McWilliams, J.C., 1990, Isopycnal mixing in ocean circulation models: Journal of Physical Oceanography, v. 20, p. 150–155, doi: 10.1175/ 1520-0485(1990)020<0150:IMIOCM>2.0.CO;2. Hamoumi, N., 1999, Upper Ordovician glaciation spreading and its sedimentary record in Moroccan North Gondwana margin, in Kraft, P., and Fatka, O., eds., Quo Vadis Ordovician?: Short papers of the 8th International Symposium on the Ordovician System: Acta Universitatis Carolinae, Geologica, v. 43, p. 111–114. Haupt, B.J., and Seidov, D., 2001, Warm deep-water ocean conveyor during Cretaceous time: Geology, v. 29, p. 295–298, doi: 10.1130/0091-7613 (2001)029<0295:WDWOCD>2.0.CO;2. Herrmann, A.D., Patzkowsky, M.E., and Pollard, D., 2003, Obliquity forcing with 8–12 times pre-industrial levels of atmospheric pCO2 during the Late Ordovician glaciation: Geology, v. 31, p. 485–488, doi: 10.1130/0091 -7613(2003)031<0485:OFWTPL>2.0.CO;2. Herrmann, A.D., Haupt, B.J., Patzkowsky, M.E., Seidov, D., and Slingerland, R.L., 2004a, Response of Late Ordovician paleoceanography to changes in sea level, continental drift, and atmospheric pCO2: Potential causes for long-term cooling and glaciation: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 210, p. 385–410, doi: 10.1016/j.palaeo.2004.02.034. Herrmann, A.D., Patzkowsky, M.E., and Pollard, D., 2004b, The impact of paleogeography, pCO2, poleward ocean heat transport and sea level change on global cooling during the Late Ordovician: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 206, p. 59–74, doi: 10.1016/ j.palaeo.2003.12.019. Holland, S.M., and Patzkowsky, M.E., 1996, Sequence stratigraphy and longterm paleoceanographic change in the Middle and Upper Ordovician of the Eastern United States, in Witzke, B.J., Ludvigson, G.A., and Day, J., eds., Paleozoic Sequence Stratigraphy: Views from the North American Craton: Geological Society of America Special Paper 306, p. 117–129. James, N.P., 1997, The cool-water carbonate depositional realm, in James, N.P., and Clarke, J.A.D., eds., Cool-Water Carbonates: SEPM (Society for Sedimentary Geology) Special Publication 56, p. 1–20. Johnson, R.G., 1997, Climate control requires a dam at the Strait of Gibraltar: Eos (Transactions, American Geophysical Union), v. 78, p. 280–281. Kolata, D.R., Huff, W.D., and Bergström, S.M., 2001, The Ordovician Sebree Trough: An oceanic passage to the Midcontinent United States: Geological
Stratigraphic change in subtropical to tropical Laurentia Society of America Bulletin, v. 113, p. 1067–1078, doi: 10.1130/0016 -7606(2001)113<1067:TOSTAO>2.0.CO;2. Lavoie, D., and Asselin, E., 1998, Upper Ordovician facies in the Lac SaintJean outlier, Québec (eastern Canada): Palaeoenvironmental significance for Late Ordovician oceanography: Sedimentology, v. 45, p. 817–832, doi: 10.1046/j.1365-3091.1998.00170.x. Lees, A., 1975, Possible influence of salinity and temperature on modern shelf carbonate sedimentation: Marine Geology, v. 19, p. 159–198, doi: 10.1016/ 0025-3227(75)90067-5. Lees, A., and Buller, A.T., 1972, Modern temperate-water and warm-water shelf carbonate sediments contrasted: Marine Geology, no. 13, p. M67–M73, doi: 10.1016/0025-3227(72)90011-4. Pacanowski, R.C., 1996, MOM 2. Documentation, User’s Guide and Reference Manual, GFDL Ocean Technical Report #3.2: Geophysical Fluid Dynamics Laboratory/NOAA, 329 p. Patzkowsky, M.E., and Holland, S.M., 1993, Biotic response to a Middle Ordovician paleoceanographic event in eastern North America: Geology, v. 21, p. 619–622, doi: 10.1130/0091-7613(1993)021<0619:BRTAMO> 2.3.CO;2. Patzkowsky, M.E., and Holland, S.M., 1996, Extinction, invasion and sequence stratigraphy: Patterns of faunal change in the Middle and Upper Ordovician of the eastern United States, in Witzke, B.J., Ludvigson, G.A., and Day, J., eds., Paleozoic Sequence Stratigraphy: Views from the North American Craton: Geological Society of America Special Paper 306, p. 131–142. Patzkowsky, M.E., and Holland, S.M., 1997, Patterns of turnover in Middle and Upper Ordovician brachiopods of the eastern United States: A test of coordinated stasis: Paleobiology, v. 23, p. 420–443. Patzkowsky, M.E., and Holland, S.M., 1999, Biofacies replacement in a sequence stratigraphic framework: Middle and Upper Ordovician of the Nashville Dome, Tennessee, USA: Palaios, v. 14, p. 301–323, doi: 10.2307/ 3515459. Patzkowsky, M.E., Slupik, L.M., Arthur, M.A., Pancost, R.D., and Freeman, K.H., 1997, Late Middle Ordovician environmental change and extinction: Harbinger of the end-Ordovician or continuation of Cambrian patterns?: Geology, v. 25, p. 911–914, doi: 10.1130/0091-7613(1997)025 <0911:LMOECA>2.3.CO;2. Pope, M.C., and Read, J.F., 1997, High-resolution stratigraphy of the Lexington Limestone (late Middle Ordovician), Kentucky, U.S.A.: A cool-water carbonate-clastic ramp in a tectonically active foreland basin, in Noel, P., and Clarke, J.A.D., eds., Cool-Water Carbonates: SEPM (Society for Sedimentary Geology) Special Publication 56, p. 410–429. Pope, M.C., and Steffen, J.B., 2003, Widespread, prolonged late Middle to Late Ordovician upwelling in North America: A proxy record of glacia-
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tion?: Geology, v. 31, p. 63–66, doi: 10.1130/0091-7613(2003)031<0063: WPLMTL>2.0.CO;2. Pratt, L.J., Johns, W., Murray, S.P., and Katsumata, K., 1999, Hydraulic interpretation of direct measurements in the Bab el Mandab Strait: Journal of Physical Oceanography, v. 29, p. 2769–2784, doi: 10.1175/1520-0485 (1999)029<2769:HIODVM>2.0.CO;2. Railsback, L.B., Ackerly, S.C., Anderson, T.F., and Cisne, J.L., 1990, Palaeontological and isotope evidence for warm saline deep waters in Ordovician oceans: Nature, v. 343, p. 156–159, doi: 10.1038/343156a0. Saltzman, M.R., and Young, S.A., 2005, Long-lived glaciation in the Late Ordovician? Isotopic and sequence-stratigraphic evidence from western Laurentia: Geology, v. 33, p. 109–112, doi: 10.1130/G21219.1. Samankassou, E., 2002, Cool-water carbonates in a paleoequatorial shallowwater environment: The paradox of the Auernig cyclic sediments (Upper Pennsylvanian, Carnic Alps, Austria-Italy) and its implications: Geology, v. 30, p. 655–658, doi: 10.1130/0091-7613(2002)030<0655:CWCIAP> 2.0.CO;2. Saylor, B.Z., Byers, C.W., Choi, Y.S., and Simo, J.A., 1997, Paleoceanography and paleoclimate of the Middle and Upper Ordovician Sinnipee Group: Geological Society of America Abstracts with Programs, v. 29, no. 4, p. 70. Scotese, C.R., and McKerrow, W.S., 1990, Revised world maps and introduction, in McKerrow, W.S., and Scotese, C.R., eds., Palaeozoic palaeogeography and biogeography: Oxford, UK, Geological Society [London] Memoir 12, p. 1–21. Scotese, C.R., and McKerrow, W.S., 1991, Ordovician plate tectonic reconstructions, in Barnes, C.R., and Williams, S.H., eds., Advances in Ordovician Geology: Geological Survey of Canada Paper 90-9, p. 225–234. Sepkoski, J.J., Jr., 1996, Patterns of Phanerozoic extinction: A perspective from global data bases, in Walliser, O.H., ed., Global Events and Event Stratigraphy in the Phanerozoic: Heidelberg, Springer-Verlag, p. 35–51. Tobin, K.J., Bergström, S.M., and De La Garza, P., 2005, A mid-Caradocian (453 Ma) drawdown in atmospheric pCO2 without ice sheet development?: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 226, no. 3–4, p. 187–204. Webby, B.D., Cooper, R.A., Bergström, S.M., and Paris, F., 2004, Stratigraphic framework and time slices, in Webby, B.D., Paris, F., Droser, M.L., and Percival, I.G., eds., The Great Ordovician Biodiversification Event: New York, Columbia University Press, p. 41–47. Yapp, C.H., and Poths, H., 1992, Ancient atmospheric CO2 inferred from natural goethites: Nature, v. 355, p. 342–344, doi: 10.1038/355342a0. MANUSCRIPT ACCEPTED BY THE SOCIETY 6 OCTOBER 2009
Printed in the USA
The Geological Society of America Special Paper 466 2010
The Upper Ordovician Guttenberg δ 13C excursion (GICE) in North America and Baltoscandia: Occurrence, chronostratigraphic significance, and paleoenvironmental relationships Stig M. Bergström† School of Earth Sciences, The Ohio State University, Columbus, Ohio 43210, USA Birger Schmitz Department of Geology, GeoBiosphere Science Centre, Lund University, Sölvegatan 12, SE-223 62 Lund, Sweden Matthew R. Saltzman School of Earth Sciences, The Ohio State University, Columbus, Ohio 43210, USA Warren D. Huff Department of Geology, University of Cincinnati, Cincinnati, Ohio 45221, USA
ABSTRACT Two prominent, and apparently globally distributed, δ13C excursions have been documented from the Upper Ordovician, namely the early Katian Guttenberg isotope carbon excursion (GICE) and the latest Ordovician Hirnantian isotope carbon excursion (HICE). The former excursion, which has lower δ13C values than the HICE, is now recorded from dozens of localities in North America and Baltoscandia, and it appears to be present also in China. In North America the GICE ranges from the uppermost Phragmodus undatus Midcontinent Conodont Zone to near the top of the Plectodina tenuis Midcontinent Conodont Zone, an interval corresponding to the lower part of the Diplacanthograptus caudatus Global Graptolite Zone. The base of the GICE lies somewhat above the Millbrig K-bentonite. In Baltoscandia the GICE occurs in the upper Diplograptus foliaceus through the lower Dicranograptus clingani Graptolite Zones, and in the upper Amorphognathus tvaerensis Conodont Zone. Its base is a few meters above the widespread Kinnekulle K-bentonite. In Baltoscandia and in Oklahoma the GICE ranges through a part of the Spinachitina cervicornis Chitinozoan Zone. In North America the GICE is regionally in a transgressive-regressive succession. The bathymetric conditions in the GICE interval in Baltoscandia were somewhat complex and have been the subject of different interpretations, but there is no obvious correlation between the GICE and apparent sea level changes. A review of the relations between the GICE and potential climatic and water temperature
†
E-mail:
[email protected].
Bergström, S.M., Schmitz, B., Saltzman, M.R., and Huff, W.D., 2010, The Upper Ordovician Guttenberg δ13C excursion (GICE) in North America and Baltoscandia: Occurrence, chronostratigraphic significance, and paleoenvironmental relationships, in Finney, S.C., and Berry, W.B.N., eds., The Ordovician Earth System: Geological Society of America Special Paper 466, p. 37–67, doi: 10.1130/2010.2466(04). For permission to copy, contact
[email protected]. ©2010 The Geological Society of America. All rights reserved.
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Bergström et al. indicators, such as lithofacies, faunas, and 18O geochemistry, does not suggest a close correlation to specific environmental conditions. The cause of formation of the GICE is enigmatic, but there is no direct evidence that it was coeval with a period of extensive glaciation in the Gondwana. The GICE is a powerful chemostratigraphic tool that is useful for detailed local and even transatlantic correlations.
INTRODUCTION Conspicuous, geographically widespread variations of δ13C values in the stratigraphic record have in recent years proved to have great utility as tools in chemostratigraphic research. Such changes, known as excursions, frequently provide new insights into stratigraphic relationships at both the local and regional scales. In the case of the Ordovician, two major, apparently global, positive δ13C excursions are currently recognized, namely the early Late Ordovician Guttenberg isotope carbon excursion (GICE) and the latest Late Ordovician Hirnantian isotope carbon excursion (HICE). Although additional Katian excursions have been recognized in both North America (Ludvigson et al., 2004) and Europe (Kaljo et al., 2004), these are mostly of smaller magnitude and are as yet not documented outside these regions. The HICE, which is a prominent positive excursion, has been studied more intensely than the GICE and is documented from North America, the British Isles, China, South America, and Baltoscandia (Bergström et al., 2006). It has been widely employed in Hirnantian chronostratigraphy. The GICE, which is a somewhat less conspicuous, but still prominent, excursion in the Ordovician δ13C curve, has so far been recorded from a substantial number of sites in North America and northern Europe, and it also occurs in China (Bergström et al., 2009). The literature on the HICE has expanded rapidly during the last decade and is considerably more extensive than that dealing with the GICE. A recent paper (Bergström et al., 2006) reviewed several aspects of the HICE, especially its global chemostratigraphic significance and its putative relations to the Gondwana glaciation. The purpose of the present paper is to summarize pertinent data on the GICE from North America and Europe, including new information from sections in Sweden, and assess its local and transatlantic chemostratigraphic significance. This review, the first to deal in some detail with both North American and European records of the GICE, also includes brief discussions of its relations to depositional environments (especially bathymetric conditions) and possible glacial periods as well as proposed hypotheses to explain its formation. DISCOVERY AND CONCEPT OF THE GICE The discovery of the GICE dates back to the mid-1980s, when Hatch et al. (1987) found elevated δ13C values in oils and rock samples from nine drill cores from Iowa and Kansas. Samples from the Guttenberg Member of the Decorah Formation (Fig. 1) showed heavy δ13C values of up to +4‰ from baseline values
ranging from –3‰ to +0.5‰. This isotope excursion was in an interval a few meters thick well above the Millbrig K-bentonite in a highly condensed section. Ludvigson et al. (1996) presented a large amount of information in a comprehensive sedimentological, chemical, and petrographic study. Later, Ludvigson et al. (2000a) published additional important δ13C data based on rock samples from three outcrops through the same stratigraphic interval in Iowa, Wisconsin, and Missouri. Further information was provided by Ludvigson et al. (1990, 2000b, 2001a, 2001b) and Smith et al. (2000). Finally, Ludvigson et al. (2004) published a comprehensive chemostratigraphic and lithologic investigation of the excursion interval in which they formally labeled the positive excursion just above the Elkport K-bentonite within the Guttenberg Member of the Chatfieldian Decorah Formation as the Guttenberg excursion. This concept of the Guttenberg excursion (GICE) has been followed herein and is the same as that used by several other authors such as Saltzman et al. (2001, 2003a, 2003b), Bergström et al. (2001, 2004), Young et al. (2003a, 2003b, 2004a, 2004b, 2005) and Barta et al. (2003, 2004). It is important to recognize this concept of the GICE because Ludvigson et al. (2004) distinguished and named four older, apparently not global and hence regionally less significant, excursions in the succession just below the GICE in eastern Iowa. These should not be confused with the GICE, and they are apparently more local in their distribution than the GICE, as will be discussed below. Likewise, Kaljo et al. (2004) and Bergström et al. (2007) recently named several excursions above the GICE in the Estonian and North American successions, respectively. GICE IN NORTH AMERICA Since the initial discovery of the GICE, this excursion has been recorded at many localities in more than a dozen states and provinces in North America (Fig. 2). It has been particularly intensely studied in the Upper Mississippi Valley, but its currently known distribution area includes a large part of the North American Midcontinent as well as the Appalachians and the Great Basin in Nevada. In all probability, future work will show that it is present in all parts of the continent where rocks of suitable age and facies are present. Format restrictions of the present paper do not permit discussion of all North American GICE records, and we have therefore selected for review only particularly important occurrences, generally 1–3 for each outcrop region. Each of these records is discussed separately below and illustrated by the pertinent δ13C curve. For a general classification of the North American stratigraphic units discussed, see Figure 3. It should be noted
Upper Ordovician Guttenberg δ13C excursion (GICE) in North America and Baltoscandia
Figure 1. GICE curve in the Guttenberg Member of the Decorah Formation in the Cominco Millbrook Farms SS-9 drill core from eastern Iowa, one of the sections in which the GICE was first recorded (Hatch et al., 1987). The δ13C curve is based on Pancost et al. (1999), as interpreted by Ludvigson et al. (2004). Note the GICE peak of ~+1.5‰ in the Guttenberg Member as compared with δ13C baseline values of ~ –1‰ to 0‰ in the Platteville and Dunleith Formations. Also note that the Deicke and Millbrig K-bentonites are not present in this core, apparently because this interval is not preserved at the PlattevilleDecorah unconformity of Ludvigson et al. (2004, Fig. 5).
that for series classification we use the recently approved global standard series in which the provincial North American Mohawkian Series is referred to the global Upper Ordovician Series. Upper Mississippi Valley To date, the GICE has been most intensely studied in the Upper Mississippi Valley, a region that includes the important Upper Ordovician outcrop areas in Iowa, Wisconsin, Illinois, Minnesota, and Missouri. More than a dozen GICE sites have been recorded, and major references include Hatch et al. (1987), Hasiuk et al. (2001), Ludvigson et al. (1996, 2000a, 2000b, 2001a,
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2001b, 2004), Smith et al. (2000), Patzkowsky et al. (1997), Pancost et al. (1998, 1999), Simo et al. (2003), Saltzman et al. (2003a, 2003b), and Witzke et al. (2000). A typical GICE curve from this region is illustrated in Figure 4. As recently described in detail by Ludvigson et al. (2004), the GICE occurs in a stratigraphic interval characterized by much condensation and the presence of significant stratigraphic gaps that locally may include the excursion interval. When present, the GICE interval is generally <5 m thick and hence is considerable thinner than in the successions in Kentucky, Pennsylvania, Virginia, and West Virginia (see below). Stratigraphically, this excursion is within the Guttenberg sequence of Ludvigson et al. (2004) and well above the level of the widespread Millbrig K-bentonite (Kolata et al., 1996). In several sections the δ13C values increase near the level of the Elkport K-bentonite (Ludvigson et al., 2000a), and the curve returns to baseline values near the level of the Dickeyville K-bentonite. Typically, the excursion is within the Guttenberg Member of the Decorah Shale, which, according to Sweet (1987), belongs to the lower part of his P. tenuis Conodont Zone. δ13C baseline values are ~ –1‰ to –2‰, and the GICE values range up to ~+3‰ but are in most cases between +1‰ and +2‰; hence the magnitude of the excursion is ~3‰. The δ13C curve returns to baseline values of about –1‰ to 0‰ in the uppermost part of the Guttenberg Member, and there are no obvious excursions in the overlying Ion Member of the Decorah Formation and the Beecher and Eagle Point Members of the Dunleith Formation. The Decorah Formation clearly represents a deepening episode in the environment of deposition of the Upper Mississippi Valley succession, and according to Ludvigson et al. (1996) and Witzke and Bunker (1996) the Guttenberg Member represents the peak of the transgression. Estimates of water depth during deposition of this member are somewhat controversial, but according to Ludvigson et al. (1996) the water depth is likely to have been several tens of meters. Hence, it appears that the excursion occurs during a marked highstand episode. Kentucky The particularly important outcrop area of Upper Ordovician rocks in north-central Kentucky and adjacent parts of Indiana and Ohio (Fig. 2) includes the reference area of the Cincinnati Series, the Upper Ordovician in the standard North American chronostratigraphic classification (Ross et al., 1982). The outcrops in Indiana and Ohio do not expose strata as old as the GICE interval, but that part of the Upper Ordovician succession is very well exhibited at many sections in Kentucky. Drill cores penetrating the GICE interval are also available, such as the stratigraphically important Minerva core from a site near Maysville in northern Kentucky (Sweet et al., 1974). Thus far, the GICE has been studied in sections near Frankfort, Franklin County (Saltzman et al., 2003a; Bergström et al., 2004), at Boonesborough, Madison County (Young et al., 2005), and in the Minerva drill core, Mason County (Saltzman et al., 2003a). Data from these sections were
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Figure 2. Sketch map of parts of the USA and adjacent Canada, showing Upper Ordovician outcrop areas (gray) and geographic location of study sections. 1—SW Wisconsin; 2—eastern Iowa; 3—NE Missouri; 4—north-central Kentucky; 5—SW Virginia; 6—eastern West Virginia; 7—central Tennessee; 8—Arbuckle Mountains, Oklahoma; 9—central Pennsylvania; 10—northern New York State; 11—Manitoulin Island, Ontario; 12—Minerva drill core; 13 (inset map)—central Nevada. No GICE has yet been recorded from Alabama, Georgia, or Arkansas, but Hatch et al. (1987) reported elevated isotope values from Kansas drill cores, indicating the presence of the GICE. Early Chatfieldian rocks are missing in the Ordovician outcrop areas in Texas.
summarized in a composite δ13C curve for Kentucky in Young et al. (2005). The conodont biostratigraphy of these sections is well established on the basis of conodont studies (Bergström and Sweet, 1966; Sweet et al., 1974; Sweet, 1979; Richardson and Bergström, 2003). Detailed sequence-stratigraphic and lithostratigraphic investigations were carried out by Holland and Patzkowsky (1996), Pope and Read (1997, 1998), and Brett et al. (2004). K-bentonites and their event-stratigraphic significance have been studied by Conkin (1991), Conkin and Conkin (1992), and Kolata et al. (1996, 1998). The lower part of our δ13C curve (Fig. 5) is based on samples from the Turinian Tyrone Limestone in the Dead Horse
Quarry near Highway 627 in the northeastern part of Frankfort, Franklin County (Conkin and Conkin, 1992). As shown in Figure 5, the δ13C values range between 0‰ and about +1‰ through the upper Tyrone Limestone, and these are interpreted to represent local baseline values for this interval. The values from more than 20 samples from the Curdsville Member of the Lexington Limestone are based on samples collected from outcrops near and along Highway 627 in the northeastern part of Frankfort (Conkin and Conkin, 1992). The δ13C values through this part of the Kentucky succession differ little from each other and average about +1‰ and do not differ substantially from the Turinian baseline values. The sudden increase to +2.5‰ a little
F O R M A T I O N
CHATFIELDIAN TURINIAN
MOHAWKIAN
KATIAN SANDBIAN
UPPER ORDOVICIAN
GLOBAL N.AM. CONODONT ZONE SER. ST. SER. ST. MIDCONT. ATL.
GRAPTOLITE ZONE U. MISS VALL. TENNESSEE OKLAHOMA KENTUCKY
B. confluens A. superbus Plectodina tenuis
Diplacanthograptus caudatus
Phragmodus Amorphognathus undatus tvaerensis Belodina compressa
Decorah
Hermitage
Viola Springs
Lexington
VIRGINIA
W. VIRGINIA
PENNSYLV.
Reedsville
Reedsville
Antes
Trenton
Dolly Ridge
Coburn Salona
NEVADA
Kings Falls
Eureka
Napanee Selby
M D
Climacograptus bicornis
NEW YORK
Platteville
Carters
Bromide
Tyrone
Eggleston
Nealmont
Nealmont
Copenhagen
Watertown
Figure 3. Stratigraphic chart showing relations between formations studied, global and North American stages, and conodont and graptolite zones. Note the position of the Millbrig (M) and Deicke (D) K-bentonites in the uppermost Climacograptus bicornis Zone.
McGREGOR QUARRY, CLAYTON CO. IOWA
ION MBR. GUTTENBERG MBR.
10
PLATTEVILLE SPECHTS FM. FERRY MBR.
? Dickeyville K-bentonite
5
GICE
DECORAH FORMATION
m 15
Elkport K-bentonite
Millbrig K-bentonite Deicke K-bentonite
0 -4
-3
-2
-1
0
1
13
C
carb.
Figure 4. GICE curve in the Guttenberg Member of the Decorah Formation at the McGregor Quarry, Clayton County, eastern Iowa (after Ludvigson et al., 2000a). Note the 2-peak appearance of the δ13C curve, the peak values of ~+1‰, and the baseline values of ~ –1‰. The GICE is within the Guttenberg Member, which represents a highstand interval.
Figure 5. GICE curve in a composite section through the Lexington Limestone, based on outcrops near and along Highway 627 at Frankfort and the Minerva drill core in northern Kentucky. Note δ13C curve baseline values of 0‰ to +1‰, and the 2-peak appearance of the δ13C curve. Also note that the excursion is within the Logana Member of the Lexington Limestone, which represents a highstand interval.
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Bergström et al.
above the base of the Logana Member is quite conspicuous and represents the beginning of the GICE. The post-Curdsville part of the δ13C curve is based on samples from the Minerva drill core. Most of the δ13C values fluctuate between +1‰ and +2‰ through the Logana Member. Slightly higher, near the top of the Logana Member, the δ13C values drop to between 0‰ and +1‰ and are considered baseline values, indicating the end of the GICE. The total thickness of the GICE is 10–15 m in Kentucky, hence more than twice as much as in the Upper Mississippi Valley. Interestingly, in both the Frankfort section and the Minerva drill core, the GICE has a two-peak appearance, as is the case for several other sections in both North America and Baltoscandia. In terms of conodont biostratigraphy, the basal Lexington Limestone clearly represents the Phragmodus undatus Zone, but the base of the superjacent Plectodina tenuis Zone is currently poorly controlled in the Lexington Limestone. In the Minerva drill core the index of the zone appears ~15 m above the base of the Lexington Limestone, which is near the middle of the GICE interval. However, indeterminate Plectodina fragments occur in slightly older strata, and it is quite possible that most of, or the entire, GICE is present in the latter zone, as is the case in other sections. In terms of sequence stratigraphy, the GICE occurs in the highstand interval in the lower half of the M5 sequence of Holland and Patzkowsky (1996) in a sequence referred to as M5A by Brett et al. (2004) and McLaughlin and Brett (2004). On the basis of chemostratigraphy, the Logana Member correlates closely with the Guttenberg Member of the Decorah Formation in the Upper Mississippi Valley, which also represents a marked highstand in sea level. Virginia The only published GICE section in Virginia is the long railroad cut at Hagan, Lee County, which is the reference profile of Cooper’s (1956) North American Wildernessian Stage (Bergström et al., 1988). It is one of the most important sections through this stratigraphic interval in the southern and central Appalachians and has been dealt with in numerous publications (see, e.g., Miller and Fuller, 1954; Miller and Brosgé, 1954; Bergström et al., 1988; Haynes, 1992, 1994; Kolata et al., 1996, 1998, 2001; Leslie and Bergström, 1997; Mitchell et al., 2004). A δ13C curve through the upper Eggleston Formation and Trenton Limestone was recently published by Young et al. (2005), and it is shown, slightly modified, in Figure 6. Apart from an ~70-m-thick interval in the upper part of the Trenton Limestone, the study interval is excellently exposed and includes the Deicke and Millbrig K-bentonites. The Hagan δ13C curve (Fig. 6) shows more fluctuations than the curves from the Upper Mississippi Valley and Kentucky but compares well in major features with the curves from these areas, although the GICE is much thicker (~120 m). The δ13C curve shows baseline values between 0‰ and +1‰ in the lower-
most 30 m of the Trenton Limestone before a conspicuous rise to >+2‰. In the next 50 m there are numerous fluctuations in δ13C values between +1‰ and near +3‰, with average values of ~+2‰. The curve, based on samples from the uppermost Trenton Limestone, above the covered interval, shows a notable decline down to values of <–1‰, which is interpreted as the end of the GICE. Although several workers (see, e.g., Bergström et al., 1988; Leslie, 1995) have investigated the conodont biostratigraphy of the Hagan section, the precise position of conodont zone boundaries remains incompletely known. However, the relations between the GICE, the Millbrig K-bentonite, and conodont zones are closely similar to those in other sections in the eastern Midcontinent. As in these sections, the M5 sequence boundary is above the Millbrig K-bentonite, and evidently the GICE is in the transgressive sequence of the Trenton Limestone. West Virginia The only current record of the GICE in West Virginia is from a relatively little known section at Dolly Ridge, Pendleton County, which was recently described by Young et al. (2005). The outcrop is along a dirt road along the slope of Dolly Ridge ~1.3 km southwest of Riverton. An ~180-m-thick succession of the Nealmont and Dolly Ridge Formations, first described by Perry (1972), consists of lime mudstones, wackestones, shales, and some beds of skeletal packstones. As many as 14 K-bentonites have been recorded from this section, including two beds in the upper Nealmont Formation, tentatively identified as the Deicke and Millbrig K-bentonites (Young et al., 2005). For further details about this section, see Perry (1972). The Dolly Ridge succession provides an excellent, and apparently complete, GICE curve (Fig. 6) in a succession representing a relatively high sedimentation rate and without recognized stratigraphic gaps. Along with the Hagan section, it is currently the stratigraphically thickest (~130 m) representation of the GICE known not only in North America but globally. The δ13C curve is based on 30 samples, but despite a rather thick average sample interval of 6 m, we expect the excursion curve to be fully representative of the GICE in this succession. As described by Young et al. (2005), the top 40 m of the Nealmont Formation, which contains the Deicke and Millbrig K-bentonites and is referable to the Phragmodus undatus Zone, has baseline δ13C values of between +1.2‰ and +0.5‰. From about the base of the Dolly Ridge Formation, there is a gradual but rapid rise in the δ13C values to about +3‰ at ~100 m above the base of the sampled interval. Values between +2.5‰ and +3‰ are present through a thickness of nearly 50 m until the curve drops conspicuously to below 0‰ in the topmost part of the Dolly Ridge Formation, just below the overlying Reedsville Formation. This indicates the end of the GICE. Although the biostratigraphic control of this section is not yet as precise as in some other North American successions, the current, largely unpublished, information at hand is in good agreement with
Upper Ordovician Guttenberg δ13C excursion (GICE) in North America and Baltoscandia
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Figure 6. GICE curve in the Trenton Limestone at Hagan, Virginia, and the Dolly Ridge Formation at Dolly Ridge,West Virginia (after Young et al., 2005). Despite the considerable stratigraphic range of the excursion, it has the same relations to the Midcontinent conodont zones at these sections as in other successions. Note the δ13C baseline values of ~+1‰ and the peak values of ~+3‰.
the biostratigraphic range of the GICE in other successions. As shown in Figure 6, the Dolly Ridge curve shows close resemblance to the less complete Hagan curve in general shape, thickness, and relations to the Millbrig K-bentonite, and the base of the Reedsville Formation. However, it is a great deal smoother than the fluctuating Hagan curve, but the reason for this requires further study. Tennessee A classical Ordovician outcrop area in Tennessee is around the City of Nashville in the central part of the state, where there are numerous exposures of Upper Ordovician strata (Wilson, 1949; Holland and Patzkowsky, 1997, 1998). The late TurinianChatfieldian succession consists of carbonates and calcareous shales referred to the Carters and Hermitage Formations. The Carters Formation includes shallow-water lime mudstones and packstones that locally contain bird’s-eye structures. The Hermitage Formation apparently was deposited in deeper water than the Carters Formation and consists mainly of wackestones and calcareous shales, mudstones, and siltstones. Lithologically, it is similar to the Logana Member of the Lexington Limestone in Kentucky. The Deicke K-bentonite is present in the upper Carters Formation, and the Millbrig K-bentonite, where present, occurs
at the top of the same formation (Haynes, 1994; Kolata et al., 1996). The upper Carters Formation contains conodonts of the Phragmodus undatus Zone, and Plectodina tenuis, the index of the P. tenuis Zone, appears near the base of the overlying Hermitage Formation (Leslie, 1995). The Carters-Hermitage formation contact is an unconformity that has been taken as the M4-M5 sequence boundary (Holland and Patzkowsky, 1998). The δ13C studies by Railsback et al. (2003) and Young (2003) showed the presence of the GICE in the Hermitage Formation. Based on three sections, Young (2003) presented a composite GICE curve (Fig. 7). The baseline δ13C values in the upper Carters Formation are between 0‰ and +1‰. At the base of the Hermitage Formation the values rise conspicuously from ~+1‰ to ~+3‰. The abruptness of this increase supports the idea of a stratigraphic gap at the top of the Carters Formation. The high δ13C values are present in only a relatively thin interval (~3 m), and the upper part of the Hermitage Formation is characterized by values in the +1‰ to +2‰ range. Owing to the shape of the δ13C curve, it is difficult to define the precise level of the end of the GICE. However, this horizon is likely to be in the upper part of the Hermitage Formation somewhat below the base of the Belodina confluens Zone. As in Kentucky and other regions of the Midcontinent, the GICE occurs in a part of the succession characterized by a flooding event.
Bergström et al.
TURINIAN Clim. bicornis Phragmodus undatus BROMIDE
HIGHWAY 99, FITTSTOWN, OKLAHOMA
GICE
C H AT F I E L D I A N
One of the economically most important outcrop areas of Ordovician rocks in North America is in and around the Arbuckle Mountains in southern Oklahoma. Because of their importance as hydrocarbon producers, Ordovician strata in this region have been the subject of intense study for more than a century, resulting in a voluminous literature. Of particular interest in the present study is the Viola Springs Formation, which includes the GICE interval (Young et al., 2005). For information about various aspects of this formation, see Alberstadt (1973), Amsden and Sweet (1983), Finney (1986), and Goldman and Bergström (1997). Gao et al. (1996) carried out a pioneer δ13C study of the Viola Springs Formation, but unfortunately they studied a stratigraphically incomplete section along Interstate 35 in which most, if not all, of the GICE interval is cut out by an unconformity. One of the stratigraphically most complete and best exposed sections of the Viola Springs Formation is along Hwy 99 ~3.3 miles south of Fittstown in Pontotoc County. This section, which
was described by Decker (1933), is a classical field trip stop. Its conodont biostratigraphy was decribed by Sweet (1983), and the graptolite biostratigraphy was dealt with by Finney (1986) and Goldman et al. (2005). A δ13C curve through the upper 10 m of the Bromide Formation and the lower 46 m of the overlying Viola Springs Formation was published by Young et al. (2005). The δ13C values from the upper 10 m of the Bromide Formation are ~0‰ (Fig. 8), whereas those from the lower 20 m of the Viola Springs Formation are between –1‰ and +1‰ with notable fluctuations. The values in the middle part of the study section reach nearly +1.5‰ before dropping to <0‰ in the upper 6 m of the sample interval. This drop, which represents the end of the GICE, is just below the base of the Belodina confluens Zone and below the first record of the important graptolite Diplacanthograptus spiniferus.
Plectodina tenuis VIOLA SPRINGS
Oklahoma
STAGE GRAPTOD. spiniferus LITE ZONE CONODONT Bel. confluens ZONE FM.
44
Millbrig K-bentonite?
Deicke -2 K-bentonite?
M5 M4
-1
0
1
2
13
C
carb.
Figure 7. Composite δ13C curve through the Hermitage Formation in central Tennessee, based on outcrops at Nashville Airport, and road cuts at Gladesville and Tanglewood (after Young et al., 2005). The M4-M5 sequence boundary is associated with a prominent gap that includes the uppermost Carters and lowermost Hermitage Formations. This results in an abrupt beginning of the GICE at the base of the Hermitage Formation, which is a highstand unit that is lithologically reminiscent of the Logana Member of the Lexington Limestone in Kentucky. Note that the δ13C curve has baseline values of ~+1‰ and peak values of ~+3‰.
Figure 8. GICE in the Viola Springs Formation at the Highway 99 road cut south of Fittstown, Oklahoma (after Young et al., 2005), where the sampled interval is 73 m thick. The excursion occurs in an ~30-mthick interval. The curve baseline values are ~0‰, and the peak values ~+1.5‰. Note that the lowermost Viola Springs Formation, below the GICE, has yielded chitinozoans of the Spinachitina cervicornis Zone (Goldman et al., 2005), which are also present below the GICE in Baltoscandia. For information about the recently discovered K-bentonites in the Bromide Formation in this section, see Leslie et al. (2006). As in the other figures, the lithologic patterns in the stratigraphic column follow those used in Bergström et al. (2006; Fig. 4).
One of the early Ordovician δ C investigations in North America was carried out by Patzkowsky et al. (1997) on the upper Nealmont, Salona, and Coburn Formations at a highway cut near Reedsville, which occupy a stratigraphic position similar to that of the Decorah Formation in the Upper Mississippi Valley on the basis of conodont biostratigraphy (Sweet, 1984) and the presence of the Deicke and Millbrig K-bentonites in the lower part of the Salona Formation (Kolata et al., 1996). Among the two sections studied by these workers, the most informative one is the Highway 144 cut just south of Reedsville, Mifflin County. Their investigated interval is ~150 m thick and ranges to near the top of the Coburn Formation. This unit is overlain by the Antes Shale, the lower part of which contains graptolites of the Corynoides americanus and Orthograptus ruedemanni Zones (Mitchell et al., 2004). The Nealmont, Salona, and Coburn formations belong to the Atlantic Amorphognathus tvaerensis Conodont Zone, and the index of the Midcontinent Belodina confluens Conodont Zone appears in the topmost part of the Coburn Formation (Richardson and Bergström, 2003). The baseline δ13C values in the Turinian Nealmont Formation are difficult to establish because values are currently available from only the topmost 18 m of the formation, but comparison with the Hagan and Dolly Ridge successions suggests that the baseline value may be ~1‰. Assuming this is correct, there is a rather distinct negative δ13C excursion in the topmost part of the Nealmont Formation, a few meters below the Deicke K-bentonite (Fig. 9). The curve from the lower 10 m of the Salona Formation shows a gradual increase in δ13C values up to ~+2‰. This is an unusual feature because in most studied sections in North America the stratigraphic interval just above the Millbrig K-bentonite has typically rather uniform, or even decreasing, δ13C values. If this is due to the fact that the Reedsville section is unusually complete stratigraphically—there appears to be no gap in the projected position of the M4-M5 sequence boundary—or due
FM.
CONODONT ZONE
B. confluens
m 150
140
120
110
L
D I A N COBURN
REEDSVILLE, MIFFLIN COUNTY, PENNSYLVANIA ANTES SHALE
130
GICE
Plectodina tenuis
I E TURINIAN
NEALMONT
13
90
80
70
60
50
Phragmodus undatus
Pennsylvania
45
100
C H A T F SALONA
This section is particularly important biostratigraphically in that it provides information about the relations between the GICE and conodont, graptolite, and chitinozoan zones. The interval beneath the GICE represents the Phragmodus undatus Conodont Zone (Sweet, 1984), the Climacograptus bicornis Graptolite Zone (Finney, 1986; Goldman et al., 2005), and the Spinachitina cervicornis Chitinozoan Zone (Goldman et al., 2005). The GICE falls largely, if not entirely, within the Plectodina tenuis Zone, which overlies the Phragmodus undatus Zone. It is also of interest to note that the M4-M5 sequence boundary of Holland and Patzkowsky (1996) is at the unconformity at the base of the Viola Springs Formation; a K-bentonite at that stratigraphic level was recently interpreted to be the Millbrig K-bentonite (Leslie et al., 2006). Hence, the relations between the GICE and various stratigraphic indicators in the Fittstown section agree well with the conditions at other Midcontinent sections.
STAGE
Upper Ordovician Guttenberg δ13C excursion (GICE) in North America and Baltoscandia
40
30
Millbrig K-bentonite Deicke K-bentonite
20
10
0
0
1
2
3
4
13
C
carb.
Figure 9. GICE in the Salona and Coburn Formations in the highway cut at Reedsville, Pennsylvania (after Patzkowsky et al., 1997). The δ13C curve baseline values appear to be ~+1‰, and the peak values ~+3‰. Note that the thickness of the GICE is of the same order (~100 m) as at Hagan and Dolly Ridge, and hence many times greater than in the Upper Mississippi Valley sections.
to some other, still unidentified reason, requires further study. The gradual increase in δ13C values continues upward in the succession to reach maximum values of somewhat more than +3‰ in the upper half of the Salona Formation and through much of the overlying Coburn Limestone. There is a gradual decline in values in the upper 15 m of the Coburn Limestone below the base of the Belodina confluens Zone, but a sudden decline to baseline values is not very obvious. However, it is important to note that there is severe tectonic deformation in the upper part of the Coburn Formation at Reedsville, and there may be even a significant gap between this formation and the overlying Antes Shale, as recently illustrated by Mitchell et al. (2004, Fig. 7). Hence, the uppermost part of the Reedsville curve may be considered problematic. At any rate, the thickness of the GICE in the Reedsville section (~100 m) appears to be comparable to that of the Dolly Ridge and Hagan sections and more than
New York State Another classical area of Ordovician rocks is around the Adirondack Mountains in northern New York State (Fig. 2). This is the type area of the Trenton Group (Kay, 1937) and of the Mohawkian Series, the upper Middle Ordovician in the standard North American Series classification. Studies since the mid-nineteenth century have resulted not only in numerous publications but also in a complex, and in some cases controversial, stratigraphic nomenclature (for recent discussions, see Goldman et al., 1994; Brett and Baird, 2002; Mitchell et al., 2004; Brett et al., 2004). Recent δ13C work in the Trenton Group (Barta, 2004; Barta et al., 2007) has shown the presence of a distinct excursion, identified as the GICE, which starts in the mid-Napanee Formation a few meters above a K-bentonite identified as the Millbrig K-bentonite (Mitchell et al., 2004; Fig. 3) and ends in the middle of the overlying Kings Fall Formation (Fig. 10). Full details of this excursion have been published by Barta et al. (2007). It should be noted that the typical two-peak GICE reaches maximum δ13C values of ~+3‰ in the upper Napanee Formation and lower Kings Fall Limestone (Barta, 2004). Slightly higher stratigraphically, in the middle Kings Fall Limestone, the isotope curve drops to baseline values of +1‰ to +2‰. Barta (2004) and Barta et al. (2007) also recorded the GICE above the Deicke K-bentonite in the Bobcaygeon Formation on Manitoulin Island, Ontario, and in the reference section of the Rocklandian Stage (or Substage) at Stewart Quarry at Rockland, Québec. The discovery of the GICE in the lower part of the Trenton Group has important implications for the regional correlation of the Rocklandian Stage (or Substage) in North America. The fact that the GICE starts about midway through the Rocklandian shows that this level correlates closely with the base of the Logana Member of the Lexington Limestone, the base of the Hermitage Formation in central Tennessee, and the base of the Guttenberg Member of the Decorah Formation in the Upper Mississippi Valley. Also, the chemostratigraphy suggests that the Selby Formation, the basal unit in the Trenton Group, correlates with the Curdsville Member of the Lexington Limestone in Kentucky. These correlations at about the base of the Rocklandian differ substantially from those recently advocated by Brett et al. (2004, Fig. 6), who considered the Curdsville Member a correlative of the Watertown Formation of the Black River Group in New York.
DEXTER QUARRY, NEW YORK
GICE
FM.
STAGE
C H AT F I E L D I A N ROCKLAND N A PA N E E
20 times as large as in the epicontinental platform successions in the Upper Mississippi Valley, illustrating the much higher rate of deposition in the Appalachian Foreland Basin compared with that in the continental interior. It should be noted also that in Pennsylvania the GICE occurs in a highstand succession. The Salona Formation was clearly deposited in deeper water than the underlying shallow-water Nealmont Formation, whereas the Coburn Limestone might represent a renewed shallowing of the depositional environment.
HULL KAY ‘37
Bergström et al.
TURINIAN WATERTOWN WATERTOWN SELBY
46
10
Not exposed
m
no data
0 Millbrig (Hounsfield) K-bentonite Pond level
0
1
2
3
13
C
carb.
Figure 10. GICE in the Napanee Limestone at the Dexter Quarry, Jefferson County, New York (after Barta, 2004). For a composite New York State δ13C curve, see Barta et al. (2007, Fig. 10). Base line δ13C values are ~+1‰ in the early Chatfieldian Selby Limestone, and excursion peak values approach 3‰ in the Napanee Limestone in the highway road cut. Although no δ13C values are available from the ~6-m-thick covered interval, it is likely that baseline values continue through that portion of the succession. Based on unpublished recent studies of other sections in this region (Barta, 2004), the excursion extends upward to the middle portion of the Kings Fall Limestone, which overlies the Napanee Limestone. The Dexter Quarry is the reference section of the well-known Hounsfield K-bentonite, which, based on recent geochemical fingerprinting (Mitchell et al., 2004), is identical with the Millbrig K-bentonite.
Nevada Until recently the GICE had not been recorded in the vast region west of the Rocky Mountains. A pioneer δ13C study on samples from the Chatfieldian interval in the Vinini Shale (Jacobson et al., 1995) failed to provide evidence of the presence of the GICE. One major reason for the lack of GICE records is undoubtedly the fact that in most regions this stratigraphic interval is missing or is represented only by clastic sediments, such as the prominent Eureka Quartzite, which are less useful for chemostratigraphic work. One of the very few, if not the only, carbonate formation that holds some promise for a δ13C study of the GICE interval is the Copenhagen Formation of central Nevada, which is considered a local carbonate facies of the lower part of the Eureka Quartzite (Harris et al., 1979). The suitability of the Copenhagen Formation for chemostratigraphic study is also enhanced by the fact that it has well-established conodont
Upper Ordovician Guttenberg δ13C excursion (GICE) in North America and Baltoscandia
ing its identification as the GICE. Interestingly, the values (~+1) of the samples from the top of the Copenhagen Formation are heavier than the baseline values in the lower part of the formation (between –1‰ and –2‰). This could be interpreted as indicating that the complete GICE excursion is not represented (Saltzman and Young, 2005), its top segment either being present in the basal part of the overlying Eureka Quartzite or cut out by the possible unconformity at the Copenhagen-Eureka contact. The GICE occurs in an interval of the Copenhagen Formation that based on conodonts (Phragmodus biofacies of Sweet, 1988) and lithology (much increased amount of shale in Member C) represents a deeper-water environment. Shallow-water hyaline conodonts of the types characteristic of the Turinian successions in the eastern Midcontinent are completely absent. This suggests that in central Nevada the GICE is coeval with a highstand event, as is generally the case in Midcontinent successions such as that in Kentucky, where the GICE appears at the base of the shaly, deeper-water Logana Member of the Lexington Limestone (see above). On the other hand, the Eureka Quartzite, with a sequence boundary at its base, is usually interpreted as a regressive deposit and appears to represent the post-GICE shallowing event. The identity and distribution of the distinctive positive δ13C excursion beneath the GICE in the lower part of Member C of the Copenhagen Formation are currently unclear. This excursion is firmly dated as representing the Baltoniodus gerdae Subzone of the Amorphognathus tvaerensis Zone by the occurrence of the subzonal index, B. gerdae, which is restricted to this subzone
ANTELOPE VALLEY, NEVADA m
Eureka Qtz.
50
B
A 0
Antelope Valley Ls.
Pl. aculeata-E. quadr. B. variabilis ? B.g.
C
100
B. com ?
P. undatus ?
GICE
137
C. sweeti P. anserinus
CHAZYAN
? TURINIAN COPENHAGEN
CHATFIELDIAN
STAGE FM.
biostratigraphy (Ethington and Schumacher, 1969; Harris et al., 1979; Sweet et al., 2005). The first δ13C curve from the Copenhagen Formation was recently published by Saltzman and Young (2005). Their data were based on samples from two sections of the Copenhagen in the Monitor and Antelope Ranges, where the unit has a total thickness of ~140 m. Most of the δ13C values from the lower 50 m of the formation are between –1‰ and –2‰ (Fig. 11). These are regarded as baseline values and are comparable to those from the underlying Antelope Valley Limestone (Saltzman and Young, 2005). Based on conodonts, this part of the Copenhagen Formation ranges through the uppermost Whiterockian Series (upper Chazyan Subseries). At about the 55-m level, near the contact between Members B and C, the δ13C values increase conspicuously to ~+1‰. High values are maintained to about the 80-m level, where the curve drops down markedly to ~ −1‰, indicating the end of this excursion. This value drop is followed by another, even more drastic, positive δ13C increase, with maximum curve values of ~+3.5‰. The values decrease again through the uppermost 25 m of the formation and are ~+1‰ at the top of the formation. The top 30–35 m of the Copenhagen Formation contains Phragmodus undatus, Polyplacognathus ramosus, and other conodonts present in the Phragmodus undatus Zone elsewhere, and this interval is referable to the Phragmodus undatus Zone that straddles the Turinian-Chatfieldian Stage boundary. These biostratigraphic data indicate that the prominent excursion in the upper 35 m of the formation occupies the same stratigraphic interval as the GICE in the eastern Midcontinent, thus confirm-
47
-2
-1
0
1 13
C
carb.
2
3
4
Figure 11. GICE in the Copenhagen Formation in the Antelope Valley area, central Nevada (after Saltzman and Young, 2005). GICE peak values approach 4‰. Note that conodont biostratigraphy (Harris et al., 1979) shows that the interval of the apparent δ13C excursion between 55 and 85 m is of Turinian age and hence significantly older than the GICE. The top of the Copenhagen Formation is a prominent sequence boundary, and it is possibly associated with a stratigraphic gap that may cut out the topmost part of the excursion. B.g.—B. gerdae.
48
Bergström et al.
(Bergström, 1971) in an interval corresponding to the upper half of this excursion. The base of this subzone has been taken as the base of the North American Turinian Stage (Ross et al., 1982), and hence the excursion is of early Turinian age and clearly older than the GICE and the late Turinian pre-GICE excursions recognized in the Upper Mississippi Valley. No obvious excursion has yet been recorded from coeval strata in the eastern Midcontinent or in the stratigraphically more condensed but well-studied successions in Estonia (Ainsaar et al., 2004a; Meidla et al., 2004; Kaljo et al., 2004; Martma, 2005). In fact, in the Estonian and Swedish successions the δ13C curve shows a dip from ~+1‰ to ~0‰ in the B. gerdae Subzone. This suggests that the apparent excursion in central Nevada may be a local anomaly and not a global feature in the δ13C curve. However, if decisive biostratigraphic data were not available, this excursion, together with the “real” GICE, could readily be incorrectly interpreted as a 2-peak GICE similar to that in many North American and European sections. This underscores the necessity of having adequate chronostratigraphic control for reliable identification of the GICE and other isotope excursions. TRANSATLANTIC CORRELATION OF THE GICE The fact that the GICE is recognizable across much of the North American continent suggests that it is a globally distributed excursion. If so, it should be present in coeval successions on other continents such as Baltica. High-resolution biostratigraphy is probably the only reliable means of establishing correlation of an excursion between continents, although its position in a global eustatic curve, and relation to widespread event–stratigraphic markers such as major K-bentonite beds, may provide supporting evidence for consanguinity. On the basis of recent investigations the biostratigraphic position of the GICE is well established in North America. As recently shown (Bergström et al., 2004; Young et al., 2005), this excursion starts near the boundary between the Phragmodus undatus and Plectodina tenuis Midcontinent Conodont Zones, and ends in the upper part of the latter zone, based on sections in the Upper Mississippi Valley, Kentucky, Tennessee, Oklahoma, New York State, and the Appalachians (Fig. 12). There is also evidence from these sections, especially those in New York State, Kentucky, and Oklahoma, that the GICE interval is in the upper part of the Amorphognathus tvaerensis Atlantic Conodont Zone and ends well below the base of the overlying Amorphognathus superbus Conodont Zone. In terms of graptolite biostratigraphy, the GICE begins near the base of the global Diplacanthograptus caudatus Zone and ends below the base of the North American Diplacanthograptus spiniferus Zone. Recent work (Goldman et al., 2005) has also shown that the beginning of the GICE at the Fittstown section in Oklahoma is just above an interval containing chitinozoans of the Baltic Spinachitina cervicornis Zone. This zone occurs in the upper Haljala and Keila Stages in Baltoscandia (Nõlvak and Grahn, 1993). Finally, depending on the local net depositional rate, the GICE starts a few decimeters to a
few meters above the Millbrig K-bentonite, which is an excellent regional marker bed in the sections in eastern North America. In Baltoscandia, a prominent δ13C excursion (the mid-Caradoc excursion of Kaljo et al., 2004) is present in the upper part of the Amorphognathus tvaerensis Zone in strata representing the Keila and Oandu Stages (Fig. 12). The graptolite correlation of the beginning of this excursion is not yet firmly established in Baltoscandia, but on the basis of sections in southern Sweden it is likely to be in the uppermost part of the Diplograptus foliaceus (formerly D. multidens) Zone. Diplacanthograptus spiniferus, the index of the North American D. spiniferus Zone, is present in the post-GICE Rakvere and Nabala Stages in Estonia (Bergström et al., 2004). In Baltoscandia, and especially in the many studied sections in Estonia, the beginning of the excursion has been recorded from the upper part of the Spinachitina cervicornis Chitinozoan Zone of the middle-upper part of the Keila Stage. Finally, in the Baltoscandian sections the excursion starts a few meters above the Kinnekulle K-bentonite, which has been correlated with the North American Millbrig K-bentonite on the basis of a variety of evidence (Huff et al., 1992; Bergström et al., 2004). Hence, the GICE and the Baltoscandian “mid-Caradoc” excursion occupy precisely the same stratigraphic interval. The possible transatlantic equivalence of this interval was first suggested by Huff et al. (1992) and has also been discussed by, among others, Ainsaar et al. (1999) and Bergström et al. (2001, 2004). Based on all the evidence at hand, it seems clear that these excursions are the same, and so it is appropriate to use the GICE designation also for the Baltoscandian excursion (see Bergström et al., 2004). For a diagrammatic illustration of the biostratigraphic relations of the GICE in North America and northern Europe, see Saltzman et al. (2003a) and Bergström et al. (2004, 2007). GICE IN SWEDEN Lower Upper Ordovician strata that represent the global Sandbian and Katian stages are present in several areas in southern and central Sweden (Fig. 13), especially in the provinces of Scania, Västergötland, Östergötland, and Dalarna. As shown in Figure 14, the GICE interval is in the Skagen and Moldå Formations and in equivalent strata in carbonate mound facies, the Kullsberg Limestone. Apart from a thin development of the Skagen Limestone, most of the excursion interval is in shaly facies in Scania (Nilsson, 1977; Bergström et al., 1999; Pålsson, 2002), and no GICE data are currently available from that region. The GICE interval is not present on the Island of Öland, and it is poorly developed, if present at all, in the subsurface of the island of Gotland in the Baltic (Nõlvak and Grahn, 1993). The only GICE data published from Västergötland (Meidla et al., 2004) are based on a few samples collected from the Skagen Limestone at the famous Gullhögen Quarry. Because their recorded stratigraphic succession of the interval just above the Skagen Limestone differs significantly from that presented from the site by Holmer (1986), further studies are needed to establish
0
1
2
13
C
carb.
3
B. compressa
C. bicornis
0
1 13
C
carb.
2
C. bicornis
C. spiniferus
GRAPTOLITE ZONE
D. clingani
RAKVERE KEILA
Kinnekulle K-b.
RANGES CHITINOOF KEY ZOAN GRAPTOZONE LITES
F. fungiformis
S. cervicornis
D. foliaceus
Deicke K-b.
S. cervicornis
A. tvaerensis
TURINIAN
A. tvaerensis
OANDU
HALJALA
GICE
P. undatus
Millbrig K-b.
C. americanus
P. tenuis
C. caudatus
C. B. A. spiniconflu- superbus ferus ens
CHATFIELDIAN
STAGE
K-BENTONITES CONODONT ZONE GRAP- CHITINOK-BENTONITES STAGES & TOLITE ZOAN & C EXCURSION C EXCURSION MIDCONT. ATL. ZONE ZONE
A. CONODONT superbus ZONE
BALTOSCANDIA
NORTH AMERICA
B. hirsuta L. dalbyensis
Figure 12. Transatlantic correlation diagram, showing the relations between conodont, graptolite, and chitinozoan biostratigraphy, K-bentonite event stratigraphy, and the early Chatfieldian δ13C excursion in North America and Baltoscandia. On the basis of these relations, the GICE is considered the same δ13C excursion as the “middle Caradocian” excursion of Kaljo et al. (2004). The δ13C curves in the North American and Baltocandian parts of the diagram are based on parts of the Kentucky composite curve and the Fjäcka curve, respectively. For C. spiniferus read D. spiniferus.
Figure 13. Sketch map of the southern part of Baltoscandia, showing study localities. For the location of other δ13C sites in Estonia and Latvia, see Kaljo et al. (2004, Fig. 1B). Black denotes Ordovician outcrop areas.
50
Bergström et al.
SERIES
S TA G E
GLOBAL BRITISH N. AM. BALTOS. GLOBAL BALTOS. N. AMER. ASHGILLIAN
CONODONT SWEDISH LITHOZONES STRATIGRAPHIC UNITS
RICHMONDIAN
CINCINN- HARJUAN ATIAN
NABALA
MAYSVILLIAN EDENIAN
A. ordovicicus
SLANDROM Amorphognathus superbus
SKÅLBERG
RAKVERE
CARADOCIAN
OANDU CHATFIELDIAN
VIRUAN KEILA
Amorphognathus tvaerensis
FREBERGA
KATIAN UPPER ORDOVICIAN
C
MOLDÅ
KULLSBERG
GICE
SKAGEN
SANDBIAN
HALJALA
Millbrig K-bent. TURINIAN
Kinnekulle K-bent. DALBY
Figure 14. Stratigraphic diagram showing series, stages, conodont zones, and Swedish formations discussed in the text.
the relations between the GICE and the Gullhögen succession. No well-exposed outcrop of the GICE interval is known from Östergötland, but it is completely penetrated by core drilling at Smedsby Gård in the northwestern part of the province (Fig. 13), which has been investigated in the present study. The best GICE sections in Sweden are in the Siljan area in Dalarna (Fig. 15), where the excursion interval is developed in carbonate facies and is well exposed at several localities. In the present study we describe new δ13C studies of three localities in the Siljan area and of the Smedsby Gård drill core. These sites were selected for study because they represent different depositional facies (inner shelf, mound, and outer shelf, respectively), and it would be of interest to determine if the GICE was environmentally controlled. Siljan Area, Dalarna A pioneer study of δ13C and δ18O in Upper Ordovician rocks of this region, and in entire Baltoscandia, was carried out by Jux and Manze (1979). This interesting but widely overlooked report was based on numerous samples through both the GICE and the HICE intervals, but the authors did not recognize any excursions. Subsequently, some δ13C data from the Siljan area were reported by Tobin and Walker (1996, 1997), Bergström et al. (2001, 2004), Saltzman et al. (2003a), Ainsaar et al. (2000, 2004a), and Tobin et al. (2005), but the only GICE curves published previously are from the Fjäcka section. Below we present new δ13C data from sections at Fjäcka, Skålberget, and Amtjärn (Fig. 15). These δ13C data, as well as those from the Smedsby Gård drill core, were obtained from limestone samples using the same standard laboratory procedures as described in Bergström et al. (2006).
Fjäcka This outcrop, which is one of the best known and most studied Ordovician sections in Sweden (see, e.g., Jaanusson and Martna, 1948; Jaanusson, 1963, 1976, 1982; Laufeld, 1967; Bergström, 1971; Holmer, 1989; Nõlvak and Grahn, 1993; Nõlvak et al., 1999), lies along the Fjäckan rivulet (Fig. 16). It provides a continuous exposure of the Skagen and Moldå Formations, as well as overlying and underlying units, in a stratigraphically condensed succession. Our new sampling started in the upper part of the Dalby Formation and ranged through the Skagen and Moldå Formations into the lowermost part of the Slandrom Limestone (Fig. 17). The δ13C values from the stratigraphically lowest 14 samples, which are from the upper Dalby Formation and the lower Skagen Formation, range from ~+0.5‰ to +1‰, which we consider to be pre-GICE baseline values. At ~2.5 m above the Kinnekulle K-bentonite, there is a marked but gradual increase in values to near +2‰. These heavy values last for several meters before dropping to baseline values ~4 m below the top of the Moldå Formation. Of note is a dip in the δ13C curve near the base of the Moldå Formation that gives the curve a 2-peak appearance similar to that in the GICE curve from some localities in North America. We consider the GICE to end ~2 m above the base of the Moldå Formation, where the values are ~+1, hence the same as the pre-GICE baseline values. The total thickness of the GICE at Fjäcka amounts to ~5 m. This is slightly larger than the thickness of the GICE interval in the cratonic sections in the Upper Mississippi Valley, but only a fraction of its thickness at Dolly Ridge, Hagan, and Reedsville. Conodonts (Bergström, 1971, 2007) as well as chitinozoans provide the best biostratigraphic control of the excursion interval. The uppermost Dalby Formation contains conodonts of the
Upper Ordovician Guttenberg δ13C excursion (GICE) in North America and Baltoscandia
51
Figure 15. Sketch map of the Siljan area, showing location of our studied GICE sections. Inset map illustrates the position of the carbonate mounds at Amtjärn, Skålberget, Kullsberg, and Unkarsheden.
Baltoniodus alobatus Subzone of the Amorphognathus tvaerensis Zone. Two important conodont species, Hamarodus europaeus and Amorphognathus complicatus, which probably represent the Amorphognathus superbus Zone, appear ~1 m above the end of the GICE (Fig. 17). Diagnostic conodont taxa have not been found in the Skagen Formation, and its conodont zone assignment remains uncertain, although it is likely to represent the upper part of the A. tvaerensis Zone on the basis of chemostratigraphy. In terms of Baltoscandic chitinozoan zones, the GICE lies in the middle-upper part of the Spinachitina cervicornis Zone (Nõlvak and Grahn, 1993; also see Laufeld, 1967), which has its reference section in the Fjäcka outcrop. Skålberget Skålberget is a relatively small, now disused and partly waterfilled, quarry in the southeastern part of the Siljan area (Fig. 15). It was quarried for the Kullsberg and Boda Limestones in the 1930s and 1940s (Thorslund, l935, l936). Although abandoned
for more than half a century, it still provides an excellent section through the Kullsberg Limestone and the overlying Skålberg and Slandrom Limestones, the Fjäcka Shale, and the Jonstorp Formation. A quarry pit formerly exposing the Boda Limestone and the overlying Llandovery shales has now been filled in, so that exposure is completely destroyed. A thickness of ~50 m of the Kullsberg Limestone is now exposed along the southern rim of the quarry (Fig. 18). The Kullsberg Limestone is a carbonate mound deposit that forms a >200-m-long elongated lens of in part poorly stratified limestone. The dip of the beds varies from very steep to vertical. Because the basalmost part of the Kullsberg Limestone is covered by Pleistocene till, its total thickness at this locality is unknown. In view of the fact that the quarrying operations at other Kullsberg sites included virtually the entire formation, it appears likely that its total thickness at Skålberget does not substantially exceed 50 m. The record of “Macrourus-Kalk und Mergel” (= Moldå Formation in current terminology) by Jux and Manze (1979, Fig. 3) beneath the Kullsberg Limestone at Skålberget is most unlikely
52
Bergström et al. rate, our Kullsberg values compare well, and in most cases they exceed the GICE values from the non-mound facies at Fjäcka, and we interpret them as indicating that the GICE ranges through the entire thickness of the Kullsberg Limestone.
Slandrom Lst.
N
sampled interval mill -ra
ce
Moldå Fm.
Ki
Skagen Fm.
nn
0
ite on nt le be ul K-
K-be nton ite
ek
Dalby Fm.
5m
Figure 16. Detailed sketch map of part of the Fjäcka outcrop and our sampled interval (marked by thick black line) in the Skagen and Moldå Formations. Note the Kinnekulle K-bentonite at the base of the Skagen Formation.
to be correct in view of the fact that the equivalent of the Moldå Formation (Skålberg Limestone) overlies rather than underlies the Kullsberg Limestone (Jaanusson, 1982). For details about the Kullsberg lithology and geochemistry, see Jaanusson (1982), Tobin and Walker (1997), and Tobin et al. (2005). Some δ13C data from Skålberget were previously published by Jux and Manze (1979), Tobin and Walker (1997), and Tobin et al. (2005). Our new data are from a series of samples collected through a 50-m-thick section along the southern rim of the quarry (Fig. 18). The sampled Kullsberg section illustrated by Jux and Manze (1979, Fig. 3) is only ~25 m thick. This small figure suggests that their samples came from the northern end of the Kullsberg lens, where the thickness is markedly smaller. The δ13C values from our sampled profile are extraordinarily uniform, being about +2‰ all through the succession. The values presented by Jux and Manze (1979) are also uniform, but strangely they are lighter, mostly ranging from 0‰ to +1‰. We have no explanation for this consistent difference, and Jux and Manze (1979) did not indicate what standard they used. At any
Amtjärn Another mound of the Kullsberg Limestone is ~1 km SW of Skålberget (Fig. 15). It is of slightly smaller dimensions than that at Skålberget, and the quarry exposing it is known as the Amtjärn Quarry (Thorslund, 1935, 1936). In the vicinity of this quarry there are outcrops, now mostly overgrown and covered, of both underlying and overlying units (Laufeld, 1967; Thorslund and Jaanusson, 1960; Jaanusson, 1982). Of special interest for the present study is a section a few meters thick of the uppermost Dalby Formation just below the Kullsberg mound near the south entrance to the quarry (Thorslund and Jaanusson, 1960, Fig. 12). This may be the only locality in the region that currently exposes the base of the Kullsberg Limestone as well as subjacent strata. Condonts from the Dalby Formation at this site show that the exposed beds represent the topmost part of the unit and the Baltoniodus alobataus Subzone of the Amorphognathus tvaerensis Zone (Tobin et al., 2005). The δ13C data from the Dalby Formation show baseline values of ~1‰ (Fig. 19), but they get heavier upward, and this increase continues into the lowermost part of the overlying Kullsberg Limestone to reach values of +1.7‰ at the end of the outcrop. We interpret this increase as marking the beginning of the GICE. Further sampling at Amtjärn may help bridge the gap in the δ13C curve between the Amtjärn and Skålberget successions, but we believe this gap to be relatively insignificant. The importance of the Amtjärn succession is that it shows that the GICE is present also in the lowermost part of the Kullsberg Limestone mound. It should be noted that the GICE interval is >50 m thick in the Kullsberg mound facies, hence some 10 times as thick as at Fjäcka that represents the inter-mound facies. However, the end of the excursion is currently poorly established in the mound facies, because at Skålberget there is a stratigraphic gap at the top of the mound lens. However, based on conodonts, and especially the important species Hamarodus europaeus, the Skålberg Limestone that overlies the Kullsberg Limestone correlates with the post-GICE part of the Moldå Formation at Fjäcka, and the end of the GICE therefore may be at the top of the Kullsberg Limestone. Smedsby Gård, Östergötland The well drilled by the Geological Survey of Sweden in 1946 at Smedsby Gård near the town of Motala (Fig. 13) penetrated the GICE interval as well as older and younger strata (Jaanusson, 1962; Wikman et al., 1982). Because the succession corresponding to the Skagen and Moldå Formations is uniform lithologically in this region, Jaanusson (1982) introduced the
Kinnekulle K-bentonite
-6
-8
0
1
2
13
C
carb.
Amorphognathus complicatus Hamarodus europaeus
80
85
Kinnekulle K-benonite
Baltoniodus alobatus
-4
SLANDROM LS.
?
Amorphognathus tvaerensis B. alobatus not named
-2
Amorphognathus tvaerensis
Baltoniodus alobatus
SKAGEN FM.
Core missing
2
0
DALBY FM.
75
GICE
?
m
74
FREBERGA FM.
4
53
SMEDSBY GÅRD CORE
DALBY FM.
6
Am. superbus
Amorphognathus complicatus Hamarodus europaeus
Depth m
GICE
MOLDÅ FM.
SLANDROM LS.
FJÄCKA
CONODONT ZONE CONODONT SUBZONE
Upper Ordovician Guttenberg δ13C excursion (GICE) in North America and Baltoscandia
90
0
1
2
13
C
carb.
Figure 17. Comparison of the GICE in the Fjäcka and Smedsby Gård sections, based on data from the present study. Note that the baseline value of the δ13C curve is ~+1‰, and the GICE peak value approaches +2‰ in both successions.
term Freberga Formation for this interval, and this designation is used herein. Although Wikman et al. (1982) briefly referred to the Freberga Formation portion of the core, this interval has not yet been described in any detail. In the lower 5 m of the Freberga Formation the δ13C values are between 0‰ and +1‰ (Fig. 20), which are considered baseline values. At about the 83 m core depth, the values show a distinct increase to nearly +2‰, and we regard this as the beginning of the GICE. For a stratigraphic thickness of ~5 m, the δ13C values range between +1‰ and +2‰ before decreasing to baseline values of ~1‰ at 78–79 m core depth, which marks the end of the GICE. As seen in Figure 17, the GICE curve is closely similar to the Fjäcka curve, and the excursion has a similar thickness. Also the biostratigraphy of the Freberga Formation resembles closely that of the Skagen and Moldå Formations at Fjäcka. The uppermost part of the Dalby Formation contains conodonts of the Baltoniodus alobatus Subzone of the Amorphognathus tvaerensis Zone. As in the Fjäcka section, the conodonts Hamarodus europaeus and Amorphognathus complicatus appear just after
the end of the GICE in the upper part of the Freberga Formation at virtually the same level as in the Fjäcka section. The Freberga Formation is slightly thicker (~13.3 m) compared with ~10 m of the combined thickness of the Skagen and Moldå Formations at Fjäcka, but the δ13C chemostratigraphy and biostratigraphy are remarkably similar. GICE IN ESTONIA AND LATVIA An exceptionally large amount of δ13C work has been carried out in recent years in the Ordovician of eastern Baltoscandia, and especially in Estonia (Ainsaar et al., 2004a, 2004b; Kaljo et al., 1999, 2004). In fact, the stratigraphical and regional δ13C variations in that system are better known in Estonia that in any other region of similar size in the world. Furthermore, biostratigraphy, lithofacies, faunas, and lithostratigraphy have been investigated in great detail in both outcrops and many drill cores. For a useful general summary of non-isotopic information, see Raukas and Teedumäe (1997).
54
Bergström et al.
N. ENTRANCE
SKÅLBERGET & AMTJÄRN
N
m SKÅLB. KULLS
50
Fjäcka Sh.
Boda Lst. Tunnel section
30 25 20 15 10 5 0
S. entrance section 0
covered
30
AMTJÄRN SOUTH
sampled section
5
DALBY FM.
m
Figure 18. Sketch map of the Skålberget locality (after Jaanusson, 1982), which is now a nature preserve. Our sampled section through the 50+-m-thick Kullsberg Limestone at the south margin of the quarry is marked by a thick black line.
GICE
Sla nd rom
Kullsberg Lst.
N. entrance section
35
SKÅLBERGET SOUTH
water-filled pit
KULLSBERG LIMESTONE
40
Skålberg Lst. Lst .
no ou tcr op s
45
0
covered
0
1
2
3
13
C
carb.
The Upper Ordovician succession in eastern Baltoscandia includes mostly calcareous epicontinental rocks deposited in shallow-water environments at temperate to subtropical latitudes. By and large, water depth increased in a southern direction in terms of present latitudes, with the shallowest environments being in northern Estonia and somewhat deeper waters in southern Estonia and Latvia. Although the Upper Ordovician succession in both northern and southern Estonia has several stratigraphic gaps (Nõlvak and Grahn, 1993; Ainsaar et al., 2004a; Kaljo et al., 2004, Fig. 2), these tend to be relatively minor, and as a whole the portion of the succession pertinent to the present study is remarkably complete stratigraphically. A significant δ13C excursion in the upper Keila and Oandu Stages, referred to as the mid-Caradoc excursion by Kaljo et al. (2004), is here identified as the GICE on the basis of biostratigraphic and other evidence. The report by Ainsaar et al. (1999) is apparently the first to compare this Estonian excursion with the Guttenberg excursion (GICE) in North America. As indicated by drill core investigations, the excursion is best developed in southern Estonia and Latvia. This may at least partly
Figure 19. Composite δ13C curve, based on our Skålberget and Amtjärn sections. The baseline δ13C values are likely to be ~+1‰, and the GICE peak values through the Kullsberg Limestone are extraordinarily uniform, being mostly between +2‰ and +2.5‰. Note that the top of the GICE is not yet well established at Skålberget, but the topmost part of the GICE curve may well be cut out by the unconformity at the top surface of the Kullsberg Limestone mound.
reflect the widespread stratigraphic gap at the Keila-Oandu contact that may cut out parts or all of the upper Keila GICE interval in many sections in northern Estonia. For southern Estonia, instructive excursion curves have been published from, for instance, the Ristiküla-174 (Meidla et al., 1999; Ainsaar et al., 2004a), Viljandi (Kaljo et al., 2004), Männamaa F-367 (Ainsaar et al., 2004a), Mehikoorma-421 (Martma, 2005), and Valga–10 (Ainsaar et al., 2004a) drill cores. In Latvia, the excursion is present in the Kandava (Brenchley et al., 1996) and Jurmala (Ainsaar et al., 2004a) drill cores. Because of its stratigraphic completeness and informative δ13C curve, we select the
Upper Ordovician Guttenberg δ13C excursion (GICE) in North America and Baltoscandia
74 75 Core missing
406.7
KEILA
80
GICE
402.5
GICE
FREBERGA FM.
401.1
85
90
0
1 C
13
2
carb.
424.6
Kinnekulle K-b.
0
HALJALA
Kinnekulle K-benonite
Baltoniodus alobatus
DALBY FM.
RAKVERE
Depth m
OANDU
m
NABALA ESTONIAN STAGES
RISTIKÜLA 174 CORE, SOUTHERN ESTONIA Amorphognathus complicatus Hamarodus europaeus
SLANDROM LS.
SMEDSBY GÅRD CORE
55
1 C
2
13
carb.
Figure 20. Comparison between the GICE in the Smedsby Gård, Sweden, and the Ristiküla, southern Estonia (Ainsaar et al., 1999), drill cores. Note the close similarity of the δ13C curves, despite the fact that these sites are more than 600 km apart. In both successions the baseline values are between 0‰ and 1‰, and the GICE curve peak values are near 2‰. Based on chemostratigraphic correlation, the base of the Oandu Stage would correspond to a level at ~79 m in the Smedsby Gård drill core.
Ristiküla-174 drill core as an instructive illustration of the GICE in eastern Baltoscandia. Ristiküla-174 This core was drilled by the Estonian Geological Survey and was studied by Ainsaar et al. (1996, 1999), who provided extensive faunal and lithologic data. In this drill core an 8–9-m-thick interval above the Kinnekulle K-bentonite, which represents the lower half of the Keila Stage, shows baseline δ13C values between 0‰ and +1‰, with most values being ~+0.5‰ (Fig. 20). Slightly higher stratigraphically there is a conspicuous rise in the δ13C curve to values near +2‰, which mark the beginning of the GICE. These high values persist for the next few meters before the δ13C curve drops to values between +1‰
and +1.5‰; these are followed, slightly higher up, by even lighter values of ~0‰, which are taken as marking the end of the GICE. As a whole, the Ristiküla δ13C curve shows a remarkably close similarity to that of the Smedsby Gård, allowing a close correlation between these successions (Fig. 20). There is also an obvious similarity between the Estonian curve and the Fjäcka one (Fig. 17). This close chemostratigraphic correlation is of special stratigraphic interest, because in this stratigraphic interval the precise relations between the Estonian standard regional stages and the Scandinavian succession have long been uncertain and even controversial (see, e.g., Jaanusson, 1976). Assuming that no significant gaps occur in the Keila-Rakvere succession in the Ristiküla drill core, the chemostratigraphic data suggest that the base of the Oandu Stage corresponds to a level 1–2 m above the base
56
Bergström et al.
of the Moldå Formation in its type section at Fjäcka, and that the top of the Oandu Stage is coeval with a level in the middle part of the Moldå Formation. As noted by several authors, for instance, Ainsaar et al. (1996), the boundary between the Keila and Oandu Stages is in large parts of eastern Baltoscandia marked by a stratigraphic gap that reflects a widespread regression. In the Fjäcka succession there is no obvious lithologic evidence of such a gap, although the more common occurrence of limestone beds with a higher carbonate content in the Moldå Formation compared with those of the Skagen Formation perhaps may be interpreted as indicating a shallowing in the depositional environment. We attribute the apparent lack of the Keila-Oandu boundary gap in the Fjäcka succession as due to the fact that the latter was deposited at a water depth large enough to prevent minor regressions to leave a recognizable imprint in the geological record. δ13C EXCURSIONS ABOVE AND BELOW THE GICE INTERVAL Several, in most cases relatively small, positive shifts in the δ13C curve above and below the GICE interval have been distinguished and named in both North America and northern Europe. However, the question as to whether or not these are local or global excursions has not been adequately addressed in the literature by regional comparisons. The detailed δ13C curves now available from a considerable number of sections provide a timely opportunity to assess the regional distribution and stratigraphic significance of these δ13C perturbations. In their detailed study of the δ13C chemostratigraphy of the Decorah Formation and associated strata in the Upper Mississippi Valley, Ludvigson et al. (2004) recognized four positive pre-GICE δ13C excursions in the late Turinian Stage, namely, in descending order, the Spechts Ferry, Quimby’s Mill, Grand Detour, and Mifflin excursions; three of these are shown in Figure 21. Presumably, owing to the existence of stratigraphic gaps in the succession, not all of these excursions are recognized in the δ13C curve of each locality, but they are best displayed at their localities 2–4 (Ludvigson et al., 2004, Figs. 7–9). The Spechts Ferry excursion, which lies typically between the Deicke and Elkport K-bentonites, is characterized by a positive shift in the δ13C curve just below a negative shift beneath the point of the beginning of the GICE (Fig. 21). However, in most sections the excursion curve differs little in this interval from the baseline values above and below the GICE. The stratigraphic position of the Spechts Ferry excursion just above the Deicke K-bentonite facilitates direct comparison with North American sections outside the Upper Mississippi Valley. For instance, the Tyrone Limestone at Boonesborough, Kentucky (Fig. 21), the Nealmont Limestone of Dolly Ridge, West Virginia (Fig. 6), and Reedsville, Pennsylvania (Fig. 9), and the Carters Formation of central Tennessee (Fig. 7) do not show a significant excursion between the two major K-bentonites. On the other hand, there is a positive shift of ~+1.5‰ between these K-bentonites at Hagan,
Virginia (Fig. 6). In Baltoscandia there is no notable positive excursion directly below the Kinnekulle K-bentonite that could be interpreted as the Spechts Ferry excursion. Hence, based on the evidence at hand, it appears that this excursion is a local, rather than global, shift in the δ13C curve. The stratigraphically next lower excursion recognized by Ludvigson et al. (2004, Figs. 9 and 10), the Quimby’s Mill excursion, was described from the uppermost part of the Platteville Formation just below the Deicke K-bentonite. This excursion is identified in only a few drill cores, but its absence in other sections in the Upper Mississippi Valley may be due to the absence of the excursion interval at the very prominent unconformity at the top of the Platteville Formation. In the Boonesborough, Kentucky, succession (Conkin and Conkin, 1983), there is a minor positive shift in the δ13C values from 0‰ to ~+1‰ within an interval a few meters thick beneath the Deicke K-bentonite (Fig. 21) that might correspond to the Quimby’s Mill excursion in eastern Iowa. However, in the coeval interval in other studied sections outside the Upper Mississippi Valley there appears to be little deviation from the δ13C baseline values, and we consider this excursion to be a local, rather than regional, feature. The lowest two excursions recognized by Ludvigson et al. (2004), the Grand Detour and Mifflin excursions, are characterized by shifts of up to ~2‰ from the baseline values in the δ13C curve (Fig. 21), and we are uncertain about their regional significance. If regionally distributed, they ought to be present in the interval of the Tyrone-Oregon Formations in Kentucky, but the biostratigraphic resolution currently available is insufficient to establish their precise positions within this interval in which there appears to be no significant shift from the baseline values (Fig. 21). At Hagan, there is a minor positive shift (up to ~+1.5 ‰) in the δ13C curve ~30–50 m below the Deicke K-bentonite that might represent one of these excursions, but further studies are needed to confirm this. The contact between the Platteville and the Decorah Formations marks a prominent gap of regionally variable magnitude in the Upper Mississippi Valley succession (Ludvigson et al., 2004, Fig. 3). In some areas, for instance, in northern Illinois, the entire GICE interval is missing. There is no evidence of the corresponding unconformity in the upper Turinian interval below the Deicke K-bentonite in Kentucky. In Baltoscandia, equivalents to the North American Turinian Stage ought to be present in the middle and upper Dalby Formation in Sweden, and in the lower Adze and Dreimani Formations in Estonia, which represent an interval referred to the Baltoniodus alobatus Subzone, and possibly the B. gerdae Subzone, of the Amorphognathus tvaerensis Zone (Bergström, 1971). This interval, below the Kinnekulle K-bentonite, is characterized by a prominent negative δ13C shift in some published δ13C curves from Sweden, Latvia, and Estonia (Ainsaar et al., 2004b, Fig. 1; Meidla et al., 2004, Fig. 1). In the recent description of the Kergula (565) drill core from northern Estonia, Martma (2006, Fig. 4) showed this interval as representing the upper part of the
Upper Ordovician Guttenberg δ13C excursion (GICE) in North America and Baltoscandia
57
Figure 21. Comparison of the Turinian δ13C excursions recognized in eastern Iowa (Ludvigson et al., 2004) and the δ13C curve from the coeval interval in the thicker, and presumably stratigraphically more complete, succession through the upper High Bridge Group at Boonesborough, Kentucky. Whereas the Iowa succession shows δ13C value fluctuations from near –3‰ to ~+0.5‰, most Kentucky values are between 0‰ and ~+1‰. This makes safe comparison of the Iowa δ13C excursions with the Kentucky succession difficult. Interestingly, correlation between the two curves is improved if one assumes that the prominent gap between the Platteville and the Decorah Formations in Iowa corresponds to ~6 m of section represented just below the Deicke K-bentonite in the Kentucky section, where there is no evidence of an unconformity.
“Upper Kukruse low” in the δ13C curve. Hence, in Baltoscandia there is currently no obvious evidence of the existence of the three Turinian positive δ13C excursions recognized in the Upper Mississippi Valley. Investigations after the initial submission of the present study (Bergström et al., 2007, 2009) have shown that the several positive δ13C excursions recognized between the GICE and HICE in Estonia (Kaljo et al., 2004) have their counterparts in the Cincinnati region reference successions of the Cincinnatian Series, the Upper Ordovician in the North American classification. These interesting results, which will be published in some detail elsewhere, indicate that not only GICE and HICE, but these excursions also, are of a global nature.
GICE RELATIONSHIP TO EUSTASY It is generally agreed that the latest Ordovician Hirnantian δ13C excursion (HICE) was associated with a eustatic regression caused by the lodging of very large amounts of water in the continental ice sheets on the Gondwana continents. Although recent studies (Bergström et al., 2006) suggest that this event might have involved more than a single glacial period, the correlation between the positive δ13C excursion and regression(s) appears well established, as it is based on investigation of a considerable number of successions on several continents. The possible relation between GICE and a Gondwana glaciation is more controversial and has been discussed in several
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recent papers (see, e.g., Kaljo et al., 2004; Tobin et al., 2005; Young et al., 2005), but finding a conclusive relationship between GICE, glaciation, and eustatic regression has been elusive. It can be assumed that the development during GICE time of a significant continental glaciation with large ice sheets would result in a glacio-eustatic regression that is likely to have left a recognizable record, especially in the shallow-water cratonic sedimentary successions. Unfortunately, most Late Ordovician sea level curves are based only on a particular region (see, e.g., Ross and Ross, 1992; Nielsen, 2004), and published global sea level curves tend to have limited stratigraphic resolution. Furthermore, it is often difficult to separate regressive events caused by more or less local epeirogeny from those resulting from eustasy. As noted by Tobin et al. (2005), in order to separate epeirogenic effects from eustasy, it is useful, if not necessary, to examine coeval sea level events on several continents. Despite these problems, in a regional study like the present one it is of interest to examine the relations between the GICE and bathymetric indications in the rock record. If it could be shown that this excursion was globally associated with a eustatic regression similar to the regression present in the HICE interval, it would strongly suggest the existence of an ice age during GICE time. Unfortunately, the evidence used in support of the existence of such a glacial period during GICE time in northern Africa is highly controversial (Hamoumi, 1999, 2003) and is not considered conclusive by most workers. Abundant evidence indicates that in the North American Midcontinent, late Turonian strata, such as the Tyrone Formation of Kentucky (Cressman and Noger, 1976), the Bromide Formation of Oklahoma (Amsden and Sweet, 1983), the Platteville Formation and coeval formations of the Upper Mississippi Valley (Witzke and Bunker, 1996), the Carters Formation of central Tennessee (Holland and Patzkowsky, 1998), the Eggleston Formation of Virginia, the Nealmont Formation of central Pennsylvania, and the Black River Group of New York–Ontario (Brett et al., 2004), represent very shallow-water (mostly peritidal) facies that has been described as having “layer-cake” stratigraphy. The upper part of this succession represents the M4 sequence of Holland and Patzkowsky (1996, 1998). Commonly, the top of this sequence is an omission surface (Kolata et al., 1998) that locally, as in Kentucky, has a relief of several meters (Cressman, 1973). The strata overlying this unconformity, which were regarded as the lower part of their M5 sequence by Holland and Patzkowsky (1996, 1998), are known by different formation names in different states, such as the Lexington Limestone in Kentucky, the Viola Springs Formation in Oklahoma, the Decorah Formation in the Upper Mississippi Valley, the Dolly Ridge Formation in West Virginia, and the Hermitage Formation in central Tennessee. All these units show evidence of having been formed in deeper water than the underlying Turinian peritidal carbonates, and clearly they represent sediments deposited during a sea level rise. The GICE starts in the Logana Member of the Lexington Limestone in Kentucky and in a coeval position in other sections near, or at, what has been described as a maximum flooding surface (Brett
et al., 2004). Hence, the major part of the GICE occurs in a highstand systems tract in Kentucky, and the same is the case in the other regions just listed. The upper part of the GICE is present in strata laid down in a shallowing depositional environment characterized by the occurrence of grainstones, etc. A similar trend can be recognized in the central Nevada succession, where the moderately deep-water Copenhagen Formation with the GICE is overlain by the very shallow-water Eureka Quartzite (Saltzman and Young, 2005). To summarize, in North America at least a significant part of the GICE appears to be present in a transgressive, rather than regressive, part of the Ordovician sedimentary record. In this respect it would seem to differ from the HICE. The bathymetric conditions in Baltoscandia were less uniform than in North America during the GICE interval, and two recently proposed interpretations of water depth are partly contradictory. Kaljo et al. (2004, Fig. 5) described the initial part of the GICE succession in the upper Keila stages of Estonia as having been deposited during a regression that culminated at the end of Keila time as shown by the unconformable Keila-Oandu contact, which is associated with a stratigraphic gap, especially in northern Estonia (Nõlvak and Grahn, 1993). The upper part of the GICE interval was interpreted as representing a minor transgression in early-middle Oandu time. Hence, in this bathymetric interpretation, assuming that the transatlantic correlation of the GICE is correct, the Estonian sea level changes would be opposite to those displayed in the North American Midcontinent. Nielsen (2004) proposed a partly different interpretation of the sea level history in Baltoscandia during this particular time interval. Based primarily on the Caledonian foreland basin succession in the Oslo region of southeastern Norway, his Keila Drowning Event in the lower-middle Keila Stage was followed by a lowstand in the upper Keila and Oandu Stages, the Frognerkilen Lowstand Event. On the other hand, Kaljo et al. (2004) interpreted most of the Oandu deposits as representing a deepening event. Most of the following, post-GICE, Rakvere Stage was shown by Nielsen (2004) to represent his Nakkholmen Deepening Event, whereas Kaljo et al. (2004) showed this as a lowstand episode. Hence the interpretations by Kaljo et al. (2004) and Nielsen (2004) are in conflict with each other in the GICE interval, and neither of these alternatives fits very well with the sea level history of the North American Midcontinent, although Nielsen’s (2004) interpretation would appear to be closest to the bathymetry of the latter region. Nielsen and Meidla (2004) tried to resolve some of the problems in the correlation of sea level curves between Norway and Estonia, but there are still differences in interpretation. In Sweden the GICE starts in the upper part of the Skagen Formation and in the coeval basal part of the Kullsberg Limestone. For a recent discussion of the general bathymetric conditions in the Siljan region, see Tobin et al. (2005). The strata of the GICE interval were deposited in substantially deeper water than coeval rocks in Estonia and hence were less affected by relatively minor bathymetric changes. Although estimating the
Upper Ordovician Guttenberg δ13C excursion (GICE) in North America and Baltoscandia precise water depth is difficult, the topographic relations between the Kullsberg mounds and the surrounding level bottom where the Fjäcka succession was deposited may provide a minimum local depth figure. As shown above, the GICE occurs through the >50-m-thick Kullsberg Limestone and through ~5 m of the non-mound Fjäcka succession. According to Tobin et al. (2005), the depth in the Kullsberg depositional environment was likely to have been several tens of meters. Assuming a relatively constant rate of deposition in both mound and non-mound facies during the GICE interval, it would seem that during at least late GICE time, the water depth at Fjäcka must have been >50 m and perhaps as much as 75–100 m, based on the elevation of the mound above the surrounding sea bottom on which the non-mound Skagen and Moldå Formations were deposited. In Scania, southernmost Sweden, the Skagen Formation with the GICE appears to have been deposited during a brief shallowing episode, as shown by the change from dark shale to limestone facies, and in some other areas there is a significant stratigraphic gap at the top of the Skagen Limestone (see Nõlvak and Grahn, 1993) that may mark the peak of the regression. Whether this gap cuts out the upper part of the GICE, as one would expect on the basis of the stratigraphically more complete succession in the Siljan area, is not yet known and requires further study. In view of these problems the only conclusion that can be drawn is that at least at the present time there is no very convincing regional evidence from Baltoscandia and North America that overall the GICE was closely correlated with a eustatic regression. This would make it unlikely that the GICE was somehow related to a significant glaciation episode in the Gondwana region. If there was a glaciation at that time, it would appear unlikely that it involved the formation of large continental ice sheets that trapped substantial amounts of seawater (Tobin et al., 2005). GICE AND CLIMATIC INDICATORS It is a widely held view that overall, the Ordovician Period, except its latest part (Hirnantian Stage), was characterized by a greenhouse-state climate with high CO2 and relatively low O2 (see, e.g., Barnes, 2004). Although this may be true as a generalization, it may not necessarily be valid for relatively short time periods such as the GICE interval. The possible presence of continental ice sheets in the Gondwana region in the Late Ordovician prior to the Hirnantian has been discussed by several authors (see, e.g., Hamoumi, 1999, 2003; Pope and Steffen, 2003; Saltzman and Young, 2005), but the supporting evidence for a glaciation has been equivocal. The apparent absence of undisputed glacial deposits in the well-exposed successions of early Katian age in the Gondwana region has been taken by some authors as important negative evidence for the existence of such a glacial period. Recently, Tobin et al. (2005), on the basis of geochemical evidence, suggested that the GICE was associated with a cooling event without formation of a continental ice sheet. In the following section we will briefly discuss some lithologic, faunal, and geochemical indications from North America and Baltoscandia that
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may have potential for interpretation of the climate, as reflected by the seawater temperature, during the GICE interval. This in turn may help clarify whether or not it is likely that the GICE was associated with a glacial period, as was the case with HICE. Lithofacies During the early part of the Katian Stage, which corresponds approximately with the beginning of the GICE, shallow seas with dominating carbonate deposition covered most of the North American Midcontinent from the Mississippi Valley region to the Appalachian foreland basin. Paleomagnetic data indicate that, for instance, Kentucky lay at ~10°S latitude (Bergström and Noltimier, 1982), and the data at hand suggest that the entire Midcontinent apparently was in the tropical zone. Many limestones are bahamitic, indicating deposition at water temperatures of >22–25 °C (Jaanusson, 1973). Evaporites are uncommon in these strata, and the diverse and abundant benthic faunas suggest normal seawater salinities and good water circulation. In central Kentucky the lithofacies of the GICE interval (essentially the Logana Member of the Lexington Limestone) has been described as being a cool-water type with phosphaterich limestones that differ in some respects from the underlying peritidal bahamitic limestones in the Tyrone Formation (Holland and Patzkowsky, 1996). It has also been proposed that the lithofacies of these carbonate sediments was affected by upwelling from the nearby Sebree Trough (Kolata et al., 1998, 2001), which was an elongate basinal structure that is thought to have had an oceanic connection to the south. But the carbonates of the GICE interval in Kentucky, as well as elsewhere in the Midcontinent, are of a tropical type and differ substantially from the subtropical-temperate cold-water limestones in Baltoscandia. Thus we conclude that the lithofacies of the Laurentian carbonates does not provide evidence of a marked lowering of the seawater temperature that could be used as evidence for the presence of a glaciation. It is also significant that, as noted above, the GICE is in that part of the Katian succession that represents a flooding event in both the Mississippi Valley and in the eastern Midcontinent, which is opposite to what would be expected from a glacioeustatic regression. The lithofacies of the sediments in Baltoscandia corresponding to the early Katian Stage in North America, especially those in Sweden and the Baltic States, have been investigated extensively, and a large amount of information has been published. During GICE time, Baltoscandia occupied a subtropical-temperate position (Spjeldnaes, 1961; Bergström and Noltimier, 1982; Cocks and Torsvik, 2005) at a latitude of 30–35°S. Jaanusson (1973) and Lindström (1984), among others, discussed various environmental parameters bearing on the deposition of the Ordovician limestones in Baltoscandia. Although these authors differed in their interpretation of the water depth in the depositional environment, they agreed that most pre-Hirnantian Ordovician limestones in Baltoscandia were deposited in relatively cold water. Bahamitic carbonates are unknown in the GICE interval in
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Sweden but are present locally in the Oandu Stage in Estonia (Vasalemma Formation) and particularly in coeval shallow-water strata (Mjøsa Limestone) in the Mjøsa region in southeastern Norway (Jaanusson, 1976). These unique successions, which include reef-like structures, were interpreted by Jaanusson (1973) as representing a brief invasion of warm waters into parts of the Baltoscandic depositional basin. Judging from the Estonian succession, this warm-water invasion occurred during the final phase of the GICE. The absence of coeval bahamitic strata in Sweden has been attributed to the fact that the depositional environment was too deep and too cold at the bottom for deposition of such carbonates. The local presence at relatively low latitudes in Baltoscandia of warm-water carbonates in the GICE interval, in a succession otherwise characterized by cold-water limestones, is not what would be expected if the GICE was associated with a cold climate and a glaciation. Hence, the lithofacies evidence from North America and Baltoscandia is not very indicative of a substantial climatic cooling. A general conclusion emerges: namely that the GICE is not restricted to a particular carbonate lithofacies but is present in a spectrum of carbonate rock types representing a variety of depositional environments from warm, tropical, shallow-water seas to subtropicaltemperate, relatively deep and cool epicontinental sites. Faunas It is a well-established fact that the Ordovician is one of the Phanerozoic periods with the most conspicuous biogeographic differentiation in the faunas (Jaanusson, 1976; Jaanusson and Bergström, 1980; McKerrow and Scotese, 1990; Webby et al., 2004). A primary reason for this is clearly the vast latitudinal spacing of the continents during this time period, with large regions at different latitudes with different climates from the equator to the poles. The GICE interval provides an instructive illustration of the faunal differentiation at different latitudes with associated differences in climate and seawater temperatures. The Midcontinent megafossil faunas in the North American GICE interval are of tropical types and are dominated by brachiopods and bryozoans and with relatively less abundant occurrences of other shelly fossils such as trilobites, pelecypods, and corals. For useful reviews of the megafossil faunas in the Upper Mississipi Valley, see Sloan (2005), and for those in Kentucky and nearby states, see Pojeta (1979, and the several other U.S. Geological Survey monographs in the series “Contributions to the Ordovician Paleontology of Kentucky and Nearby States”). Useful information is also available in Feldmann and Hackathorn (1996). These faunas have very little in common, even at the generic level, with most coeval megafossil faunas in Baltoscandia. For extensive information about the latter faunas, especially the diverse ones in Estonia, see Rõõmusoks (1970). Also the Baltoscandian microfossil faunas in the GICE interval differ strikingly from those from the North American Midcontinent (Bergström, 1990) apart from the presence of a few pandemic and in most cases long-ranging taxa. However, there is
one important exception to this statement—the conodont fauna of the Mjøsa and Svartsaetra Limestones in Norway and, to a lesser extent, that of the Saku Member of the Vasalemma Formation of Estonia (Bergström, 1997; Bergström et al., 1998). The conodont fauna of the Mjøsa Limestone is of a unique type in Europe and is closely similar to the tropical one from the Lexington Limestone in Kentucky (Bergström et al., 1998); but it is strikingly different from those of the coeval Skagen and Moldå Formations in Sweden and corresponding strata in Norway (Hamar, 1966). Interestingly, these condont faunas are associated with some shelly fossils of Laurentian type and occur in shallow-water bahamitic limestones (see above). There is no significant faunal change that correlates closely with the GICE interval. In the North American Midcontinent succession, the Turinian very shallow-water shelly and conodont faunas differ in several respects from the Chatfieldian ones, but this change occurs in the uppermost Turinian, hence well below the beginning of the GICE. For instance, in the well-known succession in the Upper Mississippi Valley, the faunal change is near the Deicke K-bentonite (Sloan, 2005), and it might be at a corresponding interval in Kentucky and adjacent states (Patzkowsky and Holland, 1996, 1997), although the paucity of fossils in the uppermost Tyrone Limestone and coeval strata makes it difficult to recognize details in the faunal succession. At any rate, this faunal change is attributed to the marked environmental change from peritidal to deeper, although not very deep, conditions at the beginning of the M5 sequence. In an informative review, Ainsaar et al. (2004a) summarized various faunal changes in the Keila-Oandu succession in northern Estonia. The most significant is that at the top of the Keila Stage, which is commonly an unconformity and a stratigraphic gap of regionally variable magnitude, not only in northern Estonia but also in parts of eastern and south-central Sweden (Nõlvak and Grahn, 1993). It is important to note that this gap, which was caused by a regression in the latest Keila sea, is in the middle to upper part of the GICE curve (Fig. 20). If one assumes that the GICE curve covers the same chronostratigraphic interval in Baltoscandia and in North America, as suggested by a variety of evidence presented above, the faunal change in Estonia occurred substantially later than the one near the top of the Turinian Stage in North America, and the North American M4-M5 sequence boundary is clearly older than the latest Keila regression in northern Estonia. Hence, these faunal and depositional changes are not coeval, as suggested by Ainsaar et al. (2004a, p. 128), but diachronous. Accordingly, they are not very useful evidence for a possible global oceanographic change in the GICE interval, or for a contemporaneous eustatic sea level change, in Baltoscandia and in North America. 18
O Geochemistry
Important information about seawater temperature in ancient environments can be obtained by investigations of the δ18O isotope in well-preserved calcareous fossils and carbonate
Upper Ordovician Guttenberg δ13C excursion (GICE) in North America and Baltoscandia sediments. This method is based on the fact that the difference in δ18O/δ16O ratios between the CaCO3 and the seawater from which the calcium carbonate is precipitated is related to the temperature, with a decrease in temperature of the water leading to an increase in the calcite δ18O value. This paleothermometer is widely used in geologically relatively young (Cenozoic and Cretaceous) deposits, and it has often proved to produce valuable, and seemingly reliable, data on past seawater temperatures (Anderson and Arthur, 1983). However, many attempts to use this technique on Paleozoic marine limestones and fossils have been unsuccessful, because the δ18O values have been reset owing to diagenesis (Marshall, 1992) and do not represent the original δ18O signature. The idea that in general, Paleozoic δ18O values are unreliable (see, e.g., Kaljo et al., 2004), if not entirely useless, has resulted from the unfortunate practice of leaving out δ18O values in many 13 C publications. Most δ18O values through the GICE interval in Kentucky are between –4.8‰ and −6‰, in the Hagan (Virginia) and Dolly Ridge (West Virginia) successions between –5.7‰ and –6.5‰, and in the Fittstown (Oklahoma) section between –4.5‰ and −5.2‰ (Young et al., 2005). Shields et al. (2003) reported similar values, based on brachiopods from Missouri. Several authors (e.g., Marshall and Middleton, 1990; Marshall et al., 1997; Shields et al., 2003) recorded δ18O values as high as ~ −3‰, or even ~ −1‰, from Hirnantian brachiopods. Such high values are interpreted to reflect low seawater temperatures during the Hirnantian Gondwana glaciation. The fact that similar high δ18O values have not been recorded from the GICE interval in North America may be interpreted in at least two ways: They may indicate that the GICE was not associated with a significant lowering of the seawater temperature in the tropical region, or that the δ18O values have been diagenetically reset so the original δ18O signature has been lost. Carbonate δ18O values from Estonia recorded by Kaljo et al. (2004), who regarded them as diagenetically reset, range between <–4‰ and >−6‰, with most between –4.5‰ and −5‰, hence values similar to those in North America. An isotopic study of the GICE interval in the Kullsberg Limestone in central Sweden was recently carried out by Tobin et al. (2005). Unlike most other investigations, this study was centered on what appeared to be diagenetically relatively unaltered carbonate cements, which in the upper part of the Kullsberg Limestone produced δ18O values as high as −3.6 ‰ to –2.8 ‰, with one value being –1.8‰. Brachiopods from the lower part of the same formation show δ18O values between –4‰ and –6‰ (Marshall and Middleton, 1990), which are similar to the baseline values in the underlying Dalby Formation. Many additional samples processed during the present study have δ18O values between –4‰ and –6‰, with a few <~4‰. Tobin et al. (2005) interpreted the δ18O values as not having been substantially reset diagenetically, and they suggested that the seawater temperature in this area during the GICE interval was of the order 10–17 °C. In view of the fact that the Kullsberg Limestone was probably deposited at a water depth of at least several tens of meters and at a latitude of ~35°S, this
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is not an unlikely seawater temperature. Recent studies of δ18O in conodont apatite in the GICE interval in Kentucky and Minnesota suggest seawater temperatures of ~25 oC that do not differ notably from those in the stratigraphic interval above and below the GICE (Buggisch et al., 2010). These values, which may have special significance because it is believed that δ18O in conodont apatite has not been affected by diagenesis as is the case in carbonates, suggest that the GICE interval was not associated with a significant seawater temperature lowering in the tropical zone. POSSIBLE CAUSES OF THE GICE Positive excursions in the δ13C of marine carbonate on time scales longer than ~105 yr may indicate an increase in the burial ratio of organic carbon to carbonate carbon, forg, or changes in the isotopic composition of riverine carbon inputs, δw, according to equation 1 (Kump and Arthur, 1999): forg = (δw - δ carb)/ΔB.
(1)
Because the isotopic difference between carbonates and organic matter (ΔB) depends on the photosynthetic fractionation of carbon isotopes, changes in atmospheric pCO2 can also have an impact on δ13Ccarb. Indeed the GICE in carbonates does correspond to a shift in the δ13C of organic matter (Patzkowsky et al., 1997), and this can be interpreted as an expression of a drop in pCO2 values, perhaps even reaching the threshold for ice buildup in the Ordovician, although this is controversial, as discussed in the above sections. If enhanced burial of organic matter in the global oceans produced the GICE, this may be linked to increased productivity, enhanced preservation under anoxic conditions, or high sedimentation rates. Productivity may increase with a greater flux of phosphate from the continents or upwelling from the deep ocean, and this may in turn lead to anoxia and an increase in the carbon-to-phosphorus ratio of buried organic matter (e.g., Schrag et al., 2002). However, a larger oceanic phosphate inventory can only sustain higher production and burial of organic matter under anoxic conditions if nitrogen concentrations keep pace (N-fixation must balance loss from denitrification Saltzman, 2005). The burial of organic matter in modern settings occurs predominantly in fine-grained siliciclastic sediments, such as deltas, where high rates of sedimentation, nutrient fluxes, and suspended loads of terrestrial organic matter are found (Schrag et al., 2002). Significant amounts of organic matter could also have been buried in the deep ocean, which may have been anoxic in the early Paleozoic (Wilde and Berry, 1984). Enhanced preservation of organic matter in the deep ocean also may have played a role in the GICE, linked to increased oceanic stratification during rising sea levels (e.g., Cramer and Saltzman, 2005). At this stage, several different models can be proposed for the cause of the GICE, and some have been discussed in the literature (see, e.g., Young et al., 2005). However, a more complete understanding, especially of the pattern of changes in global sea level,
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during the event is required to further assess its impact on carbonate and organic-carbon weathering rates and marine burial rates. CONCLUSIONS AND SUMMARY The following conclusions emerge from the above review of the Guttenberg isotope carbon excursion (GICE) in North America and Baltoscandia: 1. The GICE is a characteristic feature in a specific stratigraphic interval at the more than 40 localities studied thus far in the North American Midcontinent, the Appalachians, the Great Basin in Nevada, Sweden, Estonia, and Latvia. Biostratigraphy and K-bentonite event stratigraphy indicate strongly that the North American GICE is the same as the “mid-Caradocian” excursion in Baltoscandia, hence justifying the use of the term GICE for the Baltoscandian excursion also. Furthermore, recent studies by Bergström et al. (2009) indicate that this excursion is present also in China, hence supporting the proposal that the excursion has a global distribution. 2. On the basis of conodont, graptolite, and chitinozoan biostratigraphy, the stratigraphic range of the GICE is well established. In North America it is present in the uppermost Phragmodus undatus and in the Plectodina tenuis Midcontinent Conodont Zones, in the uppermost Amorphognathus tvaerensis Atlantic Conodont Zone, and in the lowermost Diplacanthograptus caudatus Graptolite Zone. In Baltoscandia it occurs in the uppermost Amorphognathus tvaerensis Conodont Zone and in the uppermost Diplograptus foliaceus and Dicranograptus clingani Graptolite Zones. In both these continents it ranges through the upper part of the Spinachitina cervicornis Chitinozoan Zone. In North America it occurs in the lower-middle Chatfieldian Stage, and in Baltoscandia, in the upper Keila and Oandu Stages. In North American sections containing the prominent and widespread Millbrig K-bentonite, the base of the GICE lies above this ash bed, and in terms of the eastern Midcontinent sequence stratigraphy, in the lower part of the M5 sequence. In stratigraphically complete, or nearly complete, successions, the beginning of the GICE is from less than one to a few meters above the Millbrig K-bentonite, whereas in sections with a significant stratigraphic gap at the base of the M5 sequence the GICE may begin just above the Millbrig K-bentonite. In stratigraphically virtually complete Baltoscandian sections, the GICE starts a few meters above the Kinnekulle K-bentonite. 3. Depending on the net depositional rate and the stratigraphic completeness of the section, the thickness of the GICE ranges from <1 m to >100 m in North America. In cratonic sections in North America and Baltoscandia, the GICE interval is generally ~5–10 m thick. 4. The excursion δ13C peak value shows some, apparently regional, variation. In the Midcontinent the GICE
5.
6.
7.
8.
reaches peak values of 2‰ to 3‰, in Baltoscandia 1‰ to 2‰, and in central Nevada >3‰ above baseline values. There is also a regional variation in baseline values. In the Upper Mississippi Valley these values are ~ –1‰ to ~0‰, whereas in the eastern Midcontinent they are ~0‰ to +1‰. The relations between the GICE and eustasy are not straightforward. In most North American sections the GICE starts in a transgressive interval and ends in a regressive interval. The bathymetric conditions in the coeval interval in Baltoscandia are somewhat controversial, but at least part of the GICE occurred during a prominent regression at the top of the Keila Stage and during the Oandu Stage. It appears difficult to link the GICE to a significant eustatic regression such as the one associated with a major glaciation. In this respect the GICE seems to differ from the latest Ordovician Hirnantian δ13C excursion (HICE). Lithofacies evidence indicates that the GICE is present in tropical warm-water carbonates (bahamites) in the Midcontinent as well as in subtropical to temperate carbonates in Baltoscandia. The fact that the GICE is associated with warmer-water macro- and microfossil faunas in North America, and with quite different colder-water faunas in Baltoscandia, suggests no obvious direct correlation between the excursion and the water temperature. Geochemical indications from δ18O investigations on carbonate are inconclusive, as most δ18O values probably have been diagenetically reset. However, a recent δ18O study on conodont apatite in the GICE interval in Kentucky and Minnesota indicates no marked lowering of the seawater temperature. The various data at hand do not provide strong support for the idea that the GICE was closely correlated with, and in some way perhaps due to, a climatic deterioration of a similar magnitude as that associated with significant glaciations such as those during Hirnantian time. A preliminary assessment of the minor δ13C excursions recently recognized in the lowermost Chatfieldian and upper Turinian in the Upper Mississippi Valley indicates that these positive shifts in the δ13C curve, several of which are not present in every section even in that region, are not global but regional or local in their distribution. On the other hand, the several positive excursions recognized between the GICE and the HICE in Estonia have been recently identified also in the type Cincinnatian Series in its type area in the Cincinnati region, suggesting that these excursions are of a global nature.
ACKNOWLEDGMENTS We are indebted to T. Ainsaar, G.A. Ludvigsson, and S.C. Finney for careful reviews of the mauscript and useful suggestions for its improvement, and to S.A. Young for valuable
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POSTSCRIPT Since the present paper was finally submitted in 2006, several studies of aspects of the GICE have been published. It is not possible to review the results presented in these papers in any detail, but it is appropriate to list the most important of these publications. In North America, recent GICE investigations include Panchuk et al. (2005, 2006) and Fanton and Holmden (2007), which were based on Upper Mississippi Valley sections. Young et al. (2008) compared δ13Corg and δ13Ccarb GICE curves from exposures in North America and China. Finally, in a study of the relations between several Katian δ13C excursions in North America and Estonia, Bergström et al. (2007) compared the biostratigraphic positions of the GICE in the Cincinnati region standard succession and in an important drill core in Estonia. Postscript References Bergström, S.M., Young, S., Schmitz, B., and Saltzman, M.R., 2007, Upper Ordovician (Katian) δ13C chemostratigraphy: A trans-Atlantic comparison: Acta Palaeontologica Sinica, v. 46, suppl., p. 37–39. Calner, M., Lehnert, O., and Joachimski, M., 2009, Carbonate mud mounds, conglomerates, and sea-level history in the Katian (Upper Ordovician) of central Sweden: Facies, v. 56, p. 157–172, doi: 10.1007/s10347-009 -0192-6. Fanton, K.C., and Holmden, C., 2007, Sea-level forcing of carbon isotope excursions in epeiric seas: Implications for chemostratigraphy: Canadian Journal of Earth Sciences, v. 44, p. 807–818, doi: 10.1139/E06-122. Goldman, D., Leslie, S.A., Nõlvak, J., Young, S., Bergström, S.M., and Huff, W.D., 2007, The Global Stratotype Section and Point (GSSP) for the base of the Katian Stage of the Upper Ordovician Series at Black Knob Ridge, southeastern Oklahoma, USA: Episodes, v. 30, p. 258–270. Kaljo, D., Martma, T., and Saadre, T., 2007, Post-Hunnebergian Ordovician carbon isotope trend in Baltoscandia, its environmental implications and some similarities with that of Nevada: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 245, p. 138–155, doi: 10.1016/j.palaeo.2006.02.020. Panchuk, K.M., Holmden, C., and Kump, L.R., 2005, Sensitivity of the epeiric sea carbon isotope record to local-scale carbon cycle processes: Tales from the Mohawkian Sea: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 228, p. 320–337, doi: 10.1016/j.palaeo.2005.06.019. Panchuk, K.M., Holmden, C., and Leslie, S., 2006, Local controls on carbon cycling in the Midcontinent region of North America with implications for carbon isotope secular curves: Journal of Sedimentary Research, v. 76, p. 200–211, doi: 10.2110/jsr.2006.017. Young, S., Saltzman, M.R., Bergström, S.M., Leslie, S.A., and Chen Xu, 2008, Paired δ13Ccarb and δ13Corg records of Upper Ordovician (Sandbian-Katian) carbonates in North America and China: Implications for paleoceanographic change: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 270, p. 166–178, doi: 10.1016/j.palaeo.2008.09.006. MANUSCRIPT ACCEPTED BY THE SOCIETY 6 OCTOBER 2009
Printed in the USA
The Geological Society of America Special Paper 466 2010
The Ordovician brachiopod radiation: Roles of alpha, beta, and gamma diversity David A.T. Harper† Natural History Museum of Denmark (Geological Museum), University of Copenhagen, Øster Voldgade 5-7, DK-1350 Copenhagen K, Denmark
ABSTRACT The Ordovician Period (ca. 488–444 Ma) witnessed profound changes in the biodiversity and biocomplexity of marine life, marked by the installation of a benthos dominated by suspension-feeding animals, most notably the brachiopods. The Ordovician brachiopod fauna was dominated by rhynchonelliformeans in contrast to that of the underlying Cambrian System, characterized by a diversity of various nonarticulated groups. Over an interval of some 25 m.y., accelerating γ (inter-provincial), β (inter-community), and α (intra-community) diversity was initiated by high diversities among Early Ordovician brachiopod faunas associated with the dispersal of the continents and the high frequency of volcanic arcs and microcontinents (γγ diversity). During the Early and Middle Ordovician, community types expanded particularly into deeper water and around carbonate platforms and structures (β β diversity). Moreover, during the period the α diversity of individual assemblages increased from <10 species during the Late Cambrian to ~30 in the Late Ordovician, with the canalization of ecological niches and the opportunity for more densely packed communities. The end-Ordovician extinction event was a severe crisis for the more common Ordovician taxa, the orthide and strophomenide brachiopods. Whereas some widespread taxa, characteristic of deep-water environments, survived, many from shallower-water, together with those from the deep shelf and slope, disappeared. The subsequent Silurian fauna became increasingly dominated by atrypide, athyridide, spiriferide, and rhynchonellide brachiopods.
INTRODUCTION
Paleozoic biodiversifications in most marine groups were spectacular and sustained, setting the agenda for subsequent marine life on the planet. The majority of metazoan groups appeared first at the base of the Paleozoic, increasing in diversity during the Cambrian Explosion and Ordovician Radiation to establish an ecosystem that survived some 250 m.y. of Earth history. Brachiopods were an important part of the latter process, establishing themselves as key components of a new Ordovician Earth System. The development and evolution of the Paleozoic fauna were, however, modified by a series of extinction events ascribed
Paleozoic oceans were characterized by short trophic chains dominated by suspension-feeding organisms, evolved during a sustained interval of greenhouse climate. This ecosystem contrasted with that of the subsequent Mesozoic and Cenozoic Eras, dominated by deposit-feeding communities linked to more bioturbated substrates and complex community structures. Early †
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Harper, D.A.T., 2010, The Ordovician brachiopod radiation: Roles of alpha, beta, and gamma diversity, in Finney, S.C., and Berry, W.B.N., eds., The Ordovician Earth System: Geological Society of America Special Paper 466, p. 69–83, doi: 10.1130/2010.2466(05). For permission to copy, contact
[email protected]. ©2010 The Geological Society of America. All rights reserved.
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to a variety of causes. The first test of the Paleozoic brachiopod fauna was the end-Ordovician extinction event, removing or greatly reducing many of the commoner groups, the orthides and strophomenides, but promoting the atrypides, athyridides, and spiriferides during its recovery phase (Rong and Harper, 1999). The Ordovician Period underwent unusually high sea levels and intense magmatic and tectonic activity against a background of sustained warm climates. During this period a truly massive rise in marine biodiversity (Sepkoski, 1981), accompanied by an increase in the biocomplexity of marine life (Droser and Sheehan, 1997), marked “The Great Ordovician Biodiversification” as one of the two most significant radiation events in the history of marine life. The phylum Brachiopoda was one of the most diverse and successful benthic groups, evolving a wide range of morphologies (Harper and Wright, 1996) and participating in a spectrum of community types across all the main biogeographic provinces. Biological events associated with the Cambrian Explosion involved the origins of widespread skeletalization, a range of new body plans, and the extinction of the soft-bodied Ediacara fauna, together with the appearance and diversification of the Bilateralia (see Marshall, 2006, and references therein). The Ordovician diversification, however, generated few new higher taxa, for example, phyla, but it exposed a massive increase in rolling biodiversity at the family, genus, and species levels. The radiation included members of the so-called “Cambrian, Paleozoic, and Modern” evolutionary biotas, and a train of ecological
events set the agenda for much of subsequent marine life on the planet (Sheehan, 1996). The brachiopods participated in both evolutionary faunas present during the Ordovician Radiation: The Cambrian (linguliformeans) and the Paleozoic (rhynchonelliformeans), the latter forming the core of the Ordovician diversification (Fig. 1). Many taxa counts are now available through 45 m.y. of the Ordovician Period; there are still relatively few studies of the ecological and environmental aspects and consequences of this diversification (Bottjer et al., 2001). A brief overview of brachiopod ecology through the radiation was provided by Sheehan and Harper (in Harper et al., 2004a). Significantly, the causes of the event, and its relationship to both intrinsic (biological) and extrinsic (environmental) factors, remain far from clear, although it was possibly associated with an increasing abundance of phytoplankton (Vecoli et al., 2005) and planktotrophic larvae (Peterson, 2005), an Ordovician superplume (Barnes, 2004), and the increased flux of asteroids (Schmitz et al., 2008). The unique environmental conditions through the Ordovician Period have been emphasized in a number of publications (e.g., Jaanusson, 1984). Extensive, epicontinental seas developed during sea-level highstands, driven by sustained greenhouse climates, were associated with virtually flat seafloors and restricted land areas, many probably represented only by occasional, emergent archipelagos. Magmatic and tectonic activity was intense and persistent, with rapid plate movements and widespread volcanic
Figure 1. Standing diversities of all the major rhynchonelliformean groups compared and measured as numbers of genera (after Harper et al., 2004a, Fig. 17.11).
The Ordovician brachiopod radiation: Roles of alpha, beta, and gamma diversity activity. Island arcs and mountain belts provided sources for clastic sediment in competition with the carbonate belts associated with the platforms on most of the continents. Biogeographical differentiation was extreme, affecting plankton, nekton, and benthos; and climatic zonation, particularly in the Southern Hemisphere, was pronounced. Provincial differentiation among the Brachiopoda was also marked with biogeographic differences that persisted until near the end of the period (Williams, 1973), when these were disrupted by the end-Ordovician glaciation (Rong and Harper, 1988; Owen et al., 1991). Together these conditions were, nevertheless, clearly ideal for allopatric (geographic) speciation processes together with opportunities for canalization of ecological niches (Harper, 2006). The temporal and spatial framework for the system is well established and relatively precise. A revised international chronostratigraphy of the Ordovician System was recently completed, largely through the work of the International Subcommission on Ordovician Stratigraphy (ISOS). Some of the many unique properties of the system have, in fact, made the development of a global time frame difficult; the intense biogeographical and ecological differentiation of biotas, together with a lack of reliable radiometric data and a lack of international agreement, have hindered the rapid establishment of a modern global chronostratigraphy for the Ordovician System. Nonetheless, the definition of three global series (Lower Ordovician, Middle Ordovician, and Upper Ordovician) and seven global stages (the Tremadocian, Floian, Dapingian, Darriwilian, Sandbian, Katian, and Hirnantian) is now complete (Ogg, 2004; Bergström et al., 2006). More precise and highly resolved international chronostratigraphic units have yet to be defined, although these already exist for many regions in local chronostratigraphic schemes (e.g., Fortey et al., 2000). A number of recent publications have provided a robust and testable template for the changing paleogeography through the period. Fortey and Cocks (2003) provided an up-to-date review of the paleogeography of the period, based largely on faunal evidence, whereas Cocks and Torsvik (2004) described the major terranes of the Ordovician world. For purposes of calibration and description of the Ordovician Radiation, Webby et al. (2004b) devised a sequence of shorter time slices, each corresponding to entire or partial, correlatable graptolite, conodont, and chitinozoan biozones (Fig. 1). These time slices, 19 ranging through 1a–6c, formed the basis for the majority of biodiversity charts in Webby et al. (2004a); these have proven suitably robust and unambiguous. The main radiation commenced during the early Darriwilian (late Arenig) and continued into the late Katian (mid-Ashgill), prior to the endOrdovician extinction event (Sepkoski, 1995; Sheehan, 2001a). This was a time span of some 25 m.y., relatively short in the overall 3.8 Ga of the history of life. CAMBRIAN BRACHIOPOD FAUNA The Cambrian Explosion generated a range of new and spectacular body plans together with the extinction of the soft-
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bodied Ediacara fauna and the appearance of the Bilateralia over a relatively short period of time (Conway Morris, 1998; Briggs and Fortey, 2005). The rapid appearance of such a wide range of apparently morphologically disparate animals has suggested, however, an interval of cryptic evolution of probable micro- and meioscopic organisms, operating beneath the limits of detection prior to the explosion (Cooper and Fortey, 1998). Moreover, molecular-clock estimates initially indicated that animal lineages split some 800 Ma or more before their appearance in the fossil record (Wray et al., 1996), although problems with this approach have been exposed (Peterson et al., 2004). Greater refinement of Cambrian stratigraphy, the taxonomy and phylogeny of key Cambrian taxa, and their relative appearance in the fossil record have suggested an alternative hypothesis: The current, Lower to Middle Cambrian fossil record displays the sequential and orderly appearance, albeit over a short time interval, of successively more complex metazoans (Budd, 2003). Nevertheless it is probable that the so-called superphyla (the Ecdysozoa, Lophotrochozoa, and Deuterostomia) diverged in the Proterozoic, and that the Cambrian Explosion records the orderly diversification of these clades at the base of the Cambrian. The position of the Brachiopoda within the phylogenetic framework of this diversification is still unresolved; most authorities now favor a position within the protostomous Lophotrochozoa (Valentine, 2004). Much of our knowledge of the Cambrian Explosion is derived from three spectacular, intensively studied lagerstätte assemblages: Burgess (Canada), Chengjiang (South China), and Sirius Passet (North Greenland). The diversities of the Cambrian “background” faunas are generally much lower and arguably contain more “normal,” less morphologically disparate organisms. Ordovician exceptionally preserved biotas are rarer; the Soom Shale (South Africa) contains exquisitely preserved faunas but lacks new higher taxa (Aldridge et al., 1994). Thus, whereas the Cambrian Explosion provided higher taxa, in some diversity, the Ordovician Radiation generated the sheer biomass, biodiversity, and biocomplexity that would fill the world’s oceans. Brachiopods are conspicuously rare in these exceptional deposits, but where they occur, some soft parts are spectacularly preserved, suggesting that there were already a number of different lophophore types (Zhang et al., 2003). No demonstrable Precambrian brachiopods have yet been described. The Cambrian Explosion, nevertheless, established a characteristic brachiopod fauna that included both epi- and infaunal ecotypes (Sheehan, 2001b). In general the brachiopods of the Cambrian Evolutionary Fauna were dominated by some shortlived rhynchonelliformeans—for example, the Kutorginata and Obolellata together with widespread and variable assemblages of linguliformeans together with rarer craniiformeans later in the period. Thus all three subphyla—the Linguliformea, Craniiformea, and Rhynchonelliformea—were in place. However, they were by no means diverse and were not, with a few exceptions, significant components of shell concentrations. Their influence on substrate development was probably only local, and the Brachiopoda probably played little part in mediating Cambrian
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ecosystems. Many communities and organisms were highly variable. For example, Late Cambrian brachiopod communities from the outer shelf of the Great Basin (McBride, 1976) lacked consistent and clearly definable community associations, suggesting the development of eurytopic, opportunistic taxa with broad environmental ranges. The Cambrian brachiopod fauna was not particularly diverse ecologically. The majority of taxa were biconvex, epifaunal, pedunculate forms occupying the lowest tiers of the filterfeeding benthos. The brachiopod components of the Cambrian Evolutionary Fauna persisted through the Cambrian Period into the Tremadocian, although assemblages more typical of the Paleozoic Evolutionary Fauna were already present on peri-Gondwanan terranes from the Middle Cambrian onward (Bassett et al., 2002). ORIGIN OF THE “ORDOVICIAN” BRACHIOPOD FAUNA Inevitably, at a taxonomic level, the majority of “Paleozoic” taxa were derived from Cambrian stocks, but the Ordovician Radiation was quite different from the Cambrian Explosion and in many ways more dramatic and interesting (Droser and Finnegan, 2003). In contrast to the Cambrian Explosion, with the exception of the bryozoans, no new higher taxa emerged during the Ordovician Radiation; instead the event witnessed a three- to fourfold increase at the family level, producing clades that would dominate marine life for the next 200 m.y. In particular, the orthide and strophomenide brachiopods, rare in Cambrian faunas, numerically and taxonomically dominated the Ordovician seafloor. Whereas the orthides formed a monophyletic group, the integrity of the strophomenides is less certain; the two main superfamilies, the plectambonitoids and strophomenoids, may have originated separately, and the plectambonitoids are possibly polyphyletic. Two key Cambrian brachiopod assemblages had a particular influence on the subsequent development of the Ordovician brachiopod fauna. The Protorthide Fauna, consisting of primitive orthides, protorthides, and some nonarticulates, together with the Billingsella Association, contained many of the stem groups for the Ordovician diversification (Bassett et al., 2002), the former containing stem groups for the clitambonitoids and pentamerides, whereas the latter included the roots of the strophomenides, orthotetides, and polytoechioids. Orthides were already present in Cambrian faunas, with the orthidines (impunctate) consistently present, if rare; possibly the dalmanellidine (punctate) superfamilies were also represented, as advanced endopunctate forms were described from the Tremadocian rocks of Argentina (Harper et al., 2004b). Several transitional associations of brachiopods, the Clarkella Fauna, various syntrophinidine assemblages, and Tritoechia-Protambonites associations (Bassett et al., 2002) helped form the interface between the two evolutionary faunas, although their direct links to the typical Paleozoic brachiopod fauna are unclear. In many parts of the world the ini-
tial stages of the Ordovician were marked by a dominance of orthide-pentameride assemblages, the majority represented by biconvex taxa. Thus the appearance and diversification of the “Ordovician” brachiopod fauna were strongly linked to facies and paleogeography. The typical compositions and structures of the Paleozoic evolutionary fauna were probably already in place during the endFloian (mid-Arenig) on Baltica (Bassett et al., 2002), and at the same time around Gondwana but slightly later around Laurentia and Siberia, where elements more characteristic of the Cambrian fauna were retained in nearshore, warm-water carbonate facies. BRACHIOPOD TAXONOMIC DIVERSIFICATION For nearly 100 years relatively little attention was paid to global diversity curves for the Brachiopoda and most other organisms through time. However, a number of initial surveys of the Brachiopoda (Cooper and Williams, 1952; Newell, 1952; Williams, 1957) indicated a steep diversification during the Early and Middle Ordovician (as then defined). In a detailed, quantitative investigation of the variation in diversity of higher taxa of marine invertebrates, Valentine (1969) identified a number of clear trends. During the Early and Middle Ordovician there was a steep rise in numbers of, first, classes, and then orders; the number of phyla remained stable at eight during these diversifications. During the same interval there were a marked number of appearances among the classes and orders. More recent compilations by specialists (e.g., Grant, 1980; Harper et al., 1993; Harper et al., 2004a) continued to emphasize the Ordovician spike with increasing amplitudes and increasing precision. In a series of benchmark papers, Sepkoski (1978, 1979, 1981, 1984, 1991; see also 1997) established without doubt the statistical reality of the Ordovician diversification on the basis of database compilations mainly from the Treatise; multivariate (principal components) analyses highlighted the dominance of groups of suspension-feeding organisms through the post-Cambrian Paleozoic interval. This theme was developed in detail by Droser et al. (1997) and together with Droser and Sheehan (1997) as the “Great Ordovician Biodiversification,” and three focal points of the event were emphasized: taxonomic diversity, morphological disparity, and ecological change. Miller and his colleagues have additionally dissected the various aspects of the event based on the acquisition and management of a series of extensive data sets (Miller, 1997a, 1997b, 1997c; Miller and Connolly, 2001; Connolly and Miller, 2002; Miller and Foote, 1996; Miller and Mao, 1995; Novack-Gottshall and Miller, 2003a, 2003b) developed to test particular distributional models. All these key studies helped drive publication of a comprehensive review of all the main taxonomic aspects of the event together with its environmental and geographic setting (Webby et al., 2004a). The Ordovician Radiation, in taxonomic terms, marks the greatest and most sustained interval of diversification of life on the planet, involving at least 4600 genera. A clear and consistent signal emerges, one of sustained and cascading diversification
The Ordovician brachiopod radiation: Roles of alpha, beta, and gamma diversity throughout the period. The radiation was, however, selective; some taxa diversified more than others, some earlier and some later, and some environments and regions showed preferential diversifications across selected higher taxa. This is particularly obvious at the phylum level and is clearly demonstrated in the Brachiopoda (Harper et al., 2004a). The majority of the diversity trajectories are complex, multipeaked systems, showing stepwise increases throughout the period, suggesting that the global signal is a complex time series curve. For example, the diversity patterns and trends of the nonarticulated and articulated brachiopods are, not surprisingly, quite different (Harper et al., 2004a). For example, within the rhynchonelliformean brachiopods three clear peaks are discernible for the typical “modern” articulated forms, the Orthida (Fig. 2A), with the group expanding in the mid-Arenig and fluctuating throughout the rest of the period; the Darriwilian (late Arenig–early Llanvirn), late Sandbian (midCaradoc), and late Katian (mid-Ashgill) peaks, corresponding to an initial association with the disparate continental and microcontinental configuration of that time, and a subsequent move into deeper-water environments; and finally a diversification in carbonate buildups. The pattern for the Strophomenida is quite different (Fig. 2B); the group expanded first in the Dapingian (midArenig) but did not peak until the early Katian (mid-Caradoc), with a less marked late Katian (mid-Ashgill) spike. The patterns of the other, more minor groups—the atrypides, pentamerides, and rhynchonellides—differ in detail, radiating later in the lower Darriwilian (upper Arenig), with maximum levels in the later Katian (Ashgill); the late Katian (mid-Ashgill) diversifications may have been associated with carbonate environments during the later Ordovician (Fig. 2C). The last three groups, in particular, dominated the Silurian benthos following the end-Ordovician extinction event (Harper and Rong, 2001), when carbonate facies were more prevalent. Morphological disparity, in contrast to diversity, addresses the actual morphological variation among groups of taxa. Both are thus quite different types of measurements: Morphologically disparate faunas may be of relatively low diversity, but the morphological distances between individual taxa may be large. Morphological disparity has been calculated from cladograms and dendrograms and by ordination techniques (Wills, 2001). High-profile studies of morphological disparity on Cambrian Explosion taxa have established that present-day disparity among certain groups, such as the arthropods (Wills et al., 1994) and priapulid worms (Wills, 1998), may be equal to or even exceed Cambrian levels of disparity. During the Ordovician Radiation, however, trilobite morphological disparity apparently accelerated after their peak diversity during the Early Ordovician (Foote, 1991, 1993), probably when a range of new ecomorphotypes arrived during the later Ordovician (Fortey and Owens, 1990). These new morphologies may not have been the basis for a huge taxonomic diversification, many groups having relatively few families and genera. In some ways the pattern suggests a similarity to the Cambrian Explosion, when a range of new body plans, and life modes, was not necessarily asso-
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ciated with a major hike in diversity but rather an increase in disparity values. In contrast, however, the orthide brachiopods reached both peak disparity and diversity during the Middle–Late Ordovician (Harper and Gallagher, 2001), with the orthidines (impunctates) retaining the lead over the more finely divided dalmanellidine taxa. Significantly, another sessile benthic, filter-feeding group, the crinoids, exhibited the same pattern (Foote, 1995). Diversity and morphological disparity were clearly decoupled in some groups but not in others. Disparity studies of the more morphologically diverse strophomenide clade containing the contrasting plectambonitoids and strophomenoids may show a quite different pattern. The apparent decoupling of diversity and disparity in some groups but not others is significant and requires investigation. BRACHIOPOD ECOLOGICAL DIVERSIFICATION The profound changes in biodiversity and morphological disparity were matched by dramatic changes in the planet’s marine ecosystems from the apparently less well-organized associations of the Cambrian evolutionary faunas to the more structured paleocommunities and diverse assemblages of the filterfeeding Paleozoic evolutionary faunas (Sheehan, 1991, 1996). Ordovician Earth systems were probably quite different from their Cambrian counterparts. Moreover these changes were neither necessarily coincident nor varied in the same way. There is strong evidence to suggest that the many Phanerozoic ecological events were decoupled from major changes in diversity, implying that the controls and constraints on diversity trends and ecological structures were not necessarily similar. Ecological structures can be strongly influenced by the abundance or disappearance of keystone species rather than by merely increasing or declining biodiversity (McGhee et al., 2004). For example, intense biological activity that created thick shell concentrations, hardgrounds, bioturbation, and epifaunal tiers may have been limited to a few critical species rather than to a hike in diversity; nonetheless their effects on an ecosystem can be profound. In this way a number of brachiopod taxa developed an exceptional abundance and may have helped mediate local ecosystems. The two most significant morphological changes apparent in Ordovician brachiopod assemblages were the development of planar to concavoconvex shells (Harper et al., 2004a) and cyrtomatodont dentitions (Bassett et al., 1999). Both had important implications for the life styles of the Brachiopoda (Sheehan and Harper in Harper et al., 2004a) and helped these groups develop new life strategies in a variety of new niches. A hierarchy of severity of paleoecological events has been developed for Phanerozoic marine life (Droser et al., 1997; Bottjer et al., 2001). The events or levels range from first level (appearance-disappearance of an ecosystem) to fourth level (appearances-disappearances of paleocommunities). These types of changes had a profound effect on the composition and structure of Phanerozoic life. Bottjer et al. (2001) emphasized the reality of second, third, and fourth level changes throughout
A
B
Figure 2. Standing diversity of (A) the two orthide suborders; (B) the two strophomenide superfamilies; and (C) the atrypide, pentameride, and rhynchonellide orders. Since submission of the manuscript, the third stage of the Ordovician is now named the Dapingian.
C
The Ordovician brachiopod radiation: Roles of alpha, beta, and gamma diversity the Ordovician Radiation. For example, changes in dominants, such as transitions from trilobite to brachiopod-dominated communities on soft substrates and echinoderms to bryozoans on hard substrates, the evolution of deep-mobile burrowers, and stromatoporoid reefs are cited in support of second level changes. The appearance of new Bambachian megaguilds (groups of organisms with mutually similar adaptive strategies) and the supplementation of existing megaguilds with new taxa together with the arrival of new community types (receptaculitid-macluritiddominated, orthid-dominated, and bivalve-trilobite combinations) together with a marked increase in tiering complexity signal third level changes, whereas the development of many new paleocommunities is the hallmark of fourth level changes. Key ecological changes, modified from Bottjer et al. (2001), are listed in Table 1 and are associated where possible with ecological events within the Brachiopoda, which mainly involved third level changes. For this time, five key themes have been emphasized by a number of authors: the increase in numbers of Bambachian megaguilds, tiering complexity (both above and beneath the sedimentwater interface), changes in the types of shell concentrations, development of hardgrounds and hardground communities, and changing composition and structure of carbonate buildups and their associated communities (Harper, 2006). Can these be linked to the radiations of the Brachiopoda? First, during the radiation there was a marked increase in the number of megaguilds (Bambach, 1983). In addition to those already established during the Cambrian, four new megaguilds appeared while a marked increase in numbers of sessile epifaunal (suspension-feeders: attached low, attached erect, and reclining) and mobile epifaunal (detritivores, herbivores, and carnivores) components occurred. These ecologically organized animals either filled hitherto unoccupied niches or subdivided existing
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niches (Sheehan, 2001b). Brachiopods participated in all three suspension-feeding megaguilds, where particularly the orthides (attached low and erect) and strophomenides (reclining) dominated. At lower ecological levels the radiations generated at least seven guilds (Bambach 1983; Sheehan and Harper in Harper et al., 2004a): (1) cap-shaped “nonarticulates”; (2) infaunal “nonarticulates”; (3) strongly biconvex pedunculate; (4) strongly biconvex alate pedunculate; (5) planoconvex to weakly biconvex pedunculate; (6) relatively flat free-lying; and (7) inflated freelying articulates. There was also a range of minor guilds such as cemented forms, those attached to bryozoans and crinoids in higher epifaunal tiers, and possibly cryptic faunas and minute nonarticulated taxa that may have been part of the meiofauna. Such changes were also associated with important changes in tiering complexity both above and below the sediment-water interface (Bambach, 1983; Bottjer and Ausich, 1986). It is probable that as brachiopod shell dimensions increased during the event (Sheehan and Harper in Harper et al., 2004a), brachiopods occupied higher tiers by virtue of their size. They also, however, occupied the higher tiers attached to both bryozoans (Harper and Pickerill, 1996) and crinoids (Sandy, 1996). Although brachiopod shell concentrations are known from the Cambrian, commonly dominated by billingselloids and protorthoids (Bassett et al., 2002), there was clearly a major shift in the composition of Ordovician shell concentrations at the base of the Whiterockian in North America (Droser and Sheehan, 1997; Li and Droser, 1999). These shell beds are dominated by orthide brachiopods, such as Anomalorthis, Desmorthis, Hesperonomiella, Hesperonomia, Shoshonorthis, and Orthidiella. The simplified morphology of variably biconvex, ribbed, pedunculate shells strongly influenced both biological production and the environment in the carbonate ramp settings of California,
TABLE 1. THE FOUR PALEOECOLOGICAL LEVELS AND THEIR CHARACTERISTIC SIGNALS TOGETHER WITH ASSOCIATED EVENTS DURING THE ORDOVICIAN RADIATION Level Description Signals Ordovician Radiation First Appearance-disappearance Initial colonization of an environment 1. Deep-water benthos: Foliomena of an ecosystem fauna 1. Brachiopod takeover in Second Structural changes within an 1. First appearance of, or changes sessile benthos ecosystem in, ecological dominants of 2. Development of “strophomenide” higher taxa Bambachian megaguild 2. Loss-appearance of metazoan 3. Increasing levels of predation and reefs 3. Appearance-disappearance of brachiopod responses Bambachian megaguilds 1. Appearance and/or disappearance 1. New community types associated Third Community level changes of community types with deep-water and metazoan reef within an established 2. Increase and/or decrease in environments ecological structure tiering complexity 2. Marked increase in epifaunal tiering 3. “Filling-in” or “thinning” within and bioturbation 3. Increase in the membership of Bambachian megaguilds Bambachian megaguilds Fourth Community-level changes 1. Appearance and/or disappearance 1. Extensive development of new of paleocommunities brachiopod-dominated communities 2. Taxonomic changes within a clade 2. Taxonomic diversity changes at many levels 3. Size change in brachiopod clades Note: Modified from Harper (2006, Table 1).
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Nevada, and Utah. But there is a strong environmental and geographic component to even the development of shell concentrations. Coeval but nearshore areas of the Laurentian continent, such as NE Greenland, were in fact dominated by elements of the Modern Evolutionary Fauna, with shell concentrations dominated by gastropods. Within the Brachiopoda there is a similar biogeographical component. The orthide shell concentrations, so typical of the Great Basin, were complemented by local concentrations of clitambonitoids and porambonitoids around Baltica. The widespread development of hardgrounds during the Ordovician was related to the extensive and pervasive precipitation of low-magnesium calcite on shallow-water marine seafloors (Palmer and Wilson, 2004). Dissolution of aragonite may have been the source of the calcite cement. The development of such hardgrounds, providing a hard substrate for attachment together with surfaces for cementing and boring organisms, opened up new opportunities for a specialized, yet widespread, set of communities. Aragonite dissolution and calcite precipitation (both organic and inorganic) possibly directly influenced evolution of biomineralization and skeletalization, aiding the development of the calcitic rhynchonelliforms and providing substrates for many pedunculate (orthide) and possibly co-supportive (pentameride) forms. The sudden appearance of the Paleozoic brachiopod during the mid-Arenig around Baltica has been linked to the onset of sustained carbonate deposition together with the development of hardgrounds (Bassett et al., 2002). The composition and structure of reefs and other carbonate buildups changed during the Ordovician Radiation (Webby, 2002). Critical was the transition from microbial to metazoandominated structures at and around the Middle Ordovician–Upper Ordovician boundary. The change from buildups associated with stromatolites, thrombolites, algal mats, and various microbes, together with some lithistid sponges and stromatoporoids, to the more familiar structures dominated by tabulate corals, stromatoporoids, lithistid sponges, echinoderms, and solenoporan algae, probably occurred during the Darriwilian. These metazoan buildups were a focus for much biodiversity, creating a plethora of niches around the buildup and its associated marginal facies. DIACHRONISM Early studies of regional patterns (e.g., Harper, 1986) isolated global trends, suggesting that regional studies could be used as proxies for more global patterns. However, contrasts in the diversity trajectories for brachiopods are now apparent between some paleocontinents (Harper and Mac Niocaill, 2002), those for marginal Gondwana being quite different (where the diversification was apparently delayed) from those of Baltica (Harper and Hints, 2001; Hints and Harper, 2003). Moreover, even across individual brachiopod clades there are clear partitions (Harper et al., 2004a). Within the rhynchonelliformeans, many of the Early Ordovician distinctive orthide groups were associated with the margins of Gondwana; the clitambonitoids were associated
with Baltica, whereas the pentamerides and possibly the plectambonitoids preferred the carbonate environments of Laurentia and its margins. Data from South China suggest that the radiation was initially dominated by orthide-pentameride (particularly the syntrophiids) associations during the Floian, with more typical Ordovician communities developing during the Darriwilian (Zhan and Harper, 2006). Thus the current data available suggest that in contrast to the Cambrian Explosion, the timing of changes in Ordovician biodiversity and biocomplexity was diachronous across the various taxonomic groups, environments, and regions (Webby et al., 2004a), although comparable information is arguably lacking for the Cambrian. Deconstruction of the global signal into more regional patterns has emphasized the spatial contrasts in the intensity and timing of the events within the Brachiopoda. Moreover it is clear that particular groups favored particular provinces. This implies strong extrinsic control on the amplitude and shape of the radiation. Regarding individual phyla, most marked is the strong Early Ordovician bivalve radiation around Gondwana, associated with siliciclastic environments (Cope and Babin, 1999); the group diversified outside Gondwana, later during the Middle and Late Ordovician. In contrast, gastropods were more typical of lowlatitude carbonate environments in contrast to the high-latitude, siliciclastic settings of the Ordovician bivalve fauna (NovackGottshall and Miller, 2003a, 2003b). The interdigitation of bivalve and gastropod faunas, with respect to siliciclastic versus carbonate environments, is also exemplified on a regional scale in the type Cincinnatian (Miller, 2004). Moreover bivalve-dominated faunas appear sporadically in the Ordovician successions of Avalonia, where predominantly siliciclastic environments were dominated by brachiopods. Generally brachiopods are not obvious associates of such molluscan-dominated assemblages, a relationship that continued more extravagantly after the end-Permian extinction event. GENERATION OF BRACHIOPOD DIVERSITY Global biodiversity can be resolved into three components associated with specific areas: α-diversity (intra-community), β-diversity (inter-community), and γ-diversity (inter-province) (Sepkoski, 1988), a simple concept imported from many biological studies (see, e.g., Whittaker et al., 2001). This simple dissociation of the three key building blocks of global biodiversity, however, highlights the geographic, environmental, and communitylevel factors that underpin any changes in the diversity of life. Studies in the late 1960s and early 1970s had already addressed the biogeographic (e.g., Williams 1969, 1973), environmental (e.g., Bretsky, 1969), and community (e.g., Williams 1963, 1974) dimensions of Ordovician brachiopod diversity. And arguably some of these issues had already been exposed in the 1930s. In simple terms, the radiation can be modeled in terms of its strong association with the dispersal of the Early Ordovician continents and terranes (γ-diversity), the occupation of new ecospace
The Ordovician brachiopod radiation: Roles of alpha, beta, and gamma diversity (β-diversity), and the closer packing of species within existing communities (α-diversity). Sepkoski (1988) noted that increased α- and β-diversity only accounted for ~50% of the observed genus-level diversity; thus the surplus may be hidden within the γ-diversity and diversities associated with hardground and reef communities not originally investigated by Sepkoski (Droser and Finnegan, 2003). Nevertheless, a model involving the sequential and overlapping increase in these three component diversities can elegantly explain many aspects of the event (Harper and Mac Niocaill, 2002) and provide a number of testable hypotheses. First, the correspondence of the Early Ordovician brachiopod diversification with continental dispersion, together with magmatic and tectonic events, is obvious (Neuman, 1984). The Early Ordovician continental and terrane configuration in the Southern Hemisphere was dispersed (Fortey and Cocks, 2003), whereas virtually nothing is known about continental assembly in the Northern Hemisphere (Fig. 3). Neuman (1964, 1972) first noted the importance of Early Ordovician island arcs as centers for endemism and migration, with such arcs developing their own biogeographically distinct units, quite separate from the platform provinces (Neuman, 1984; Neuman and Harper, 1992). For example, the low-latitude Toquima–Table Head province was quite separate from the Celtic province (Harper et al., 1996). These units were characterized not only by their own endemics but also by mixtures of taxa from adjacent continental provinces. These marginal and ocean terranes in the greater Iapetus region probably acted both as cradles and museums, the cradles providing taxa to drive subsequent radiations on adjacent platforms,
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and the museums by holding refugia for otherwise extinct taxa (Harper and Mac Niocaill, 2002). Similar processes were in force at the same time in the central Asian terranes (Holmer et al., 2000, 2001) and later in the Ordovician in eastern Australia (Webby et al., 1997). Global statistical analyses of the Early Ordovician brachiopod faunas demonstrate a range of disparate faunas that can, however, be grouped into continental and marginal provinces. Endemism was widespread, particularly within the Celtic group of faunas (Harper and Mac Niocaill, 2002). Moreover, Miller and Mao (1995) and Botting (2002) directly related tectonism and volcanism to diversification within the context of the Ordovician Radiation. Significantly γ-diversity was not, apparently, a strong component of Cambrian diversifications and may have been a phenomenon of the Paleozoic and Modern faunas. The onshore-offshore expansion of the Ordovician marine fauna is well established (Sepkoski and Sheehan, 1983) and probably occurred as a stepwise process (Sheehan, 2001b). It can be demonstrated in detail at regional (Bassett et al., 2002) and local (Mergl, 1999; Hansen and Harper, 2008) levels among brachiopod faunas. This expansion exploited vacant ecospace or habitats occupied by relict elements of the Cambrian fauna, creating a range of new communities that culminated during the Late Ordovician, with the well-established deep-water brachiopod Foliomena fauna (Harper et al., 1999; Rong et al. 1999) and the trilobite cyclopygid fauna occupying Benthic Assemblage Zone 5 and beyond. The velocity of such onshore-offshore expansion was probably rapid, with the transitions, in deeperwater environments, from the Cambrian fauna to the Paleozoic
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Figure 3. Terrane development during (A) the Floian to (B) the Mid-Darriwilian interval; obvious are the many marginal, dispersed Gondwanan terranes that together form the Celtic group of faunas (after Harper and Mac Niocaill, 2002, Fig. 2). Am—Alpine Massifs; Arm—Armorica; ATA—Armorican terrane assemblage; Boh—Bohemia (Perunica); DA—Dashwoods Block; Ib—Iberia; Høl.—Hølonda; M—Meguma.
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fauna commencing in the Floian and finishing in the Katian (Fig. 4). There were regional variations, faster around Gondwana and slower around Laurentia (Bassett et al., 2002). Such deepwater faunas attained a relative stability and longevity, with, for example, the Dicoelosia-type community ranging from the Late Ordovician to the Early Devonian (Watkins et al., 2000). These new community types potentially generated β-diversity, although some authors have challenged this (e.g., Miller and Mao, 1998), pointing out that the marked environmental and geographic-range extension of taxa during the later Ordovician may have buffered this effect. Nevertheless, relatively few members of the Foliomena fauna are found in shallow-water environments. However, it was not only the offshore movement of taxa that created new communities; both microbial and metazoan buildups provided a wide range of ecological opportunities for taxa. For example, during the Middle Ordovician, carbonate buildups provided the focus for the origin and diversification of a range of rhynchonelliformean brachiopod taxa (Bassett et al., 1999). In this context the late Katian (mid-Ashgill) global warming event (Boucot et al., 2003) possibly created a range of new carbonate environments, for example, the Boda (Swe-
den), Keisley (England), and Kildare (Ireland) mud-mounds, which provided a range of new opportunities for the diversifying Ordovician brachiopod fauna. The Boda event generated many endemic taxa and prompted major migrations of taxa out of the tropics to higher latitudes and deeper-water habitats (Fortey and Cocks, 2005). During the radiation it is clear that communities became more tightly packed; α-diversity increased. For example the majority of Late Cambrian brachiopod-dominated assemblages contained <10 taxa, whereas by the Late Ordovician such assemblages contained ~30 genera. On a regional scale, Lockley (1983) provided a review of the brachiopod-dominated paleocommunitites of the Anglo-Welsh basin through time. The paleocommunities developed in progressively deeper water through time, but individual associations attained higher diversities through the period. A possible explanation is the increasing canalization within community structures (Valentine, 1969). Increased specialization was matched with narrower niches in communities undergoing greater competition and interaction among their taxa. An alternative approach (Waisfeld et al., 2003) addresses the expansion of guilds within community structures. For example, 11 separate
Figure 4. Succession and environmental distribution of key Cambrian–Early Ordovician rhynchonelliformean brachiopod assemblages across different paleoplates (after Bassett et al., 2002, Fig. 2).
The Ordovician brachiopod radiation: Roles of alpha, beta, and gamma diversity brachiopod guilds are recognized across the three main Early Ordovician basins in Argentina; each guild served to further subdivide ecological space, providing additional opportunities for diversification (Fig. 5). Ecospace was partitioned to allow competition within guilds rather than between guilds. LONG-TERM TRENDS IN THE PALEOZOIC BRACHIOPOD FAUNA The many facets of this complex diversification event are now beginning to be mapped and understood by the disassembly of the strong global signal into ecological, taxonomic, and regional components (Miller, 2004). The ecological and taxonomic amplitudes of the biodiversification may indeed be decoupled, and there are important feedback loops in the process (Fig. 6). The hike in biodiversity and marked change in biocomplexity significantly changed the planet’s seafloors and provided a new agenda
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for much of subsequent marine life (Harper, 2006). Ordovician benthic systems were clearly quite different from those of the Cambrian. Within this frame of diversification the main players and their mutual relationships changed through time, particularly within the Brachiopoda. The Modern brachiopod fauna, consisting of relatively low-diversity assemblages of rhynchonellides and terebratulides, scattered around the world, contrasts strongly with the abundance, activity, and diversity of the Ordovician phases of the Paleozoic evolutionary fauna. In general terms the Phanerozoic brachiopod advanced through the organophosphatic, nonarticulate brachiopod biofacies of the Cambrian through the deltidiodont (orthide-strophomenide) faunas of the Ordovician (Fig. 7), the cyrtomatodont (atrypide-athyridide-rhynchonellidespiriferide) faunas of the mid-Paleozoic to the highly specialized cyrtomatodont and super-recumbent (productide) faunas of the late Paleozoic (Harper and Rong, 2001). The move of the phylum into deeper-water and possibly more cryptic habitats was already
Figure 5. Rhynchonelliformean guilds recognized in the Ordovician basins of Argentina (after Waisfeld et al., 2003, Fig. 4).
Initial graptolite radiations Diversifications of trilobites and nonarticulated brachiopods
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Figure 6. Summary of key ecological and taxonomic events through the Ordovician diversification (modified from Harper, 2006, Fig. 10).
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Figure 7. Brachiopod diversity through time, displayed as the four main rhynchonelliformean “morphological” groups (after Harper and Rong, 2001, Fig. 2). Simple deltidiodonts—Billingsellida + Protorthida + Orthida + Strophomenida + Chonetida + Orthotetida + Productida; complex deltidiodonts—Pentamerida; spire bearers—Athyridida + Atrypida + Spiriferida; crura-loops—Rhynchonellida + Thecideida + Terebratulida.
The Ordovician brachiopod radiation: Roles of alpha, beta, and gamma diversity established during the Ordovician. The longevity of the phylum was secured, however, during the Ordovician Radiation with the development of the cyrtomatodont dentition (Fig. 7); and these taxa, and not the deltidionts, thrived in the carbonate environments that developed after the end-Ordovician extinction event (Harper and Rong, 2001). ACKNOWLEDGMENTS I thank the Carlsberg Foundation (Denmark) for financial support. This paper benefited from constructive reviews by Leonid Popov and Rong Jiayu; Stan Finney added important editorial comments. This publication is a contribution to IGCP 503 “Ordovician Palaeogeography and Palaeoclimate.” REFERENCES CITED Aldridge, R.J., Theron, J.N., and Gabbott, S.E., 1994, The Soom Shale: A unique Ordovician fossil horizon in South Africa: Geology Today, v. 10, p. 218–221, doi: 10.1111/j.1365-2451.1994.tb00993.x. Bambach, R.K., 1983, Ecospace utilization and guilds in marine communities through the Phanerozoic, in Tevesz, M.J.S., and McCall, P.L., eds., Biotic Interactions in Recent and Fossil Benthic Communities: New York, Plenum Press, p. 719–746. Barnes, C.R., 2004, Was there an Ordovician Superplume Event? in Webby, B.D., Paris, F., Droser, M.L., and Percival, I.G., eds., The Great Ordovician Biodiversification: New York, Columbia University Press, p. 77–80. Bassett, M.G., Popov, L.E., and Sokirin, E.V., 1999, Patterns of diversification in Ordovician cyrtomatodont rhynchonellate brachiopods: Acta Universitatis Carolinae: Geologica, v. 43, p. 329–332. Bassett, M.G., Popov, L.E., and Holmer, L.E., 2002, Brachiopods: CambrianTremadoc precursors to Ordovician radiation events, in Crame, J.A., and Owen, A.W., eds., Paleobiogeography and biodiversity change: The Ordovician and Mesozoic-Cenozoic Radiations: Geological Society [London] Special Publication 194, p. 13–23. Bergström, S.M., and Finney, S.C., Chen Xu, Goldman, D., and Leslie, S.A., 2006, Three new global stage names: Lethaia, v. 39, p. 287–288. Botting, J., 2002, The role of pyroclastic volcanism in Ordovician diversification, in Crame, J.A., and Owen, A.W., eds., Palaeobiogeography and Biodiversity Change: The Ordovician and Mesozoic-Cenozoic Radiations: Geological Society [London] Special Publication 194, p. 99–113. Bottjer, D.J., and Ausich, W.I., 1986, Phanerozoic development of tiering in soft substrata suspension-feeding communities: Paleobiology, v. 12, p. 400–420. Bottjer, D.J., Droser, M.L., Sheehan, P.M., and McGhee, G.R., Jr., 2001, The ecological architecture of major events in the Phanerozoic history of marine life, in Allmon, W.D., and Bottjer, D.J., eds., Evolutionary Paleoecology: The Ecological Context of Macroevolutionary Change: New York, Columbia University Press, p. 35–61. Boucot, A.J., Rong Jia-yu, Chen Xu, and Scotese, C.R., 2003, Pre-Hirnantian Ashgill climatically warm event in the Mediterranean region: Lethaia, v. 36, p. 119–132. Bretsky, P.J., 1969, Upper Ordovician ecology of the central Appalachians: Peabody Museum of Natural History Bulletin, v. 34, 150 p. Briggs, D.E.G., and Fortey, R.A., 2005, Wonderful strife: Systematics, stem groups, and the phylogenetic signal of the Cambrian radiation: Paleobiology, v. 31, Supplement, p. 94–112, doi: 10.1666/0094-8373(2005)031 [0094:WSSSGA]2.0.CO;2. Budd, G.E., 2003, The Cambrian fossil record and the origin of the phyla: Integrative and Comparative Biology, v. 43, p. 157–165, doi: 10.1093/icb/ 43.1.157. Cocks, L.R.M., and Torsvik, T.H., 2004, Major terranes in the Ordovician, in Webby, B.D., Paris, F., Droser, M.L., and Percival, I.G., eds., The Great Ordovician Biodiversification: New York, Columbia University Press, p. 61–67.
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Connolly, S.R., and Miller, A.I., 2002, Global Ordovician faunal transitions in the marine benthos: Ultimate causes: Paleobiology, v. 28, p. 26–40, doi: 10.1666/0094-8373(2002)028<0026:GOFTIT>2.0.CO;2. Conway Morris, S., 1998, The evolution of diversity in ancient ecosystems: A review: Philosophical Transactions of the Royal Society of London, v. B353, p. 327–345, doi: 10.1098/rstb.1998.0213. Cooper, A., and Fortey, R.A., 1998, Evolutionary explosions and the phylogenetic fuse: Trends in Ecology & Evolution, v. 13, p. 151–156, doi: 10.1016/S0169-5347(97)01277-9. Cooper, G.A., and Williams, A., 1952, Significance of the stratigraphic distribution of brachiopods: Journal of Paleontology, v. 26, p. 326–337. Cope, J.C.W., and Babin, C., 1999, Diversification of bivalves in the Ordovician: Geobios, v. 32, p. 175–185, doi: 10.1016/S0016-6995(99)80029-1. Droser, M.L., and Finnegan, S., 2003, The Ordovician Radiation: A follow-up to the Cambrian Explosion: Integrative and Comparative Biology, v. 43, p. 178–184, doi: 10.1093/icb/43.1.178. Droser, M.L., and Sheehan, P.M., 1997, Palaeoecology of the Ordovician Radiation; resolution of large-scale patterns with individual clade histories, palaeogeography and environments: Geobios, v. 30, p. 221–229, doi: 10.1016/S0016-6995(97)80027-7. Droser, M.L., Bottjer, D.J., and Sheehan, P.M., 1997, Evaluating the ecological architecture of major events in the Phanerozoic history of marine invertebrate life: Geology, v. 25, p. 167–170, doi: 10.1130/0091-7613(1997)025 <0167:ETEAOM>2.3.CO;2. Foote, M., 1991, Morphological patterns of diversification: Examples from trilobites: Palaeontology, v. 34, p. 461–485. Foote, M., 1993, Discordance and concordance between morphological and taxonomic diversity: Paleobiology, v. 19, p. 185–204. Foote, M., 1995, Morphological diversification of Palaeozoic crinoids: Paleobiology, v. 21, p. 273–299. Fortey, R.A., and Cocks, L.R.M., 2003, Palaeontological evidence bearing on global Ordovician–Silurian continental reconstructions: Earth-Science Reviews, v. 61, p. 245–307, doi: 10.1016/S0012-8252(02)00115-0. Fortey, R.A., and Cocks, L.R.M., 2005, Late Ordovician global warming—The Boda event: Geology, v. 33, p. 405–408, doi: 10.1130/G21180.1. Fortey, R.A., and Owens, R.M., 1990, Trilobites, in McNamara, K.J., ed., Evolutionary Trends: London, Belhaven Press, p. 121–142. Fortey, R.A., Harper, D.A.T., Ingham, J.K., Owen, A.W., Parkes, M.A., Rushton, A.W.A., and Woodcock, N.H., 2000, A revised correlation of the Ordovician rocks in the British Isles: Geological Society [London] Special Report 24, p. 1–83. Grant, R.E., 1980, The human-face of the brachiopod: Journal of Paleontology, v. 54, p. 499–507. Hansen, J., and Harper, D.A.T., 2008, The late Sandbian-earliest Katian (Ordovician) brachiopod immigration and its influence on the brachiopod fauna in the Oslo Region, Norway: Lethaia, v. 41, p. 25–35, doi: 10.1111/j.1502 -3931.2007.00038.x. Harper, D.A.T., 1986, Distributional trends within the Ordovician brachiopod faunas of the Oslo Region, south Norway, in Racheboeuf, P.R., and Emig, C.C., eds., Les Brachiopodes Fossiles et Actuels: Brest, Biostratigraphie du Paleozoique, v. 4, p. 465–475. Harper, D.A.T., 2006, The Ordovician biodiversification: Setting an agenda for marine life: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 232, p. 148–166, doi: 10.1016/j.palaeo.2005.07.010. Harper, D.A.T., and Gallagher, E., 2001, Diversity, disparity and distributional patterns amongst the orthide brachiopod groups: Journal of the Czech Geological Society, v. 46, p. 87–93. Harper, D.A.T., and Hints, L., 2001, Distribution and diversity of the Ordovician articulated brachiopods in the east Baltic, in Brunton, C.H.C., Cocks, L.R.M., and Long, S.L., eds., Brachiopods Past and Present: Systematics Association Special Volume, v. 63, p. 315–326. Harper, D.A.T., and Mac Niocaill, C., 2002, Early Ordovician rhynchonelliformean brachiopod diversity: Comparing some platforms, margins and intra-oceanic sites around the Iapetus Ocean, in Crame, J.A., and Owen, A.W., eds., Paleobiogeography and Biodiversity Change: The Ordovician and Mesozoic-Cenozoic Radiations: Geological Society [London] Special Publication 194, p. 25–34. Harper, D.A.T., and Pickerill, R.K., 1996, Mid Ordovician commensal relationships between articulate brachiopods and a trepostome bryozoan from eastern Canada: Atlantic Geology, v. 32, p. 181–187.
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Harper, D.A.T., and Rong Jia-yu, 2001, Palaeozoic brachiopod extinctions, survival and recovery: Patterns within the rhynchonelliformeans: Geological Journal, v. 36, p. 317–328, doi: 10.1002/gj.897. Harper, D.A.T., and Wright, A.D., 1996, Brachiopods, in Harper, D.A.T., and Owen, A.W., eds., Fossils of the Upper Ordovician: Palaeontological Association Field Guide to Fossils 7, p. 63–94. Harper, D.A.T., Brunton, C.H.C., Cocks, L.R.M., Copper, P., Doyle, E.N., Jeffrey, A.L., Owen, E.F., Parkes, M.A., Popov, L.E., and Prosser, C.D., 1993, Brachiopods, in Benton, M.J., ed., Fossil Record 2: London, Palaeontological Association and Chapman & Hall, p. 427–462. Harper, D.A.T., Mac Niocaill, C., and Williams, S.H., 1996, The palaeogeography of early Ordovician Iapetus terranes: An integration of faunal and palaeomagnetic constraints: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 121, p. 297–312, doi: 10.1016/0031-0182(95)00079-8. Harper, D.A.T., Rong Jia-yu, and Zhan Ren-bin, 1999, Late Ordovician development of deep-water brachiopod faunas: Acta Universitatis Carolinae: Geologica, v. 43, p. 351–353. Harper, D.A.T., Cocks, L.R.M., Popov, L.E., Sheehan, P.M., Bassett, M.G., Copper, P., Holmer, L.E., Jin Jisuo, and Rong Jia-yu, 2004a, Brachiopods, in Webby, B.D., Paris, F., Droser, M.L., and Percival, I.G., eds., The Great Ordovician Biodiversification: New York, Columbia University Press, p. 157–178. Harper, D.A.T., Villas, E., and Ortega, G., 2004b, Lipanorthis Benedetto from the Tremadocian of NW Argentina reidentified as a dalmanellidine: Significance for the origin and early radiation of the punctate orthide brachiopods: Lethaia, v. 37, p. 271–279, doi: 10.1080/00241160410006537. Hints, L., and Harper, D.A.T., 2003, Review of the Ordovician rhynchonelliformean Brachiopoda of the East Baltic: Their distribution and biofacies: Bulletin of the Geological Society of Denmark, v. 50, p. 29–43. Holmer, L.E., Popov, L.E., and Bassett, M.G., 2000, Early Ordovician organophosphatic brachiopods with Baltoscandian affinities from the Alay Range, southern Kyrgystan: GFF, v. 122, p. 367–375, doi: 10.1080/ 11035890001224367. Holmer, L.E., Popov, L.E., Koneva, S.P., and Bassett, M.G., 2001, Cambrian– Early Ordovician brachiopods from Malyi Karatau, the Western Balkash region, and Tien Shan, Central Asia: Special Papers in Palaeontology, v. 65, 180 p. Jaanusson, V., 1984, What is so special about the Ordovician?, in Bruton, D.L., ed., Aspects of the Ordovician System: Oslo, Universitetsforlaget, p. 1–3. Li Xing and Droser, M.L., 1999, Lower and Middle Ordovician shell beds from the Basin and Range Province of the Western United States (California, Nevada, and Utah): Palaios, v. 14, p. 215–233, doi: 10.2307/3515435. Lockley, M.G., 1983, A review of brachiopod dominated palaeocommunities from the type Ordovician: Palaeontology, v. 26, p. 111–145. Marshall, C.R., 2006, Explaining the Cambrian “explosion” of animals: Annual Review of Earth and Planetary Sciences, v. 34, p. 355–384, doi: 10.1146/ annurev.earth.33.031504.103001. McBride, D.J., 1976, Outer shelf communities and trophic groups in the Upper Cambrian of the Great Basin: Brigham Young University Studies, v. 23, p. 139–152. McGhee, G.R., Jr., Sheehan, P.M., Bottjer, D.J., and Droser, M.L., 2004, Ecological ranking of Phanerozoic biodiversity crises: Ecological and taxonomic severities are decoupled: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 211, p. 289–297, doi: 10.1016/j.palaeo.2004.05.010. Mergl, M., 1999, Inarticulated brachiopod communities in Tremadoc-Arenig of Prague Basin: A review: Acta Universitatis Carolinae: Geologica, v. 43, p. 337–340. Miller, A.I., 1997a, Comparative diversification dynamics among Palaeocontinents during the Ordovician Radiation: Geobios (Lyon, France), Mémoire Special, v. 30, p. 397–406, doi: 10.1016/S0016-6995(97)80044-7. Miller, A.I., 1997b, Dissecting global diversity trends: Examples from the Ordovician radiation: Annual Review of Ecology and Systematics, v. 28, p. 85–104, doi: 10.1146/annurev.ecolsys.28.1.85. Miller, A.I., 1997c, A new look at age and area: The geographic and environmental expansion of genera during the Ordovician radiation: Paleobiology, v. 23, p. 410–419. Miller, A.I., 2004, The Ordovician Radiation: Towards a new global synthesis, in Webby, B.D., Paris, F., Droser, M.L., and Percival, I.G., eds., The Great Ordovician Biodiversification: New York, Columbia University Press, p. 380–388.
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MANUSCRIPT ACCEPTED BY THE SOCIETY 6 OCTOBER 2009
Printed in the USA
The Geological Society of America Special Paper 466 2010
Ordovician paleogeography and tectonics of the major paleoplates of China Chen Xu† Zhou Zhi-yi† Fan Jun-xuan† State Key Laboratory of Palaeobiology and Stratigraphy, Nanjing Institute of Geology and Palaeontology, Chinese Academy of Sciences, Nanjing 210008, China
ABSTRACT The tectonics and paleogeography of Ordovician rocks in China record the four major paleoplates—the South China, North China, Tarim, and Xizang (Tibet) Blocks. New paleogeographic maps of South China for the Tremadocian, Darriwilian, Sandbian–early Katian, and late Katian–Hirnatian time intervals display lithofacies and biofacies belts that depict continuous changes from the Yangtze Platform through the Chiangnan (Jiangnan) Slope to the Zhujiang Basin. North China was dominantly a carbonate platform during the Ordovician. Facies belts, particularly the trilobite biofacies belts, change westward from the platform edge to the slope along the west margin of the platform. In Tarim, Ordovician rocks provide the main source and reservoir rocks for oil and gas. The vast expanse of the block was a northward-deepening, shallow-water platform that was fringed by peripheral, deeper water facies belts developed along the northern side of the South Tianshan. Facies analysis indicates that the paleogeographic setting varied in response to eustatic sea level changes. North of the Tibet Block lies a long mobile belt that crosses more than half of China from west to east. The Tibet Block is mainly composed of two units, the north Qiantang region, which remained as a separate, small paleoplate through the late Paleozoic, and the south Gandise-Himalaya region, which persisted as a distinct paleoplate through the late Paleozoic and well into the Mesozoic. Ordovician rocks consisting mainly of carbonates with shelly faunas have been recorded from regions of the Tibet Block.
INTRODUCTION—THE CRUSTAL BLOCKS OF CHINA
with explanations in his significant book On Major Tectonic Forms of China, and he recognized five successive orogenic cycles, which were either monocyclic or polycyclic and which he termed Caledonian (early Paleozoic), Variscan (Late Devonian to end of Permian), Indosinian (Late Triassic to Early Jurassic), Yenshanian (Late Jurassic to end of Cretaceous), and Himalayan (Oligocene to Pleistocene); the terms Caledonian, Hercynian,
Study of the tectonics of China began with Richthofen’s three famous volumes on China from 1887 to 1912, which were followed by the work of Grabau (1924, 1928), Wong (1926), and Lee (1939). Huang (1945) published a series of tectonic maps †
E-mails:
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[email protected].
Chen Xu, Zhou Zhi-yi, and Fan Jun-xuan, 2010, Ordovician paleogeography and tectonics of the major paleoplates of China, in Finney, S.C., and Berry, W.B.N., eds., The Ordovician Earth System: Geological Society of America Special Paper 466, p. 85–104, doi: 10.1130/2010.2466(06). For permission to copy, contact
[email protected]. ©2010 The Geological Society of America. All rights reserved.
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and Variscan are used commonly for orogenic belts in China that formed contemporaneously with the classical orogenic belts in Europe to which these terms were first applied and are hereafter written as “Caledonian,” “Hercynian,” and “Variscan” to prevent confusion. Huang’s concept of orogenic cycles was followed by Chinese geologists until the 1970s. When the theory of plate tectonics was introduced to China, Huang and his successors revised their explanation of the major tectonic belts of China as features produced by polycyclic plate tectonic processes (Huang and Jiang, 1962; Huang et al., 1977), and this revised concept of the orogenic belts was then followed by Ren et al. (1999) when they published the new tectonic map of China and its explanation. Chen and Rong (1992) and Chen et al. (1995a) published maps illustrating the distribution of paleoplates of China during the Ordovician Period. Four major blocks—South China, North China, Tarim, and Xizang (Tibet)—were recognized with metamorphosed Precambrian basement rocks of different ages overlain by Sinian sedimentary rocks. The fundamental basement structures influenced the subsequent tectonic regimes, which, in turn, significantly controlled the types and distributions of sedimentary rocks that accumulated on the paleoplates. The fundamental basement complexes of each of the four major blocks were described by Wang (1985) and serve as a framework for the Paleozoic tectonic units (Fig. 1). The boundaries of the Ordovician plates lie along suture zones that can be traced on current geological maps. However, the ages of the suture zones differ significantly from one another (Ren et al., 1999). Fourteen
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Figure 1. Major paleoplates of China during the Ordovician; modified from Wang (1985) and Chen et al. (1995a). KZ—Kazakhstan Block; HG—Hinggan Block; TA—Tarim paleoplate; NC—North China paleoplate; CH—Chaidam Block; SG—Songpan-Garze (Ganzi) Block; TB—Tibet paleoplate; SB—Sibumasu Block; SC—South China paleoplate; YK—Yukai Block.
suture zones of Salairian (Middle Cambrian) and “Caledonian” (Late Silurian) ages are distributed in China and its neighboring regions (Fig. 2). Among them, one suture zone (no. 9 of Fig. 2) developed from the early Neoproterozoic to the “Caledonian,” one (no. 2 of Fig. 2) from the Neoproterozoic to the Yenshanian, two (nos. 3 and 4 of Fig. 2) from the Salairian, one (no. 1 of Fig. 2) from the Salairian to the Yenshanian, seven (nos. 5–7, 10–13 of Fig. 2) during the “Caledonian,” and two (nos. 8 and 14 of Fig. 2) are possibly “Caledonian” in age. The 14 Salairian and “Caledonian” suture zones are characterized either by ophiolite belts or other oceanic crustal rocks, and they define the boundaries of four major cratonic blocks during the Ordovician or early Paleozoic. The tectonic history of China is complicated. The Chinese Paleozoic blocks converged with Gondwana during the Permian, when they were located in eastern Pangaea. Subsequently, the Pacific oceanic plate and the Indian plate converged with the Chinese blocks during the Yenshanian and Himalayan orogenic cycles. The resulting collisional zones are recognized today as major orogenic regions termed the Paleoasian (northern China), Tethyan (southern China), and Peri-Pacifican (eastern China) orogenic belts. The history of convergence and accretion appears to have advanced progressively from Siberia in the north to India in the south. However, each orogenic belt might have been active many times from the Proterozoic to the Cenozoic. Besides the four major paleoplates of South China, North China, Tarim, and Xizang, we also recognize numerous smaller blocks (microplates) that likely existed in the early Paleozoic and were subsequently included into the China mainland with the major orogenic belts (Fig. 3). During the Ordovician, two small blocks, the Tuva-Mongolia Block (no. 1 of Fig. 3) and the Central Mongolia-Ergun Block (no. 2 of Fig. 3) were in the Salairian orogenic belt directly south of the Siberian craton. The Kokchetav-Issyk Block (nos. 3–4 of Fig. 3), the Balkhash-Illi Block (no. 5 of Fig. 3), the Junggar Block (no. 6 of Fig. 3), the Dariganga Block (no. 7 of Fig. 3), the Zalantun Block (no. 8 of Fig. 3), the Xingxingxia Block (no. 9 of Fig. 3), the Hanshan Block (no. 10 of Fig. 3), the Yagan Block (no. 11 of Fig. 3), the Totoshan Block (no. 12 of Fig. 3), the Xilinhot Block (no. 13 of Fig. 3), the Songhuajiang Block (no. 14 of Fig. 3), the BureyaJiamusi Block (no. 15 of Fig. 3), and the Xingkai Block (no. 16 of Fig. 3) are along the “Hercynian” Tianshan-Hinggan orogenic belt. Ordovician marine strata and volcanic rocks in these small blocks indicate that they have existed since the early Paleozoic, although three of them (nos. 14–16 of Fig. 3) were older lands without Ordovician deposits. The Central West Kunlun Block (no. 17 of Fig. 3), the Altun Block (no. 18 of Fig. 3), the Jinshuikou Block (no. 19 of Fig. 3), the Lenghu Block (no. 20 of Fig. 3), the Olonbulag Block (no. 21 of Fig. 3), the Central Qilian Block (no. 22 of Fig. 3), the Central East Qinling Block (no. 23 of Fig. 3), and the Wudang Block (no. 24 of Fig. 3) were distributed along the Kunlun-Qilian-Qinling “Caledonian-Hercynian” polycyclic orogenic belt (Ren et al., 1999). Ordovician rocks in these small blocks consist mostly of marine deep-water clastics
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Figure 2. Salairian and “Caledonian” suture zones of China and its neighboring regions, after Ren et al. (1999). IN—Indian plate; SB—Siberian plate; (1) Mongolia-Okhotsk suture zone (Salairian-Yenshanian polycyclic suturing); (2) Kunlun-Qinling suture zone (Yangtzeian-Yenshanian polycyclic suturing); (3) Bayanhongor suture; (4) Ob’-Lake Area suture zone (Salairian); (5) Naiman-Dzhalair suture zone (“Caledonian”); (6) Xiaohuangshan suture zone (“Caledonian”); (7) Jiamusi-Mudanjiang suture zone (“Caledonian”); (8) West Kunlun suture zone (“Variscan”); (9) Jinyanshan suture zone (Yangtzeian and “Caledonian”); (10) North Qilian suture zone (“Caledonian”); (11) Northern Qaidam margin suture (“Caledonian”); (12) Tengtiaohe-Song Ma suture zone (“Caledonian”); (13) Nan suture zone (“Caledonian”); 14—Lungmuco-Lancangjiang suture zone (“Caledonian”).
and shallower water carbonate deposits. One small early Paleozoic block, the Zongza Block (no. 26 of Fig. 3), lies within the Tethyan Indosinian to Yenshanian orogenic belt. The Tibetan paleoplate might have existed as four subblocks, the Qamdo (no. 25 of Fig. 3), Qiangtang (no. 28 of Fig. 3), Lhasa (no. 30 of Fig. 3), and Himalaya (no. 31 of Fig. 3), during the early Paleozoic. The Sibumasu Block (no. 29 of Fig. 3, i.e., the Shan Thai paleoplate of Chen et al., 1995a), was separated from other major blocks in China, and the Yunkai Block (no. 27 of Fig. 3) likely was far away from South China during the Ordovician (Chen et al., 1995b).
ORDOVICIAN PALEOGEOGRAPHY OF MAJOR CHINESE PALEOPLATES Three different versions of Ordovician paleogeographic maps published by Liu (1955), Guan et al. (1984), and Wang (1985) have been widely used in China. Recently, Feng et al. (2003a, 2000b, 2004) also published paleogeographic maps for the Lower, Middle, and Upper Ordovician rocks of China. Unfortunately, their tripartite division of the Ordovician and their regional stages from South China, North China, and Northwest China are very different from the global standard series and stages recently
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approved by the International Subcommission on the Ordovician System and the International Commission on Stratigraphy. In this chapter we describe new paleogeographic maps that are based on correlations using the new global Ordovician chronostratigraphygeochronology. Paleogeographic maps for South China paleoplates are shown for the Tremadocian, Darriwilian, Sandbian–early Katian, and late Katian–Hirnantian Ages. The paleogeography of the North China paleoplate is represented by maps for the Dapingian, early to mid-Darriwilian, late Darriwilian–early Katian, and late Katian time intervals, and that of the Tarim paleoplate is illustrated for the Tremadocian–Dapingian, Darriwilian–early Katian, and late Katian time intervals. We construct only one paleogeographic map for Xizang (Tibet).
South China The Ordovician rocks and fossils of South China are well studied, and the four new paleogeographic reconstructions are based on data from 203 sections or localities (Fig. 4). The geologic evidence from these localites is derived from Chen et al. (1995b), Zhou et al. (2000, 2003), Yuan et al. (2000), Zhang et al. (2002), Zhou et al. (1993), Du and Xu (1990), Li (1995), Nan and Zhou (1996), Compiling Group of Regional Stratigraphy Chart of Sichuan (1978), and a series of publications of the Bureau of Geology and Mineral Resources of Jiangxi (1984), Jiangsu (1984), Guangxi (1985), Guizhou (1987), Hunan (1988), Zhejiang (1989), Shaanxi (1989), Yunnan (1990), Hubei (1990), and
Ordovician paleogeography and tectonics of the major paleoplates of China Sichuan (1991). The late Katian–Hirnantian map is after Chen et al. (2004). Huang (1945) described South China as composed of three parts: the Yangtze Platform, the Chiangnania (Jiangnan transitional belt of Chen and Rong, 1992; Jiangnan slope of Chen et al., 1995b), and the Cathaysia. He considered Cathaysia as a “Caledonide” orogenic belt. Ren et al. (1999) generally followed Huang’s concept but divided South China into only two parts: the Yangtze Paraplatform and the South China “Caledonides.” We redefine the South China Block to include three parts: the Yangtze Platform and the Cathaysian Land with the Zhujiang Basin in between. The Chiangnan (Jiangnan) transitional slope belt is the northwest slope of the Zhujiang Basin, and it is characterized by biofacies and lithofacies that are transitional between those of the Yangtze Platform and the Zhujiang Basin (Rong and Chen, 1987; Chen and Rong, 1992; Chen et al., 1995b). On the basis of sedimentary successions and facies distributions of pre-Devonian rocks in Hunan Province, Chen et al. (1997) and Chen and Mitchell (1996) interpreted the tectonic development of the Cathaysian Land, which is fundamental to reconstructions of the paleogeographic evolution of South China. The initiation of uplift and erosion of an immature, low-grade metamorphic source region in southern Hunan Province was recognized in strata of the Upper Ordovician Diplacanthograptus spiniferus Biozone in the Taojiang area, Central Hunan, and the sedimentary record is progressively younger to the north. These clastic strata are attribubted to a flexurally mediated foreland basin associated with a northwestward-verging collision between the South China paleoplate and an as yet unidentified block to the southeast. As a result of this tectonic event, the Cathaysian Land increased gradually northwestward, and the Yangtze Platform was entirely uplifted by the end of the Early Silurian Telychian Age (Chen and Rong, 1996; Rong et al., 2003). Tremadocian (Fig. 5) The Yangtze Sea covered most of the South China paleoplate during the Tremadocian Age except for three adjacent land areas: Longmen Land to the northwest, Dian-Qian Land to the southwest, and Cathaysian Land in the east. Clastic tidal flat deposits are distributed around the north edge of Dian-Qian Land. The facies change gradually to inner shelf dolomitic deposits and then to widely distributed open shelf carbonate deposits that occupied the central Yangtze Platform. The representative rock units along this lithofacies transition are the Tangchi Formation (mostly shale; type locality near Kunming, loc. 38, Fig. 4) in the tidal flat belt, the Hsihsiangchih Formation (mostly dolomites; type locality at loc. 28, Fig. 4) in the inner shelf belt, and the Tungtzu Formation (mostly limestone and shale; type locality at loc. 64, Fig. 4) as well as the Nantsinkuan and Fenhsiang Formations (mostly limestone; type locality at loc. 111, Fig. 4) in the open shelf region (Chen et al., 1995a). Southwestward from the central Yangtze Platform to the Zhujiang Basin, lithofacies and biofacies belts are parallel to the slope of the Chiangnan (Jiangnan) transitional belt (Zhang et al., 2002; Yuan et al., 2000). The
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Yangtze Platform open shelf region was characterized by dolomitic limestone and bioclastic limestone, and the platform edge was the site of deposition of banded limestone and oolitic limestone in the upper Yangtze region (Chen et al., 1995a). Stromatolites and Calathium microbial reefs occur in this biofacies at Dongzhi (loc. 133, Fig. 4) of the lower Yangtze region (Li et al., 2004). Micritic limestone was deposited on the upper slope, and mudstone and silty mudstone accumulated along the lower slope. Biofacies gradually change along with the lithofacies. Yuan et al. (2000) defined the trilobite Asaphellus-Dactylocephalus biofacies (loc. 111, Fig. 4), Dikelokephalinid biofacies (loc. 111), and Tungtzuella-Asaphopsis biofacies (loc. 111 to loc. 107, Fig. 4) from the open shelf lithofacies, and the Psilocephalina biofacies (loc. 94, Fig. 4) along the platform-edge lithofacies in Hunan. The Metayuepingia-shumardiid biofacies occurs in the upper slope lithofacies (loc. 95, Fig. 4, and southward), and finally the Olenid biofacies is distributed along the low slope lithofacies (loc. 181, Fig. 4, and southeastward). The Cathaysian Land was likely the primary source area of clastic sediments for the Zhujiang Basin (Zhou et al., 1993; Xu et al., 1996). Xu et al. (1996) reported that the grain size of these clastic sediments decreases westward from the southeast China coast, and Zhou et al. (1993) subdivided the Zhujiang Basin into different lithofacies. However, because evidence for chronocorrelation is virtually absent from the stratigraphic succession of the Zhujiang Basin—the Ordovician age of these strata is based on a single locality where Ordovician graptolites occur (loc. 204, Fig. 4)—facies reconstructions for the basin are not possible. Darriwilian (Figs. 6, 7) The lithofacies and biofacies patterns of the Tremadocian persisted through the Floian and Dapingian and into the Darriwilian. Narrow, nearshore dolomitic and detrital facies occur around the southwestern lands of the Yangtze Platform (Zhang et al., 2002). An inner-shelf carbonate and mudstone belt occupied the western Yangtze Platform with Martelia-Lapidorthis (lower Darriwilian, upper part of the Meitan Formation, loc. 61–70, 90–92, Fig. 4), and Saucrorthis (upper Darriwilian, Dashaba Formation, loc. 54–58, Fig. 4) brachiopod faunas in the clastic or mudstone facies. An endemic Meitanoceras nautiloid fauna occurs in this area (Nanjing Institute of Geology and Palaeontology, 1974; loc. 72–73, Fig. 4). On the eastern part of the Yangtze Platform shallow open shelf carbonates accumulated (Kuniutan Formation) with a Birmanites-Nileus-Lonchodomas trilobite fauna (loc. 97–113, Fig. 4), and a Doderoceras wahlenbergi nautiloid fauna (Nanjing Institute of Geology and Palaeontology, 1974, loc. 97–113, Fig. 4). Conodonts from the Kuniutan Formation (Darriwilian) in W Hubei and NW Hunan (loc. 94–113, Fig. 4) are of Atlantic Province faunal affinity and represent the Lenodus antivariabilis, L. variabilis, Yangtzeplacognathus crassus, Dzikodus tablepointensis, Eoplacognathsu suecicus, Y. foliaceus, and Y. protoramosus Biozones (Zhang, 1998). In contrast to the abundant shelly faunas of the Yangtze Platform (Fig. 6), the strata of the Chiangnan (Jiangnan) transitional
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Figure 4 (on preceding page). Localities in South China at which Ordovician strata have been described. (1) Zhaojiaba, Ningqiang, SW Shaanxi; (2) Maoping, Ningqiang, SW Shaanxi; (3) Shuanghe, Wangchang, N Sichuan; (4) Qiaoting, Nanjiang, N Sichuan; (5) Zhongliangcun, Nanzheng, S Shaanxi; (6) Shangliangshan, Nanzheng, S Shaanxi; (7) Xihe, Nanzheng, S Shaanxi; (8) Fucheng, Nanzheng, S Shaanxi; (9) Sanlangpu, Xixiang, S Shaanxi; (10) Xiaoyang, Zhenba, S Shaanxi; (11) Qingsi, Zhenba, S Shaanxi; (12) Zihuang, Ziyang, SE Shaanxi; (13) Bajiaokou, Ziyang, SE Shaanxi; (14) Caojia, Wanyuan, N Chongqing; (15) Liaozikou and Datangkou, Chengkou, NE Chongqing; (16) Yangjiaba and Tianba, Chengkou, NE Chongqing; (17) Sanleiba, Guangyuan, N Sichuan; (18) Shangshi, Guangyuan, N Sichuan; (19) Maniaogou, Pingwu, NW Sichuan; (20) Guiyuan, Beichuan, NW Sichuan; (21) Wuan, Beichuan, NW Sichuan; (22) Wulangmiao, Anxian, NW Sichuan; (23) Longwangmiao, Mianzhu, NW Sichuan; (24) Erlangshan, Tianquan, W Sichuan; (25) Binlingxiang, Hongya, W Sichuan; (26) Jiaodingshan, Hanyuan, W Sichuan; (27) Jinkouhe, Hanyuan, W Sichuan; (28) Gaodongkou, Emeishan, W Sichuan; (29) Pogongxiang, Ganluo, SW Sichuan; (30) Liuhong, Meigu, SW Sichuan; (31) Tiezhufeike, Butuo, SW Sichuan; (32) Shuhe, Yanyuan, SW Sichuan; (33) Dapingzi, Yanbian, SW Sichuan; (34) Genhaizi, Huidong, SW Sichuan; (35) Beicaoping, Ninglang, NE Yunnan; (36) Wande, NE Yunnan; (37) Wuding, NE Yunnan; (38) Ercun, Kunming, E Yunnan; (39) Qujing, E Yunnan; (40) Xindian, Qiaojia, NE Yunnan; (41) Xiaohejie, Ludian, NE Yunnan; (42) Huanggexi, Dakuan, NE Yunnan; (43) Sandaoshui, Yongshan, NE Yunnan; (44) Zhonghechang, Yanjing, NE Yunnan; (45) Yanjing, NE Yunnan; (46) Huangjuecao, Yanjing, NE Yunnan; (47) Yangchang, Zhenxiong, NE Yunnan; (48) Shiziba, Zhenxiong, NE Yunnan; (49) Mianbugeng, Zhenxiong, NE Yunnan; (50) Sanqiutian, Zhenxiong, NE Yunnan; (51) Yujingshan-Beiguoshu, Weixing, NE Yunnan; (52) Weiyuan well, Central Sichuan; (53) Yanggaosi, Luzhou, S Sichuan; (54) Shuanghe, Changning, S Sichuan; (55) Fuxingxiang, Changning, S Sichuan; (56) Shizitan, Gongxian, S Sichuan; (57) Gusong, Xingwen, S Sichuan; (58) Daping, Gulin, S Sichuan; (59) Yanzikou, Bijie, W Guizhou; (60) Yejuo-Daye, Bijie, W Guizhou; (61) Yangliugou, Renhuai, N Guizhou; (62) Shichang, Renhuai, Guizhou; (63) Donggongsi, Zunyi, N Guizhou; (64) Honghuayuan, Tongzi, N Guizhou; (65) Sancha, Tongzi, N Guizhou; (66) Sangmuchang, Xishui, N Guizhou; (67) Wenshui, Xishui, N Guizhou; (68) Tongguyuan-Liangfongya, Tongzi, N Guizhou; (69) Hanjiadian, Tongzi, N Guizhou; (70) Guanyinqiao, Qijiang, S Chongqing; (71) Xikou, Yuechi, E Sichuan; (72) Niuchang, Meitang, N Guizhou; (73) Wulipo, Meitan, N Guizhou; (74) Leijiatun-Qishuwan, Shiqian, N Guizhou; (75) Guanyintang, Fenggang, N Guizhou; (76) Balixi, Fenggang, NE Guizhou; (77) Dongkala, Fenggang, NE Guizhou; (78) Yingwuxi, Sinan, NE Guizhou; (79) Yangjiazhai, Heshui, Yinjiang, NE Guizhou; (80) Longjingpo, Wuchuan, NE Guizhou; (81) Ganxi, Yanhe, NE Guizhou; (82) Siqu, Yanhe, NE Guizhou; (83) Ludiping and Huangban, Songtao, NE Guizhou; (84) Lailongshan, Songtao, NE Guizhou; (85) Datianba, Xiushan, SE Chongqing; (86) Rongxi, Xiushan, SE Chongqing; (87) Xiaoxian, Youyang, SE Chongqing; (88) Jiangkou, Wulong, S Chongqing; (89) Jielongchang, Wulong, S Chongqing; (90) Jiangkou, Daozheng, NE Guizhou; (91) Podu, Nanchuan, S Chongqing; (92) Ganhegou, Qianjiang, S Chongqing; (93) Wentang, Zhangjiajie, NW Hunan; (94) Maocaopu, Reshi, Taoyuan, N Hunan; (95) Jiuxi, Taoyuan, N Hunan; (96) Erfangping, Cili, N Hunan; (97) Longchihe, Nishi, Shimeng, NW Hunan; (98) Gaoluo, Xuan’en, SW Hubei; (99) Lianghekou, Xuan’en, SW Hubei; (100) Taiyanghe, Enshi, SW Hubei; (101) Sanmuliangzi, Jianshi, SW Hubei; (102) Changliang teagarden, Jianshi, SW Hubei; (103) Siyangchiao, Badong, W Hubei; (104) Huaqiao, Changyang, S Hubei; (105) Qianhe, Wufeng, S Hubei; (106) Xiejiuping, Songzhi, S Hubei; (107) Liujiachang, Songzhi, S Hubei; (108) Xizhai, Songzhi, S Hubei; (109) Maohutang, Yidu, S Hubei; (110) Huanghuachang, Yichang, W Hubei; (111) Fenxiang, Yichang, W Hubei; (112) Wangjiawan and Tangya, Yichang, W Hubei; (113) Xintan, Zigui, W Hubei; (114) Xujiaba, Wuxi, NW Chongqing; (115) Shenlongjia, NW Hubei; (116) Qingquan, Fangxian, NW Hubei; (117) Maliangping, Baokang, N Hubei; (118) Limiao, Nanzhang, N Hubei; (119) Huitingshan, Jingshan, E Hubei; (120) Yangloudong, Puxi, SE Hubei; (121) Huangmachong, Chongyang, SE Hubei; (122) Liuzuiba, Tongshan, SE Hubei; (123) Zhoujia, Xianning, SE Hubei; (124) Changshouban, Xianning, SE Hubei; (125) Shiyucun, Yangxin, SE Hubei; (126) Zhangshan, Huangshi, SE Hubei; (127) Longshan, Huangmei, SE Hubei; (128) Guantangyuan, Wuning, NW Jiangxi; (129) Lishuwo-Xilong, Wuning, NW Jiangxi; (130) Dianbei, Wuning, NW Jiangxi; (131) Longshan, Susong, SE Anhui; (132) Tuolongshan, Taihu, SE Anhui; (133) Lijia, Dongzhi, S Anhui; (134) Zhangjiatan-Chengshanzhai, Shitai, S Anhui; (135) Liyang, Shitai, S Anhui; (136) Dawugan, Shitai, S Anhui; (137) Baian, Guichi, Central Anhui; (138) Wujiagou, Guichi, Central Anhui; (139) Zhangcunxu, Qingyang, Central Anhui; (140) Beigongli, Jingxian, Central Anhui; (141) Changhongguan, Guangde, SW Anhui; (142) Tangxing, Guangde, SW Anhui; (143) Donggushan, Lujiang, Central Anhui; (144) Yuanshan, Wuwei, Central Anhui; (145) Shanaoting, Hanshan, Central Anhui; (146) Xiangquan, Hexian, Central Anhui; (147) Sanyuanzi, Chuxian, N Anhui; (148) Tangshan, Nanjing, SE Jiangsu; (149) Kaoluncun, Jurong, SE Jiangsu; (150) Shanghai; (151) Duyun, S. Guizhou; (152) Sandu, S. Guizhou; (153) Yupin, NE Guizhou; (154) Tongren, NE Guizhou; (155) Fenghuang, W. Hunan; (156) Zhijiang, W. Hunan; (157) Huihua, W. Hunan; (158) Liue, Penglai, Dayouping, Linggui, NW Guangxi; (159) Shangpingjie, Guanyang, NW Guangxi; (160) Shengping, Xingan, NW Guangxi; (161) Baimaowu, Quanzhou, NW Guangxi; (162) Daoxian-Xintian, S Hunan; (163) Fuxiangshan, Xintian, Hunan; (164) Chengxian, SE Hunan; (165) Jicun, Dayu, SE Jiangxi; (166) Sishun, Chongyi, SE Jiangxi; (167) Xiankou, Shuichuan, SE Jiangxi; (168) Pulong, Sanwan, and Heyeping, Ninggang, SE Jiangxi; (169) Hanjiang, Qixiling, and Shikou, Yongxin, SE Jiangxi; (170) Damiaokou, Dong’an, S Hunan; (171) Daliangtuo, Chengbu, SE Hunan; (172) Guanxia, Shuining, SE Hunan; (173) Yuexi, Dongkou, SE Hunan; (174) Shuangfuting, Qidong, S Hunan; (175) Xiaoyuanchong, Hengyang, Central Hunan; (176) Tianmashan, Shuangfong, Central Hunan; (177) Dapingxi, Xinhua, Central Hunan; (178) Yenxi, Anhua, Central Hunan; (179) Jintian, Anhua, Central Hunan; (180) Dafuping, Anhua, Central Hunan; (181) Nanshichong, Xiangtaoyuan, Taojiang, Central Hunan; (182) Yanzitan, Yiyang, Central Hunan; (183) Yuanjiang, N Hunan; (184) Zhuzhou, E Hunan; (185) Longtan, Tongcheng, SE Hubei; (186) Hongtan, Yixian, S Anhui; (187) Kuocun, Taiping, S Anhui; (188) Taokeng, Taiping, S Anhui; (189) Shibixia, Taiping, S Anhui; (190) Shilibai, Jixi, S Anhui; (191) Hulo, Ningguo, S Anhui; (192) Zhufongpu, Ningguo, S Anhui; (193) Shangjincun, Ningguo S. Anhui; (194) Xiaofeng, Anji, NW Zhejiang; (195) Huangshu, Anji, NW Zhejiang; (196) Sanqiaopu, Deqing, NW Zhejiang; (197) Changhua, Linan, NW Zhejiang; (198) Yankou, Yuqian, Linan, NW Zhejiang; (199) Banqiao, Linan, NW Zhejiang; (200) Yindianjie, Zhuji, N Zhejiang; (201) Wenchang, Chunan, NW Zhejiang; (202) Zitang, Longyou, W Zhejiang; (203) Jiangshan-Changshan-Yushan (JCY) area, Zhejiang-Jiangxi border area; (204) Weifang, Dahu, Yongan, Fujian; (205) Longtuozai, Qujiang, N Guangdong.
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Figure 5. Tremadocian paleogeographical map of South China.
belt and the Zhujiang Basin are rich in graptolites. The Chiangnan belt is characterized by slope mudstone-clastic deposits, and the Zhujiang Basin by basinal flysch sediments. The transition of Darriwilian biofacies and lithofacies from the Yangtze Platform to the Zhujaing Basin are noteworthy. Graptolite diversity is very high (41 genera) in strata deposited on the Chiangnan (Jiangnan) Slope (41 genera), indicating that graptolites flourished in the overlying waters as opposed to the more shallow sea of the platform (only 9 genera) or the oceanic water of the deep basin (19 genera) (Fig. 7). The faunal differentiation and diversity change may be related to nutrient availability. Finney and Berry (1998) suggested that graptolites flourished in relatively narrow bands of upwelling waters along, and extending somewhat oceanward from, certain continental margins. Cooper (1999) came to the same conclusion on the basis of the distribution of Tremadocian graptolites. The fossil preservation potential of various lithofacies also may have influenced the apparent diversity of graptolites in sedimentary successions from the carbonate dominated Yantze Platform to the turbidite or flysch dominated deposits of the Zhejiang Basin. However, the primary factor of diversity control is not dependent on the regional lithofacies.
Most of the graptolite localities in the Zhujiang Basin are in Hunan and eastern Jiangxi. Only one locality (loc. 204, Fig. 4) in central Fujian yields graptolites (Expansograptus sp. aff. E. hirundo) as well as the chitinozoan Bursachitina oelundica (Zhu and Zhou, 1990). These taxa occur in turbidite deposits of the Weifeng Formation and are composed of feldspathic and quartzose graywacke, metamorphosed siltstone, and phyllitic mudstone (Chen et al., 1995b). Expansograptus sp. aff. E. hirundo indicates a Dapingian to early Darriwilian Age, and the Weifeng Formation is nearest to the Ordovician Cathaysian Land source region, which contains biostratigraphic and paleogeographic fossil evidence. Sandbian–Early Katian (Fig. 8) Ordovician marine transgression reached its climax during the mid-Sandbian Age (Chen and Qiu, 1986; Zhou et al., 2000). The Sandbian–early Katian deposits on the Yangtze Platform are represented by two major units, the lower Datianba-Miaopo Formation and the middle-upper part of the Pagoda Formation. The Miaopo Formation includes mainly graptolitic black shale with limestone concretions; the graptolite fauna represents the
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Figure 6. Darriwilian paleogeographical map of South China.
lowest Sandbian Nemagraptus gracilis Biozone (Geh, 1963a, 1963b). This graptolite-bearing black shale was deposited in three depressions on the Yangtze Platform (Fig. 8), and each one is surrounded by carbonate srata of the Datianba Formation. The black shale indicates anoxic bottom waters, and the presence of the limestone concretions in the Miaopo Formation indicates the alternation of CaCO3 saturated and unsaturated bottom waters. Microerosion surfaces on the limestone concretion layers indicate the presence of bottom currents (Chen and Qiu, 1986). Carbonate strata of the overlying Pagoda Formation were deposited across the entire Yangtze Platform. These carbonates are characterized by a polygonal reticulate structure that apparently formed as shrinkage cracks on the sea bottom (Chen and Qiu, 1986), as organic traces in the limestone (Lindström et al., 1991), or as penecontemporaneous deformation structures (Zhou and Xue, 2000). Nautiloids are the dominant fossils in both the Datianba Formation (Lituites fauna) and the Pagoda Formation (Sinoceras fauna). In strata that correlate with the uppermost Darriwilian to Sandbian Pygodus anserinus–Prioniodus variabilis conodont Zone interval, Zhou et al. (2000, 2003) recognized a set of trilobite biofacies that from northwest to southeast form
the shallow inner shelf of the northwestern Yangtze Platform (A, loc. 1, 2, Fig. 8) across a deeper, steep outer shelf (B, loc. 5, Fig. 8) and into the Zhujiang Basin (Zhou et al., 1999, 2001). These included a Remopleurides-Ovalocephalus biofacies from the Yichang area (loc. 110–114, Fig. 4), indicating a shallow offshelf environment; an Ovalocephalus-Cyclopyge biofacies (loc. 93, 96–97, 107, Fig. 4) and a Cyclopyge-tricucleid biofacies in the border area between Hubei and Hunan (loc. 94, Fig. 4), indicating a deeper shelf environment; a Cyclopygid biofacies in NW Hunan (loc. 177–183, Fig. 4), indicating a slope environment; and a Cyclopygid-Girvanopyge biofacies from Central Hunan (loc. 174 and southward, Fig. 4), indicating a basin environment. In the P. insculptus Biozone the Ovalocephalus-Cyclopyge biofacies belt migrated to the Yichang area, reflecting the maximum extent of marine transgression during the Katian. Late Katian–Hirnantian (Fig. 9) Chen et al. (2004) published three paleogeographic maps for the latest Katian P. pacificus graptolite Biozone, the Hirnantian, and the earliest Silurian–post-Hirnantian (A. ascensus and P. acuminatus graptolite Biozones) of South China. The latest
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Pseudobryograptus Kinnegr. ? D. (Didymogr.) Acrograptus Xiphograptus D. (Expansogr.) Pterograptus Undulograptus Hustedogr. Tetragraptus Dichograptus Loganograptus Goniograptus Wuninograptus Aulograptus Allograptus Sinograptus Tylograptus Nicholsonograptus Kalpinograptus Pseudisograptus Arienigraptus Abrograptus Protabrograptus Pseudophyllograptus Pseudotrigonograptus C. (Procardiograptus) Cardiograptus Exigraptus Skiagraptus ? Cryptograptus Glossograptus Paraglossograptus Amplexograptus Eoglyptograptus Bergstromaegraptus Haddingograptus Archiclimacograptus Proclimacograptus Normalograptus Dicaulograptus Retiograptus
shelly fauna from Nanjing and eastward. Rong and Chen (1987) and Rong et al. (2003) described a brachiopod benthic assemblage (BA) distribution pattern with water depth increasing from southwest to east across the Yangtze region. Fine grained detritial clastic sediments were deposited along the south edge of the upper Yangtze Platform and the Hubei-Hunan Submarine High, while thick clastic wedges accumulated along the southeast edge of the Yangtze Platform (Fig. 9). North China
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Figure 7. Darriwilian graptolite distribution and diversity changes from the Yangtze Platform through the Chiangnan (Jiangnan) Slope to the Zhujiang Basin.
Katian and Hirnantian maps are combined here (Fig. 9) to illustrate the major sedimentological and tectonic events. The Cathaysian Land increased rapidly during the late Katian and Hirnantian (Chen and Mitchell, 1996). Most of the Zhujiang Basin was uplifted during the Katian, and this uplift migrated northwestward such that the entire Yangtze Platform was drained of its epeiric sea and exposed at the end of the Llandovery (Chen and Rong, 1996; Rong et al., 2003). During the latest Katian, the main part of the Yangtze Platform was the site of deposition of the Wufengian black graptolitic shale and chert, which were replaced in the Hirnantian by muddy limestone, to be followed, in turn, by deposition of the Lungmachian black graptolitic shale in the latest Hirnantian to earliest Silurian. These lithofacies changes coincided with and record the expansion and contraction of the Hirnantian glacial in Gondwana and the associated glacio-eustatic fall and rise of sea level. The expansion of the Hubei-Hunan Submarine High also may be related to sea level fall. Hirnantia faunas were mostly distributed in the upper Yangtze region, and a Paramalomena fauna occupied the lower Yangtze region. In addition, a graptolite facies replaced the Hirnantian
The North China paleoplate includes mainly the North China Platform and its western margin. The north margin is not well studied, and the eastern marginal belt was cut off by the Tan-Lu fault system during the Mesozoic (Chen et al., 1995b). The southern margin is also poorly exposed and might have been destroyed during the collision with the South China paleoplate. The North China Platform is a typical shallow-water carbonate platform with evaporite deposits in the central part and in the northwest Ordos region. Tremadocian to Sandbian carbonates within the platform are uniform and not well studied; thus lithofacies and biofacies have not been differentiated. Katian and Hirnantian strata are nearly absent throughout the North China Block except for the Ordos region, which remained a relatively shallow-water depositional environment throughout the Ordovician. Zhou et al. (1989) and Fu et al. (1993) studied Ordovician lithofacies and biofacies in the western margin of the North China paleoplate. Although Tremadocian and Floian strata are absent in this region, Dapingian to lower Katian strata are extensive and well studied. Two trilobite biofacies are recognized in Dapingian strata, the Eoisotelus biofacies across the North China Platform, and the Basilicid biofacies along the western margin (Fig. 10). The Eoisotelus biofacies is dominated by Eoisotelus, which constitutes up to 95% of the fauna. A few Basilicus species occur in the western open shallow-water platform region, and an IllaenidCheriruid facies belt occurs on the platform edge. The occurrence of Basilicus species in the platform region might be related to the disappearance of the Ordos Floian uplift. The lower-middle Darriwilian Basilicid trilobite biofacies occupied most of the North China Platform region, where Basilicus (Basilicus) and B. (Basiliella) occur in 95% or more of the outcrop area (Zhou et al., 1989). The Illaenid-Cheiriruid biofacies lay along the platform margin, accompanied by carbonate mudmounds. The Nileid biofacies occupies the slope facies with a trilobite fauna composed of nileids (50%), raphiophorids (20%), asaphids (15%), and endymioniids (less than 5%; Fig. 11). During the Sandbian, the Basilicid biofacies occupied most of the platform region (Zhou et al., 1989). The Illaenid-Cheiriruid biofacies (70% Illaenus) occurs along the southwest edge of the North China Platform in the shallow-water carbonate bank deposits. The Nileid biofacies (including 30% nileids, 12% Ovalocephalus, 10% raphiophorids, 10% Pseudostygina) is recorded from the shelf slope, and a graptolite facies is recorded from the shelf basin and westward down the shelf slope (Fig. 12).
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Figure 8. Sandbian–early Katian paleogeographical map of South China, showing locations of (1) Remopleurides-Ovalocephalus biofacies; (2) Ovalocephalus-Cyclopyge biofacies; (3) Cyclopyge-tricucleid biofacies; (4) Cyclopygid biofacies; (5) Cyclopygid-Girvanopyge biofacies. (A) Datianba Formation nautiloid-bearing limestones; (B) Miaopo Formation black graptolitic shales.
The distribution of trilobite biofacies of the Katian Stage record a regression (Zhou et al., 1989). The Basilicid facies occurs within a limited semi-restricted shallow water platform. The Illaenid-Cheriruid facies occurs within a platform margin carbonate mud-mound belt (Fig. 13). Most areas of the North China Platform were exposed land surfaces. Zhou et al. (1989) concluded that the distribution of the Ordovician trilobite biofacies indicates a depth control pattern. From platform to slope or basin, endemic trilobites decreased, while cosmopolitan forms increased. Faunal diversity increased from the platform region to the platform margin and reached its maximum in the slope facies and then decreased again to the basin. All the platform trilobites are endemic. The platform margin forms are infaunal, as opposed to epifaunal taxa in the slope facies. Warm water forms mainly occurred on the platform and platform margin regions, whereas a cool water fauna occurred in the open sea and deeper water regions. Zhou et al. (1989) concluded that rising sea level reached its maximum in the late Darriwilian–Sandbian because the lower-middle Darriwilian Illaenid-Cheriruid biofacies was replaced by the Nileid biofacies. However, as sea level fell during the Katian, the Illaenid-Cheriruid facies migrated westward to the platform margin.
Tarim As indicated by Zhou and Dean (1996), the Paleozoic history of Northwest China is essentially one of rifting, dispersal, and collision that involved the paleoplates of Siberia and Tarim and the orogenic belts of Kazakhstan and Tarim. Although Tarim has an appreciable cratonic core, the margins of the paleoplate are not well defined or preserved. No distinct early Paleozoic suture zones are recognized south of the Tarim paleoplate, but the Kegang fault and Qiemo-Xingxingxia strike-slip fault may form the western and eastern sections of the southern boundary, respectively (Zhou et al., 1995a; Ni et al., 2001). The northern boundary of the paleoplate evolved and changed position throughout the early Paleozoic (Zhou et al., 1996a); thus we reconstruct the Ordovician paleogeography only for the area south of the Xingeer fault (Ni et al., 2001), where stratigraphic correlations are well established (Zhou et al., 1990), allowing for confident reconstructions. Faunal evidence suggests that the Tarim paleoplate was part of northeastern Gondwana and within the tropical zone during the Ordovician (Zhou et al., 1990, 1992; Zhou et al., 1998a). Ordovician rocks are exposed widely along the northern margin
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Figure 9. Late Katian–Hirnantian paleogeographical map of South China. HR—Hirnantia fauna distribution area; PA—Paramalomena fauna distribution area; GR—continuous graptolite sequence through Ordovician–Silurian transition; FD—fine grained detrital sedimentation during Hirnantian; HH—Hunan-Hubei Submarine High; DT—Dongting Submarine High; PY—Poyang Submarine High; CW—thick clastic wedges through late Katian and Hirnantian; MU—late Katian algal-reef facies and manganiferous limestones.
of the Tarim Basin, and faunal and stratigraphic sequences in the Bachu, Kalpin, Queerqueke, and Uligezitag areas have been studied extensively (Zhou et al., 1990, 1992; Ni et al., 2001). Ordovician paleogeography was first reconstructed stage by stage on the distribution of lithofacies and biofacies as known from surface outcrops (Zhou et al., 1990, 1992). Subsequently, Zhou et al. (1995a, 1996b) and Ni et al. (2001) further elaborated on this paleogeographic framework by making use of seismic and subsurface data from a large number of borehole sections in the hinterland of the basin. In general, during the Ordovician, the vast expanse of the Tarim paleoplate was a northward-deepening, shallow-water platform, which was fringed to the north by peripheral, deeperwater facies that developed along the South Tianshan. The paleogeographic framework was strongly influenced by eustatic changes in sea level, and it is presented here in three maps (Figs. 14–16).
Tremadocian–Dapingian (Fig. 14) During this time span the southern part of Tarim was a highland. Biofacies and lithofacies of the Upper Qiulitag Group in the Bachu and Kalpin areas indicate that from the Tremadocian to the Dapingian the carbonate platform evolved from semirestricted to open marine, with a gradual reduction in dolomite and an increase in limestone (chiefly thick-bedded biocalcarenites) as sea level rose (Zhou et al., 1998b). Seismic data provide evidence of a narrow slope along the east and northeast sides of the platform (Zhou et al., 1991, 1992). As indicated by the KN1 borehole log the deposits in this region consist of thin-bedded calcilutite intercalated with allochthonous biocalcarenite (upper part of Torsuqtagh Group). To its north the slope was bounded by a stagnant shelf-basin that was developed from the late Tremadocian to the earliest Darriwilian (interval of the austrodentatus graptolite zone). Correlative strata in the central Queerqueke Mountains were named the Heituao Formation
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Figure 10. Early Middle Ordovician (Dapingian) trilobite biofacies distribution in North China, modified from Zhou et al. (1989, text-figs. 3 and 4). Localities: (1) Helanshan; (2) Miboshan; (3) Xiji; (4) Liupanshan; (5) Longxian.
by Zhong and Hao (1990). These units are characterized by carbonaceous shale and chert that are rich in graptolites. Northeast of the basin in the Uligezitag area the strata are similar to those penetrated by borehole KN1, including mainly argillaceous-banded and nodular calcilutite and biocalcilutite of the upper part of the Tremadocian–lower Darriwilian Torsuqtagh Group, and mudstone and tuffite of the upper Darriwilian-Sandbian Selikedaban Formation. Associated trilobites were assigned to the Nileid biofacies, suggesting a shelf-slope environment (Zhou et al., 1990, 1992). Geologically, this northern slope facies may have rimmed a small stable massif (Zhou et al., 1995a, 1996b).
Semi-restricted shallow-water platform Basilicid Biofacies Shallow-water carbonate mounds (platform edge) Illaenid-cheirurid Biofacies Shelf slope Nileid Biofacies
Figure 11. Late Middle Ordovician (early–middle Darriwilian) trilobite biofacies distribution in North China, modified from Zhou et al. (1989, text-figs. 5 and 6). Localities: (1) Helanshan; (2) Yinchuan; (3) Shibangou; (4) Miboshan; (5) borehole H14; (6) Xiji; (7) Liupanshan; (8) Longxian.
Darriwilian–Early Katian (Fig. 15) During the Darriwilian-Katian, the platform deepened from continued transgression, and deposition of dolomite was replaced with that of limestone, largely biocalcarenites. The northern margin of the platform shifted southward and developed carbonate buildups, including those exposed in the Bachu area. These buildups include Calathium-Receptaculites reefs in the lower Darriwilian Yijianfang Formation, nodular biocalcarenites
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Figure 12. Late Middle to early Late Ordovician (late Darriwilian to early Katian) trilobite biofacies distribution in North China, modified from Zhou et al. (1989, text-figs. 7 and 8). Localities: (1) Helanshan; (2) Yinchuan; (3) Shibangou; (4) borehole H14; (5) Chedao; (6) Xiji; (7) Guyuan; (8) Liupanshan; (9) Longxian; (10) Qishan; (11) Dongzhuang; (12) Yaoxian; (13) Fuping; (14) Xi’an.
of reef foreslope facies in the upper Darriwilian lower Tumuxiuke Formation, bioclastic banks in the Sandbian upper Tumuxiuke Formation, and algal mounds and oolitic banks in the Katian Lianglitag Formation (Zhou et al., 1990, 1992). Drill logs (Zhao et al., 2000) indicate that this platform-margin facies extends eastward as far as the Luntai area, and in the Yingmaili area (e.g.,
Shallow-water carbonate mounds (platform edge) Illaenid-cheirurid Biofacies
Figure 13. Late Ordovician (late Katian) trilobite biofacies distribution in North China, modified from Zhou et al. (1989, text-figs. 9 and 10). Localities: (1) Helanshan; (2) Xiji; (3) Guyaun; (4) Longxian; (5) Yaoxian; (6) Liquan; (7) Liupanshan.
borehole YM2) the Katian Daxikumu Formation contains mudmounds (Ni et al., 2001). The slope facies migrated southward into the Kalpin area during the early Darriwilian, as evidenced by deposition of biocalcilutite of the Dawangou Formation, which is rich in nileid and illaenid trilobites, indicating a water depth of >70 m (Zhou et al., 1998a). During the late Darriwilian–early Sandbian, as the transgression progressed, the stagnant shelf-basin migrated southward to the Kalpin area, resulting in deposition of the Saergan Formation, which is a condensed succession of carbonaceous and
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Figure 16. Late Katian (Late Ordovician) paleogeography of Tarim, modified from Zhou et al. (1995a, 1996b) and Ni et al. (2001). STL— Southern Tarim Land; LB—littoral belt; PCF—platform carbonate flat; S—slope; ST—shallow trough; PM—platform margin.
siliceous shale with abundant graptolites and trilobites. However, following the climax of transgression during the middle–late Sandbian, stagnant conditions ended abruptly, to be replaced by deposition of the partly rhythmic sediments of biocalcilutite, shale, and siltstone of the slope facies, recorded in the Kanling and Qilang Formations. The trilobite fauna in the biocalcilutite, dominated by nileids and cyclopygids, indicates water depths to 100 m. By the early Katian the slope became shallower, with deposition of the Yingan Formation, which consists of alternating beds of carbonaceous, silty, and calcareous shale intercalated with biocalcilutite with graptolites and trilobites of the Raphiophorid biofacies, which indicate a water depth of <100 m (Zhou et al., 1995b). North of the slope in the Queerqueke area, the earlier stagnant shelf-basin underwent extension and subsidence from the middle Darriwilian on and evolved into a bathyal trough that accumulated a thick flysch succession of rhythmic siltstone, shale, and thin-bedded calcarenite and calcirudite of the lower and middle Charchaq Group. Trilobites collected from the Queerqueke area are mostly mesopelagic forms (Zhou et al., 1994), which were assigned to the Cyclopygid biofacies, indicating that deposition was at depths <200 m (Zhou et al., 1990, 1992). The Uligezitag area underwent tectonic uplift and shallowing at the end of the Sandbian, and bioclastic banks of the Uligezitag Group were
intermittently developed. Trilobites in these strata belong to the Illaenid-Cheirurid biofacies, indicating a platform-margin setting (Zhou et al., 1990, 1992). Late Katian (Fig. 16) With continued regression in the Late Katian, the southern Tarim upland expanded northward, and the platform became shallower. It was fringed by a littoral facies on its southern and western marginal area, receiving deposits of sandstone and silty shale of the lower Kalpintag Formation. A complete stratigraphic succession at Tierekeawati and Kalpin was described by Ni et al. (2001). Various Ordovician sections described from the subsurface in this area by Zhao et al. (2000) indicate that the vast hinterland was covered by a carbonate tidal flat where biocalcarenite, mudstone, calcilutite, calcareous siltstone, and, uncommonly, dolomite of the Sangtamu and Yimaili Formations were deposited. To the northeast the platform was bounded by a narrow slope that was the site of turbidite deposition, as revealed by the rhythmic deposits of arenite and argillite of the lower Kalpintag Formation in cores from borehole CH1 (Ni et al., 2001). The deep trough north and east of the slope accumlated flysch deposits of the upper Charchaq Group, while carbonate buildups of the upper Uligezitag Group continued to develop in the Uligezitag region to the north. During the late Katian the trough shallowed to water
Ordovician paleogeography and tectonics of the major paleoplates of China depths of <100 m, as indicated by the appearance of the Nankinolithus trilobite fauna dominated by trinucleids (Zhou et al., 1990, 1992). Reliably dated latest Ordovician strata have been found only in a few borehole sections (e.g., TH1 and CH1) and in a few localities in the Kalpin area (Ni et al., 2001). They are mostly siliciclastic deposits of the Kalpintag Formation. Although stratigraphic evidence is sparse, it suggests that at maximum regression in the latest Ordovician, the sea covering the platform and the trough retreated substantially and shrank considerably following further northward expansion of the Southern Tarim Land and uplift of the Saarming Massif north of Tarim.
it includes Kunlun, which is the south mobile belt of Tarim, Chaidam with the Qilianshan mobile belt, and the Qinling mobile belt between South and North China. Tibet is composed of two main units, the north Qiantang Block and the south Gandise-Himalaya Block. The relationships between these crustal units reflect accretionary processes. The Qing-Qi-Kun mobile belt, which contains Ordovician to Silurian island arcs, converged with the Tarim and North China plate margins near the end of the early Paleozoic, as indicated by a Devonian terrigenous facies belt. The Qiantang Block of Tibet then converged with Tarim and Chaidam during the latest Paleozoic to earliest Mesozoic, followed by the convergence of the Gandise-Himalaya Block. Finally, the collision of the India paleoplate and the amalgamated Tibet Block during the Neogene uplifted the whole of Tibet (Geological Survey of China and Chengdu Institute of Geology and Mineral Resources, 2004). Ordovician rocks have been recorded only from the GandiseHimalaya Block, and from only two areas: One of these is at Nyalam, where Mu et al. (1973) reported Middle Ordovician strata of the 823-m-thick Jiacum Formation, which comprise shallowwater carbonates with common fossils being the nautiloids Ordosoceras, Manchuroceras, Sinoceras, and Michelinoceras, and gastropods like Maclurites. Both the nautiloids and gastropods are similar to those of the North China and Yangtze regions of South China. Similar nautiloids such as actinoceratids and Michelinoceras were recovered from the Kangmar area (Geological Survey of China and Chengdu Institute of Geology and Mineral Resources, 2004). These data indicate that Tibet was a typical east Gondwana paleoplate during the Middle Ordovician.
Tibet The Ordovician rocks and fossils of Tibet were reported first in Mu et al. (1973) from the Neilamu area of the Himalaya region. Recently, geological maps at a scale of 1:250,000 from most areas of Tibet have been completed. Based on these large projects, organized by the Ministry of Geology and Mineral Resources, a geological map at a scale of 1:1,500,000, with an explanation, was published by the Geological Survey of China and the Chengdu Institute of Geology and Mineral Resources (2004). The report and geological map provide significant new data on the nature and distribution of Ordovician rocks in Tibet. Here, we summarize the main tectonic units of Tibet and its neighboring region (Fig. 17). The long Qing-Qi-Kun mobile belt lies north of Tibet and crosses more than half of China (Wang, 1985). From west to east, 72
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Figure 17. Major tectonic units of Tibet and its neighboring region. TA—Tarim paleoplate; CH—Chaidam Block; Chn—Qilianshan mobile belt; Sq—Qinling mobile belt; SG—Songpan-Garze Block; SM—Sibumasu Block; QT— Qiangtang Block; GH—Gangdise-Himalaya Block.
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A 70-m-thick, brown, nonfossiliferous shale overlies the Jiacun Formation and might be of Late Ordovician age. It is succeeded by sandstone, black graptolitic shale, and limestone of the Silurian Shiqipo Formation (Mu et al., 1973). An Ordovician to Silurian stratigraphic succession has also been reported from Xainza, north of Lhasa. The Comprehensive Prospecting Team (1980) recognized three lithologic units in ascending order, as follows: (1) the Keerduo Formation, composed of calcareous shale, siltstone, and thin-bedded limestone with the nautilods Sinoceras and Discoceras, which indicate a Katian age; (2) the Gangmushang Formation, also composed of calcareous shale and thin-bedded limestone but with Hirnantian Normalograptus extraordinarus and N. persculptus Biozone graptolites (Mu and Ni, 1983; Zheng, 2005) and Hirnantia brachiopods (Rong and Xu, 1987); and (3) the Dewukaxia Formation, composed of graptolitic shale with early Silurian (Rhuddanian to early Aeronian) graptolites (Huang and Lu, 1983; Ni, 1987). Obviously, the Ordovician and Silurian strata from the Xainza area are generally similar to those of the Nyalam area, although the Nyalam and Xainza areas may be within different lithofacies. We conclude that the Gangdise-Himalaya Block was a platform during the Ordovician and Silurian, which we name the Lhasa Platform (Fig. 17). The oldest Ordovician rocks and fossils from the Lhasa Platform were recently reported from the Xainza area by Cheng et al. (2005), who described a 687-m-thick section of slightly metamorphosed siliciclastic strata that yielded the Early Ordovician (Floian) graptolites Tetragraptus approximatus and Didymograptellus protoindentus. This report indicates that shallow-water marine environments with a low-diversity graptolite fauna existed during the Early Ordovician in the Xainza area. The Ordovician rocks from both the Nyalam and Xainza areas of the Gandise-Himalaya Block rest unconformably on Precambrian basement (Zhu et al., 2003; Cheng et al., 2005). ACKNOWLEDGMENTS The present study was supported by the National Natural Science Foundation of China (grants 40532009, 40772002). The authors thank Prof. A.J. Boucot and Dr. Dan Goldman for valuable reviews of early drafts of the manuscript, and S.C. Finney for final editing. This is a contribution to the IGCP 503 project “Ordovician Palaeogeography and Palaeoclimate” and the Geobiodiversity Database (GBDB, www.geobiodiversity.com) project. REFERENCES CITED Bureau of Geology and Mineral Resouces of Guangxi Zhuang Autonomous Region, 1985, Regional Geology of Guangxi Zhuang Autonomous Region: Geological Memoirs of Ministry of Geology and Mineral Resources, People’s Republic of China: Beijing, Geological Publishing House, v. 1, no. 3, 853 p. (in Chinese with English summary). Bureau of Geology and Mineral Resources of Guizhou Province, 1987, Regional Geology of Guizhou Province: Geological Memoirs of Ministry of Geology and Mineral Resources, People’s Republic of China: Beijing,
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The Geological Society of America Special Paper 466 2010
Ordovician of the Siberian Platform Alexander V. Kanygin† Institute of Oil and Gas Geology, Siberian Branch of the RAS; pr. Koptuga 3, 630090 Novosibirsk, Russia Tatiana N. Koren† A.P. Karpinsky Geological Research Institute, Sredny pr. 74, 199106 St. Petersburg, Russia Anastasia G. Yadrenkina Siberian Research Institute of Geology, Geophysics and Mineral Resources, Krasny pr. 67, 630091 Novosibirsk, Russia Alexander V. Timokhin Institute of Oil and Gas Geology, Siberian Branch of the RAS; pr. Koptuga 3, 630090 Novosibirsk, Russia Oleg V. Sychev Siberian Research Institute of Geology, Geophysics and Mineral Resources, Krasny pr. 67, 630091 Novosibirsk, Russia Tatiana Yu. Tolmacheva A.P. Karpinsky Geological Research Institute, Sredny pr. 74, 199106 St. Petersburg, Russia
ABSTRACT The Ordovician of the Siberian Platform exhibits a wide range of shallow-water facies in thick, normal marine and lagoonal carbonate-clastic sequences. It is illustrated herein by two west–east geological profiles across the Siberian Platform. Stratigraphic and paleontologic studies in this area led to subdivision of the Ordovician into 12 regional stages. Their boundaries correspond to changes in sedimentology and in taxonomic composition of a benthic fauna dominated by brachiopods and ostracodes. In most cases the precise correlation of regional with global stages is difficult, mainly because of the endemic character and sparse distribution of the East Siberian fauna in shallow-water facies, as well as the sporadic occurrences of graptolites. Two important events at the beginning of the Volginian (the Hustedograptus teretiusculus Zone) and Chertovskian (Nemagraptus gracilis Zone) are documented in the Ordovician of the East Siberian Basin. They include large-scale transgressions and significant biotic changes, and are used as the most reliable levels for correlation in the vast area of the Siberian Platform and the Verkhoyano-Chukotka region. The distribution of lithofacies and dominant members of faunal communities are figured on three maps for the Nyaian (Early Ordovician), Volginian (Middle Ordovician), and Baksanian (Late Ordovician). Comparative analysis of the Ordovician benthic faunas of the Siberian Platform, Taimyr, Novosibirsk Islands, and the Verkhoyano-Chukotka
†
E-mails:
[email protected];
[email protected].
Kanygin, A.V., Koren, T.N., Yadrenkina, A.G., Timokhin, A.V., Sychev, O.V., and Tolmacheva, T.Yu., 2010, Ordovician of the Siberian Platform, in Finney, S.C., and Berry, W.B.N., eds., The Ordovician Earth System: Geological Society of America Special Paper 466, p. 105–117, doi: 10.1130/2010.2466(07). For permission to copy, contact
[email protected]. ©2010 The Geological Society of America. All rights reserved.
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Kanygin et al. region shows that they all belong to the Kolymo-Siberian paleobiogeographic province. Available data confirm the location of the Siberian paleocontinent in the Southern Hemisphere during most of the Ordovician and an inverted orientation relative to its present position.
INTRODUCTION In East Siberia the Ordovician is widely distributed in several large tectonic megastructures. These include the Siberian Platform, the adjacent Altai-Sayan and Verkhoyano-Chukotka areas, the Taimyr, and the Novosibirsk and Severnaya Zemlya Islands (Fig. 1). Ordovician sedimentary successions of the Siberian Platform are represented by a wide spectrum of shallow-water facies that range from carbonate and terrigenous marine deposits to restricted lagoonal facies. Ordovician sedimentary rocks are well exposed on the margins of the Tunguska and Vilyuy Syneclises and in the Irkutsk Amphitheater, where the key Ordovician sections are present. A majority of the key sections are confined to the area between the Yenisei and Lena Rivers, where they crop out along small rivers and brooks. In the central part of the Siberian Platform, as well as in the Yenisey-Khatanga Depression, the Ordovician is overlain by Upper Paleozoic and Mesozoic sedimentary rocks, and can be studied only from boreholes (Fig. 1). The first data on the Siberian lower Paleozoic, including the Ordovician, resulted from early geographical expeditions in the middle of the nineteenth century. During the next hundred years the accumulation of information on the lower Paleozoic was slow and irregular, and the data from that period are only of historical interest. Large-scale mapping and drilling projects in the 1950s–1960s stimulated intensive geological and stratigraphical studies in the region and led to a series of monographs on Ordovician stratigraphy and paleontology. A comprehensive summary of Ordovician paleontology and stratigraphy of the Siberian Platform was published by Nikiforova and Andreeva (1961). It made possible the compilation of the first regional chart for this territory based on the monographic description of brachiopods. It also evaluated the biostratigraphic potential of such benthic groups as tabulate corals, stromatoporoids, crinoids, bryozoans, pelecypods, gastropods, nautiloids, ostracodes, trilobites, and some rare exotic fossils. This monograph remains a highly useful synthesis of regional paleontology and biostratigraphy. The most numerous and widely distributed groups were described later. Among these are the bryozoans (Nekhoroshev, 1961), crinoids (Eltysheva, 1960), trilobites (Maksimova, 1962; Rozova, 1968), gastropods (Vostokova, 1962), nautiloids (Balashev, 1962), ostracodes (Ivanova, 1959a, 1959b), and graptolites (Obut and Sobolevskaya, 1967; Obut et al., 1984). The first conodont studies go back to the mid-1960s (Moskalenko, 1970, 1973; Abaimova, 1975). A summary of the Ordovician stratigraphy and paleontology of the Siberian Platform was published in a book edited by Sokolov and Tesakov (1975). Since 1972, new integrated studies of the Ordovician of the Siberian Platform were carried out by specialists on paleontol-
ogy, stratigraphy, and sedimentology from different institutions in Novosibirsk and Saint Petersburg. These studies resulted in a series of monographs and reports (Moskalenko et al., 1978; Yadrenkina et al., 1978; Kanygin et al., 1980, 1982, 1984a, 1984b, 1988, 1989, 2007; Sokolov (ed.), 1982; Sennikov, 1996; Tesakov et al., 2003; and others). They contain new data on the stratigraphic distribution of different groups of macro- and microfossils, rare finds of graptolites, and the first information on chitinozoan and acritarch occurrences in the Ordovician of the Siberian Platform. A vast amount of data on Ordovician paleontology, stratigraphy, and sedimentology was accumulated from studies of numerous sections that crop out along the rivers and that are known from drill cores (Fig. 1). However, more extensive taxonomic and biostratigraphic studies are needed for more precise dating and correlation of the regional biostratigraphic and lithological units with global subdivisions. The main goal of this report is to provide a general summary of the Ordovician stratigraphy, facies, and biota of the Siberian Platform. REGIONAL CHRONOSTRATIGRAPHY Ordovician carbonate-terrigenous sequences represent a significant part of the thick sedimentary cover of the Siberian Platform. The Cambrian-Ordovician boundary interval is characterized by monotonous carbonates, whereas the Lower Silurian boundary is easily recognizable in all sections, as it coincides with a remarkable hiatus that has a different duration within the Upper Ordovician to lower Llandovery in different parts of the Siberian Platform. The Ordovician of the Siberian Platform is subdivided into twelve regional stages (“stratohorizons” in Russian stratigraphical nomenclature) (Fig. 2). In general, the boundaries of a majority of the regional stages correspond to sedimentological changes and to changes in taxonomic composition of benthic faunas dominated by brachiopods and ostracodes. The precise correlation of regional and global stage boundaries is difficult owing to the endemic character and localized distribution of the East Siberian fauna in shallow-water facies. Sporadic occurrences of graptolites in the Ordovician of the Siberian Platform are the other serious impediment for establishing precise ties between local and global chronostratigraphic units. This problem can be illustrated by difficulties in definition of the Cambrian-Ordovician boundary in the Siberian sections and boreholes. The global standard for definition of the base of the Ordovician coincides with the lowest appearance of the conodont Iapetognathus fluctivagus Nicoll et al., which is 4.8 m below the earliest planktic graptolites at the standard in western Newfoundland
Ordovician of the Siberian Platform
Norilsk
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Tunguska syneclise
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Reference sections
Ba ke La
AltaiSayan fold area
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Anga
Irkutsk
Exposed Ordovician deposits
Russian Frontier
Boreholes 0 Â
VerkhoyanoChukotka area
Yenisey-Khatanga depression
Ust’Kut West-Siberian plate
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Kotel’ny Island
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Figure 1. Sketch map showing distribution of Ordovician rocks and tectonic structures of the Siberian Platform and surrounding areas.
(Cooper et al., 2001). For zonal subdivision and correlation of the boundary interval, successive species of Cordylodus, C. andresi, C. proavus, C. lindstroemi, and C. angulatus, are used as an additional biostratigraphic aid. According to this sequence, the lower boundary of the Ordovician is placed within the lower part of the C. lindstroemi Zone. In the shallow-water carbonate sequences of the Cambrian-Ordovician boundary interval of East Siberia conodonts are rare, and I. fluctivagus has not been found. All representatives of the Cordylodus lineage are known in the Siberian sections. C. proavus occurs in the Loparian and in the lower Nyaian regional Stages; C. angulatus is identified from the Nyaian
and Kimaian regional Stages in northeast Siberia (Tesakov et al., 2003). Recently C. lindstroemi was recovered from sediments of the upper Nyaian regional Stage in the key section along the Kulumbe River (Tolmacheva and Abaimova, 2009). In 1962 a single specimen of dendroid graptolite, later described as Dictyonema flabeliforme kulumbeense Obut et Sobolevskaya, was found in carbonates assigned to the Mansian regional Stage at the Kulumbe River section (Obut and Sobolevskaya, 1967). This species has also been identified from other localities, mainly from drill cores, in a stratigraphic interval correlated with the Mansian and Loparian regional Stages. These
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British graptolite zones
Katian Sandbian
D. clingani
Upper
P. linearis
D. caudatus
Darriwilian Dapingian Floian
Tremadocian
Lower
Bellimurina sibirica
Dicellograptus, Glyptograptus, Orthograptus
Acanthodina nobilis
Paraorthograptus ?pacificus
Dolborella bifunctata
Ozarkodina dolborica Acanthocordylodus festus
Maakina parvuliformis Leptellina carinata
Parajonesites notabilis
Belodina compressa Culumbodina mangazeica
Bodenia aspera
Cahabagnathus sweeti - Phragmodus inflexus
Quadrilobella recta
Ptiloconus anomalis
Baksanian
Chertovskian
Mimella panna
Isalaux
KirenskoKudrian Lenatoechia
Mukteian
D. artus
Vikhorevian
Kimaian I. gibberulus
Evenkina
Soanella maslovi
Phragmodus flexuosus
Rhyselasma
Apheorthis submelita
Tetralobula
Mansian
Mastigograptus datzenkoi Orthograptus propinquus
Coleodus mirabilis
Biolgina
Aparhites clivosus
Histiodella angulata
Shumardia Paenebeltella
Scandodus warendensis - Scandodus pseudoquadratus
Eoapatokephalus -Ijacephalus
Loxodus bransoni - Acodus oneotensis
C. angulatus
Loparian
Glyptograptus siccatus, Oepikograptus bekkeri, Amplexograptus fallax, Glyptograptus euglyphus
Cherskiella Cardiodella lyrata notanilis Ventrigyrus Polyplacognathus intricatus angarense
Leontiella
Ugorian
Nyaian
Homotelus
Graptolite assemblages
Aphelognathus pyramidalis
Boreadorthis
A. cucullus (E. hirundo)
Cambrian Upper
T
Dolborian
D. murchisoni
T. phyllograptoides H. copiosus A. murrayi Kiaerograptus Adelograptus Rh. f. anglica A. matanensis Rh. flabelliforme s.l.
Conodonts
Evenkorhynchia dichotomians Bumastus evenkiensis
Volginian
E. simulans C. varicosus
Brachiopods Trilobites Ostracodes
Nirundian
D. multidens
N. gracilis
Biostratigraphic zones
Sea level curve R
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D. morrisi
H. teretiusculus
Middle
Regional Stage
N.? persculptus
N. extraordinarius P. pacificus D. complexus D. complanatus
Ordovician
Regional stratigraphic chart
Beds with Angarella
Series
Hirn. Stage
System
International Stratigraphic scale
LoparellaNotaiella DolgeulomaSaukiella
Dictyonema omnutachense
Callograptus kravtsovi, Syringotaenia bystrovi
Rb. flabelliforme anglica
Dictyonema kulumbense, Cordylodus Airograptus furciferus, proavus Callograptus staufferi, Proconodontus - Dendrograptus hallianus Eoconodontus
Figure 2. Regional chronostratigraphy and sea-level curve for Ordovician strata of the Siberian Platform, showing correlation with the new global chronostratigraphic classification for the Ordovician System.
finds were taken as the main biostratigraphic criteria for considering both regional stages to be of Tremadocian Age. A taxonomic revision of dendroid graptolites and their assignment to the genera Dictyonema (benthic forms) and Rhabdinopora (planktic forms) by Erdtmann (1982, p. 129) concluded that D. flabeliforme kulum-
beense is a benthic dendroid graptolite and of Cambrian age. The planktic Rhabdinopora anglica, R. parabola (Bulman), and R. flabeliformis graptolithinum Kjerulf; the benthic Callograptus kravtsovi (Obut and Sobolevskaya); and others of Tremadocian Age were identified in several drill cores on the Siberian Platform
Ordovician of the Siberian Platform (Obut et al., 1984; Sennikov, 1996). However, in all cases, a correlation of the graptolite-bearing deposits of the drill cores with the Loparian and Mansian regional stages was uncertain. We came to the conclusion that the lower boundary of the Ordovician on the Siberian Platform can be defined at a level not lower than the base of the Nyaian regional Stage. The occurrence of only one species of Cordylodus, C. proavus, in the Loparian and Mansian regional Stages in the stratotype section of these units on the Kulumbe River is used as the criterion for assigning them to the Upper Cambrian. However, much work is still needed on the taxonomic revision of Siberian Cordylodus species to solve this problem. Recognition of the global Lower-Middle Ordovician boundary based on established biostratigraphical markers (Wang et al., 2005) is not possible in shallow-water facies of the Siberian Platform, owing to the presence of highly endemic conodonts and an absence of graptolites. This level roughly corresponds to the boundary of the Ugorian and Kimaian regional Stages. The most distinct correlative level within the Middle Ordovician of the Siberian Platform and surrounding areas is the base of the Volginian regional Stage. It corresponds to the beginning of a transgression that brought significant biotic changes across the East Siberian Basin. The lower Volginian boundary can be traced in such different parts of the Verkhoyano-Chukotka region as the Verkhoyan Mountains, Sellenjakh Ridge, and the Kolyma River Basin, as well as on the Taimyr Peninsula and Kotel’ny Island. In sections on the Siberian Platform, this level is marked by the appearance of numerous new taxa of brachiopods (e.g., Evenkina lenaica [Girardi], E. anabarensis Andreeva, Atelelasma peregrinum Andreeva, Hesperorthis ignicula [Raum], H. brachiophorus [Cooper]) and ostracodes (Soanella maslovi Ivanova, Sibiritella rara Ivanova, S. costata Ivanova, and Egorovella defecta Ivanova), which are widely used for regional correlation. For the first time the key species of other faunal groups appear in shelf communities. Among the latter are the trilobites Homotelus lenaensis Maksimova and Ceraurinella biformis Maksimova; the crinoid Kalgacrinus kalgiensis (Yeltysheva); the tabulate coral Billingsaria lepida Sokolov; the bryozoans Ceramopora spongiosa Bassler, Dianulites petropolitana (Pander), and Hallopora dubia Loeblich; as well as the conodont Phragmodus flexuosus Moskalenko. Orthograptus propinquus (Hadding) is the only graptolite found in the Volginian regional Stage. The Volginian and Kurensko-Kudrian association of brachiopods and ostracodes in the carbonate facies of the Siberian Platform shows very close affinities to those of the Lachug regional Stage of the Verkhoyano-Chukotka region, and this allows an interregional correlation. The correspondence of this stratigraphic interval with the Hustedograptus teretiusculus graptolite Zone of the upper Darriwilian is based on the occurrences of Hustedograptus aff. teretiusculus (Hisinger) in the terrigenous deposits of the Lachug Stage in the Kolyma River Basin (Oradovskaya, 1988). The base of the Upper Ordovician corresponds with the lower boundary of the Chertovskian regional Stage and coincides
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with the most distinctive transgressive event in the Ordovician of East Siberia (Fig. 2). The Chertovskian boundary is characterized by the appearance of a diverse assemblage that includes different faunal groups and is persistent in different facies settings. Brachiopods and ostracodes are the most important fossils for stratigraphical subdivision and for regional and interregional correlation. The lower boundary of the stage is well defined by the appearance of numerous brachiopods (Mimella panna Nikiforova, Oepikina tojoni Andreeva, Rostricelulla transversa Cooper) and ostracodes (Bodenia aspera Ivanova, Egorovella captiosa Ivanova, Coellochilina laccochilinoides, and other species). Some trilobites and conodonts (e.g., Isalaux stricta [Khramova], Monorakos lopatini Schmidt, Phragmodus inflexus Stauffer, and Polyplacognathus sweeti Bergström) also appear. The lower Chertovskian boundary coincides with the base of the C. sweeti–Ph. inflexus Zone (Moskalenko, 1994). The Chertovskian regional Stage is the best stratigraphic interval for interregional correlation, as it is based on a characteristic association of brachiopods and ostracodes that is widely distributed in northeastern Asia. The fact that the Chertovskian faunal assemblage is readily recognizable in sections of the lower Kharkindzha regional Stage (Oradovskaya, 1988) in the Verkhoyano-Chukotka region is of particular importance. In the Kolyma River Basin this fauna is associated with graptolites of the Nemagraptus gracilis Zone. The correlation with this globally traceable graptolite zone suggests that the base of the Chertovskian Stage coincides with the Middle-Upper Ordovician boundary. The recognition and regional correlation of the Baksanian, Dolborian, Nirundian, and Burian Stages is based on distinct changes in benthic communities in shallow-water carbonate facies. However, even indirect correlation of the Upper Ordovician Siberian stages with global chronostratigraphic units is difficult. The Upper Ordovician graptolite successions established in the upper Kharkindzha, Padun, and Tirekhtjakh regional stages of the Kolyma River Basin (Oradovskaya, 1988) cannot be used for correlation because of the extremely rare and patchy graptolite occurrences on the Siberian Platform (Sennikov, 1996), and different benthic associations are characteristic for this interval in different parts of eastern Siberia and northeastern Russia. REGIONAL FACIES The Siberian Platform was covered by a warm sea characterized by shallow-water carbonate sedimentation and the accumulation of contrasting terrigenous-carbonate deposits in littoral and sublittoral facies. Two especially distinctive transgressiveregressive phases are recognized in the Ordovician of the Siberian Platform. The first corresponds with the Early and early Middle Ordovician, and the second to the late Middle and Late Ordovician. The boundary between these two phases coincides with the beginning of the first large-scale transgression at the base of the Volginian regional Stage (the base of the H. teretiusculus graptolite Zone). The Ordovician sediments that accumulated during these phases differ in lithology, thickness, and facies, and
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these differences were previously used to subdivide the Ordovician of the Siberian Platform into two series (“otdel,” Sokolov and Tesakov, 1975). Lithology and sedimentary structures of local stratigraphic units, as well as their lateral relationships, are shown in two geological profiles that cross the Siberian Platform from west to east and from northwest to southeast (Figs. 1, 3A, 4A). In addition to these profiles, other figures show the correlation of lithostratigraphic units with regional stages (Figs. 3B, 4B). In the Early and early Middle Ordovician, carbonate sedimentation was dominant, whereas red-colored deposits and evaporites were subordinate (Figs. 3 and 4). The rate of accumulation of compacted sedimentary rocks during that time was 5–10 times higher than in the second phase of the Ordovician. The general character of the Ordovician sedimentary patterns was inherited from the Cambrian. The similarity is seen in the same bathymetric differentiation of the basin, the lithology of the sedimentary rocks, and their depositional rates. This resulted in continuous carbonate sedimentation with no distinctive lithological changes through the Cambrian–Ordovician transition on the Siberian Platform. The Lower Ordovician fauna is also quite similar to that of the Upper Cambrian, and short-ranged species and monospecific associations are widely distributed. During the Early Ordovician, cyanobacteria produced abundant stromatolites and bioherms. In the Early Ordovician before the Volginian taxonomic explosion, some exotic fossils of uncertain systematic position appeared almost everywhere in the shallow-water areas of the Siberian Platform. Among these are Clyptolichinaria, Tolmachovia, Soanites, and Angarella. Some of these short-lived taxa occurred on other continents at the same time (Zhe, 1983; Gutiérrez-Marco, 1997). The distribution of lithofacies and dominant members of fossil assemblages in the Early and early Middle Ordovician are shown in Figure 5A. This was a time that recorded a prevalence of shallow-water marine environments with numerous nearshore islands. The Ordovician landmasses were situated to the south and southwest of the Siberian Platform, where coarse-grained sandy deposits with numerous inarticulate brachiopods, stromatolitic limestones, and clayey-silty dolostones were widespread (Fig. 5A). The beginning of the Nyaian regional Stage was characterized by a small-scale transgression. Along the southern and southwestern margins of the Siberian Platform, mainly terrigenous sediments accumulated (Fig. 5A). These sediments yield rare and monotonous faunal assemblages with a strong dominance of inarticulate brachiopods. These successions were formed in relatively shallow-water environments (from several to tens of meters in depth) characterized by low salinity owing to fresh-water runoff from the land. Northward, the Lower Ordovician terrigenous sediments were gradually replaced by carbonates deposited in open shelf environments. Here, the sedimentary successions are composed of limestones with numerous and diverse trilobites, conodonts, brachiopods, gastropods, and monoplacophorans. In this part of the basin the water depth was not more than a few tens
of meters, and hydrodynamic activity was generally low. Similar environments were typical for the northern Siberian Platform, where thick successions of carbonates accumulated. However, these carbonates contain faunal assemblages different from those of the western part (Fig. 5A). Even shallower environments with variable salinity and hydrodynamic conditions were typical for the central Siberian Platform in the Early Ordovician. Shallowwater carbonates with a subordinate amount of intertidal sediments dominate in the successions with numerous stromatolite buildups. Terrigenous strata became more abundant toward the Khatanga Land (Fig. 5A). The Ugorian and Kimaian regional stages are characterized by a continued large-scale regression that started in the early Nyaian (Fig. 2). The area of terrigenous sedimentation significantly enlarged and occupied almost the entire territory of the Irkutsk Amphitheater. The insignificant transgression during the early Kimaian resulted in the brief appearance of marine environments in the northern part of the platform. The significant drop in sea level that followed the Kimaian transgression led to denudation of the Kimaian sediments in the eastern and northeastern Siberian Platform (Fig. 5B). During the second transgressive-regressive phase, which corresponded to the late Middle Ordovician and Late Ordovician, the Siberian Platform was covered by a shallow-water, epicontinental, semiclosed sea that, in general, occupied the same region as in the Early Ordovician. However, from the beginning of the Volginian, the sedimentation of the basin significantly changed from the previous phase. Carbonate deposition decreased, whereas terrigenous sedimentation became prevalent (Fig. 5B). Stromatolite buildups and dolostones became rare and completely disappeared by the end of the Middle Ordovician. Two large-scale regressions during the Kirensko-Kudrian and late Chertovskian to Baksanian led to a denudation of the vast Khatanga Land in the southern Siberian Platform. The sedimentary succession of these time intervals contains numerous weathering crusts that usually coincide with more or less continuous stratigraphic gaps. Levels corresponding to the subsequent transgressive series at the bases of the Chertovskian and Dolborian regional Stages contain phosphorite nodules, pebbles, and phosphatizated organic remains. The base of the Volginian regional Stage is marked by the taxonomic “explosion” of the East Siberian biota (Fig. 6). From this time, the shallow-water environments became densely populated by dendroid graptolites, ostracodes, conodonts, and nautiloids. Neritic communities also underwent an increase in species diversity and population density. The rare exotic groups disappeared, whereas new groups of such suspension-feeders as bryozoans, crinoids, tabulate corals, and stromatoporoids colonized the bottom habitats. The evolution of the ostracodes was characterized by an expansion of monospecific assemblages into shallow-shelf biotopes. The ostracodes reached a high level of taxonomic differentiation and became dominant in benthic faunal communities. The taxonomic “explosion” in the lowermost part of the Volginian regional Stage can be traced through different parts of the Siberian Platform as well as in the Verkhoyano-Chukotka region
Ordovician of the Siberian Platform
A
Cht-3
Ng-1 Pm-1,2 Tt-1 Uch-1 Ttn-1 Vv-1 To-2 Nirunda Fm Dolbor Fm
Nerutchanda Fm
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Nirunda Fm
Oyusut Fm
Mangazea Fm
Stan Fm
Tura Fm
LEGEND
Kholokhit Fm
LEGEND
Ust’Munduika Fm
Facies boundaries
a b c
Stratigraphic unconformity
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Tt-1 Uch-1 Ttn-1 Vv-1 To-2
Ng-1 Pm-1
Ordovician
Nirundian Dolbo1431 1075 987 rian 283 276 Neruchanda Fm Baksanian 302 303 1047 1512 Cherto1153 vskian
1267 1199 1423 2200
1047
1085
792 656
1100
817 677
1145
855 717
Baitah Fm Oyusut Fm Khar’yalakh Fm
326 364
1283
upper subform.
968 833
Stan Fm
lower subform.
2350
464
Ust'stolbovaya Fm
485
1512
1372 1267
Malykaya beds
2815
1503 2302
1244 1244
1534 1342 1561 1534 13421561
Tura Fm 2500
2985 1283
1093
Babkino Fm 1577 1443 1940
2638
3157
968 833
1592
1200 1066
1998
1388 1131
485
Balyktakh Fm
1297
Ust'munduika Fm
303
2392 2453
2325
Nyaian
Mansian
Dolbor Fm
1231 1457 2262 2550
1347 1347
Baikit Fm
Loparian
US-292
57S
427
1166 1093 1607
2355
2392
Gypsiferous beds Red beds
Mangazea Fm
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Ugorian
Cambrian
1240
KirenskoKudrian
Volginian
1113 1272 1790
Clayey and carbonate deposits
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Nirunda Fm
1100 1241 1752
Sandy, silty, clayey deposits
Ogogut Fm
Chamba Fm
Burian
Kimaian
Ks-1 Uil-251 Hl-1
Kr-1
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a-limestone b-dolomite c-clayey dolomite Sand dominated deposits oncoids (oolites) stromatolites
1200 m
0 km
ds aya be Malyk Sokhsolokh Fm
Oldondo Fm
Babkino Fm
Boreholes
Mukteian Vikhorevian
lakh Fm
Khar’ya
Balyktakh Fm
400 m
S1
US-292
Ust’Stolbovaya Fm
Baikit Fm
Regional Dk-1 Sht-3 stage Moierokian
B
57s
Ks-1 Uil-1 Hl-1
Kr-1
Sokhsolokh Fm 2662
Oldondo Fm
Kholokhit Fm 321
1172
1701
1540
Ust'peladka Fm
1805 1950 2218 2920
2023 3588
1420 1171
Evenkia Fm
712
2956
Markha Fm
Figure 3. (A) Cross section of Ordovician strata of the Siberian Platform along line A–B (see Fig. 1), showing lithological composition and lateral relationships of lithostratigraphic units. (B) Correlation of lithostratigraphic units, with regional stages (Fig. 3A).
(Sette Daban Range, Selenjakh Ridge, Kolyma River Basin), in Taimyr, and in the Novosibirsk Islands. The base of the Chertovskian regional Stage is the next stratigraphic level that shows an extensive transgressive event accompanied by changes in the East Siberian biota at the species level (Fig. 6). Large bioherms built by algae, tabulate and rugose
corals, bryozoans, stromatoporoids, and crinoids occurred in many regions of East Siberia. Trilobites became significantly less abundant than earlier. On the vast territory of the Siberian Platform, normal marine environments were partly reduced, while the lagoon settings, populated only by monospecific ostracode associations, became widespread (Fig. 5C).
C Krivaya Luka Fm
Khar’yalakh Fm
R. Lena vil. Polovinka D’yukte
Oyusut Fm
Gr4/2, 13-146
Totchinoye Fm
Sytykan Fm
Ch-1
Sokhsolokh Fm
upper subform.
Boreholes
middle subform.
Facies boundaries
Reference sections Sandy, silty, clayey deposits Clayey and carbonate deposits Gypsiferous beds Red beds
Regional R. Lena stage vil. Polovinka D'yukte Melichan Fm S1 Moierokian
13/146
0 km
Ogogut Fm
74 80
Mukteian
108
Kylakh Fm
40 81
Zagornyi Fm
?
Dzerom Fm
110 223
?
? 135
? 301
193
Ugorian
Cambrian
Nyaian Loparian
upper subform. 70 m
Gr-4/2
lower subform. Mansian 75 m Verkhnelenskaya Fm
193
102
0
0
Sokhsolokh Fm
upper subform.
m hF
59
260
k
kta
ly Ba
upper subform.
lower subform.
145
Kholomolokh Fm
170
Angir Fm 207
Guragir Fm 350
lower subform.
Il’tyk Fm
? Tchirinda Fm
Oldondo Fm
Amarkan Fm
middle subform.
893
52
Totchinoye Fm
Kotchakan Fm
Vikhorevian Kimaian
?
Moyerokan Fm
Sytykan Fm 150
80
183 183
114
lower subform. 91 m
Krivaya Luka Fm
Ordovician
Volginian
Chamba Fm
Tchanganda Fm
160
upper subform. Stan Fm
Dolborian
KirenskoKudrian
Kharyalakh Fm
SP-3
40
Nirundian
Baksanian Chertovskian
Tukalanda Fm IVmember
Moyeronskaya Fm
Meik Fm
Oyusut Fm
lower subform.
Ch-1
10 Burian
Uyigur Fm
200 km
415N
D 1-4
Il’tyk Fm lower subform. upper subform.
upper subform. middle subform. lower subform.
Tchirinda Fm
a-limestone b-dolomite c-clayey dolomite Sand dominated deposits oncoids (oolites) stromatolites
upper subform.
lower subform.
Stratigraphic unconformity
a b c
Fm Zagornyi n Fm Fm ka ar Am ir Ang Guragir Fm
Moyerokan Fm
Kotchakan Fm
LEGEND
400 m
SP-3
Dzerom Fm
Oldondo Fm
Balyktakh Stan Fm Fm
D
Tchanganda Fm
415N
102, 103
Stan Fm
Kylakh Fm
Uyigur Fm
1014 1366
213
Mapkha Fm.
550
1223
Sanat mb.
Tukalanda Fm
1419
Figure 4. (A) Cross section of Ordovician strata of the Siberian Platform along line C–D (see Fig. 1), showing lithological composition and lateral relationships of lithostratigraphic units. (B) Correlation of lithostratigraphic units, with the regional stages shown on the cross section (Fig. 4A).
Norilsk
Norilsk
Anabar anticlise
Yen i
Yen i
sey
sey
Ko
tu
y
Anabar anticlise
Tunguska syneclise
Tunguska syneclise
Mirny
Tura
Vilyuy syneclise
Vilyuy syneclise Mirny
Tura
Khatanga Land
Khatanga Land Angara
Angara 0
L
0
300 km
a en
L
Krasnoyarsk
300 km
a en
Krasnoyarsk
N
Irkutsk amphitheater
Irkutsk amphitheater
A
Irkutsk
Denudation of carbonates
sey
ka
Kurey
Anabar anticlise
B
Irkutsk
Denudation of terrigenous deposits
Norilsk
N
Dark gray limestones of outer shelf
Facies boundaries
Yen i
Coastal sands
Tunguska syneclise
Oolitic, stromatolitic and clayey-silty limestones and dolostones of shallow-water shelf
Tura
Vilyuy syneclise Mirny
Silty and clayey dolostones, dolomitic marls (domerites) of inner shallow-water shelf
Khatanga Land
Limestones, marls and siltstones of shallow shelf with low-energy hydrodynamic regime
Angara 0 Le
Krasnoyarsk
Irkutsk amphitheater
Irkutsk
300 km
na
Variegated sand-siltclayey deposits of shallow-water shelf Variegated limestones, bioclastic marls and siltstones of shallowwater shelf
N
C
Dark gray claystones, marls and limestones of deep-water shelf
Biofacies distribution, presumable biofacies distribution
Inarticulate brachiopods Trilobites Bivalves Ostracodes Conodonts Brachiopods Graptolites Stromatolitic build-ups
Figure 5. Schematic maps showing facies distribution and dominant fossils in the Siberian Platform for Ordovician time intervals: (A) Nyaian, (B) Volginian, (C) Baksanian.
Kanygin et al.
Chitinozoans
Acritarchs
Graptolites
Conodonts
Nautiloids
Trilobites
Ostracodes
Dolborian
Gastropods
Nirundian
Brachiopods
Katian
Burian
Rugose corals
Moierokan
Hirnantian
Bryozoans
Llandovery
FAUNAL GROUPS
Tabulate corals
Regional stages
Crinoids
Stage/ Series
Stromatoporoids
SIL. System
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Baksanian Chertovskian KirenskianKudrinian Darriwilian
ORDOVICIAN
Sandbian
Volginian Mukteian Vikhorevian
CAMBRIAN
Dapingian
Kimaian
Floian
Ugorian
Tremadocian
Nyaian Loparian Mansian
Figure 6. Biodiversity dynamics of selected taxonomic groups through the Ordovician of the Kolyma-Siberian Province.
In the Late Ordovician, numerous tabulate and rugose corals, heliolitids, and bryozoans colonized the shallow-water basin. Trilobites, ostracodes, and conodonts were numerous and diverse, however, most of them were endemic. By the end of the Ordovician the central part of the Siberian Platform was an area with a wide distribution of evaporite deposits represented by intercalated gypsum and siltstone beds. The Silurian history of the region started with a major sea-level rise and an accumulation of carbonate successions with abundant and diverse benthic faunas. At the same time, siltstones with Rhuddanian graptolites were deposited in depressions. PALEOBIOGEOGRAPHY By the middle of the twentieth century, it became clear that the Ordovician fauna of Eastern Siberia differs significantly from the contemporaneous biotas of Europe and North America. At that time the Siberian Platform was regarded as a separate Siberian biogeographic province that embraced only the platform (Nikiforova, 1955, Nikiforova and Andreeva, 1961). Later, a wider dis-
tribution of Siberian faunal assemblages became evident, and the Taimyr, the Novosibirsk Islands, and the Verkhoyano-Chukotka region were included in the Siberian biogeographic province. The wider areal extent of the Siberian biogeographic province was initially demonstrated by ostracode studies and later by brachiopod, trilobite, and conodont data (Kanygin, 1967, 1971; Rozman, 1970). Ostracodes are generally numerous and diverse in all carbonates of the East Siberian Ordovician. Detailed studies of the taxonomic composition of these assemblages and their relative species abundance show a consistent dominance of the same taxa in all facies zones. Figure 7 shows the distribution of dominant ostracode species at several localities on the Siberian Platform and surrounding fold belts in the H. teretiusculus and N. gracilis graptolite zones. The ecological structure of populations changed insignificantly during the Ordovician. In all regions the ostracode associations show a persistent predominance of particular species in association with other diverse but rare elements. At present, there is no doubt that the Siberian Platform, Taimyr, the Novosibirsk Islands, and the Verkhoyano-Chukotka region belonged to one biogeographic province in the Ordovician. However, the rank
Ordovician of the Siberian Platform
115
Chukotsk Peninsula Ko lym
Laptev Sea
a Ri ve
Barents Sea
Pacific Ocean
r
Taimyr
na
Ri
ve
r
Okhotsk Sea
e nis Ye
yR
Le
ive
r
Dominated species Coellochilina laccochilinoides Bodenia aspera Egorovella captiosa
A
Ko
Chukotsk Peninsula
lym
Laptev Sea
a ve Ri
Barents Sea
Pacific Ocean
r
Taimyr
na
Ri
ve
r
Okhotsk Sea
Le
r ive yR e nis Ye
Dominated species Soanella maslovi Egorovella defecta Egorovella cuneata
B Figure 7. Schematic maps showing diagrams of structures of ostracode communities in several localities of the Siberian Platform and surrounding fold areas for the H. teretiusculus (A) and N. gracilis (B) graptolite zones.
of the province is questionable. It is considered either as a separate Kolymo-Siberian biogeographic province (Kanygin et al., 1984b; Kanygin, 2001) or as a part of the Canadian-Siberian biogeographic belt (Rozman, 1970). Studies of the faunal distribution within the region show that the biogeographic ties within the Kolymo-Siberian province were relatively stable during the Ordovician. The similarity of different parts of the province increased during transgressive episodes when an identical benthic faunas populated similar bathymetric zones. The transgressive maxima coincide with the H. teretiusculus and N. gracilis graptolite zones (Fig. 2). The resemblance of some components of the brachiopod, trilobite, and conodont
assemblages of Siberia with contemporaneous faunas in Laurentia was the basis for recognition of the Canadian-Siberian biogeographic belt (Rozman, 1970). During the regressions the faunas became more endemic, and this can be shown by the endemic ostracode assemblages typical of the Kurensko-Kudrian and Baksanian regional Stages. Dissimilarities in faunal composition between different parts of the Kolymo-Siberian province are best explained by ecological rather than climatic parameters. The biogeographical unity of the Siberian Platform and the Verkhoyano-Chukotka fold area during the Early and Middle Ordovician indicates that the whole territory was a single sedimentary basin.
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PALEOMAGNETIC DATA AND LOCATION OF SIBERIAN PLATFORM Paleomagnetic studies of the Ordovician of the Taimyr, Verkhoyano-Kolyma, and Altai-Sayan regions are somewhat limited, and the available data are of low quality. Hence, they cannot be used to constrain the paleogeographic position of these regions. All known paleomagnetic data were obtained from the Ordovician of the Siberian Platform. During the last decade the magnetostratigraphy of the most important Cambrian–Ordovician sections of the Siberian Platform was well established (Gallet and Pavlov, 1996; Pavlov and Gallet, 1998; Rodionov et al., 2001; Pavlov and Gallet, 2005). In general, the new data are in very good agreement with earlier data compiled by Khramov et al. (1987). The data allow definition of a detailed apparent polar wander path (APWP) for Siberia from the Middle Cambrian to the earliest Silurian. Data show that throughout the Cambrian and for most of the Ordovician the Siberian paleocontinent lay in the Southern Hemisphere and was oriented 180° relative to its present position. At the beginning of the Ordovician, the northern part of the modern Siberian Platform was at ~20°S, while its southern margin lay at near-equatorial latitude. At that time the Siberian Platform was slightly (~30° clockwise) rotated relative to its modern orientation. During the Ordovician the Siberian Platform constantly moved northward, and by the end of the period it lay in near-equatorial latitudes of the Northern Hemisphere. Thus, it kept the same position relative to modern longitude as it had at the beginning of the Ordovician. Its southern margin crossed the equator approximately at the end of the Darriwilian. ACKNOWLEDGMENTS The work was assisted by Russian Foundation grants 05-0564509 and 05-05-64832-a and 07-05-01035a, as well as by Ministry of Education and Science grant NSH 1569.2003.5 and RAS Presidium Program 25. The authors are indebted to V. Pavlov (Schmidt Institute of Physics of the Earth RAS, Moscow) for providing information on paleomagnetic studies of the Ordovician of the Siberian Platform. Many thanks are given to M. Oradovskaya for useful discussions. A. Syarkova is thanked for help in technical preparation of the manuscript. This paper is a contribution to ISCP 503—Ordovician Palaeogeography and Palaeoclimates. REFERENCES CITED Abaimova, G.P., 1975, The Early Ordovician Conodonts of the Middle Watercourse of the Lena River: Novosibirsk, Nauka, 140 p. (in Russian). Balashev, Z.G., 1962, Nautiloids of the Ordovician of the Siberian Platform: Leningrad, Izdatel’stvo LGU, 205 p. (in Russian). Cooper, R.A., Nowlan, G.S., and Williams, S.H., 2001, Global Stratotype Section and Point for base of the Ordovician System: Episodes, v. 24, p. 19–28. Eltysheva, R.S., 1960, Ordovician and Silurian Crinoids of the Siberian Platform: Leningrad, Nedra, 40 p. (in Russian).
Erdtmann, B.D., 1982, A reorganization and proposed phylogenetic classification of planktic Tremadoc (early Ordovician) dendroid graptolites: Norsk Geologisk Tidsskrift, v. 62, p. 121–144. Gallet, Y., and Pavlov, V., 1996, Magnetostratigraphy of the Moyero River section (north-western Siberia): Constraint on the geomagnetic reversal frequency during the early Paleozoic: Geophysical Journal International, v. 125, p. 95–105, doi: 10.1111/j.1365-246X.1996.tb06536.x. Gutiérrez-Marco, J.C., 1997, Tolmachovia babini nov. sp. (Rostroconchia), a new ribeirioid mollusk, from the Middle Ordovician of the Central Iberian Zone: Geobios-Mémoire Spécial, no. 20, p. 291–298. Ivanova, V.A., 1959a, Some Ordovician ostracodes of the Siberian Platform: Paleontologicheski Zhurnal, no. 4, p. 130–142 (in Russian). Ivanova, V.A., 1959b, New and previously unknown in USSR ostracode genera from the Ordovician sediments of the Siberian Platform: Material to the fundamental paleontology, vyp. 3, Moscow, p. 71–83 (in Russian). Kanygin, A.V., 1967, Ordovician Ostracodes of the Cherskii Mountains: Moscow, “Nauka,” 123 p. (in Russian). Kanygin, A.V., 1971, Ordovician Ostracodes of the Sette-Daban Range (Verkhoyansk Mountains): Moscow, “Nauka,” 65 p. (in Russian). Kanygin, A.V., 2001, The Ordovician explosive divergence of the Earth’s organic realm: Causes and effects of the biosphere evolution: Russian Geology and Geophysics, v. 42, p. 599–633. Kanygin, A.V., Moskalenko, T.A., and Yadrenkina, A.G., 1980, Lower and Middle Ordovician boundary deposits on the Siberian Platform: Russian Geology and Geophysics, no. 6, p. 13–18. Kanygin, A.V., Moskalenko, T.A., Yadrenkina, A.G., et al. (Sokolov, B.S., ed.), 1982, The Ordovician of the Siberian Platform: Reference Section on the Kulumbe River: Moscow, Nauka, 224 p. (in Russian). Kanygin, A.V., Moskalenko, T.A., Divina, T.A., et al., 1984a, The Ordovician in the Western Part of the Irkutsk Amphitheatre: Moscow, Nauka, 159 p. (in Russian). Kanygin, A.V., Moskalenko, T.A., Obut, O.T., and Sennikov, N.V., 1984b, Ordovician, in Yanshin, A.L., ed., Phanerozoic of Siberia, T.I. Vendian, Paleozoic: Novosibirsk, Nauka, p. 60–88 (in Russian). Kanygin, A.V., Moskalenko, T.A., and Yadrenkina, A.G., 1988, The Siberian Platform, in Rubeun, J., Ross, Jr., and Talent, J., eds., The Ordovician System in Most of Russian Asia, Correlation Charts and Explanatory Notes: International Union of Geological Sciences (IUGS), no. 26, p. 1–27. Kanygin, A.V., Moskalenko, T.A., Yadrenkina, A.G., et al., 1989, The Ordovician of the Siberian Platform, Fauna and Stratigraphy of the Lena Facies Zone: Novosibirsk, Nauka, 216 p. (in Russian). Kanygin, A.V., Yadrenkina, A.G., Timokhin, A.V., et al., 2007, Stratigraphy of Oil and Gas Basins of Siberia, Ordovician of Siberian Platform: Novosibirsk, Publishing House of Siberian Branch of RAS, Department “GEO,” 269 p. (in Russian). Khramov, A.N., Goncharov, G., Komissarova, R., Pisarevsky, S., and Pogarskaya, I., 1987, Paleomagnetology: Berlin, Springer-Verlag, 308 p. Maksimova, Z.A., 1962, Trilobites of the Ordovician and Silurian of the Siberian Platform: Moscow, Gosgeoltekhizdat, 184 p. (in Russian). Moskalenko, T.A., 1970, Conodonts of the Krivolutchkii Regional Stage (Middle Ordovician) of the Siberian Platform: Moscow, Nauka, 118 p. (in Russian). Moskalenko, T.A., 1973, Conodonts of the Middle and Upper Ordovician of the Siberian Platform: Novosibirsk, Nauka, 144 p. (in Russian). Moskalenko, T.A., 1994, Zonal distribution of conodonts in the Middle and Upper Ordovician of the Siberian Platform: Russian Geology and Geophysics, v. 35, p. 312–337. Moskalenko, T.A., Yadrenkina, A.G., Semenova, V.S., and Yaroshinskaya, A.M., 1978, The Ordovician of the Siberian Platform, Upper Ordovician Reference Sections (Biostratigraphy and Fauna): Moscow, Nauka, 164 p. (in Russian). Nekhoroshev, V.P., 1961, Ordovician and Silurian Bryozoans of the Siberian Platform: Moscow, Gostoptekhizdat, 246 p. (in Russian). Nikiforova, O.I., 1955, New data on stratigraphy and paleontology of the Ordovician and Silurian of the Siberian Platform: Materials on the geology and source minerals of the Siberian Platform: Leningrad, Izdatel’stvo VSEGEI, p. 50–106 (in Russian). Nikiforova, O.I., and Andreeva, O.N., 1961, Siberian Platform Ordovician and Silurian Stratigraphy and Its Paleontological Substantiation: Leningrad, Gostoptekhizdat, 412 p. (in Russian). Obut, A.M., and Sobolevskaya, R.F., 1967, Some stereostolonates of the late Cambrian and Ordovician of the Norilsk Region, in Ivanovsky, A.B., and
Ordovician of the Siberian Platform Sokolov, B.S., eds., Materials on Geology and Source Minerals of the Siberian Platform: Moscow, “Nauka,” p. 45–64 (in Russian). Obut, A.M., Sennikov, N.V., and Zaslavskaya, N.M., 1984, Siberian graptolite and chitinozoan assemblages on the Cambrian/Ordovician boundary: Geologia i Geophizica, no. 3, p. 3–7 (in Russian). Oradovskaya, M.M., 1988, Ordovician and Silurian Biostratigraphy and Facies in the Northeast USSR: Moscow, “Nedra,” 160 p. (in Russian). Pavlov, V., and Gallet, Y., 1998, Upper Cambrian to Middle Ordovician magnetostratigraphy from the Kulumbe River section (northwestern Siberia): Physics of the Earth and Planetary Interiors, v. 108, p. 49–59, doi: 10.1016/S0031-9201(98)00087-9. Pavlov, V., and Gallet, Y., 2005, Third superchron during the Early Paleozoic: Episodes, v. 28, p. 78–84. Rodionov, V.P., Pavlov, V.E., and Galle, I., 2001, Magnitopolar structure of the type section of the Kirensko-Kudrinian and Chertovskian regional stages of the Middle Ordovician (upper watercourse of the Lena River, behind Kirensk town): To the problem of the Ordovician geomagnetic superchron: Phisika Zemli, v. 6, p. 67–71. Rozman, Kh.S., 1970, Stratigraphy and brachiopods of the Middle and Upper Ordovician of the Sette-Daban and Seleniakhinsk Mountain Ridge: Biostratigraphy of the upper Ordovician of the North-East of USSR: Moscow, “Nauka,” p. 8–44 (in Russian). Rozova, A.V., 1968, Biostratigraphy and trilobites of the Upper Cambrian and Lower Ordovician of the Northwestern Part of the Siberian Platform: Moscow, Nauka, 196 p. (in Russian). Sennikov, N.V., 1996, Paleozoic Graptolites of East Siberia (Systematics, Phylogeny, Biochronology, Biologic Affinity and Paleozoogeography): Novosibirsk, izdatel’stvo SO RAN, 225 p.
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The Geological Society of America Special Paper 466 2010
Early–Middle Ordovician conodont paleobiogeography with special regard to the geographic origin of the Argentine Precordillera: A multivariate data analysis Guillermo L. Albanesi† CONICET—Museo de Paleontología, Universidad Nacional de Córdoba, C.C. 1598, 5000 Córdoba, Argentina Stig M. Bergström† School of Earth Sciences, The Ohio State University, Columbus, Ohio 43210-1397, USA
ABSTRACT Statistical comparisons of conodont faunas from many parts of the world were carried out in an effort to shed light on one of the most discussed, and most controversial, problems in the lower Paleozoic geology of South America, namely, the geographic origin of the exotic terrane in western Argentina known as the Precordillera. The similarity between the conodont faunas from the Precordilleran La Silla, San Juan, Gualcamayo, and Yerba Loca Formations and many coeval faunas from Laurentia, as well as from other parts of the world, was assessed using the Jaccard Index. The analysis of faunas from six biostratigraphic intervals in the Lower and Middle Ordovician shows that the earliest Ordovician (Tremadocian) faunas cluster with those of Laurentia, whereas slightly younger faunas show less obvious provincialism. The conodont faunas of the Middle Ordovician (early Darriwilian) of the Precordillera again show dominantly Laurentian affinities. The hypothesis that the Precordillera rifted from the Ouachita embayment and moved across part of the Iapetus Ocean to dock with western Gondwana in Ordovician time is not clearly supported by the conodonts (and other non-conodont phosphatic microfossils). The similarity with faunas in southern Laurentia (mainly from the El Paso area of Texas and southern New Mexico) is high in the Tremadocian. The expected similarity decrease with presumed increase in distance from Laurentia later in the Early Ordovician is not evident in the conodont faunas. Similarity between the two regions (mainly the Precordillera and the Marathon area of Texas) remains about the same through the early Middle Ordovician. It is concluded that the conodonts, the best known and most widespread fossil group in the study areas, do not provide conclusive evidence of the geographic origin of the Precordillera.
†
E-mails:
[email protected];
[email protected].
Albanesi, G.L., and Bergström, S.M., 2010, Early–Middle Ordovician conodont paleobiogeography with special regard to the geographic origin of the Argentine Precordillera: A multivariate data analysis, in Finney, S.C., and Berry, W.B.N., eds., The Ordovician Earth System: Geological Society of America Special Paper 466, p. 119–139, doi: 10.1130/2010.2466(08). For permission to copy, contact
[email protected]. ©2010 The Geological Society of America. All rights reserved.
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INTRODUCTION Intercontinental biostratigraphical correlations tend to be problematical owing to the pronounced regional differences in faunal composition that are known as provincialism. Marine fossil groups of the Ordovician System (ca. 488–443 Ma) show particularly marked provincialism. Conodonts are an extinct group of marine chordates that have become useful as biostratigraphical and paleoenvironmental tools (e.g., Clark, 1984; Sweet, 1988; Bergström, 1990; Pohler and Barnes, 1990; Sweet and Donoghue, 2001) as are some other Paleozoic invertebrates such as graptolites, trilobites, and brachiopods (Ross, 1975; Barnes et al., 1996; Fortey and Cocks, 2003; Benedetto, 2004). As the global database for conodonts and some other widely distributed fossil groups such as brachiopods, bivalves, ostracodes, sponges, bryozoans, and palynomorphs constantly increases, comparative faunal studies are becoming fundamental in assisting, or serving as a base for, paleogeographical reconstructions (e.g., Finney and Chen Xu, 1990; Fortey and Mellish, 1992; Erdtmann and Kraft, 1998; Lees et al., 2002). A particularly complex and controversial paleobiogeographic history is the original geographic position of the Precordillera region of western Argentina. This geological province, which is now located at the foothills of the Andes, has been the focus of intensive studies during the last two decades. These have resulted in its recognition as an exotic terrane, which, according to some authors, was derived from Laurentia—an idea that has been extensively discussed in many papers in Cooper et al. (1995), Kraft and Fatka (1999), and Albanesi et al. (2003a). Because Cambrian carbonate deposits in the Precordillera contain olenellid trilobites similar to those of North America (Poulsen, 1958), a probable Precordilleran-Laurentian connection was suggested by Bond et al. (1984). Based on a variety of geological evidence, Ramos et al. (1986) considered, for the first time, the Precordillera to be a suspect terrane derived from Laurentia. The concept of the Precordillera as a far-traveled lithospheric fragment rifted off the Ouachita embayment of southern Laurentia was further developed by several later authors, and various paleogeographic models based on this concept have been proposed (e.g., Pankhurst and Rapela, 1998; Ramos and Keppie, 1999; Ramos, 2004). Other authors have proposed a model in which the Precordillera was a paraautochthonous terrane that drifted to its present position along the Gondwanan margin from a starting point close to Antarctica and South Africa (Aceñolaza et al., 2002; Finney et al., 2003, 2005), basically following the hypothetical SAMFRAU geosyncline of Keidel (1916) and Du Toit (1927). Fossils are among the best tools for evaluating timing of rifting, migration, and accretion of tectonic plates, and compared with other evidence, they often permit a far more critical assessment of the dynamic plate tectonic history of the Iapetus Ocean during the Cambrian, Ordovician, and Silurian (McKerrow and Cocks, 1986; Williams et al., 1995). Data from brachiopods, trilobites, sponges, and bryozoans suggest a significant geographical separation of the Precordillera from Laurentia across the Iape-
tus Ocean (Benedetto, 1998, 2003, 2004; Benedetto et al., 1999) since the late Early Ordovician (mid-Arenig). A maximum spatial separation distance apparently ended in the Late Ordovician, when the Precordillera became an integral part of the Gondwanan margin. The latter is suggested by a variety of evidence, including the Precordilleran occurrence of Sacabambaspis, the earliest armored fish in South America, which is a typical Gondwanan species that is unknown in North America (Albanesi et al., 1995a, 2005b), as well as invertebrate faunas of Gondwanan affinity (Benedetto, 2003, 2004). Recent paleobiogeographic data from conodonts (Lehnert et al., 1997) and shelly faunas (e.g., Bordonaro and Banchig, 1995; Vaccari, 1995; Holmer et al., 1999; Astini et al., 2004) indicate a close similarity between Precordilleran and Laurentian faunas during the Cambrian and earliest Ordovician, but a more complex similarity pattern developed in the late Early Ordovician (Albanesi, 1998a). A comprehensive analysis of the relations between Early and Middle Ordovician conodont faunas from the Precordillera and those from key localities in previously defined provinces has the potential to furnish critical data for clarifying the paleobiogeographical origin and development of the Precordillera through time and space. A detailed comparison of the extensively studied Precordilleran conodont faunas with those from the southern Laurentian margin has the potential to provide information in support of, or against, the hypothesis of a Laurentian origin of the Precordillera. The Marathon region of southwestern Texas (Berry, 1960; Bradshaw, 1969; Bergström, 1978; Izold, 1993) has been recognized as a possible geological counterpart of the Precordillera (Dickerson and Keller, 1998), and this region is therefore particularly critical for the evaluation of the evolving paleobiogeographical relations between the Precordillera and Laurentia. The purpose of this study is to carry out such a precise comparison, at the species level, between conodont faunas from the La Silla, San Juan, Gualcamayo, and Yerba Loca Formations from the Argentine Precordillera and the coeval Marathon Limestone, Fort Peña Formation, and associated strata in West Texas and southern New Mexico in order to determine whether conodonts can clarify the changing paleogeographic relations between these regions during Early and Middle Ordovician time. One would expect that if the Precordillera came from southern Laurentia, the similarity between the conodont faunas from these regions would show a marked decrease with increased geographic separation of these regions during the move of the Precordillera from Laurentia to Gondwana. Obviously, factors other than geographic separation probably would have affected the similarities and differences between these faunas, and some of these are discussed below. Our detailed assessment of the conodont-faunal relationships, at the species level, between the Argentine Precordillera and the Marathon area of West Texas and coeval strata in New Mexico is achieved by comparative analyses of conodont collections throughout the critical Lower and Middle Ordovician interval (Fig. 1) using cluster analysis. For comparative purposes, we have also included data from other well-documented localities in
dianae
Tripodus
swe.
borealis
deltifer
deltifer
pristinus
angulatus
manitouensis
North America, the Baltoscandian region, and from more sparse collections from China and Australia. The analysis of all these collections provides an outline of the global provincial differentiation of the conodont faunas in the study interval. The extensive database used (Tables 1 and 2) is similar to that assembled by Albanesi and Bergström (2004). GEOLOGICAL AND FAUNAL BACKGROUND Geological Aspects In its original sense, the term Precordillera included the region from Mendoza northward to the Guandacol area. A much larger region called Cuyania (ca. 28°–35°S lat, 67°–69°W long),
TIME SLICES
MILLION YEARS
4b 3b
4a 3a 2c
SINUOSASUECICUS
amoenus
manitouensis fluctivagus
com.
elong./delt.
elong./delt. gracilis
low diversity interv. deltifer
angulatus
elegans
var.
1d
elegans
olds.
Figure 1. Stratigraphic chart of the Ordovician System, showing the conodont biozonal schemes of the North American Midcontinent and North Atlantic Realms, the Argentine Precordillera biozones, and the composite biostratigraphical intervals subjected to multivariate data analysis. After Albanesi and Bergström (2004), Webby et al. (2004), Albanesi et al. (1998, 2006b), and Wang et al. (2005, 2009).
1a
deltatus/ costatus
evae
2b
evae
andinus
communis
laevis intermedius
2a
navis
navis flabellum/laevis triangularis
1c
parva
LAEVISNORRLANDICUS
gla.
ANDINUSEVAE
altifrons
variabilis
antivariabilis norrlan./parva originalis
COMMUNISELEGANS
sinuosa
pseudoplanus crassus variabilis hor.
MANITOUENSISDELTIFER
holodentata
ani. kri.
4c
lind. rob. rec. fol.
1b
lind.
rob. serra rec. friendsvillensis serra fol. polonicus suecicus suecicus
DELTATUSPROTEUS
sweeti
BIOSTRAT. INTERVALS
ARGENTINE
N. ATLANTIC PRECORDILLERA
proteus
YUSHANIAN ICHANGIAN
LANCEFIELDIAN
N. AMERICAN MIDCONTINENT
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ANALYZED INTERVALS
CONODONT ZONES
proteus
CHINA
AUSTR.
GLOBAL
DAPINGIAN DARRIWILIAN DARRIWILIAN DAWANIAN ZHEJIANGIAN FLOIAN TREMADOCIAN
BENDIG.CHEW. CASTL. YA.
N AMER.
WHITEROCKIAN IBEXIAN
TREMADOC
489
LOWER
ORDOVICIAN
ARENIG
471
LLANVIRN BRITAIN
GLOBAL
SERIES STAGES
MIDDLE
SYSTEM
460 MILLION YEARS
Early–Middle Ordovician conodont paleobiogeography
which includes not only the Precordillera but also Grenvillian rocks of the western Sierras Pampeanas, and a series of isolated lower Paleozoic outcrops in the Mendoza and La Pampa Provinces, has been considered to be the drifted terrane (Ramos, 1995). Because our analysis is mostly based on collections from the Precordillera in the old, restricted sense, we use this term in our discussions even though we realize that the Cuyania rather than the Precordillera was the exotic terrane unit. An outstanding feature of the Precordillera is its important succession of carbonate rocks, which is more than 2.2 km thick and which was deposited in shallow-water platform environments during Cambrian to Middle Ordovician time. The development of this succession began in the earliest Phanerozoic, with a first order sea-level rise. This event is recorded also in peripheral parts
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TABLE 1. ANALYZED LOWER ORDOVICIAN CONODONT RECORDS FROM SELECTED LOCALITIES WORLDWIDE Map Lithostratigraphic unit and/or Reference Biostratigraphic Biozone and/or unit (N°) section interval Precordillera (Cuyania Terrane), western Argentina 1 San Jorge Formation Albanesi et al. (2003b) M-D Rossodus manitouensis R. manitouensis, Paltodus deltifer 1 La Silla Formation Lehnert (1995); Albanesi et al. (1998) M-D Paroistodus proteus D-P 1 San Juan Formation, Ponón Serpagli (1974); Lehnert (1995); Albanesi et al. D-P P. proteus Trehué Formation (1998, 2006b); Lehnert et al. (1998) C-E Prioniodus elegans A-E Oepikodus evae, O. intermedius Northwestern Argentina 15 Volcancito Formation Albanesi et al. (2005b) M-D Paltodus deltifer 15 Suri Formation, Famatina Albanesi and Astini (2000a) A-E O. evae, O. intermedius 14 Santa Victoria Group, Eastern Rao et al. (1994); Rao and Flores (1998); R. manitouensis, P. deltifer M-D Cordillera Albanesi and Aceñolaza (2005); Zeballo et al. Acodus deltatus, P. proteus D-P (2005) O. evae A-E Baltoscandian region 13 Lava River section, West Russia Tolmacheva (2001) P. deltifer M-D P. proteus D-P P. elegans C-E O. evae A-E 11 Talubäcken section, Sweden Bergström (1988) P. proteus D-P P. elegans C-E O. evae A-E 12 Furuhäll section, Central Öland, Bagnoli et al. (1988) P. deltifer M-D Sweden P. proteus D-P P. elegans C-E O. evae A-E 12 Horns Udde sections, North Bagnoli and Stouge (1997) A-E O. evae, Trapezognathus diprion, Öland, Sweden Microzarkodina n. sp. A 11 Brattefors, Västergötland Löfgren (1997) M-D P. deltifer 11 Hunneberg, Västergötland, Löfgren (1993a) M-D P. deltifer Sweden D-P P. proteus C-E P. elegans 10 Siljan district, Central Sweden Löfgren (1993b, 1994) M-D P. deltifer D-P P. proteus C-E P. elegans A-E O. evae North America 2 Marathon Limestone, West Texas Izold (1993) M-D R. manitouensis D-P A. deltatus C-E P. elegans A-E O. evae 2 El Paso Formation, West Texas, Repetski (1982); Izold (1993) R. manitouensis, Ulrichodina quadraplicatus M-D New Mexico A. deltatus D-P Fahraeusodus marathonensis, Oepikodus C-E communis A-E Jumudontus gananda–Reutterodus andinus 3 Ouachita Mountains Repetski and Ethington (1977); Ethington et al. M-D R. manitouensis (1987); Repetski and Ethington, in Stone et al. D-P P. proteus (1994) A-E O. evae 8 Green Point Formation, Cow Head Johnston and Barnes (1999, 2000) D-P Paracordylodus gracilis Group, Saint Pauls Inlet section, C-E P. elegans Newfoundland A-E O. evae 8 The Ledge Point of Head section, Stouge and Bagnoli (1988) Prioniodus gilberti M-D Cow Head Group, Newfoundland Prioniodus adami, P. oepiki, D-P P. elegans C-E O. evae A-E 8 Watts Bight, Boat Harbour, Ji and Barnes (1994) M-D Cordylodus angulatus Catoche formations, Saint George D-P Drepanoistodus nowlani, Macerodus dianae Group, Newfoundland C-E A. deltatus, Acodus? primus A-E O. communis, P. simplicissimus Ross et al. (1997) 5 The Ibexian Series composite M-D R. manitouensis Macerodus dianae, A. deltatus – stratotype section and adjacent D-P strata, House-Confusion Range C-E Stultodontus costatus, Area, West-Central Utah O. communis A-E Reutterodus andinus 6 Ninemile Formation, Whiterock Finney and Ethington (2000) A-E R. andinus Narrows section, Monitor Range, Nevada 7 Little Falls and Tribes Hill Landing et al. (1996) M-D R. manitouensis Formations, East New York 9 Cape Weber Formation, Ella Smith (1991) R. manitouensis M-D Section, East Greenland Fauna D (middle-upper part) D-P C-E O. communis 4 Manitou Formation, Colorado Seo and Ethington (1993) M-D R. manitouensis Australia 16 Allochthonous limestones, Webby et al. (2000) C-E P. elegans Hensleigh Siltstone, NSW
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TABLE 2. ANALYZED MIDDLE ORDOVICIAN CONODONT RECORDS FROM SELECTED LOCALITIES WORLDWIDE Map Lithostratigraphic unit and/or section Reference Biostratigraphic Biozone and/or unit (N°) interval Precordillera (Cuyania Terrane), Western Argentina 1 San Juan, Gualcamayo, Los Azules, Lehnert (1995); Albanesi et al. (1995b, L-N Tripodus laevis, Baltoniodus Yerba Loca, Sierra de La Invernada, 1998, 1999); Ottone et al. (1999); S-S navis, Microzarkodina parva Las Aguaditas, and Las Chacritas Albanesi and Astini (2000b); Heredia Lenodus variabilis, Formations et al. (2005) Eoplacognathus suecicus Western Argentina 2 Suri Formation, Famatina Albanesi and Vaccari (1994) L-N B. navis Baltoscandian region 10 Mäekalda, North Estonia Viira et al. (2001) L-N B. navis, B. norrlandicus 11 Lava River, West Russia Tolmacheva (2001) L-N B. triangularis, B. norrlandicus 8 Sjurberg, Siljan district, Central Löfgren (1994) L-N B. triangularis, B. navis Sweden 8 Jämtland, Central Sweden Löfgren (1978, 1993b) L-N B. triangularis, B. navis S-S L. variabilis, E. suecicus 9 Gillberga, North Öland, Sweden Löfgren (2000) L-N B. navis, B. norrlandicus S-S L. variabilis, Yangtzeplacognathus crassus, Eoplacognathus pseudoplanus 9 Hagudden, North Öland, Sweden Stouge and Bagnoli (1990) L-N B. navis, M. parva S-S L. variabilis 9 Horns Udde sections, North Öland, Bagnoli and Stouge (1997) L-N B. triangularis, B. norrlandicus Sweden 7 Huk Formation, Slemmestad, South Rasmussen (1991) L-N B. navis, M. parva Norway S-S L. variabilis 7 Scandinavian Caledonides, NorwayRasmussen (2001) L-N Microzarkodina flabellum, Sweden B. norrlandicus 12 Mójcza Limestone, Poland Dzik (1994) S-S L. variabilis North America 6 Saint George Group, Western Ji and Barnes (1994) L-N Pteracontiodus cryptodens Newfoundland 6 Saint Pauls Inlet section, Western Johnston and Barnes (1999) L-N T. laevis Newfoundland 6 Table Head Formation, Newfoundland Stouge (1984) S-S L. variabilis, E. suecicus 3 Fort Peña Formation, Marathon area, Bradshaw (1969) S-S L. variabilis, E. suecicus Texas 5 Whiterock Narrows section, Monitor Finney and Ethington (2000) L-N T. laevis Range, Roberts Mountains, Nevada S-S Histiodella sinuosa 5 Antelope Valley Limestone, Eagan Harris et al. (1979) S-S L. variabilis, E. suecicus Range, Martin Ridge–Monitor Range, Copper Mountain–March Spring–Ikes Canyon, Groom Range, Test Site, Meiklejohn Peak sections, Nevada 4 Ibex area, Utah Ethington and Clark (1981); Ross L-N T. laevis et al. (1997) S-S H. sinuosa, H. holodentata Australia 16 Amadeus Basin, Central Australia Cooper (1981) L-N B. triangularis, B. navis 15 Canning Basin, Western Australia Watson (1988) S-S L. variabilis, E. suecicus China 14 East-Central China Wang Zhi-hao and Bergström (1999) L-N Paroistodus originalis, B. norrlandicus 13 South-Central China Zhang Jianhua (1998) S-S L. variabilis, E. suecicus
of Laurentia. The lithofacies heterogeneity of the oldest deposits (Baldis and Bordonaro, 1982) changed progressively to rhythmic deposition in the Late Cambrian, when recurrent peritidal cycles were established, as is the case also in the central and southern Appalachians (Astini, 1998; Glumac and Walker, 2000; Cañas, 1999). In the Early Ordovician, extensive muddy carbonate platforms, with open marine faunas and associated biohermal reef structures, were developed in the Precordillera in a similar architecture as in marginal settings of Laurentia (Beresi and Rigby, 1993; Carrera and Cañas, 1996; Alberstadt and Repetski, 1989). In the former region, the carbonate succession is overlain by diachronous, dominantly siliciclastic units, which reflect basinal
instability caused by regional tectonics combined with eustatic sea-level changes. These began near the beginning of the Middle Ordovician (e.g., Albanesi et al., 1998, 1999; Astini, 1998; Astini and Thomas, 1999). In the Precordillera, this Middle Ordovician time interval is also characterized by a significant influx of volcanic ashes (K-bentonites). Although the geographic location of the source volcano(es) of the many ash beds remains elusive, it is most unlikely to have been along the Laurentian margin because no correlative ash bed complexes are known from southern Laurentia (Huff et al., 1998). Comparison of successive stages of the development of miogeoclinal sequences in southern Laurentia shows a broad
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agreement with coeval ones of the Precordillera. Their history of subsidence, paleoenvironments, and depositional sequences seem to be consistent with a parallel evolution of both passive margins (Benedetto, 1993; Astini et al., 1995; Cañas, 1999). Two current alternative hypotheses for PrecordilleranLaurentian relationships are (1) that the significant facies changes from carbonates to siliciclastics were due to the initiation of a rifting episode that led to the separation of Laurentia and Gondwana (Dalla Salda et al., 1992; Dalziel et al., 1996; Keller, 1999; Cañas, 1999), and (2) that such a stratigraphical framework resulted from the rearrangement and extensional relaxation after the accretion of the Precordillera with the Gondwanan margin (Astini, 1998; Astini and Thomas, 1999; Thomas et al., 2002, 2004). An alternative conceptual model for the original and changing location of the Precordillera suggests that it was a paraautochthonous Gondwanan terrane, which was displaced from a hypothetical intermediate sector between South America, Africa, and Antarctica. This idea was first proposed by Aceñolaza and Toselli (1988) and further elaborated by Baldis et al. (1989), Aceñolaza and Toselli (2000), and Aceñolaza et al. (2002). A similar model implying that the Precordillera was not a part of Laurentia but a separate microcontinent originally located at tropical latitudes off the coast of Gondwana was recently proposed by Finney et al. (2003, 2005), based primarily on geochemical evidence from basement rocks. Finney (2007) also critically reviewed the rather limited megafossil evidence used by previous authors for assessments of the paleogeographic position of the Precordillera, and he concluded that the evidence was more consistent with a Gondwanan origin. Faunal Aspects Current interpretations of Precordilleran-Laurentian faunal relationships are mainly based on benthic organisms such as trilobites, brachiopods, bivalves, sponges, bryozoans, and ostracodes (Benedetto, 1993, 1998; Benedetto et al., 1999; Waisfeld et al., 1999), although interpretations of paleogeographic relations based on nektonic-planktonic groups, such as graptolites and palynomorphs, have also been published (Maletz and Ortega, 1995; Ottone et al., 1999; Peralta and Finney, 2002). Based on faunal evidence, Benedetto et al. (1999) distinguished four successive gradual stages in the paleobiogeographic evolution of the Precordillera that were consistent, in general terms, with geological evidence: (1) Laurentian stage (Cambrian–Tremadocian), (2) isolation stage (Floian–Middle Ordovician), (3) pre-accretion stage (Middle–Late Ordovician), and (4) Gondwanan stage (Hirnantian–Silurian). According to Benedetto et al. (1999), trilobite data suggest a gradual separation between the Precordillera and Laurentia beginning in the Middle Cambrian. Through the Floian–Middle Ordovician interval, Laurentian faunal influence diminished and Baltic-Avalonian genera increased in number. The isolation of the Precordillera within the Iapetus Ocean is supported by the appearance of endemic forms through allopatric speciation processes (Carrera and Rigby, 1999; Albanesi and
Barnes, 2000). Typical Gondwanan faunas colonized the Precordillera during the presumed accretion in late Middle to latest Ordovician time (e.g., Albanesi et al., 1995a, 2005b; Benedetto, 2003, 2004). Paleogeographical interpretations of the Precordillera on the basis of its conodont faunas were first published by Serpagli (1974) and were subsequently dealt with by, e.g., Hünicken, (1989), Pohler and Barnes (1990), and Bagnoli and Stouge (1991). However, these contributions discussed only the provincial affinity of the conodont faunas in relation to major faunal regions or realms. Only recently has a specific conodont connection between the Cuyania terrane and the southern margin of Laurentia been considered. Hence, Lehnert et al. (1997) concluded that a particular conodont species assemblage from the La Silla Formation (Late Cambrian–early Tremadocian), which is characterized by the presence of the warm, shallow-water conodont Clavohamulus hintzei Miller, was closely related to those from southern Laurentia. Owing to its restricted geographic distribution, this species assemblage was interpreted to indicate geographic proximity between these regions. In a more comprehensive study, involving numerical analysis of conodont biofacies, Albanesi (1998a) recognized the close similarity of these faunas but noted the progressive turnovers of the nekto-benthic conodont communities through time, which are related to environmental changes in the La Silla, San Juan, Gualcamayo, and Las Plantas Formations. He recognized a gradual immigration of conodonts from colder regions and interpreted the Precordillera as a separate temperate province of the Atlantic Realm in Middle Ordovician time, but he also recognized a continuous exchange of genera with Laurentian regions (cf. Keller and Lehnert, 1998; Lehnert et al., 1999b). The known Silurian conodont faunas from the Precordillera include only cosmopolitan taxa, precluding paleogeographic assessment (Lehnert et al., 1999a; Albanesi et al., 2006a). The scope of the present report does not allow the listing of the entire conodont database now available for comparative studies on the Argentine Precordillera and the southern Laurentia, although it is appropriate to mention some references. Since the first descriptions of conodonts from the Precordillera by Hünicken (1971), numerous papers deal with this group, most of which are focused on Ordovician conodont biostratigraphy (see review by Albanesi et al., 1995b). A few monographic contributions summarize the most significant information about conodont taxonomy and biostratigraphy in the Ordovician of the Precordillera (Serpagli, 1974; Lehnert, 1995; Albanesi et al., 1998; Albanesi, 1998b). In North America several studies provide information about the Ordovician conodont successions in the southern Appalachians and the Ouachita embayment. For example, Bergström (1973) and Bergström and Carnes (1976) reported on faunas from Appalachian sequences of East Tennessee, Sweet and Bergström (1962) studied conodonts from the Pratt Ferry Formation (Middle Ordovician) of Alabama, and Repetski (1992) reported conodonts from the Knox and Stones River Groups in Georgia. There are fewer, but not less important, reports on Lower-Middle
Early–Middle Ordovician conodont paleobiogeography Ordovician conodonts from the Ouachita Mountains (Repetski and Ethington, 1977; Ethington et al., 1987; Repetski et al., in Stone et al., 1994). Graves and Ellison (1941) described the first Ordovician conodonts from the Marathon Basin in Texas, and more comprehensive studies were later carried out by Bradshaw (1969) and Bergström (1978). Significant records of conodont faunas from the Ouachita orogen in Arkansas, Oklahoma, and West Texas were summarized by Ethington and Repetski (1984), Ethington et al. (1989), and Derby et al. (1991). The monographic work by Repetski (1982) on Lower Ordovician conodonts from the El Paso Group of westernmost Texas and southern New Mexico, and the unpublished master’s thesis of Izold (1993), which is based on unpublished collections from the Marathon area, are particularly important for the present study. MATERIAL AND METHODS Complex rearrangement of the lithosphere, such as the formation of geographic barriers by plate divergence, transform fault displacement of terranes and collision, incidental volcanism, and secular recirculation of ocean currents (also related to atmospheric changes), may produce two types of biogeographic patterns recognizable in the fossil record: disjunct and temporal. In the former, patterns of endemism and diversity that are already part of the fossil record may be fragmented, dispersed, or brought from distant places and juxtaposed. In the latter, existing patterns of endemism and diversity can be affected by the establishment and destruction of geographic barriers by plate tectonic processes leading to speciation, radiation, and extinction events (Smith, 1988). Correctly piecing together the mixed information of the geological record may address two complementary objectives, namely, to reconstruct a lost paleogeography and to clarify the biological, environmental, and historical controls on past living environments. Vicariance biogeography and panbiogeographical concepts assume that, if the history of life has paralleled the history of the Earth, then congruent biological and geological patterns of relationships should result (Newton, 1992; Craw et al., 1999). With recent updates, the database used here (Tables 1 and 2) is that used by Albanesi and Bergström (2004), who compiled data on the presence-absence of Ordovician conodont species from numerous published and unpublished sources. Although extensive, the coverage of geographic sites is not uniform. It includes diverse localities in Laurentia, Baltica, the western Gonwanan margin, and the Precordillera, as well as localities in Australia and China (Figs. 2 and 3). We have restricted our analysis to the Lower (but not basal) and lower Middle Ordovician (Fig. 1), as represented in the same six biostratigraphic intervals used by Albanesi and Bergström (2004). We have made an attempt to provide adequate coverage of both the North American Midcontinent and the North Atlantic provinces or realms (Bergström, 1973, 1990; Barnes et al., 1973; Sweet and Bergström, 1974; Barnes and Fåhraeus, 1975), even if particular localities with detailed studies were not considered for
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cluster analysis because their locations are beyond direct influence of Precordilleran faunal dynamics (e.g., An, 1987; Pyle and Barnes, 2002). Herein, we refer to these major biogeographic units as Atlantic and Midcontinent Realms, following Bergström (1990). A conodont province, according to the definition of Pohler and Barnes (1990, p. 411), represents “an assemblage of conodont communities existing within a geographical area.” This is an ecological definition that differs from provincialism as a biogeographic concept. Following the recent reconsideration of Ordovician conodont paleobiogeography by Zhen and Percival (2003, p. 360), “…provincialism can only be accurately assessed through analysis of endemism…” (cf. Longhurst, 1998; for modern ecological geography of the sea). Accordingly, a biogeographical province is a region with coincident occurrence of relatively large numbers of endemic elements. A region possessing 25%– 50% endemicity is classified as a province (Kauffman, 1973). Zhen and Percival (2003) reconsidered the hierarchy of biogeographic units, recognizing the Shallow-Sea Realm (depth of <200 m), which includes the tropical (similar to the Midcontinent Realm), temperate, and cold domains with different provinces, and the Open-Sea Realm with deep-cold waters in the same three domains (equivalents to the Atlantic Realm). The evolutionary faunal trends through time and space can be envisaged by means of an integrated approach, using ecological and historical methods. Consequently, in our study we consider both cosmopolitan and endemic conodont species. Cosmopolitan taxa are important for biostratigraphic correlation, and the restricted distribution of endemic taxa permits reconstruction of biogeographic provincial patterns (Rasmussen, 1998). Previous cluster analysis investigations have suggested that more than five species per sample site and more than five areas are necessary for the resulting dendrograph to adequately express changes in the similarity index and hierarchical patterns of provinciality (Flessa and Hardy, 1988). We have used these premises in our statistical analysis and have excluded any queried, aff., or cf. species designations in the lists for every sample site. Incidentally, 157 conodont species are listed for the Argentine Precordillera, and their ranges are distributed through the six biostratigraphic intervals. The locations of the study sites are shown in Figures 2 and 3. The index of faunal similarity used is the Jaccard Index: C/ (N1 + N2 – C), where C is the number of species in common to a pair of areas, N1 is the number of species in the first region, and N2 is the number of species in the second region. Jaccard is a sensitive index for similarities, and it provides a more conservative resemblance than other applied indices. Six matrices of faunal similarity were subjected to cluster analysis using the Statistica software (Statsoft: http://www.statsoftinc.com/). This program adopts the unweighted pair-group method of clustering, whose resulting dendrographs have the property of arranging the localities in such a way that the most distant ones along the vertical axis belong to those regions most unlike each other. A useful introduction to general applications of cluster analysis is given by
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480 Ma
16
2 1 3 15 14
6 5 4
Figure 2. Paleogeographic map of the Lower Ordovician, with study sites as indicated in Table 1. After Scotese (2001).
9 7
10 11 13 12
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Figure 3. Paleogeographic map of the Middle Ordovician, with study sites as indicated in Table 2. After Scotese (2001).
Middle Ordovician Davis (1986), and several authors provide examples of its application to conodont biogeography; see, e.g., Bitter (1972), Sweet and Bergström (1984), Flessa and Hardy (1988), and Rasmussen (1998). Limitations of clustering analyses are well known, as are their sensitivity to different similarity indices and to different clustering algorithms, as well as the tendency to reduce dimensionality (Hazel, 1970; Hubalek, 1982). Being acutely aware of the danger of over-interpreting dendrographs, we have chosen a more conservative approach by eliminating questionable species and low diversity localities, and have chosen a measure of faunal similarity that tends to underestimate affinity, so that if affinity is found, it has, arguably, greater significance. Nevertheless, the dendrographs express in a quantitative and efficient mode the relationships between areas, although the final interpretation of the relationships is the responsibility of the authors. The paleogeographic maps of the Scotese (2001) paleomap project were adopted for the geographic location of the different study sites. Although the paleogeographic location of some
areas is not fully resolved, these updated maps provide a widely accepted reconstruction of the position of larger Ordovician plates where most of the study areas are located, i.e., Laurentia and Baltica. As discussed above, the position of the Precordilleran exotic terrane prior to its present situation in South America is controversial (cf. Astini et al., 1995; Aceñolaza et al., 2002), although paleobiogeographical data of megafossils support the hypothesis of the Precordillera as a drifted terrane (Benedetto, 2004). Diverse names have been used for faunal regions or paleobiogeographic units on the basis of Ordovician conodonts (e.g., Barnes and Fåhraeus, 1975; Lindström, 1975; Dzik, 1983; Charpentier, 1984; Bergström, 1990; Pohler and Barnes, 1990; Zhen and Percival, 2003). In the present study we employ an informal, geographically based nomenclature, because we are not interpreting the precise biogeographic category of the groups recognized in the clusters, although the two major previously recognized conodont realms (the Midcontinent and Atlantic, or Shallow-Sea and Open-Sea Realms) are identifiable in most
Early–Middle Ordovician conodont paleobiogeography intervals. Present methodology is congruent with the panbiogeographic method that emphasizes the analysis of raw locality and broader distribution data for taxa, and stresses the importance of the geographic context for the understanding of the history of life (Craw et al., 1999). CONODONT FAUNAL PATTERNS THROUGH TIME AND SPACE In our cluster analyses (Figs. 4–9) we have considered localities with lithostratigraphic units that represent diverse environments and that contain endemic species. The platform and slope settings of Laurentia are represented separately in various plotted localities, even from the southern margin of Laurentia. The conodont data from the Marathon and El Paso Formations were combined for the purpose of homogenizing the faunas from shallow- and open-sea environments in a critical region of southern Laurentia. These faunas are compared with faunas from different settings of the Argentine Precordillera, namely those represented by the La Silla (shallower water, lagoon), San Juan (open platform), Gualcamayo and lateral equivalents (slope), and Yerba Loca (deeper water–slope–ocean basin) Formations. This gives us the opportunity to make a biogeographically controlled double comparison at the same time. First, we can compare the Precordilleran faunas with coeval faunas from different regions and similar or diverse facies. Second, we can compare particular faunas of the Precordillera with mixed Marathon–El Paso faunas for each chrono-interval. This type of comparison makes it possible to recognize the appearance of a Precordilleran endemic bias through time. This faunal bias should differentiate the Precordillera terrain from others and locate it within one of both realms (shallow- and open-sea) and closer to particular localities within its group, as shown in different clusters. Furthermore, the similarity analysis (Fig. 10) allows us to compare the affinity of different Laurentian and Baltoscandian conodont communities with coeval communities of the Precordillera. For this comparison the El Paso platform faunas are separated from the Marathon slope faunas in order to evaluate Precordilleran faunas in equivalent environments. Early to Middle Ordovician Conodont Provinciality Although conodont provincialism can be distinguished already in the Late Cambrian (Jeong and Lee, 2000), the differentiation into the Midcontinent and Atlantic Realms became more marked in the late Tremadocian (Miller, 1984). In view of this, our analysis commences with Tremadocian faunas that represent a significant tool for paleogeographic assessments in the Iapetus Ocean region. Our time intervals follow those used by Albanesi and Bergström (2004) in the analysis of conodont diversity. The manitouensis-deltifer Interval Dendrograph, Figure 4. The manitouensis-deltifer interval dendrograph documents two well defined groups of sites.
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Although the clustering level that separates the two main groups is rather low, it significantly increases for the subclusterings of particular localities. Almost all North American sites cluster together with the Argentine Precordillera, whereas the western Gondwanan sites of the Eastern Cordillera and the Famatina System, both Argentine basins, cluster together with the Baltic sites. These well defined groups appear to demonstrate that at this time the two widely recognized Midcontinent and Atlantic Realms were already established (cf. Ethington and Repetski, 1984). The only North American locality linked to the Atlantic Realm is that of Cow Head, Newfoundland, although the low level of clustering to this group could be related to sampling bias or biofacies differentiation rather than reflecting a vicariance. The disjunct pattern of conodont provinciality shown by the Argentine Precordillera and the currently adjacent basins of the Famatina System and the Eastern Cordillera can be explained by the hypothesis that the Precordillera was a terrane exotic to the Gondwanan margin, and that its faunas were more directly related to those of Laurentia, as the dendrograph links suggest in Figure 4. The deltatus-proteus Interval Dendrograph, Figure 5. In this interval the provinciality is not as evident as in the previous interval, which was more directly linked to paleocontinents and their epicontinental seas. In the deltatus-proteus interval the differentiation into two groups is apparent, but these groups include “mixed” localities that belong to either Laurentia or Baltica. One of these clusters groups Swedish localities together with the Marathon–El Paso areas, Texas, the Ouachita Mountains, and the Precordillera. This suggests that during the Ceratopyge regressive eustatic event (Erdtmann, 1986) that took place during this interval, an oceanic corridor connecting low and high latitudinal belts was efficient enough to permit a faunal exchange between Laurentia and Baltica, similar to that demonstrated by Löfgren et al. (1999). The steplike appearance of the other cluster, which includes localities as distant from each other as the Eastern Cordillera, Baltic Russia, Cow Head in Newfoundland, the Ibex area in Utah, and East Greenland more likely represents a variety of similar biofacies than a paleobiogeographic unit (Ethington et al., 1987). The Saint George, Newfoundland, locality is disconnected from both clusters because of sampling bias or biofacies difference rather than vicariance. The apparent weak provincialism during this time interval is probably related to the complex Ceratopyge regressive event, with its various phases (Nielsen, 2004), that promoted more uniform habitats during a sea-level lowstand, as well as possibly more uniform climatic conditions across the various paleolatitudinal belts. The communis-elegans Interval Dendrograph, Figure 6. This interval includes the beginning of the second sea-level highstand in the Ordovician. Its dendrograph differs markedly from the random clustering displayed in the previous dendrograph in that more resolved provinciality groups are displayed. The drowning of continental platforms, the establishment of widespread epeiric seas, and the presence
Matrix: Jaccard coefficient Unweighted pair-group average Euclidean distances Precordillera, Argentina Ouachita Mountains, North America Ibex Area, Utah, North America Manitou, Colorado, North America East New York, North America Saint George, Newfoundland Marathon-El Paso, North America East Greenland Famatina, Argentina Central Sweden Baltic Russia Öland, Sweden South Sweden Eastern Cordillera, Argentina Cow Head, Newfoundland
Figure 4. Dendrograph from cluster analysis of the manitouensis-deltifer interval.
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Figure 5. Dendrograph from cluster analysis of the deltatus-proteus interval.
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Figure 6. Dendrograph from cluster analysis of the communis-elegans interval.
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Matrix: Jaccard coefficient Unweighted pair-group average Euclidean distances Precordillera, Argentina Ninemile, Nevada, North America Marathon-El Paso, Texas, North America Cow Head, Newfoundland Öland, Sweden Central Sweden Ouachita Mountains, North America Baltic Russia Saint George, Newfoundland East Greenland Ibex Area, Utah, North America Famatina, Argentina Eastern Cordillera, Argentina
Figure 7. Dendrograph from cluster analysis of the andinus-evae interval.
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Matrix: Jaccard coefficient Unweighted pair-group average Euclidean distances Precordillera, Argentina West Newfoundland Nevada, North America Central Australia Ibex Area, Utah, North America South Central Sweden South Central China South Norway Öland, Sweden Baltic Russia
Figure 8. Dendrograph from cluster analysis of the laevis-norrlandicus interval.
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Matrix: Jaccard coefficient Unweighted pair-group average Euclidean distances Precordillera, Argentina Marathon Area, Texas, North America Table Head, Newfoundland Antelope Valley, Nevada, North America
Figure 9. Dendrograph from cluster analysis of the sinuosa-suecicus interval.
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of connecting ocean currents were favorable for biogeographical differentiation (Christiansen and Stouge, 1999). The dendrograph displays two major groups that probably represent the Atlantic and Midcontinent Realms. The first coherent group includes the Precordillera, the Marathon–El Paso area, and Cow Head, Newfoundland. A fourth site, in New South Wales, Australia, is linked to this group by a higher Euclidean distance, making the rela-
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Figure 10. Comparison of conodont faunal composition of the Argentine Precordillera with selected sites. Sea-level changes after Nielsen (2004). Stratigraphic references as in Figure 1.
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tionship less resolved than that for the other three localities. A most coherent group includes the four Baltic localities. The third group, linking Saint George, Newfoundland, with the Ibex area, Utah, and East Greenland, shows a relatively weak similarity. It connects areas in the same circumequatorial belt and may reflect a particular biofacies, or a separate province, in the Midcontinent Realm (Ji and Barnes, 1994).
Early–Middle Ordovician conodont paleobiogeography The andinus-evae Interval Dendrograph, Figure 7. The dendrograph of this interval shows a complex clustering. The weak linkages between areas are probably due to decreased provincialism, especially for the Atlantic Realm. This time is characterized by one of the highest sea levels in the Ordovician (Fortey, 1984; Nielsen, 2004), and this tended to diminish the physical barriers that caused separation and development of geographical endemism. Although a few strong similarities are evident in the dendrograph, associations tend to be only partly coherent, as expected from the conditions of the previous interval. The strongest apparent affinities are between the Precordillera and Ninemile, Nevada, Marathon– El Paso, Texas, and Cow Head, Newfoundland. Another tight subcluster within the same grouping integrates the Swedish localities. At this time the Ouachita Mountains and Baltic Russia were probably linked to the same group, reflecting biofacies similarities. Another isolated group is represented by the Famatina System and Eastern Cordillera, which appear to represent a distinct biogeographic unit because of the extreme position of their cluster branches. Although the remaining localities, including Saint George, Newfoundland, East Greenland, and the Ibex area, Utah, are grouped together with a low level of clustering; they maintain the consistent linkage from the previous interval, making it plausible that they represent a particular biogeographic unit or a similar biofacies within the Midcontinent Realm (Bagnoli and Stouge, 1991). It could also mean that both intervals are not similar to those of the other areas. The laevis-norrlandicus Interval Dendrograph, Figure 8. Two groups are apparent in this cluster. One group consists of the Precordillera, West Newfoundland, Nevada, Central Australia, and Ibex area, Utah. The other cluster integrates south-central Sweden, south-central China, south Norway, Öland, Sweden, and Baltic Russia. After the highest level of inundation, there is a return to a more normal sea level, which is associated with the redevelopment of a well-differentiated conodont paleobiogeography. The high affinity linkages are demonstrated by the low Euclidean distances in the cluster, although the hierarchical structure appears to be weak. The two major recognized groups apparently represent the Midcontinent and Atlantic Realms, respectively. The sinuosa-suecicus Interval Dendrograph, Figure 9. This early Darriwilian interval is the youngest of the biostratigraphical units to be analyzed. A clear hierarchical structure is present in the dendrograph with three well-defined groups. As in previous intervals, the Precordillera clusters with North American localities; in this case it groups with the Marathon (Fort Peña Formation), the Table Head, Newfoundland, and the Antelope Valley, Nevada, regions. A second cluster includes Baltic localities and south-central China, which shows the recurrent association of high latitude localities. In a third group, the Ibex area of Utah links to the Canning Basin of west Australia. This probably reflects the particular location of
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these study sites in shallow, warm-water epeiric seas (Rasmussen and Stouge, 1995). Oceanographic conditions related to the fact that a significant sea-level lowstand was followed by a sealevel highstand is the probable cause of a dendrogram hierarchy with closer matches between different groups than in the previous dendrogram. Co-Evolution of Precordilleran Conodont Faunas Using the similarity index, trends of faunal changes in seven different regions are plotted against that of the Precordillera (Fig. 10). Measures are expressed in percentages of common species after the application of the Jaccard similarity coefficient for the pair of faunas compared from each biostratigraphical interval. This plot does not match sample sites in a hierarchical structure as does the cluster analysis (where linkage tends to occur at higher levels than in pair-group methods on the basis of their highest correlation with any object already in the cluster) so that the direct application of the Jaccard formula reveals only a similarity expressed in percentages of common components but not the geometric relationships among different study sites. For the manitouensis-deltifer interval, the El Paso area reaches the highest similarity, 40%, with Precordilleran faunas, followed by the Ibex area, Utah, more than 30%. Both are typical Laurentian sites. The similarity between the Marathon area and the Precordillera is ~25%, which is a high percentage considering that this comparison is between different facies (slope versus platform facies). A 10% similarity with a North American site is shown by Saint George in Newfoundland, which is characterized by restricted shallow, warm-water faunas. An identical percentage is presented by Öland in Baltoscandia, with where there are cold-water faunas. Other localities from this region, and the deep, cold-water biofacies of Cow Head, show the lowest similarities. A dramatic change of trends takes place within the deltatusproteus interval, where four of seven similarity curves change their directions, from lower to higher percentage values and vice-versa. This is probably related to the global regressive Ceratopyge Event that had an extended duration with several pulses and was associated with an important extinction process, explained below. This tendency is reversed in the communis-elegans interval, with the recurrence of a sea-level highstand. At this time, there was maximum conodont provincialism expressed as a temporal pattern of endemism with ~56% of dissimilarities between extreme biogeographic sites. At this time the Precordilleran fauna reached one of the highest similarities with the Marathon and El Paso faunas (56% and 42%, respectively. During the peak drowning event of the Early Ordovician, which is in the andinus-evae interval, several curves cross, illustrating the breakdown of biogeographic barriers. This is likely to have been caused by the global transgression that was associated with other factors such as rarefaction flux and extinction or evolution of endemic species. The Marathon and El Paso faunas continue to be most similar (~60% and 46%, respectively) to that
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of the Precordillera. The Cow Head fauna also shows high values (35%) of similarity. Other curves maintain approximately the previous values. Through the laevis-norrlandicus and the sinuosa-suecicus intervals, the Marathon fauna maintains high values of similarity (~48%) to the Precordilleran fauna, whereas the Cow Head similarity drops significantly to <20%. It is important to note that the Cow Head slope succession includes mixed slope-platform (from allochthonous blocks) faunas. A similar decline in faunal similarity is shown by the Ibex area curve. Records for Swedish localities fluctuate within a 5% range. In the sinuosa-suecicus interval the sea-level lowstand ended. This probably coincided with another period of maximum provincialism in the conodont paleobiogeography, as shown by the well-separated similarity percentages of the different study sites with respect to the Precordillera. Because particular data from a spectrum of conodont biofacies, from shallowest to deepest, were not filtered in this statistical analysis, the highest curve fluctuations could be related to particular biofacies effects at each study site (Tolmacheva et al., 2004). Nevertheless, it is interesting to note that among all the curves the El Paso curve most consistently shows Precordilleran affinities in that it never drops below 40% for platform facies and maintains similar high values after the lowstand crisis in the deltatus-proteus interval. The Marathon faunas are strikingly similar (50%–60%) above the communis-elegans interval, when a deeper facies was established in the Precordillera, and similar environments can be compared. It is remarkable that in the lower part of the lowstand interval the direction of the Marathon curve (a Laurentian slope site) abruptly changes. The effect of environmental biases in testing affinities is shown by this curve, which represents deep, cold-water faunas in the circumequatorial position of Laurentia compared with shallow, warm-water faunas of the Argentine Precordillera. Paleobiogeographic Indications from Non-Conodont Phosphatic Microfossils Conodont samples frequently yield other phosphatic microfossils such as buttonlike or studlike skeletal elements, some of which may provide useful evidence for paleobiogeographic interpretations. Many of these elements are likely to have been dermal sclerites in complex exoskeletons; other elements, such as the Ordovician genus Ptiloncodus Harris from Laurentia, remain enigmatics. For instance, phosphate sclerites of infaunal animals, such as hadimopanellids or paleoscolecidans, which are now interpreted to be priapulid worms (Conway Morris, 1997; Huang et al., 2004a, 2004b; Vannier and Chen, 2005) have records ranging from the Early Cambrian to the present and were common in benthic marine assemblages of the early Paleozoic. In Cambrian and Lower Ordovician strata, priapulid sclerites have been recorded under such generic names as Palaeoscolex, Protoscolex, and Hadimopanella (e.g., Whittard, 1953; Conway Morris, 1977; Kraft and Mergl, 1989). The phosphatic fossils had a moderate diversity, as shown by the presence of additional taxa such as
Utahphospha and Milaculum, which are frequently associated with conodonts in samples of diverse facies from Laurentia, Baltica, China, Siberia, Australia, and Turkey (e.g., Müller, 1973; Bengtson, 1977; Gedik, 1977; Ethington and Clark, 1981; Boogard, 1989; Müller and Hinz-Schallreuter, 1993; Kraft and Lehnert in Hints et al., 2004). Soft bodies assigned to priapulid worms, preserved as casts on bedding plane surfaces, were recently found in Cambrian strata of the Eastern Cordillera, NW Argentina (García-Bellido and Aceñolaza, 2005) and in Middle Ordovician rocks of the Peruvian Cordillera (Gutiérrez-Marco, 2006, personal commun.). Early Paleozoic priapulid worms were predatory, infaunal, probably carnivorous animals that may have exploited the watersediment interface layer, where meiobenthic organisms provided abundant prey. However, they also were mud-feeders capable of intense burrowing in muddy bottoms (Huang et al., 2004a, 2004b; Vannier and Chen, 2005). The fact that these organisms were bottom-dependent probably makes them significant facies fossils. It may be important paleogeographically that paleoscolecidan priapulids are widespread in Lower and Middle Ordovician strata of Laurentia, including those in the El Paso–Marathon areas (Repetski, 1981; Izold, 1993), but so far they have not been found in the Precordillera. This is in apparent conflict with our statistical analysis of conodont faunas, which suggests a strong linkage between the Argentine Precordillera and southern Laurentia during the earliest Ordovician. However, the absence of priapulid records in the Argentine Precordillera may be due to sampling bias or may be taken as a suggestion that in Early Ordovician time the Precordillera was an exotic terrane that was sufficiently separated from both the Laurentian and Gondwanan margins to prevent migration of these infaunal organisms. Phosphatic vertebrate remains (tesserae, plates, scales) found associated with conodonts may provide valuable information for the interpretation of the paleogeographic position of the Precordillera in the late Middle Ordovician. If one accepts the hypothesis that the Precordillera was an exotic terrane, the part of the Iapetus Ocean that separated the Precordillera from the Gondwanan margin must have been nearly closed in late Middle Ordovician time. This is suggested by the presence of Sacabambaspis janvieri, the oldest armored fish from South America (Gagnier et al., 1986), which is recorded both in the Precordillera and the neighboring Gondwanan basins (Albanesi et al., 1995a; Albanesi and Astini, 2002). This typical Gondwanan fish, which lived in shallow-water environments, is unknown in Laurentia and was apparently unable to swim across oceans. It provides the oldest unequivocal paleontological signal that in late Middle Ordovician time the Precordillera was connected in shallow-water environments with the rest of Gondwana (Smith et al., 2002; Albanesi et al., 2005a). SUMMARY AND CONCLUSIONS The patterns of provinciality shown in the present analysis may be related to plate tectonics, water mass distributions, sealevel changes, and other oceanographic conditions as well as
Early–Middle Ordovician conodont paleobiogeography climate changes. Different causes, single or in combination, may have led to a complex sequence of changing environmental and paleobiogeographic settings during the Early and Middle Ordovician. Despite the fact that the evidence of some causes has been lost in the geological and paleontological records, it is important to try to identify a few potential causes in order to identify the dominant controls. The series of paleogeographic reconstructions of Scotese (2001) illustrate tectonically induced paleogeographic changes, which might have been the cause for biogeographic changes. For instance, if one assumes that the Precordillera was an exotic terrane derived from Laurentia, which began to drift through the Iapetus Ocean in the Cambrian and was accreted to the Gondwanan margin in the Middle Ordovician, its changing position (Figs. 2 and 3) can be used as an explanation for its biogeographic patterns, which notably differ from those of the parts of Gondwana that border the closing Iapetus. Other profound effects may be related to the variable chemical conditions of the oceans, changes in ocean currents, and changes in climatic belts through time (Barnes, 2004). These factors are not always easily recognizable in the rock record, and they could be crucial, considering that biogeographic ranges may be determined more by where pelagic species can maintain viable populations than by dispersal (Norris, 2000). A model of changing sea level used to explain changing patterns of provinciality basically depends on the validity of the eustatic sea-level curve used for the interval studied. In our study we interpret the results of cluster analysis using Nielsen’s (2004) estimates of the relative sea-level changes. For comparative interpretations, the work by Narkiewicz and Szulc (2004) was useful in suggesting how controls on the diachronous migration of conodont faunas in peripheral oceanic areas, through tectonically predisposed pathways and relative sea-level changes, are recorded in the Middle Triassic Muschelkalk in the northern peri-Tethys Ocean. We conclude that during the manitouensis-deltifer interval, provinciality among localities is strong, with two well-defined groups that show tendencies for subdivision into different paleobiogeographic units of lower order. The most outstanding features of this interval are these units, some of which probably correspond to different provinces within the Shallow-Sea and the Open-Sea conodont Realms. In particular, the Precordillera cluster shows similarity to that of the Tropical domain (Midcontinent Province), as represented by the El Paso and Ibex areas. This is in agreement with the idea that the Precordillera was a terrane exotic to Gondwana that originally rifted from Laurentia (Astini and Rapalini, 2003; Rapalini, 2005). The alternative idea that the Precordillera was positioned close to the Gondwanan margin of South America is not supported by our analysis, although it has been suggested (Cocks and Torsvik, 2004, Fig. 5.1; Finney et al., 2003, 2005) that the Precordillera might have been at about the same latitude as southern Laurentia. The dendrograph of the deltatus-proteus interval clearly segregates two clusters. One of the clusters maintains the two North American localities, i.e., the Ouachita Mountains and the
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Marathon–El Paso region, as being akin to the Precordillera Ordovician. Three Swedish localities are gathered together in the same cluster, demonstrating a connection between Laurentian sites and typical Baltic sites, along the borders of the Iapetus Ocean. This phenomenon could be explained by the presence of an oceanic current that allowed an efficient dispersal of the conodont faunas (Löfgren et al., 1999; Christiansen and Stouge, 1999; Barnes, 2004). The other cluster is a weak linkage of diverse localities from distant tectonic plates, such as Laurentia, Baltica, and Gondwana, which probably reflects different conodont biofacies. In the deltatus-proteus interval the Ceratopyge regressive event most probably exerted significant changes in faunal dynamics, such as the well-documented extinction of a significant number of conodont taxa and a following radiation of new lineages during the late Tremadocian (Ethington et al., 1987; Ji and Barnes, 1993). Interestingly, the similarity between the faunas of the Precordillera and the El Paso region increased (Fig. 10) to a coefficient of ~47%, accompanying major evidence of endemicity during this regression. In the evolving pattern of endemism the number of endemic species apparently continues to increase during the communiselegans interval. The higher levels of endemism could be attributed to the development of faunas restricted to discrete epeiric seas when these basins developed at the beginning of a new sea-level highstand. Within this interval, three well-differentiated groups are evident. The Precordillera is strongly linked to the Marathon– El Paso region, and to a lesser extent to Cow Head, Newfoundland, and New South Wales, Australia, forming one distinctive temperate-water group of conodont faunas. Baltoscandian sites constitute a consistent cold-water group, and a third group clusters Saint George, the Ibex area, and East Greenland, which typically correspond to shallow, warm-water facies. The strong linkage between the Precordillera and the Marathon area is surprising (see Fig. 10). That is, if the Precordillera became increasingly removed from Laurentia, one would expect a decrease, rather than a strong increase, in faunal similarity between these regions, although these two sites border the same ocean and perhaps also share a similar latitude. During the andinus-evae interval a higher degree of cosmopolitanism resulted from the breaching of migration barriers as the evae transgression proceeded. A mosaic distribution of study sites in a highly hierarchical structure apparently indicates a common exchange of particular taxa, while previously defined groups maintained their separate identities. This situation probably continued into the laevis-norrlandicus interval with a complex interplay of the Iapetus conodont faunas. The last dendrogram, which represents the sinuosa-suecicus interval, shows the tripartite division of the previously defined warm-, temperate-, and cold-water groups, and an internal partitioning within each group. This reflects the transition from a sea-level lowstand to a sea-level highstand when water mass signatures became less well defined. The recurrent position of the Precordillera, the Marathon area, and the Ouachita Mountains (when comparing similar biofacies) in the same dendrogram group
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is remarkable. The close similarity between the Precordillera and the Marathon area continued to be about as high as before (Fig. 10); hence the presumed greatly increased distance between these regions is not reflected in the conodont faunas. It is important to note that whereas other authors have focused on the identification of endemic and cosmopolitan species associations, we have examined the biogeographic relationships among the sample sites themselves, making it possible to recognize and correlate biogeographic units with specific geographic areas. The non-conodont phosphatic microfossil remains, such as priapulids in Laurentia and the oldest armored fish in Gondwana, provide some constraints on the position of the Precordillera in the Early to Middle Ordovician. Priapulid indications suggest that the Precordillera was well separated from Laurentia in the Early Ordovician, and the fish suggests that it was close to, if not a part of, Gondwana in the late Middle Ordovician. Our analysis shows that the close relationship between the southern Laurentian and Precordilleran conodont faunas in the earliest Ordovician prevails, with relatively little change, into the early Middle Ordovician. The successive long-distance migration of the Precordillera from Laurentia to Gondwana is not supported, nor is it strongly refuted, by the analysis using the Jaccard similarity coefficients of conodont data. The evidence provided by conodonts for the successive long-distance migration of the Precordillera from Laurentia to Gondwana suggested by many authors is not conclusive. However, the faunal data are not inconsistent with the idea of the existence of two areas of platform sedimentation on either side of a spreading ocean in the same climatic belt, which then were greatly affected by the changes in oceanic circulation that accompanied the dramatic tectonism and plate movements in the early Middle Ordovician. ACKNOWLEDGMENTS G.L. Albanesi acknowledges financial support from CONICET, Argentina, for his program on conodont studies. This author received primary support for this project from the Fulbright Commission, the Antorchas Foundation, and the National Agency for Science and Technology (ANPCyT, Argentina; FONCyTPICT 07-11822/15076). Much of the research for the project was carried out at The Ohio State University, Columbus, during the senior author’s stay there in 2001–2002. Special thanks are extended to reviewers J. Repetski and S. Leslie for helpful suggestions for improving the original version of the manuscript. REFERENCES CITED Aceñolaza, F.G., and Toselli, A.J., 1988, El Sistema de Famatina, Argentina: Su interpretación como orógeno de margen continental activo: Congreso Geológico Chileno, 5th, Actas 1, p. 55–67. Aceñolaza, F.G., and Toselli, A.J., 2000, Argentine Precordillera: Allochthonous or autochthonous Gondwanic?: Zentralblatt für Geologie und Paläontologie: Teil, v. 1, p. 743–756. Aceñolaza, F.G., Miller, H., and Toselli, A.J., 2002, Proterozoic–Early Paleozoic evolution in western South America—A discussion: Tectonophysics, v. 354, p. 121–137, doi: 10.1016/S0040-1951(02)00295-0.
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The Geological Society of America Special Paper 466 2010
Black shales: An Ordovician perspective William B.N. Berry† Department of Earth and Planetary Science, University of California, Berkeley, Berkeley, California 94720, USA
ABSTRACT Ordovician black shales have been investigated extensively for at least three basic reasons: (1) they bear one of the system’s most useful biostratigraphic tools, graptolites; (2) many of them are richly petroliferous; and (3) they are relatively widespread. Review and analysis of environments in which modern analogues of black shales may form indicate that they accumulate under hypoxic to anoxic conditions. Investigations of these environments in modern oceans indicate that they form in settings where an abundance of organic matter is present, and oxygen supply and possible resupply is shut off, or those in which oxygen resupply is at a slower rate than oxygen consumption. Ordovician black shales likely accumulated in environments relatively similar to modern hypoxic-anoxic environments under oceanic oxygen minimum zones as well as in shelf sea basins of varying depths and shallow shelf seas in which density stratification shuts off oxygen supplies. INTRODUCTION Black shales have a special significance for those interested in the Ordovician, because graptolites, one of the system’s most useful biostratigraphic tools, occur in many black shale successions. Berry and Wilde (1978) indicated that black shales were relatively more widespread during the Ordovician than during most other intervals in the Phanerozoic, and they suggested that the relative abundance of Ordovician and other lower Paleozoic black shales may be an indication of relatively lower atmospheric oxygen concentrations early in the Paleozoic than later. Furthermore, certain Ordovician black shales are highly productive petroleum source rocks. Because black shales have relatively unique biological and geochemical characteristics, they have been discussed and analyzed extensively (Tyson and Pearson, 1991; Arthur and Sageman, 1994; Wignall, 1994). The many investigations into the origins and biological constituents of black shales indicate that, in general, they accumulated in severely oxygen-deficient †
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(hypoxic to anoxic) environments (Tyson and Pearson 1991; Arthur and Sageman, 1994; Wignall, 1994). Sediments in these oxygen-deficient environments range from organic-rich with total organic carbon (TOC) contents of 8%–10% or more to as little as 1% (Arthur and Sageman, 1994). Tyson and Pearson (1991, p. 1) commented in the introductory remarks to their symposium volume on modern and ancient continental shelf anoxia, “the phenomenon of severe oxygen depletion in continental shelf waters is of great significance to both geologists and marine biologists.” They (Tyson and Pearson, 1991, p. 1) went on to say that geologists seek organic-rich deposits that accumulated in anoxic environments for their petroleum source rock potential, and marine biologists examine hypoxic to anoxic environments, especially those in the inner shelf areas, to determine how and why these environments form. Because when they do form in a number of highly organically productive marine environments, mass mortalities among fish and other commercially valued marine organisms have significant economic consequences. Ferber (2004, p. 1557) described the consequences of development of seasonal coastal zone hypoxia-anoxia in striking terms: “Every summer, death stalks the waters of the northern Gulf of Mexico. A New
Berry, W.B.N., 2010, Black shales: An Ordovician perspective, in Finney, S.C., and Berry, W.B.N., eds., The Ordovician Earth System: Geological Society of America Special Paper 466, p. 141–147, doi: 10.1130/2010.2466(09). For permission to copy, contact
[email protected]. ©2010 The Geological Society of America. All rights reserved.
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Jersey–size swath of sea becomes depleted of oxygen, suffocating millions of crabs and other denizens of the sea floor.” He (Ferber, 2004) drew attention to the economic losses that oxygen depletion in these waters (called the Gulf of Mexico Dead Zone) cause. The economic importance of the Gulf of Mexico so-called Dead Zone and other, similar oxygen deficient environments has led to numerous analyses of them (Rabalais and Turner, 2001; Diaz, 2001; Rabalais et al., 2002a; Ferber, 2004; Service, 2004; Neretin, 2006). These studies of modern shelf sea oxygen-deficient environments provide useful insights into an understanding of environmental conditions in which ancient black shales accumulated. In addition to the many investigations and analyses of modern shelf sea hypoxic-anoxic environments, critical evaluation of the development of permanent lowoxygen environments at the outer edges of certain continental shelves (oxygen minimum zones or OMZs) and the upper part of adjacent continental slopes was carried out by Kamykowski and Zentara (1990), Levin (2003), and Helly and Levin (2004). Hypoxic-anoxic depositional environmental conditions present in the modern environments reviewed here are used as clues to suggest depositional environments for a number of Ordovician black shales. HYPOXIC-ANOXIC ENVIRONMENTS IN MODERN SHELF SEAS Studies of the development of hypoxic to anoxic waters that form in the Gulf of Mexico during the summer months have enhanced understanding of shallow-shelf sea oxygen-deficient environments (Rabalais and Turner, 2001; Diaz, 2001; Rabalais et al., 2002a, 2002b). In a review of the global occurrence of these environments, Diaz (2001) pointed out that their formation depends upon two primary factors: (1) decomposition of a relatively large quantity of organic matter, and (2) isolation of waters at some depth beneath the surface so that oxygen is not resupplied to these waters after the available oxygen has been consumed by decomposition of organic matter. Diaz (2001, p. 276) stated that when oxygen supply is cut off or consumption rate exceeds resupply, oxygen concentrations decline beyond an availability that can sustain most animal life. This condition of low oxygen is known as hypoxia. The point at which animals suffocate varies, but generally effects start to appear when oxygen drops below 2 mg of oxygen per liter. Rabalais et al. (2002a, 2002b) discussed development of oxygen-poor environments in the Gulf of Mexico, noting that these environments include a broad area, which their studies show expands and contracts yearly. Hypoxia within that area “may encompass from 10% to over 80% of the water column, and may reach to within 2 m of the surface in a 10-m water column, or to within 6 m of the surface in a 20-m water column” (Rabalais et al., 2002a, p. 132). Hypoxic waters extend from the shoreline to “as far as 125 km offshore and in waters up to 60 m deep” (Rabalais et al., 2002a, p. 133).
Gulf of Mexico hypoxia develops in nearshore waters (Diaz, 2001; Rabalais and Turner, 2001; Rabalais et al., 2002a, 2002b) when these waters become stratified during the spring and summer as fresh river water flowing into the Gulf is warmed by the sun to form a less dense layer overlying colder, saltier, deeper waters. This density stratification prevents both oxygen penetration and oxygen resupply to waters beneath the surface layer. The several investigations of seasonal development of modern oxygen-deficient shelf sea environments have identified two fundamental factors that are essential for documenting ancient black shale depositional environments: (1) a rich supply of organic matter, and (2) well-developed density stratification that shuts off oxygen supply and resupply. When these two basic conditions are in place, the result is development of hypoxic to anoxic environments in waters above the water-sediment interface as well as in sediments below that interface. When all available oxygen is consumed in the waters above the water-sediment interface and in sediment below it by bacterial decomposition of organic matter, then bacterial nitrate reduction of organic matter, followed by bacterial sulfate reduction, takes place. In environments in which organic matter is so abundant that bacterial sulfate reduction consumes the available sulfate, then methanogenesis takes place. At sites where bacterial sulfate reduction occurs in the sediments, sulfur that forms during sulfate reduction may be used by chemosynthetic organisms. These organisms may generate microbial mats at or within a few centimeters of the watersediment interface (Gallardo, 1977; Levin et al., 2002). De Diego and Douglas (1999) investigated development of oxygen-deficient shelf sea and adjacent continental shelf margin sediments in the Gulf of California. Their study provides insight into seasonal development of an alternation between oxygenpoor and oxic depositional environments and thus may be especially useful for interpretations of ancient black shale sequences. De Diego and Douglas (1999, p. 143–144) collected and analyzed “dark colored, thinly-bedded, organic-rich muds accumulating under low oxygen conditions” in a continental shelf basin in the Gulf of California as well as under an OMZ along the adjacent continental slope, and they (De Diego and Douglas, 1999, p. 453) commented that these muds have “features similar to black shales found in the Cretaceous and offer insights into the depositional processes recorded in these mudrocks.” De Diego and Douglas (1999, p. 457 and their fig. 4) found that they could relate sediment microfabrics to bottom water oxygen levels. Benthic foraminifers occur in finely laminated sediments deposited under waters with 0.1 mL of oxygen per liter, and small polychaetes and tiny nematodes were found in “semi-laminated” sediments deposited under waters with 0.1–0.2 mL of oxygen per liter (Douglas, 2003, oral commun.). De Diego and Douglas (1999, p. 462) stated that an oxygen minimum zone developed along the continental slope and that seafloor oxygen values under that oxygen minimum zone ranged from ~0.1 mL per liter to essentially anoxic over the past 600 yr. De Diego and Douglas (1999, p. 462) concluded that their analysis of oxygen-related sediment microfabrics indicates “that the thickness and intensity
Black shales: An Ordovician perspective of the OMZ has varied in cyclical fashion on decadal to centennial time scales.” MODERN OXYGEN MINIMUM ZONES UNDER UPWELLING CONDITIONS Helly and Levin (2004) recorded investigations of hypoxicanoxic sediments accumulating under OMZs that form along continental margins of the eastern Pacific Ocean, the southeastern Atlantic Ocean, the Arabian Sea, and the Bay of Bengal. Helly and Levin (2004, p. 1160) noted that these regions constitute “all the major open-ocean, upwelling-induced OMZs of the world ocean.” Anderson et al. (1982) documented the decline in oxygen concentration with depth from the ocean surface in OMZ waters off the Peru coast in the eastern Pacific. From their analyses of oxygen in OMZ waters, Helly and Levin (2004, p. 1164) cited latitudinal variation “in the eastern Pacific upper and lower OMZ boundaries,” which, they suggested, “reflect combined effects of surface productivity, water mass age, and circulation.” They also suggested (p. 1164) that an OMZ may be thicker and have lower oxygen concentration under conditions of “increased upwelling, high productivity, and a greater oxygen demand along with sluggish circulation.” In addition, they commented (p. 1165) that only rarely have investigations been made into organisms living in modern OMZ environments. Levin (2003) did note that foraminifers and nematodes have been found in environments in which the oxygen level is <0.2 mL per liter. Childress and Seibel (1998) described several adaptations that marine animals use to survive in OMZ waters. Thuesen (2002) and Kirouac et al. (2002) described “jelly fish blooms” in hypoxic waters in Puget Sound, Washington, and Thuesen (2002) reported finding several species of scyphomedusae, hydromedusae, and siphonophores that survive without oxygen for several hours. Jorgensen and Gallardo (1999) recorded the presence of widespread mats of sulfide-oxidizing bacteria covering sediments accumulating under the core of OMZ zone waters. Levin et al. (2002) examined sediments under the Peru margin OMZ and recorded a bottom water oxygen value of 0.02 mL/L under one site or station. They (Levin et al., 2002, p. 21) commented that one surprising result of their investigation is that all benthic faunal groups found at the site (including agglutinated and calcareous foraminifers and metazoan meiofauna and macrofauna) with the most minimal oxygen availability “exhibited their highest densities,” and available food resources were greatest there. Based on this relationship, Levin et al. (2002, p. 210) stated “that food rather than oxygen, controls infaunal densities.” Helly and Levin (2004, p. 1165) commented that among organisms living under OMZs around the world, “many OMZ species are new to science and some have exhibited unusual evolutionary novelty. An inconspicuous gutless oligochaete discovered in the Peru OMZ, for example, supports more types of nutritional bacterial symbionts (3 for certain and possibly 5) than reported for any
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other invertebrate.” Levin (2003) concluded that a number of chemosynthetic organisms appear to be living in OMZ-related environments in which sulfide oxidation occurs, which led Levin et al. (2002, p. 21) to suggest that “chemoautotrophy may represent a significant trophic pathway for metazoans within OMZs.” These observations of modern OMZ organisms provide clues to possible food relationships among organisms and the environments in which they live, as, for example, graptolites, that lived near or within ancient OMZs and OMZ-related environments. Helly and Levin (2004) pointed out that their studies showed that the upper OMZ boundary may shift seasonally or interannually. For example, El Nino events have led to oxygenation of the upper part of the Peru OMZ. Global warming, Helly and Levin (2004, p. 1166) suggested, can lead to lowered oxygen content in the oceans, which could result in OMZ expansions. OMZs can be barriers to movements of species and “OMZ expansion or shrinkage may promote the evolution of species and genetic diversity.” Furthermore, positions and extent of OMZs will change over time with changes in ocean circulation, productivity, and temperature. For example, Service (2004) described seasonal changes in dissolved oxygen concentration in upwelling waters off the Oregon coast. The changes in position, extent, and duration of OMZs may have been significant influences in the evolution of certain organisms (such as the graptolites) that lived in ancient oceans. The descriptions of modern hypoxic-anoxic environments in both shelf and oxygen minimum zone settings provide clues to understanding development of hypoxic-anoxic environments in ancient oceans as well as the accumulation of both fossil-bearing (dominantly graptolitic in the Ordovician) and nonfossiliferous ancient black shales. The oceanographic and biological processes discussed by De Diego and Douglas (1999) within the Gulf of California that influence shelf basin and shelf margin environments may be analogues of processes that took place during the Ordovician along many shelves and adjacent shelf margins. Analyses of the sediments that accumulated in the California borderland basins (Gorsline, 1996) and those deposited under OMZ waters across the Peru shelf (Emeis et al., 1991; Levin et al., 2002) provide certain insights into interpreting environments of deposition of Ordovician black shales that accumulated under similar depositional environments. All studies of sedimentary layers encountered in cores of basin sediments and cores of sediments that accumulated under oxygen minimum zones show that, over time, layers that accumulated in anoxic conditions are interlayered stratigraphically with those deposited under oxic conditions in a pattern documented by De Diego and Douglas (1999, p. 458) in the Gulf of California shelf basin. ORDOVICIAN BLACK SHALES: SELECTED EXAMPLES In turning to the geologic record of environments in which black shales may form, the modern record suggests that most attention be directed toward paleogeographic reconstructions that indicate ancient oceanic upwelling sites and both shallow shelf
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settings and shelf basins. In regard to locating sites of oceanic upwelling in the geologic record, Wilde (1991) and Wilde et al. (1990) pointed out that oceanic upwelling under which OMZs may form is controlled primarily by the major wind patterns and related wind driven currents. Wilde et al. (1997) amplified that statement by noting that the planetary oceanic vertical advection systems are fixed latitudinally by the rotation of the earth and the resultant planetary wind patterns in a band centered about 30 degrees N and S (convergences, downsinking and low productivity) and the Equator and 60 degrees N and S (divergences, upwelling and high productivity). Additionally, upwelling phenomena are observed at 30 degrees N and S on western coastlines. Seasonal conditions due to Monsoonal or local weather patterns or unique land-ocean boundary conditions can produce either upwelling or downsinking which may blur or override the planetary signature. For seasonal cases, the geologic record will preserve conditions during major sedimentary events.
This discussion indicates that understanding plate positions and paleogeographies are significant in locating ancient black shale sequences. Wilde (1991) used fundamental oceanographic circulation principles, major Ordovician lithofacies patterns, and plate positions proposed by Scotese (1986) to interpret major Ordovician ocean currents (see his Fig. 2 for the Middle Ordovician). He concluded that the western side of Laurentia, the eastern side of Australia, and areas within South China probably were sites of continental margin upwelling during the Ordovician. Certain well-studied Ordovician black shale sequences that bear abundant graptolites appear to have formed in the environments within or beneath an OMZ that developed as a result of upwelling along a continental shelf margin (Scotese, 1986, his Fig. 2). Examples include the Vinini Formation in Nevada (Finney and Berry, 1997; Finney et al., 1997), the Phi Kappa Formation in Idaho (Dover et al., 1980), the Moffat Shale Group in the South Uplands of Scotland (Leggett, 1980; Zalasiewicz, 2001), and Cambrian-Ordovician black shales such as the Alum-Dictyonema in Scandinavia (Quinby-Hunt and Wilde, 1994, 1996; Wilde et al., 1997). The first three of these well-known graptolite-rich successions accumulated on Laurentian plate margins. Zalasiewicz (2001) reviewed the many investigations of the Scottish South Uplands Ordovician-Silurian graptolitic succession, noting that the graptolite-bearing strata there occur as discrete black shale units with medium- to coarse-grained graywackes in thrust fault– bounded “packages.” He (Zalasiewicz, 2001, p. 244) noted that “analysis of the graptolites between thrust slices showed an overall younging direction for each thrust slice.” He (p. 244) went on to state that “the Moffat Shales represent slowly accumulating pelagic and hemipelagic (partly radiolarian) oozes on a deep ocean floor, the common presence of fossilized plankton (graptolites) associated with a lack of benthos or bioturbation indicating sea floor anoxia.” Although many black shale layers in these successions bear abundant graptolites, these graptolite-bearing shales commonly
are interbedded with gray to black shales that bear little, if any, fauna. Zalasiewicz (2001, p. 238) commented from his review of British graptolite-bearing sequences that many of them include black, partly organic-rich, finely laminated shale layers bearing graptolites interbedded with lighter colored, weakly laminated to massive beds barren of fossils. These Ordovician graptolitebearing sequences are relatively similar to the sedimentary record found in cores from the Gulf of California shelf slope (De Diego and Douglas, 1999) and under the Peru OMZ (Levin et al., 2002). The light-colored, weakly laminated to semi-laminated sediments accumulated in these depositional sites at times when oxic waters flowed into the depositional site to resupply oxygen to hypoxicanoxic environments. Return of anoxic conditions led to renewed black mud accumulation in fine laminae (De Diego and Douglas, 1999). Similar ocean processes that result in a change from hypoxic-anoxic environments to oxic may be invoked to account for interbedded black shales, which commonly bear graptolites in Ordovician successions, with lighter colored shales that are semilaminated and may bear traces of bioturbation. Wilde et al. (1997, p. 126) noted that the Alum-Dictyonema black shales suggest that “the influence of latitude is demonstrated by an examination of the concentrations of the anoxia indicators” in these shales. Chemical analyses of these Lower Ordovician black shales from the Oslo region, Norway, indicates that they accumulated in an anoxic environment (Wilde et al., 2004). During the Late Cambrian and Early Ordovician, the Oslo region lay within the area of planetary upwelling, leading to development of an OMZ and accumulation of organic-rich black shales under it. By the Middle Ordovician, that area had drifted outside the area of planetary upwelling, as the shale geochemistry indicates that high organic productivity and resultant anoxic environments were no longer present (Wilde, 1991; Wilde et al., 1997). BASINS WITHIN THE PLATFORM Many Ordovician shelf seas included areas tectonically similar to the California borderland and the Gulf of California. Basins of different depths occur on these shelves today (De Diego and Douglas, 1999, their Fig. 1). Tectonism during the Ordovician along shelf margins resulted in development of relatively similar basins. Pysklywec and Mitrovica (2000) described how widespread marine transgressions occur at a time when foreland and intracratonic basins form as a result of mantle flow under platform margins. As a subducting slab begins to descend underneath a platform or craton, a foreland basin develops between the platform and an orogenic land. Continued subsidence of the subducting slab leads expansion of the basin across the platform as well as to development of shallow basins on the interior part of the platform (Pysklywec and Mitrovica, 2000, their fig. 6). Taconic foreland basin development on the eastern side of Laurentia as a result of Taconian tectonism in the latter part of the Ordovician, and formation of basins in the interior of the Laurentian plate, are consistent with the Pysklywec and Mitrovica (2000) model for formation of sedimentary basins on cratons or plates.
Black shales: An Ordovician perspective The black shales that constitute the Middle to Upper Ordovician Utica Shale (discussed as the Utica magnafacies by Lehmann et al., 1995) were deposited in the Taconic foreland basin, and then they spread widely across a broad area of Laurentia. These shales have been identified as the source rocks for ~600 million barrels of oil and >1.5 trillion cubic feet of natural gas obtained from Upper Cambrian, Ordovician, and Lower Silurian reservoir strata (Ryder et al., 1998). Lehmann et al. (1995) recognized five organic-rich black shale units deposited in hypoxic to anoxic environments within their Utica magnafacies, and a unique graptolite zonal fauna occurs in each one. Each black shale unit formed as organic-rich mud on the basin slope following an interval of basin subsidence and a steepened slope (Lehmann et al., 1995, p. 718). Witzke (1987, p. 231) indicated that turbulence and related upwelling may occur along a basin margin at times when basin slopes are abruptly steepened. Upwelling of nutrientrich waters at such sites may lead to increased organic productivity and, ultimately, anoxic conditions and black mud deposition (Witzke, 1987, p. 231). Lehmann et al. (1995, p. 720) pointed out that the western flank of the Taconic foreland basin shifted westward during the Ordovician, and that as it did, organic-rich mud deposition moved westward as well. Westward migration of the Taconic foreland basin during the latter part of the Ordovician is consistent with the Pysklywec and Mitrovica 2000) model for basin development within cratons. The Maquoketa Formation occurs widely across the central part of the present-day United States Midcontinent. Witzke (1987) described three facies belts within the Maquoketa, one of which is an organic-rich, graptolite-bearing shale. Laminated brown to black shales within this facies seem to have been deposited in anoxic environments that developed when density stratified waters developed across the area. Witzke (1987) suggested that density stratification formed during what he termed quasiestuarine circulation. In that circulation pattern, an influx of fresh water and siliclastic sediment from the Taconic land sources spread out from the source. Surface fresh water would have overlain denser oceanic water, which, Witzke (1987, p. 239) suggested, could have included OMZ water flowing in from shelf or platform margin sites (his Fig. 3a). Witzke (1987) also noted that hypoxic to anoxic environments spread widely across the Laurentian platform during a Late Ordovician transgression. Black shales, many of which are organic rich, accumulated as the basal deposits of marine transgressions across shelves and into shelf basins during the Ordovician (Witzke, 1987; his Fig. 3 indicates a mechanism for this depositional pattern). Many of these basal shales bear graptolites (Leggett, 1980; Sharma et al., 2003). Zalasiewicz (2001, p. 248) noted that “graptolites seemed to flourish at times of transgression, when ocean waters tended to anoxia.” Wignall (1994, p. 87) noted a close relationship of certain black shales with marine transgressions and stated that “many of the world’s best source rocks belong to the transgressive black shale category.” Potter et al. (2005) discussed black shale deposition during marine transgressions, pointing out that as marine waters spread across shelf lands, water depths initially
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were shallow and within the photic zone. Accordingly, bottom life initially was prolific and the demand for dissolved oxygen was high, both by growing organisms and by those involved in decomposition of dead organisms. Potter et al. (2005) also indicated that even shallow depressions on the shelf could trap decomposing organic matter and thus promote development of hypoxic-anoxic environments. Potentially, OMZ waters may flow across a subsiding shelf during a transgression and so lead, eventually, to development of hypoxic-anoxic environments. The Williston Basin, in the north-central United States and adjacent Canada, formed as a broad, essentially semicircular depositional basin on the Laurentian plate in the Late Cambrian when siliciclastic deposition took place in shallow marine environments across a basin that had modest topographic relief (Dow, 1974; Heck et al., 2005). Initial marine sediment deposition ceased early in the Ordovician, and some erosion took place (Dow, 1974; Heck et al., 2005). Siliciclastic deposition in shallow marine environments redeveloped (the Black Island Formation of the Winnipeg Group) during the Middle Ordovician. Organic-rich shales (the Icebox Formation of the Winnipeg Group) were deposited stratigraphically above the Black Island siliciclastics (Dow, 1974; Heck et al., 2005). These organicrich shales have been identified as the source rocks for lower Paleozoic petroleum reservoirs (Dow, 1974; Williams, 1974). Most likely, these organic-rich shales were deposited in shallow hypoxic to anoxic environments under density-stratified ocean waters. Fresh waters from surrounding lands would have sealed off oxygen supply, as is seen in the modern Gulf of Mexico. The kerogen in these organic-rich shales is “dominated by small, thin-walled Prasinophyte alginite and a weakly autofluorescing amorphous matrix” (Seibel and Bend, 2001, p. 082-1). These authors (2001, p. 082-1–082-2) also stated that “studies elsewhere have shown that the presence of Prasinophyte alginite is strongly correlated with the occurrence of marine sediments deposited under dysoxic to anoxic conditions.” Limestones and dolomites followed stratigraphically by anhydrite (the Big Horn Group with the basal Red River Formation) overlie the organic-rich shales (Dow, 1974; Williams, 1974; Heck et al., 2005). Where fractured and overlain by anhydrite, these limestones and dolomites are petroleum reservoirs (Dow, 1974; Williams, 1974). Chen et al. (2004) described Late Ordovician facies patterns and paleogeography of the Yangtze Platform in South China. The Late Ordovician Wufeng Shale in that area is black, finely laminated, organic rich, and graptolite bearing. It accumulated in a shallow basin on the platform. Analyses of the brachiopod faunas on the shelf suggest that the water depth in this basin could have been about as deep as the deeper parts of the modern Gulf of Mexico Dead Zone. The Late Ordovician graptolite-bearing black shales on the Yangtze Platform seem to have accumulated in a setting that was deep enough for development of relatively stable density stratification, which led to sediment accumulation in anoxic conditions. The presence of nearby lands from which fresh waters probably flowed into the shelf sea seems to have
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been a factor in the development of density stratification, as it is in the Gulf of Mexico. During the Ordovician, certain organic-rich petroliferous shales accumulated in a number of relatively shallow marine depositional environments on platforms in which organic productivity was high and a density-stratified water mass persisted. Perhaps the best studied of such successions is the Estonian oil shales, which have been the subject of detailed studies for decades, including a 1968 United Nations symposium on the development and use of these shales. The shales developed from organic-rich mud deposition in shallow marine environments (Lille, 2003). The term kukersite has been given to the kerogen in these shales following studies by the botanist Zalensky in 1917 (Lille, 2003, p. 253). Lille (2003) reviewed analyses of the kerogen in the Estonian black shales and suggested that the shales formed under hypoxic to anoxic conditions and that the kerogen developed from growth of cyanobacteria (Lille, 2003, p. 255). He stated (p. 255) that “we speculate here that the dominant role in its (the kerogen) formation could have been played by green sulfur bacteria using in the process of photosynthesis hydrogen sulfide as electron donor.” Lille’s (2003) study suggests that the Estonian oil shales accumulated in an area of high organic productivity, which was reflected primarily in the growth of prasinophyte algae, and that extensive microbial mats developed on or in the sediments in which bacterial sulfate reduction was taking place in anoxic environments in a relatively shallow shelf sea. Density stratified waters in that sea probably caused development of these anoxic environments, as described and illustrated by Witzke (1987, p. 231, his Fig. 1). Interestingly, Lille (2003, p. 256) commented in regard to microbial mat productivity that “multilayered microbial communities (microbial mats) along with tropical rain forests and coral reefs are among the most productive ecosystems in the world.” CONCLUSIONS Analysis of modern oceanic hypoxic-anoxic environments has demonstrated that the two fundamental conditions requisite for development of these environments are (1) an abundance of organic matter, and (2) a water mass to which oxygen supply and any resupply are greatly diminished or cut off. The latter requirement commonly is facilitated by density stratification. Density stratification in most modern hypoxic-anoxic settings is seasonal, as it is, for example, in the Gulf of Mexico and the Gulf of California. Modern OMZs change over time with changes in oceanographic circulation and stratification. Sediments beneath them range from finely laminated black muds that may have mats formed by sulfur-oxidizing bacteria at or near the water-sediment interface to light-colored, bioturbated muds that accumulated as oxic waters flowed over the site of accumulation. The study of modern hypoxic-anoxic environments has generated clues to an understanding of ancient (in this study, Ordovician) black shales, some of which bear rich graptolite faunas. Many well-studied Ordovician black shales formed under OMZs that developed along continental or plate margins, as do modern OMZs. Other
Ordovician black shales formed in a variety of shelf sea settings that range from relatively deep to shallow basins on the shelf to shallow to moderate-depth shelf sea environments that formed under density stratified waters. The Ordovician stratigraphic record of shelf sea shale successions in which black shales are prominent is similar to that recorded for the sedimentary column in the modern Gulf of California in which finely laminated black and light-colored muds are interbedded. The light-colored sediments formed during an influx of oxygen into the depositional environment. Algae and bacteria seem to have provided the greatest quantity of organic matter to surface ocean waters during the Ordovician. Chemosynthetic, commonly sulfur-oxidizing bacteria may have been prominent in Ordovician hypoxic-anoxic environments, as they are in similar modern environments. Certain graptolites, for example, may have had bacterial symbionts, as does a gutless oligochaete in the Peru OMZ. Others may have been enmeshed in gelatinous material similar to that of modern gelatinous zooplankton, and relished low oxygen waters as do certain modern gelatinous zooplankton. Investigations of the many features, both biological and chemical, of modern hypoxicanoxic environments suggest that closely similar environments led to development of Ordovician black shale sequences. REFERENCES CITED Anderson, J.J., Okubo, A., Robbins, A.S., and Richards, F.A., 1982, A model for nitrite and nitrate distributions in oceanic oxygen minimum zones: DeepSea Research, v. 29, p. 1113–1140, doi: 10.1016/0198-0149(82)90031-0. Arthur, M.A., and Sageman, B.B., 1994, Marine black shales: Depositional mechanisms and environments of ancient deposits: Annual Review of Earth and Planetary Sciences, v. 22, p. 499–551. Berry, W.B.N., and Wilde, P., 1978, Progressive ventilation of the oceans— An explanation for the distribution of the Lower Paleozoic black shales: American Journal of Science, v. 278, p. 257–275. Chen Xu, Rong Jia-yu, Li Yue, and Boucot, A.J., 2004, Facies patterns and geography of the Ytangtze region, South China, through the Ordovician and Silurian transition: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 204, p. 353–372. Childress, J.J., and Seibel, B.A., 1998, Life at stable low oxygen levels: Adaptations of animals to oceanic oxygen minimum layers: Journal of Experimental Biology, v. 201, p. 1223–1232. De Diego, T., and Douglas, R.G., 1999, Oxygen-related sediment microfabrics in modern “black shales,” Gulf of California, Mexico: Journal of Foraminiferal Research, v. 29, p. 453–464. Diaz, R.J., 2001, Overview of hypoxia around the world: Journal of Environmental Quality, v. 30, p. 275–281. Dover, J.H., Berry, W.B.N., and Ross, R.J., Jr., 1980, Ordovician and Silurian Phi Kappa and Trail Creek Formations, Pioneer Mountains, Central Idaho—Stratigraphic and Structural Revisions, and New Data on Graptolite Faunas: U.S. Geological Survey Professional Paper 1090, 54 p. Dow, W.G., 1974, Application of oil-correlation and source-rock data to exploration in Williston Basin: American Association of Petroleum Geologists Bulletin, v. 58, p. 1253–1262. Emeis, K.-C., Whelan, J.K., and Tarafa, M., 1991, Sedimentary and geochemical expressions of oxic and anoxic conditions on the Peru Shelf, in Tyson, R.V., and Pearson, T.H., eds., Modern and Ancient Continental Shelf Anoxia: Geological Society [London] Special Publication 58, p. 155–176. Ferber, D., 2004, Dead zone fix not a dead issue: Science, v. 305, p. 1557, doi: 10.1126/science.305.5690.1557. Finney, S.C., and Berry, W.B.N., 1997, New perspectives on graptolite distributions and their use as indicators of platform margin dynamics: Geology, v. 25, p. 919–922, doi: 10.1130/0091-7613(1997)025<0919:NPOGDA> 2.3.CO;2.
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Rabalais, N.N., and Turner, R.E., 2001, Hypoxia in the northern Gulf of Mexico: Description, causes and change, in Rabalais, N.N., and Turner, R.E., eds., Coastal Hypoxia: Consequences for Living Resources and Ecosystems: Washington, D.C., American Geophysical Union, Coastal and Estuarine Studies, no. 58, p. 1–36. Rabalais, N.N., Turner, R.E., and Scavia, D., 2002a, Beyond science into policy: Gulf of Mexico hypoxia and the Mississippi River: Bioscience, v. 52, p. 129–142, doi: 10.1641/0006-3568(2002)052[0129:BSIPGO]2.0.CO;2. Rabalais, N.N., Turner, R.E., and Wiseman, W.J., Jr., 2002b, Gulf of Mexico hypoxia, A.K.A. “The Dead Zone”: Annual Review of Ecology and Systematics, v. 33, p. 235–263, doi: 10.1146/annurev.ecolsys.33 .010802.150513. Ryder, R.T., Burruss, R.C., and Hatch, J.R., 1998, Black shale source rocks and oil generation in the Cambrian and Ordovician of the central Appalachian Basin, USA: American Association of Petroleum Geologists Bulletin, v. 82, p. 412–441. Scotese, C.R., 1986, Phanerozoic Reconstructions: A New Look at the Assembly of Asia: University of Texas Institute for Geophysics, Technical Report 66, 54 p. Seibel, C., and Bend, S., 2001, Organofacies and source potential of the Middle Ordovician Winnipeg Formation within southern Saskatchewan (Abstract): Rock the Foundation Convention, 18–22 June 2001, Canadian Society of Petroleum Geologists, p. 083–084. Service, R.F., 2004, New dead zone off Oregon coast hints at sea change in currents: Science, v. 305, p. 1099, doi: 10.1126/science.305.5687.1099. Sharma, S., Dix, G.R., and Riva, J.F.V., 2003, Late Ordovician platform foundering, its paleooceanography and burial, as preserved in separate (eastern Michigan Basin, Ottawa Embayment) basins, southern Ontario: Canadian Journal of Earth Sciences, v. 40, p. 135–148, doi: 10.1139/e02-099. Thuesen, E.V., 2002, Ecophysiology of gelatinous zooplankton in estuarine hypoxia: Abstract for 2002 Annual Meeting of Pacific Estuarine Research Society in Portland, Oregon, p. 1. Tyson, R.V., and Pearson, T.H., 1991, Modern and Ancient Continental Shelf Anoxia: Geological Society [London] Special Publication 59, 470 p. Wignall, P.B., 1994, Black Shales: Oxford, UK, Clarendon Press, 124 p. Wilde, P., 1991, Oceanography in the Ordovician, in Barnes, C.R., and Williams, S.H., eds., Advances in Ordovician Geology: Geological Survey of Canada Paper 90-9, p. 283–298. Wilde, P., Quinby-Hunt, M.S., and Berry, W.B.N., 1990, Vertical advection from oxic or anoxic water from the main pycnocline as a cause of rapid extinctions, in Kauffman, E.G., and Walliser, O.H., eds., Lecture Notes in Earth Sciences, v. 30: Berlin, Springer-Verlag, p. 85–98. Wilde, P., Quinby-Hunt, M.S., and Erdtmann, B.-D., 1997, Prediction of potentially metalliferous organic-rich shale locales using paleo-oceanographic and paleo-geographic techniques, in Papunen, H., ed., Mineral Deposits: Research and Exploration, Where Do They Meet?: Rotterdam, A.A. Balkema, p. 125–128. Wilde, P., Lyons, T.W., and Quinby-Hunt, M.S., 2004, Organic carbon proxies in black shales: Molybdenum: Chemical Geology, v. 206, p. 167–176, doi: 10.1016/j.chemgeo.2003.12.005. Williams, J.A., 1974, Characterization of oil types in Williston Basin: American Association of Petroleum Geologists Bulletin, v. 58, p. 1243–1252. Witzke, B.J., 1987, Models for circulation patterns in epicontinental seas applied to Paleozoic facies of North American craton: Paleoceanography, v. 2, p. 229–248, doi: 10.1029/PA002i002p00229. Zalasiewicz, J., 2001, Graptolites as constraints on models of sedimentation across Iapetus: A review: Proceedings of the Geologists’ Association, v. 112, p. 237–251.
MANUSCRIPT ACCEPTED BY THE SOCIETY 6 OCTOBER 2009
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The Geological Society of America Special Paper 466 2010
Paleogeographic, paleoceanographic, and tectonic controls on early Late Ordovician graptolite diversity patterns Daniel Goldman† Wu Shuang-Ye Department of Geology, University of Dayton, 300 College Park, Dayton, Ohio 45469, USA
ABSTRACT The Katian Age (early Late Ordovician) was a time of significant decline in marine biodiversity, but whether this decline was a real phenomenon or an artifact of the relatively few studies devoted to this interval requires further research. We examined the pattern of graptolite faunal changes across the boundary between the Climacograptus bicornis and Diplacanthograptus caudatus graptolite zones in North America and on several other continents. A sharp decline in species diversity occurs in the Appalachian Basin. Scores for normalized diversity dropped from 20 in the C. bicornis Zone to 7 in the D. caudatus Zone. Only 11% of the species present in the C. bicornis Zone carry over into the D. caudatus Zone. A similar pattern occurs in central Oklahoma. Regions at higher paleolatitude, such as Wales and Baltoscandia, exhibit low graptolite diversity in lower Katian strata, and then diversity declines further in higher strata. In other regions at low paleolatitude, such as Australasia and Scotland, however, diversity is fairly constant across this interval (although the percentage of carryover taxa remains low). We conclude that seawater temperature change or disruption of the oceanic density structure, which might accompany temperature change, provides explanations for the similarity between Laurentian and higher paleolatitude diversity patterns. Flooding of the Laurentian craton through the Sebree Trough by cool, subpolar Iapetus seawater may have adversely affected graptolite diversity there. Regions at high paleolatitudes likely underwent cooling associated with Katian climate deterioration. Thus seawater cooling, albeit driven by different mechanisms, may have produced similar diversity patterns at different paleolatitudes. INTRODUCTION Ordovician biodiversity studies have overwhelmingly focused on two remarkable features of the fossil record, the nearly fourfold increase in marine biodiversity that occurred during the Ordovician and the catastrophic mass extinction that marked the end of the period. Considerably less attention has been paid to †
E-mail:
[email protected].
the smaller peaks and valleys that make up the overall curve, although Sepkoski (1995) pointed out an interesting upper Caradoc (Katian Stage) diversity drop among marine invertebrates (Fig. 1) and noted that more research was needed to discover if this decline was a real phenomenon or an artifact of fewer studies and fewer workers devoted to this interval. Over the past decade a series of studies has examined faunal turnover and biodiversity changes during the late Sandbian and early Katian. In North America, Patzkowsky and Holland (1993,
Goldman, D., and Wu Shuang-Ye, 2010, Paleogeographic, paleoceanographic, and tectonic controls on early Late Ordovician graptolite diversity patterns, in Finney, S.C., and Berry, W.B.N., eds., The Ordovician Earth System: Geological Society of America Special Paper 466, p. 149–161, doi: 10.1130/2010.2466(10). For permission to copy, contact
[email protected]. ©2010 The Geological Society of America. All rights reserved.
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Number of Genera
Late Caradoc Diversity Drop
1000
Cambrian Fauna 0
M
U
Cambrian
T
A
L
L
Ordovician
C
A
L
W
L
Silurian
Figure 1. Lower Paleozoic marine diversity patterns. Note the prominent diversity decline prior to the Late Ordovician mass extinction. X-axis division abbreviations are (in order) M—Middle; U—Upper; T—Tremadoc; A—Arenig; L—Llanvirn; L—Llandeilo; C—Caradoc; A—Ashgill; L—Llandovery; W—Wenlock; L—Ludlow. Adapted from Sepkoski (1995).
1996) documented a regional brachiopod extinction in the lower Chatfieldian strata of the Appalachian Basin (Fig. 2), and Emerson et al. (2001) described a similar decline in brachiopod faunal diversity in rocks of similar age from the Upper Mississippi Valley. In Baltoscandia, Kaljo et al. (1995, 1996) and Ainsaar et al. (1999) recognized a precipitous decline in organic-walled microfossils across the Keila-Oandu regional stage boundary and referred to this diversity crash as the Oandu Crisis (Fig. 3). All of these regional events occur in a similar stratigraphic position relative to a suite of K-bentonite beds (Bergström et al., 2004; Huff et al., 1992), the prominent Guttenberg carbon isotope excursion (GICE; Fig. 4 herein; Young et al., 2005), a widely correlated stratigraphic sequence boundary (M3-M4 sequence boundary of Patzkowsky and Holland, 1996), and a graptolite zonal boundary (Goldman, 2004). Indeed, the similar timing of these local events and their proximity to a δ13C excursion suggests that they may be part of a global bio-event. All of the previously mentioned studies examined the diversity patterns of benthic fossil organisms from carbonate platform successions. In this study we examine Upper Ordovician diversity data from a group of planktonic fossil organisms generally found in offshore, deeper water deposits—graptolites. Graptolites are an ideal group for Ordovician biodiversity studies because they are widely distributed around the globe, are well represented in numerous relatively continuous black shale sections, and have been intensively sampled for biostratigraphy (Cooper et al., 2004). In addition to gathering data from a group of fossils that had a different lifestyle (planktonic) and are found in a different biofacies, this study examines graptolite faunal turnover across regions, in most of the key Upper Ordovician successions around the world.
Cooper et al. (2004) compared graptolite diversity changes through the entire Ordovician in three regions (Australasia, Baltica, and Avalonia) that represented different paleolatitudes. The present study focuses on a much shorter stratigraphic interval, the upper Sandbian to Katian Stages (lower Upper Ordovician), but more comprehensively surveys graptolite diversity across the globe (eight regional successions) in that interval. The localities examined in this study are (1) the Appalachian Basin (Laurentia); (2) Ouachita Mountains, southeastern Oklahoma (Laurentia); (3) Southern Uplands, Scotland (Laurentia); (4) Newfoundland (Laurentia); (5) Trail Creek region, Idaho (Laurentia); (6) Victoria, Australia (northeast Gondwana); (7) Scania, Sweden (Baltica); and (8) South Wales (Avalonia) (Fig. 4). Breaking down the global pattern of lower Upper Ordovician graptolite biodiversity into regional patterns that can be compared and contrasted may help answer the question of whether the upper Caradoc biodiversity decline represents a global bio-event or a composite of different regional patterns that have regional explanations. Finally, this study examines the possible links between graptolite faunal turnover and paleoenvironmental changes during the late Sandbian to early Katian. Several (and not mutually exclusive) explanations for the individual regional declines in lower Katian benthic diversity have been proposed. These include eustatic sea-level change (Patzkowsky and Holland, 1996), extensive volcanic ash deposition (Sloan, 1997), ocean temperature changes (Jaanusson, 1973; Patzkowsky et al., 1997), and changes in paleoceanographic circulation (Zalasiewicz et al., 1995). The pattern of graptolite faunal change provides additional data that can be used to evaluate these competing hypotheses, and also fuels some speculation on the relationship of biodiversity to preHirnantian Late Ordovician global climate. METHODOLOGY Measuring biodiversity through geologic time and across different geographic regions presents a number of difficulties that need to be taken into consideration. Some of the problems stem from sampling biases and a lack of taxonomic consistency in data sets compiled by different workers, whereas others result from the process of converting stratigraphic range data derived from biostratigraphic studies into diversity measures. For many of the successions in this study we have personally examined the graptolite collections, thus reducing the problem of taxonomic inconsistency. In other cases, however, species lists were compiled from the literature with taxonomic updates and revisions where possible. Sampling differences among regions can be the most difficult bias to overcome. No attempt was made to standardize the sampling. However, as we demonstrate below, because the most densely sampled region in this study (in terms of both sections and collections), the northern Appalachian Basin of the eastern United States, also showed the greatest diversity decline across the upper Sandbian to lower Katian (Chatfieldian), we conclude that diversity decline in less intensively sampled sections is not the result of under-sampling.
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Figure 2. Pattern of lower Upper Ordovician brachiopod extinctions in eastern North America. Adapted from Patzkowsky and Holland (1996). M— Millbrig K-bentonite; D—Deicke Kbentonite.
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30 60 90 120 150 30 40 50 60 70 80% 5 10 15 20 25 30 40 50 60 70% % Change in Faunal % Change in Faunal # of Taxa per Stage Composition at Stage # of Taxa per Stage Composition at Stage Boundary Boundary
Figure 3. Lower Upper Ordovician diversity changes in organic-walled microfossils from the Rapla core in Estonia. Note the close correlation of diversity changes to the brachiopod extinction in eastern North America (Fig. 2). Approximate position of Guttenberg carbon isotope excursion (GICE) on δ13C curves indicated by horizontal dashed lines. Adapted from Kaljo et al. (1995).
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Figure 4. Locality map for regional graptolite diversity studies and lower Upper Ordovician carbon isotopic (δ13C) excursions. Black arrows indicate Late Ordovician oceanic circulation patterns and are adapted from Barnes (2004). Graptolite localities are (1) Appalachian Basin (Laurentia); (2) Ouachita Mountains, southeastern Oklahoma (Laurentia); (3) Southern Uplands, Scotland (Laurentia); (4) Newfoundland (Laurentia); (5) Trail Creek region, Idaho (Laurentia); (6) Victoria, Australia (northeast Gondwana); (7) Scania, Sweden (Baltica); and (8) South Wales (Avalonia). Base map produced from the paleogeographic map software ESH-GIS 1.0, The Paleomap Project (www.scotese.com). Carbon isotopic (δ13C) excursion data adapted from Bergström et al. (2004). K-b. indicates K-bentonites.
Converting taxon range data into diversity measures presents another type of problem. Cooper (2004) noted that when estimating the mean standing diversity (MSD) of species over a specified interval of time, there are several different ways in which taxa can be counted and biases introduced. The simplest method, total
diversity, is to just tally the total number of taxa within a time interval. Alternatively, one can divide the total diversity by the interval duration to produce species per time unit measures. Foote (2000) noted that counting taxa that do not cross the study interval boundaries but are restricted to the interval itself (“singletons”)
Early Late Ordovician graptolite diversity patterns
GRAPTOLITE DIVERSITY PATTERNS Appalachian Basin Upper Ordovician black shale crops out along the extent of the Appalachian Mountains from Newfoundland to Alabama in series of northeast- to southwest-trending foreland basins that developed along the eastern margin of Laurentia during the Taconic orogeny (Finney et al., 1996; Hatcher et al., 1990). We have divided the Appalachian Basin into two regions on the basis of geography and faunal differences. These are the Appalachian Basin outcrops in the United States and Quebec (Fig. 6A), and those in Newfoundland that are treated separately (Fig. 6D). Graptolite diversity data for the United States and Quebec comes from the classic Utica Shale and Mount
Australasian Stages
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produces a number of undesirable effects and should be avoided. Unfortunately, adopting this recommendation would mean eliminating a substantial portion of the available data. Sepkoski (1975) proposed a diversity measure (called normalized diversity by Cooper, 2004) that assigns a full score to a species whose range crosses both the lower and upper boundaries of the study interval, a half score to species whose ranges cross only the upper or lower boundary of the study interval, and a half score to “singletons.” Cooper (2004) used model data sets to compare three diversity measures—total diversity, species per time unit, and normalized diversity—with actual MSD. Generally, Cooper (2004) found that total diversity measures tend to overestimate MSD, species per time unit measures tend to underestimate MSD, and normalized diversity scores best estimate MSD. Hence, we follow Cooper (2004) in using normalized diversity to calculate graptolite species diversity in lower Upper Ordovician strata. Last, we must be sure that we are actually comparing diversity scores in coeval intervals at each locality. Thus, the study intervals must be correlated precisely enough around the globe to be sure that the diversity measures are calculated from coeval stratigraphic intervals. The diversity decline in benthic faunas in North America and Baltoscandia occurs just above the Millbrig and Kinnekulle K-bentonite beds, respectively (Bergström et al., 2004), and nearly coincident with the prominent GICE carbon isotope (δ13C) excursion (Young et al., 2005). This stratigraphic level also closely approximates the Climacograptus bicornis– Diplacanthograptus caudatus graptolite zonal boundary (Bergström et al., 2004) or the boundary between the Sandbian and Katian Stages (Fig. 5). Thus, whenever possible, graptolite diversity is calculated in four subequal intervals that symmetrically span the C. bicornis–D. caudatus graptolite zonal boundary (two below and two above). These intervals precisely correlate with the Gisbornian (N. gracilis and O. calcaratus Zones) and lower Eastonian (D. lanceolatus and D. spiniferus Zones) Stages in the Australasian succession (Fig. 5). The key horizon in each section is the base of the Diplacanthograptus caudatus graptolite Biozone, a level that can be confidently correlated around the world (Goldman, 2004).
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Gi 2
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bicornis
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gracilis
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Figure 5. Lower Upper Ordovician chronostratigraphic chart, showing global and Australasian stages, eastern United States and Australasian graptolite zones, and interpolated radiometric dates. The interval of our diversity study is shaded. Adapted from Sadler and Cooper (2004) and Bergström et al. (2006).
Merino Shale outcrops in the Mohawk and Hudson River Valleys of New York State and the Saint Lawrence Lowlands of Quebec (Berry, 1962; Hall, 1847, 1865; Goldman et al., 1994; Riva, 1969, 1972, 1974; Ruedemann, 1908, 1912, 1947), and from the Athens Shale of the southern Appalachian Mountains (Decker, 1952; Finney et al., 1996; Grubb and Finney, 1995). Graptolite diversity is constant in the Nemagraptus gracilis and Climacograptus bicornis Zones (normalized diversity scores of 20) and then declines precipitously in the overlying Diplacanthograptus caudatus Zone (normalized diversity score of 7). In addition to the steep decline in taxonomic diversity, morphologic diversity declines as well. Reclined (dicellograptid) and partially reclined (dicranograptid) taxa, as well as species with cladia and scopulae, all disappear or are greatly reduced in number at the base of the D. caudatus Zone. Diversity declines further in the Orthograptus ruedemanni Zone before rebounding in the Diplacanthograptus spiniferus Zone (Fig. 6A). Graptolite diversity never attains lower Upper Ordovician levels higher in the Appalachian Basin succession; as the Taconic orogeny waned, black shales bearing graptolites were replaced by coarse clastic deposits derived from the eroding Taconic highlands (Ruedemann, 1925; Fisher, 1977). Because no single section spans the Climacograptus bicornis–Diplacanthograptus caudatus zonal boundary in the Applachian Basin, the possibility exists that the dramatic drop in graptolite diversity is in part an artifact of an unconformity at that level. However, a complete section that is conformable through this interval does exist in the Ouachita Mountains of southeastern Oklahoma.
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Figure 6. Early Late Ordovician diversity histories from eight key graptolite successions (see Fig. 4). Histograms illustrate normalized diversity scores (see text for explanation of normalized diversity), and the superimposed line graph shows the percentage of carryover taxa from one zone to another. The dashed vertical line in each histogram represents the boundary between the Sandbian and Katian Stages. Data such as regional species lists, scores for total diversity, normalized diversity, and holdover taxa are available from the authors upon request.
Early Late Ordovician graptolite diversity patterns Ouachita Mountains of Oklahoma In southeastern Oklahoma, Upper Ordovician strata are exposed along Black Knob Ridge, a low, narrow ridge at the extreme western end of the Ouachita Mountains (Hendricks et al., 1937; Finney, 1988). These strata are composed primarily of graptolite-rich shales associated with deep-water limestones and cherts (Ethington et al., 1989) and were deposited in the deep marine environment of the Ouachita Geosyncline off the southern margin of Laurentia (Finney, 1988). The units exposed along Black Knob Ridge are, in ascending order, the Womble Shale, Bigfork Chert, and Polk Creek Shale. Graptolites from the Womble Shale and Bigfork Chert at Black Knob Ridge have been comprehensively described by Finney (1986) and more recently by Goldman et al. (2007). Only graptolites from the C. bicornis and overlying D. caudatus zones were counted at this section. At Black Knob Ridge the Womble Shale contains a diverse assemblage of C. bicornis Zone graptolites (normalized diversity score of 20) that is nearly identical in its faunal composition to that of the Appalachian Basin. The boundary between the C. bicornis and overlying D. caudatus Zones occurs 4 m above the base of the Bigfork Chert, which conformably overlies the Womble Shale. With a pattern similar to that of the Appalachian Basin, normalized diversity in the D. caudatus Zone at Black Knob Ridge drops by nearly 50% to 10.5 (Fig. 6B). Southern Uplands, Scotland Another important low paleolatitude graptolite locality in the environs of Laurentia is Hartfell Score in the Southern Uplands of Scotland. It was in this region that Lapworth’s (1876, 1878) pioneering work demonstrated the stratigraphic utility of graptolites by using them to help work out the geologic structure of the uplands. In the Southern Uplands, Upper Ordovician to Silurian mudstones and graywackes are exposed in a series of structurally repeated, thrust-bounded slices that dip steeply to the northwest (Leggett et al., 1979; Rushton et al., 1996). At Hartfell Score near Moffatt, ~20 m of graptolite-rich black mudstones conformably span the Climacograptus wilsoni (= C. bicornis)–Diplacanthograptus caudatus zonal boundary (Williams, 1982; Zalasiewicz et al., 1995). The graptolite biostratigraphy was recently updated by Zalasiewicz et al. (1995), and we have used their detailed range chart for our diversity analysis (Fig. 6C). The normalized diversity score for the C. wilsoni Zone at Hartfell Score is 22.5, and the score for the overlying D. caudatus Zone is 21.5. In stark contrast to the Appalachian Basin, graptolite diversity in southern Scotland remains nearly constant across the C. bicornis–D. caudatus zonal boundary. Also noteworthy is the fact that dicellograptinids and dicranograptinids remain an important part of the D. caudatus Zone fauna in southern Scotland, whereas they disappear almost entirely from the Appalachian Basin in the strata above the C. bicornis Zone. Although we only counted species from the single section at
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Hartfell Score, the data indicate a much more diverse D. caudatus Zone than in the Appalachian Basin or Ouachita Mountains of Oklahoma. Newfoundland Late Ordovician graptolites from the Lawrence Harbour Formation (Exploits Zone) in central Newfoundland were described and figured by Erdtmann (1976) and Williams (1995). We have used the species list and range chart provided by Williams (1995), because it covers a greater number of collection localities and uses updated graptolite taxonomy. The Lawrence Harbour Formation outcrops are structurally complex, and no single exposure contains a complete succession of zones; hence Williams’ (1995) range chart does not illustrate the stratigraphic ranges of taxa, only the zones in which they occur. Normalized diversity scores for graptolites from the Lawrence Harbour Formation are N. gracilis Zone, 9; C. bicornis Zone, 12; D. clingani Zone, 11; P. lineariz Zone, 8 (Fig. 6D). As in Scotland, graptolites in central Newfoundland exhibit no notable decrease in species diversity across the C. bicornis–D. caudatus zonal boundary. Trail Creek, Idaho The Trail Creek Summit section within the Phi Kappa Formation of central Idaho, northwestern United States, exposes an ~200-m-thick succession of black siliceous shale and argillite that has yielded biostratigraphically important graptolites for nearly a century. Surveys of the local biostratigraphy in this area, conducted by Churkin (1963), Carter (1972), and Carter and Churkin (1977), established a set of zones that have been employed as a standard reference for the North American Cordillera. More recent studies by Mitchell et al. (2003), Maletz et al. (2005), and Motz et al. (2006) have updated the graptolite taxonomy and substantially revised the biostratigraphic zonation. The new data indicate that the Trail Creek succession is similar to the Australasian and South China (Chen et al., 2005) successions (low paleolatitude, tropical regions) and is also more incomplete than previously thought. At Trail Creek the normalized diversity score for the C. bicornis Zone is 14 (Fig. 6E). Mitchell et al. (2003) noted that a succeeding D. caudatus Zone fauna could not be unambiguously identified at Trail Creek. The strata that overlie the C. bicornis Zone were called passage beds by Carter and Churkin (1977), who were unable to precisely define their biostratigraphic age. Although D. caudatus does occur within the “passage beds,” its first appearance is above that of D. spiniferus (which was misidentified as C. bicornis by Carter and Churkin), the index for the overlying D. spiniferus Zone. Additionally, J.F.V. Riva (2005, personal commun.) claims to have found evidence of a prominent unconformity just below the first appearance of D. spiniferus at the summit section. The normalized diversity score for the zone succeeding the C. bicornis Zone is 6, indicating a large drop in diversity, but because we cannot differentiate between a
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D. caudatus and a D. spiniferus Zone at Trail Creek the data must be viewed with caution. Victoria, Australia The Australasian region of Gondwana lay in low latitudes during the Ordovician (Scotese and McKerrow, 1990), and its graptolite succession is one of the most complete and finely divided in the world. It is considered to be the standard for the Pacific Faunal Province (VandenBerg and Cooper, 1992). Graptolite biostratigraphy and biodiversity in the Victorian succession were comprehensively reviewed by VandenBerg and Cooper (1992) and Cooper et al. (2004), and we relied on their range charts for our data. These data, which are similar to Williams’ (1995) Newfoundland data, do not illustrate the exact stratigraphic range of each species from measured sections but only the biozones in which they occur. Graptolite diversity increases from the Gisbornian 1 (Nemagraptus gracilis Zone, normalized diversity of 12) to a nearly constant level in the three succeeding zones—Gisbornian 2 (nearly equivalent to the C. bicornis Zone), normalized diversity of 15; Eastonian 1 (nearly equivalent to the D. caudatus Zone), normalized diversity of 14.5; and Eastonian 2 (D. spiniferus Zone), normalized diversity of 15 (Fig. 6F). A detailed survey of low paleolatitude graptolite localities reveals that the precipitous declines observed in the Appalachian Basin and the Ouachita Mountains of Oklahoma are clearly not pervasive phenomena but are restricted to certain parts of Laurentia. One important low paleolatitude locality that we have not analyzed is the Marathon region of West Texas (Berry, 1960). Riva (1969), Bergström (1978), and Goldman et al. (1995) all noted that a prominent unconformity corresponds to much of the Katian Stage (upper Caradoc) in the Marathon succession. Scania, Sweden Scania (southern Sweden) belonged to the paleocontinent Baltica, which was at mid-latitudes during the early Late Ordovician (Scotese and McKerrow, 1990; Mac Niocall et al., 1997). The Ordovician rocks of Baltoscandia are subdivided into subparallel, generally SE-NW–trending confacies belts that maintained fairly constant geographic positions through time (Jaanusson, 1976, 1995). The Middle and Upper Ordovician strata in Scania (studied from both core and outcrop) are predominantly black to gray shales and mudstones representing outer shelf or foreland basin deposition (Bergström et al., 2000). Our diversity data come from the Koängen and Fågelsång cores described by Nilsson (1977) and Hede (1951), respectively, and Pålsson (2001). The senior author of this paper has personally examined the graptolites from both cores and has revised and updated the taxonomy and taxon ranges. Graptolite diversity in Scania increases from the Nemagraptus gracilis Zone to the Diplograptus foliaceous Zone (= C. bicornis Zone) and drops dramatically into the Dicranograptus clingani Zone (approximately equivalent to the D. caudatus Zone), with
normalized diversity scores of 11.5, 21, and 10, respectively (Fig. 6G). Thus, the graptolite diversity pattern in Scania is similar to that found in the Appalachian Basin and the Ouachita Mountains. South Wales (Avalonia) Another classic location in the history of graptolite studies is Wales, which, along with England, was part of Avalonia, a midto high-latitude paleocontinent during the early Late Ordovician (Mac Niocall et al., 1997). Elles and Wood (1901–1918) provided the first comprehensive descriptions of Welsh graptolites, and Elles (1939) later elucidated the stratigraphic ranges of many Llandeilo and Caradoc graptolite species. These early classic works were updated by Hughes (1989). The graptolite-bearing strata in Wales are generally older than the intervals examined in this paper, but Zalasiewicz et al. (1995) described a new section in Whitland, South Wales, that has its base in the Dicranograptus clingani Zone. Although we cannot compare diversity changes across the base of the Katian Stage because the underlying Diplograptus foliaceous Zone is absent from this section, we believe that the pattern of graptolite faunal change that is exhibited in these rocks provides some insight into understanding the global pattern of Katian (late Caradoc) diversity changes. At Whitland, nearly 60 m of laminated dark-gray graptolitic mudstones and silty mudstones of the Mydrim (or Dicranograptus) Shales is exposed (Zalasiewicz et al., 1995). In its upper part, the Mydrim Shales exhibit less pronounced lamination and become interbedded with thin muddy limestone bands before finally grading upward into the Sholeshook Limestone. Based on brachiopods and trilobites, the Sholeshook Limestone at Whitland is Ashgillian in age (Zalasiewicz et al., 1995). Between 12.5 and 39 m above the base of the section, the Mydrim Shales contain a low diversity Dicranograptus clingani Zone fauna. This fauna contains elements (Diplacanthograptus spiniferus, Neurograptus margaritatus, Dicranograptus nicholsoni, Dicellograptus flexuosus, and Orthograptus quadrimucronatus) that indicate a correlation with the D. caudatus to D. spiniferus zones in North America and Eastonian 1–Eastonian 2 in Australia. Above 39 m, faunal diversity decreases even further, and the fauna is completely dominated by several species of Normalograptus (Fig. 7). Zalasiewicz et al. (1995) note that the Mydrim to Sholeshook transition represents a shallowing upward sequence and that even in the lowermost beds most offshore or deep water taxa (deep water biotope of Cooper et al., 1991) and the cosmopolitan mesopelagic biotope of Goldman et al. (1995) are absent from the Whitland section. They also suggest that the marked diversity decline in post–D. clingani– age rocks might be related to either decreasing water depth or water temperatures, postulating that the transition from laminated graptolitic mudstones to bioturbated, nearly barren mudstones, and then to shallow shelf carbonates may record a preglacial climatic deterioration.
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Early Late Ordovician graptolite diversity patterns
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Ainsaar et al., 1999). Thus, each region has its own local diversity history, and although the sum of these regional histories likely produces the global decline recognized by Sepkoski (1995), no global pattern that might warrant a comprehensive explanation is immediately evident. Nevertheless, we do see an interesting relationship between the regional diversity histories and paleolatitude, as well as a possible connection between these patterns and the changing Late Ordovician climate.
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% Normalograptus Species Figure 7. Percentage of Normalograptus species in the Whitland, South Wales, graptolite succession. Note the increasing percentage of normalograptids with time. Adapted from data in Zalasiewicz et al. (1995).
Summary of Regional Upper Sandbian to Katian Graptolite Diversity Patterns The preceding regional review of graptolite diversity change through the upper Sandbian and Katian Stages reveals a varied pattern. In the low paleolatitudinal localities of the eastern and southern United States, graptolite diversity declines precipitously across the Climacograptus bicornis–Diplacanthograptus caudatus zonal boundary. However, other former tropical localities such as Scotland, Newfoundland, and Australia, exhibit nearly static diversity across this same interval. The mid-paleolatitude succession in Scania records a sharp diversity drop, and the mid- to high-paleolatitude Whitland section in South Wales contains a low diversity lower Katian assemblage followed by nearly complete domination by a single genus, Normalograptus. These results, however, are consistent with data from other fossil groups—regional brachiopod extinctions in the Appalachian Basin and Upper Mississippi Valley (Patzkowsky and Holland, 1996; and Emerson et al., 2001, respectively) and organic-walled microfossil declines in Baltoscandia (Kaljo et al., 1995, 1996;
Graptoloids are one of the fossil groups that were decimated by the end-Ordovician mass extinction. In several regions, particularly those at higher latitudes, graptolite diversity began declining in the Katian, well before the Hirnantian climate deterioration and glaciation. Zalasiewicz et al. (1995) proposed that because Wales lay at a higher latitude than Scotland during the Late Ordovician it may have been subject to more rapid climatic deterioration during a cooling event. Additionally, the density-stratified setting that seems to have favored graptolite proliferation and fossilization (Berry et al., 1987) may have been disrupted by an influx of cooler, oxygen-rich polar waters (Zalasiewicz et al., 1995). Cooper et al. (2004) also noted that graptolite diversity in Avalonia (and perhaps Baltica) began to decline well before it declined in Australasia, and they speculated on similar causal mechanisms (that higher latitude localities may have undergone surface water temperature changes or breakdowns in ocean density structure from an earlier onset of global cooling than at lower latitude localities). They also noted, based on the paleoplate reconstructions of Cocks and Torsvik (2004), that Avalonia and Baltica had already drifted into near-tropical latitudes by the end of the Katian, making latitude an unlikely determinant of diversity change. It appears to us, however, that on the basis of the paleocontinental reconstructions of Scotese and McKerrow (1990), Mac Niocall et al. (1997), and even Cocks and Torsvik (2004), a distinct latitudinal gradient did exist among the regions discussed here during the early Late Ordovician. We re-propose the hypothesis that a pre-Hirnantian episode of global cooling took place and that this climate change negatively affected graptolite diversity at middle and high latitudes but was not a major factor at low latitudes. However, the steep decline in diversity exhibited by several Laurentian successions requires a separate causal mechanism. Several workers (e.g., Berry et al., 1987; Cooper, 1998) proposed that graptolites thrived during times of sluggish oceanic circulation and that their preferred habitat was in a nutrient-rich layer above an oxygen minimum zone (OMZ) that produced the anoxic conditions necessary for graptolite preservation—generally deep shelf to slope settings. During the end-Ordovician regression and glaciation, the formation of cold, oxygen-rich polar waters and their sinking and spreading toward the equator would have intensified, increasing the ventilation of the ocean and disrupting the density-stratified water that
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graptolites favored (Berry et al., 1987). Normalograptus, the only graptolite genus to survive the Ordovician mass extinction and diversify in the Silurian was an epipelagic species and occupied one of the shallowest, most oxygen-rich zones of the upwelling region. Finney and Berry (2001) and Chen et al. (2005) noted that the normalograptids therefore were the species least likely to be affected by deep ocean ventilation and disruption of a density stratified water column. As noted above, the lower Katian (D. clingani Zone) in South Wales has a low-diversity graptolite fauna. Near the top of the D. clingani Zone the graptolite fauna becomes almost entirely dominated by the genus Normalograptus. Thus, both the sedimentological changes that occur up section at Whitland—transition from laminated, graptolitic mudstones to bioturbated, barren mudstones and finally shelf carbonates—and the faunal changes are consistent with the end-Ordovician extinction model, even though they occurred during the Katian. Is there any independent evidence that an episode of global cooling occurred prior to the Hirnantian? Patzkowsky et al. (1997) suggested that positive excursions in carbon isotope compositions in both carbonate and organic carbon from lower Katian strata in the eastern and midcontinental United States—the prominent GICE carbon isotope (δ13C) excursion (Fig. 4)—indicate increased productivity and rates of organic carbon burial that may have drawn down atmospheric pCO2 and precipitated global cooling. Pope and Steffen (2003) noted that Upper Ordovician carbonates from the southern and western margins of Laurentia contained abundant chert and phosphate, which they interpreted as evidence of widespread upwelling and vigorous thermohaline circulation related to early Gondwanan glaciation. They also point out that this period of upwelling corresponded with cool surface waters in the Appalachian Basin (Railsback et al., 1990), a northward expansion of cool-water trilobite faunas in North America (Shaw, 1991), and a transition to cooler water benthic faunas across eastern North America (Patzkowsky and Holland, 1993). Finally, several workers claimed that Katian (middle to upper Caradoc) siliciclastic deposits in northwestern Gondwana display clear evidence of glacial transport and deposition (e.g., Barnes, 1986; Theron, 1994; Hamoumi, 1999). It is also worth noting, however, that Brenchley et al. (1994) and Marshall et al. (1997) consider there to be little substantive isotopic evidence for a long-lived glacial episode that substantially preceded the Hirnantian. If climatic cooling occurred during the early Katian it clearly did not affect graptolite faunal diversity in all the low latitude tropical regions. Southern and eastern Laurentian graptolite faunas were severely depleted, but faunas from northeast Laurentia (Scotland and Newfoundland) and northeast Gondwana (Australia) do not exhibit a pronounced diversity decline. Interestingly, the percentage of taxa that carry over from the upper Sandbian into the lower Katian is similarly low in all areas, suggesting that where a diversity drop is observed, it is the result of reduced origination rate and not elevated extinction rate. Zalasiewicz et al. (1995) and Finney (1986, 1988) noted that prominent regressions, as evidenced by distinct facies changes,
occurred in South Wales and southwestern Laurentia during the early to mid-Katian. Perhaps the different regional diversity patterns could be attributed to unique water-depth histories of individual basins. Although we agree that sea level was an important factor that affected the graptolite biotope, as a single causal mechanism, sea level alone could not be responsible for Katian graptolite diversity decline in Laurentia because the Appalachian Basin underwent a relative sea-level rise at that time (Holland and Patzkowsky, 1996). Indeed, the nadir of graptolite diversity in the northern Appalachian Basin occurred within the Orthograptus ruedemanni Zone, an interval considered by Joy et al. (2000) to represent a transgressive system. Patzkowsky and Holland (1996) and Pope and Steffen (2003) provided evidence for the incursion of cool oceanic waters onto the Laurentian continent. Kolata et al. (2001) suggested that cool, subpolar Iapetus seawater may have flooded the Laurentian craton through a narrow depressed corridor, the Sebree Trough. This trough is filled with a succession of dark brown to gray shales that contain a very low diversity graptolite fauna of Katian age (Kolata et al., 2001; Mitchell and Bergström, 1991). These shales are similar in age and faunal composition to the classic Utica Shale of the northern Appalachian Basin (Mitchell and Bergström, 1991). We think that if cool, well-oxygenated subpolar waters flooded the Laurentian craton, it may have adversely affected graptolite diversity. Thus, physical and chemical seawater changes, albeit driven by different mechanisms, may have produced similar diversity patterns at different paleolatitudes. Finally, Finney and Berry (1997) proposed that graptolites thrived in a denitrification layer above an OMZ that developed where vigorous upwelling occurred along continental margins. They attribute changes in graptolite diversity to fluctuations in upwelling conditions and that the OMZ loss of upwelling conditions resulted in destruction of the preferred graptolite habitat and declines in both abundance and diversity. The upwelling model was proposed to explain the relationship between vertical changes in facies patterns and graptolite diversity and abundance in the Vinini Formation of north-central Nevada (Finney and Berry, 1997). This model is seemingly at odds with the interpretation of widespread Late Ordovician upwelling in North America (Pope and Steffen, 2003) and coincident graptolite diversity decline. We agree that upwelling produces the nutrient-rich conditions necessary for high productivity, but it is not clear that plankton abundance and species diversity respond similarly to circulation changes. The phosphate-rich, dark shale of the Late Ordovician Maquoketa Group crops out extensively in the Upper Mississippi Valley of the United States (Templeton and Willman, 1963). Witzke (1987) noted that in its southeastern outcrop area the lower Maquoketa was deposited in an epicontinental seaway with well developed density and oxygen stratification. Interestingly, the lower Maquoketa Formation is rich in graptolites but exhibits very low species diversity (Goldman and Bergström, 1997). We think that more research on the relationship between plankton abundance, species diversity, and upwelling is still needed.
Early Late Ordovician graptolite diversity patterns CONCLUSIONS The early Late Ordovician (Katian) was a time of significant decline in benthic marine biodiversity. In this study we compiled Upper Ordovician diversity data for graptolites, a group of fossil organisms that had a different lifestyle (planktonic) and are generally found in a different biofacies. We investigated the pattern of graptolite faunal changes across the Climacograptus bicornis– Diplacanthograptus caudatus graptolite zonal boundary in North America and on several other continents. In the Appalachian Basin a sharp decline in species diversity occurred. Scores for normalized diversity dropped from 20 in the C. bicornis Zone to 7 in the D. caudatus Zone. Only 11% of the species present in the C. bicornis Zone carry over into the D. caudatus Zone. A similar pattern occurs in the Ouachita Mountains of Oklahoma. High and middle paleolatitude regions such as Wales and Baltoscandia exhibit low diversity in the lower Katian and then further decline in graptolite diversity at higher stratigraphic levels. In other low paleolatitude regions such as Australasia and Scotland, however, diversity is fairly constant across this interval (although the percentage of carryover taxa is low). Climatic cooling during the Katian may have directly affected graptolite faunas at mid- to high latitudes but not at low latitudes. Parts of Laurentia were affected only because a depressed corridor, the Sebree Trough, allowed deep, cool Iapetus Ocean water to spread across the craton. Thus, different combinations of local tectonics, climate change, and water-depth history may have produced similar diversity patterns at different paleolatitudes. ACKNOWLEDGMENTS Research for this study was supported by U.S. National Science Foundation grant EAR-0106844 to D. Goldman. This paper was improved by constructive and thoughtful reviews from Roger Cooper and Jan Zalasiewicz. Discussions with Jörg Maletz and editorial comments by Stan Finney were also helpful and appreciated. REFERENCES CITED Ainsaar, L., Meidla, T., and Martma, T., 1999, Evidence for a widespread carbon isotopic event associated with late Middle Ordovician sedimentological and faunal changes in Estonia: Geological Magazine, v. 136, p. 49–62, doi: 10.1017/S001675689900223X. Barnes, C.R., 1986, The faunal extinction event near the Ordovician-Silurian boundary: A climatically induced crisis, in Wallister, O.H., ed., Global Bioevents: Lecture Notes in Earth Sciences, v. 8, p. 121–126. Barnes, C.R., 2004, Ordovician oceans and climate, in Webby, B.D., Paris, F., Droser, M.L., and Percival, I.G., eds., The Great Ordovician Biodiversity Event: New York, Columbia University Press, p. 72–76. Bergström, S.M., 1978, Middle and Upper Ordovician conodont and graptolite biostratigraphy of the Marathon, Texas graptolite zone reference standard: Palaeontology, v. 21, p. 723–758. Bergström, S.M., and Finney, S.C., Chen Xu, Pålsson, C., Wang Zhi-hao, and Grahn, Yngve, 2000, A proposed global boundary stratotype for the base of the Upper Series of the Ordovician System: The Fågelsång section, Scania, southern Sweden: Episodes, v. 23, p. 102–109. Bergström, S.M., Huff, W.D., Saltzman, M.R., Kolata, D.R., and Leslie, S.A., 2004, The greatest volcanic ash falls in the Phanerozoic: Trans-Atlantic
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MANUSCRIPT ACCEPTED BY THE SOCIETY 6 OCTOBER 2009
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The Geological Society of America Special Paper 466 2010
Origin of Late Ordovician (mid-Mohawkian) temperate-water conditions on southeastern Laurentia: Glacial or tectonic? Frank R. Ettensohn† Department of Earth and Environmental Sciences, University of Kentucky, Lexington, Kentucky 40506, USA
ABSTRACT In mid-Mohawkian time at the Turinian–Chatfieldian (Blackriverian–Trenton) transition, the extensive, shallow-, warm-water Blackriverian carbonate platform, in a low-latitude, subtropical setting across east-central Laurentia, underwent an abrupt change to temperate-water, sedimentary and faunal regimes. The change occurs across a regional unconformity, the formation of which is coincident with inception of a major Taconian tectophase and the related breakup of the Blackriverian Platform along old zones of structural weakness into smaller and higher platforms and shelves separated by a linear low area called the Sebree Trough. This breakup is interpreted to reflect far-field, foreland deformation, and temperate-water carbonates soon predominated across the resulting shelves and platforms, some 1000 km from the nearest open continental margin. Yet, paleogeographic models, the presence of warm-water carbonates distal to the trough as well as warm-water faunal and lithologic inliers in protected back-shoal settings on the Lexington Platform still indicate the presence of warm surface waters in a subtropical setting. Most interpretations support the necessity of upwelling from the Sebree Trough to explain this seeming anomaly, reflecting the significance of coeval tectonism in generating necessary conditions for the upwelling. Others have called upon glaciation for generating cooler oceanic waters and the enhanced oceanic circulation that would have developed coastal upwelling, sufficiently intense and widespread to have penetrated far across the open continental margin. Although latest Ordovician Hirnantian glaciation is well supported, evidence for the timing, extent, and likelihood of earlier Late Ordovician (Chatfieldian) glaciation is uncertain and contradictory. Although it is not possible to preclude the effects of such glaciation, the synergetic effects of tectonics and paleogeography may offer a more plausible explanation for upwelling during the time of syn-Taconic, far-field, foreland deformation on southeastern Laurentia; these effects may have even accentuated any glacial influence. Inasmuch as the Late Ordovician was a time of global tectonism, there may have been other low-latitude Ordovician cratonic settings in which changes in nearby orogens offer explanations for abrupt changes in sedimentary and faunal regimes on the adjacent forelands.
†
E-mail:
[email protected].
Ettensohn, F.R., 2010, Origin of Late Ordovician (mid-Mohawkian) temperate-water conditions on southeastern Laurentia: Glacial or tectonic?, in Finney, S.C., and Berry, W.B.N., eds., The Ordovician Earth System: Geological Society of America Special Paper 466, p. 163–175, doi: 10.1130/2010.2466(11). For permission to copy, contact
[email protected]. ©2010 The Geological Society of America. All rights reserved.
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INTRODUCTION Lower to lower Upper Ordovician (Ibexian–Mohawkian) carbonates are common across large parts of Laurentia (e.g., Cook and Bally, 1975) and seem to reflect deposition in the southern, subtropical, high-pressure belt (evaporative trade-wind belt; e.g., Scotese, 2003), where warm tropical currents bathed continental margins (Wilde, 1991) during a period of high atmospheric pCO2 (Berner, 1994; Crowley and Baum, 1995; Berner and Kothavala, 2001). The rocks here are generally characterized by faunal and lithologic associations that are analogous to the recent, warm-water Chlorozoan association of Lees (1975) or Photozoan assemblage of James (1997), although exact Ordovician analogues for all recent components are not present. Judging from the continuation of paleoclimatic and atmospheric conditions into Late Ordovician time, warm-water carbonate deposition should have occurred throughout most of Late Ordovician time. However, in mid-Mohawkian time (Turinian–Chatfieldian or Blackriverian–Trenton transition), and perhaps earlier, across large parts of Laurentia warm-water associations abruptly gave way to deposits interpreted to be temperate-water carbonates (e.g., Patzkowsky and Holland, 1993, 1996; Holland and Patzkowsky, 1996, 1997; Pope and Read, 1997a; Kolata et al., 2001; Ettensohn et al., 2002; Pope and Steffen, 2003; Pope, 2004). The cause of this abrupt change is uncertain, but on the Lexington Platform in southeastern Laurentia the change occurred across a regional unconformity that marks the initiation of a major Taconian tectophase, and Ettensohn et al. (2002) suggested that the change, at least in part, must have reflected the tectonic restructuring of the Laurentian craton that occurred at that time. Others, however, have suggested development of enhanced thermohaline circulation and upwelling related to the position of preexisting bathymetric lows and the paleogeographic position of the southern Laurentian margin (Witzke, 1980, 1987; Kolata et al., 2001) or to the advent of early Late Ordovician glaciation (Pope and Read, 1997a; Pope and Steffen, 2003; Saltzman and Young, 2005) and its continuation into latest Ordovician (Hirnantian) time. Those models that support the importance of preexisting bathymetric lows and paleogeography in explaining the changed faunal and lithologic associations at the Turinian–Chatfieldian or Blackriverian–Trenton transition are similar to what will be proposed herein, except that they downplay the underlying significance of Laurentian tectonics. This paper, however, is a review of evidence from the Lexington Platform that is germane in understanding the relative importance of tectonic versus glacial hypotheses. THE EARLY LATE ORDOVICIAN LEXINGTON PLATFORM During much of early Late Ordovician (Chazyan– Blackriverian) time, southern parts of Laurentia were largely characterized by extensive peritidal to very shallow, open-marine carbonate environments (Ross, 1976; Keith, 1989) in a setting
that has become informally known as the Blackriverian carbonate platform. By late Turinian (Blackriverian) time, the “Blountian” tectophase of the Taconian orogeny (Kay and Colbert, 1965) had begun along the extreme southeastern margin of Laurentia (e.g., Ettensohn, 1991), but it does not seem to have influenced platform development much beyond the narrow foreland basin. Aside from the foreland basin, this preexisting, largely peritidal platform extended across southeastern and central Laurentia (Keith, 1989; Scotese, 2003), and based on sedimentary features (Lindström, 1984; Scotese, 2003), paleogeographic position (Scotese, 2003), and an abundance of carbonate mud with sparse, Chlorozoan-like fauna and flora (e.g., Lees, 1975), it is interpreted to have been a site of subtropical warm-water carbonate deposition. Abruptly, however, at the Turinian–Chatfieldian (Blackriverian–Trenton) transition, the Blackriverian carbonate platform underwent stratigraphic and structural differentiation that resulted in breakup of the platform along old basement zones of structural weakness (Ettensohn et al., 2002) and in development of a temperate-water (Pope and Read, 1997a), mixed carbonateclastic facies mosaic (Keith, 1989) (Fig. 1). This transition is also marked by the regional Black River–Trenton unconformity that has been interpreted to represent inception of the “Taconic” tectophase of the Taconian orogeny near the New York promontory on southeast-central parts of the Laurentian margin (Ettensohn, 1991, 1994; Ettensohn et al., 2002). Breakup of the Blackriverian Platform generated new carbonate shelves to the north, the rectangular Lexington Platform and the intervening Sebree Trough, a structurally lowered, deepwater trough that connected the southern continental margin of Laurentia with the Taconic foreland basin (Fig. 1). The rectilinear nature of the Lexington Platform, moreover, reflects the basement zones of weakness along which the breakup occurred (Ettensohn et al., 2002). Inception of the breakup co-occurred with development of the Black River–Trenton unconformity and a bentonite complex (Kolata et al., 1996, 1998), which suggest links to the initial eastward subduction of the Laurentian margin in the Taconic tectophase and the resulting tensional, far-field forces in the foreland (Ettensohn et al., 2002, 2004). Dating of the bentonite at or near the unconformity (Mud Cave–Millbrig K-bentonite) by Tucker and McKerrow (1995) indicates that these apparently coeval events all transpired at ca. 454 Ma, although more recent dating of the bentonites may indicate an age as young as 448 Ma (Min et al., 2001). Also at this same point in the stratigraphic record, Holland and Patzkowsky (1997) and Pope and Read (1997a) interpreted a correspondingly abrupt change from warm-water Blackriverian or Turinian carbonates to temperate-water Chatfieldian carbonates. Among the most important lines of evidence noted by Holland and Patzkowsky (1997) and Pope and Read (1997a) in support of temperate-water conditions on the new Lexington Platform are a nearly omnipresent brachiopod-bryozoan association (similar to the Bryomol association); abundance of phosphorite; abundance of marine-cemented, mineralized hardgrounds; and absence of warm-water features, such as Chlorozoan fauna,
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Figure 1. Schematic map of east-central United States showing sedimentary and tectonic differentiation by early Edenian time of the former warm-water Blackriverian carbonate platform into the temperate-water Galena and Trenton Shelves, Sebree Trough, and Lexington Platform. Breakup and differentiation of the Blackriverian platform occurred along a structural zone of weakness at and following the Turinian–Chatfieldian (Blackriverian–Trenton) transition. The Tanglewood buildup (T) and Louisville high (L) were areas of local uplift in mid- to late Chatfieldian time that generated thicker areas of shallow-water, upper Lexington carbonates. PP—Point Pleasant basin; PE—Pennsylvania embayment; V—Virginia promontory; N— New York Promontory (from Ettensohn et al., 2002).
algal laminites, oolites, and evaporites. The Lexington Limestone itself is also characterized by an abundance of trilobite fragments and locally abundant red algae, both of which are also indicators of cool-water conditions (Lindström, 1984; James, 1997). Coeval phosphate-rich, temperate-water carbonates are also reported from the related Galena and Trenton Shelves and other parts of the Lexington Platform (Wilson, 1962; Brookfield, 1988; Lavoie, 1995; Holland and Patzkowsky, 1996; Kolata et al., 2001). Interestingly, coeval warm-water facies are reported in the Upper Mississippi Valley ~400 km northwest of the Lexington Platform (Delgado, 1983; Kolata et al., 2001) and in later Ordovician facies as far east as the Maritime provinces of Canada (Holland and Patzkowsky, 1997). Moreover, in the midst of all the temperate-water Lexington Limestone facies, the localized Perryville and Devils Hollow Members (Fig. 2) exhibit warm-water carbonate characteristics similar to those encountered in underlying Blackriverian carbonates. “Warm-water” parts of these members, however, are everywhere associated with coarse calcarenites and calcirudites of the Tanglewood and Devils Hollow Members, which are interpreted to have formed protective shoals or beachbarrier complexes behind which warm-water facies could be maintained (Mackey, 1972; Ettensohn et al., 2004) (Fig. 2). Moreover, on some of these shoal complexes, zones of stromatoporoids and corals, which are considered to be warm-water fauna (James, 1997), repeatedly developed, suggesting development at the top of highstand sequences, when the shoals projected upward into very warm surface waters (Pope and Read, 1997a). ORDOVICIAN GLACIATION CO2 levels during most of Late Ordovician time, 14–16 times higher than present-day values (Berner, 1991), would seem
to have precluded glaciation of any type. However, Crowley and Baum (1995) have demonstrated that the unique geographic configuration of Gondwana and the presence of major relief on certain parts of that continent may well have triggered glaciation at this time, whereas others have called upon a very shortlived drawdown in atmospheric CO2 (e.g., Brenchley et al., 1994; Kump et al., 1995; Marshall et al., 1997; Gibbs et al., 1997). In fact, end-Ordovician (Hirnantian, Gamachian) glaciation is now widely accepted and well supported on the basis of several lines of geochemical (e.g., Brenchley et al., 1994, 2003; Shields et al., 2003), paleontological (e.g., Spjeldnaes, 1961; Brenchley, 1984, 1989; Eyles, 1993; Marshall et al., 1997; Brenchley et al., 2003), sedimentologic-stratigraphic (e.g., Hambrey, 1985; Hambrey and Harland, 1985; Frakes et al., 1992; Eyles, 1993; Crowell, 1999), and eustatic (Brenchley and Newell, 1980) information. Moreover, the detailed work of Brenchley et al. (1994, 2003) suggests that this episode of glaciation commenced at the beginning of the Hirnantian Age, or at ca. 446 Ma (Shields et al., 2003; Gradstein and Ogg, 2004). Although Brenchley et al. (1994) suggested that major Late Ordovician glaciation was no more than a million years in duration and confined to Hirnantian time, Frakes et al. (1992), Eyles (1993), and Crowell (1999) suggested that the glaciation probably began in early Late Ordovician (Mohawkian–early Cincinnatian) time and climaxed at the end of Hirnantian (Gamachian) time, and modeling may support this conclusion (Herrmann et al., 2004). Frakes et al. (1992) even questionably suggested that glaciation might have extended back to Middle Ordovician (late Whiterockian) time. There does, however, seem to be some consensus that Ordovician glaciation had begun on Gondwana by Mohawkian time, but the age of inception is poorly constrained with cited ages of ca. 458 Ma (Frakes et al., 1992), 455 Ma
Figure 2. Generalized stratigraphic column for the Lexington Limestone in central Kentucky. The “upper Lexington” or Tanglewood buildup includes parts of the unit above the sub–Sulfur Well unconformity, which are present only in central Kentucky. Parts of the unit below the unconformity are informally known as the “lower Lexington” and represent the Lexington-Trenton Limestone as it occurs across most of the Lexington Platform. In contrast to the temperatewater nature of most of the Lexington Limestone, the Perryville and Devils Hollow Members include warm-water components that apparently developed in protected settings behind Tanglewood shoal complexes.
Controls on Late Ordovician temperate-water conditions (Theron, 1994), 447 Ma (Crowell, 1999), and 440 Ma (Crowley and Baum, 1995); these ages range from near the beginning of Late Ordovician (latest Whiterockian) to earliest Silurian time (Shields et al., 2003; Gradstein and Ogg, 2004). The possible development of glaciation as early as Mohawkian time could have been important, because several recent workers have used evidence for its occurrence as means of explaining the presence of temperate-water carbonates in Ordovician subtropical settings (e.g., Lavoie, 1995; Pope and Read, 1997a, 1998; Pope and Steffen, 2003; Pope, 2004) and for explaining the prominent shifts in Mohawkian geochemical (δ13C) signatures (e.g., Ainsaar et al., 2004; Saltzman and Young, 2005). In contrast, however, Fortey and Cocks (2005) presented evidence for a Late Ordovician (Richmondian) period of global warming just prior to Hirnantian glaciation, and they point out the apparent anomaly of early cooling scenarios. In fact, Fortey and Cocks (2005, p. 407) stated relative to Laurentia, “There is no independent evidence for cooling, or for oceanic overturn and reorganization, before the end of the Ordovician extinction.” Clearly, there is still some uncertainty about the nature and extent of preHirnantian cooling and glaciation. ORDOVICIAN TECTONICS ON LAURENTIA The Middle to Late Ordovician was a time of major tectonism and high volcanism, based on patterns of sedimentation (Ronov et al., 1980; Kolata et al., 1996), geochemistry (Shields et al., 2003), and mountain building (e.g., Rodgers, 1971; Nikitin et al., 1991; Scotese and McKerrow, 1991; Scotese, 1997). In fact, if the reconstructions of Scotese and McKerrow (1991) and Scotese (1997) are correct, from Middle Ordovician to Middle Silurian time, active subduction zones surrounded most of the major continents, including Laurentia and Gondwana, so that plate reorganization that had begun on one margin may have caused nearly synchronous tectonism on other margins as well. In particular, major coeval mountain-building events, which Eyles (1993) merely labeled “Taconic,” were ongoing at the eastern margin of Laurentia (Taconian orogeny), on the margins of Gondwana, and on the northern margins of Siberia and Kazakhstania (Scotese and McKerrow, 1991). Moreover, Eyles (1993) and Crowley and Baum (1995) not only attributed inception of Ordovician glaciation to the high-latitude position of Gondwana as it crossed the South Pole, but also to high elevations on the western margin of the Afro-Arabian Platform related to tectonism and Gondwanan mountain building. On the southeast margin of Laurentia, Taconian tectonism began near the Early–Middle Ordovician transition (ca. 472 Ma) and proceeded diachronously northeastward up the Laurentian margin in three separate tectophases that reflect convergence events at three successive continental promontories (Kay and Colbert, 1965; Ettensohn, 1991, 1994; Ettensohn and Brett, 2002). The first or Blountian tectophase reflects collision of the southeastern margin of Laurentia with an island arc and/or various crustal fragments or microcontinents (Vick et al., 1987; Faill,
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1997; McClellan et al., 2005a, 2005b). Except for formation of the nearly continent-wide, post-Sauk or Knox unconformity and a relatively local, but deep and narrow, foreland basin (Sevier basin; see Shanmugam and Walker, 1980), the first or Blountian tectophase had little effect on cratonic sedimentation, and there is no evidence of upwelling (Holland and Patzkowsky, 1997). The absence of effect may reflect a relatively narrow belt of deformation and the inability of the deformational load to fully surmount the steep Laurentian coastal margin in the area of the Alabama and Virginia promontories, and the fact that the deep Blountian foreland basin (Sevier basin) acted as an intercepting clastic sink between the Blountian uplands and the craton. The Sebree Trough had not yet formed, and there is no evidence that the Blountian foreland basin opened directly into the open ocean. By early Late Ordovician (Chatfieldian; 454 Ma) time, however, the Taconian orogeny had progressed northward, and in the area of the New York promontory a different kinematic style developed, which involved eastward subduction of the Laurentian plate below an island arc, with resulting obduction onto the Laurentian coastal margin (e.g., Malpas and Stevens, 1977; Rodgers, 1987; Faill, 1997; Hatcher, 1999). Some workers interpret this to be the main deformational event that marked the Taconian orogeny (e.g., Rodgers, 1971, 1987), or as Kay and Colbert (1965) suggested, the Taconic phase of the Taconian orogeny. In the foreland area, moreover, inception of the Taconic tectophase is marked by the regional Black River–Trenton unconformity, a series of bentonites, regional transgression, collapse of the Blackriverian carbonate platform and development of isolated platforms and shelves (Fig. 1), as well as abrupt development of a mixed carbonateclastic facies mosaic that is interpreted to have accumulated in temperate waters across much of east-central Laurentia (Keith, 1989; Ettensohn et al., 2002). Formation of the temperate-water Lexington Platform and shelf areas to the north seems to have paralleled generation of the Sebree Trough in little more than three million years of Chatfieldian time (Ettensohn et al., 2002). Collapse of the Blackriverian Platform, formation of the Sebree Trough along old structural lows, regional deepening, and structurally related depositional lows with thickened carbonate accumulations, reflected in the lower Lexington Limestone (Kulp, 1995; Ettensohn and Kulp, 1995; Pope and Read, 1997b; Ettensohn et al., 2002, 2004), all support an overall, far-field, foreland, tensional regime related to Taconic kinematics at the subduction zone (Fig. 3A). Ettensohn et al. (2002) suggested that the abrupt change from warm-water to temperate-water carbonates at the TurinianChatfieldian boundary was also related to the collapse of the Blackriverian Platform and the resulting development of structural highs like the Lexington Platform and Galena-Trenton Shelf and intervening lows like the Sebree Trough (Fig. 1). The highs acted as foundations for the extensive buildup of carbonates that would become the Lexington Platform and Galena-Trenton Shelf. The intervening Sebree Trough, however, was sufficiently depressed so that it must have made contact with deep, openmarine waters in the Ouachita Sea south of the Laurentian margin
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Figure 3. Schematic diagrams showing kinematic regimes during parts of the Taconic tectophase of the Taconian orogeny and the likely origin of foreland, far-field forces on the Lexington Platform relative to subduction polarity in the orogen. No scale intended; positions of Lexington Platform, foreland basin, and orogen are relative only. Thick, dark arrows represent likely directions of predominant stresses. (A) Extensional regime during early to mid-Chatfieldian time with east-dipping subduction in the orogen; coeval with breakup of Blackriverian carbonate platform; infilling of foreland basin and local pull-apart basins in central Kentucky. (B) Compressional regime from mid-Chatfieldian to early Edenian time with west-dipping subduction in the orogen; coeval with generation of sub–Sulfur Well unconformity and Tanglewood buildup, regional deepening, and cratonic tilting; foreland basin overflows so that the Lexington Platform, shelves, and Sebree Trough are flooded with siliciclastic sediments (from Ettensohn et al., 2004). SL—sea level.
(Holland and Patzkowsky, 1997; Kolata et al., 2001; Ettensohn et al., 2002) (Fig. 4). With this open-ocean contact, the paleogeographic and paleoclimatic setting of east-central Laurentia at the time was such that a quasi-estuarine circulation would have been established west of the Taconian mountains (e.g., Witzke, 1987), moving warm surface waters to the southwest while funneling deep, mineral-rich oceanic waters northeastwardly through the Sebree Trough to replace them on platform and shelf areas through upwelling (Fig. 4). Upwelling onto the shallow shelves and platform could then explain the abrupt change to temperatewater conditions, the phosphorite-rich nature of the carbonates on either side of the Sebree Trough, and the rapid aggradation of skeletal carbonates on platform and shelf highs (Kolata et al., 2001; Ettensohn et al., 2002). While skeletal carbonates were accumulating on the Lexington Platform, erosional remnants from the Sebree Trough (Keith and Wickstrom, 1993; Hohman, 1998) indicate that dark shales interbedded with micrograined limestones were being deposited there, but the only manifestation of these shales on the platform are shale-rich units like the Logana and Brannon Members (Figs. 2 and 5), deposited there during major flooding events. The above situation became possible only in Late Ordovician time because of large-scale tectonic and paleogeographic changes that had been set in motion by Early Cambrian time. In particular, by Early Cambrian time Iapetan rifting had separated a large rectangular Laurentian block, now called the Precordillera terrane,
from the southeastern margin of Laurentia and sent it moving southeastward for eventual collision with the western margin of Gondwana (Thomas and Astini, 1996). This rifting event opened the large rectangular embayment called the Ouachita Sea in Figure 4 and generated a series of abortive rifts from the continental margin into the craton interior (Thomas and Astini, 1996; Kolata et al., 2001; Ettensohn et al., 2002). At this time the long axis of Laurentia was oriented nearly east-west, parallel to the equator, but by Early Ordovician time, when easternmost parts of Laurentia began approaching the Taconian subduction zone, the continent started a counterclockwise rotation toward the Middle–Late Ordovician long-axis orientation (Scotese and McKerrow, 1991) shown in Figure 4. Although upwelling could have occurred in any intervening continent orientation, the Middle–Late Ordovician paleogeographic orientation is especially efficacious for craton-interior upwelling, because the direction of resultant offshore water movement was nearly parallel to that of deep influent flow via the Sebree Trough (Fig. 4), potentially generating a stronger, more persistent exchange of waters and upwelling. Although paleogeographic conditions would have also favored coastal upwelling into nearby peri- and epicontinental seas on the southern continental-platform margin of Laurentia throughout the entirety of Ordovician time, not until the advent of the Taconic tectophase in mid-Mohawkian (early Chatfieldian) time did cratoninterior upwelling become important. By this time, eastward subduction in the Taconic tectophase had generated a tensional
Controls on Late Ordovician temperate-water conditions
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Figure 4. Early Late Ordovician (Chatfieldian) paleogeographic reconstruction of Laurentia, showing how resultant surface-water movement (interaction of trade-wind vector and Coriolis force) promoted quasi-estuarine circulation, forcing deep, cool, oceanic waters from Ouachita Sea northeastward into Sebree Trough to upwell onto the Lexington Platform and the cross-trough shelves. Coastal upwelling may have contributed to similar cool-water carbonates in the Texas–Oklahoma–New Mexico and Nevada areas. Note the subtropical setting of southern Laurentia; the southeastern parts of the continent were probably influenced by the warm Southern Tropical Current near the eastern coast (from Ettensohn et al., 2002).
regime in eastern Laurentia that reactivated older Keweenawan, Grenvillian, and Iapetan structures (Fig. 3A), causing collapse of the Sebree Trough area and opening the craton interior to the influx of deep oceanic waters from the Ouachita Sea (Ettensohn et al., 2002, 2004). Subsidiary upwelling onto the craton may have periodically come from the foreland basin to the east (Holland and Patzkowsky, 1997), but upwelling there probably could not have been sustained over long periods because bottom waters may have been warm and saline (Railsback et al., 1990), and the
foreland basin apparently lacked a direct, deep-water connection to the open ocean for replenishing cold bottom waters (Fig. 4). This situation persisted until mid- to late Chatfieldian (Shermanian) time, when the region underwent abrupt deepening, regional tilting, structural reactivation, subaqueous erosion of the Trenton-Maquoketa or sub–Sulfur Well unconformity, and inundation of platforms and shelves with fine-grained clastic sediments from the east, now represented by the Clays Ferry, Kope, Point Pleasant, “Utica,” and Maquoketa Formations (Fig. 5).
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Figure 5. NW-SE chronostratigraphic section from the Galena Shelf in central Illinois to the Tanglewood buildup of the Lexington Platform, based on five cores and one exposure and showing possible temporal relationships among the Galena Shelf, Sebree Trough, and Lexington Platform. Vertical lines below the section reflect relative positions of basement structures that influenced the above stratigraphic configuration; diagonal hatching represents missing section along unconformities. The Black River–Trenton unconformity apparently represents bulge moveout accompanying inception of Taconic tectophase; Trenton-Maquoketa or sub–Sulfur Well unconformity probably represents regional tilting and flooding accompanying change in subduction polarity (see Fig. 3). Logana (Oll) and Brannon (Olb) shale tongues represent major flooding events from an earlier Sebree Trough filling event. T—Turinian; PP—Point Pleasant Formation; CF—Clays Ferry Formation. Members of the Lexington Limestone: Oll—Logana Member; Olg—Grier Member; Olt—Tanglewood Member; Olsw—Sulphur Well Member; Olm—Millersburg Member (adapted from Ettensohn et al., 2002).
Except for the latest Chatfieldian and Edenian carbonates of the upper Lexington Limestone (those above the sub–Sulfur Well unconformity; Figs. 2 and 5) in central Kentucky, which exist nowhere else on the Lexington Platform, carbonate deposition was terminated across the platform by abrupt deepening and clastic influx. The area of central Kentucky where carbonate deposition (upper Lexington Limestone) persisted into early Edenian time has been called the Tanglewood buildup (Ettensohn, 1992), and it is interpreted to represent a structurally elevated carbonateshoal complex (Tanglewood Member, Lexington Limestone) that developed on a series of nearly triangularly arranged, reactivated basement fault zones in the area (Ettensohn, 1992; Ettensohn and Kulp, 1995; Ettensohn et al., 2002, 2004) (Figs. 1, 2, 3B, and 5). In contrast to the 15–70-m thicknesses of the Lexington-Trenton Limestones that occur elsewhere on the platform, in the area of the Tanglewood buildup the Lexington Limestone is >100 m thick, and upper parts of the unit pinch out into deeper water shales and micrograined limestones (Clays Ferry, Kope, and Point Pleasant Formations) in all directions (Cressman, 1973) (Figs. 2 and 5). These mid- to late Chatfieldian events across the platform and the Sebree Trough are approximately coeval with an east-to-
west shift in Taconic subduction polarity (Coakley and Gurnis, 1995; Karabinos et al., 1998; Karabinos and Hepburn, 2001) and probably reflect a corresponding change to a far-field, foreland compressional regime related to a new kinematic regime at the subduction zone (Ettensohn et al., 2004) (Fig. 3B). DISCUSSION AND CONCLUSIONS Proximity of the Lexington Platform to the Taconian orogen (Fig. 1) would have almost certainly guaranteed major tectonic influence on platform development, but the case for glaciogenic influence on its development has only recently been made through the indirect use of carbonate rocks interpreted to be temperatewater deposits as proxies for glaciation through upwelling (Pope and Steffen, 2003). In fact, all workers seem to agree that deposition of temperate-water carbonates in shallow-shelf or platform settings, particularly in a subtropical belt like that occupied by southern parts of Laurentia in Late Ordovician time (Fig. 4), requires some form of upwelling to bring in cool waters. However, glaciation is not a necessary condition for upwelling; more important are features of atmospheric circulation relative to
Controls on Late Ordovician temperate-water conditions paleogeography (Parrish, 1982; Kolata et al., 2001). Nonetheless, Pope and Read (1997a, 1998), Pope and Steffen (2003), and Saltzman and Young (2005) have used the evidence of early Late Ordovician (Chatfieldian) temperate-water carbonates and equivalent sediments to support the idea of initial Late Ordovician glaciation at the Turinian-Chatfieldian boundary, ca. 10 Ma earlier than the widely accepted terminal-Ordovician, Hirnantian glacial maximum. They reasoned that because glacial periods are times of more vigorous thermohaline circulation, the likelihood of upwelling would be greater, and temperate-water carbonates should be more extensive. This could have well been the case in continental-platform margin settings where coastal upwelling was the only practical form of upwelling (Parrish, 1982), but in Chatfieldian time interpreted temperate-water carbonates were present in places like Quebec, Ontario, western New York, Ohio, Indiana, and Illinois, including locales >1000 km away from the nearest open continental platform margin. Although the influence of possible Chatfieldian glaciation cannot be precluded, simple coastal upwelling (Parrish, 1982) is probably adequate to explain upwelling in the New Mexico–Texas–Oklahoma study area of Pope and Steffen (2003) and in the Nevada study area of Saltzman and Young (2005); both areas were within a few hundred kilometers of open continental margins, as were southern parts of the Lexington Platform (Fig. 4). Saltzman and Young (2005) also looked for evidence of geochemical and sea-level changes that might reflect the initiation of glaciation. Near the base of their Chatfieldian section they noted the prominent positive δ13C excursion, known as the Guttenberg isotope carbon excursion (GICE) or the Chatfieldian excursion (Hatch et al., 1987), which may be global in nature (Ainsaar et al., 2004). Saltzman and Young (2005) further suggested that this excursion reflected enhanced organic-carbon burial in response to the inception of glaciation and upwelling. Although this is a plausible explanation, it is not the only possible explanation, and Ainsaar et al. (2004) suggested the additional possibility of increased Late Ordovician orogenic activity as yet another way to change global ocean chemistry, a possibility that was similarly suggested by Shields et al. (2003) on the basis of Sr- and O-isotope changes at the Middle–Late Ordovician transition. In addition, Tobin and Walker (1997) and Tobin et al. (2005) suggested that the excursion may reflect global salinity and temperature changes based on δ18O values, whereas Ludvigson et al. (1996) and Simo et al. (2003) called upon increased photosynthetic activity near the sea surface to explain the excursion. Simo et al. (2003) and Panchuk et al. (2005) also cautioned that these excursions may be the product of different carbon sources in different places, rather than indicating any one source, global or otherwise. As glaciation should be accompanied by a major drop in sea level, Saltzman and Young (2005) looked for evidence of a eustatic drop that coincided with the δ13C excursion and other proxies, and found evidence of a mid-Chatfieldian truncating sequence boundary coupled to a regressive event that had previously been interpreted to reflect tectonic adjustment in central
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Nevada (Cooper and Keller, 2001). They, in contrast, thought that the event was more likely related to glacioeustatic drawdown. Again, this could be a plausible interpretation, but on a global scale (e.g., Vail et al., 1977; Hallam, 1984), and on a more regional scale in the Appalachian Basin (e.g., Dennison and Head, 1975; Dennison, 1989) and in the Midcontinent (Ross and Ross, 1992; Shutter, 1992; Holland and Patzkowsky, 1996; Witzke and Bunker, 1996; Kolata et al., 2001), a similar eustatic drop is not seen. In fact, all the above sources generally point to a time of latest Turinian–Chatfieldian sea-level rise. A similar event, represented by the Black River–Trenton unconformity, apparently occurred at the Turinian-Chatfieldian boundary at the base of the Lexington-Trenton Limestones (Figs. 2 and 5) on the Lexington Platform, on the Galena Shelf, in the Sebree Trough, and in parts of the Appalachian Basin. Distribution of this unconformity (Ettensohn, 1994) and its coincidence with inception of the Taconic tectophase have led many workers to interpret the event and unconformity as tectonic in origin (e.g., Kolata et al., 1998; Ettensohn, 1994; Ettensohn et al., 2002). Also occurring on the Lexington Platform is the TrentonMaquoketa or sub–Sulfur Well unconformity (Figs. 2 and 5), which is late Chatfieldian in age. As already discussed, this break in sedimentation corresponds to a major flooding event, cratonic tilting, clastic inundation of the carbonate platform, and structural reactivation, resulting in the Tanglewood buildup (Ettensohn et al., 2002). Although this event was not regressive across most of the Lexington Platform, it reflects a major change in subduction vergence, and with it a corresponding change in cratonic kinematics to a far-field compressive regime, which was apparently expressed differently on different parts of Laurentia, depending on the local structural framework. Moreover, if the continent was surrounded by connected subduction zones, as suggested by Scotese and McKerrow (1991), kinematic reorganization on one margin may have initiated nearly coeval tectonism on other margins, and, as Karner and Watts (1983) and Ziegler (1987) pointed out, far-field stresses can be transmitted up to several thousand kilometers across continental lithosphere via zones of crustal weakness. So the possibility of widespread Laurentian cratonic responses to Taconic tectonism with related responses in sea level and accommodation space cannot be ignored. The point of the above discussion, however, is not to argue against possible glacial influence in the development of sedimentary and chemostratigraphic features on the southern and southwestern margins of Laurentia but to suggest that other viable possibilities should also be considered. In fact, near continental margins where coastal upwelling would have normally been effective, glaciation may well have enhanced upwelling and related temperate-water conditions as suggested, and any tectonic effects could have augmented them. However, in a craton-interior setting more than a thousand kilometers from an open continental margin, as, for example, New York and southern Quebec were at the time (Fig. 4), where normally warm-water conditions would have prevailed, it is difficult to call on glacial enhancement of circulation alone to explain the sudden development of temperate-water
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conditions. Directly subjacent warm-water Blackriverian or Turinian units, as well as paleogeographic reconstructions that indicate that the area was being bathed by the warm Southern Tropical Current (Wilde, 1991), suggest that southeastern Laurentian seas were “warm” at the surface for most of Ordovician time. Despite the predominance of temperate-water carbonates across the area after the Blackriverian–Trenton or the Turinian– Chatfieldian transition, the continued growth of warm-water stromatoporoids high on shoal trends, where they would have contacted warm surface waters throughout Lexington deposition, as well as the presence of warm-water inlier units like the Perryville and Devils Hollow Members (Fig. 2), the depositional setting of which behind protective shoals or beaches obviated the effects of cool waters at depth, indicate that the surface regime was definitely warm water. Moreover, Kolata et al. (2001) indicated that the cool-water mixing zone extended only ~400 km into the craton from the Sebree Trough before cool-water, bryozoanbrachiopod-echinoderm grainstones graded into coeval warmwater mudstones. Not only does this indicate that warm-water carbonate deposition was proceeding elsewhere in the area, it also shows that the Sebree Trough was the source of the cool waters, and that the extent of cool-water influence was ~400 km from the source of upwelling. At present the figure of ~400 km, derived from the work of Kolata et al. (2001), is the only estimate for the extent of cool-water influence from upwelling, glacially enhanced or not, onto a broad Paleozoic shelf before its effects dissipated. Cressman (1973), Kolata et al. (2001), and Ettensohn et al. (2002) have all indicated that the Sebree Trough was the source of Chatfieldian upwelling. The fact that apparent upwelling and the inception of temperate-water carbonate deposition on either side of it occurred coevally, with the origin of the trough at the beginning of the Taconic tectophase, points to the likely significance of Taconian far-field tectonism in the abrupt changeover to temperate-water conditions at the Turinian–Chatfieldian transition (Ettensohn et al., 2002). Tectonism generated the conduit (Sebree Trough) from the continental margin that allowed ingress of cool, deep waters into the continent (Fig. 4) as well as the later foreland tilting and deepening that probably enhanced the effect and extent of the upwelling. Stratigraphically and biostratigraphically, the early inception of Late Ordovician glaciation is poorly constrained such that the best current consensus is that glaciation probably began in late Mohawkian–Cincinnatian (Chatfieldian–Edenian) time (Eyles, 1993; Crowell, 1999). However, also by late Mohawkian time tectonism had become widespread across the globe and was prominent on Laurentia, where, by the beginning of Chatfieldian time, the major phase of the Taconian orogeny had begun. The orogeny had major consequences in the structural and stratigraphic differentiation of the formerly adjacent warm-water (Blackriverian) carbonate platform and its abrupt breakup into several smaller shelves and platforms with new temperate-water sedimentary and faunal regimes in what should have been a lowlatitude, warm subtropical belt. Although it is neither possible nor
reasonable to exclude the influence of glaciation on low-latitude carbonate platforms in early Late Ordovician (mid-Mohawkian or Chatfieldian) time, it is important to note that structure, tectonics, and paleogeography may have combined to generate similar regime changes. In fact, in the circumstances described for southeast Laurentia (Fig. 4), it seems plausible that just this situation transpired to form temperate-water conditions across southern Laurentia. However, perhaps even more probable at this juncture, is the likelihood that incipient glacial and tectonic agencies acted synergetically to produce the noted changes. Although glaciation is always an attractive explanation for seemingly anomalous temperate-water conditions in low-latitude settings, in the right situation coeval changes in tectonic framework can generate similar conditions or act to accentuate glacial effects. East-central and southern Laurentia may have provided such a setting, and, based on current understanding of Ordovician paleogeography and tectonics, suggests that other Ordovician continental settings may have existed where similar warm-to-temperate-water changes might be expected. ACKNOWLEDGMENTS This paper was initially presented as part of a symposium on “The Global Ordovician System” at the 32nd International Geological Congress. I wish to thank the editors, S. Finney and W.B.N. Berry, for their invitation to submit this paper. I also wish to thank John Coates for helpful discussion and important references, as well as S.A. Leslie and an anonymous reviewer for comments that helped to greatly improve the manuscript. REFERENCES CITED Ainsaar, L., Meidla, T., and Martma, T., 2004, The Middle Caradoc facies and faunal turnover in the Late Ordovician Baltoscandian paleobasin: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 210, p. 119–133, doi: 10.1016/j.palaeo.2004.02.046. Berner, R.A., 1991, A model for atmospheric CO2 over Phanerozoic time: American Journal of Science, v. 249, p. 339–376. Berner, R.A., 1994, GEOCARB II, a revised model of atmospheric CO2 over Phanerozoic time: American Journal of Science, v. 301, p. 56–91. Berner, R.A., and Kothavala, Z., 2001, Geocarb III: A revised model of atmospheric CO2 over Phanerozoic time: American Journal of Science, v. 301, p. 182–204, doi: 10.2475/ajs.301.2.182. Brenchley, P.J., 1984, Late Ordovician extinctions and their relationship to the Gondwana glaciation, in Brenchley, P.J., ed., Fossils and Climate: New York, John Wiley, p. 291–315. Brenchley, P.J., 1989, Late Ordovician extinction, in Donovan, S.K., ed., Mass Extinctions: Processes and Evidence: London, Bellhaven Press, p. 104–132. Brenchley, P.J., and Newell, G., 1980, A facies analysis of Upper Ordovician regressive sequences in the Oslo region, Norway—A record of glacioeustatic changes: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 31, p. 1–37, doi: 10.1016/0031-0182(80)90002-4. Brenchley, P.J., Marshall, J.D., Carden, G.A.F., Robertson, D.B.R., Long, D.G.F., Meidla, T., Hints, L., and Anderson, T.F., 1994, Bathymetric and isotopic evidence for short-lived Late Ordovician glaciation in a greenhouse period: Geology, v. 22, p. 295–298, doi: 10.1130/0091-7613(1994) 022<0295:BAIEFA>2.3.CO;2. Brenchley, P.J., Carden, G.A.F., Hints, L., Kaljo, D., Marshall, J.D., Martma, T., Meidla, T., and Nõlvak, J., 2003, High-resolution stable isotope stratigraphy of Upper Ordovician sequences: Constraints on the timing of
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Printed in the USA
The Geological Society of America Special Paper 466 2010
Correlations across a facies mosaic within the Lexington Limestone of central Kentucky, USA, using whole-rock stable isotope compositions John W. Coates Noble Energy Incorporated, 100 Glensborough Drive, Houston, Texas 77067, USA Frank R. Ettensohn Department of Geological Sciences, University of Kentucky, Lexington, Kentucky 40506, USA Harold D. Rowe University of Texas Arlington, 500 Yates Street, Arlington, Texas 77019, USA
ABSTRACT The chemostratigraphic (δ13Ccarb) record of the Lexington Limestone in central Kentucky is a high-resolution record in Chatfieldian and lower Edenian strata. Parts of the Lexington Limestone in this region reflect an anomalous structurally controlled carbonate buildup with complex facies relationships. Chemostratigraphic analysis of buildup and off-buildup composite sections, which exemplify the Lexington facies mosaic, reveals an overall decreasing trend in 13C compositions throughout the formation and into the overlying Clays Ferry Formation. Superimposed on this trend are four locally correlative excursions. The most significant excursion is found within the Logana Member. The Guttenberg excursion is locally expressed as two prominent δ13C peaks (maxima +2.58‰). The excursion is important because of its stratigraphic position above the regional and globally correlated Millbrig K-bentonite, which allows for correlation with chronostratigraphically equivalent successions. The ubiquity of the Guttenberg excursion has been recognized throughout the eastern and central United States and the Baltic region. Additional excursions have been detected and are common in other sections, showing that the chemostratigraphic record of the Lexington Limestone is similar in buildup and off-buildup sections, even across complex lithofacies boundaries. The similarity of the record across complex facies, moreover, suggests that these localized excursions are “events,” which can be used to constrain correlations and augment known facies relationships. Furthermore, the overall chemostratigraphic trend in the Lexington Limestone appears similar to published δ13C values from equivalent strata in Baltoscandia, emphasizing the global correlative value of chemostratigraphic trends in Upper Ordovician strata.
Coates, J.W., Ettensohn, F.R., and Rowe, H.D., 2010, Correlations across a facies mosaic within the Lexington Limestone of central Kentucky, USA, using wholerock stable isotope compositions, in Finney, S.C., and Berry, W.B.N., eds., The Ordovician Earth System: Geological Society of America Special Paper 466, p. 177– 193, doi: 10.1130/2010.2466(12). For permission to copy, contact
[email protected]. ©2010 The Geological Society of America. All rights reserved.
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Coates et al.
Central Baltoscandia Confacies Belt Rägavere Fm.
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Hirmuse Fm.
Kinnekulle K-bentonite
Kahula
Logana Mbr.
Lexington Limestone
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Tyrone Formation
Hermitage Formation Carters Limestone
Guttenberg Formation Spechts Ferry Formation
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Upper Central Mississippi Tennessee Valley
Decorah
Chatfieldian
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Caradocian
Mohawkian
Sandbyan
Katyan
Series
Formation
Carbon-isotope excursions are abrupt changes in carbonisotopic composition relative to underlying and overlying strata. In particular, positive carbon-isotope excursions have been widely interpreted to reflect times of increased deposition of 12 C-enriched organic matter (Kump and Arthur, 1999). One of the most studied intervals of the Paleozoic has been the Chatfieldian Guttenberg isotope carbon excursion (GICE), which has recently been referred to simply as the Chatfieldian excursion. Although the Guttenberg excursion is well defined, additional isotope excursions during Chatfieldian time have now been recognized, so that the term Chatfieldian excursion will become increasingly inappropriate. In the following parts of this paper, the GICE will be referred to as Guttenberg in order to distinguish it from other Chatfieldian excursions. This excursion has been identified from isotopic analysis of organic shales as well as the components of carbonate rocks, and apparently has global correlative value. The Guttenberg excursion was a significant chemostratigraphic and paleoceanographic event ~10 m.y. prior to the end of Ordovician time. This event was first observed in the Guttenberg Member of the Decorah Formation in Iowa by Hatch et al. (1987) and has been widely recognized in chronostratigraphically equivalent strata in Illinois, Iowa, and Wisconsin (Ludvigson et al., 1996, 2000, 2004; Pancost et al., 1999), Kentucky (Coates et al., 2004; Coates, 2006), Tennessee (Saltzman et al., 2000), Nevada (Saltzman and Young, 2005), Pennsylvania (Patzkowsky et al., 1997), New York (Bergström et al., 2001), Oklahoma and Virginia (Young et al., 2003), as well as in the Baltic states (Ainsaar et al., 1999, 2004). The excursion is stratigraphically a few meters above the Millbrig K-bentonite (ca. 453.7 ± 1.3 Ma; Tucker and McKerrow,
1995), which has been used for regional correlations based on wireline and geophysical logs (Haynes, 1994; Kolata et al., 1996) as well as through chemical fingerprinting techniques. As a result of these techniques, the excursion has also been correlated with the Kinnekulle K-bentonite (Fig. 1) from the Baltic states of Estonia and Latvia (Huff et al., 1992; Bergström et al., 2004). Coupling the Guttenberg excursion with the stratigraphic proximity of the widespread Millbrig-Kinnekulle K-bentonites supports the apparent global nature of the chemostratigraphic signal. Moreover, recent studies suggest that the Guttenberg excursion is the result of enhanced organic-carbon burial that may have reduced pCO2 (Kump et al., 1995) to levels that would have triggered ice buildup (Saltzman and Young, 2005) and initiated a pre-Hirnantian glacial event during late Middle to early Late Ordovician time (Pope and Steffen, 2003). New and independent data from the Lexington Limestone of central and northern Kentucky (Coates et al., 2004; Coates, 2006) is presented here to describe the chemostratigraphic record on the Lexington Platform, as well as to augment the middle to upper Chatfieldian–lower Edenian chemostratigraphic record on southern Laurentia. A similar, coeval, positive excursion has been recognized in the Logana Member of the Lexington Limestone in central Kentucky (Saltzman et al., 2000; Coates et al., 2004; Coates, 2006), and its stratigraphic position with respect to the Millbrig K-bentonite (Fig. 1) (Haynes, 1994; Huff et al., 1996) suggests correlation with other Trenton limestones and equivalent strata that record similar δ13C-enriched zones (i.e., GICE). However, on the eastern and southern parts of the Lexington outcrop area the Logana pinches out and is replaced by a thicker underlying Curdsville Member and equivalent parts of the overlying Grier Member and their equivalents. One of the goals of
Salona
INTRODUCTION
Curdsville Mbr.
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Figure 1. Correlation of units in the Eastern United States at the Turinian-Chatfieldian boundary relative to the Baltic States. Modified from Leslie and Bergström (1995).
Correlations across a facies mosaic using whole-rock stable isotope compositions this research was to test the utility of correlating the Guttenberg excursion across the Lexington Platform, even in areas where the Logana (micrite-rich limestones) does not exist, to determine whether or not the excursion is facies-dependent (i.e., recorded in non-micritic limestones) and to detect other possible isotopic trends that may have correlative value. Data obtained from whole-rock analyses of samples from the upper Lexington Limestone and the overlying Clays Ferry Formation also provide a more complete record of middle to upper Chatfieldian and lower Edenian chemostratigraphic trends, which have not been presented before at such a high degree of resolution. Two composite sections, constituting the entire Lexington Limestone, as well as lower and upper parts of bounding units, were sampled for carbon-isotope analysis. The overall chemostratigraphic δ13C signature of the Lexington Limestone appears to partially cover the gap within the Chatfieldian-Edenian δ13C record and appears to correspond with published trends during this interval. Overall, the focus of this paper is to introduce new data for the Chatfieldian-Edenian δ13C chemostratigraphic record and present data for perspective regional correlations. STRATIGRAPHIC-DEPOSITIONAL FRAMEWORK The Lexington Limestone (Fig. 2) is a facies mosaic composed of generally shallow, temperate-waterargillaceous limestones and shales that were deposited across the Lexington Platform (Ettensohn, 1992; Ettensohn and Kulp, 1995; Pope and Read, 1997). The Lexington Platform (Fig. 3) was a large, rectangular foreland block formed during the breakup of the extensive Blackriverian carbonate platform from Taconian far-field forces (Ettensohn et al., 2002). The platform was bound on all sides by deeper water environments, including the Martinsburg foreland basin to the east, a probable outlet to the Ouachita Sea to the south, and the Sebree Trough to the north and west (Fig. 3). The occurrence of many phosphorite-rich beds and hardgrounds throughout the Lexington Limestone and elsewhere on the platform probably reflects the upwelling of cool, nutrient-rich waters from the Sebree Trough (Pope and Read, 1997; Ettensohn et al., 2002; Kolata et al., 2001). The Lexington Limestone is separated from the warm-water calcilutites of the underlying Tyrone Formation of the High Bridge Group by the Black River–Trenton unconformity (Kolata et al., 1998; Cressman, 1973; Holland and Patzkowsky, 1996; Ettensohn et al., 2002). Based upon measured sections and biostratigraphy, Cressman (1973) and Ettensohn (1992) showed that parts of the Lexington Limestone grade laterally and vertically into the deeper water shales with interbedded micrograined limestones of the Clays Ferry, Kope, and Point Pleasant Formations (Figs. 2 and 4). The Lexington Limestone (Fig. 2) is subdivided here informally into lower and upper units. The lower Lexington is composed of the Curdsville through Brannon Members (Figs. 2 and 4). Across the Lexington Platform the lower Lexington is largely coeval to what most workers have called the Trenton Limestone
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and its stratigraphic equivalents, which typically range from 15 to 70 m in thickness (Shaver, 1985). The lower Lexington is dominated by cross-bedded, coarse-grained calcarenites in the Curdsville Member and by nodular to irregularly bedded argillaceous calcarenites in the Grier Member. Where present, even-bedded calcisiltites and shales of the Logana Member represent a deepramp, southeast-thinning unit that pinches out proximal to the Lexington Fault System (Fig. 5). A bed of brachiopod coquina is typically present in the middle of the Logana and is laterally persistent where the Logana pinches out. The brachiopod coquina is typically phosphatic and locally bounded by hardgrounds. Cressman (1973) suggested that the Logana is transgressive below and regressive above the coquina unit, whereas Ettensohn et al. (2004) suggested that the unit represents an easterly transgression from the Sebree Trough. The hardgrounds bounding the coquina suggest that the unit is correlative with a maximum-flooding surface in the lower, third-order cycle of the upper MohawkianCincinnatian supersequence (Pope and Read, 1997). Based on isopach maps and measured sections (e.g., Butler Quarry), the Logana reflects flooding from the west and northwest from the Sebree Trough. Past studies (Cressman, 1973; Borella and Osborne, 1978) showed that the Curdsville, Grier, and their stratigraphic equivalents are laterally extensive but grade laterally to the north and west into dark interbedded shales and calcisiltites of the Maquoketa and possibly Point Pleasant Formations in the Sebree Trough (Keith, 1989; Ettensohn et al., 2002). The middle to upper Grier intertongues complexly and is essentially a facies mosaic in itself. The upper Grier represents a depositional continuum of Logana-like, deep-ramp environments (Macedonia and Cane Run beds; Fig. 6) to intermediate-ramp, nodular-bedded argillaceous Grier calcarenites and shallowramp calcarenite shoal complexes of the lower Tanglewood (Fig. 6). In southwestern parts of the central Kentucky study area a well-developed lower Tanglewood shoal (Mackey, 1972) was deposited around reactivated structures and formed a protective barrier, behind which thick warm-water calcilutites represented by the Faulconer and Salvisa beds of the Perryville Member were deposited (Figs. 2 and 4). The overlying even-bedded calcisiltites and shales of the Brannon Member reflect a local transgression event, in which deep-water carbonates inundated the Lexington Platform. The Brannon Member becomes coarser and eventually grades northward into the calcarenites and calcirudites of the Tanglewood (Kulp, 1995; Ettensohn and Kulp; 1995) (Figs. 2 and 4). The upper part of the lower Lexington is truncated by the sub– Sulphur Well unconformity, which marks the upper boundary of the lower third-order cycle in the upper Mohawkian-Cincinnatian supersequence (Pope and Read, 1997) (Figs. 2 and 4). Separating lower and upper Lexington units is the sub– Sulphur Well (Ettensohn et al., 1986, 2002; Hohman, 1998) or Trenton-Maquoketa (Keith and Wickstrom, 1993) unconformity. The upper Lexington includes the units above the sub–Sulphur Well unconformity and largely represents a regressive, shoalingupward pile of skeletal carbonates called the Tanglewood buildup (Ettensohn, 1992) (Figs. 2–5). The upper Lexington occurs only
Series
British Stage
Global Stages
North American Stages Midcontinent Conodont Biozones
Winchester SE
Formation
Frankfort CS N-NW
Informal units
Depositional Sequences (Holland and Patzkowsky, 1996)
Danville CS SW
Clays Ferry
CF
PDHU Devils Hollow
Stamping Ground Sub-Sulphur Well unconformity
Brannon
Tanglewood (lower tongue)
Brannon
Cane Run
Cornishville
Tanglewood (lower tongue)
Grier Mac
Grier
Lower
Faulconer
Salvisa
Chatfieldian
Tanglewood (middle tongue)
Lexington Limestone
M6
Sulphur Well
Upper Ordovician
Millersburg
Upper
SC
Caradocian
C1
Edenian
MB
P. tenuis
Tanglewood (upper tongue)
Clays Ferry
B. confluens Katyan
Clays Ferry
Logana
Curdsville Curdsville Millbrig
M4
Blackriverian-Trenton unconformity
P. undatus Sandbyan
M5
Figure 2. Generalized stratigraphy of the Lexington Limestone in central Kentucky. Conodont biostratigraphy based upon Sweet (1984) and Richardson and Bergström (2003). Depositional sequence stratigraphy based upon Holland and Patzkowsky (1996). Note occurrence of the Blackriverian-Trenton unconformity at the base and the sub–Sulphur Well unconformity. Not drawn to scale. SC—Strodes Creek; CS—Composite Section; MB—Millersburg Member; CF—Clays Ferry Formation; and PDHU—Post–Devils Hollow Unconformity.
Correlations across a facies mosaic using whole-rock stable isotope compositions
181
H
PE
U
LA
O
FO
M R O
IA ON
P
XI LE
Argillaceous carbonates or interbedded carbonates and shale
TA C
Shale
V
Coarser siliciclastic rocks
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G
SE
TO
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R
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EE
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Clean carbonate rocks
RE
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0
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0
500 km
Figure 3. Schematic map view showing the sedimentary and tectonic subdivision of the previous Blackriverian carbonate platform into the Galena and Trenton Shelves, Sebree Trough, and Lexington Platform. PE—Pennsylvania Embayment; T—Tanglewood buildup; L—Louisville high; PP—Point Pleasant Basin. Adapted from Keith (1989).
Siliciclastic rocks overlying carbonates Approximate limit of marine rocks
EDENIAN
in central Kentucky as a roughly triangular buildup (Figs. 3 and 5). The Frankfort composite section (Fig. 6) is representive of the buildup and serves as our “type” chemostratigraphic section; it is present near the thickest part of the buildup (Fig. 7). The Danville and Butler quarry sections (Figs. 8 and 9) represent offbuildup sections to the south and north, respectively. Because of the buildup in central Kentucky, the Lexington Limestone (lower and upper) attains a thickness >100 m (Cressman, 1973; Shaver, 1985). The limestone unit is thickest in the Frankfort-Lexington area (Fig. 7) near the apex of the Jessamine Dome and thins through intertonguing in the upper part in all directions away from the apex (Figs. 4 and 7). The upper Lexington Limestone is dominated by coarse-grained, shoal-related calcarenites and calcirudites represented by the middle and upper tongues of the Tanglewood Member. Hraber et al. (1971) and Cressman (1973) noted the thickening of the Tanglewood in the upper Lexington, whereas Borella and Osborne (1978) suggested that the Tanglewood shoals were structurally related. On the basis of the mapped distribution of the Tanglewood, Ettensohn (1992), Ettensohn and Kulp (1995), and Ettensohn et al. (2002) further supported the idea that the area of triangular shoal facies (Fig. 5) was a regressive carbonate buildup (e.g., Heckel, 1974) formed by uplift along now largely extant fault zones, whose basement precursors were reactivated by Taconian far-field forces. In addition, the lateral distributions of many Lexington members are interpreted as having been controlled by the reactivation of basement structures in the area (Mackey, 1972; Grossnickle, 1985; Ettensohn et al., 1986; Kulp, 1995; Ettensohn and Kulp, 1995; Kasl, 2001; Ettensohn et al., 2004; Jewell, 2001; Jewell and Ettensohn, 2004; Ettensohn et al., 2004). Drahovzal et al. (1992) and Drahovzal and Noger (1995) discuss many of these controlling basement
structures. Owing to limited vertical and lateral facies distribution of representive members of the Lexington Limestone, three sections from different regions on the Lexington Platform were chosen for chemostratigraphic analysis (Figs. 6–10). SAMPLING LOCALITIES Complete exposures of the Lexington Limestone are rare, but by using a few closely located exposures to form composites and a few key partial exposures, nearly complete sets of samples from representive buildup and off-buildup sections have been collected for analysis. Figure 4 illustrates the relative position of the sampled chemostratigraphic sections in relation to stratigraphy in the central Kentucky region. Sections A and B on U.S. 127 (Fig. 7) were used to develop the Frankfort composite section (Fig. 6) and contain the complete succession of units near the thickest area on the buildup. In contrast, the Danville composite section (sections C and D; Fig. 8) in Danville, Kentucky, is on the margin of the buildup. There, the Tanglewood shoal facies is replaced by the Sulphur Well Member, which is a transitional facies between the shallowwater Tanglewood shoal facies and the deeper water, off-buildup Clays Ferry Formation (Figs. 2 and 4). In addition, the section exhibits some unusual structurally controlled facies development in the Perryville Member (Mackey, 1972), which is not present elsewhere in the Lexington Limestone (Figs. 2 and 4). The northernmost section in Pendleton County, Kentucky, at the Butler Quarry (Fig. 9), is off the buildup and on the northern margin of the Lexington Platform. The upper part of the Lexington Limestone, reflecting the buildup, is completely absent here, having been replaced by deeper water shales and micrograined
182
Coates et al. SSW
NNE
Clays Ferry
Kope Formation Clays Ferry Formation
DH Sulphur Well Mbr.
pDHU
Millersburg Member
Tanglewood Member SG
sSW
Brannon Member
TWB
Tanglewood Member Point Pleasant Formation
Upper Ordovician Chatfieldian Lexington Limestone
Edenian
Garrard Formation
Tanglewood Member Perryville Mbr. Grier Member
Logana Member Curdsville Member
BRTU
Tu T Dcs
Fcs
M D
GQ
BQ
Figure 4. Schematic cross section across the Tanglewood buildup in central Kentucky, along C–E illustrated from Figure 2, showing the complex intertonguing relationship of various members, the Tanglewood buildup and its relationship to deeper water units, and the relative positions of sections studied. Datum is the Blackriverian-Trenton unconformity (BRTU). BQ—Butler Quarry; GQ—Georgetown Quarry; Fcs—Frankfort composite section; Dcs—Danville composite section; T—Tyrone Formation; Tu—Turinian Stage; M—Millbrig K-bentonite; D—Deicke K-bentonite; pDHU—Post–Devils Hollow Unconformity; DH—Devils Hollow Member; SW—Sulphur Well Member; SG—Stamping Ground Member.
limestones of the Point Pleasant Formation (Fig. 4). In addition, the deeper water incursion reflected in the Logana Member is well represented here (Fig. 4). METHODS In order to adequately characterize the Lexington Limestone’s chemostratigraphic profile, this study utilized whole-rock samples, but micritic-rich samples were used where possible. Consequently, some samples were obtained from coarse-grained calcarenites consisting of variable amounts of allochems, which have variable original carbonate chemistries (i.e., aragonite, highMg and low-Mg calcite). The resulting data exhibit consistent
low standard deviations and similar overall chemostratigraphic trends throughout the study area. This suggests that these carbonates were probably rock-buffered during diagenesis (Banner and Hanson, 1990), which retained the dominant δ13C values, but that other chemical components of these limestones were found to consist of variable compositions, indicating that they were chemically altered (i.e., δ18O) (Coates, 2006). Carbonate samples (n = 428) were recovered from six roadcut exposures along a north-south transect from Danville to Butler, Kentucky (Fig. 7, C–E). Hand samples were collected at 2-ft (0.6-m) intervals for detailed chemostratigraphic analysis. Additional samples from the same stratigraphic zones were collected for testing the reliability of δ13C compositions.
Correlations across a facies mosaic using whole-rock stable isotope compositions
0
7.5
15
miles
0
2.3
4.6
km
183
N
c
b a b
Figure 5. Map of central Kentucky, illustrating the major fault systems and outline of the Tanglewood buildup on the Lexington Platform. Darkened linear blocks parallel basement structural trends and some extant faults, and define the outer limits of the Tanglewood buildup; a—Kentucky River Fault System; b—Lexington Fault System; c—Centerville Fault. Modified from Ettensohn (1992).
All samples were crushed and pulverized using the methods of Kaljo et al. (2001). Sample aliquots were analyzed for carbonand oxygen-isotope composition using a ThermoFinnigan Gas Bench II peripheral coupled to a DeltaPlusXP mass spectrometer housed at the University of Kentucky Environmental Research and Training Laboratory. Samples were weighed to 400 µg. Standard deviations were determined from every 15 samples. Samples were purged with research-grade helium and subsequently reacted in 0.25 mL of 100% phosphoric acid for 10 h at 60 °C. All results were reported with respect to the V-PDB (Vienna Peedee belemnite). The average precision for the δ13C and δ18O measurements was ±0.06‰ and ±0.12‰, respectively. The precision for the NBS-19 standard was always better than 0.10‰ for both δ13C and δ18O. Only the δ13Ccarb compositions will be discussed, because the large lateral and vertical variability in δ18O compositions suggests postdepositional alteration, which is typical of many carbonate sequences (Marshall, 1992).
DATA Frankfort Composite Section (Localities A and B) The Frankfort composite section (Fcs) includes the Frankfort 1 (N 38° 13.232′, W 84° 51.219′) and Swallowfield (N 38° 18.589′, W 84° 50.619′) road-cut exposures, measured along U.S. 127, north of Frankfort, Franklin County, Kentucky. This section encompasses the upper Tyrone Formation; lower Lexington Limestone members, including the Curdsville, Logana, Grier (including the Macedonia and Cane Run Beds); and the lower tongue of the Tanglewood, as well as upper Lexington members including the Stamping Ground, lower and upper tongues of the Millersburg and Tanglewood, Devils Hollow, and Clays Ferry Formation (Figs. 4 and 6). The Fcs features an overall decreasing trend in δ13C values (Fig. 6). Superimposed on this negative trend are two prominent
Midcontinent Conodont Biozone
Series
Member/ Formation
100 –2
Belodina confluens
Upper Upper Millersburg Tanglewood Mbr. Mbr.
Clays Ferry
Edenian
Relative Sea Level
m’s –1
+2
+3
Deflection E Deflection D
90
Deflection C 80 Deflection B Post-Devils Hollow unconformity
Devil Lower Millersburg
+1
0
70
M-TW Mbr. Stamping Mbr.
Sub-Sulphur Well unconformity 60
Cane
Grier Mbr.
50
40 Mac
Grier Mbr.
Lexington Limestone
Plectodina tenuis
Chatfieldian
Lower Tanglewood Mbr.
h f
g
Excursion A
30
Logana Mbr.
c
Guttenberg Excursion (GICE)
b 10
a
CC
Blackriverian-Trenton unconformity
0 –5
+2
+3
Unconformity
Interbedded Calcisiltite and Shale
Nodular Bedded CoarseGrained Calcarenite and thin shale beds interbeds
Calcilutite
Nodular, Argillaceous Coarse-grained Calcalrenite with interbedded shale
Bioclastic Coarse-Grained Calcarenite/Calcirudite w/ shale partings
Basinal
+1 13 δ Ccarb
0
Deep Ramp
–1
Sandbelt/Shoal
–2
Open Marine - Intermediate Ramp
D
Platform/Shelf Lagoon
Curdsville Mbr.
e
d
Shallow Ramp - Tidal Flat
B. compressa
Tyrone
Turinian
Phragmodus undatus
20
Correlations across a facies mosaic using whole-rock stable isotope compositions
185
Figure 6. Carbon-isotope data from the Frankfort carbon-isotopic composite section; Farkfort; Franklin County, Kentucky. Conodont biozones are approximated from Bergström and Sweet (1966). Mac—Macedonia Bed of the Grier Member; Cane—Cane Run of the Grier Member; Devil—Devils Hollow Member; Stamping—Stamping Ground Member. The unconformity between the Tyrone Formation of the High Bridge Group and the Curdsville Member of the Lexington Limestone will be referred to as the sub-Lexington and regionally correlated with the Blackriverian-Trenton unconformity (Kolata et al., 1998). The unconformity at the base of the Stamping Ground Member is referred to herein as the sub–Sulphur Well unconformity and correlated with the sub-Trenton–Maquoketa unconformity (Keith and Wickstrom, 1993). This section was developed from whole-rock samples measured every 2 ft (0.6 m) from the Frankfort 1 (base of composite to sub–Sulphur Well unconformity) and Swallowfield (above the Sulphur Well unconformity) sections (Fig. 2). Bentonite beds are the Deicke (D) and Capitol City (CC) K-bentonites. Relative sea level diagram (right) is modified from Ettensohn et al. (2002).
0
5
10
15
20 mi
140
N
E
160 180 200 220 240 260 280
B
300 320
A
320 b
260 280 240 220 200 180
D
300 a b
C
Figure 7. Isopach map of the Lexington Limestone (thickness in feet) in north-central Kentucky; from Cressman (1973). Thick area in the center is the Tanglewood buildup (Ettensohn, 1992). Section locations: A, B—Danville composite section; C, D—Frankfort composite section; E— Butler Quarry; a—Kentucky River Fault System; b—Lexington Fault System. Inset map shows the approximate boundaries of the research area. Cross section C–E is shown schematically in Figures 4 and 10.
m’s Conodont Member/ Series Biozone Formation
70 +1
0
60
+2
+3
Relative Sea Level
Deflection C
Sulphur Well Mbr.
Lex.
–1
Deflection B
50
Sulphur Well Mbr.
Sub-Sulphur Well unconformity
Tanglewood
Corn
40
Salvisa
PerryvilleFaulconer Bed Lower Tanglewood Mbr.
k 30
j i
Excursion A
h
g
f 20
Guttenberg Excursion (GICE)
c a
b
Grier Mbr.
Blackriverian-Trenton unconformity
0
0
+1
δ13Ccarb
+2
+3
Unconformity
Calcilutite
Nodular Bedded CoarseGrained Calcarenite and thin shale beds interbeds
Interbedded Calcisiltite and Shale
Fossiliferous calcilutite
Nodular, bryozoan calcirudite/calcarenite with interbedded shales
Nodular, Argillaceous Coarsegrained Calcalrenite with interbedded shale
Basinal
–1
Deep Ramp
–2
Open Marine - Intermediate Ramp
–5 Sandbelt/Shoal
M
10
Platform/Shelf Lagoon
Lexington Limestone
Brannon
Shallow Ramp - Tidal Flat
B. compressa
Clays Ferry
Tyrone Formation Curdsville Mbr.
Plectodina tenuis
Belodina confluens
Clays Ferry Formation
Phragmod usundatus
Turinian
Chatfieldian
Edenian
–2
Bioclastic Coarse-Grained Calcarenite/Calcirudite w/ shale partings
Figure 8. Carbon-isotope data from the Danville composite section, Danville, Boyle County, Kentucky. Lower Tanglewood and Tanglewood refer to the lower tongue of the Tanglewood Member. Salvisa is the middle bed of the Perryville Member. Corn is the Cornishville Bed of the Perryville Member. At the base of the Lexington Limestone, M refers to the Millbrig K-bentonite. Note a negative shift toward the lighter isotope values across both unconformities, similar to the Frankfort composite section. Relative sea level diagram (right) is modified from Ettensohn et al. (2002).
Correlations across a facies mosaic using whole-rock stable isotope compositions m’s 50
Relative Sea Level +2
+3
40
30
20
10
Guttenberg Excursion
e
d
c b
a
+1
δ13Ccarb
+2
Bioclastic Coarse-Grained Calcarenite/Calcirudite w/ shale partings
+3
Nodular Bedded CoarseGrained Calcarenite and thin shale beds interbeds
Basinal
0
Deep Ramp
–1
Open Marine - Intermediate Ramp
0 –2
Interbedded Calcisiltite and Shale
+1
0
Sandbelt/Shoal
Grier Mbr. Logana Mbr. Curdsville Mbr.
–1
Platform/Shelf Lagoon
Point Pleasant Formation Lower Tanglewood Mbr
Lexington Limestone
Belodina confluens Plectodina tenuis Phragmodus undatus
Chatfieldian
Edenian
–2
Shallow Ramp - Tidal Flat
Conodont Member/ Series Biozone Formation
187
Figure 9. Carbon-isotope data from the Pendleton Quarry section, Pendleton County, Kentucky. A thicker Logana Member at Butler, Kentucky, records three positive peaks (points a, c, and e) similar to those at the Frankfort composite section. The sampling density is sparse compared with the Frankfort and Danville composite sections, but enough data are present to determine the δ13Ccarb Logana excursion. Relative sea level diagram (right) is modified from Ettensohn et al. (2002).
positive excursions and four conspicuous positive and negative deflections. The prominent characteristic signature of the Lexington Limestone is the δ13C excursion within the Logana Member (up to +2.58‰), interpreted to be the Guttenberg (GICE) excursion. An additional excursion was recognized close to the stratigraphic position of the Macedonia Bed of the Grier Member (up to +1.4‰), referred to here informally as excursion A. Three minor positive δ13C deflections were recognized and are defined as deflection C (up to −0.2‰), and deflection E (−0.7‰). Moreover, two minor negative deflections were also noted; these
are defined as deflections B (down to −1.6‰) and D (down to –1.3‰). Danville Composite Section (Localities C and D) The Danville composite section (Dcs) includes the Caldwell Stone Company Quarry section (N 37° 37.50′, W 84° 45.17′; Boyle County, Kentucky) and the Herrington Lake section (N 37° 40.910′, W 84° 40.797′; Boyle-Garrard county line). The Dcs contains the upper Tyrone Formation, Curdsville, Grier, lower
SW
C,D
M –2
–2
–1
–1
i
e
+2
+1
b
g
+2
c a
+3
BRTU
+3
D
CC
–2
–2
–5
O lc
O ll
O lg
O t
–1
lgm
O
lg
O
–1
lt O lgcr
O
O lt O lm O ldh O lsg
lm
O
lt
O
O cf
d b
+1
+1
δ13Ccarb
0
g
0
e
h f
+2
+2
c
+3
a
+3
Frankfort composite section
A,B
E –2
lc
ll
–2
–1
–1
d
+1
+1
δ13Ccarb
0
b
0
e
+2
a
+2
c
Butler Quarry Section
O
O
O lg
O lt
O pp
m’s
+3
+3
10 m’s
NE
Figure 10. Comparative chemostratigraphic cross section C–E, showing correlation through δ13C signatures between the Danville and Butler sections in Kentucky. FAD—first appearance datum; BRTU—Blackriverian unconformity; sSU—sub–Sulphur Well unconformity; M—Millbrig K-bentonite; Ocf—Ordovician Clays Ferry Formation; Olt— Ordovician Lexington Limestone Tanglewood Member; Olm—Ordovician Lexington Limestone Millersburg Member; Oldh—Ordovician Lexington Limestone Devil’s Hollow Member; Olsg—Ordovician Lexington Limestone Stamping Ground; Olgcr—Ordovician Lexington Limestone Grier Member Cane Run Bed; Olg—Ordovician Lexington Limestone Grier Member; Olgm—Ordovician Lexington Limestone Grier Member Macedonia Bed; Olc—Ordovician Lexington Limestone Curdsville Member; Opp—Ordovician Point Pleasant Formation; Oll—Ordovician Lexington Limestone Logana Member.
d
f
h
sSU
+1
δ13Ccarb
0
0
Sowerbyella rugosa FAD
Danville composite section
Tyrone Fm.
Curdsville Mbr.
Grier Mbr.
Tanglewood Mbr.
Faulconer Bed
Tanglewood Mbr. Brannon Mbr. Cornishville Bed Salvisa Bed
Sulphur Well Mbr.
Clays Ferry Fm.
Sulphur Well Mbr.
Clays Ferry Fm.
Sowerbyella rugosa FAD
Perryville Mbr.
Correlations across a facies mosaic using whole-rock stable isotope compositions tongue of the Tanglewood, Perryville (including the Faulconer, Salvisa, and Cornishville Beds), Brannon, and Sulphur Well Members, and the Clays Ferry Formation (Figs. 4 and 8). Overall, the Dcs shows an overall decreasing trend in δ13C values similar to the Fcs. A major excursion similar to that at the Fcs was recognized within the lower tongue of the Tanglewood (up to +2.1‰). Based on its stratigraphic position relative to the Millbrig K-bentonite and an abrupt increase in δ13C values, we recognize this as the Guttenberg (GICE) excursion. In addition to the Guttenburg excursion, two minor, positive carbon deflections and one prominent negative deflection were recognized. The lowermost positive deflection recognized in the Faulconer Bed is assumed to be the same as excursion A (up to +1.3‰) in the Fcs. The uppermost positive deflection, C, falls close to the Sulphur Well (Lexington Limestone) and Clays Ferry Formation contact (up to −0.1‰) and is the same as the similarly designated deflection from the upper Tanglewood Member at the Fcs. The only negative carbon-isotope deflection noted is in the middle of the Sulphur Well Member (down to –1.6‰), and it is correlated with deflection B. Also noteworthy is the fact that the trend in the upper Tyrone Formation starts at an initial baseline d13C value of +1.2‰ and steadily decreases upward toward the BlackriverianTrenton unconformity, a trend similar in Tyrone Formation equivalents noted by Ludvigson et al. (2001, 2002, 2004).
189
and Kolata, 1990). Hence, the term Mud Cave will be abandoned hereafter and merely addressed as the Millbrig K-bentonite. Similarly, in the Nashville Dome region, the Millbrig is found at or near an unconformable contact separating the Carters Limestone and the overlying Hermitage Formation. Leslie and Bergström (1995) suggested that Turinian-Chatfieldian limestones in the Nashville and Jessamine dome areas are chronostratigraphically equivalent on the basis of the stratigraphic position of the Millbrig. Moreover, the Millbrig has been correlated with the Kinnekulle K-bentonite, exposed in the Baltoscandia Region of northern Europe (Huff et al., 1992; Bergström et al., 1995; Bergström et al., 2004). The Millbrig and Kinnekulle beds were correlated by phenocryst mineralogy, discriminant analysis of trace-element compositions (Huff et al., 1992), similar 40Ar/39Ar and U-Pb isotope data (Huff et al., 1992; Bergström et al., 1995), and by conodont-based graphic correlation methods (Sweet, 1984, 1988). It should be noted, however, that new research has put the transatlantic correlation of these two bentonites in doubt (Min et al., 2001; Samson, 1996; Haynes et al., 1995). In summary, the occurrence of the Millbrig K-bentonite at the base of the Lexington Limestone confirms the position of the unit within the larger, global chronostratigraphic framework (Fig. 1). The Logana Excursion and Its Chemostratigraphic Equivalents
Butler Quarry (Locality E) The Butler Quarry (N 38° 47.797′, W 84° 20.347′; Figs. 4 and 9), in Pendleton County, Kentucky (Fig. 7), exposes the upper Curdsville, Logana, Grier, and Tanglewood (lower tongue) Members, as well as the overlying Point Pleasant Formation. Similar to the Frankfort and Danville composite sections, the overall trend of the δ13C values decreases upward through the exposure. In addition to the overall negative trend, a major positive excursion (up to +2.2‰) within the lower half of a stratigraphically thicker Logana Member is recognized as the Guttenberg excursion. Unfortunately, the high vertical walls of the quarry restricted sampling to widely spaced intervals so that high-resolution sampling could not be completed. Generally, the Lexington Limestone chemostratigraphic signature in northern Kentucky records the overall negative δ13C trend and the Guttenberg excursion. DISCUSSION Fitting the Lexington Limestone into a Chronostratigraphic Framework In the Jessamine Dome region of central Kentucky, the Lexington Limestone unconformably overlies the Tyrone Formation, and locally the “Mud Cave” bentonite. The Mud Cave bentonite was correlated with the Millbrig K-bentonite and other equivalent bentonites in the eastern United States through physical tracing in outcrop, wireline logs, and chemical fingerprinting (Huff
The Logana Member is underlain by a thin Curdsville Member in the Frankfort area and thins southward to pinchout at Danville, where it is replaced by intermediate-ramp (Grier) to shallow-ramp (Curdsville and Tanglewood) carbonates (Fig. 4). Examination of the chemostratigraphic record of both the Frankfort and Danville composite sections generally displays similar trends through this stratigraphic interval (Figs. 6 and 8). The Logana at the Fcs records two prominent peaks designated points a and c (Fig. 6), having δ13C values of 2.5‰ and 2.2‰, respectively. The Danville composite section (Fig. 8) has, within a thin tongue of the Tanglewood Member, two positive peaks (a and c) having δ13C compositions of +2.0‰ and +2.1‰, respectively. In both composite sections, these excursions are relatively large positive deflections (>+1.5‰), coming from an average baseline of +0.6‰ throughout the Curdsville and the lower Grier (Danville only) members. Based on position above the base of the Lexington Limestone, we suggest that deflections on both curves a and c are, respectively, coeval (Fig. 10). The position near the base of the Lexington Limestone and earlier correlations (Hatch et al., 1987; Ludvigson et al., 1996; Smith et al., 2000; Patzkowsky et al., 1997) indicate that our Logana excursion is in fact equivalent to the Guttenberg (GICE) δ13C excursion of Young et al. (2003). The calcisiltites (micrite dominated) and shales of the Logana have been interpreted to represent deep-ramp to basinal carbonates (Cressman, 1973; Ettensohn et al., 2002) deposited well below storm wave base. The Logana has also been interpreted to reflect a period of sea-level rise such that deep waters
190
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from the Sebree Trough encroached on the Lexington Platform from the north and west (Ettensohn et al., 2002). The rising arm of the Guttenberg excursion at Danville appears to rise just above the Curdsville-Logana (Fcs) or Grier-Tanglewood (Dcs) contact. Because the Logana reflects major flooding from the Sebree Trough, it also probably reflects an influx of nutrients related to upwelling (Ettensohn et al., 2002). While the northern and western parts (Fcs and Butler Quarry) of the Lexington Platform were being inundated, eastern and southern parts (Dcs) of the platform were probably structurally higher and underwent increased rates of skeletal deposition. Thus, relative to the base of the Lexington, the excursion is stratigraphically higher by nearly 5 m more in the Danville area than in Frankfort and comparable northern areas of the platform. Local and Regional Patterns of Isotope Data The Frankfort and Danville composite sections exhibit two positive carbon isotope excursions (GICE and excursion A) and four prominent carbon-isotope reflections (deflections B–E) within the overall Lexington Limestone δ13C chemostratigraphic profile. Both composite sections have similar trends with respect to their chemostratigraphic records. The apparent GICE correlation across complex facies boundaries, similar overall negative trends of δ13C values, and low standard deviations in respective carbonate samples from these two composite sections suggest that the deflections record the same events as do chronostratigraphically equivalent strata in Iowa (Ludvigson et al., 1996), Pennsylvania (Patzkowsky et al., 1997), and the Baltic states (Ainsaar et al., 1999) and indicate a global signature reflecting a possible synchronous “event.” Hence, not only are there global implications for correlation with the Lexington Limestone but also for correlations across complex facies within the formation. The chemostratigraphic correlations in the Lexington Limestone also show that the δ13C excursions and prominent deflections are not facies (i.e., lithologic) dependent. Conversely, because the post-GICE interval has not been studied in detail, excursion A and the other “deflections” have no correlatives except locally. Thus, until there are additional data from Lexington Limestone– equivalent strata (post-Guttenberg intervals), the chemostratigraphic profile of the Lexington Limestone can only be explained as local effects of carbon cycling on the Lexington Platform. A smaller, positive carbon-isotope excursion is also recognized in the mid-Grier (Fcs) (Fig. 6) and in probably equivalent parts of the Faulconer Member at Danville (Fig. 8). Two δ13C peaks (f and h) define what is referred to as excursion A, but an additional peak is recognized at Danville. These carbon-isotopic peaks appear to correlate in relation with the stratigraphic position above the Guttenberg excursion and below the sub–Sulphur Well unconformity (Fig. 10). At Danville, excursion A occurs in the Faulconer Bed, a lagoonal unit that formed behind the topographically higher Tanglewood shoals (Mackey, 1972). In the Frankfort section the same excursion apparently occurs in the Macedonia Bed of the Grier Member, which represents a flood-
ing event from the Sebree Trough onto the Lexington Platform during a period of elevated sea level (Ettensohn et al., 2002). Although the apparently equivalent excursion A from the Macedonia Bed at Frankfort exhibits higher δ13C values (up to +1.35‰) than the Faulconer values (up to +1.30‰), these two sections represent the same chemostratigraphic event. The deeper marine carbonates of the Macedonia Bed reflect a period of deeper water deposition throughout most of the Lexington Platform (except in the southwest areas of the Lexington Platform). Although the Macedonia Bed is thin compared with the Logana Member, this suggests that this was a short-term flooding event, probably reflecting a period of increased upwelling of nutrients from the Sebree Trough and causing the positive excursion similar to the GICE. Moreover, excursion A is approximately in the middle to upper Plectodina tenuis Biozone. Fanton and Holmden (2001) recognized three positive excursions in the same zone within the Dunleith Formation (Upper Mississippi Valley), and Kaljo et al. (2004) recognized similar expressions they refer to as the “first-late Caradoc” excursion from Baltoscandia. On the basis of stratigraphic position and biozones, it is likely that these Baltoscandia and Dunleith Formation excursions are correlative and are possibly correlative with the δ13C excursion recognized in the Faulconer and Macedonia beds (Figs. 6, 8, and 10). Additional stratigraphic constraints need to be addressed, but it is likely that the excursions from these three separate areas may reflect yet another global signal. The chemostratigraphic record from the overlying parts of the Lexington, subsequentially follows a progressive shift in δ13C just above excursion A through and across the sub–Sulphur Well unconformity and continues to its lowest δ13C values of –1.66‰, which is labeled as deflection B near the Chatfieldian-Edenian boundary (Figs. 6 and 8). This decreasing trend in δ13C appears to coincide with a time of deeper water environments. Moreover, the development of a mineralized, phosphate-rich sub–Sulphur Well unconformity probably reflects a subaqueous omission surface (Hohman, 1998; Ettensohn et al., 2002), which truncates the Brannon Member and equivalent members. This surface represents the top of the lower Lexington depositional sequence and the base of the upper Lexington. The upper Lexington section appears to coincide with regional transgression characterized by expansion of deeper water Maquoketa muds from the Sebree Trough from the west and by siliciclastic flux from the east (Clays Ferry and Kope Formations) onto most of the Lexington Platform (Ettensohn et al., 2004). At the same time, in central Kentucky, uplift on reactivated structures initiated a carbonate buildup of calcarenitic Tanglewood shoals in shallow-ramp environments with transitional environments (Millersburg and Sulphur Well Members) into deep-ramp and basinal environments (Clays Ferry, Kope, and Point Pleasant Formations) away from this buildup (Figs. 2 and 4). Concomitant with this depositional scheme change, a polarity shift in subduction (Coakley and Gurnis, 1995; Karabinos et al., 1998) related to the ongoing Taconic orogeny to the east, may have played a role in subsidence of the Lexington Platform and/or reactivation of local bathymetric structural highs
Correlations across a facies mosaic using whole-rock stable isotope compositions and lows on the platform (Ettensohn, 1991; Ettensohn et al., 2002, 2004). The overall decreasing δ13C trend of the Lexington Limestone appears to be related to marine transgression or siliciclastic influx from the east. Lower Lexington members appear to have consistent δ13C values, averaging +0.40‰ (Figs. 6, 8, and 9), whereas the upper Lexington Limestone members average –0.37‰. This depletion trend probably reflects the local-scale carbon cycling within the semi-restricted marine environments upon the Lexington Platform. Between these two baselines, which separate the lower and upper Lexington units, the sub–Sulphur Well unconformity represents an incursion of deeper water environments as terrigenous silts and clays of the Clays Ferry and Kope Formations from the east (via erosion of the Taconic highlands) inundated the Lexington Platform. The arrival of depleted δ13C (~ –0.5, Kump and Arthur, 1999) silicates indicates that this was probably the primary factor of an overall decreasing δ13C trend of the upper Lexington Limestone. Panchuk et al. (2006) and Kump and Arthur (1999) discussed a variety of causes for positive or negative excursions. The addition of weathered carbonates and siliciclastics from the Taconic highlands, resulting in adding dissolved inorganic carbon via continental runoff, would have effectively decreased the 13C composition of the Lexington Limestone carbonates. The resulting overall negative chemostratigraphic shift appears to have followed the decreasing δ13C value trend in the lower Chatfieldian (Saltzman and Young, 2005). Additional data need to be collected in post-Guttenberg intervals in order to determine if this chemostratigraphic signature from the Lexington Limestone is correlatable on the local, regional, or possibly global scale. Implications of Chemostratigraphic δ13C Signature of the Lexington Limestone The chemostratigraphic record of the Lexington Limestone, recorded in two different composite sections, exhibits similar trends and suggests that negative and positive maxima, or peaks of isotopic excursion, are “event-like” beds. Although isotope compositions do vary slightly from locality to locality, the overall trends across different lithofacies suggest pronounced excursions that were in chemical equilibrium with the ocean water at the time of deposition. The most conspicuous chemostratigraphic trend is the consistent decrease of δ13C values during the Grier-Tanglewood transition, culminating in negative excursion B. If trends like this can be shown to be synchronous, or at least penecontemporaneous, then δ13C signatures have tremendous correlative value in helping to constrain the vertical extent of complex lithofacies boundaries in facies mosaics like the Lexington Limestone in central Kentucky (Figs. 2 and 4). For example, the Sulphur Well Member (Figs. 2 and 4) is an argillaceous, nodular-bedded, bryozoanrich calcarenite and calcirudite with interbedded shales that attains a maximum thickness of 11 m (Cressman, 1973). Based on cross sections, it is a transitional, intermediate-ramp unit on
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the flank of the buildup that grades northward and eastward into the middle tongue of the Tanglewood and thins southward and westward into the Clays Ferry Formation (Fig. 4). If changes in δ13C composition like that in excursion B and C can be utilized as “event” beds, then the chemostratigraphic signatures from Danville and Frankfort composite sections (Fig. 10) suggest that the upper boundary of the Lexington Limestone is stratigraphically higher and formed temporally later in Frankfort than in Danville, as might be expected, because the Frankfort section developed on the Tanglewood buildup, which was a site of carbonate deposition lasting much longer than at the Danville locality, which was on the southern flank of the buildup. Nonetheless, additional field work and biostratigraphy will be required to definitely prove that the chronostratigraphic excursions can be defined as synchronous “events.” CONCLUSIONS The chemostratigraphic record of the Lexington Limestone reflects an overall decreasing trend of δ13C compositions with two positive excursions (i.e., the Guttenberg and the A). In addition, two prominent positive (C and E) and two negative (B and D) deflections have local correlative value. The close proximity to the Sebree Trough and the abundance of phosphorite-rich hardgrounds throughout the formation suggest that the driving mechanism for the positive excursions was increased bioproductivity from periods of enhanced upwelling in which nutrients were transferred from the deep, cold waters of the Sebree Trough onto the Lexington Platform. Conversely, the chemostratigraphic record of the Lexington Limestone is an overall negative trend toward decreasing δ13C values, which probably resulted in the delivery of depleted δ13C weathered carbonates and siliciclastics onto the semi-restricted waters of the Lexington Platform. Owing to its stratigraphic proximity to the globally correlated Millbrig K-bentonite, the pronounced δ13C positive excursion within the Logana Member suggests and supports correlation to chronostratigraphically equivalent strata regionally and globally via the Guttenberg (GICE) excursion. The similarity of the chemostratigraphic record in different stratigraphic localities on the Lexington Platform suggests that δ13C signatures are not facies dependent and can be used to correlate units across complex facies regimes. This suggests that the 13C composition of the carbonates was probably in chemical equilibrium with paleoceanic waters during Chatfieldian and early Edenian time. Moreover, the apparent relative timing of these excursions and deflections probably reflects “events” that can help to constrain correlations among complex facies and to augment field studies where biostratigraphic data are unavailable or limited. REFERENCES CITED Ainsaar, L., Meidla, T., and Martma, T., 1999, Evidence for a widespread carbon isotopic event associated with late Middle Ordovician sedimentological and faunal changes in Estonia: Geological Magazine, v. 136, no.1, p. 49–62, doi: 10.1017/S001675689900223X.
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