CONTOURITES
Edited by
M. Rebesco and
A. Camerlenghi
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Elsevier Science Radarweg 29, PO Box 211, 1000 AE Amsterdam, The Netherlands The Boulevard, Langford Lane, Kidlington, Oxford OX5 1GB, UK First edition 2008 Copyright Ó 2008 Elsevier B.V. All rights reserved No part of this publication may be reproduced, stored in a retrieval system, or transmitted in any form or by any means, electronic, mechanical, photocopying, recording, or otherwise, without the prior written permission of the publisher. Permissions may be sought directly from Elsevier’s Science & Technology Rights Department in Oxford, UK: phone (+44) (0) 1865 843830; fax (+44) (0) 1865 853333; email:
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PREFACE
Contourites are sediments deposited or substantially reworked by bottom currents. The study of contourites is nowadays crucial for several fields of fundamental and applied research: 1. palaeoclimatology and palaeoceanography, since these fairly continuous and relatively high-resolution sediments hold the key for priceless information on the variability in ocean circulation patterns, current velocities, oceanographic history and basin interconnectivity; 2. hydrocarbon exploration, since accumulation of source rocks may be favoured by weak bottom currents, whereas ‘‘clean’’ deep-sea sands may be formed by robust flows; 3. slope stability, since fine-grained, low-permeability, high pore-water content contourites facilitate the formation of overpressurized gliding planes when rapidly loaded, or when their rigid biosiliceous microfabric collapses due to diagenetic alteration. Despite its significance, this group of sediments is poorly known by the majority of non-specialists. Notwithstanding the growing interest and the now intensive research in contourites, a textbook that serves as a reference book on contourites was missing until now. This book addresses all aspects of the knowledge in the field of contourites and provides a comprehensive and cross-referenced coverage of the subject. It can serve as a standard reference work for non-specialists, and in particular postgraduate students, university teachers and lecturers, researchers and professionals who are seeking an authoritative source of information about contourites. It reviews both theoretical topics and case histories, and it points to approaches that may help tackle open problems. Divided into a wide and interdisciplinary spectrum of topical sections, it provides practical advice on multidisciplinary research techniques. The ample use of illustrations, diagrams, photographs and maps, complemented by a CD-ROM including all illustrations (as many as possible in colour), provides a helpful tool for researchers in the preparation of classroom lectures and training courses, journal articles and meeting presentations. The figures presented in this book are partly new, partly adapted from previous works, and partly reproduced from previous works. Wherever appropriate, permission has been obtained for the reproduction or adaptation, but not all copyright holders could be traced. We would appreciate if such copyright holders would be so kind as to contact us. The authors invited to contribute to this book are distinguished specialists in the field. They have been invited not simply to produce a research paper on the topic, but to critically review the current state of knowledge. They have been asked to accomplish their review in a clear and concise way, including as much factual.
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Preface
information as possible. They have been encouraged to submit as many tables and illustrations (diagrams, photographs and maps) as are necessary to complement the text of their chapter to present the information in the best possible way. The authors have been recommended to keep the language simple, and those whose native language is not English have been encouraged to choose co-authors whose mother tongue is English and/or to propose native English reviewers for their manuscripts. The editors and the authors are grateful to the reviewers who contributed to improve the scientific quality of each chapter and to the book as a whole: A. Arche, C. Bjerrum, A. Camerlenghi, A. Cattaneo, X.D. de Madron, G. Evans, F. Eynaud, J.-C. Fauge`res, M. Frenz, Z. Gao, M. Gardner, J. Jones. D.W. Kirkland, P. Knutz, J.S. Laberg, R. Larter, E. Llave, D. Long, J. Lo´pez-Go´mez, L. Lo¨wemark, L. Masse´, T. Mulder, H. Nelson, M. Rebesco, C. Silva, P. Talling, D. Thornalley, R. Urgeles, A.J. van Loon, M. Vanneste, A. Wiewio´ra. Michele Rebesco and Angelo Camerlenghi
LIST oF CONTRIBUTORS
A. Camerlenghi ICREA, c/o Universitat de Barcelona, Departament d’Estratigrafia, Paleontologia i Geocie`ncies Marines, GRC Geocie`ncies Marines, C/Martı´ i Franque`s, s/n, E-08028 Barcelona, Spain.
[email protected] B. Chaco´n Fachbereich Geowissenschaften, Universita¨t Bremen, D-28334 Bremen, Germany.
[email protected] A. Crise Istituto Nazionale di Oceanografia e Geofisica Sperimentale, B.go Grotta Gigante 42/c, I-34010 Sgonico (TS), Italy.
[email protected] T. Duan Marathon Oil Company, Houston, Texas, USA.
[email protected] J.-C. Fauge`res De´partement de Ge´ologie et Oce´anographie, Universite´ Bordeaux1, UMR CNRS 5805 EPOC, Avenue des Faculte´s, F-33405 Talence cedex, France.
[email protected] M.A. Fregenal-Martı´nez Department of Estratigrafı´a Facultad de Ciencias Geolo´gicas, Universidad Complutense, E-28040 Madrid, Spain.
[email protected] Z. Gao School of Geosciences, Yangtze University, Jingzhou, Hubei province, 434023, China.
[email protected] P. Giresse Laboratoire d’E´tudes des Ge´o-Environnements Marins, Universite´ de Perpignan, 52, Avenue Paul Alduy, F-66860 Perpignan, France.
[email protected] E. Gonthier De´partement de Ge´ologie et Oce´anographie, Universite´ Bordeaux1, UMR CNRS 5805 EPOC, Avenue des Faculte´s, F-33405 Talence cedex, France. .
[email protected] xix
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List of Contributors
Y. He School of Geosciences, Yangtze University, Jingzhou, Hubei Province, 434023, China.
[email protected] F.J. Herna´ndez-Molina Facultad de Ciencias del Mar, Universidad de Vigo, E-36200 Vigo, Spain. f
[email protected] J.A. Howe Scottish Association for Marine Science & UHI Millennium Institute-Dunstaffnage Marine Laboratory, Oban, Argyll, Scotland, PA37 1QA, UK.
[email protected] K.J. Hsu¨ Oakcombe, Marley Common Haslemere, Surrey, GU27 3PT, UK.
[email protected] H. Hu¨neke Institute of Geography and Geology, University of Greifswald, D-17487 Greifswald, Germany.
[email protected] S. Hunter National Oceanography Centre, Southampton (NOCS), Waterfront Campus, Southampton, SO14 3ZH, UK.
[email protected] P.C. Knutz Geological Survey of Denmark and Greenland (GEUS), Øster Voldgade 10, 1350 Copenhagen, Denmark.
[email protected] A. Kuijpers Geological Survey of Denmark and Greenland (GEUS), Øster Voldgade 10, 1350 Copenhagen, Denmark.
[email protected] J.S. Laberg Department of Geology, University of Tromsø, N-9037 Tromsø, Norway.
[email protected] E. Llave Instituto Geolo´gico y Minero de Espan˜a, Rı´os Rosas, 23, E-28003 Madrid, Spain.
[email protected] A. Maldonado Instituto Andaluz de Ciencias de la Tierra. C.S.I.C./Universidad de Granada, Campus de Fuentenueva, s/n, E-18002 Granada, Spain.
[email protected] List of Contributors
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J. Martı´n-Chivelet Department of Estratigrafı´a Facultad de Ciencias Geolo´gicas, Universidad Complutense, E-28040 Madrid, Spain.
[email protected] D.G. Masson National Oceanography Centre, Southampton (NOCS), Waterfront Campus, Southampton, SO14 3ZH, UK.
[email protected] I.N. McCave Department of Earth Sciences, University of Cambridge, Downing Street, Cambridge, CB2 3EQ, UK.
[email protected] T. Mulder De´partement de Ge´ologie et Oce´anographie, Universite´ Bordeaux1, UMR CNRS 5805 EPOC, Avenue des Faculte´s, F-33405 Talence cedex, France.
[email protected] T. Nielsen Geological Survey of Denmark and Greenland (GEUS), Øster Voldgade 10, 1350 Copenhagen, Denmark.
[email protected] M. Rebesco Istituto Nazionale di Oceanografia e Geofisica Sperimentale (OGS), Borgo Grotta Gigante 42/C, I-34010 Sgonico (TS), Italy.
[email protected] S. Salon Istituto Nazionale di Oceanografia e Geofisica Sperimentale, B.go Grotta Gigante 42/c, I-34010 Sgonico (TS), Italy.
[email protected] G. Shanmugam Department of Earth and Environmental Sciences, The University of Texas at Arlington, Box 19049 – Arlington, TX 76019 USA.
[email protected] M.S. Stoker British Geological Survey, Murchison House, West Mains Road, Edinburgh, EH9 3LA Scotland, UK.
[email protected] D.A.V. Stow National Oceanography Centre, Southampton (NOCS), Waterfront Campus, Southampton, SO14 3ZH, UK.
[email protected] xxii
List of Contributors
F. Trincardi ISMAR-CNR, Via Gobetti 101, I-40129, Bologna, Italy.
[email protected] A.J. Van Loon Geological Institute, Adam Mickiewicz University, Mako´w Polnych 16, 61-606 Poznan, Poland.
[email protected] T. van Weering Royal Netherlands Institute for Sea Research (NIOZ), P.O.Box 59, Texel, 1790 AB Den Burg, and Department of Paleoclimatology and Geomorphology, Free University, de Boelelaan 1085, 1081 HV Amsterdam, The Netherlands.
[email protected] G. Verdicchio ISMAR-CNR, Via Gobetti 101, I-40129, Bologna, Italy and EDISON SpA, Foro Buonaparte 31, I-20121, Milano, Italy.
[email protected] A.R. Viana Petrobras, Research Center, R&D Exploration, Rio de Janeiro, Brazil.
[email protected] F. Werner Institut fu¨r Geowissenschaften, Universita¨t Kiel, Olshausenstrasse 40-60, D-24118 Kiel, Germany.
[email protected] A. Wetzel Geologisch-Pala¨ontologisches Institut, Universita¨t Basel, Bernoullistrasse 32, CH-4056 Basel, Switzerland.
[email protected] D. Wilkinson National Oceanography Centre, Southampton (NOCS), Waterfront Campus, Southampton, SO14 3ZH, UK. R.B.Wynn National Oceanography Centre, Southampton (NOCS), Waterfront Campus, Southampton, SO14 3ZH, UK.
[email protected] W. Zenk Leibniz Institute of Marine Sciences at the University of Kiel (IFM-GEOMAR), Du¨sternbrooker Weg 20, D-24105 Kiel Kiel Germany.
[email protected] P A R T
1
CONTOURITE RESEARCH
C H A P T E R
1
C ONTOURITE R ESEARCH : A F IELD IN F ULL D EVELOPMENT M. Rebesco1, A. Camerlenghi2 and A.J. Van Loon3 1
Istituto Nazionale di Oceanografia e Geofisica Sperimentale (OGS), Sgonico (TS), Italy ICREA, c/o Universitat de Barcelona, Departament d’Estratigrafia, Paleontologia i Geocie`ncies Marines, Barcelona, Spain 3 Geological Institute, Adam Mickiewicz University, Poznan, Poland 2
Contents 1.1. 1.2. 1.3. 1.4. 1.5.
Bottom Currents Contourites Drifts Sedimentary Structures Prospects
4 6 8 9 10
Contourites were first recognized and described only 40 years ago when photographs of the deep-sea bottom showed distinct current ripples (Heezen and Hollister, 1964; Hollister, 1967); much progress has been made since, but the current status in this field shows clearly that this relatively new research topic is still in full development: a surge of new research results (reflected in, for instance, the publication of several special volumes of scientific journals) go still hand in hand with uncertainties and with insufficient knowledge (there is, for instance, a lack of indisputable diagnostic criteria). New techniques are continuously developed and implemented (e.g., swath bathymetry and 3-D seismics). All this indicates a great potential for further research, in the same way as the now mature research on turbidites developed in the 1960s. Comparing the research on turbidites with that on contourites, it seems that the latter is presently facing a transition from adolescence to maturity. One might consider such a period of transition not to be the right moment to compile a reference book, which is aimed at providing both experts in the field a state-of-the-art overview, and the non-specialist readers an introduction that may help starting research in this field. However, we consider it just crucial that this transitional stage of contourite research is documented. It is the right moment to increase the recognition of contour currents as important transport and sedimentary phenomena that control much of the deep-sea sedimentation; in addition, this seems the right moment to bring new forces into play. The aim of this book is therefore not to present an inventory of the fossilized knowledge of a fully developed field, but rather to provide a basis for future Developments in Sedimentology, Volume 60 ISSN 0070-4571, DOI: 10.1016/S0070-4571(08)00201-X
Ó 2008 Elsevier B.V. All rights reserved.
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Contourite Research: A Field in Full Development
development, a platform for ample exchange of complementary expertise, ideas and experience. The insight into the role of contourites will certainly benefit from input by new researchers within the – so far – fairly restricted circle of the current specialists in this area of research. A good understanding of contourites and of the sedimentary processes behind them is often complicated by a lack of clear, unambiguous and commonly accepted diagnostic criteria, and by discussions about conflicting views about the sedimentary structures. The different views that still exist are particularly confusing for nonspecialists and students, who may easily become disoriented by such complex situations. With this book, and with this introduction in particular, we want to offer to non-specialist readers a simplified, often descriptive, overview of contourites. We are aware of the risk implied in any simplifications. For this reason, we neither hide the problems nor provide oversimplified explanations. We asked the authors of the various chapters to accomplish in a clear and concise way a critical review of the current state of knowledge regarding the subject of their chapter. This implies that the book does not contain contributions reflecting new research. Nor is it our intention to present, in this introduction, new syntheses or untested new hypotheses. In contrast, we provide a kind of glossary that may help the reader to find his way, with special reference to key terms. These terms include the ‘‘generator’’ (bottom currents), the sediments (contourites), the types of accumulation (drifts) and the types of sedimentary structures (traction, bioturbation). A comprehensive text on contourites is, in our opinion, justified also because of the incomplete coverage of the subject by reference books, to which the non-specialist readers are referred before entering the details of specific journal articles: even the Encyclopedia of Sediments and Sedimentary Rocks (Middleton et al., 2003) does not include ‘‘Contourites’’ among its entries! The Glossary of Geology (Bates and Jackson, 1987) lists 43 entries for ‘‘drift’’, of which the only one referring specifically to sedimentology lists wind or river currents as main ‘‘generators’’ of sedimentary drifts.
1.1.
BOTTOM C URRENTS
‘‘Bottom current’’ is our preferred term to refer to the water-mass flows that control the deposition of contourites. Many other terms are often used, for which reason we need to clarify the meaning. Among geologists, the most adhered to definition of contourites is that they are sediments deposited or substantially reworked by the persistent action of bottom currents (e.g., Stow et al., 2002c; Rebesco, 2005). However, the term ‘‘bottom currents’’ is not usually employed by physical oceanographers. It is obvious that the time span of the processes involved in most geological processes is quite different from that analysed by physical oceanographers. In fact, the semi-quantitative inferences derived by geologists from the impact of ocean currents on the sea-floor sediments represent an integral over a poorly defined time interval. A detailed discussion about bottom currents from the point of view of physical oceanographers is given by Zenk (2008). However, we state as a general simplification that any ‘‘persistent’’ water current near the seafloor may be called a ‘‘bottom current’’.
M. Rebesco et al.
5
Such a current has the capability to affect the sea floor by re-suspending, transporting and/or controlling the deposition of sediments. This type of current is influenced by a number of processes (tides, internal waves, barotropic waves, dynamic instabilities) that modulate the speed and the instantaneous direction. Though typically affecting the sea floor, episodic flows that are not in equilibrium conditions do not belong to this category. Bottom currents do therefore not include turbidity currents, which are – in contrast to contour currents – rapid, dense (loaded with suspended sedimentary particles) currents driven by gravity, which fail to develop equilibrium conditions, even if they perhaps have more impact on the sea floor than bottom currents. For the same reason, the specific type of thermohaline circulation (THC) called ‘‘submarine overflows’’ (e.g., Denmark Strait Overflow Water) or ‘‘cascading currents’’, that are generated by transients in water-mass distribution in response to the local effect of evaporation, cooling or freezing in the surface layer over the continental shelf (Shapiro et al., 2003; Ivanov et al., 2004), should not be strictly considered as a bottom current. These currents are produced by the sinking – in relatively confined ‘‘flow pipes’’ – of important volumes of sea water, typically with high speed and intermittent flow; they occur every few years. The flow is controlled by morphological elements such as sills, submarine canyons and narrow straits. They are included among the engines of the THC at a basin-wide scale because the intermittent flow is – in the long term – compensated by the persistence of the process over very long time intervals, often coinciding with long-term climatic cycles. Known as purely physical oceanographic processes, these currents are now known to have an effect on the sea-floor morphology and to be able to transport important quantities of sediment in suspension (Canals et al., 2006; Trincardi et al., 2007). Future work should address the sedimentological importance of these currents. Shallow-water motion produced by surface waves, storms and tides (which are intermittent and not in equilibrium conditions), does not typically develop long-lasting bottom currents, though they affect the sea floor of continental shelves. Nevertheless, bottom currents can be affected by a number of distinct forces acting at different water depths. The term ‘‘bottom current’’ should be considered as a generic term that embraces different types of current. For clarity, we list here the (sometimes coexisting) types of current included within this term: • wind-driven currents, which originate by horizontal movement of the superficial layers due to wind shear stress, and subsequent propagation of the motion through the water column down to greater depth; • thermohaline currents, driven by gravity (uneven density distributions due to variable temperature and salt content of water masses), which are the most common type of bottom current; the predominantly horizontal large-scale transport of a water mass (advection) is called ‘‘THC’’ or ‘‘meridional overturning circulation’’ (MOC) in the Atlantic Ocean and determines the so-called oceanic conveyor belt; • geostrophic currents, which are characterized by long-lasting equilibrium conditions between the horizontal pressure gradient and the Coriolis force. By definition they have zero vertical velocity, which implies that the flow is forced
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Contourite Research: A Field in Full Development
to follow the bathymetric contours (therefore these currents are also named ‘‘contour currents’’); • contour currents, which are currents that have a net flow along-slope, sub-parallel to the topographic isobaths (contours); in spite of that, for certain parts of their path they can also flow upslope, down-slope, around and over topographic obstacles or irregularities; • boundary currents, which are currents the direction of which is controlled by sea-floor morphology (a wide canyon, continental slopes, but also flanks of major submarine mountain chains); due to the meridional distribution of continents and the Coriolis force, they are typically intensified along the western boundaries of the oceans (western boundary currents); • abyssal currents (sometimes also indicated as ‘‘ocean currents’’), which may be considered synonyms of large-scale bottom currents flowing in the deep sea (below the edge of the continental shelf) where topography and Coriolis force play a major role in determining the pathway of the current. Two more processes are important in the context of the processes related to contourite formation: • downwelling, which is the downward transport of cooled water masses from the surface; it occurs in restricted regions, mainly at polar latitudes, resulting in bottom currents; note that downwelling does not necessarily occur at the ocean margin, so it not necessarily affect the sea floor; • upwelling, which involves – in contrast – deep-water masses, and which produces bottom currents against topographic barriers and continental margins, mainly at low latitudes. Bottom currents can also be forced by deep-water tidal currents (in submarine canyons), long-wave baroclinic currents (internal waves and tides, solitary waves) and tsunami-related traction currents (Shanmugam, 2008). In conclusion, most (though not all) bottom currents start as density currents sinking to their equilibrium level (quasi-geostrophic balance). They are predominantly unidirectional subsurface currents in contact with the sea floor. They show a quasi-steady flow, though possibly affected by tides, seasonal changes and/or migrating eddies. They are controlled by topography, and when their direction largely parallels the topography, they are called ‘‘contour currents’’.
1.2.
C ONTOURITES
The term ‘‘contourite’’ was originally introduced to define the sediments deposited in the deep sea by contour-parallel thermohaline currents. Early pioneering work documented the strong influence of the deep Western Boundary Undercurrent along the continental margin of eastern North America (Heezen and Hollister, 1964; Heezen et al., 1966; Schneider et al., 1967). However, a rigorous restriction to this definition would prevent the application to ancient deposits, where both depth and direction of the currents can rarely be precisely reconstructed.
M. Rebesco et al.
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The definition was consequently widened to embrace a larger spectrum of sediments that are affected to various extent – and in a wide range of water depths – by different types of current. According to a now widely accepted recent definition, contourites are sediments deposited or significantly affected by bottom currents (Stow et al., 2002c; Rebesco, 2005; Stow and Fauge`res, 2008). Yet, the different types of bottom current are known to influence to a greater or lesser extent many depositional environments. They affect various types of sediment, both during and after deposition. This implies the risk of an excessively wide application of the term ‘‘contourite’’, and consequently of a loss of significance. On the contrary, where the dominant action of a bottom current is ascertained, the presence of other types of sediment is not excluded. Turbidites, for example, occur frequently even where the bottom-current influence is large enough to control the overall geometry of the deposits and to generate a (contourite) sedimentary drift. It is therefore not possible to restrict the use of the term ‘‘contourites’’ to the sediments contained in a sediment drift. In contrast, it is commonly accepted that contourite sedimentary facies also include sediments, usually interbedded, that are not deposited under the influence of bottom currents. This is especially true where the persistence of bottom-current action has not been enough to determine the geometry of the deposit. Another way to restrict the use of the term ‘‘contourites’’ is to set a minimum depth above which contourites cannot reliably be distinguished from shallow-water shelf-current deposits. This water depth is suggested to be around 300 m, according to the definition by Stow et al. (2008), even though this limit should not be applied rigidly. Shallow-water contourites may reflect also other hydrodynamic factors (shelf currents, tides and waves, storms) capable of impinging the sea floor, but the influence of which is still negligible compared to a dominant (though not absolutely steady) bottom current. As discussed in detail by Verdicchio and Trincardi (2008a), the processes involved in the formation of shallow-water contourites are generally more varied and less steady than in the case of deep-water contourite deposits. Shallow-water bottom currents derived from a stable geostrophic circulation can nevertheless form contourite drifts resembling the typical deep-water contourite drifts in morphology, internal geometry and sedimentary facies. Therefore, water depth alone does not seem to be an effective criterion to restrict the use of the term ‘‘contourites’’. As is often stressed in literature, the process of contourite deposition is not a simple one: it often involves multi-phase entrainment, long-distance transport, and interaction among depositional processes induced by the various types of bottom current described above (e.g., He et al., 2008; McCave, 2008; Stow et al., 2008). In addition, contourites are generally not easy to recognize because of the lack of simple, unambiguous diagnostic criteria. A composite triple-stage approach (e.g., Fauge`res et al., 1999; Rebesco and Stow, 2001; Nielsen et al., 2008) is recommended for the identification of sediments deposited by bottom currents as can be determined beyond reasonable doubt from seismic-reflection data: the analysis must include the overall architecture of the deposit (gross geometry and large-scale depositional units), the internal architecture (structure and sub-units) and seismic attributes and facies in each sub-unit.
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Contourite Research: A Field in Full Development
When samples are available, the problem of the lack of universally recognized unequivocal diagnostic criteria for the sedimentary characteristics of contourites must be faced. Though most workers suggest that pervasive bioturbation is the most diagnostic criterion (Stow and Fauge`res, 2008; Wetzel et al., 2008), others suggest a combination of traction structures (Martı´n-Chivelet et al., 2008; Shanmugam, 2008). Distinctive characteristics that are not diagnostic by themselves may be provided by early diagenesis (Giresse, 2008) and physical properties (Laberg and Camerlenghi, 2008). We therefore suggest that contourites as such should be identified on the basis of analysis of the characteristics of their facies and facies associations (cf. Stow and Fauge`res, 2008). Considering the above, we suggest that the well-established term ‘‘contourite’’ should be used as a generic term, in the same way as, for example, ‘‘mass-wasting deposits’’ or ‘‘gravity-flow deposits’’. These generic terms may be considered family names that simply describe a process that affects a certain kind of sediment: bottom currents in the case of contourites, down-slope mass transport of sediment and water for mass-wasting deposits. These family names include several kinds of sediment that have more specific names (e.g., turbidites, debris-flow deposits or mudflow deposits). In the case of contourites, such specific names do not exist; they should be defined in the time to come on the basis of research aimed at distinguishing between the various types of transport and depositional processes involved.
1.3.
D RIFTS
In sedimentology, ‘‘drift’’ is a general term used to describe ‘‘unconsolidated rock debris transported from one place and deposited to another’’ (Bates and Jackson, 1987). Its use is taken from one of the many definitions that can be found in a dictionary: ‘‘motion or action under external influence’’. Its use has traditionally been related to sediment transport by river currents, wind and glaciers. It is important to note that the ‘‘drifting’’ implied by the term does not refer to the deposit, but to the particles that move, transported by a flow, before settling. The first to use the term in association with bottom currents were Heezen and Johnson (1963), who stated that ‘‘scour and drift of sediments due to current activity are the most reasonable explanation of sediment moats’’ that were identified at that time next to seamounts and other striking morphological structures in the Atlantic and Pacific. The drift of sediment particles by the geostrophic circulation was later recognized as responsible not only for sediment moats, but also for sediment knolls (Heezen et al., 1966) as the drift could focus sedimentation in certain areas of the continental rise. The use of the term ‘‘drift’’ then shifted from that of a process for sediment transport by deep contour currents, to the end product of that process: the sedimentary deposit. The Blake–Bahama Outer Ridge was defined by Heezen and Hollister (1971) as a migratory sediment drift built by bottom currents. The sediments deposited by the action of bottom currents, whether or not forming a sediment drift, were termed ‘‘contourites’’ by the same authors. Because the flow responsible for the ‘‘drift’’ of particles from one place to another and the resulting sedimentary deposit is implicitly a steady one, mass-flow deposits resulting from surges of density currents were
M. Rebesco et al.
9
distinguished in principle from contourites since the very beginning (see the early definitions by Hollister and Heezen, 1971, p. 421). Bottom currents are capable of building thick and extensive accumulations of sediments not only on the continental rise, but also elsewhere along continental margins, from the abyssal floor to the outer shelf, if a significant sediment input is available (Fauge`res and Stow, 2008). These sediment bodies have received various names, including ‘‘outer ridges’’, ‘‘sedimentary ridges’’, ‘‘sedimentary mounds’’, ‘‘sediment drifts’’ and ‘‘contourite drifts’’. In our view, all these terms are synonyms, referring to accumulations of sediment deposited – or significantly affected – by bottom currents. Strictly speaking, the term ‘‘contourite drifts’’ should be specifically used for sediment accumulations deposited by currents flowing along the contours. However, we highlighted already above that different types of bottom current exist, and that even the so-called ‘‘contour currents’’ do not always follow the contours. An additional complication is provided by sedimentary accumulations for which the current direction is inferred only indirectly. We hence think that the term ‘‘contourite drifts’’ should be used for sediment accumulations deposited by bottom currents in general. Further, because of the presence of ‘‘contourite’’ in this term, we think that this term is to be preferred. In many cases, especially where mixed turbidite/contourite systems are involved, terms that not belong to the contourite terminology in a strict sense (e.g., mounds, levees, fans, lobes, channels) are used in the literature for deposits that have been significantly affected by the interaction with bottom currents. We recommend for such situations that these terms be preceded by the prefix ‘‘contourite’’ (e.g., contourite levees). Not only depositional structures are produced by bottom currents: there are a number of erosional and non-depositional structures. The classification and precise terminology of these structures that are related to bottom currents have not been fully developed yet; only some attempts have been made (Herna´ndez-Molina et al., 2008a, b).
1.4.
S EDIMENTARY STRUCTURES
What are the diagnostic sedimentary structures of contourites? The scientific community has not come yet to complete agreement. There are two viewpoints, with contrasting ways of reasoning: one is in favour of traction structures, whereas the other one is in favour of bioturbation. The initial finding in the 1950s of current ripples and other traction structures in deep-marine deposits affected by bottom currents (Hollister, 1967; Hollister and Heezen, 1967, 1972; Bouma, 1972a, b, 1973; Bouma and Hollister, 1973) suggested that contourites show lamination as a result of fluid-flow processes and depositional sorting mechanisms. Distinctly laminated deposits are still interpreted as contourites by some authors (see Hu¨neke and Stow, 2008; Martı´n-Chivelet, 2008; Shanmugam, 2008). These authors stress that traction structures are abundant on modern ocean floors that are influenced by bottom currents; they also interpret ancient coarsegrained deposits in some large outcrops as contourites. They wonder whether the general absence of such structures in the case studies based on sedimentary cores may
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Contourite Research: A Field in Full Development
result from a bias imposed by the impossibility of larger-scale observations. They also suggest that bioturbation in contourites should not be considered a diagnostic criterion, since bioturbated mud is equally abundant in areas unaffected by bottom currents and since turbidites may be extensively bioturbated as well. According to most contourite workers (e.g., Stow and Fauge`res, 2008), crosslamination is only rarely described from modern contourites, whereas extensive bioturbation is generally dominant (Wetzel et al., 2008). Their observations are derived from a huge number of data – published over the past 15 years – from modern drift systems using conventional coring techniques. They suggest that the lack of clear lamination in contourites is due to several reasons, including the commonly low current velocity that is insufficient to produce primary lamination, the relatively low accumulation rate allowing bioturbation to destroy primary lamination, and the small sediment input that is insufficient to allow a sorting mechanism to develop lamination. The controversy regarding diagnostic criteria (traction structures versus bioturbation) for the identification of contourites is still a problem. It was not possible to provide here a solution, considering the present state of the knowledge. We therefore purposely took care that both views are adequately expressed in this volume by inviting authors belonging to the two ‘‘schools’’. In our opinion, they honestly offered their own evidence, and we hope that the comprehensive scenario provided by this scientific, non-personal confrontation will provide a sound basis for a new step forward in the ongoing struggle to clarify this controversial issue.
1.5.
P ROSPECTS
More research is obviously needed to define a universally acceptable set of diagnostic criteria for contourites. Such research should be aimed at obtaining a much deeper insight into the processes involved. Simultaneously, the existing terminology for all aspects related to contour currents (deposits, processes, morphology, etc.) needs more consistency and logic, and new terminology has to be developed for aspects for which no real terminology exists as yet. This is a prerequisite for progress, because more precise terms and a stricter application of the terminology are needed for unambiguous interpretation of descriptions, as well as for better comprehension of the numerous processes involved. Improved terminology is, as mentioned before, particularly important to distinguish between the various types of contourite, and for classification of erosional structures that are related to bottom currents. The growing attention for these sediments – because they can be hydrocarbon seal rock and reservoirs, because they form a significant part of the palaeoceanograpic record, and because they form a source for potentially hazardous submarine landsides – will hopefully give rise to an ever increasing amount of high-quality data sets that will improve the understanding of these sediments, especially on the basis of their sedimentology and their seismic characteristics. An intensified access to industrial-quality hydrocarbon-exploration data in particular may play an essential role in an increased understanding of contourites, just like was the case in the 1960s for turbidites.
C H A P T E R
2
P ERSONAL R EMINISCENCES ON THE H ISTORY OF C ONTOURITES ¨ K.J. Hsu Oakcombe, Marley Common Haslemere, Surrey, UK
Contents 2.1. 2.2. 2.3. 2.4.
Introduction Not All Deep-Sea Sands are Turbidites Turbidophiles and Turbidophobes Erosional Unconformity Misinterpreted
2.1.
11 12 14 15
INTRODUCTION
I studied metamorphic petrology at the University of California (UCLA) in the early 1950s, and my ambition was to become a teacher of petrology at an American university. That was 10 years after the repeal of the Chinese Exclusion Act and there were not yet voices in the Congress on equal employment. I had little luck in finding a job. The nearest I came was a chance to become an acting lecturer at the Oregon State University as a 1-year substitute for somebody with sabbatical leave. However, my application was rejected. Eventually I found work for Shell Research, because it was a Dutch company. No American company would employ me because it would be impossible for Orientals to make business trips to the Deep South, where hotels existed only for the white or the coloured. I was not well prepared for my new job, not having taken any course on sedimentology. I asked for some advice, when I went to say goodbye to Jerry Winterer. Winterer was a classmate, but also he taught sedimentary petrography. I did not know him very well, because we, ‘‘hardrock guys’’, had our coffee room on the fourth floor, while the ‘‘softies’’ congregated on the fifth. Winterer was very pleased to learn that I was going to work for Shell. He just read an article in the AAPG Bulletin by someone there called Nanz, on the geometry of Oligocene sandstone reservoirs of the Seeligson Field in south Texas. It was brilliant, Winterer told me. I should look him up; perhaps he might teach me a thing or two on the art and science of basin analysis. I did not have to look Nanz up: I was assigned to him as a trainee. Taking into account of my experiences in California, Nanz asked me to make a study of the Developments in Sedimentology, Volume 60 ISSN 0070-4571, DOI: 10.1016/S0070-4571(08)00202-1
Ó 2008 Published by Elsevier B.V.
11
12
Personal Reminiscences on the History of Contourites
Pliocene sandstone reservoirs of the Ventura Field, California. There had been rumors that those might be turbidites. Nanz was ready to consider all scientific evidence, while most of his colleagues were downright hostile to the innovative suggestion. I went to Ventura in 1954, and the Pliocene sands are obviously turbidites, as can be deduced from their sedimentary structures. Applying Nanz’s technique of mapping the geometry of genetic units, I recognized on the electric logs a group of sand layers within a thick succession of shale. With easy correlation, I could map the distribution of this genetic unit. It had been thought, on the basis of the Grand Banks study by Heezen and Ewing (1952), that turbidites were blanket sands. I found, to the surprise of everyone, that the Ventura turbidites are shoestring bodies (Hsu¨, 1977). In hindsight, the discovery should not have been surprising. The event started from a point, and the loci of points constitute a line. Turbidites originated from a single locality, and should thus be linear sand bodies. My Shell report was internally distributed in 1954, but it was not published for more than two decades (Hsu¨, 1977), and even then only in part. Nevertheless, the impact in the oil industry was significant. After the Saticoy Field of the Ventura Basin had been discovered in 1955, the knowledge that the ‘‘turbidite reservoirs are string-sands’’ has become very helpful to production geologists. The recognition that linear sand bodies have up-dip pinch-out edges has contributed to the discovery of gas fields by the Occidental Company in the Sacramento Valley. I suspected that one of my former assistants at Ventura made the knowledge transfer; he received a generous bonus from his company after he changed employers.
2.2.
N OT ALL D EEP -SEA SANDS ARE T URBIDITES
‘‘Turbidite’’ soon became a ‘‘household word’’ in the Shell Oil Company, and the word was synonymous with ‘‘deep-sea sand’’. I did not like the word because of the genetic implication of deposition from the turbid waters of a suspension. When I was a post-doc in Switzerland, I found evidence in the Alpine flysch. The existence of groove casts as bottom markings clearly indicates that the flow depositing the laminated bottom layer (Bouma interval A) is not a turbidity current; it consists of debris in viscous motion (Hsu¨, 1959). Neither is graded bedding (also Bouma A) necessarily an evidence of settling from a suspension. The grading may have resulted from a decrease of the current-transport velocity (Kersey and Hsu¨, 1976). Finally, the cross-laminated silt (Bouma C), above the parallel laminated Bouma B interval, may or may not be a deposit from suspension. Years later my students and I did carried out experiments to show that the bedform of suspension deposits can be rippled (Hsu¨ et al., 1980). At that time, however, I was speculating on the possibility that the so-called Bouma C unit has been reworked by a bottom current. Bruce Heezen, who had acquired a fame because of his paper with Ewing on the turbidity-current deposition after the Grand Banks Event, came in 1955 to Houston to give a talk on deepsea sedimentation. I was his host. We soon found ourselves in agreement that not all deep-sea sands are turbidites. He showed me beautiful pictures of deep-sea ripples.
¨ K.J. Hsu
13
I could tell him, from my observations at Ventura, that those ripples are underlain by cross-laminated sand or silt. I had a field season, in the summer of 1955, to study the Ventura sediments. Together with Jim Valentine, who was collecting foraminifera samples for ecological studies, we climbed a steep slope covered by a dense growth of sage brush. Below a vertical cliff of a thick turbidite ledge was a mudstone deposit. Intercalated in the pelitic sediment is an ash unit which crops out everywhere in the Ventura Basin; it is the boundary ash between the Pliocene Pico Formation and the Pleistocene Santa Barbara Mudstone. The volcanic ash had settled on a deep-sea bottom, as Valentine identified the typical Uvigerina peregrina assemblage in the pelitic sediments. On a close inspection, we found that the tuff unit consists of two ash layers, with a rippled horizon between the two (Figure 2.1). Instead of a laterally continuously rippled layer, the ripples are ‘‘starved’’ micro-dune features underlain by a cross-laminated, very well-sorted coarse silt. That silt cannot be a turbidite, because the speed of a turbidity current would have disturbed or eroded away the underlying ash layer.
Figure 2.1 Deep-sea sedimentary facies. Cross-laminated silt deposited by bottom currents is found intercalated in the sediments of thin-bedded sand facies, which is found on the fringe of main turbidite sand bodies (Hsu« et al., 1980, reprinted with permission from AAPG, whose permission is required for further use).
14
Personal Reminiscences on the History of Contourites
2.3.
T URBIDOPHILES AND T URBIDOPHOBES
I wrote up the results of the Ventura Basin first as an internal report. I identified four lithofacies in the Pliocene of the Ventura Basin that represent four environments of deposition (Figure 2.2), namely: a mudstone facies deposited on the slope of the basin, a conglomerate facies in the canyons cutting across the slope, a turbidite sand facies in the deepest part of the basin trough, and a thin-bedded facies on the fringes of the turbidite sand. My interpretation of the genesis of the cross-laminated or thin-bedded sand represents little competitive value to Shell. The work was released for publication (Hsu¨, 1964) after I had presented the results at the 1963 SEPM meeting at Houston. My talk was attended by two persons, among others. Glenn Bartle was the President of the Harpur College; he came to the talk because I was being considered for a staff position at the College, or SUNY Binghamton. He did not make any comment of my interpretation, but he was very critical of my presentation. I did not hold the microphone steadily, and the variable volume of my voice was very irritating to the audience. The other person was Ken Emery, and he was furious. Emery, then teaching at the Southern California University, was a ‘‘turbidophile’’. He had convinced himself that all deep-sea sands are turbidites. He was angry because I was reinforcing the doubt of the ‘‘turbidophobes’’, and there were many of those in the oil industry. As he said, he had been working for
Tuff
Cross-laminated sands and silts
Shales
Figure 2.2 Cross-laminated silt between two ash falls.When I encountered the cross-laminated silt layer between the ash falls for the first time, I began to doubt if all deep-sea sands are turbidites (Hsu«, 1964, with permission from the Society for Sedimentary Geology).
15
¨ K.J. Hsu
(a)
(b)
Figure 2.3 Morphological similarity of cross-laminated silts (Hsu«, 1964, with permission from the Society for Sedimentary Geology). (a) Cross-laminated silt in the Pliocene, Pico Formation, California. (b) Cross-laminated silt in Holocene tidal-flat deposits of the Wadden Sea, the Netherlands.The same bedform can be present in two different environments.
more than a decade to educate the ignorant. And now I was giving them a new excuse. I was thus obstructing the progress of science, while driving a wedge between the academics and the industry. Ripples are ripples, and the same bedform can be found in different depositional environments. Deep-sea ripples are morphologically not distinguishable from ripples on tidal flats. I gave a comparison of the sedimentary structures of the deposits from the two very different environments (Figure 2.3). Emery was right, my after-dinner talk at the St Louis SEPM was a hit and gave much comfort to the ‘‘turbidophobes’’ of the oil industry. I was also right, not all deep-sea sands are turbidites. I learned then that Heezen was continuing to develop his idea on deep-sea sand deposition. Eventually, the rippled and cross-laminated sands and/or silts are called ‘‘contourites’’, because they have been deposited by deep-sea currents flowing parallel to submarine contours.
2.4.
EROSIONAL UNCONFORMITY M ISINTERPRETED
I left Shell in 1963. While I continued to be interested in sedimentology, I became almost a full-fledged geological oceanographer after I joined the JOIDES Deep Sea Drilling Project. As the Chairman of the South Atlantic Group of the Paleoenvironment Panel, I was an avid reader of reports and cruise proposals on deepsea circulations, and became acquainted with the latest investigations on the Antarctic Bottom Current (AABW) and the North Atlantic Deep Water (NADW). Particularly interesting were the results of drillings on the West African Margin. The power of the NADW caused deep erosion of the slope sediments. In places, Middle Miocene hemipelagic deposits of the South Atlantic overlie slope deposits as old as the Cretaceous.
16
Personal Reminiscences on the History of Contourites
When I was, as they called it, ‘‘on the beach’’, I taught Alpine tectonics at the Swiss Federal Institute of Technology. I encountered in the geology of the Alps a century-old puzzle that had been called ‘‘Wang Transgression’’. Unlike the common transgressive deposits of sand or gravel, the Wang is a Maastrichtian shale formation, containing a deep-sea fauna indicative of marine deposition on a Cretaceous slope. The relation was considered transgressive, because the Maastrichtian formation lies directly upon Campanian pelagic deposits in the North Helvetic palaeogeographic realm. The Wang sediments of the higher ultrahelvetic nappes were deposited farther offshore; they overlie unconformably progressive older (Santonian, Turonian, Cenomanian, Early Cretaceous, Late and Middle Jurassic) hemipelagic deposits (Figure 2.4). The orthodox interpretation postulates uplift, followed by subaerial erosion, followed by subsidence, and finally followed by a shallow-marine transgression. There is in fact no evidence of uplift, of subaerial erosion, of subsidence, or of a shallow-marine transgression. Using the modern West African Margin as an analogue, we see the obvious fact that the Wang ‘‘transgression’’ was not a transgression: it is an overlap of Maastrichtian contourites above older slope sediments. The unconformity under the Wang Formation signifies a period of very active submarine erosion by contour currents in the Helvetic realm of an Alpine Basin during the pre-Maastrichtian. This erosional progress is comparable to that on the West African Margin by the NADW during the pre-Middle Miocene. An American student, S. Diefenbach, completed a master thesis on the ‘‘Wang Transgression’’. She found clear evidence that the North Helvetic realm was a marine slope environment far south of the European continental coast. The gradient was steep enough to have caused widespread slumping of hemipelagic deposits, and the overshore slope was cut by submarine gullies or canyons in which coarse clastics were accumulated. Helvetic
Ultrahelvetic Internal Prealps Plaine-Morte Decke
Wildhorn Decke
Wang Leimer Habkern Schurfling Schuppen Wildflysch
External Prealps
Tothorn Laubhorn Decke Decke Leissigen Schlieren Flysch Flysch Habkern
Gurnigel Flysch
swell
Sea level
Priabonian
on
Luteti
SH Swell
Wang beds Amdener Schichten Turonian
n
Wang
mania
c
Malm Dogger nian rias Aale T
granite
oi oz
Lowe
o Pale beds
Lu tet ian Yp re Pa sion leo ce ne es
us
ceo r Creta
Habkern e cen
M
Ceno
n
esio
Ypr
Figure 2.4 The Wang ‘‘Transgression’’ (Hsu«, 1960, with permission from the Geological Society of America). A reconstruction of the paleogeographic relations suggests that the MaastrichtianWang Formation consists of transgressive deposits overlying sub-aerially eroded older formations. This classic interpretation is wrong. The Wang Formation consists of contourite beds.
¨ K.J. Hsu
17
The Wang is clearly a contourite formation. The interpretation was not published until my book The Geology of Switzerland was printed after my retirement (Hsu¨, 1995). Meanwhile, I committed an indiscretion when I taught the idea in my class on The Geology of Switzerland. For that, students and colleagues alike chastised me. I was told that I should not have taught Ken Hsu¨’s crazy ideas to beginning students. I often wonder, if Isaac Newton was similarly reproached when he taught gravity to his Cambridge students before the publication of the Principia. If the academic establishment of the 17th-century England had been as dogmatic as the Swiss geological community of the 20th century, Newton would probably have had to tell his students, contrary to his conviction, that the Sun went around the Earth. I recalled that one of my students had to delete the word ‘‘me´lange’’ when he intended to publish, in the early 1970s, his thesis on the Wildflysch in the Eclogae Geologicae Helvetiae. Now that a book on contourites is published, I am hopeful that the editors of that illustrious journal would permit the use of the word ‘‘contourite’’ in articles on Alpine palaeoceanography.
C H A P T E R
3
M ETHODS FOR C ONTOURITE R ESEARCH J.A. Howe Scottish Association for Marine Science & UHI Millennium Institute-Dunstaffnage Marine Laboratory, Oban, Argyll, Scotland, UK
Contents 3.1. Introduction 3.2. Oceanographic Measurements 3.3. Geophysical Methods 3.4. Sampling Strategies 3.5. Analytical Methods 3.6. Onshore Studies of Ancient Sequences 3.7. Summary of Multidisciplinary Techniques Acknowledgements
3.1.
19 20 21 27 29 31 31 33
INTRODUCTION
The physical effect of a persistent bottom current on deep-sea sediment can be studied using a wide variety of oceanographic, geophysical and sedimentological techniques. Contourite workers need to be able to study both the modern and palaeo-deep-sea floor. A definition of the term ‘‘contourite’’ is provided by Rebesco et al. (2008); this chapter aims to present a summary of the range of techniques that can be employed to investigate contourite sedimentation in the deep sea. The range of processes contributing to current-influenced sedimentation in the deep sea can be extremely diverse, and at a variety of scales from the localised to the global, such as biogenic fluxes from the water column, changes in ocean chemistry, the localised dynamics of the benthic boundary layer and the variation of thermohaline flow in response to bathymetry. Since Wu¨st (1936) first proposed the idea of thermohaline flow in the deep oceans, bottom current and perhaps more specifically, contourite researchers have needed to take a broad, multi-disciplinary approach when examining the deep sea bed for evidence of contourites. Early workers have, understandably, been restricted by the variety of techniques available to them, and used standard approaches such as locating rippled sands at abyssal depths in deep-sea photographs (Hollister and Elder, 1969; Heezen and Hollister, Developments in Sedimentology, Volume 60 ISSN 0070-4571, DOI: 10.1016/S0070-4571(08)00203-3
2008 Elsevier B.V. All rights reserved.
19
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Methods for Contourite Research
1971) or identifying regions of enhanced sedimentation on a margin from seismicreflection profiles (Hollister et al., 1978) as evidence for persistent along-slope bottom currents. Modern contourite researchers have considerably more in their investigative arsenal, including multi-beam bathymetry, side-scan sonar, high-resolution seismicreflection profiles, as well as the ability of recovering longer, less disturbed samples from drifts as coring technologies have developed. Progress has not simply been restricted to the collection of data from deep-water basins and margins: once ashore, core samples are increasingly subjected to ever more detailed sedimentological and geochemical examination. This drive for very detailed, high-resolution core data has resulted in some spectacular discoveries that have considerably increased our knowledge of how significant global thermohaline flow is to climate and, more importantly, how changes in global thermohaline flow can influence climate change. Most, if not all, of these discoveries have arisen from detailed multidisciplinary work on regions of contourite sedimentation in the deep sea.
3.2.
O CEANOGRAPHIC MEASUREMENTS
Oceanographic measurements of modern, persistent thermohaline or winddriven flow in the deep sea are the most fundamental data needed when beginning to investigate bottom-current deposits (see also Salon et al., 2008 and Zenk, 2008). For reasons of cost, only major drilling and coring projects can undertake to obtain new, detailed current-meter data over drift sites. Ideally, two cruises are needed, one to deploy the current meters and one to recover them, usually after enough time has elapsed to measure the hydrographic conditions at the site (e.g. 6–12 months later). Typically, contourite workers have made use of pre-existing hydrographic data sets, which are not always ideally located on major drift sites and collected for other purposes. One of the earliest and certainly one of the most important multidisciplinary contourite studies showed how the deep sea can experience high-energy bottom-current events or ‘‘benthic storms’’. These storms are characterised by short-duration flow reversal and enhancement, produced in regions of high eddy kinetic energy. The HEBBLE (High Energy Benthic Boundary Layer Experiment) site (Nowell et al., 1982; McCave et al., 2002) on the Nova Scotia Rise identified these events through the combined use of current measurements, sea-floor photography as well as high-resolution seismics and coring. Studies such as the Project Mudwaves in the Argentine Basin (Manley and Flood, 1993a; Flood et al., 1993) also used direct physical measurements from current meters placed across a wave crest, flank and trough to deduce the bottomcurrent velocities and structure across the features. Measurements such as these can lead to mathematical modelling of bottom-current flow across waves (Hopfauf et al., 2001). Camerlenghi et al. (1997) used two current-meter moorings across a giant sediment drift west of the Antarctic Peninsula to ascertain present-day bottom-current conditions. In an extensive study from
J.A. Howe
21
the Campos Basin of the southwestern Brazilian Margin, Viana et al. (1998a) utilised current-meter moorings, wave measurements and infrared satellite images to investigate the origin of contouritic sand deposits on the margin. Hydrological processes considered in this study included the influence of surface currents, counter-currents, waves, tide and eddies resulting in the offshore transport of contourite sand by a combination of shelf spillover and internal waves. The direct physical measurement of bottom currents is of vital importance in relating the modern benthic conditions to ancestral bottom-current flows, inferred from core-sediment texture. Another significant activity that can be carried out in the water column above sites of bottom-current influence is the collection of sediment from the water column, including from the nepheloid layer. Such data are unusual but very helpful. It is notable from high-latitude studies where significant sediment supply can originate from the surface through iceberg rain-out. The presence of a nepheloid layer can be a useful indicator of near-sea-floor current activity. These layers of suspended sediment can be several hundred metres thick and consist of material derived from both primary productivity from the surface and in situ resuspension. These measurements can provide information on the sediment flux to a site, either by along-slope transport or directly from biogenic productivity from the overlying water masses. A simple sediment trap on a current-meter mooring can collect sediments for a prolonged period of time, typically months. The sediment record can show seasonal variations not usually detectable in core section. Some sediment traps utilise a sequential carousal of collecting bottles, permitting samples to be obtained over a seasonal cycle, such as pre-planktonic bloom, bloom and postbloom. Water samples collected during conductivity, temperature and depth (CTD) casts can reveal the nature, and – if combined with transmissometer data – the extent of suspended sediment in the nepheloid layer. Pudsey (1992) and later Gilbert et al. (1998) used transmissivity profiles collected during CTD transects in the northwestern Weddell Sea to identify an active nepheloid layer present on the rise, the product of the along-slope transport of Weddell Sea Bottom Water. In the North Atlantic, across the Bjorn and Gardar Drifts, Bianchi and McCave (2000) collected water samples to examine suspended particulate material and transmissometer readings to determine the thickness of the benthic nepheloid layer. These data were combined with pre-existing current-meter data from the region and used to calibrate the palaeo-records of bottom-current velocities from the cores, resulting in a very detailed record of the flow of Iceland–Scotland Overflow Water moving as a Deep Western Boundary Current in the Iceland Basin across this major drift complex.
3.3.
GEOPHYSICAL METHODS
In any multidisciplinary study of bottom-current-influenced sedimentation in the deep sea, geophysical – specifically acoustic – surveys play a key role (see also Nielsen et al., 2008). For simplicity, acoustic surveying is here divided into
22
Methods for Contourite Research
two main techniques: sea-floor mapping (for surface expression and morphology) and seismic-reflection profiling (for internal acoustic character, geometry and hence interpretations of the depositional environment). Some inferences can also be made on the relative ages of the sediments (seismostratigraphy). It could be argued that it is only through the use of seismic-reflection profiling and echo sounders surveys that the basin-scale extent and nature of bottomcurrent sedimentation has been recognised (Johnson and Schneider, 1969; Hollister et al., 1978). A wide range of acoustic tools are now available that can penetrate the sediment cover, depending on the frequency used. Frequencies around 12 kHz and above provide little sediment penetration and are therefore used for surface mapping, 3.5 kHz profiling systems can provide penetration up to 100 m below sea floor and are suited to detailed process-related studies, and, finally, 10–200 Hz airgun seismic sources give very deep penetration (km) and are much more suited to deep geological studies of ancient drifts and basement. Even at its most basic, sea-floor mapping purely for morphology can provide one of the most graphic illustrations of the dynamic nature of bottom-current systems available. Through detailed bathymetric surveying, it is possible to acoustically image the sea floor where sediments have become moulded into drifts and waves or sculpted into moats around obstructions. Even with basic side-scan sonar instruments, some information on sediment morphology, grain size and sea-floor roughness can be obtained (Figure 3.1). More advanced, towed side-scan systems such as TOBI (Towed Ocean Bottom Instrument) and GLORIA (Geological Long-Range Inclined Asdic) have provided stunning acoustic images of contourite systems from most of the major ocean basins. The modern development of multibeam echo sounders has revolutionised mapping of the sea floor. These systems provide wide coverage (typically in a water depth of 3000 m the ‘‘footprint’’ could be as wide as 12 km). Regional sediment processes can be identified from the morphology of the sea floor. At their best, multibeam systems can easily be used to identify drifts, moats and waves characteristic of bottom-current-influenced sedimentation (Figure 3.2a–c). Using an echo sounder can provide a wealth of information on sea-floor morphology, acoustic character and even some limited sediment thicknesses (isopach maps) in the surface sediments. Sea-floor mapping using 3.5 kHz echo sounders or, more recently, with modern parametric sub-bottom profilers (such as ‘‘TOPAS’’ and ‘‘Parasound’’) has the added advantage of being able to cover large areas of sea floor quickly and cheaply (Figure 3.3a and b). This technique has been used on numerous occasions in contourite studies. Initially pioneered by Damuth (1978, 1980), this technique classifies the echo type based on sea-floor microtopography, sub-bottom penetration and reflectors. The use of 3.5 kHz survey data has extended the range and scope of contourite study, enabling workers to occasionally observe a basin-wide, more regional picture of sedimentation. Examples of this technique can be found in Pudsey and Howe (1998), where a simple surface acoustic-character map was produced for the Central Scotia Sea using data sets from numerous cruises from the mid-1980s to the present day.
J.A. Howe
23
Water column
Ship track/ transmit pulse
First bottom return
10 m
N
Figure 3.1 Side-scan sonar image of bottom-current-reworked sands and outcropping rocks from approximately 500 m water depth southeast of the Falkland Islands, South Atlantic (Image from the British Antarctic Survey cruise 29 on RRS James Clark Ross).
Seismic-reflection profiling (Figures 3.4a, b and 3.5) has provided the basis for a number of classification schemes on the various types of drift morphology and internal configuration (McCave and Tucholke, 1986; Fauge`res et al., 1999; Rebesco and Stow, 2001). Regional seismic-reflection profiles can also be used to ‘‘backstrip’’ or map laterally continuous reflectors or mega sequences and produce models of palaeo-bathymetry (and hence palaeo-current pathways) and sediment thicknesses. Wold (1994) used this technique highly effectively producing models of North Atlantic Drift development and hence sediment fluxes during the Cenozoic. This technique is typically underused in contourite research, due to the high density of seismic profiles needed across a drift system. In addition, the reflectors within the modelled drifts need to be related to some direct age constraint, typically from a drill site or long piston core.
24
Methods for Contourite Research
(a)
Minor drifts Sediment drift
Wave field
20 km
Wave field Parasitic cones
Terraces
Concave slide scar
Moat Water depth (m)
(b)
800.00
1500.00
Sediment drift
5 km
2000.
Slide scar
Sediment waves
Parasitic cones
Moat
Terrace Slide scar
Slide Scar
(c)
63°N
20 km Sediment wave field
60°N
N
57°N
54°N
51°N
Sediment drift Multiple moat Sediment wave field
For caption see next page.
25°W
Moat –2300 m
20°W
15°W
10°W
5°W
0°W
J.A. Howe
25
Figure 3.2 Multibeam bathymetry of the Rosemary Bank Seamount, North Atlantic. (a) Shaded relief (from the northeast) and colour contoured multibeam bathymetry (Kongsberg EM120 12 kHz, 1° 1° beams, collected during cruise 99 of the RRS James Clark Ross) of the Rosemary Bank Seamount, North Atlantic Ocean. (b) Detailed view of the western spur region of Rosemary Bank showing the sediment wave-field, multiple moat and drift complex. Also indicated is a concave slide scar on the southern flank. (c) Perspective view of the seamount viewed from the northwest. Inset map shows the location of the seamount in the northern Rockall Trough (from Howe et al., 2006, with permission from Elsevier). A multicolour version of this figure is on the enclosed CD-ROM.
(a)
(b)
Figure 3.3 TOPAS sub-bottom profiles from the Fram Strait. (a) TOPAS sub-bottom profile from the western Svalbard Margin hemipelagic and contouritic sediments interrupted by debris flows and diapirs from the Storfjorden Trough Mouth Fan. (b) Small, current-controlled sediment drift developed on the Vestnesa Ridge, western Svalbard Margin, a region associated with abundant gas-escape-related pockmarks (adapted from Howe et al., 2008). Inset map shows the location of the figures on the western Svalbard Margin. (Profiles collected on the Scottish Association for Marine Science Arctic cruise 75 on RRS James Clark Ross.) A multicolour version of this figure is on the enclosed CD-ROM.
26
Methods for Contourite Research
SW
(a)
NE Northern Rockall Trough
Distal edge of Sula Sgeir Fan
Subsidiary drift
Elongate drift
Wyville-Thomson Ridge
Moat-related drift
0.8
Plio–pleistocene 1.0 Miocene 1.2 Palaeogene BGS 83/04-60
1.5 km NE Hebrides Slope
Multi-crested sediment drift
(b)
Elongate drift
SW
Northern Rockall Trough
TWT (s)
0.8 TWT (s)
Elongate drift
Elongate drift
1.0
1.2
BGS 84/06-24
Plio–pleistocene
1.4
Miocene
1.5 km
Palaeogene
Figure 3.4 Seismic-reflection profiles from the northern Rockall Trough, North Atlantic. (a) British Geological Survey single-channel airgun seismic-reflection profile 83/04 -60 illustrating the development of elongate, subsidiary and moat-related drifts adjacent to the Wyville-Thomson Ridge, northern Rockall Trough (adapted from Howe et al., 1994). (b) British Geological Survey sparker seismic-reflection profile 84/06 -24 showing a complex multicrested elongate sediment drift developed parallel to the Hebrides Slope, northern Rockall Trough (adapted from Howe et al., 2002).
TWT Northern s 0.5 NW 1.0 1.5 2.0 2.5 3.0 3.5
Rockall Trough
(b)
Rosemary Bank
Northern Rockall Trough SE
(c)
SBM SBM
(d)
SBM
(a) BGS 03/03/1
s 0.5 1.0 1.5 2.0 2.5 3.0 3.5
5 km
Figure 3.5 Major seismostratigraphic units occurring in the drift sequences surrounding the Rosemary Bank Seamount. Megasequences: RPa = Pliocene ^ Holocene; RPb = midMiocene to early Pliocene; RPc = late Eocene to early Miocene. Reflectors: C10 = early Pliocene angular unconformity; TPu = late Oligocene to mid-Miocene; C20 = late-early to early-mid-Miocene; C30 = late Eocene unconformity (after Stoker et al., 2001). TL =Top of Late Cretaceous ^ early Paleogene lavas; SBM = sea bed multiple (adapted from Howe et al., 2006). The location of the Rosemary Bank Seamount is shown in Figure 3.2d (from Howe et al., 2006, with permission from Elsevier). (a) British Geological Survey seismic-reflection profile 03/03/1 (with locations of insets b ^ d). The profile displays the distinctive moat-drift association at the base of the seamount produced by enhanced bottom-current flow.
J.A. Howe
27
(b)
Rosemary Bank Seamount
TL
Elongate drift Cut-and-fill
Onlap RPa
Drape of post-C10
RPb
Moat
0.5 s C10
?C20 RPc C30
3 km Eocene
Neogene drift
(c)
Eocene prograding wedge with Neogene veneer
Rosemary Bank Seamount
TL
0.5 s
3 km SBM
TL
? debris flow uncertain age TL
Rosemary Bank Seamount
Elongate drift RPa C10
Moat
(d)
RPb TPU
0.5 s
RPc C30 3 km TL
Eocene
Figure 3.5 (Continued) (b) Detail of drift ^ moat association showing the main mid-Miocene ^ Pliocene age construction of the drifts. (c) Eocene age prograding wedge overlain by Neogene drift sediments on the flanks of the seamount. (d) Deeply incised moat at the base of the seamount produced by sustained, enhance bottom-current flow.
3.4.
SAMPLING STRATEGIES
In most multi-disciplinary studies of contourite systems, after the oceanographic and geophysical surveys, it will usually be essential to sample the sea bed. The collection of any undisturbed contourite sediment samples from the sea bed presents no more problems than in any other deep-sea sediment facies. It is worth noting, however, the
28
Methods for Contourite Research
local variation in facies across drift complexes, not always appreciated from conventional geophysical data. To examine small-scale depositional changes, the use of sea-floor photography or video transects is very useful (Figure 3.6a–f ). The use of sea-floor photography is not a new technique: indeed, it was one of the first, and most widely used to examine the deep sea. As described earlier, workers such as Heezen and Hollister (a)
(b)
(c)
(d)
(e)
(f)
Figure 3.6 Sea-floor photographs showing contouritic sediments from the North Atlantic and Arctic Oceans (all images courtesyof David Hughes, the Scottish Association for Marine Science). (a) Sea-floor photograph of inclined glass sponges bending under the influence of Norwegian Sea Deep Water flowing west along the Faeroe ^Shetland Channel, north of the Wyville-Thomson Ridge, 1100 m. (b) Rippled, bioturbated contourite sandy muds in the northern Rockall Trough, south of theWyville-Thomson Ridge,1050 m. (c) Glaciomarine hemipelagites under seasonal sea ice on the Yermak Plateau, northern Fram Strait, 803 m. (d) Rippled contourite silty sands from 700 m on the Hebrides Slope, northern Rockall Trough. (e) Gravelly and sandy contourites from 1000 m on the Hebrides Slope, northern RockallTrough. (f ) Rippled silty/sandy contourites with emerging clasts from 1000 m on the Hebrides Slope, northern Rockall Trough. A multicolour version of this figure is on the enclosed CD-ROM.
J.A. Howe
29
(1971) and Heezen et al. (1966) made extensive collections of sea-floor photographs from the ocean basins and showed, visually, that the deep-sea floor was not a wholly tranquil place but was swept by persistent bottom currents. Photographs across drift sites were obtained by McCave (1982), Scha¨fer and Asprey (1982) and Carter and Schafer (1983) from the Orphan Knoll, Labrador Slope. These showed that rippled sands and silts are predominant where current velocities are highest (e.g. at drift margins and moats), with mud deposited marginally to the axis of the flow (such as the drift crest), in this case the Western Boundary Undercurrent. A wide variety of sampling, coring and drilling devices are available to collect sea-floor sediments, but not all of these are suited to contourite studies. Simple grabs tend to provide a homogenised sample, destroying the fabrics and structures that are invaluable to contourite interpretation. Short corers such as multiple and mega corers as well as Sholkovitch, Craib and box corers are useful in that they provide undisturbed samples from the sediment/water interface, but – being short (under 0.5 m) – can lack the temporal range needed to examine any change in the bottom-current system. Longer soft-sediment corers such as Kasten, gravity, vibrocore and piston corers are much more useful in providing longer records from drift sites, and are also heavy enough to sample some sands that may be present. Very long coring systems are now available, such as the spectacular Giant Piston Corer, the Advanced Piston Corer and the French STACOR (stationary piston corer) and Calypso corer, which can routinely collect undisturbed core samples of over 30 m. Problems associated with these corers have been highlighted by Skinner and McCave (2003) as including ‘‘over-sampling’’ (due to cable rebound) and ‘‘under-sampling’’, whereby the basal or mid-core sections are deformed or lost. These problems aside, the techniques of piston and gravity coring have provided a wealth of short-term, high-resolution palaeo-oceanographic records on bottom-current variability from drift sites across the world. Longer drilling projects have been, thus far, the domain of the original Deep Sea Drilling Project (DSDP), Ocean Drilling Program (ODP) and its present incarnation, the Integrated Ocean Drilling Program (IODP). This program has routinely drilled drift sites over the past few decades for the geological record of bottomcurrent activity with an age range of Quaternary–Cretaceous and extending from the high-latitude Arctic Ocean and Southern Ocean to the equatorial Atlantic: see Stow et al. (1998a) and Rebesco (2005) for a review.
3.5.
ANALYTICAL M ETHODS
Contourite facies collected in core samples are recognised by the criteria of structure, texture, fabric, composition and their sequence arrangement (see also Stow and Fauge`res, 2008). Extensive work has been conducted on contourite facies, notably by Stow (1979), Fauge`res et al. (1984) and Gonthier et al. (1984); it is summarised in Stow et al. (2002c). A wide variety of contourite facies have been recognised, from the initial classification of muddy, silty–muddy and sandy proposed by Stow and Holbrook (1984). Contouritic gravels, volcaniclastic, calcareous, siliceous and even manganiferous facies have now been described (Stow et al., 1998a). Identifying these facies in core samples can be difficult and workers have gone
30
Methods for Contourite Research
to considerable lengths to identify useful criteria. Meticulous core logging is necessary, with particular care needed when describing lithological contacts, grain size, structure, level of bioturbation and sequence relationships. The ability to differentiate between down-slope turbidite and along-slope contourite deposition was considered by Stow (1979) and Stow and Lovell (1979). A multi-disciplinary approach is required with a combination of morphological evidence (i.e. are there features on the sea floor indicative of a depositional pathway, such as furrows, moats and waves, visible on acoustic data?) and sedimentological information (sediment fabric and structure). Typically, contourites result from persistent or semi-persistent flows as opposed to the catastrophic down-slope movement of turbidites. Contourites therefore might posses irregular grading and sharp contacts, heavy-mineral fabric and lack the reworked shallow-water microfossils of turbidites. Interbedded turbidites and contourites as well as contourites derived from the reworking of turbidites remain difficult to distinguish, although combined sedimentological and morphological (including sub-bottom profiles) approaches should provide evidence of local scale (sedimentological) and regional scale (morphological) along-slope pathways. Flood et al. (1985) examined the magnetic fabrics of sediments from the HEBBLE site to determine along-slope fabrics in muddy contourites. Other techniques useful for contourite recognition include X-radiography of cores for internal structure, ice-rafted debris and bioturbation (Figure 3.7a–c), grain-size analysis and an examination of any microfaunal assemblages. (a)
(b)
Silty lamination
(c) D/St
Glaciogenic contourites
Silty lamination Sandy turbidites
Sandy–gravelly contourites
Glaciogenic contourites X-radiograph positive BGS core 59-08/41 Upper Hebrides Slope 470 m 0.35–0.70 m
X-radiograph negative SAMS core 067 From Strait Western Svalbard Margin Fram Strait 1226 m 1.70–2.00 m
X-radiograph positive BGS core 58-14/34 Northwestern Rockall Trough 1486 m 2.80–3.00 m
Figure 3.7 X-radiographs of contourite facies. (a) X-radiograph positive image of silty lamination in bioturbated muddy contourites from the upper Hebrides Slope (British Geological Survey core 59-08/41). (b) Negative X-radiograph image of sandy turbidites interbedded with glaciogenic contourites from the western Svalbard Margin, Fram Strait. D/St indicates dropstone clast (Scottish Association for Marine Science core 067). (c) Positive X-radiograph image of sandy/ gravellycontourites fromthe northern RockallTrough (British Geological Surveycore 58-14/38).
J.A. Howe
31
Recent developments include the use of non-destructive, pass-through core-logging systems; using these, a great deal of information can be obtained whilst at sea, such as magnetic susceptibility and other physical properties such as bulk density, acoustic velocity and natural gamma-ray emissions. More elaborate onshore logging systems use X-ray fluorescence to determine bulk geochemistry, which can be useful in examining changing sediment sources and bottom-water chemistry. The most useful sedimentological technique is grain-size analysis. The use of grain size, and its link to bottom-current flow was first demonstrated by Ellwood and Ledbetter (1977) from the Vema Channel, South Atlantic. This study demonstrated the direct relationship between subtle changes in silt grain size and bottomcurrent velocities (in this case, Antarctic Bottom Water). McCave et al. (1995a) took this technique a step further, identifying the ‘‘sortable silt’’ interval; this is the mean of the terrigenous silts between 10 and 63 mm, which are most susceptible to transport by bottom currents. Through a meticulous analysis of this sediment fraction, subtle bottom-current variations can be inferred. A number of important palaeo-oceanographic studies have been undertaken using this technique and have revealed the wider, climatic signals preserved in fine-grained contouritic sediments. Geochemical proxies are becoming more widely used, notably the use of rare-earth elements and isotopes, particularly Nd with Sr and Pb, when combined with grain size as a tracer of palaeo-bottom-current velocities (Revel et al., 1996). Other work has used stable isotopes such as 234Th, 210Pb and 210Po for particle cycling (Murray et al., 2005) and 18O and 13C to develop a detailed stratigraphy of bottom-current variations (Rasmussen et al., 2002).
3.6.
ONSHORE STUDIES OF ANCIENT SEQUENCES
Studies from outcrop as well as from boreholes of fossil contourite sequences are much rarer than the offshore, modern counterparts (see also Hu¨neke and Stow, 2008). Stow et al. (1998a, 2002c) summarise one of the main problems as being confusion with fine-grained turbidites. Many of the claimed contourite cases originate from reworked turbidite successions. The recognition of fossil contourites uses a three-stage approach: (1) small-scale studies (from field, laboratory or borehole evidence); (2) regional (drift, formation and region) trends in facies; and (3) fitting the conclusions from (1) and (2) with any independent palaeooceanographic reconstruction of the region showing a deep-water, persistent bottom-current pathway in the geological past.
3.7.
S UMMARY OF MULTIDISCIPLINARY TECHNIQUES
Contouritic sedimentation remains somewhat enigmatic, although it is becoming clear that bottom-current-influenced sedimentation is a highly significant process in the deep ocean. Outlined in this chapter are a number of approaches to investigating along-slope processes, at a number of scales, using hydrographic,
32
Methods for Contourite Research
geophysical and sedimentological and outcrop studies (Table 3.1). Developments continue, especially in the areas of geophysics and palaeo-oceanography using contourites. New high-resolution bathymetric surveys are now being routinely collected from drift and wave sites around the world and new proxies for palaeo-ocean chemistry and bottom-current pathways investigated. Deep-ocean technologies and instruments are being developed which are smaller and cheaper, and hence easier to use on research cruises. Satellite data are now more widely available, showing the distribution of surface currents and sea-surface heights, enabling regions of enhanced bottom-current activity to be better understood. Table 3.1 The recognition of contouritic sedimentation in the deep sea and investigative techniques used Scale Small-scale: field, borehole and laboratory • Non-turbiditic characteristics or origin • Mixed pelagite/hemipelagite setting with strong evidence for bottom-current sedimentation • Any cyclicity is related to bottomcurrent velocities, not to terrigenous input or biogenic productivity
Methods
• • • • • • • •
Medium-scale: drift, formation or region • Regional trends in facies, and current/ palaeo-current directions, texture, mineraological or geochemical indicators • Unconformities, condensed sequences, regional variation in thickness, drift geometry present • Shape and geometry of the sediment body, indicating an along-slope trend • Contouritic environment/palaeo-environment, including accumulation rates and in situ faunal assemblages Large-scale: system, ocean or continental margin • A bottom-current-influenced system from environmental/palaeo-environmental evidence • Modern and ancestral bottom-current system
Source: Adapted from Stow et al. (2002c).
• • • •
Detailed core logging Facies analysis Gravity, kasten or box coring Physical properties Grain-size analysis Extensive site survey data (multi-beam bathymetry, sub-bottom profiles, sea bed photography, side-scan sonar) X-radiographs hydrographic survey (current moorings, CTD) Seismic-reflection profiling Multibeam bathymetry Sea bed video/photography Onshore section logging/wireline logging
• Hydrographic surveys • Regional multi-beam bathymetric mapping
• Satellite imagery • Seismic-reflection profiling • Drilling/long-piston-coring investigations
J.A. Howe
33
As techniques for investigating contourites are continuing to be developed, this forward progress is given the added impetus of the knowledge that contourites may, in the future, provide significant discoveries in the fields of petroleum exploration, thermohaline circulation and climate studies.
ACKNOWLEDGEMENTS This review could not have been completed without the invaluable help of Peter Morris of the British Antarctic Survey for advice on the geophysics and multi-beam, and Martyn Stoker of the British Geological Survey for allowing access to the Rockall Trough seismic-reflection profiles. The text was greatly improved by the critical reviews of Rob Larter of the British Antarctic Survey and David Long of the British Geological Survey.
P A R T
2
BOTTOM CURRENTS
C H A P T E R
4
A BYSSAL AND C ONTOUR C URRENTS W. Zenk Leibniz Institute of Marine Sciences at the University of Kiel (IFM-GEOMAR), Kiel, Germany
Contents 4.1. Introduction 4.2. Abyssal Currents in the Global Thermohaline Circulation 4.3. Contour Currents 4.3.1. Physical properties and modes 4.3.2. Entrainment and modeled contour currents 4.4. Conclusion and Summary Acknowledgments
4.1.
37 40 42 42 51 55 57
INTRODUCTION
In marine geography, the large temperature difference at low latitudes of over 20°C between the sea surface and the bottom of the ocean remained a mystery for more than two centuries (cf. Schiermeier, 2006). Although in the tropics the ocean’s surface is exposed to its strongest solar irradiance, temperatures exceeding 20°C are found only in a thin top layer of some tens of meters. Roughly in the upper kilometer, temperatures decrease to about 5°C. This uppermost part of the oceanic water column is called the thermocline or, occasionally, the warm water sphere. Its well-defined upper interface is in permanent exchange with the atmospheric boundary layer, enabling free exchange of heat, mass, and energy. In contrast to the warm water sphere, the cold water sphere below is characterized by minimal vertical temperature gradients. At depths >2000 m, temperatures are usually 1 m s1) are rare. They may occur at the surface as well as at depth. In a series of generalized bottom sketches (Figure 4.1), Heezen and Hollister (1971) demonstrate the impact of increasing abyssal currents on the underwater seascape. It reaches from marine organisms gently deflected by a few millimeters per second current to bare rocks with pockets of gravel caused by very strong bottom currents ( >1 m s1). We note that such quantitative descriptions of the strength of ocean currents represent an integral over a rather unspecified time interval. Wind-driven currents at the surface are subject to varying weather and climatic conditions. At depth, however, the three-dimensional shape of bottom contours is one of the most prominent factors that influence the direction and the strength of internal currents. They can be forced by long-wave motion (tides, internal waves, seiches) or by horizontal density differences (gradient currents). Topographic control of oceanic currents in deep passages and straits defines the abyssal motion in substantial parts of the world ocean. Whitehead (1998) distinguishes between unidirectional currents over saddle points between neighboring deep ocean basins, depending on the source of the densest bottom water, and bidirectional currents in ocean straits with exchange of water masses between oceans and marginal seas. Marginal seas are defined as basins isolated by topography from the rest of the oceans. In Figures 4.2 and 4.A1 we show selected subsurface passages and straits that subdivide basins of the Atlantic Ocean and the Mediterranean Sea, principally allowing over- and outflows of bottom waters between them. In the long run, the deep drain of a semienclosed source basin is balanced by a refill with strongly cooled water masses from the surface. The latter phenomenon is called ‘‘convection’’ and is generally restricted to relatively small but highly effective regions of the oceans primarily located at polar latitudes. In extreme cases, observation of this process has revealed downwelling velocities of up to 0.1 m s1 (Marshall and Schott, 1999). The onset of convection is very sensitive to surface salinity. The deep cold North Atlantic limb of the thermohaline circulation (THC), occasionally called the ‘‘conveyor belt’’ (Broecker, 1991), is directed toward the equator. Farther to the south, 1
Equivalent color figures for the Indian and Pacific Oceans are given in Figure 4.A_(2b) and 4.A_(2c). Figures identified by letters are found in the enclosed CD-ROM (Compact Disk-Read Only Memory) along with the pertinent caption.
39
° De
40
o
n
ar –F
d
lan
Ice
ait ssa Str e Pa nel oe
nk
60
60
ge
rk
a nm
°
W. Zenk
a Ch
Ba
r Fa
°
40
°
Strait of Gibraltar
Discovery Gap
20°
it ra
20°
St
W ind An ward eg Pa ad s Pa a– J sage ss un ag gf e ern
Charlie Gibbs FZ
of Si c il y
Vema Gap
ara
Ce
0°
nche
Roma
0°
FZ
lain al P
yss Ab
20°
°
40
°
0°
30
°
°
60
°
90
40
60
°
South Sandwich Island Arc Gap
Sh a Pa g Ro ss ck ag s e
V Ch ema an ne l
20°
Figure 4.2 Representative selection of deep-ocean passages in the Atlantic Ocean and the Mediterranean Sea (courtesy Dr J.Whitehead,WHOI,Woods Hole, MA, USA).Water exchange across sills is called ‘‘overflow.’’ Straits show out- and inflows separated by a variable shear zone between the surface and the bottom. Light gray indicates depths 5000 m. A multicolor version of this figure is on the enclosed CD-ROM.
the Southern Ocean with connections to all three oceans enables a global distribution of abyssal currents on a secular timescale with diffused slow upwelling into the thermocline (Figure 4.3). The belt is closed by the interoceanic surface and nearsurface current system above the abyssal layers in the tropics and subtropics. It transports water masses for new convective sinking in polar regions of both hemispheres (Rahmstorf, 2002). Two of the prime propelling engines of the THC with the upright standing ‘‘U-turns’’ (Schiermeier, 2006) are situated in the Labrador and the Nordic Seas. Figure 4.B demonstrates the complex interaction between surface and abyssal currents in the North Atlantic, coupled by deep convection (vertical arrows). The shown cartoon by V. Byfield (personal communication, 2006) strongly simplifies the deep-sea basins by showing a uniform depth and square-edged ocean margins. Because dominant parts of the Atlantic THC2 follow a meridional direction, one often refers to the displayed advection processes (large horizontal arrows in Figure 4.B) as the ‘‘meridional overturning circulation’’ (MOC). More generally, 2
In the THC scenario, the wind-driven part of the global oceanic circulation is excluded per definition.
40
Abyssal and Contour Currents
Figure 4.3 Highly simplified cartoon of the global thermohaline circulation (THC), modified from the original presentation of the oceanic ‘‘conveyor belt’’ by Broecker (1991). For further details, see the multicolor version of Figure 4.3 on the CD-ROM (courtesy Dr S. Rahmstorf, PIK, Potsdam, Germany). Surface and near-surface waters flow toward convection regions in the Labrador and Nordic Seas, and in the Weddell and Ross Seas. They recirculate as deep and abyssal currents, and participate in basin-scale slow upwelling in the interior of all three oceans. Sea-surface salinitycontrols the convection process at high latitudes decisively.
physical oceanographers define the term ‘‘advection’’ as the transport of a water mass and its properties as a current in a three-dimensional velocity field. The former describes predominantly the horizontal large-scale flow, whereas ‘‘convection’’ refers to locally induced vertical motion driven by buoyancy forces.
4.2.
A BYSSAL C URRENTS IN THE G LOBAL T HERMOHALINE C IRCULATION
Historically, the basic concept of the THC was developed by Stommel (1958) and Stommel and Arons (1960): in the oceans’ interiors, rising abyssal water from a flat bottom is laterally replaced by freshly ventilated water from a limited number of sinks in polar regions. On the northern hemisphere, the transformation from light surface waters to heavy bottom water that starts the deep convection process is restricted to late winter time in semienclosed polar seas (S1 in Stommel’s simplified diagram (Figure 4.4) such as the Greenland, Norwegian and Labrador Seas. In the south, the Weddell (S2) and Ross Seas play analogous roles in deep-water renewal. The Stommel–Arons theory postulates the existence of a confined current in a 100–200 km wide corridor in the form of the Deep Western Boundary Current (DWBC) and a remaining, much larger, uniform upwelling regime with minimal horizontal drift. Boundary currents have been observed in all oceans (cf. Imawaki et al., 2001). They are not necessarily restricted to western continental slopes, but can also be guided by flanks of major submarine mountain chains such as the Mid-Atlantic Ridge (Figure 4.C; Machı´n et al., 2006).
41
W. Zenk
S1
S2
Figure 4.4 Classic model of the depth-integrated currents of the thermohaline circulation below 2000 m depth (after Stommel, 1958). S1 and S2 symbolize convective source regions at polar latitudes. The circulation in the interior is fed and driven by boundary currents on the western sides of the ocean basins.
In contrast to the spatially limited downwelling regions, Stommel and Arons assume for the upwelling at moderate and lower latitudes fairly broad lateral and long-term scales. Calculations of the global integral upwelling speed based on steady-state downward heat-flux estimates (Hogg, 2001) result clearly in the submillimeter per second range. Experiments to directly observe the very slow upward drift in the open ocean have so far been unsuccessful. Instead, the horizontal dominance of internal waves, ubiquitous eddies, and boundary currents have been repeatedly documented in numerous trials since early current measurements in the western North Atlantic (Swallow and Worthington, 1961; Crease, 1962). The vertical velocity of the widely spread upwelling process implies vortex stretching on a basin scale. Conservation of potential vorticity on the rotating earth induces the poleward return current beneath the base of the thermocline. Up till now, dedicated observations have failed to remove all inconsistencies of the Stommel–Arons theory. Even after the end of the decade-long World Ocean Circulation Experiment (WOCE; Siedler et al., 2001) two such inconsistencies remain: 1. Observations of tracer concentrations and float trajectories, both with high spatial resolution, show a clear dominance of deep zonal current directions. The classical theory, however, favors the conventional picture of advection in the form of large gyres in each of the major ocean basins without a preference of zonal currents (Hogg, 2001). 2. In theory the diapycnal3 diffusivity, required to support the thermocline circulation, appears to be one order of magnitude too high in comparison with observations (Webb and Suginohara, 2001). 3
Direction normal to the local isopycnal surface.
42
Abyssal and Contour Currents
0 –500 –1000
Water depth (m)
–1500 –2000 –2500 –3000 –3500 –4000 –4500 –5000 –5500 –6000 –38
0.0
–36
0.1
–34
0.2
–32
0.3
0.4
–30
0.5
–28 –26 Longitude
0.6
Diffusivity
0.7 (10–4
–24
0.8
–22
0.9
2.0
–20
5.0
–18
8.0
–16
22.0
m2 s–1)
Figure 4.5 Directly observed distribution of diffusivity in the Brazil Basin of the South Atlantic (from Polzin et al., 1997; with permission from The American Association for the Advancement of Science). The zonal section runs from the continental rise off Brazil (left side) toward the western flank of the Mid-Atlantic Ridge. Note the nonlinear scale for diffusivity. High correlation between diffusivity and bottom roughness is found on the slope of the ridge on the right side. The white line marks the observed depth of the 0.8°C potential isotherm. See also the multicolor version of Figure 4.5 on the CD-ROM.
Such concerns can be tested in general circulation models. Energy–budget studies suggest that the role of tidal mixing may have been underestimated in the past (Munk and Wunsch, 1998). Also varying bottom topography and roughness have a significant impact on the spatial distribution of density and velocity microstructure even hundreds of meters above the sea floor. A confirming example for the heterogeneous distribution of diapycnal diffusivity based on velocity microstructure observations in the Brazil Basin is shown in Figure 4.5 (Polzin et al., 1997). Above the bottommost 150 m of the rough Mid-Atlantic Ridge, their observations reveal an increased diffusivity that is about two orders of magnitude higher then above the smooth abyssal plain.
4.3. 4.3.1.
CONTOUR C URRENTS
Physical properties and modes
In physical oceanography, an abyssal current is generally defined as a flow of water masses beneath the bottom of the main thermocline or within the cold water sphere
43
W. Zenk
(cf. Zenk, 2001). In most cases, contour currents can be classified as a particular mode of abyssal currents, although their depth level is not exclusively associated with abyssal depths. Actually, most of the contour currents investigated so far occur at intermediate levels. They are controlled by topography and, as long as they have not yet reached their equilibrium level, by gravitation. Other forces acting on contour currents include pressure gradients, Coriolis and inertial forces, and – where applicable – bottom drag. Deflection against a side wall of a basin is an elementary property of all contour currents. Friction between the bottom and contour currents is part of the marine exogenic4 transformation processes. Besides the required minimum current velocity for erosion, transportation, and resuspension sedimentation depends critically on the cohesiveness of the ground and on the grain size of the available material. The simplified diagram in Figure 4.6 describes some aspects of the complex interaction of near-bottom currents versus sedimentation. Note the logarithmic scales on both axes of the experimentally determined interactions (Heezen and Hollister, 1971). A more extended version of Figure 4.6 was published later by Hollister and Heezen (1972). The term ‘‘contourite’’ was first specified by Hollister and Heezen (1972) for ‘‘contour current-deposed sediment’’ to contrast markedly with turbidite or
Current velocity (cm s–1) 0.01
0.1
1.0
10
100
Pebbles
e dim
l) E ohe rosion sive mat eria
Silt
(inc
Tr an
sp
or
ta
tio
n
Se
1.0
0.1
Grain diameter (mm)
tion
nta
Sand
(coh Erosion esive mate rial)
Granules
0.01
Clay
0.001
Figure 4.6 Simplified representation of near-bottom current velocities required for erosion, transport, and deposition (after Heezen and Hollister, 1971; with permission from WHOI, Woods Hole, MA, USA). Erosion velocities for fine particles are uncertain because they depend on the degree of sediment cohesion.
4
Originating outside the lithosphere.
44
Abyssal and Contour Currents
‘‘turbidity-current-deposed sediment’’. It is irrelevant where contour currents touch a side wall. The frictional interface between moving water layers and the sea floor may consist of the continental rise or slope, the shelf edge, or a bottom of a broad canyon. Contour currents differ from more event-controlled turbidity currents in so far that they show a quasi-steady flow, though potentially interrupted or enhanced by seasonal changes or migrating eddies. At the continental margins, contour currents run parallel to the slopes. Their counterparts, i.e. turbidity currents, follow paths perpendicular to the outer-shelf margins. Both currents are visibly involved in the morphological diversity of continental margins (Figure 4.D, after Blondel, 2003). Optical properties of contour currents, along with other physical parameters such as potential temperature, are often suitable to trace them over larger distances. As an example, the bottom part of Figure 4.7 shows a vertical section of light
Pressure (dbar)
3600
1.5
2
1.5
1
1
3800
1.5
0.5
4000
0
0.5
1
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4400
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4600 Potential temperature (°C)
4800
–39.5
–39.4
Pressure (dbar)
3600
–39.3 –39.2 Longitude (°W)
–39.1
–39
Clear water minimum 5.8
3800 5.9
5.9
4000
5.8 5.8
4200 4400 5.9
4600 4800
–38.9
6
Attenuation (0.1 m–1)
–39.5
–39.4
–39.3 –39.2 Longitude (°W)
–39.1
–39
–38.9
Figure 4.7 Vertical distribution of abyssal potential temperature (top) and of red-light attenuation (bottom) in a zonal section between the Santos Plateau and theVema Channel of the South Atlantic at approximately 29°S. See also the multicolor version of Figure 4.7 on the enclosed CD-ROM (Dr A. Macrander, AWI, Bremerhaven, Germany, 2006, personal communication). The Vema Channel represents a throughflow channel between the Argentine and Brazil Basins. The graph intersects the channel extension transporting cold ( 27.8.
to more advanced models with a stream tube approach (Smith, 1975) and parameterization of bottom drag and entrainment (Price and Baringer, 1994). Figure 4.17 displays the schematic by Price and Baringer’s outflow model on a rotating earth (f = Coriolis parameter) and associated density () profiles. represents the excess density due to the outflow. Outflows are relatively thin in comparison to the whole depth (D) of the water column. Their thickness (H) amounts to about 1% of their width (W). The density profile shows a stratified ocean with the underlying homogeneous outflow plume beneath H. The simple geostrophic balance between two opposing forces (Coriolis and buoyancy) is supplemented by the two parallel forces (bottom drag and entrainment) acting perpendicularly to both others. The resulting triangle of forces causes a slight downhill angle between isobaths on the slope and the spreading velocity (u). Price and Baringer (1994) were among the first who modeled a whole suite of out- and overflows in the North Atlantic and the Southern Ocean. Figure 4.G shows an example from the simulated pathway of the Mediterranean Water plume in the Gulf of Cadiz. The general course of the core layer around Cape St Vincent is quite satisfying when compared with selected RAFOS6 float trajectories shown in Figure 4.10. 6
‘‘SOund Fixing And Ranging’’ spelled backwards.
53
W. Zenk
C6A (21.3 Ma)
(a)
50°W
40°W
30°W
20°W
56°S
Present
60°S
64°S (b)
Figure 4.15 Sketch of bottom flows in the eastern Scotia Sea and northern Weddell Sea showing pathways of the Antarctic Circumpolar Current (ACC), Weddell Sea Deep Water (WSDW), and Weddell Sea Bottom Water (WSBW) (from Maldonado et al., 2003; with permission from Elsevier). A reconstruction from 21.3 Ma ago (top) can be compared with the present-day situation (bottom).The interaction of the various current systems generates a wide variety of contourites. A multicolor version of this figure is on the enclosed CD-ROM.
Further progress was achieved by the development of a hydrostatic, reduced gravity, two-dimensional primitive equation model with an application to DSOW along the slope of East Greenland (Jungclaus and Backhaus, 1994). Although models are increasingly successful in reconciling observations and simulations, the treatment of turbulent entrainment remains of primary importance
54
Abyssal and Contour Currents
Figure 4.16 The hypothetical extension of the westward counter (contour) current (see Figure 4.E on the CD-ROM) beneath the Antarctic Circumpolar Current (ACC) in Drake Passage (see inset on the lower right side) is shown as a black line. Its local existence was documented by moored current meters (D), in sediment cores (•), and by hydrographic observations (Giorgetti et al., 2003). The figure demonstrates the occasional co-existence of contour currents that parallel the continental slope and turbidity currents documented by the shown canyons or turbidity channels. Abbreviations: CTD = Conductivity, Temperature, Depth recorder; DSDP = Deep Sea Drilling Program; ODP = Ocean Drilling Program. For further details, see Rebesco et al. (1998b, 2007).
in large-scale ocean modeling. The descent and spreading of model overflow waters injected from high latitude and marginal seas, is extremely sensitive to the parameterization of the entrainment process. Mixing and acceleration of ambient waters in boundary currents strongly control transport and slow deepening of contour
55
W. Zenk
f/2
Depth
Density ρ
H
Outflow
δ+ρ
u Coriolis
u
Oceanic nt me g ain Dra r t En ttom Bo
W
δρ
H
D
Buoyancy
Figure 4.17 Schematic Mediterranean outflow in the Gulf of Cadiz (from Price and Baringer, 1994; with permission from Elsevier). The outflow plume (streamtube) leaves the Strait of Gibraltar (upper left corner of the left figure) and spreads westward on the bottom of the modeled gulf. The equilibrium of Coriolis, buoyancy, and frictional forces (bottom drag and entrainment) controls the down-slope direction of the modeled plume with velocity (u). The right figure shows the underlying density differences between the sea surface and a homogeneous outflow layer.
currents. For a better understanding of these processes in ocean models, further field observations and highly resolving numerical experiments are equally necessary (Price, 2002).
4.4.
C ONCLUSION AND SUMMARY
To the best of our knowledge, the term ‘‘contour current’’ as a particular mode of abyssal currents was first introduced in the (marine geological) literature by Heezen et al. (1966). Their original definition comprises near-bottom currents that ‘‘flow along isopycnals which are approximately parallel to the bathymetric contours’’. Due to a lack of definition in the physical oceanography literature,7 we suggest a more focused approach for ‘‘contour currents’’: Contour currents are predominantly unidirectional subsurface currents that are in contact with a sidewall. They are in quasi-geostrophic balance and are controlled by the local bottom topography. Their kinetic energy is attenuated by friction generating an exogenic stress on the sea floor. In marine sedimentology, the topographic control of contour currents gives rise to the generic name ‘‘contourite.’’ Contourites are sediments deposited or substantially reworked by
7
Even the extensive index of the comprehensive synthesis work of the World Ocean Circulation Experiment (Siedler et al., 2001) contains no entry on ‘‘contour currents.’’
56
Abyssal and Contour Currents
the action of contour and plain bottom currents, commonly in the vicinity of a continental rise. On descending slopes, contour currents start as density or gravity currents. They represent the confined advection of one fluid through another fluid (background water). While slowly sinking to their equilibrium level, their direction parallels largely the topography. They become mixing agents exchanging momentum and properties with ambient waters by turbulent entrainment. Where they are in contact with the sea floor, their bottom drag transforms kinetic energy into exogenic action on the sea floor. As a result, a large variety of contourite modes can be established. As a vicarious example, we reproduce in Figure 4.18 prime conceptual factors and processes that are engaged in contour currents and material transport along the Iberian Peninsula. The material has been compiled by H. de Haas (personal communication, 2006) for the cover of a synthesis volume of the Ocean Margin Exchange project (OMEX).8 We recognize the continental slope off Portugal with two stacked contour currents: the (seasonally) alternating slope current (Pingree et al., 1999) and the lateral northward transport of the MOW. Abyssal BNLs and INLs interact with downwelling of slope waters along the continental margin. Narrow canyons are natural obstacles in the course of bathymetryfollowing contour currents. In contrast, event-triggered turbidity currents in
La and teral re- trans sus p pen or t sio n
Win d upw -drive n elli ng
W MO
INL NL B
0
Ib
Sl
op
e
cu
e
ri
a
rre
20
00
wn Do ling l we L BN
40
00 41
°
nt m 200 0m
100
9°
n nyo
Ca
4 La 0° titu de (
N)
)
39
10°
e (W
tud ngi
Lo
°
Figure 4.18 Cartoon of the Iberian continental margin reflecting major processes and currents affecting particle transfer over the shelf edge and off the slope (from Pingree et al., 1999; with permission from Elsevier).The shown European Slope Current and the advection of Mediterranean Outflow Water (MOW) represent contour currents. Canyons act as pathways for direct transport of particles (occasionally by turbidity currents) from the shelf to the deep sea. A multicolor version of this figure is on the enclosed on the CD-ROM (Courtesy: Dr H. v. Haas, NIOZ,Texel,The Netherlands).
8
Funded between 1997 and 2000 by the European Commission under GD XII.
W. Zenk
57
canyons act as off-slope pathways for direct particle transport from the shelf right to an abyssal basin. Numerical simulations of outflow plumes as an approximation to the oceanographic problem of deep-water production and circulation show encouraging results. They also help us to better understand the dynamics of contour currents in three dimensions.
ACKNOWLEDGMENTS Many observational results shown in this chapter have been made possible by grants of the Deutsche Forschungsgemeinschaft, Bonn (Sonderforschungsbereich 460 – Dynamics of Thermohaline Circulation Variability). The German work in the Vema Channel has been funded by the Bundesministerium fu¨r Bildung und Forschung, Berlin, under CLIVAR-marine 2 (Climate Variability and Predictability).
C H A P T E R
5
D EEP - WATER B OTTOM C URRENTS AND T HEIR D EPOSITS G. Shanmugam Department of Earth and Environmental Sciences, The University of Texas at Arlington, Arlington, TX, USA
Contents 5.1. Introduction 5.1.1. Bottom currents versus turbidity currents 5.1.2. Genetic nomenclature 5.2. Thermohaline-Induced Geostrophic Bottom Currents 5.2.1. Antarctic and Artic bottom currents 5.2.2. Velocity 5.2.3. Deposits 5.2.4. Reservoir potential 5.3. Wind-Driven Bottom Currents: The Loop Current 5.3.1. Velocity 5.3.2. Deposits 5.3.3. Reservoir potential 5.4. Deep-Water Tidal Bottom Currents 5.4.1. Previous studies 5.4.2. Velocity 5.4.3. Deposits 5.4.4. Facies associations in submarine canyons 5.4.5. Reservoir potential 5.5. Internal Waves and Tides (Baroclinic Currents) 5.5.1. Nomenclature 5.5.2. Velocity 5.5.3. Deposits 5.6. Conclusions Acknowledgments
5.1.
59 59 61 63 63 63 64 66 66 67 68 70 72 72 72 75 76 76 77 77 79 80 81 81
I NTRODUCTION
The primary objective of this chapter is to discuss deep-water bottom currents and their deposits from oceanographical and sedimentological standpoints. The general term deep water is used here because bottom currents operate in deep waters of Developments in Sedimentology, Volume 60 ISSN 0070-4571, DOI: 10.1016/S0070-4571(08)00205-7
Ó 2008 Published by Elsevier B.V.
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Deep-water Bottom Currents and Their Deposits
both marine and lacustrine environments (Shanmugam, 2006a). The general term bottom current is used in this chapter because it covers a variety of bottom currents of different origins, flow directions, and velocities (Shanmugam et al., 1993a, p. 1242). Southard and Stanley (1976) recognized five types of bottom currents at the shelf break based on their origin. These currents are generated by (1) thermohaline differences, (2) wind forces, (3) tidal forces, (4) internal waves, and (5) surface waves. In addition, tsunami-related traction currents have been speculated to occur in bathyal waters (Yamazaki et al., 1989), but the mechanics of such currents are not yet understood (Shanmugam, 2007; 2008a). This chapter discusses four types of deepwater bottom currents, namely (1) thermohaline-induced geostrophic bottom currents (i.e., contour currents), (2) wind-driven bottom currents, (3) deep-water tidal bottom currents, and (4) internal waves and tides (baroclinic currents).
5.1.1.
Bottom currents versus turbidity currents
In discussing deep-water bottom currents, it is imperative that a clear distinction be made between bottom currents and turbidity currents. Turbidity current is a sediment flow with Newtonian rheology and turbulent state, in which sediment is supported by fluid turbulence and from which deposition occurs through suspension settling (Dott, 1963; Sanders, 1965; Middleton and Hampton, 1973; Shanmugam, 2000). Distinguishing deposits of deep-water bottom currents from those of turbidity currents has been and still is a challenge (Bouma and Hollister, 1973; Stow, 1979; Mulder et al., 2009). However, bottom currents and their deposits differ from turbidity currents and their deposits in the following respects: 1. bottom currents may occur on the shelf, slope, and in basinal environments, whereas turbidity currents are more common on the slope and basinal environments (Figure 5.1); 2. bottom currents are driven by thermohaline, wind, or tidal forces, whereas turbidity currents are driven by sediment gravity; 3. bottom currents may flow parallel to the strike of the regional slope, may flow in circular motions (gyres) unrelated to the slope, or may flow up and down submarine canyons (tidal), whereas turbidity currents always flow down-slope (Figure 5.2), though flow parallel to the strike of the regional slope may occur due to local morphology (e.g., marginal troughs); 4. bottom currents persist for long periods of time and can develop equilibrium conditions, whereas turbidity currents are episodic or surge-type events that fail to develop equilibrium conditions (Allen, 1973, 1985); 5. bottom currents can exist without the presence of entrained sediment and, for this reason, they are termed ‘‘clear water currents’’ (Bouma and Hollister, 1973, p. 82), whereas turbidity currents cannot exist without entrained sediment (Middleton and Hampton, 1973); 6. bottom currents show oscillating energy conditions, whereas turbidity currents exhibit waning energy conditions (Sanders, 1965); 7. bottom currents transport sand primarily by traction (i.e., bed load movement by sliding, rolling, and saltation; Allen, 1984), whereas turbidity currents generally transport fine-grained sand and mud in suspension;
G. Shanmugam
61
Figure 5.1 Schematic diagram showing complex deep-marine sedimentary environments occurring at water depths greater than 200 m (shelf ^ slope break). In general, shallow-marine (shelf ) environments are characterized by tides and waves, whereas deep-marine (slope and basin) environments are characterized by mass movements (i.e., slides, slumps, and debris flows), bottom currents, and pelagic/hemipelagic deposition. Turbidity currents may be common in basinal settings. Note up- and down-tidal bottom currents in submarine canyons (opposing arrows). Along-slope movement of contour-following bottom currents (contour currents) and circular motion of wind-driven bottom currents are important processes outside of the canyon. After Shanmugam (2003); with permission from Elsevier.
8. traction structures (e.g., parallel laminae, ripple laminae, and cross-beds) should be originally common in bottom-current sands (Shanmugam, 1997a; Martı´nChivelet et al., 2008) (though some authors suggest that this should be successively destroyed by pervasive bioturbation allowed by the continuous nature of bottom currents), whereas normal grading is the norm in turbidites that are deposited by relatively catastrophic episodic events of waning energy (Kuenen and Migliorini, 1950); 9. bottom-current deposits generally exhibit sharp upper contacts (Hollister, 1967), whereas turbidites show gradational upper contacts; 10. bottom currents can result in well-sorted sand with good porosity and permeability because of reworking and winnowing away of mud (Shanmugam et al., 1993a), whereas turbidites are poorly sorted, commonly mud-rich deposits with low porosity and permeability (Pettijohn, 1957; Sanders and Friedman, 1997).
62
Int erc ha nn el
Deep-water Bottom Currents and Their Deposits
Channel margin slump
Channel axis
Levee
Axial turbidity currents Overbank “turbidity” currents Contour currents
Figure 5.2 Conceptual model showing the spatial relationship between down-slope turbidity currents and along-slope bottom currents (contour currents). After Shanmugam et al. (1993a); with permission of the American Association of Petroleum Geologists.
5.1.2.
Genetic nomenclature
In science, words should have clear and consistent meanings. In geology, however, this is not always the case (Shanmugam, 2006b). The tradition of genetic nomenclature in sedimentary geology began with the introduction of the term turbidite for a deposit of a turbidity current in deep-water environments (Kuenen, 1957). Kuenen and Migliorini (1950, p. 99) and Kuenen (1967, p. 212) suggested that normal grading of a turbidite bed was a consequence of deposition from a single waning turbidity current. For a genetic term to succeed, (1) it must be based on sound fluid dynamic principles, (2) its usage must be accurate (relying on sedimentological description), precise (referring to a single process), and consistent (requiring a steady and a uniform application in time and space), and (3) it must imply a diagnostic flow behaviour. However, genetic terms of turbidites, which include (1) atypical turbidites, (2) fluxoturbidites, (3) hemiturbidites, (4) high-concentration sandy turbidites, (5) megaturbidites, (6) problematica turbidites, (7) seismoturbidites, and (8) undaturbidites, fail to reveal a clear flow behaviour (see Table 2.2 in Shanmugam, 2006a). Like turbidites, genetic terms of bottom-current-reworked sands also fail to divulge a diagnostic flow behaviour. 1. The term contourite emphasizes current orientation with respect to bathymetric contours (Hollister 1967), not the flow behaviour. 2. The term laminite represents sedimentary structure (i.e., lamina) (Lombard, 1963), not the flow behaviour. 3. The term tractionite impies traction deposition from bottom currents (Natland, 1967), but traction deposition has also been attributed to turbidity currents (Middleton and Hampton, 1973).
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4. The term tidalite impies alternating units of traction and suspension deposition from shallow-water tidal currents (Klein, 1971), but deep-water tidal bottom currents can also develop alternating units of traction and suspension deposition (Klein, 1975; Shanmugam, 2003). 5. The term winnowite impies winnowing action of bottom currents (Shanmugam and Moiola, 1982), but this is a reworking rather than a depositional process. 6. The term tsunamite represents tsunami-induced tractive bottom-currentreworked sediment (Yamazaki et al., 1989), but the term tsunami is not a self-defining expression of a single depositional process (Shanmugam, 2006b). In summary, genetic terms are ineffective for communicating depositional mechanics in process sedimentology (Shanmugam, 2006a). Therefore, I have adopted the general term ‘‘bottom-current-reworked sands.’’
5.2. 5.2.1.
T HERMOHALINE-I NDUCED GEOSTROPHIC BOTTOM C URRENTS
Antarctic and Artic bottom currents
Aspects of thermohaline-induced bottom currents are discussed by Zenk (2008). These deep-marine bottom currents in modern oceans became popular when Heezen et al. (1966) reported deep-water masses that flow along the ocean floor. An example of such bottom currents is the Antarctic Bottom Water (AABW). AABW was first identified by Brennecke (1921) in the northwest corner of the Weddell Sea in the Antarctic region (Figure 5.3). The origin of the AABW was attributed to the formation of ice from surface freezing over the Antarctic continental shelves. When sea ice forms, seawater experiences a concurrent increase in salinity due to salt rejection and a decrease in temperature. The increase in the density of cold saline (i.e., thermohaline) water directly beneath the ice triggers the sinking of a water mass down the continental slope and the spread to other parts of the ocean. The Western Boundary Undercurrent (WBUC or WBU), the Arctic counterpart to AABW, originates as a cold dense-water mass from the Norwegian Sea off Greenland (Worthington and Volkman, 1965). It flows along the western margin of the North Atlantic (Figure 5.3). These thermohaline currents tend to flow parallel to the slope, that is, along the slope at right angle to down-slope flowing gravity-induced currents (Figure 5.2). The WBUC is deflected in the northern hemisphere to the west as a result of the Coriolis force. Because of its tendency to flow parallel to bathymetric contours, the WBUC is known as a contour current (Heezen et al., 1966). These currents are commonly known as geostrophic contour currents, because they strike a balance between the Coriolis and the gravity forces.
5.2.2.
Velocity
Measured current velocities usually range from 1 to 20 cm s–1 (Hollister and Heezen, 1972); however, exceptionally strong, near-bottom currents with
64
Deep-water Bottom Currents and Their Deposits
80°
160°
100°
80°
0°
Arctic ocean
80°
Norwegian Sea
C
BU
40°
Ocean
W
ater
G GS
Antarctica
Indian ocean
Atl
BW AA
w
Pacific
p
40°
antic
Pacific ocean de e
0°
Weddell Sea
W
AAB
Antarctica
80°
Figure 5.3 Simplified circulation patterns of major thermohaline bottom currents (contour currents). Most contour currents originate from the Weddell Sea and from the Norwegian Sea. WBUC = Western Boundary Undercurrent. AABW =Antarctic BottomWater. GSG = Gulf Stream Gyre. Compiled from several sources (Wu«st, 1950; Stommel, 1958; Heezen and Hollister, 1971; Stow and Lovell, 1979).
maximum velocities of up to 3 m s 1 were recorded in the Strait of Gibraltar (Gonthier et al., 1984). Bottom-current velocities of 73 cm s 1 were measured at a water depth of 5 km on the lower continental rise off Nova Scotia (Richardson et al., 1981). Stow and Lovell (1979) summarized velocities of contour currents. Because of their high velocities, bottom currents in the deep sea are quite capable of erosion, transport, and redeposition of fine to coarse sand. Regional erosional unconformities in the deep sea throughout thousands of square kilometers of sea floor have been attributed to erosion by bottom currents (Berggren and Hollister, 1977; Tucholke and Embley, 1984; Shanmugam, 1988). In the Rockall Trough region, for example, bottom currents associated with the North Atlantic Deep Water (NADW) have caused an erosive area extending over 8500 km2 in water depths of 500–2000 m (Howe et al., 2001).
5.2.3.
Deposits
The following general features of bottom-current-reworked deposits have been discussed by Hubert (1964), Hollister (1967), Hollister and Heezen (1972), Bouma and Hollister (1973), Unrug (1977), Stow and Lovell (1979), Lovell and Stow (1981), Shanmugam (2000), and Ito (2002): • fine-grained sand and silt; • thin-bedded to laminated sand (usually less than 5 cm) associated with deepmarine mud;
65
G. Shanmugam
• • • • • • • • • • • •
rhythmic occurrence of sand and mud layers; sharp to gradational bottom contacts; sharp, non-erosional, upper contacts; well-sorted sand; little depositional mud matrix (clean sand); horizontal laminae; low-angle cross-laminae; ripple cross-laminae; lenticular bedding or starved ripples; inverse size grading (Figure 5.4); mud-offshoots in ripples; mud-draped ripples.
No single criterion by itself is unique to bottom-current-reworked sands. Although many of the criteria listed above can be attributed to processes other than bottom-current reworking, the association of several of the above criteria in a given deep-water example, along with the knowledge of the regional depositional setting, greatly enhances the chance of recognizing bottom-current-reworked facies. It is difficult to establish that a given sedimentary structure in the rock record was originated by contour-following thermohaline currents without establishing the paleowater circulation pattern independently. Therefore, the general term ‘‘bottom-current-reworked sands’’ is appropriate.
cm 5 4 3 2 1 0
Figure 5.4 Core photograph showing inverse grading and sharp upper contacts of sand layers (arrow), interpreted as bottom-current-reworked sands. Paleocene, North Sea.
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Deep-water Bottom Currents and Their Deposits
A contourite model, first proposed by Gonthier et al. (1984), consists of a basal inversely graded unit overlain by a normally graded unit (Stow et al., 1998a; Stow and Fauge`res, 2008). Criticism has been addressed in detail by Shanmugam (2000, 2006a). This facies model, for example, strongly suggests that bioturbation is a diagnostic feature. This suggestion is based on the belief that active contour currents would increase the oxygen concentration of the water mass, and thereby would increase the activity by benthic organisms. Tucholke et al. (1985), however, suggested that the degree of preservation of bioturbation is a function of bottom-current intensity; strong bottom currents do not favour preservation of biogenic structures. Bioturbated mud in the deep sea is equally abundant in areas that are unaffected by contour currents. Even if bioturbation were prevalent in areas of contour currents, it would not directly reveal anything unique about current orientation (i.e., contour-following currents). The bioturbation criterion for ‘‘contourites’’ is criticized because ancient deep-water turbidites (e.g., in the Late Cretaceous Point Loma Formation near San Diego, California) are extensively bioturbated and contain the trace fossil Ophiomorpha (Nilsen and Abbott, 1979). In contrast, convincing cases of ‘‘contourites’’ without bioturbation have been documented in the rock record (Dalrymple and Narbonne, 1996). In short, bioturbation cannot substantiate the contour-following current orientation, which is the basic tenet of the ‘‘contourite’’ model (Hollister, 1967). Finally, analogous to the five divisions (Ta, Tb, Tc, Td, and Te) of the turbidite facies model (i.e., the Bouma Sequence), Stow and Fauge`res (2008, their Figure 13.9) have introduced five divisions (C1, C2, C3, C4, and C5) for their contourite facies model. In defense of their facies model, Stow and Fauge`res (2008, p. 240) have argued that process models derived from ancient strata are less reliable in comparison to their contourite model derived from modern sediments. Their argument is specious because the Bouma Sequence, which they have used as the analogy for their model, was derived strictly from a study of ancient strata. Furthermore, no one has ever documented the complete Bouma Sequence from modern sediments. Idealistic turbidite and contourite facies models, despite their popularity, are a step backward in the pragmatic science of process sedimentology (Shanmugam, 2008b).
5.2.4.
Reservoir potential
Aspects of economic relevance of bottom-current-reworked sands are discussed by Viana et al., (2007). From a reservoir point of view, thermohaline bottom currents are important because of their ability to intrude and rework sand. A thick prism of calciclastic reworked sands (middle Miocene to Pleistocene) off Little and Great Bahama Banks has been studied in detail (Mullins et al., 1980). These calciclastic reworked sands were lithified by early submarine cementation. Measured maximum porosity and permeability values were 40% and 9880 mD, respectively. High permeability values are attributed to winnowing away of mud by vigorous bottom currents. Such bottom-current-reworked sands are potential petroleum reservoirs.
67
G. Shanmugam
5.3.
W IND-D RIVEN BOTTOM C URRENTS: THE LOOP C URRENT
The Loop Current in the eastern Gulf of Mexico is a wind-driven surface current (Figure 5.5). It enters the Gulf of Mexico through the Yucatan Straight as the Yucatan Current; it then flows in a clockwise loop in the eastern Gulf as the Loop Current, and exits via the Florida Strait as the Florida Current (Neumann and Pierson, 1966; Nowlin, 1972; Mullins et al., 1987). Finally, this current merges with the Antilles Current to form the Gulf Stream. The Loop Current also propagates eddies into the north-central Gulf of Mexico, where the Ewing Bank area, a case study used in this chapter, is located (Figure 5.5).
5.3.1.
Velocity
Velocities in eddies that have detached from the Loop Current have been recorded as high as 2 m s 1 at a depth of 100 m (Cooper et al., 1990). The Loop Current and related eddies pose significant problems for deep-water drilling (Koch et al., 1991).
36° m
N
f St
200 m
An
a Current
op
Gulf of
Atlantic Ocean
Gul
a
xico Me
Lo
tille
sC
urr
en
Cu
t
rre
r id
28°
rid
Texas
Flo
Ewing Bank 826 Garden Banks
rea
United States 32°
nt
Flo
24°
Yu c
ata
nC
20°
Mexico
ur
re
m 200
16°
nt
Nicaragua
Caribbean Current Rise
Pacific Ocean 10°
102°
0
98°
Isthmus of Panama
400 km
94°
South America 90°
86°
82°
78°
74°
70°
66°
Figure 5.5 Present circulation pattern of the Loop Current in the Gulf of Mexico.This winddriven surface current is considered to be affecting the sea floor (Pequegnat, 1972). Note the detached eddies from the Loop Current in the Ewing Bank area. After Shanmugam et al. (1993a); with permission of the American Association of Petroleum Geologists.
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Deep-water Bottom Currents and Their Deposits
For example, drilling operations in the Green Canyon 166 area were temporarily suspended in August of 1989 because of high current velocities that reached nearly 150 cm s 1 at a depth of 45 m, and 50 cm s 1 at a depth of 250 m. These intense bottom currents affect the ability of a drilling rig to hold station over a wellhead (Koch et al., 1991). Current-velocity measurements, bottom photographs, high-resolution seismic records, and GLORIA side-scan sonar records indicate that the Loop Current influences the sea floor at least periodically in the Gulf of Mexico (Pequegnat, 1972). Computed flow velocities of the Loop Current vary from nearly 1 m s 1 at the sea surface to more than 25 cm s 1 at 500 m water depth (Nowlin and Hubert, 1972). This high surface velocity suggests a wind-driven origin for these currents. Flow velocities measured using a current meter reach up to 19 cm s 1 at a depth of 3286 m (Pequegnat, 1972). Such currents are capable of reworking fine-grained sand on the sea floor. Current ripples, composed of sand at a depth of 3091 m on the sea floor, are clear evidence of deep bottom-current activity in the Gulf of Mexico today (Pequegnat, 1972). These current ripples are draped by thin layers of mud. If these mud drapes on sand ripples were preserved in the rock record, they would be termed ‘‘mud-offshoots.’’
5.3.2.
Deposits
Deposits of the Loop Current have been interpreted in the cores from the Ewing Bank 826 Field, Plio Pleistocene, Gulf of Mexico. The Ewing Bank Block 826 Field is located nearly 100 km off the Louisiana coast in the northern Gulf of Mexico (Figure 5.5). It contains hydrocarbon-producing clastic reservoir sands that have been interpreted as bottom-current-reworked sands (Shanmugam et al., 1993a, b). The Ewing Bank cores exhibit the following features: • predominantly fine-grained sand and silt; • thin-bedded to laminated sand (usually less than 5 cm) intercalated with deepwater mud (Figure 5.6); • rhythmic sand and mud layers; • numerous layers (50 or more per 1 m of core); • sharp (non-erosional) upper contacts of sand layers (Figure 5.6); • sharp to gradational bottom contacts; • internal erosional surfaces; • external truncation surfaces; • megascopic inverse size grading (Figure 5.7); • microscopic inverse size grading; • horizontal lamination and low-angle cross-lamination (Figure 5.6); • cross-bedding; • lenticular bedding or starved ripples at core scale (Figure 5.6); • current ripples with preserved crest or with eroded crest; • ripple forms with curved bases; • flaser bedding; • mud-offshoots (Figure 5.8); • well-sorted sand and little depositional matrix.
1 cm
Figure 5.6 Core photograph showing discrete thin sand layers with sharp upper contacts. Traction structures include horizontal laminae, starved ripples (arrows), and low-angle crosslaminae. Dip of cross-laminae to the right suggests current from left to right. Note rhythmic occurrence of sand and mud layers. Middle Pleistocene, Gulf of Mexico. After Shanmugam et al. (1993a); with permission of the American Association of Petroleum Geologists.
Figure 5.7 Core photograph showing megascopic inverse size grading. Arrow shows the gradational nature of basal contact from mud (dark color) at the bottom to sand (light color) at the top. Each scale division is 3 cm. Middle Pleistocene, Gulf of Mexico. After Shanmugam et al. (1993a); with permission of the American Association of Petroleum Geologists.
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Deep-water Bottom Currents and Their Deposits
cm 5 4 3 2 1 0
Figure 5.8 Core photograph showing discrete sand units with current ripples and mud offshoots (arrow). Note sigmoidal configuration of ripples and truncated tops. Middle Pleistocene, Ewing Bank Block 826, Gulf of Mexico. After Shanmugam et al. (1993a); with permission of the American Association of Petroleum Geologists.
Most of the features listed above are interpreted as the products of deposition by traction or combined traction and suspension (Figure 5.9). Sand layers with traction structures occur in discrete units, but not as part of a vertical sequence of structures. These features are interpreted here as evidence for bottom-current reworking. Because traction structures are also observed in ‘contourite’ deposits discussed before, caution must be exercised in classifying a deposit as a ‘contourite’ based solely on traction structures without independent evidence for contour-following bottom currents in the area.
5.3.3.
Reservoir potential
Details of reservoir quality of reworked sands in the Ewing Bank area have been discussed by Shanmugam et al. (1993a). Sedimentary structures in cores reveal that the reservoir is composed of a lower unit with a turbidite channel sand, and an upper unit with a bottom-current-reworked sand. Although both the turbidite and the reworked sands show porosity values in the range of 35–42%, their permeability values are strikingly different. The basal turbidite channel sand unit shows a distinct upward decrease in permeability (with
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Mud offshoots
Climbing ripple cross-bedding
Suspension (mud offshoot)
Traction and suspension
5 cm
5 cm
Traction
Flaser bedding
Lenticular bedding
Suspension Traction
Suspension
5 cm
5 cm
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Horizontal bedding
Rhythmic bedding Sharp upper contact
Traction
5 cm
1 cm
Traction Suspension
Cross-bedding
Sharp upper contact Sharp upper contact
Inverse grading
Erosion
10 cm
5 cm
Traction
Fine sand
Graditional Lower contact
Mud
Figure 5.9 Summary of traction features interpreted as indicative of deep-water bottomcurrent reworking. After Shanmugam et al. (1993b); with permission of the Geological Society of America.
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Deep-water Bottom Currents and Their Deposits
maximum values of 5000 mD in the turbidites and about 500 mD in the bottom-current-reworked sands).
5.4.
D EEP-W ATER T IDAL BOTTOM C URRENTS
In this chapter, I maintain a distinction between ‘‘deep-water tidal bottom currents’’ as a general category and ‘‘baroclinic currents’’ as a special category. This is necessary because deep-water tidal bottom currents may include baroclinic currents. However, not all deep-water tidal currents are baroclinic types. In addition, our understanding of baroclinic currents associated with internal tides is still in its infancy. Controversies and complexities associated with internal tides and ocean mixing have been addressed by Garrett (2003).
5.4.1.
Previous studies
Deep-marine tidal bottom currents in submarine canyons and in their vicinity are one of the best-studied and most extensively documented modern geologic processes (e.g., Shepard et al., 1969, 1979; Shepard, 1976; Beaulieu and Baldwin, 1998; Petruncio et al., 1998; Xu et al., 2002). During the past four decades, an understanding of deep-marine tidal bottom currents has been developed by synthesizing a great wealth of published information. This information includes direct observations from deep-diving vehicles, dredging from canyon floors, underwater photographs of canyon floors, photographs and X-radiographs of box cores, seismic profiles, current-velocity measurements (Shepard and Dill, 1966; Shepard et al., 1969, 1979; Dill et al., 1975; Shepard, 1976), and from study of conventional cores and outcrops (Shanmugam, 1997b, 2002). Selected examples of studies that dealt with tidal processes and/or their deposits in modern and ancient deep-water environments were reviewed by Shanmugam (2003).
5.4.2.
Velocity
Tidal currents are significant processes in many modern submarine canyons (Shepard et al., 1979). The interaction of the canyon topography with the tidal current is particularly important. In the modern Zaire (formerly the Congo) Canyon in West Africa, the canyon head can be traced 25 km up the estuary on land (Heezen et al., 1964b; Shepard and Emery, 1973; Droz et al., 1996). The deep Zaire Canyon is simply a deep-water extension of the Zaire estuary. The width and the relief of the canyon increase seaward from the estuary reaching a maximum width of 15 km and a maximum relief of 1300 m near the shelf break (Babonneau et al. 2002). The mean tidal range in the Zaire Canyon is 1.3 m (Shepard et al., 1979). Shepard et al. (1979) documented systematically that up- and down-canyon currents closely correlated with timing of tides (Figure 5.10). Shepard et al. (1979) measured current velocities in 25 submarine canyons at water depths
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Hueneme Sta. 28 Lo.1 Depth 448 m 3 mAB V vs Time
Down-canyon–Up-canyon cm s–1
30
30 cm s–1
20
20 cm s–1
10
10 cm s–1 Data from tide tables
0
Tide
2m
–1
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20
20 cm s–1
30
30 cm s–1
1
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2000 h 2/12/73
12
24
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48
60
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h
Figure 5.10 Time/velocity plot of data obtained at 448 m in the Hueneme Canyon, California, showing excellent correlation between the timing of up- and down-canyon currents and the timing of tides obtained from tide tables (solid curve). 3 mAB = velocity measurements were made 3 m above sea bottom. Modified after Shepard et al. (1979); with permission of the American Association of Petroleum Geologists.
ranging from 46 to 4200 m by suspending current meters, usually 3 m above the sea bottom (Figure 5.11). These canyons include the Hydrographer, Hudson, Wilmington, and Zaire in the Atlantic Ocean; and the Monterey, Hueneme, Redondo, La Jolla/Scripps, and Hawaii canyons in the Pacific Ocean. Maximum velocities of up- and down-canyon currents commonly ranged from 25 to 50 cm s 1. Keller and Shepard (1978) reported velocities as high as 70–75 cm s 1, velocities sufficient to transport even coarse-grained sand, from the Hydographer Canyon. In the Niger Delta area of West Africa, five modern submarine canyons (Avon, Mahin, Niger, Qua Iboe, and Calabar) have been recognized (Figure 5.12). In the Calabar River, the tidal range is 2.8 m and tidal flows with reversible currents are common (Allen, 1965). In the Calabar estuary, maximum ebb-current velocities range from 60 to 280 cm s 1, and flood current velocities range from 30 to 150 cm s 1. These velocities are strong enough to transport particles of sand and gravel size. The Calabar estuary has a deep-water counterpart that cuts through sediments of the outer shelf and slope, forming the modern Calabar submarine canyon (Figure 5.12). Thus, as they do in the Zaire Canyon to the south, tidal currents are likely to operate in the Calabar and Qua Iboe canyons.
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Deep-water Bottom Currents and Their Deposits
Submarine canyon Depth: 46–4200 m
Sea level
Ebb flood
Velocity range: 25–50 cm s–1
Current meter: 3 m above base
Figure 5.11 Conceptual diagram showing cross-section of a submarine canyon with ebb and flood tidal currents (opposing arrows). Shepard et al. (1979) measured current velocities in 25 submarine canyons at water depths ranging from 46 to 4200 m by suspending current meters commonly 3 m above the sea bottom. Measured maximum velocities commonly range from 25 to 50 cm s 1. After Shanmugam (2003); with permission from Elsevier.
Av on
Opuama Canyon
6°N
hin
Gu lf
Afam Canyon
Shelf Slope
Ma
Qua lboe Canyon
lab Ca
Ancient canyon 100 km 2°E
4°
ar
bo e Qu
al
Gu in ea
Recent submarine canyons
Niger
of
2°N 4°
6°
8°
10°E
Figure 5.12 Location of modern and ancient submarine canyons in the Gulf of Guinea,West Africa. Outside submarine canyons, the shelf ^ slope break (dashed line) is not only an important physiographic boundary between shelf and slope, but also a major controlling factor of processes on the shelf (e.g., tides and waves) and on the slope (e.g., mass transport). However, within submarine canyons (e.g., recent Calabar Canyon), the shelf ^ slope break does not control processes. Map modified after Petters (1984); with permission from Blackwell Publishing.
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5.4.3.
Deposits
Sedimentary features indicative of tidal processes in shallow water environments have been well established (e.g., Reineck and Wunderlich, 1968; Klein, 1970; Visser, 1980; Terwindt, 1981; Allen, 1982; Banerjee, 1989; Nio and Yang, 1991; Dalrymple, 1992; Archer, 1998; Shanmugam et al., 2000). Traction structures that develop in shallow-water estuaries also develop in deep-water canyons and channels with tidal currents (Shanmugam, 2003). General characteristics of deep-water tidal deposits are • • • • • • • • • • • • •
heterolithic facies; rhythmic alternation of sandstone/shale couplets (tidal rhythmites); thick (spring)/thin (neap) bundles; alternation of parallel and cross-laminae; double mud layers (Figure 5.13); climbing ripples; cross-beds with mud-draped foresets; bidirectional (herringbone) cross-bedding; sigmoidal cross-bedding (i.e., full-vortex structures) with mud drapes and tangential basal contacts; reactivation surfaces; crinkled laminae; elongate mudstone clasts; flaser bedding;
cm 5 4 3 2 1 0
Figure 5.13 Core photograph showing double mud layers (arrow) in Pliocene sand. Edop Field. After Shanmugam (2003); with permission from Elsevier. A multicolor version of this figure is on the enclosed CD-ROM.
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• wavy bedding; • lenticular bedding; • alternating traction and suspension structures.
5.4.4.
Facies associations in submarine canyons
Submarine canyons are not only unique for providing a protected environment for focusing tidal energy from shallow-marine estuaries to deep-marine canyons, but also prone for generating mass movements (e.g., slides, slumps, grain flows, and debris flows) due to failure of steep canyon walls. The Edop Field is located in the ancient Qua Iboe Canyon (Figure 5.12). The Pliocene Intra Qua Iboe (IQI) reservoir in the Edop Field is a major hydrocarbonproducing siliciclastic reservoir. Based on recognition of a 3-km-wide erosional feature observed on a seismic time slice, a submarine canyon for the Edop reservoir has been documented (Shanmugam, 1997b, his Figure 21). The bulk of the cored interval is interpreted to be deposits of sandy and muddy slumps and debris flows. Some cored intervals are composed of fine to very fine sand with well-developed double mud layers (Figure 5.13), mud-draped ripples, and tidal rhythmites with thick and thin sand layers. These features have been interpreted as products of deep-water tidal currents. Such a close association of mass-flow deposits and deep-marine tidal deposits is an indication of canyon-fill facies (Figure 5.14).
5.4.5.
Reservoir potential
Recognition of tidal facies in deep-water successions has implications for reservoir potential. For example, in channel-mouth environments, down-slope turbidity currents are likely to develop depositional lobes, whereas bi-directional tidal bottom currents are likely to develop elongate bars (Shanmugam, 2003). Turbidite Double mud layers (tidal) Floating mudstone clasts and quartz granules (debris flows)
Contorted layers and floating quartz granules (slumps) Double mud layers (tidal) Mud-draped ripples (tidal)
Figure 5.14 Facies association showing interbedded occurrence of double mud layers (tidal origin), floating mudstone clasts and quartz granules (debris-flow origin), double mud layers (tidal origin), contorted layers and floating quartz granules (slump origin), and double mud layers with mud-draped ripples (tidal origin). In the rock record, such a facies association may be used as evidence for deposition within submarine canyons. Modified after Shanmugam (2003); with permission from Elsevier.
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lobes are aligned perpendicular to the channel axis, whereas tidal bars are aligned parallel to the channel axis. Depositional lobes are likely to be much larger than channel width, whereas tidal sand bars are thought to be much smaller than channel width. Deep-water elongate tidal bars are speculated to be analogous to tidal bar sands that develop in shallow-water estuarine environments (see Shanmugam et al., 2000). In frontier exploration areas, an incorrect use of a turbidite-lobe model (with sheet geometry) instead of a tidal bar model (with bar geometry) will result in an unrealistic overestimation of sandstone reservoirs.
5.5. 5.5.1.
I NTERNAL W AVES AND T IDES (BAROCLINIC C URRENTS)
Nomenclature
A plethora of nomenclature is in use for internal waves and tides. In particular, the term baroclinic has been used by different authors with different meanings (Wunsch, 1996). Thus it is useful to explain these terms to minimize confusion. Surface waves, caused by wind (meteorological force) blowing over the water surface (Komar, 1976), develop at the interface between water and air. Internal waves, first reported by Ekman (1904), are gravity waves that oscillate along the interface between two water layers of different densities (Figure 5.15). These waves are common phenomena in coastal seas, fjords, lakes, and the atmosphere. In shallow-water shelf environments, waters can range from well mixed to density-stratified types. Most shelf waters are vertically well mixed. In deep-water environments, however, most of the ocean is vertically stratified, with an upper low-density layer and a lower high-density layer (Figure 5.15). The interface between layers of different densities (i.e., pycnocline) can be caused either by differences in temperature (i.e., thermocline) or by salinity (i.e., halocline). Navrotsky et al. (2004) made observations of internal waves and spatial inhomogeneities of a thermohaline structure during the warm season on the continental slope and shelf of the Sea of Japan. Santek and Winguth (2005) documented internal waves, induced by the 2004 Indian Ocean tsunami, along the eastern continental slope off Sri Lanka. Deep water
Shallow water Surface waves (meteorological) Shelf
Sea level
Internal waves and tides (astronomical)
1. Low-density (warm) layer 2. High-density (cold) layer
1 2
Slo
Thermocline
pe
Basin
Figure 5.15 Schematic diagram showing meteorological surface waves in shallow-water shelf environments and astronomical internal waves in deep-water slope and basinal environments. Internal waves, occurring along a thermocline, can extend onto shelf environments. Internal waves typically have much higher amplitudes than surface waves. Diagram based on concepts of Inman et al. (1976) and Maxworthy (1979). Not to scale.
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Figure 5.16 Satellite image of internal waves in the Sulu Sea between the Philippines (to the northeast) and Malaysia (to the southwest). Sunlight highlights delicate curving lines of internal waves moving to the north toward Palawan Island. The Sulu Sea is stratified with water layers of differing densities. Unlike surface waves, however, internal waves can stretch tens of kilometers in length and move throughout the ocean for several hours. This true-color Aqua Moderate Resolution Imaging Spectroradiometer (MODIS) image was acquired on April 8, 2003. Image courtesy Jacques Descloitres, MODIS Land Rapid Response Team at NASA/GSFC. Uniform Resource Locator (URL): http://earthobservatory.nasa.gov/ Newsroom/NewImages/images.php3?img_id=15334 (accessed May 12, 2007). A multicolor version of this figure is on the enclosed CD-ROM.
Internal waves typically have much lower frequencies and higher amplitudes than surface waves because the density difference between two water layers is typically much less than the density difference between water and air. Unlike surface waves, internal waves can stretch tens of kilometers in length (Figure 5.16). They can move throughout the ocean for several hours. They have their greatest wave height at intermediate depths and their greatest velocities at the bottom (LaFond, 1962). Internal waves are of significance not only for maintaining ocean circulation by downward mixing of heat, but also for sustaining biological productivity by supplying nutrients. Internal waves that correspond to periods of tides are called internal tides (Shepard, 1975). It is important to distinguish the surface (barotropic) tide from the internal (baroclinic) tide. This is because sedimentological aspects of deposits of surface tides are well established (Alexander et al., 1998), whereas those of internal tides are not. The passage of tropical cyclone Bobby in 1995 over the Western Australian Shelf influenced the generation of internal tides at 300 m depth (Davidson and Holloway, 2003). The currents associated with internal tides are
G. Shanmugam
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termed baroclinic currents or motions. These tidal currents tend to create turbulent mixing in the stratified water column as well as at the sea floor. The term near-inertial wave has been used for internal waves that correspond to periods of upwelling (Federiuk and Allen, 1996). Internal solitary waves, consisting of a single isolated wave, are common in stratified fluids. The term soliton is commonly used as a synonym for solitary waves. Although the term ‘‘soliton’’ is strictly a mathematical solution of a non-linear internal-wave theory, it has become common practice in the offshore drilling industry to use the term to describe observations of large-amplitude nonlinear internal waves in the ocean (Hyder et al., 2005). The origin of internal solitary waves has been attributed to a number of causes, including sea-floor topography (Maxworthy, 1979; Farmer and Armi, 1999). Because internal solitons are hazardous to offshore drilling operations (Fraser, 1998), a clear understanding of these internal waves has practical implications for the cost and the safety of offshore drilling operations.
5.5.2.
Velocity
Deep-water bottom currents have been attributed to internal waves in offshore California (Emery, 1956). Wunsch (1969) proposed amplification of near-bottom velocities as internal waves propagate over a shoaling bottom. Brandt et al. (2002) reported results of high-resolution velocity measurements carried out by means of a vessel-mounted acoustic Doppler current profiler on the 12 November 2000 in the equatorial Atlantic, at 44°W between 4.5°N and 6°N. The data showed the presence of three large-amplitude internal solitary waves. The pulse-like intense solitary disturbances propagated, perpendicular to the Brazilian Shelf, toward the north–northeast. These internal waves were characterized by maximum horizontal velocities of about 2 m s 1 and maximum vertical velocities of about 20 cm s 1. Shepard et al. (1979), who studied bottom currents in submarine canyons, documented that internal waves advance in both up- and down-canyon directions. Measured values of velocity reach up to 1 m s 1 in the up-canyon direction and 265 cm s 1 in the down-canyon direction. Shepard (1975) suggested that internal waves, which occur in canyon depths of up to 3500 m,were mostly tidal in origin (i.e., internal tides). In a stratified ocean, internal tides are generated commonly above an area of steep bathymetric variation, such as the shelf break. An example is the Bay of Biscay, where the internal tides are the most energetic (Baines, 1982). Internal tides travel slowly compared with surface gravity waves. Hyder et al. (2005) made observations of internal solitons that occurred between January and April 1998 at a water depth of 440 m northeast of the Andaman Islands, Bay of Bengal. Their observations indicated the occurrence of internal solitons with thermocline depressions of up to 50 m and an upper-layer current velocity of up to 120 cm s 1. These solitons only occurred during spring tides, when the tidal range exceeded 1.5 m. The advances in tidal mapping afforded by the Topex/Poseidon satellite have allowed Egbert and Ray (2000) to answer some longstanding questions about tidal energetics. Analyses of the altimeter-derived cotidal charts reveal that most tidal energy is dissipated in shallow seas, but about 25 30% of the global energy
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dissipation occurs generally near rugged bottom topography in the open ocean. St. Laurent et al. (2003) also reported that internal tides are generated commonly at regions where the barotropic tidal current encounters variations in bottom topography. Semi-diurnal internal tidal currents are likely major factors in shaping continental slopes. Cacchione et al. (2002) have discussed the intensification of near-bottom water velocities, caused by reflection of semi-diurnal internal tides, and their role on sedimentation patterns and bottom gradients of continental slopes off northern California and New Jersey. Oceanographic studies have shown that the Makassar Strait in Indonesia is dominated by strong tidal currents, internal waves/tides, and solitons (Ray et al., 2005). The Indonesian Throughflow (ITF), which passes through the Makassar Strait, is a series of ocean currents that flow from the tropical Pacific Ocean through the Indonesian Seas into the Indian Ocean. It transports nearly 10 Sv (1 Sv or Sverdrup = 106 m3 s 1) of the Pacific Ocean water into the Indian Ocean. The ITF, a thermocline flow, is stratified along the Makassar sill depth of 680 m (Gordon, 2005). Current meter measurements from two moorings in the Labani Channel recorded velocities in excess of 50 cm s 1 at 250 m (Wajsowicz et al., 2003). At these velocities, medium-grained sand can be eroded and transported by baroclinic tidal currents.
5.5.3.
Deposits
The potential significance of shoaling internal waves for causing sediment movement on continental shelves and slopes has been discussed by Cacchione and Southard (1974). Laboratory experiments confirmed that shoaling interfacial waves could generate ripples and larger bedforms in artificial sediment (Southard and Cacchione, 1972). Stride and Tucker (1960) attributed the development of modern sand waves near the shelf edge to internal waves. Karl et al. (1986), using sparker profiles, documented sand waves in the heads of submarine canyons of the Bering Sea. In this case, a surface sediment sample (C1) was composed of 19% gravel, 76% sand, and 5% mud. The modal class of this sample was fine sand. However, no sedimentary structures were described from the cores from these sand waves. Karl et al. (1986) speculated that internal waves were responsible for the origin of sand waves. They also suggested that delivery of large volumes of fresh water and large quantities of sediment directly to the heads of submarine canyons during periods of low sea level might have enhanced the propagation of highfrequency internal waves. Gao et al. (1998) interpreted ancient strata with bidirectional cross-bedding, flaser bedding, wavy bedding, and lenticular bedding as deposits of internal tides based on associated deep-water turbidite and slump facies. The key to interpreting deposits of ‘‘internal tides’’ or baroclinic currents in the rock record is the evidence for tidal currents in a stratified deep ocean. Without that evidence for density stratification, there is no difference between a tidal deposit formed by surface (barotropic) tide in a shallow-marine shelf and a tidal deposit formed by internal (baroclinic) tide in a deep-marine slope or canyon environment. Furthermore, tidal bottom currents in submarine canyons may be unrelated to density stratification. Until we develop objective sedimentological criteria for
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recognizing deposits of internal tides of stratified water bodies, it is preferred to classify deep-water deposits with tidal signatures as products of ‘‘deep-water tidal bottom currents’’ rather than of ‘‘baroclinic currents’’ (i.e., internal tides). This is a topic for future research.
5.6.
C ONCLUSIONS
Four basic types of deep-water bottom currents and their deposits are discussed: (1) thermohaline-induced geostrophic bottom currents, (2) wind-driven bottom currents, (3) deep-water tidal bottom currents, and (4) internal waves and tides (baroclinic currents). • A distinctive attribute of reworked sands by thermohaline-induced and winddriven bottom currents is their traction structures. Common sedimentary features of these traction currents include cross-bedding, current ripples, horizontal lamination, sharp upper contacts, and inverse grading. These sands also exhibit internal erosional surfaces and mud-offshoots indicating oscillating energy conditions. • Deposits of deep-marine tidal bottom currents are sand/mud rhythmites, double mud layers, climbing ripples, mud-draped ripples, alternation of parallel and cross-laminae, sigmoidal cross-bedding with mud drapes, internal erosional surfaces, lenticular bedding, and flaser bedding. These features represent alternating events of traction and suspension deposition. • There are no objective sedimentological criteria to recognize deposits of internal tides in the rock record. • Sedimentary structures of bottom-current deposits occur in discrete units, not as a predictable vertical sequence.
ACKNOWLEDGMENTS I thank M. Rebesco, A. Camerlenghi, A. J. van Loon, and D. W. Kirkland for helpful comments.
C H A P T E R
6
D YNAMICS OF THE B OTTOM B OUNDARY L AYER S. Salon1, A. Crise1 and A.J. Van Loon2 1
Istituto Nazionale di Oceanografia e Geofisica Sperimentale, Sgonico (TS), Italy Geological Institute, Adam Mickiewicz University, Poznan, Poland
2
Contents 6.1. Introduction 6.2. Spectral Windows 6.3. Characteristics of the Bottom Boundary Layer 6.3.1. General characteristics 6.3.2. Influence of wind stress 6.3.3. Interactions with the sea floor and suspended particles 6.4. Analytical Approach of the BBL 6.4.1. The Ekman layer 6.4.2. The viscous sub-layer (bed layer) 6.4.3. The logarithmic layer 6.5. Conclusions Acknowledgements
6.1.
83 85 87 87 88 88 88 89 94 95 96 97
I NTRODUCTION
The water column of the ocean can be roughly divided into the following three main layers, which differ from one another because of the role played by friction: (1) the surface boundary layer, which is due to atmospheric forcing that is expressed by a frictional force (wind stress) that is felt by the ocean surface; (2) the ocean interior, which is essentially frictionless and which is characterized by the balance, commonly defined as geostrophic balance, between the horizontal pressure gradients and the Coriolis force (which forces water flows on the northern hemisphere to move to the right, and on the southern hemisphere to move to the left)1; and (3) the boundary layer near the sea floor, where friction between the water and 1
The term ‘‘geostrophy’’ comes from Greek or = Earth, and o = turning. The attribute ‘‘geostrophic’’ is always related to flows characterized by an equilibrium between the horizontal pressure gradient and the Coriolis force. In the ocean’s interior, currents with horizontal dimensions of over a few tens of kilometres and durations of more than a few days are in geostrophic balance. Equations describing the geostrophic balance can be used to evaluate currents at depth (see, for example, Stewart, 2005).
Developments in Sedimentology, Volume 60 ISSN 0070-4571, DOI: 10.1016/S0070-4571(08)00206-9
Ó 2008 Published by Elsevier B.V.
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the bottom dissipates the energy from the ocean’s interior (Pedlosky, 1987; Cushman-Roisin, 1994). The ocean currents flowing in the boundary layers near the surface and the bottom meet discontinuities in the form of the air/water interface and the sea floor, respectively. The presence of such a frictional boundary affects the current velocity of the ocean’s interior and the turbulent structures. At the contact with the sea floor, the current velocity is forced to diminish to zero, so that the vertical profile of the velocity undergoes shearing (‘‘shear’’ is defined in this case as the way the horizontal component of velocity changes with distance from the boundary), particularly in the water layer close to the sea floor; the thickness of such a layer is a dynamic property of this boundary layer. The turbulent processes transfer and mix the physical properties of the near-boundary flows (momentum, energy and materials such as living organisms, atmospheric gases and sediments), thus giving rise to well-mixed boundary layers at steady state. The surface and the bottom boundary layers are therefore the pathways of communication between, respectively, the atmosphere and the sea, and the sea floor and the sea; exchange of particles, chemicals and organisms takes place in these boundary layers. The simple picture of a stratified ocean that is well-divided into three layers is valid for many regions of the open ocean, but it becomes, obviously, more complex when density gradients, a complex topography, lateral boundary currents or other features are involved. A different approach, based not on a steady description of phenomena along the vertical dimension but on their temporal variability then is required to understand the underlying processes and their morphological, sedimentary and granulometric effects. Moreover, the transport processes near the sea floor that are associated with the sediment dynamics are determined by the turbulent structures that are related to the dimensions (spatial scales) and duration (time scales) of the ocean currents (Le Couturier et al., 2000). The dynamics of the bottom boundary layer itself is driven by the ocean currents: the kinetic energy, which is mainly contained in the large-scale currents and which is ultimately derived from the solar irradiation that drives the global thermohaline circulation,2 is transferred along the so-called turbulent energy cascade down to the smallest dimensions (molecules) and, finally, is converted by molecular dissipation into heat. Turbulent phenomena are characterized by a wide range in dimensions, namely from molecular size to extremely large eddies. It should be noted in this context that eddies, identified as vortexes in fluid dynamics, are described in ocean circulation as areas of closed swirling motion. Their diameters range from a few tens of kilometres (mesoscale eddies) to a few hundreds of kilometres (synoptic eddies), and they can develop both at the surface and at great depth. Turbulence can be seen as a superposition of these eddy structures, which tend to mix the ocean water masses, thus homogenizing them with respect to 2
The thermohaline circulation is due to processes at the sea surface – such as heating and cooling, influxes of fresh water, evaporation, precipitation, river runoff, sea-ice formation and melting – that induce differences in the temperature and salinity – and thus the mass – of the oceanic waters. The thermohaline circulation extends from the upper oceanic layer to the bottom (see also Zenk, 2008, this volume).
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temperature, momentum, particle concentration, etc. The efficiency of mixing is related to the energy transfer by the various processes that take place at different scales: kinetic energy is supplied by large-scale processes, not affected by viscosity, whereas the kinetic energy of flows of ever smaller dimensions is eventually – as mentioned above – converted into heat by molecular (viscous) dissipation. The concept of such a ‘‘turbulent cascade’’ was clearly illustrated by Richardson (1922): Big whirls have little whirls Which feed on their velocity, And little whirls have lesser whirls, And so on to viscosity. Hence, the energy of the turbulent structures that characterize the dynamics of the bottom boundary layer, in combination with their interaction with organisms and sediments, is governed by the turbulent energy cascade. The energy transfer through the various sizes of turbulent eddies is commonly described in terms of wave number (or frequency) spectra: energy production, by means of mechanical instabilities in the fluid, occurs at small wave numbers, whereas viscous dissipation occurs at large wave numbers. After a brief discussion on the concept of spectral windows, we will deal with the description of the general characteristics of the bottom boundary layer, whose analytical investigation constitutes the core of this chapter.
6.2.
SPECTRAL W INDOWS
Spectral windows (Nihoul and Djenidi, 1987) represent typical scales for time and space relative to a certain process under investigation. This concept has been adopted to define hydrodynamic phenomena in marine systems. A schematic distribution of spectral windows in time and space for the most relevant marine dynamical processes is shown in Figure 6.1. In this case, two main spectral windows can be defined, sub-divided by the diurnal time scale: the upper part of Figure 6.1 includes physical processes such as surface and internal tides, internal and inertial waves (the latter are periodic, circular motions related to the Coriolis force, with periods depending on the latitude; at mid-latitudes the inertial period is 17 h). The lower part of Figure 6.1 encompasses processes associated with mesoscale phenomena and synoptic storms, meanderings of fronts (an oceanic front is defined as the transition zone between water masses of different density; frontal currents are driven by the strong temperature and/or salinity gradients that occur in the frontal zone), eddies and gyre circulations developed within the major currents (scales of weeks to months) and at the ocean basin scale (scales of several years). This spectral window can be extended down to the thermohaline circulation at the global scale, which may take 1000 years to overturn completely. A sub-window can be further identified for processes shorter than few hours (uppermost part of Figure 6.1). This includes, among others, some types of waves and three-dimensional turbulent structures (surface waves, Langmuir cells). The processes belonging to this sub-window are commonly considered as ‘‘stirring mechanisms’’
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1 mm 1s
1 cm
1 dm
1m
Molecular processes
10 m
100 m
1 km
10 km
100 km 1000 km
Surface waves
1 min Individual movement
Turbulent patch size
Langmuir cells
1h Inertial/internal waves 1 day
Internal tides
Surface tides
Diurnal Plankton migration
Synoptic storms
1 week Phytoplankton patch 1 month Zooplankton patch 1 year
Coastally trapped waves Fronts and eddies
Seasonal MLD and biomass cycles Gyre circulation
Mesoscale phenomena 10 years
Figure 6.1 The relevant time and length scales of several physical and biological processes in oceans (from Dickey, 1991; with permission from American Geophysical Union).
that increase the contact frequency between water parcels, and therefore increase the mixing rate of these parcels, and consequently of the water properties. As shown in Figure 6.1, spectral windows approach can be adapted also to biological phenomena, that are tightly related with physical processes. The individual movement and the turbulent patch size characterizing the micro-organisms belong to the upper spectral window, while phenomena related with plankton migration, phytoplankton and zooplankton patches have time scales longer than the diurnal one. The turbulent cascade of energy therefore proceeds from the bottom right corner of the figure to the top left corner, where molecular processes are involved in the turbulent dissipation due to viscosity. Following this approach, bottom currents variability (Shanmugam, 2008, this volume; Zenk, 2008, this volume) can have time scales of about 20 days (e.g. Camerlenghi et al., 1997) and persist for a couple of months (e.g. Giorgetti et al., 2003). Together with deep cyclones, that may have lifetimes of 6–9 weeks (e.g. Savidge and Bane, 1999), they belong to the lower part of Figure 6.1. Conversely, turbidity currents were observed to develop in few hours (e.g. Khripounoff et al., 2003; Xu et al., 2004) and therefore belong to the upper spectral window of Figure 6.1. Jankowski et al. (1996) studied the behaviour of the sediment that is brought into suspension due to deep-sea mining, with the primary objective to understand and model the mesoscale processes that are superimposed upon the stable geostrophic component. Fourier analysis of the experimental data collected
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from a bottom boundary layer at a depth of over 4000 m showed the following three components: (1) a slowly varying geostrophic current that could be considered constant over a duration of a week; (2) persistent long waves characterized by an inertial and a sub-inertial period (in the area of the experiment, the inertial period is 98 h); and (3) semi-diurnal tides. Spectral analysis of the velocity field responsible for tidal propagation showed that the main energy peaks were found at near- to super-inertial radial frequencies. The dynamical view of a combination of different components co-existing in the same oceanic region, whether in the surface boundary layer or in the bottom one, is therefore much more appropriate than one dividing the ocean in static layers. The latter is, however, extremely useful for an analytical description of the characteristics of each layer, as we will show below.
6.3. 6.3.1.
CHARACTERISTICS OF THE B OTTOM BOUNDARY L AYER
General characteristics
As pointed out by Kanta and Clayson (2000), the boundary layer just above the sea floor should, strictly speaking, be called ‘‘bottom boundary layer’’ in the shallowmarine environments of the continental shelf (cf. Soulsby, 1983; Grant and Madsen, 1986) and ‘‘benthic boundary layer’’ in the deep ocean (cf. Wimbush and Munk, 1970; Armi and Millard, 1976). It should be noticed that the bottom boundary layer (BBL) can extend over the entire water column on shallow coastal shelves, thus being coupled with the surface boundary layer. In such a case, the turbulent flow is very complex due to the presence of a wide variety of processes, and is beyond the scope of this chapter. In the following, we will hence describe a generic BBL, no matter whether in the deep ocean or on the continental shelf. The BBL described will be considered as separated from the surface dynamics by a well-defined geostrophic interior, with dynamics that results mainly from influences by the current in the ocean interior, by the sea-floor topography, and by sediment-transport processes. In general, a BBL is characterized by currents of a few centimetres per second (in deep water usually not over 20–40 cm s1), but at the contact with the sea floor the current velocity is zero (no-slip condition). This is because the velocity shear in the BBL is particularly strong near the sea floor, and diminishes upwards, to become zero when approaching the ocean interior, where the velocity field is approximately constant and determined by the geostrophic balance. The thickness of the BBL is of the order of 10 m in the deep ocean (depths up to 4000 m; Lueck, 2001), but under high-velocity current conditions it may reach a thickness of 40 m (Kanta and Clayson, 2000) or even involve the whole water column in shallow-water areas, where friction and currents are relatively strong compared to the deep ocean. When the dynamics of the BBL is considered over a nearly horizontal sea floor, far from mid-ocean ridges or other relief forms, a homogeneous fluid,
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vertically uniform in temperature and salinity, may safely be assumed. Kanta and Clayson (2000) have already pointed out that the benthic heat fluxes from the sea floor to the water column are so small (0.1 W m2) that they do not influence the dynamics, and therefore may be neglected. This assumption is supported by observations made over large regions of the oceans more than two decades ago (D’Asaro, 1982). Similar data have also been presented by Jankowski et al. (1996).
6.3.2.
Influence of wind stress
Grant and Madsen (1986) have discussed the characteristics of a continental-shelf BBL driven by wind stress. They deal with the shallow BBL with a spectralwindow approach, in terms of wave propagation at different frequencies, from the surface down to the bottom. They describe the currents as low-frequency wavy motions that can penetrate the geostrophic interior and then influence the BBL, whereas the higher-frequency surface waves remain confined in the surface boundary layer. Their approach to the BBL phenomenon, as representing a response to wind forcing and pressure differences may be applied also to a deep-sea BBL (on which the present contribution is focused), though generally decoupled from the surface dynamics. Surface perturbations can, in fact, also be transferred to depth through internal waves, as these propagate both horizontally and vertically (Gill, 1982; D’Asaro, 1989; D’Asaro et al., 1995).
6.3.3.
Interactions with the sea floor and suspended particles
The abyssal circulation and, as a consequence, the dynamics of the BBL are far from constant, and this should be taken into account when the interaction between currents and sediment is considered. In particular, the HEBBLE experiment (HighEnergy Benthic Boundary Layer Experiment: Hollister and Nowell, 1991) showed a dynamic BBL that plays a significant role in the mixing of the water column and that largely affects the sedimentary processes. A peculiar feature of the near-bottom currents is their sensitivity for the seafloor roughness that is caused by the micro-topography resulting from biological activity and sediment characteristics that are due to current-induced suspension and resettling (Grant and Madsen, 1986). In particular, in the deep ocean, large amounts of sediment can be removed from the sea floor and re-suspended into the water column when the bottom stress associated to strong currents exceeds a certain threshold value that depends on the nature and density of the sediments (Kanta and Clayson, 2000). If the concentration gradient of suspended sedimentary particles surpasses a critical value, it can change the characteristics of the density field, introducing a stratification within the water column and consequently influencing the dynamics of the BBL. Sedimentary particles may remain in suspension because turbulent processes in the water column act against the gravitational force that otherwise would cause resettling of the sediments.
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6.4.
A NALYTICAL APPROACH OF THE BBL
The following summary describes the BBL from an analytical point of view. It is largely based on the reviews by Soulsby (1983) and Lueck (2001). The BBL is usually described to consist of three different layers (Figure 6.2): (1) a bottom Ekman layer that is due to the equilibrium between the Coriolis force, the pressure gradient and the turbulent friction; (2) a very thin (order of magnitude: millimetres) viscous layer, close to the boundary with the solid substratum, where only molecular friction and the pressure gradient are significant; and (3) a transitional ‘‘logarithmic’’ layer between the two layers just mentioned, where the turbulent friction prevails over the other forces and is in equilibrium with the pressure gradient.
6.4.1.
The Ekman layer
Scientific literature regarding atmospheric or oceanic boundary layers commonly refers to a seminal paper published in 1905 by V.W. Ekman. He analytically described the way in which the surface or bottom stress is transferred away from the boundaries towards the oceanic or atmospheric interior. The Ekman equations are an approximation of the equations of motion, and are formally based on the
Ocean interior
Z
Bottom Ekman layer
Z Y
X
Viscous layer Logarithmic layer
Z =0 U
Figure 6.2
Sketch of the vertical structure of the bottom boundary layer.
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balance between the pressure gradient that drives the flow, the Coriolis force and turbulent friction. The Ekman layer constitutes the part of the BBL where these three forces are in equilibrium. The Ekman equations are valid under steady conditions (no dependence on time) and are based on the eddy-viscosity assumption, which can be summarized as follows. In analogy with molecular viscosity, that directly relates the viscous stresses with the mean shear of a flow, eddy (or turbulent) viscosity directly relates the turbulent stresses to the mean shear (Boussinesq, 1877), thus defining the so-called first-order turbulence closure. Whereas the molecular viscosity, , is a parameter of the fluid ( 106 m2 s1 for the water), the eddy viscosity, T, is a property of the flow: it may change in space and with time, and its value can be as high as four orders of magnitude larger than that of . It is very difficult, however, to conform to such an assumption under actual environmental conditions, because turbulence and stratification influence the water column to different degrees.
6.4.1.1. The Ekman solution The Ekman solution is extremely important to understand the dynamics of boundary layers, whether the boundary is the ocean’s surface (surface Ekman layer) or the ocean’s floor (bottom Ekman layer). The way in which this solution is achieved, can be summarized as follows. We consider here for the purpose a coordinate system where the positive X-axis is directed along the bed shear stress, the Z-axis is perpendicular to the bed (Z = 0) and directed to the ocean interior, and the Y-axis is perpendicular to the plane defined by X and Z in a right-hand frame. Due to the choice of the reference frame, the bed shear stress is thus defined as 0 ð X ; Y ÞjZ ¼ 0 ¼ ð X ð0Þ; 0Þ: It may be worthwhile to note here that the shear stress, XZ, is the rate of transfer of momentum in the X-direction across the X–Y-plane (similar definitions can be applied to the other components of the shear tensor, XY and YZ). It basically describes the frictional force that water at one level exerts on the level below. If the lower level is the sea floor, the shear stress exerted on it is referred as 0, and if it is higher than a certain value, it can induce movement of the bottom sediment. The shear stress ( XZ) is defined – from the equations of motion – as the sum of the viscous stresses due to the molecular friction resulting from the mean vertical shear and the Reynolds shear stresses that are due to the fluctuations in turbulent velocity: XZ ¼
dU u 0 w 0 dZ
where ðdU=dZ Þ is the viscous stress that depends on the molecular viscosity ( ) and the mean velocity profile U = U(Z), while u 0 w 0 is the Reynolds stress. The latter is identified as the covariance of the velocity fluctuations in the X and the Z directions (equal to the average of the product of the velocity fluctuations, u 0 and w 0 ), multiplied by the fluid density, . The Reynolds stress equals the shear stress, XZ, except in a very thin layer near the bed where viscous stresses dominate (the
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viscous sub-layer, see above). The eddy-viscosity assumption relates the Reynolds stress with the mean vertical shear: u 0 w 0 ¼ T ðdU=dZ Þ and therefore XZ ¼ ð þ T Þ
dU : dZ
In the following, we will denote XZ and YZ by X and Y, respectively. At a certain distance from the sea floor, ocean currents are affected by the bottom friction, t, and depart from the geostrophic equilibrium that holds in the ocean interior. The value of this departure determines the thickness of the Ekman layer, where a balance among the Coriolis force, the pressure gradient and the friction is established. The equations of motion in the Ekman layer, written in the horizontal coordinates, read as follows: f V ¼ fU¼
1 @P 1 @ X þ @X @Z
1 @P 1 @ Y : þ @Y @Z
The left-hand sides of the above equations represent the Coriolis force (divided by the fluid density, ), while the right-hand sides are the sum of the pressure gradient and the friction, expressed as the variation of the stress in the vertical direction. The Coriolis force is formally given by the cross-product f U, where U is the velocity field and f is the Coriolis parameter, which has only a component parallel to Z, and the amplitude of which is defined as f = 2 sin (where is the angular velocity of the Earth, equal to 2/Trot, where Trot is the period of Earth rotation, and the latitude; at mid-latitudes, f 104 rad s1). The Coriolis force tends to deviate the flow to the right on the northern hemisphere (to the left on the southern hemisphere), and the pressure gradient acts to restore the original direction. One may safely consider the density field in the boundary layer as homogeneous; this is consistent with a pressure gradient that is independent of the height within the layer. Vertical velocities are much smaller than the horizontal current velocity and may be neglected. The boundary conditions of the Ekman layer far above the bottom (Z ! 1), i.e. in the ocean interior, constrain the flow (U, V) to match the geostrophic flow (Ug, Vg), and the bottom stress ( X, Y) to vanish; at Z = 0, in contrast, the no-slip condition U = V = 0 is imposed. Combining the equations of motion with the geostrophic balance (that is determined by the equations of motion at Z ! 1), f Vg ¼ f Ug ¼
1 @P @X
1 @P @Y
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the equations governing the Ekman layer are thus obtained: 1 @ x f V Vg ¼ @Z 1 @ y f U Ug ¼ : @Z The solution to these equations, originally proposed by Ekman (1905), is based on the eddy-viscosity assumption, i.e. X ¼ T dU=dZ and Y ¼ T dV =dZ, where the eddy viscosity, T, is constant and the viscous stresses are negligible. Considering, as an example, a geostrophic flow with Vg = 0, we obtain the following solution for the current in the bottom Ekman layer (Cushman-Roisin, 1994): Z Z=D U ¼ Ug 1 e cos D Z V ¼ Ug e Z=D sin D pffiffiffiffiffiffiffiffiffiffiffiffiffi where D ¼ 2 T =f represents the distance over which the solution approaches the geostrophic flow, i.e. the layer thickness above the sea floor for which frictional forces are much smaller than the Coriolis force, and is therefore a measure of the thickness of the Ekman layer. The vertical profile of the velocity is known as the ‘‘Ekman spiral’’ (Figure 6.3): the horizontal velocity U ¼ ðU; V Þ in the Ekman layer rotates, due to the effect of friction that acts against the current, leftward (on the northern hemisphere) with respect to the geostrophic flow, Ug, with increasing water depth, diminishing to zero at Z = 0. The eddy-viscosity assumption is, however, still far from sufficient to describe the conditions in a fully turbulent oceanic boundary layer in an adequate way. The mass transport in fluid dynamics is defined as the integral of the velocity field along the vertical dimension: in the case of the bottom Ekman layer, this
U(Z1)
U(Z2)
0.2
V
U(Z = 0) U(Z3) Ug
0 0.2
0.4
0.6
0.8
1
U
Figure 6.3 The Ekman spiral. Polar diagram of the current-velocity vector, U, in the bottom Ekman layer. The current-velocity vector is plotted at different distances from the sea floor: Z1 < Z2 < Z3. This figure is drawn for the northern hemisphere ( f > 0); a reversed plot is valid for the southern hemisphere.
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transport is always perpendicular to the geostrophic flow direction (to the left on the northern hemisphere, the opposite on the southern one; Cushman-Roisin, 1994). The Ekman transport generated by contour currents in a stratified ocean on a sloping bottom can influence the flow within the bottom boundary layer. The bottom Ekman transport advects the initial density gradient and creates a buoyancy force that opposes the cross-slope flow. This force gives rise (in absence of external forces) to the exponential decay of the transport from its initial value and in turn, reducing the speed-dependent bottom drag, limits the slowdown of the interior flow (see for details MacCready and Rhines, 1993; Durrieu de Madron and Weatherly, 1994). 6.4.1.2. Thickness of the Ekman boundary layer According to Lueck (2001) and numerous earlier authors (see among others Wimbush and Munk, 1970; Weatherly, 1972; Weatherly and Martin, 1978; Bird et al., 1982; Soulsby, 1983; Thorpe, 1988; Cushman-Roisin, 1994; Jankowski et al., 1996), the thickness of the turbulent Ekman boundary layer (or Ekman height, hE) is proportional pffiffiffiffiffiffiffiffiffiffi to the ratio u*/f, where u* (which is called ‘‘friction velocity’’) is equal to 0 =, and where f is the Coriolis parameter.3 The friction velocity, u*, is a characteristic of the flow and is basically determined by the turbulent fluctuations of the current velocity close to the solid boundary. To estimate the height of the Ekman layer, it is necessary to formulate the friction velocity u* (or the bed shear stress 0) in terms of variables that can be actually measured or calculated, such as the free-stream velocity, a vertically averaged velocity or a reference velocity at some fixed distance from the bottom. The simplest approach to evaluate the bed shear stress, 0, is to consider it as a quadratic function of an appropriate velocity, U: 0 ¼ CD U 2 , where CD is a drag coefficient that depends on the bottom characteristics. To estimate the order of magnitude of hE, Lueck (2001) adopted a geostrophic velocity Ug 0.1 m s1 as the reference velocity (U) and a drag coefficient CD 0.002 (values also used by Soulsby, 1983, and Thorpe, 1988), obtaining hE u* =f ¼ 45 m (with a midlatitude value for f ), corresponding roughly to 1% of the average ocean depth. It is noteworthy that such a formulation results also in u* =Ug ¼ CD½ 1=22; values between 1/20 and 1/30 have often been mentioned (Weatherly et al., 1980; D’Asaro, 1982; Bird et al., 1982; Gust and Weatherly, 1985; Jankovski et al., 1996). Weatherly and Martin (1978), Bird et al. (1982), Thorpe (1988) Cushman-Roisin (1994) and Jankovski et al. (1996) report hE ¼ u* =f for the case of neutral stratification, with the von Ka´rma´n constant, = 0.4, as a proportionality constant. For the case of stratified conditions with a specific value of the buoyancy frequency, N, Weatherly and Martin (1978) suggested the formulation hE ¼ 1:3u* =f ð1 þ N =f Þ1=4 (in a stratified fluid, the buoyancy frequency, N – also known as the Brunt–Va¨isa¨la¨ frequency – is the natural frequency of the vertical 3
As previously defined, a measure of the thickness of the Ekman layer is D. Considering that the eddy viscosity can be expressed as a product of a velocity and a distance, we can choose, respectively, the friction velocity and the Ekman depth, and express T u*D. Therefore we obtain hE = D u*/f.
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oscillations of fluid particles; it is expressed as N 2 ¼ ð g=Þðd=dZ Þ, where g is the acceleration due to gravity and = (Z) is the density that depends on the vertical coordinate, i.e. the height in the atmosphere or the depth in the ocean; in the absence of stratification, N = 0). The Ekman height, hE, that is a measure of the Ekman layer, can thus be considered as the thickness of the water column where the bottom friction is important; for Z > hE, the Coriolis force dominates over friction (geostrophic balance), whereas for Z > ; rough turbulent regime). The coupling between the sediment and the flow characterizes the flow dynamics near the sea floor, thus leaving the role of viscosity insignificant (see below). In
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particular, the thickness of the layer closest to the sea floor is mainly determined by the spacing, rather than the height, of the sedimentary roughness elements (Soulsby, 1983) that cause horizontal variations in the flow properties. The term ‘‘bed layer’’ is also used to identify this sub-layer either in case of a smooth bed where viscosity plays a major role, and in case the roughness dominates the dynamics.
6.4.3.
The logarithmic layer
The vertical profile of the velocity departs from linearity at Z >> : the Reynolds stresses then become much larger than the viscous ones and neither the viscous length nor the Ekman height can adequately describe the dimensions of the flow dynamics. The shear stress in the layer where