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OCEANOGRAPHY AND MARINE BIOLOGY AN ANNUAL REVIEW Volume 20 OCEANOGRAPHY AND MARINE BIOLOGY AN ANNUAL REVIEW Volume 20...

Author: Harold Barnes

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OCEANOGRAPHY AND MARINE BIOLOGY AN ANNUAL REVIEW Volume 20

OCEANOGRAPHY AND MARINE BIOLOGY AN ANNUAL REVIEW Volume 20

HAROLD BARNES, Founder Editor MARGARET BARNES, Editor The Dunstaffnage Marine Research Laboratory Oban, Argyll, Scotland

ABERDEEN UNIVERSITY PRESS

FIRST PUBLISHED IN 1982 This edition published in the Taylor & Francis e-Library, 2005. To purchase your own copy of this or any of Taylor & Francis or Routledge’s collection of thousands of eBooks please go to www.eBookstore.tandf.co.uk. This book is copyright under the Berne Convention. All rights reserved. Apart from any fair dealing for the purpose of private study, research, criticism or review, as permitted under the Copyright Act, 1956, no part of this publication may be reproduced, stored in a retrieval system, or transmitted, in any form or by any means, electronic, electrical, chemical, mechanical, optical, photocopying, recording or otherwise, without the prior permission of the copyright owner. Enquiries should be addressed to the Publishers. © Aberdeen University Press 1982 British Library Cataloguing in Publication Data Oceanography and marine biology. Vol. 20 1. Oceanography—Periodicals 2. Marine biology—Periodicals 551.46′005 GC1 ISBN 0-203-40060-7 Master e-book ISBN

ISBN 0-203-70884-9 (Adobe eReader Format) ISBN 0-08-028460-4 (Print Edition)

PREFACE Once again it has not been possible to enclose all the manuscripts offered for the present volume. Such enthusiasm to obtain publication in the Annual Reviews is greatly appreciated and must be the best indication of their value to marine scientists. The congenial relations with the contributors and the care of the publishers have again made the editor’s task more rewarding than arduous. The help of everybody is gratefully acknowledged.

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CONTENTS

PREFACE

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Interlinking of Physical and Biological Processes in the Antarctic Ocean D.J.TRANTER The Mediterranean Water Outflow in the Gulf of Cadiz M.R.HOWE The Rôle of Bacteria in the Turnover of Organic Matter in the Sea I.R.JOINT AND R.J.MORRIS Particulate Matter in the Oceans—Sampling Methods, Concentration, Size Distribution, and Particle Dynamics W.R.SIMPSON Biology and Ecology of Marine Oligochaeta, a Review OLAV GIERE AND OLAF PFANNKUCHE The Biology of Sandy-beach Whelks of the Genus Bullia (Nassariidae) A.C.BROWN Recent Studies on the Biology of Intertidal Fishes R.N.GIBSON Aspects of the Bioluminescence of Fishes PETER J.HERRING The Biological Importance of Copper in Oceans and Estuaries A.G.LEWIS AND W.R.CAVE

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AUTHOR INDEX SYSTEMATIC INDEX SUBJECT INDEX

38 74 129

197 351 420 472 534

788 894 929


INTERLINKING OF PHYSICAL AND BIOLOGICAL PROCESSES IN THE ANTARCTIC OCEAN* D.J.TRANTER Division of Fisheries Research, CSIRO Marine Laboratories, PO Box 21, Cronulla, N.S.W. 2230, Australia

Oceanogr. Mar. Biol. Ann. Rev., 1982, 20, 11–35 Margaret Barnes, Ed. Aberdeen University Press

INTRODUCTION Present understanding of the remote and inhospitable Antarctic Ocean has developed in an episodic fashion. The pioneering voyages were those of Cook, in the late eighteenth century, and those of Bellingshausen, Wilkes and Ross half a century later. The CHALLENGER, VALDIVIA, BELGICA, GAUSS, SCOTIA and TERRA NOVA played a leading rôle before and after the turn of the century. This was followed by the era of DISCOVERY II when intensive studies were made of Antarctic baleen whales and their environment. After World War II there was the International Geophysical Year (IGY) in which the Soviet vessels OB and VITIAZ were prominent. The momentum of the IGY was continued by the U.S. National Science Foundation vessel ELTANIN. Meanwhile, at laboratories on shore, studies proceeded at a steadier pace, mainly on breeding colonies of birds and mammals and on nearshore communities on the sea floor, in the water column, and embedded in the ice. From this knowledge and understanding has arisen the concept of an Antarctic ecosystem, (Baker, 1954; Currie, 1964; Holdgate, 1967; Knox, 1970; Hedgpeth, 1977). The physical structure of the system has been described by Sverdrup (1933), Deacon (1937), Brodie (1965), Gordon (1971), and Gordon, Taylor & Gingi (1974); plankton productivity by El-Sayed (1968, 1970a,b, 1972), El-Sayed, Mandelli & Sugimura (1964); El-Sayed & Jitts (1973), El-Sayed & Turner (1974), and Holm-Hansen, El-Sayed, Franceshini & Cuhel (1977); the benthos by Knox (1970), Dell (1972), Gruzov (1977), and Knox & Lowry (1977); the ice community by Meguro (1962), Burkholder & Mandelli (1965), Andriashev (1966), Bunt (1966), Buinitsky (1974), and Horner (1974); and the vertebrates by Laws (1977a,b). The references listed here are but an indication of a much wider body of literature. The purpose of the present work is to consider how physical and biological processes interact in the Antarctic Ocean and to discover from the information in hand what forces

Interlinking of physical

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are likely to drive the system. It is directed, in particular, towards “BIOMASS” (Biological Investigations of Marine Antarctic Systems and Stocks), an international collaborative study of the Southern Ocean which is at present in progress. * CSIRO Marine Laboratories Reprint No. 1178.

THE ANTARCTIC ECOSYSTEM The influence of the polar ice cap is all pervasive. Except for those marine birds and mammals that come on shore to breed, there is little life of any kind on the continent itself (Holdgate, 1977); lichens grow on rocky outcrops and, in summer, the snow is sometimes stained by algal growth. Further north, in

Fig. 1.—Distribution of pack-ice about the Antarctic continent in winter (Sept.) and summer (Mar.) 1974: from satellite photography (Budd, 1979).

Oceanography and marine biology

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Fig. 2.—Seasonal variation in Antarctic sea ice cover (Mackintosh, 1972): in winter and spring, the pack-ice stretches halfway to the Antarctic Convergence.

Fig. 3.—Collection of invertebrates from a bottom trawl in the Scotia Sea, WALTER HERWIG January 1978: the collection includes starfish, brittle-stars, pycnogonids, isopods, gastropods, sponges, octopus, and coral.


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Fig. 4.—The Antarctic krill, Euphausia superba: this species forms the staple diet of many Antarctic animals (photograph by U.Kils).


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milder latitudes, as in the islands of the Scotia Arc, there is a mossy carpet which forms a microvegetation, but this does not sustain any known grazing system (Holdgate, 1967). For the main, the Antarctic ecosystem is the sea. As in the Arctic, the sea is separated from the atmosphere each winter by a layer of pack-ice, which reduces the light available for photosynthesis, shelters the water column from the wind, and cuts off surface aeration. The freezing-melting cycle of the ice controls the dynamics of the water column beneath, through changes in temperature and salinity. A community of algae, photosynthesizing at low light intensities, grows within the pack-ice, providing forage close to shore for a benthic ice community and, out to sea, perhaps, an annual bonanza for pelagic grazers. The seasonal expansion and contraction of the ice cover (Fig. 1) constitute a variable annual pulse to which the components of the system are closely synchronized. The edge of the ice represents a wandering coastline in an open sea (Fig. 2). Most of the species in the benthos are endemic to the Antarctic (Dell, 1972). The fauna has been isolated for a long time, the Scotia Arc providing the main avenue for intermixing. Despite the absence of certain taxa, such as crabs, the benthos is generally abundant and diverse (Tressler, 1964; Knox, 1970; Gruzov, 1977; Richardson, 1977). It is in the shallows where moving ice abrades and scours the bottom of the sea, and in areas bathed by waters from beneath the ice-shelves (Dayton & Oliver, 1977) that the fauna is relatively sparse. Brittle stars, starfish, isopods, amphipods, and sponges are well represented (Fig. 3). The pycogonids, once thought to be rare, are abundant. Most of the fish belong to a single family, the Notothenidae. There are characteristics common to a number of Antarctic benthic species. Most are sessile, sluggish, or slow growing (Holdgate, 1967; Knox, 1970), many are relatively large (Andriashev, 1966), and a high proportion brood their young (Dell, 1972). Their biomass is high but their productivity is low. The hub of the system is Euphausia superba, the Antarctic krill (Fig. 4). This representative of the planktonic Euphausiacea (Crustacea) forms the staple diet of a wide range of fish, cephalopods, birds, and mammals, such as the baleen whales, crabeater seal, fur seal, and Adelie penguin. It is taken even by bottom-living trawl fish (Permitin & Tarverdieva, 1972). Krill is generally considered to feed on phytoplankton but there is evidence (Ikeda, pers. comm.) that it feeds equally well on animal food and, on occasions, is cannibalistic. The system is illustrated in a simple way in Figure 5. One of the components of the ecosystem is Man. Note that both baleen whales and their principal food species (krill) are harvested, but not their principal competitor for forage (crabeater seal).


488

LIMITS TO GROWTH The major influence of Man on the Antarctic ecosystem has been through whaling (Gulland, 1976). The stocks of whales are generally much lower now than they were fifty years ago when Antarctic whaling was at its peak (Fig. 6). It is presumed that reduction in whale numbers has lowered the grazing pressure on Antarctic krill. There is evidence that competitors have reaped a benefit from the whales’ demise (Laws, 1977b). Between 1930 and 1960, while the stocks of baleen

Fig. 5.—Simplified Antarctic pelagic food-web, emphasizing the central rôle of Antarctic krill: the crabeater seal is the principal competitor of baleen whales; the Adelie penguin is the principal penguin predator of krill.


15

Fig. 6.—Decline in catch of Antarctic baleen whales over the past 50 years (after Gulland, 1976): the fishery moved in sequence from the larger species to the smaller.

whales declined, the number of fur seals at South Georgia increased 1000-fold (Payne, 1977) (Fig. 7). Adelie penguins show a similar trend (Sladen, 1964). After 10 years of commercial whaling in “The Sanctuary”, an area west of the Antarctic peninsula previously closed to whaling, the crabeater seal


490

Fig. 7.—Increase in the number of fur seals at South Georgia in relation to the decline of whales during the period 1930–1970 (after Laws, 1977b): the upward trend in fur seal numbers is due in part to reduced exploitation.

began to mature at an earlier age (Fig. 8). Note that neither the Adelie penguin nor the crabeater seal had previously been harvested; the response was indirect. Also, the Sei whale arrived earlier in the summer, penetrated farther south (Gambell, 1968), and its pregnancy rate increased (Gambell, 1973) before the species had been subjected to large scale exploitation. The

Fig. 8.—Progressive lowering of age at first maturity of a population of crabeater seals, following ten years of commercial whaling (after Laws, 1977b).

inference is that this species had begun to benefit from an increasing food supply resulting from the mortality of its larger competitors. In short, the components of the system appear to be food-limited. Baleen whales and crabeater seals are essentially pelagic animals and, as such, have unlimited breeding space—the crabeater seal hauls out on the pack-ice immediately above its food supply and whales migrate to warmer waters. The situation is different with many of their competitors. The breeding sites available to the Adelie penguin, for instance, are those ice-free coasts within easy reach of their staple food supply (krill). Such sites are few and very crowded (Fig. 9). Here, breeding space may well constitute the ultimate limit to population size. Cold per se does not appear to be a major hazard for Antarctic animals. Poikilotherms


17

maintain an internal freezing point lower than that of sea water. Invertebrates achieve this by hyperosmotic regulation (Rakusa-Suszczewski & McWhinnie, 1976). Fish which live among the ice have glycoprotein anti-freeze and a high degree of metabolic cold adaptation (Fig. 10) (Feeney, 1974; De Vries & Yuan Lin, 1977, De Vries, 1978). Their problem is heat not cold. They differ from temperate fish which can survive low temperatures in winter and high temperatures in summer. Birds and mammals are well insulated, their problem is how to lose heat in summer on land rather than how to retain heat in winter in the sea. The Antarctic gull solves this problem by using the blood supply to its feet as a heat exchanger (Murrish & Guard, 1977). As far as is known, the main controls on primary production in the sea are nutrients and light. Their availability in the Antarctic for phytoplankton growth is now considered.

NUTRIENTS Figure 11 from the oceanographic atlas of the International Indian Ocean Expedition (Wyrtki, Bennett & Rochford, 1971) illustrates the main features of nutrient distribution in Antarctic waters. Surface nitrate concentration increases with latitude (Fig. 11a), from 10 µg-at.1−1 in the Antarctic. Vertical profiles of inorganic phosphate (Fig. 11b) and silicate (Fig. 11c) between Australia and Antarctica show a similar pattern. The Southern Ocean is relatively rich in nutrients. Concentrations are reduced in summer (Arzhanova, 1974) but there is little evidence that primary production is limited by nutrient availability even though there is little regeneration within a single growing season (Arzhanova, 1976). The source of such high nutrient concentrations is of special interest. When Antarctica and Australia separated and drifted progressively further apart, an uninterrupted passage opened up to the sea (Fig. 12). Within this passage flows the West Wind Drift (Fig. 13), a broad continuous current driven by the westerlies. On its southern flank, adjacent to the Antarctic continent, is a narrow current flowing west, the East Wind Drift, which becomes covered by pack-ice in the winter. Between the West Wind Drift and East Wind Drift is the Antarctic Divergence (Fig. 14) where, under the influence of atmospheric cyclones, warm, deep water, poor in oxygen and


492

Fig. 9.—Colony of Adelie penguins, Estneralda Base, Antarctica, January 1978: there are few suitable sites available for penguin rookeries and these are very crowded.


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Fig. 10.—The model proposed by De Vries & Yuan Lin (1977) to explain the antifreeze in the blood of Antarctic notothenid fishes: the glycoprotein component limits freezing by attaching to the surface of the ice crystal.

rich in nutrient, upwells. This is divided by the annual freezing-melting cycle into a lighter fraction (Antarctic Surface Water) and a heavier fraction (Antarctic Bottom Water), both with a drift component to the north. Antarctic Surface Water submerges at the Polar Front beneath the warmer water on its northern flank and continues further northward as Antarctic Intermediate Water. Thus, the water masses of the Southern Ocean consist of a set of more or less concentric circles whose major anomalies are due to bathymetry, coastline topography, and mesoscale features such as eddies. Sverdrup (1933) and Currie (1964) have drawn attention to an “intermediate return current” which could well act as a feedback mechanism concentrating nutrients in Antarctic surface waters (Fig. 15). Low oxygen


21

Fig. 11.—Distribution of plant nutrients in the Indian Ocean sector of the Antarctic: surface nitrate (a), phosphate profile to 400 m (b), and


496

silicate profile (c) along the transect shown in (a) (dashed line); the units are µg-at.·1−1; concentrations increase progressively towards the Antarctic; after Wyrtki, Bennett & Rochford (1971).

and high phosphate concentration in waters descending sharply north of the Polar Front (Antarctic Intermediate Water) indicate active decomposition and regeneration. Vertical mixing in this area transfers regenerated nutrient across the interface with the southward moving warm deep water and

Fig. 12.—The birth of the Antarctic ecosystem: Australia and Antarctica separate, establishing an open corridor for the Circumpolar Current; after Kenneth (1978).

eventually restores it to the surface by way of upwelling at the Antarctic Divergence. Neshyba (1977) proposed that upwelling from melting icebergs increases nutrient flux through the pycnocline. The experiments of Huppert & Turner (1978) with model icebergs indicate, however, that melt-water


23

Fig. 13.—Circumpolar circulation (after Mackintosh, 1973): the prevailing water movement at lower latitudes is towards the east under the influence of the westerly winds; at higher latitudes, close to the Antarctic continent, the movement is toward the west under the influence of the prevailing easterlies; at the interface between the two lies the Antarctic Divergence.


498

Fig. 14.—Meridional circulation (after Gordon & Goldberg, 1970): the prevailing circumpolar movement of Antarctic waters has meridional components; cold saline water forms beneath the pack-ice and rolls down the continental slope towards the north while cold, low salinity water spreads northward at the surface; these northward flows are balanced by a southward influx at intermediate levels.


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Fig. 15.—The intermediate return current of Sverdrup (1933): north of the Antarctic Convergence (A.C.), mixing is intense, and there is much mortality and decomposition of Antarctic phytoplankton; part of this Antarctic Intermediate water is entrained by deeper water moving south, thus constituting a nutrient feedback loop.


500

Fig. 16.—Mixing and spreading pattern of melt-water from a model iceberg (a small block of fluorescein-impregnated ice in salt-stratified water) (after Huppert & Turner, 1978): melt-water does not rise towards the surface but spreads out laterally at intermediate levels.


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mixes with adjacent sea water and spreads out at intermediate levels instead of rising to the surface (Fig. 16).

LIGHT Of the several somewhat related factors that determine how much light is available for phytoplankton growth, the most obvious is day length. Figure 17 shows the seasonal cycle in day length at various latitudes in the Southern

Fig. 17.—Seasonal variation in day-length at various latitudes: there is little light for photosynthesis in the Antarctic winter.

Ocean. The Antarctic winter appears as a dark wedge interrupting the continuity of photosynthesis. The days get longer and the nights get shorter until, round about the


502

September equinox, a phytoplankton bloom develops (Fig. 18). From then until the equinox in March there is as much light available for photosynthesis at high latitudes as there is at low (Fig. 19). In fact, at the surface of the sea, towards the middle of the day, in the Antarctic summer, there is more light available than the plants can safely absorb (Fig. 20). Photosynthesis is highest not at the surface (except on cloudy days) but 5–10 m below (Fig. 21). The amount of light available for photosynthesis is determined not only by how high the sun is in the sky but how high the autotrophs are located in the water column. Because light attenuates with depth, plant production is limited by the depth of surface mixing (Sverdrup, 1953). Evidence is accumulating that phytoplankton production in the Antarctic summer is determined by the stability of the water column (Hart, 1934; Hasle, 1956, 1969; Saijo & Kawashima, 1964; Fogg, 1977).

Fig. 18.—Seasonal phytoplankton cycle as g C·m−2 in surface waters south of the Antarctic Convergence (after Currie, 1964).

It is not the surface light regime that is the proximate driving force but surface turbulence. In turbulent seas, particles are distributed throughout the surface mixed layer (Fig. 22). As the layer deepens, extending farther into the dark, the mean ambient light intensity available to algae living in this habitat progressively declines (Fig. 23). The main determinant of surface turbulence is wind stress. The Southern Ocean is well known for its rough weather and turbulent seas. Figure 24 from the records of the Australian Bureau of Meteorology (Wearn & Baker, 1980) shows the zonal wind stress. The


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Fig. 19.—Diurnal variation in incident solar radiation in the Antarctic compared with regions further north: despite the lower altitude of the sun, the total amount of light on a summer’s day at 64° S is greater than at 31° S; after Holm-Hansen et al. (1977).

Fig. 20.—Diel surface photo-inhibition in the Antarctic summer (Tranter unpubl. data, WALTER HERWIG, Scotia Sea, January 1978, continuous monitoring by Turner design fluorometer): absorption efficiency declines each day at noon (N).

eastern Indian Ocean and west Pacific sectors (south of Australia) have high and variable winds, which induce very deep (≈600 m) mixed layers (Fig. 25). This is the region where “Subantarctic Mode” water is formed (McCartney, 1977), which retains its isothermal character as it drifts northwards at subsurface levels. Such intensive mixing reduces a broad, nutrient-rich, and (superficially) well-lit part of the Southern Ocean to a virtual desert, compared for example, to nutrient-impoverished subtropical gyres. Many workers


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(e.g. Slawyk, 1979) have commented on the poverty of the area

Fig. 21.—Vertical profile of photosynthesis in the Antarctic summer showing subsurface maximum resulting from surface photoinhibition (after Holm-Hansen et al., 1977, Ross Shelf, February 1972).


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Fig. 22.—Vertical profile through the surface mixed layer in the Subantarctic summer, in relation to the light available for photosynthesis (after Slawyk, 1979): paniculate organic matter (including phytoplankton) mixed uniformly to the bottom of the euphotic layer; PN, particulate nitrogen; PC, particulate carbon; T, temperature; NO3, nitrate.

north of the Antarctic Convergence and wondered why production should be lower there than in waters further north where nutrients are limiting. Whereas there is a strong downward component to water movement in the Subantarctic, in some areas of the Antarctic (e.g. the divergence) there is a strong upward component, leading to upwelling. Associated with these upwellings, but on the downstream side (Beklemishev, 1959) there is high


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Fig. 23.—Relative irradiance (473 nm) in various sectors of the Antarctic south of Tasmania and New Zealand (after Matsuike & Sasaki, 1968): the waters south of Australia and New Zealand are extremely clear presumably because there is little phytoplankton.


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Fig. 24.—Zonal wind stress in dynes·cm−2 (after Wearn & Baker, 1980, source:


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Australian Bureau of Meteorology): the areas of greatest (a) and most variable (b) wind stress are stippled—and are most pronounced in the Subantarctic zone, particularly southward of Australia.

primary production. This appears to be a response not to nutrient enrichment as in low latitude areas of upwelling, but to light enhancement associated with the net upward flux. The time lag between upwelling and production is not fully understood (Beklemishev, 1959; Slawyk, 1979). Upwelling occurs not only at the divergence but also in clockwise eddy

Fig. 25.—Vertical section through the Subantarctic (SA) Zone south of Australia showing the deep mixing process which leads to the formation of subsurface “thermostads” (subsurface isothermal layers) (after McCartney, 1977): this process causes extensive phytoplankton mortality; temperatures in °C).

systems such as those in the Weddell Sea (Foster, Carmack & Neshyba, 1976) to the northeast of the Ross Sea, and in the vicinity of the Kerguelen-Gaussberg Ridge (Gordon, 1971). Walsh (1969), El-Sayed & Turner (1974) and Holm-Hansen et al. (1977) question the “proverbial richness” of Antarctic open-ocean waters. It is usually nearshore that high productivity is observed (Bienati, Comes & Spiedo, 1974).

THE PACK-ICE SYSTEM The layer of pack-ice which forms on the surface of the open ocean around the Antarctic continent adds a new dimension to the Antarctic ecosystem. The pack-ice provides a relatively stable substratum for micro-algal growth which, in turn, radically alters the physical properties of the substratum. Thus the pack-ice is a system within a system, or a seasonal phase of an essentially two-phase ecosystem: the summer phase, predominantly pelagic, the winter phase with strong benthic or “continental” overtones. Although the growth of micro-algae in polar pack-ice has been known for over 100


35

years, the full significance of its rôle in Antarctic productivity has only recently begun to emerge (Meguro, 1962; Burkholder & Mandelli, 1965; Horner, 1974). As icebreakers move through the pack-ice in the summer, this algal layer is seen as a brown discolouration almost everywhere. The chlorophyll concentration of the layer is of the order of 0.1 g·m−2 (≈3.7 g C·m−2, assuming a carbon: chlorophyll ratio of 37:1, Bunt & Lee, 1972), and the area covered by the annual expansion and contraction of the ice is ≈ 15+106 km2. As this is the habitat of Euphausia superba, the rôle of “ice” algae in the krill ecosystem may well be significant. Daily phytoplankton production in the ice-free water column of the open ocean during the growing season (≈150 days) is of the order of 0·1 g C·m−2·day−1 (Burkholder & Mandelli, 1965; Holm-Hansen et al. 1977) i.e. 15 g C·m−2·yr or four times the standing crop of algae in the ice. Because this is released into the water column in a concentrated form, as the pack-ice breaks up and retreats toward the Antarctic continent, it represents forage that is available for grazing at low energy cost. Whether this is done, and how, and by what grazing populations is not yet known. There is, however, historic evidence that this is where the pre-war stocks of baleen whales used to aggregate. The rôle of “ice” algae in the krill ecosystem merits urgent investigation. The most plausible model of the development of algal growth in the pack-ice is that of Meguro (1962) (Fig. 26). In autumn, the air is much colder than the sea and pancake ice forms on the sea surface. This collects and retains the

Fig. 26.—Model for algal growth in Antarctic sea ice (after Meguro, 1962): Pancake ice forms on the surface of the sea in autumn (1) and collects the winter snow (2) which submerges the ice allowing sea water to penetrate at the ice-snow interface (3); within this matrix an algal mat (A) develops, when there is enough light for photosynthesis (4); in summer, the algal mat absorbs so much heat (5) that the “pack-ice” melts and breaks apart liberating the entire algal crop into the water column (6).

autumn snowfall which in the course of time becomes heavy enough to submerge the ice. Thus, high-nutrient sea water is allowed to percolate along the interface between ice and snow forming a porous matrix in which there is enough light and nutrients for photosynthesis. In spring, growth resumes again and absorption of sunlight by the brown pigmented layer so weakens the pack-ice at the algal interface that the top half breaks away releasing the algal crop into the water column. There are other ways in which algal growth can develop in the ice, particularly near to shore, but the areas involved are so much less than the area of the pack-ice that this


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contribution to overall Antarctic production cannot be as great. In addition, the growth close to shore frequently takes place on the lower surface of the ice where there is less light available for photosynthesis. This “epontic community” first described in detail by Bunt (1964, 1966) is adapted for photosynthesis at low light intensities. There may be a “spring epontic community”, which develops on the lower surface of the nearshore ice in spring, and an “autumn epontic community”, which becomes incapsulated in the ice (Hoshiai, 1974). Buinitsky (1974) found that the concentration of algae in the inner layer of the ice is often greater than that in the lower surface layer of newly formed ice, presumably because higher up more light is available. Volume for volume, cell concentrations in the ice throughout the entire period of ice cover (April–January) were 1–2 orders of magnitude higher than in the water column below. Andriashev (1966) and Gruzov (1977) describe graphically how these nearshore layers form. The ice crystals rising from the bottom are colonized by diatoms as they rise and become incorporated in the autumn ice as an autumn algal layer. With the summer thaw, this rich organic mass is liberated into the water and much of it falls out on the shallow bottom, providing forage for detritivores and grazers. Now more light is available for the growth of algae on the sea floor. These survive until the following autumn when ice forms once again and scours the bottom clean. In the Weddell Sea, layers of algae are found within the ice, associated with brine maxima (Ackley, Buck & Taguchi, 1979). These are thought to form in response to seasonal changes in the porosity (and, therefore, the buoyancy) of the ice, where nutrients are available for photosynthesis. Little is known about the nutrient supplies available to “ice” algae. Meguro’s model (Fig. 26) implies that nutrients reach the “snow community” with sea water percolating laterally through the algal matrix. It is likely that, in the long summer days, growth is limited by nutrients rather than by light. By contrast, as salt is excluded from the surface ice, the water column below continues to mix by convective overturn (Gordon, Taylor & Gingi, 1974) and, simultaneously is shaded by algal growth in the ice above. Here, light is more likely to be limiting than nutrients. Burkholder & Mandelli (1965) found production in this shaded habitat to be only one thirtieth that of the algae in the ice. The effect of “ice” algae on the physical properties of the ice in which they are embedded is described by Buinitsky (1974). Shipboard tests of ice strength under compression and under horizontal stress showed that ice with algae in it was only half as strong as ice taken from the same block without any algae, the degree of weakness being a function of the algal concentration. Meguro (1962) has drawn attention to the further weakening which takes place in summer as the pigmented algal layers absorb heat faster than the layers between which they are sandwiched. The effect of algae on the chemistry of the ice is described by Dradovskiy (1974). Nitrite concentration is high in the surface (snow-ice) algal layer of the pack-ice, suggesting that there is growth and remineralization of organic matter in situ. This is consistent with the Arctic studies of Grainger (1974) who showed that nutrient levels in the ice are progressively reduced, presumably as a result of algal growth. The “ice” algae observed by Buinitsky (1974) were alive and, in some cases, actively reproducing. Horner (1974) has reviewed the recent work in this interesting and important field.


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THE HABITAT Within the circumpolar circulation there has developed a unique Antarctic fauna (Hedgpeth, 1969, 1977). Its northern boundary is the Subtropical Convergence, corresponding approximately to 40° S latitude. This varies considerably from season to season, and from one meridian to another. The Antarctic Convergence or Polar Front is a meandering interface embedded in the West Wind Drift where there is active interchange of heat and salt, in part by way of eddies (Gordon et al., 1974; Gordon, Georgi & Taylor, 1977; Joyce & Patterson, 1977; Sievers & Emery, 1978). The productivity here is high, due perhaps to the fact that although the water column is structurally complex it is also relatively stable. Despite a strong meridional component to the predominantly circumpolar circulation, this Polar Front is a major biogeographic barrier. Mackintosh (1937) has shown that some Antarctic macrozooplankton migrate seasonally between surface waters which have a northerly drift component (summer) and deeper waters which have a southerly component (winter). According to Vladimirskaia (1975), 60–70% of the winter zooplankton biomass is concentrated below 200 m. Such a seasonal migration pattern conserves plankton populations south of the Antarctic Convergence and adds to its effectiveness as a biogeographic barrier. The region south of the Antarctic Convergence consists of the “low Antarctic”, within the West Wind Drift, and the “high Antarctic”, within the East Wind Drift. It is in the East Wind Drift and its associated gyres and eddies that Euphausia superba breeds (Marr, 1962), i.e. in the area covered by pack-ice in the winter. The most important of these mesoscale features is the cyclonic Weddell Gyre, formed in the south Atlantic sector by the interaction of the East Wind Drift with the northward sweep of the Antarctic Peninsula and the islands of the Scotia Arc (Beklemishev, 1959; Carmack & Foster, 1974; Deacon, 1976; Foster et al., 1976). According to Marr (1962), krill eggs sink, and hatch at depth, the larvae swimming upward through the warm deep current at the nauplius stage, and reaching Antarctic Surface Water at the first calyptopis stage. Thus they would pass first through waters with a northerly drift component (Antarctic bottom water) and continue their development in waters with a southerly drift component (warm deep water). Spawning has usually been observed to take place at fronts (Makarov, 1972), areas of upwelling (Mackintosh, 1972), for example the Antarctic Divergence, and on the continental shelf (Hempel, Hempel & Baker, 1979). Voronina (1974) has drawn attention to the consequence of eggs falling beyond the point of no return, i.e. beyond the depth where larval yolk reserves can sustain the nauplius throughout its long climb back towards the surface where algal food is available. The expatriation range of the species extends northward into the West Wind Drift, e.g. to the Kerguelen-Gaussberg Ridge and, in the eastern Pacific, into Chilean fjords. In winter, the entire breeding range of the species, and the greater part of its expatriation range, is covered by pack-ice. Hence, the populations are exposed for only a few months in the summer. The period of exposure is shorter in the Pacific than in the Atlantic or Indian sectors, and particularly short to the south of Australia. Fordyce (1977) presents evidence which indicates that evolution of the baleen whales


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took place in mid-Oligocene times when the Circumpolar Current was born. Today they follow the annual retreat of the pack-ice towards the Antarctic coast, and feed on krill newly exposed to predation (Mackintosh, 1973). Pre-World War II stocks of whales moved on, in March, into the East Wind Drift; post-war stocks remain in the West Wind Drift, where, presumably, the forage is now adequate for their needs. As the ice breaks up, the algal crop is liberated into the water column where it is available to the krill and to the summer crop of phytoplankton. It is not yet known whether Euphausia superba make use of it; this constitutes one of the most important questions in the field of Antarctic productivity which need to be addressed. Antarctic krill has generally been regarded as an open ocean species, but it is as much a creature of the ice. The Antarctic ecosystem has been thought to be principally pelagic. Its benthic character (Bunt, 1966) merits wider recognition. Animals over 20 mm in length frequently occur in swarms, detectable by echo sounder and sometimes visible from the surface. According to Marr (1962) these swarms range in extent from several hundred metres to several kilometres. Although the shape of the swarm continuously changes, in Marr’s opinion it behaves as a unit, analogous to a shoal of fish. Very little is known about the physical and biological factors which control swarms of krill. As these may be the ultimate grazing units for baleen whales, seals, and even birds, their ecology is of particular importance. Of the total zooplankton biomass in the water column in winter, 60–70% is concentrated below 200 m (Vladimirskaia, 1975). If Euphausia superba share this habit, it is likely that upwelled waters would bring them to the surface in the summer. According to Elizarov (1971), however, krill swarms coincide not with the actual sites of upwelling but with adjacent areas where waters sink. There is a burst of biological activity followed by a population explosion as krill increase in size and reproduce (Latogurskii, Naumov & Pervushin, 1975). Is the observed sparsity of phytoplankton at upwelling sites due to grazing, then, or to time lags in primary production or both?

CONCLUSIONS The ocean south of the Antarctic Convergence contains a circumpolar community of plants and animals characterized by Euphausia superba, the Antarctic krill. Birds, seals, whales, and other animals use krill for food, and there is evidence (for example increased reproduction rate) that present populations now enjoy a more generous food supply than when the stocks of baleen whales were large. This indicates that the biological resources of the Antarctic must be understood and managed as a multi-species system. Primary production in the Antarctic is limited primarily by light. There is little light for photosynthesis in winter because the days are short and, south of latitude 60°, the surface of the sea is covered by ice and snow. In summer, where the sea is ice-free, the wind stress is frequently so great that the phytoplankton is mixed well beyond the compensation depth where gains by photosynthesis are lost in respiration. There is, however, an extensive growth of “ice” algae living near the waterline of the sea ice, at the interface between ice and snow. Release of this organic matter into the water column when the ice breaks up in summer constitutes a major event in the annual production


39

cycle of Antarctic waters. It is here that the large pre-World War II stocks of baleen whales used to feed. Nutrient concentrations in Antarctic waters are so high that they are unlikely to limit primary production except, perhaps, in the algal matrix of sea ice and in the surface layer of melt-water. The high phytoplankton production sometimes observed in areas of divergence appears less likely to be due to nutrient enrichment than to light-associated factors related to the wind regime. The faunas on either side of the Antarctic Convergence are distinctly different, their separate identities persisting despite the prevailing northerly drift of Antarctic water at intermediate levels. There is evidence that this is due to seasonal vertical migration between surface waters with a northerly drift component, and relatively warm deep waters with a southerly component. Sverdrup’s “intermediate return current” could constitute a feedback mechanism enhancing nutrient concentrations south of the Antarctic Convergence. It is important to abandon the concept that the Southern Ocean is universally productive and to concentrate on processes taking place in the pack-ice and at the Antarctic Divergence and Convergence. Attention should be focused on the light regime, both in relation to the stability of the ater column and to the growth of “ice” algae, on the ecological balance between baleen whales and their smaller competitors such as seals, and on the effect of the emerging fishery for krill upon the recovery of the remaining stocks of baleen whales.

DISCUSSION The main conclusion arising from this review is that the Antarctic ecosystem is foodlimited and the driving force is light. Light, in turn, has two main determinants, daylength and wind stress, the latter influencing light through turbulence and mixing. It would be rewarding to explore the effect of light in the Antarctic ecosystem, not only as a driving force, governing primary production, but also as a control mechanism synchronizing life history cycles with the seasonal production cycle. There is evidence, for instance (Griffiths, Seamark & Bryden, 1979; Bryden, pers. comm.), that the rate of change of day-length in the spring is the factor that stimulates gonad development in the elephant seal, the effect being mediated via the pineal gland which is well developed in polar animals (Cuello, 1970; Piezzi, 1973). Fine tuning of this kind may well be crucial for those Antarctic animals which “can feed where they wish but must breed where they can” (Murphy, 1964). Not only do they have to survive from one summer to the next and store up reserves for breeding but, in addition, they must time their breeding cycle so that their young can gather enough food in their first summer to survive the following winter. The methods used by Antarctic animals to survive the winter are of special interest. Whales commute annually between the Antarctic and the tropics. Some species store food reserves, hibernate or feed on the survivors (Gruzov, 1977). Their reproductive pattern reflects the seasonal variation of their habitat. Many species brood their young (Dell, 1972). Antarctic animals are generally larger than their low latitude counterparts and longer lived (Andriashev, 1966; De Broyer, 1977); they have the chance to breed a


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second time, a property of great survival value in a highly seasonal habitat. They escape in time, as it were, as other (migratory) species escape in space. For the main, the Antarctic winter, however, constitutes a blank in our understanding of the Antarctic ecosystem, particularly the world beneath the ice, the winter habitat of krill. There is urgent need to fill this gap. The problems are logistic. Observations must be made through the pack-ice well out from land, preferably near the ice edge. A suitable platform would be a vessel built like the FRAM overwintering in the ice. Fifty years ago, pelagic whaling got underway in the Antarctic on a massive scale and scientific research was done to monitor its effects. Now we are at the threshold of another exploitative era, the target species including not only whales but also their staple food supply, Antarctic krill. The science of managing multispecies systems is in its infancy. It would be difficult to nominate an optimum harvesting plan for even a simple system involving only whales, crabeater seals, and krill (Fig. 5). Table I suggests that of the various options involving these three species, the least desirable is the one that is most likely to eventuate. This illustrates the need for gathering more information about how the Antarctic system functions, and how its major biological constituents interact with each other and with their common physical environment. If the North Pacific salmon now established in Chilean waters (Joyner, 1980) should start to feed on krill, a major new variable could enter the equation.

TABLE I Harvesting options for Antarctic marine resources in order of their likelihood of maximizing and sustaining the stocks. 1.

Harvest crabeater seals

2.

No harvest of seals, krill or (baleen) whales

3.

Harvest seals+krill

4.

Harvest krill

5.

Harvest whales+seals

6.

Harvest whales

7.

Harvest whales+krill+seals

8.

Harvest whales+krill

The current international collaborative research programme known as “BIOMASS” is directed at such questions. This programme is being coordinated by a working group of specialists under the auspices of SCOR, SCAR and ACMRR. The first major field work (“FIBEX”=First International Biomass Experiment) took place in the Austral summer of 1980/81. The management Commission foreshadowed in the Convention for the Conservation of Antarctic Living Resources will depend heavily on the information and understanding generated by BIOMASS.


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THE MEDITERRANEAN WATER OUTFLOW IN THE GULF OF CADIZ M.R.HOWE Oceanography Department, University of Liverpool, Liverpool, England

Oceanogr. Mar. Biol. Ann. Rev., 1982, 20, 37–64 Margaret Barnes, Ed. Aberdeen University Press

INTRODUCTION If one were to refer to the Meteor Atlas (Defant, 1936), which illustrates so impressively the hydrographic observations that were made by the German Atlantic Expedition of 1925–1927, the tongue-like spread of the Mediterranean water into the North Atlantic would seem to emanate from an apparent “epicentre” near Cape St Vincent (37°N:9°W) rather than from the Strait of Gibraltar itself (36° N:6° W). The maximum values of salinity and temperature that are quoted for this “transposed source” occur at a depth of 1000m and are somewhat in excess of 36.4‰ and 12 °C. This representation of the Mediterranean water intrusion into the Atlantic therefore suggests that, in the initial stages of the outflow there is a preference for the undercurrent to follow a fairly direct and restricted route towards Cape St Vincent with little lateral dispersion, and it is only beyond this point that the more uniform tongue-like divergence continues into the open ocean. This review will concentrate on recent research, particularly during the period 1970– 1980, which has attempted to determine the characteristics of the Mediterranean undercurrent in these initial stages of the outflow, and our attention will, therefore, be confined to the results of work that has been done in a region extending from the Strait of Gibraltar, through the Gulf of Cadiz, to the south of Cape St Vincent (Fig. 1). Previously there had been several expeditions by ships of different nationalities to this area, but they made comparatively few measurements and these were only at certain isolated locations. In 1958 the British R.R.S. DISCOVERY II, and more significantly in 1967 the French research vessel JEAN CHARCOT, undertook what must be considered as the first large scale surveys that were designed to study the overall flow pattern of the Mediterranean undercurrent. A preliminary report of the French cruise by Lacombe, Madelain & Gascard (1968) summarizes the principal aims of that ex-pedition, which were to make hydrographic sections and current measurements at certain strategic positions that would hopefully intercept the main flow paths of the Mediterranean water in its passage through the Gulf of Cadiz. The results were interpreted by Madelain (1970)


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and in a very comprehensive discussion he was able to include data from both the above cruises as well as from several shorter surveys that were made by other ships during the period 1957–1967. Madelain’s account begins with a description of the undercurrent as it emerges from the Strait at about 6°20′ W where he recorded maximum salinities and temperatures of the order of 38.28‰ and 13.25 °C. His main conclusion was that the outflow, after turning sharply to the right due to the effect of the Coriolis force, would thereafter be considerably influenced by the unusual nature of the sea floor topography. This was known to be quite rugged due to the presence of several prominent submarine channels, shelf canyons, and sea mounts. As a result Madelain proposed that the undercurrent would be divided into two main streams. The first remains in contact with the Spanish continental shelf, whereas the other is subsequently formed by the rapid offshore flow of much of the original water mass down two of the shelf canyons. These branches of the flow would then merge at about 8° W (Fig. 1) to produce the main westward moving current that is

Fig. 1.—The Mediterranean outflow in the Gulf of Cadiz (after Madelain, 1970): the principal hydrographic observations and current meter data were recorded along Sections I to VII during the cruise of JEAN CHARCOT in 1967.


522

eventually responsible for the general spread of this water mass into the open Atlantic. There was, of course, no opportunity to investigate any variability in the outflow regime from such a set of widespread and non-synoptic data, nor indeed was it possible to appreciate the degree of inhomogeneity in the Mediterranean water structure. Madelain did note, however, with much interest, a distinct subdivision of the outflow at about 8°30′ W that took the form of separate maxima in the temperature and salinity profiles. For example, he recorded values of 3640‰ and 12.33 °C, which constituted a maximum at the relatively shallow depth of 858 m. These can be compared with a salinity of similar magnitude, 36.60‰, but lower temperature, 11.82°C, which were typical of the values that he observed in the main Mediterranean core at about 1300 m. Similarly Swallow (1969), in an analysis of the Discovery hydrographic sections, was able to intermittently identify Mediterranean water of variable intensity between depths of 675 m and 775 m, where again the salinity values were sometimes found to be equivalent to those normally observed in the deeper layers. These apparent anomalies in the structure of the undercurrent were soon to attract considerable attention. There is, however, no doubt that Madelain’s overall assessment of the outflow regime has been used as the basis and inspiration for the more detailed investigations that have taken place since 1970. These have included similar large scale surveys as well as longer hydrographic time series and current measurements in more advantageous positions. In some cases the interpretation of the results has been accompanied by certain modifications to the flow pattern while, at the same time a great deal of new information has become available concerning the structure and dynamics of the undercurrent. This in turn has initiated an interesting discussion regarding the causes for the high degree of inhomogeneity in the outflow and thereby focused attention on the various mixing processes that may occur within or near the Strait of Gibraltar itself.

OBSERVATIONS IN THE MEDITERRANEAN OUTFLOW 1970–1980 It was the availability of continuously recording salinity-temperature-depth systems that first revealed the true complexity of the Mediterranean outflow structure. This was originally appreciated when measurements were being made, not within the Gulf of Cadiz itself, but in an area about 40 miles to the west of Cape St Vincent. In 1967 German oceanographers, using a bathysonde type of instrument (temperature-electrical conductivity-pressure), quite frequently observed a division of the Mediterranean water into two cores which could be identified as separate maxima in the vertical profiles of temperature and salinity. In addition, the instrumentation was able to resolve a considerable amount of fine structure of varying vertical scales through a large part of the water column. The first reports of these new results by Zenk (1970) and Gieskes, Meincke & Wenk (1970) described some of the physical characteristics of the upper and lower cores which were located at depths of 750 m and 1170 m, respectively. During a ten-day time series at an anchor station Zenk noted a considerable variation in the heat and salt content within the layers occupied by the Mediterranean water, but basically it was possible to distinguish in his mean temperature-salinity profile values of 11.8 °C and


49

36.18‰ for the upper maximum and 11.3 °C and 36.38‰ for the lower maximum. At the same time Howe & Tait (1972) were also recognizing the existence of these separate temperature-salinity maxima in this region. They made an analysis of the fine structure that was associated with both the Mediterranean cores and which was represented in the profiles by a large range of inversions of different vertical scales. The stability of such anomalous features in the ocean was of particular interest. In a comparison of the physical characteristics of the inversions in the two maxima it was concluded that the stability generally increased for the smaller scales, but when considering equivalent scales it was consistently greater for the inversions in the upper core. This tendency was similarly reflected in the mean stability profile for the region where the values at a depth of 600 m were almost twice those at depths of about 1200 m. It was, therefore, evident from the above studies that two separate cores, with quite discernible differences in their physical properties, could be easily distinguished in the Mediterranean undercurrent at a distance of 200 miles from the Strait of Gibraltar, and any explanation of this phenomenon would necessarily entail more precise measurements nearer the source. In the following years various researchers were, therefore, motivated by the thought that observations within the Gulf of Cadiz itself would be particularly productive, especially because the modern continuous recording salinity-temperaturedepth instrumentation was now readily available. As a result several cruises have been made to the area, each concentrating on different aspects of the outflow. In 1971 the F.S.METEOR conducted a large scale current measurement programme while at the same time the R.R.S. DISCOVERY was carrying out some variability studies. In 1973 the R.R.S. SHACKLETON and in 1976 the R.R.S. CHALLENGER both undertook large scale hydrographic surveys, and it can be claimed that all these expeditions made major contributions to our present knowledge of the intrusion of the Mediterranean water into the Gulf of Cadiz. It is probably true to state, however, that these efforts relied a great deal in their planning stages on Madelain’s (1970) observations and interpretation. This review will now consider the results from these cruises, as well as other data and theoretical work, not necessarily in chronological order but rather when appropriate to the discussion of such topics as the overall flow pattern, mixing, variability, and velocities in the undercurrent.

HYDROGRAPHIC SURVEYS R.R.S.SHACKLETON CRUISE 1973 This was a significant cruise in two respects. First, it provided an opportunity to confirm some of the principal features in Madelain’s (1970) outflow pattern while using a modern salinity-temperature-depth system. Secondly, by collecting water samples at certain depths and applying the latest analytical techniques it was hoped to establish for the first time the general nutrient salt distribution throughout the region. The results of the combined analysis of both the physical and chemical surveys proved to be particularly interesting. Four sections of hydrographic stations were worked, with three positioned along the submarine canyons which are a prominent feature of the bottom topography


524

between the Strait of Gibraltar and Cape St Vincent. These were selected primarily because Madelain had decided that they were exerting a considerable influence on the flow paths of the Mediterranean water. In fact the stations in these sections coincided very closely with those of the Jean Charcot cruise which are shown in Sections IV, V, and VII of Figure 1. In general, the large scale characteristics of the outflow, as recorded by both the Jean Charcot survey in 1967 and the Shackleton survey in 1973, appear to be virtually constant. Ambar, Howe & Abdullah (1976) have given a full account of the chemical and physical data from the Shackleton cruise and Figure 2 from that paper, together with Figure 3 (Madelain, 1970) have been reproduced here to demonstrate the apparent long term consistency in the temperature and salinity values along Section IV (Fig. 1). Furthermore, by reference to the original publications, the reader can


51

Fig. 2.—Distribution of salinity and temperature along a section coincident with IV in Fig. 1 (after Ambar et al., 1976).


526

Fig. 3.—Vertical sections of salinity and temperature along Section IV (see Fig. 1) (after Madelain, 1970).

make a similar comparison of the respective sets of observations that were made along Section VII (Fig. 1) during these cruises. The sudden turn to the right as the outflow leaves the Strait and its subsequent movement along the Spanish continental slope with little offshore spreading, is easily recognizable. It is perhaps not surprising that with such a substantial and continuous outflow of this kind the temperature-salinity values near the Strait should remain so constant. The less predictable fact that the offshore spreading appears to remain so restricted must, however, also contribute greatly to the consistency of the flow characteristics. The first difference in interpretation arises when Ambar et al. (1976) consider the large scale structure of the undercurrent. They refer to the layer of Mediterranean water that settles out at depths above 900 m as the “upper core” and then proceed to treat it as a separate water mass as it moves westward. This was justified to a large extent by the extraordinary variation in the nutrient values that was revealed in the chemical analysis of the water samples. Consequently, Howe, Abdullah & Deetae (1974) have suggested that the source of this shallower Mediterranean core might be more appropriately associated with depths within the Strait of Gibraltar that are significantly different to those from which the main outflow usually descends. This conclusion was


53

based on the combined study of the physical and chemical data which essentially showed that whenever this upper core was sampled in the Gulf of Cadiz, there were anomalously low concentrations of the nutrient salts nitrate, silicate, and phosphate. So much so that when the chemical samples coincided with the actual upper temperature-salinity maximum at depths between 600 and 800 m, the nutrient values were not only less than those of the usual source of the outflow (that is, between depths of 150–300 m in the Strait) but they were also less than the values that can be normally observed in the layers of the Atlantic through which any mixing would occur. In other words, there seemed to be no way of accounting for these abnormally low concentrations in the upper maximum by reference to either the source water or the ultimate mixing environment. Howe et al. (1974), therefore, decided that an explanation of this chemical anomaly might provide the means of identifying the source of the upper Mediterranean core. The relevant information is conveniently presented in Figure 4 where the salinity-nutrient relationships have been plotted for the various water masses that are likely to be involved in the intermixing. These include a profile for the source waters, which were sampled just within the Mediterranean Sea itself; the computed regression lines for the samples that were collected at the standard depths of 600 and 800 m in the Atlantic where the upper core is to be found; and the average salinity nutrient values for the normally undisturbed layers in the Atlantic through which the upper core would have to mix in its initial stages of intrusion. The authors then focus attention on the high salinity (low nutrient) values labelled X and Y (Fig. 4) which represent the properties of the upper Mediterranean core as it was observed in its most original or pure state. They argue that when the possible mixing lines between the various layers of the water column in the Mediterranean Sea and the Atlantic are considered (see caption to Fig. 4) the points X and Y, with their extremely low nutrient concentrations, cannot be produced by outflow water emanating from the usual depths of 150–300 m, nor from a completely mixed water column between the surface and 300 m. The only suitable source water, with concentrations most appropriate to the likely mixing lines (broken in Fig. 4), is that from depths above 150 m, and more precisely from a depth of 110 m. This water would have to pass through the Strait with virtually no vertical mixing in order to account for the nutrient concentrations that were observed in the upper core. This interpretation, however, is not entirely compatible with a theory that Siedler (1968) proposed, and which Zenk (1970) applied to his earlier observations of the double temperature-salinity maxima off Cape St Vincent. Siedler postulates that such a double maxima can be readily generated by the tidal mixing processes within the Strait of Gibraltar. In his model he uses published data to represent approximately the current and salinity structure, and by superimposing the effects of a parabolic velocity profile, a surface tide and an internal tidal boundary wave, he produces a frequency distribution of water types with two preferred salinity values. In Zenk’s application of this theory it is assumed that these two Mediterranean water types will emerge from the Strait and, having mixed with North Atlantic Central Water at an initial depth of 500 m in similar ratios, will then form the two distinct cores in the outflow which he had observed to the west of Cape St Vincent. Unfortunately it is the degree of tidal mixing within the Strait, which is so implicit in Siedler’s model, that is inconsistent with the chemical interpretation suggested by Howe et al. (1974). In spite of this, the information being accumulated from these studies strongly


528

indicated that a Mediterranean upper core can be formed, and thereafter it would survive as a separate water mass at distances of at least


55

Fig. 4.—The regression relationship (solid lines) for salinity-nutrients in the Atlantic at 600 and 800 m and the salinity-nutrient profile (heavy line) at a station within the Mediterranean Sea: average salinitynutrient content of the Atlantic (A) and Mediterranean (M) water columns between the following depths: A1:0–800 m; A2:0–600m; M1:0–150m; M2:50–150 m; M3:0–300 m; M4:150–300 m; for the relevance of the broken lines and points X and Y see p. 43; (after Howe et al., 1974).

200 miles from its probable source. Accordingly Ambar, Howe & Abdullah (1976) proceed to examine their Shackleton hydrographic data with the principal aim of identifying, wherever possible, this kind of stratification in the outflow. They start by again referring to the canyon section in Figure 2, and discuss the ultimate fate of the ‘anomalous’ mass of Mediterranean water near the bottom of Station 92, where a salinity of 36.88‰, a potential temperature of 12.95 °C and a potential density of 27.88 were recorded at a depth of 870 m. More typically such densities are observed in the deeper water further offshore where, for example, at 1200 m in Station 98 the salinity was 36.52‰ and the potential temperature was 11.63 °C. They then pose the question: does part of this water flow down the canyon and thereby initiate the main deep core of the outflow or does it remain isolated at the shallower depth and make no contribution to the colder and less saline water mass that is to be observed between Stations 97 and 98? Ambar et al. (1976) point out that the salinity-nutrient relationship at a depth of 800 m in Station 92 is closely correlated with that of the upper core water mass in the other sections further to the west (Fig. 5). They, therefore, propose that because of the prevailing ambient conditions this water mass will lose salt fairly rapidly due to the large horizontal salinity gradients, while maintaining its relatively high temperature, and ultimately it will achieve a potential density of about 27.49. This will then account for its presence at a depth of 793 m in Station 71 (Fig. 5) with a salinity of 36.52‰ and a potential temperature of 13.48 °C. In contrast, the lower Mediterranean core at a depth of 1270 m in Station 73 of this same section has a salinity of 36.59‰, a potential temperature of 12.09 °C, and a potential density of 27.82. Finally, and to emphasize the amount of stratification that appeared in these Shackleton sections, it was possible to discern yet another temperature-salinity maximum in the profiles. This occurred at the relatively shallow depth of 400–500 m but only in the inshore stations at the head of the shelf canyons (Figs 2 and 5). If these profiles had been examined in isolation and without reference to the properties of the surrounding waters, it would have been tempting to speculate that this maximum was due to a further subdivision of the shallow core of the Mediterranean undercurrent. When this feature was, however, considered in relation to the general temperature and salinity distribution, it seemed more likely that it had been caused by the sinking of the surface layers which, according to Ambar et al. (1976), had been induced by the winter cooling of the shelf waters and facilitated by the presence of the canyons. If one were to summarize the evidence that has been presented up to now, it would be reasonable to assume, first, that there is indeed a strong preference for the outflow to maintain contact with the Spanish continental slope in its progress towards Cape St


530

Vincent. Secondly, the long term constancy in the downstream values of the temperature and salinity denotes very little variation or patchiness in the overall flow regime. The most significant new development in the study of the undercurrent has been to establish the existence of a separate Mediterranean water upper core, which is characterized by a consistently higher temperature than that of the main lower core, and abnormally low nutrient concentrations. It can be initially identified fairly near to the source in Section IV (Fig. 1) and then clearly distinguished in Section VII. This presumably ensures its permanent presence in the area and, therefore, readily accounts for the upper temperaturesalinity maximum in the profiles that were originally observed in the open Atlantic to the west of Cape St Vincent (Zenk, 1970; Howe & Tait, 1972).


57

Fig. 5.—Distribution of salinity and temperature along a section coincident


532

with VII in Fig. 1 (after Ambar et al., 1976).

R.R.S. CHALLENGER CRUISE 1976 It was acknowledged that the Shackleton 1973 cruise was somewhat exploratory and so the same scientific personnel considered it important that the Challenger 1976 survey should at least confirm some of the previous concepts concerning the structure and the general flow pattern of the Mediterranean undercurrent. The two cores were now believed to be a permanent feature of the outflow and so the zone in which the separation occurs needed to be identified, as well as their preferred routes through the Gulf of Cadiz. In addition, their rate of mixing and any significant changes in the thermohaline characteristics would be of interest, together with estimates of the velocities along the flow paths. With these aims a very comprehensive grid of stations was worked from Cape St Vincent to the entrance of the Strait of Gibraltar, and again this survey included a series of sections along the prominent canyons. The results have been published by Ambar & Howe (1979a, b) and later they will be discussed in some detail under more appropriate headings. It was, however, immediately obvious that there was no significant intermittency in the outflow pattern and that the maximum temperature-salinity values associated with both cores were easily discernible throughout the entire region. For example, in a north-south section off Cape St Vincent (Fig. 6) which was, of course, at the western extremity of our area of interest, the large scale stratification was plainly visible. Here Ambar & Howe (1979a) recorded maximum values of 36.57‰ salinity and 12.18°C in the lower core at a depth of 1315 m with a potential density of 27.83, whereas in the upper core the maxima at 755 m were 36.43‰ and 13.07°C with a potential density of 27.52. In addition, they also noted the presence of an even shallower temperature-salinity maximum in the profiles of the inshore station at the edge of the continental slope, where the values at 580 m were 36.03‰ and 12.90°C with a potential density of 27.24. This water mass is something of an enigma because although in the previous Shackleton survey, which was made during April 1973, this feature had been attributed by Ambar, Howe & Abdullah (1976) to winter cooling and the subsequent sinking of water from the shelf, the appearance in August of a similar subsurface layer makes their explanation more dubious. Instead, Ambar & Howe (1979a) raise the possibility that this might be a further shallow subdivision of the undercurrent. The idea gained support when they were able to identify the probable point in the outflow, at about 7°W and near Section III (Fig. 1), where the separation of the two main cores occurs. Figure 7b shows this section and it can be conveniently compared with a section through the undivided outflow (Fig. 7a) as it leaves the Strait. In Figure 7b Ambar & Howe assign maximum values of 37.42‰, 13.16°C, and a potential density of 28.28 to the main lower core at 756m; 37.07‰, 13.72°C, and a potential density of 27.88 for the upper core at 650 m; and values of 36.65‰, 13.61°C, with a potential density of 27.57 at 510 m in the inshore station where the shallower subdivision of the upper core might be occurring. If so, this shallow water mass would then have to follow closely the 500 m isobath in order to appear as the anomalous maximum at 580 m in the section south of Cape St Vincent (Fig. 6).


59

THE FORMATION OF THE TEMPERATURE-SALINITY MAXIMA IN THE OUTFLOW The hydrographic surveys have demonstrated quite convincingly that the double, and perhaps triple maxima, are formed very near the Strait and, therefore, it is now more readily accepted that the bottom topography in the


534

Fig. 6.—Distribution of temperature and salinity along a section running south


61

from Cape St Vincent (after Ambar & Howe, 1979a).

Gulf of Cadiz can no longer be regarded as being wholly responsible for the inhomogeneity in the structure of the outflow. As well as the evidence from the Shackleton and Challenger data, this hypothesis is also supported by Zenk’s (1975a) interpretation of some observations that were made very near to Section III (Fig. 1), in which he represented the outflow in terms of three maxima. Zenk’s schematic outline of these different water types is shown in


536


63

Fig. 7.—Distribution of temperature and salinity in the outflow (a) as it leaves the Strait at about longitude 6°18′ W and (b) through a section near III in Fig. 1 at about 6°55′ W; (after Ambar & Howe, 1979a).

Figure 8 where the properties of Layers 1 (salinity=37.50‰, density= 28.25) and 3 (salinity=36.90‰, density=27.60) can be related to the corresponding layers in Figure 7b of the Challenger data. Whereas Ambar &


538


65

Fig. 8.—Schematic section through the Mediterranean outflow near III in Fig. 1 showing three different water types (1, 2, and 3) (after Zenk, 1975a).

Howe (1979a) chose to subdivide Layer 3, Zenk showed by means of repeated CTD casts (conductivity-temperature-depth) through Layers 1 and 2 that in this particular section Layer 2 appeared in the temperature-salinity diagrams as a discernible maximum which remained isolated from a highly variable mixing part of the water column. In Figure 9 the more stable core of Layer 2 is at a depth of 550 m with a salinity of 36.15‰ and a density of 27.45.

Fig. 9.—T-S diagram of repeated CTD records through the centre of the outflow in a section near III in Fig. 1 (after Zenk, 1975a).

Even although the interpretations of these different sets of observations have not been entirely consistent, the complexity of the structure of the outflow is now better appreciated, and an explanation of the origin of the multiple maxima will require a more comprehensive model than that suggested by Siedler (1968). The inhomogeneity in the undercurrent can probably be attributed to mixing processes occurring within the Strait of Gibraltar, and it is now known that these can be influenced by a number of short term factors, as well as the more obvious large scale seasonal changes. Defant (1961) has discussed the tidal response and the internal oscillations that take place in the transition boundary between the inflowing surface Atlantic water and the Mediterranean outflow. Significant variations in the rate of the Atlantic inflow have been attributed by Lacombe (1961) and Crepon (1965) to changes in the mean atmospheric pressure over the western Mediterranean. There have been numerous measurements of internal wave activity, either of semi-diurnal period (Frassetto, 1960; Lacombe, Tchernia, Richez & Gamberoni, 1964) or short period (Ziegenbein, 1969, 1970). Boyce (1975) has included some of these results in a theoretical discussion of the two layer flow regime within the Strait in which


540

he advocates the frequent presence of steepening long internal waves and internal shock waves. Apart from this kind of internal activity, another possible mixing process was revealed after Ambar & Howe (1979a) had analysed a time series of temperature and salinity fluctuations in the boundary layer between the North Atlantic Central Water and the Mediterranean outflow in the section represented in Figure 7b. By applying the criteria outlined by Pingree (1972), which compares the slope of the regression line that is obtained by correlating the temperature and salinity fluctuations at fixed levels with the ratio of the vertical gradients of the mean temperature and salinity at the same level, they concluded that since these quantities were in good agreement then the fluctuations must be the result of either internal wave activity or vertical mixing. The time series was dominated by a semi-diurnal period, and it was recorded in a flow regime in which Zenk (1975b) has reported a mean monthly speed of 68 cm·s−1 and a significant 12.4-h period in both the velocity and temperature records. In view of the obvious slope of the surface of discontinuity between the water masses in this section, Ambar & Howe (1979a) considered an application of the Margules formula:

which, for relevant values of

and

gave in this case

. h/L is the gradient of the slope, i.e. ratio of vertical height (h) to the horizontal distance (L); f is the Coriolis factor; g is acceleration due to gravity; p and p′ are the densities of upper and lower water masses, respectively; and U and U′ are the mean velocities of the upper and lower water masses, respectively. In Figure 7b, with L=20 km, a change in h of 20 m would be easily achieved by a change in the differential velocity of 10 cm·s−1 and because of the large vertical temperature and salinity gradients a ∆h of 20 m would then readily account for the range of the semi-diurnal oscillation that was observed in the time series. The discussion then turned to the possibility that such a mixing or entrainment process might also occur within the Strait where there is a similar slope of the isohalines (Fig. 7a). Here again a depression of the boundary surface, which may be induced by quite reasonable changes in the differential velocity, will allow shallower and warmer water to be entrained and then released as part of the normal outflow. There are, therefore, a number of modulation processes that may be responsible for the double, or perhaps a multiplicity of temperature-salinity maxima in the structure of the undercurrent. Recent observations certainly testify to the degree of inhomogeneity but as yet there is not sufficient evidence to indicate which of the dynamical effects may be the most important.

MIXING IN THE OUTFLOW In a theoretical study of several overflow situations Smith (1975) developed a stream-


67

tube model for bottom boundary currents which he was able to apply to the Mediterranean outflow by making use of Madelain’s (1970) data. In general, Smith concludes that the dynamics of this undercurrent would be mainly affected by friction near the source, and by entrainment further downstream. Furthermore, the external stratification and entrainment would limit its descent to 1200 m and, when clearing the continental slope, the volumetric transport will have increased from about 1.0×106m3·s−1 to 10×l06 m3.s−1. From an observational analysis of small scale thermohaline structures in the deep ocean, Pingree (1972) believes that the outflow must have mixed quite vigorously to depths of at least 1600 m on the continental slopes in the Gulf of Cadiz before spreading horizontally along isopycnal surfaces into the open Atlantic. Thereafter evidence of much deeper mixing to depths as great as 3000 m has been provided by Worthington & Wright (1970). Any attempts to quantify the rate of mixing of the undercurrent within the Gulf of Cadiz have been rare. Zenk (1975b) assumed an initial volumetric transport of 1.0×106 m3·s−1 leaving the Strait, with a water type salinity value of 38.4‰, which would then proceed to mix with North Atlantic Central Water of salinity type 35.6‰ along the normal mixing lines that are usually represented on a temperature-salinity diagram. In a two-box assessment of the cascade water budget Zenk estimated that there would be a transport of 1.75×106 m3·s−1 through Section III (Fig. 1), and that this would be composed of 54% Mediterranean water and 46% Atlantic water. As the outflow moved westward, either by following a northern route around the Spanish shelf or by flowing down several southwesterly submarine channels (Fig. 12), he calculated that this transport of 1.75×106 m3·s−1 would eventually contribute 60% to the product of a mixture with Atlantic water (40%) to produce a total outflow from the Gulf of Cadiz of 2.92×106 m3·s−1. From direct current measurements and observations of the salinity distribution in these channels, Zenk, however, calculated a total outflow transport of 2.03×106 m3·s−1. He attributed the difference to an inadequate knowledge of the mean current directions and the salinity distribution, as well as the likelihood of losses by other routes which were not monitored. By applying a quite different technique Ambar & Howe (1979a), having established the preferred flow paths of the two main Mediterranean cores, made separate estimates of the mixing rates of each of these water masses. The upper core was assumed to be a mixture of Mediterranean Water and the North Atlantic Central Water which occupies the depths between 200 m and 900 m. The main lower core was regarded as an admixture of Mediterranean water, North Atlantic Central Water and North Atlantic Deep Water which occupies depths between 400 m and 1500 m. It was necessary to decide on the relevant indices to represent these water types before the appropriate “mixing triangles” could be constructed on a temperature-salinity space. The theoretical implications and application of this technique have been fully discussed by Mamayev (1975). After a detailed analysis of their data Ambar & Howe (1979a) designated specific values to the particular sources of Mediterranean water in the outflow that supplies the two cores. Their source water type for the upper core was allocated a salinity of 37.18‰ and a potential temperature of 14.02°C with corresponding values of 38.12‰ and 13.35°C for the lower. These can be compared with Zenk’s (1970) values of 37.2‰, 13.0 °C, and 38.2‰, 13 °C, respectively which he used in his original model to account for the double maxima. The percentage


542

value for the amount of Mediterranean water that was associated with each core were then calculated independently from the appropriate mixing triangles. Figure 10 shows a comparison of the results that were obtained in a section coincident with III (Fig. 1) and in a section that was made to the south of Cape St Vincent. The maximum percentage of Mediterranean water in the cores was, of course, deduced from the actual temperaturesalinity maxima that were observed in each section, and the general distribution of these values

Fig. 10.—Percentages of the Mediterranean water in the upper (Mu) and lower (Ml) cores in section (a) along III in Fig. 1 and (b) south of Cape St Vincent (after Ambar & Howe, 1979a).

throughout the region is represented by Figure 1.1. The authors concluded from this, that of the two cores the lower appears, to have mixed more vigorously in its descent from the Strait, because at this initial stage it only constitutes 70% of the mixture compared with a value of 90% for the upper core. Thereafter there seems to be a similar rate of mixing in


69

both layers of the outflow as it proceeds towards Cape St Vincent. Finally, Ambar & Howe (1979b), without the benefit of any direct current measurements, deduced from geostrophic considerations a volumetric transport of 1.51×106 m3·s−1 for the undercurrent as it leaves the Strait. This is in reasonable agreement with other estimates of the outflow which include 1.65×106 m3·s−1 (Defant, 1961), 0.66 to 1.38×106 m3.s−1 (Lacombe et al., 1964), 0.72 to 1.57×106 m3·s−1 (Lacombe, 1971), 1.60×106 m3·s−1 (Bethoux, 1979) and the previously mentioned value of 1.0×106 m3·s−1 (Zenk, 1975b) which he used in his assessment of the cascade water budget. At about longitude 7°30′ W the transport had increased to 2.74×106 m3·s−1. Although this trend appears to have been reversed in the section near Cape St Vincent (2.60×106 m3·s−1), it was noted that the estimates in this area would be subject to much larger errors due to the tendency for an offshore divergence in the flow, which would not


544

Fig. 11.—Isopleths of the maximum percentages in the (a) upper (Mu) and (b) lower (Ml) cores as they were observed through the Gulf of Cadiz (after Ambar & Howe 1979a).

have been monitored by the observations. Near the source, an estimate of the entrainment factor, based on the concept of mass continuity in the downstream direction, gave values between 0.02 km and 0.05 km and these were comparable with the predicted optimum value of 0.05 km from Smith’s (1975) stream-tube model.


71

CURRENT VELOCITIES IN THE OUTFLOW Prior to 1970 there had been very few successful attempts to record the current speeds in the outflow and invariably these were of short duration, usually lasting just one or two days. The most relevant to this discussion were two sets of measurements, using quite different techniques, which were made at the opposite extremes of the region. Bøyum (1963) has reported results from three current meter moorings that were deployed near the Strait between 6°20′ W and 6°45′ W. Here he recorded mean speeds of 80–90 cm·s−1 over a 24-h period. In contrast, Swallow (1969), who was working off Cape St Vincent with neutrally buoyant floats, measured a westward flow of 20–30 cm·s−1 during a 2-day experiment near 36°30′ N:8°45′ W. This was associated with the main high salinity core at a depth of 1080 m. Furher offshore and centred at about 36°20′ N, he observed over a period of 7 days an anticlockwise gyre at a depth of 1400 m with speeds of 10–20 cm·s−1. This interesting effect occurred well below the salinity maximum. Even further south at 35°47′ N, and after two days of tracking, the speed at a depth of 1260 m was estimated to be 2.4 cm.s−1 towards 130 °T. It, therefore, seems likely that there may be a high degree of variability in the flow pattern of the intermediate waters in this area. More recently Madelain’s (1970) current meter measurements, although widely distributed (Stations C1–C10 in Fig. 1), were also of short duration, usually lasting just 24 h or even less. Near the Strait between 6°20′ W and 6°45′ W speeds of 100 cm.s−1 were recorded, and a substantial flow rate was maintained as far as Section III where the velocity was 50 cm·s−1. Further west the current meter moorings were positioned in the submarine canyons along Sections IV and V where there was a southwesterly flow of about 25 cm·s−1, but at the offshore mooring (36° N) this had decreased to 100

20

3.2–3.9

90

5.2–8.2

360

>100

1800

>100

Banoub& Williams (1972)

14C glucose

Kinetic

50

250

14C aspartate

Kinetic

50

417

U.S.A., Estuary

14C glucose

Kinetic

Surface 0.01–0.93 Crawford et al. (1974)

Canada, Fjord

14C glucose

Kinetic

Surface 0.2–1.75

Sibert& Brown (1975)

California,

14C amino acids

Tracer

25

1.12–14

Williams et al. (1976)

100

7.8–>100

N. Pacific

Coastal Baltic

14C amino acid

Tracer

Surface 0.6–4.2 100

North Sea

14C alanine

Tracer

Seki et al. (1972)

Dawson & Gocke (1978)

2.7–13.6

Surface 0.91–97

14C glucose

Surface 0.95–18.9

14C acetate

Surface 0.34–69

14C lactate

Surface 2.1–104

Billen et al. (1980)

It is assumed in studies of organic turnover that uptake of labelled substrate is entirely by


576

heterotrophic microbes and that phytoplankton are not capable of utilizing dissolved organic compounds at the concentrations present in the sea. Autoradiographic studies have generally shown no uptake of dissolved organics by phytoplankton (Munro & Brock, 1968; Horner & Alexander, 1972; Paerl & Goldman, 1972; Paerl, 1974) but there have also been reports of low rates of uptake by autotrophic organisms (Hoppe, 1976, 1977; Pollingher & Berman, 1976). Hoppe (1978) concluded, however, that heterotrophic activity of algae was very low and would have a negligible effect on measurements of bacterial production, but clearly, individual algal cells could benefit considerably from the assimilation of useful organic compounds which they did not have to synthesize. Other evidence for the importance of bacteria in heterotrophic processes in the sea comes from size fractionation. Williams (1970) found that 80% of added substrate was incorporated by small organisms which passed through 8 µm mean pore size filter and 49% was in organisms passing through 1.2 µm pore size. Derenbach & Williams (1974) found 80–97% of the heterotrophic production passed through 3 µm mean pore size filters and Larsson & Hagström (1979) found 95% passed through a 3 µm NucleporeR filter. This association of heterotrophic activity with the smallest size fractions suggests that bacteria, rather than protozoans or algae, are involved in the turnover of organic matter in the sea. MINERALIZATION In terrestrial ecosystems, the amount of primary production consumed by herbivores is much less than is generally assumed in marine environments; Gray & Williams (1971) estimated that only 6–40% of the autochthonous organic matter is eaten, the rest reaching the soil where it is available to decomposing microbes. Terrestrial microbes are, therefore, traditionally considered as re-mineralizers of organic matter. This rôle has also been ascribed to marine bacteria. Data obtained from 14C organic studies in the last decade cast doubts, however, on this assumption. Williams (1970) first pointed out that, in contrast to data obtained using bacterial cultures, a relatively small proportion of organic matter taken up by bacteria in the sea is respired and the major part of the carbon is assimilated into cellular material. That is, bacteria appear to be very efficient utilizers of organic substrates, which results in high growth yields, but are poor mineralizers. There is now considerable evidence in the literature to support Williams’ initial observation (Table III); generally, the percentage of glucose taken up which is respired is about 30%, giving a growth yield of about 70%. The growth yield from individual amino acids varies, with valine and leucine giving the highest growth yields (about 90%) and alanine, glutamate, asparate, serine, glycine, and arginine giving lower growth yields of 60–80% (Williams, Berman & Holm-Hansen, 1976). The data presented by Billen et al. (1980) are at variance with the other reports since they found growth yields of only 20–50% on glucose and 16–50% on a mixture of three amino acids. Using a different approach, Newell, Lucas & Linley (1981) have attempted to quantify the conversion of organic matter into bacterial biomass in culture by measuring the decrease in DOC and POC concentrations and concomitant increase in bacterial biomass; they estimated a conversion efficiency for carbon of only 10%, which is much lower than


103

any estimate using 14C. Is it possible that the short incubation times employed in the 14C experiments are resulting in low estimates of bacterial respiration? How relevant are the laboratory experiments of the type done by Newell et al. (1981) to the situation in the deep-sea? Further experimentation, both in the laboratory and in situ, is required to clarify the question of growth efficiency of marine bacteria. Why should marine bacteria have such apparently high growth yields? Williams (1970, 1973b) suggested that the requirement for biosynthesis is reduced because bacteria have a range of organic compounds available which can be directly incorporated into macromolecules and biomass; in these circumstances, less energy is required for biosynthesis resulting in reduced catabolism and increased growth yields. High growth yields have implications for the traditionally conceived rôle

TABLE III Respiration of added substrates as a percentage of gross uptake

Region

Substrate

% respiration

Reference

English Channel

Glucose

24–27

Williams (1970)

Amino-acid mixtures

34

Glucose

33–46

Amino-acid mixtures

13–22

Glucose

28–49

Amino-acid mixtures

23–30

Glucose

8–17

Leucine

3–37

Alanine

34–47

Glutamate

42–57

Arginine

22–47

Valine

10–29

U.S.A.. Estuary

Glutamate

30–43

Carney& Colwell (1976)

North Sea

Glucose

28.9

Gocke (1976)

Acetate

34.8

Lactate

37.6

Malate

59.2

Amino-acid mixture

25.3

Amino-acid mixture

25.5

Glucose

20

Mediterranean

N.E. Atlantic

U.S.A., Estuary

Bahamas

Crawford et al. (1974)

Williams & Yentsch (1976)


California, Coastal

Baltic

North Sea

Amino-acid mixture

15

Leucine

2–4

Alanine

31–35

Glutamate

20–30

Amino-acid mixture

19–31

Glucose

13–24

Glucose

50–80

Acetate

40–80

Lactate

40–85

Amino acids

50–84

578

Williams et al. (1976)

Dawson & Gocke (1978)

Billen et al. (1980)

of bacteria in recycling inorganic nutrients for subsequent utilization by phytoplankton. Terrestrial ecologists are well aware that bacteria can have a significant demand for inorganic nutrients and may be responsible for the removal of soluble inorganic nutrients; in the terminology of the soil microbiologist, such nutrients are immobilized. Decomposition of organic matter results in bacterial utilization of inorganic nutrients and incorporation into biomass; the C:N:P ratio of the organic matter influences the bacterial demand for inorganic N and P (Alexander, 1961) and hence the degree of mineralization or immobilization. Immobilization of nutrients by marine bacteria has rarely been considered (Thayer, 1974) but clearly it may be a significant process, especially if the high growth yields for carbon also apply to nitrogen. If the C:N and C:P ratios of organic matter are greater than those of the bacteria, then it will be necessary to supplement that organic matter with inorganic nutrients before there can be complete utilization. So under these circumstances, bacteria may compete with phytoplankton for available nutrients. Competition for phosphate has been reported by Rhee (1972), Faust & Correll (1976), and Friebele, Correll & Faust (1978) and competition for nitrogen and phosphorus by Thayer (1974). Some experiments have been done to measure nitrogen regeneration by marine bacteria. Hollibaugh (1978) supplied individual amino acids and measured the release of ammonia by an enrichment culture of marine bacteria; he found that the nitrogen regeneration ratio was about 0–8, that is, growth efficiencies in terms of nitrogen of only 20%. His estimates for carbon growth yield were also low (about 30%) which contrasts with the majority of reports of bacterial substrate utilization with their high growth yields. Hollibaugh (1978) used quite high concentrations of amino acid so that the ammonia produced could be chemically measurable and this may have influenced the results. Subsequent work by Hollibaugh, Carruthers, Fuhrman & Azam (1980) in enclosed water columns, gave high rates of bacterial nitrogen regeneration from natural organic matter. The bacterial production was closely coupled with, and appeared to be limited by, primary production. These experimental water columns were periodically supplied with inorganic nutrients and it is difficult to extrapolate from this situation to the nutrientlimited ocean. So, the data of Hollibaugh (1978) and Hollibaugh et al. (1980) do suggest


105

that bacteria are important in re-mineralization of nitrogen but the data of others (Table III) suggests a low mineralization rate of carbon compounds. Clearly, further studies on nitrogen and carbon turnover, in both nutrient-rich and nutrient-depleted sea water, are required to resolve this discrepancy. COUPLING BACTERIAL AND PRIMARY PRODUCTION Within the euphotic zone, production of DOC by phytoplankton may be a major source of readily metabolizable substrate for bacteria and the interactions occurring between bacteria and phytoplankton have been the subject of several studies. Initially, experiments involved supplying bacterial populations with known products of phytoplankton excretion, such as glycollate (Wright, 1970; Tanaka, Nakanishi & Kadota, 1974) but in recent years, several studies have measure simultaneous phytoplankton production of DOC and DOC uptake by bacteria. Experimental approaches have been quite varied; Derenbach & Williams (1974) and Larsson & Hagström (1979) used differential filtration through membrane filters of different pore size to separate bacteria from phytoplankton. Larsson & Hagström (1979) found, after 4 h incubation, that 65% of the labelled carbonate supplied was found in phytoplankton cells, 27% in bacterial cells, and 8% was DOC; the very rapid release of organic matter by phytoplankton cells was, therefore, simultaneously taken up by bacteria. Berman (1975) and Iturriaga & Hoppe (1977) attempted to differentiate heterotrophic and autotrophic production by supplying antibiotics to inhibit bacterial activity and found that DOC release was between 1 and 10% higher in the presence of antibiotics; they assumed this increase was due to the suppression of bacterial utilization. These studies are, however, less than clear-cut because of doubts over the efficiency of antibiotics in immobilizing marine bacteria; it is not possible to estimate the flux of carbon from phytoplankton to bacteria using this method. Another approach was adopted by Wiebe & Smith (1977) and by Schleyer (1980) which involved supplying 14C-labelled algal extract as tracers. Working in an Australian estuary, Wiebe & Smith (1977) found very close coupling between phytoplankton exudation and bacterial uptake and suggested that the rapid bacterial utilization would make it impossible for phytoplankton exudation ever to be directly responsible for the DOC found in marine waters. Interestingly, most of the uptake of 14C-DOC was by particles of 100–124 µm in diameter, suggesting either that heterotrophs responsible for uptake in this estuary were not bacteria, which is unlikely, or that bacteria attached to particles were responsible. Schleyer (1980) adopting the kinetic approach of Wright & Hobbie (1965), used labelled algal extract instead of a defined substrate. There were, however, several problems which made interpretation difficult; he found that the procedure over-estimated utilization of natural DOC because an unknown proportion of organic compounds are resistant to microbial degradation. Also, the production estimates made by this method were much greater than the increase in biomass measured by epifluorescence microscopy. Schleyer obtained growth yields on labelled algal organics which were much higher than those obtained with glucose; only 1.5% of the algal DOC was respired by the bacteria, giving a growth yield of 98.5%; such a high growth yield seems improbable and, even if the assimilated organic compounds were utilized in


580

anabolism as Williams (1970) suggested, it would be remarkable if the energy maintenance requirements of the bacteria required a respiration rate of only 1.5% of uptake. Clearly, the bacteria must have been utilizing other, unlabelled organic compounds to meet their maintenance requirements. Another method of determining bacterial utilization of phytoplankton DOC was used by Smith, Barber & Huntsman (1977) who measured the different rates of accumulation of DOC in the light and dark in a coastal upwelling off northwestern Africa; their carbon flux model suggested that bacteria remove 18% of the DOC excreted by phytoplankton per hour, but this represented only 1.7% of the total autotrophic production of organic matter. Clearly, this approach is open to criticism if, as Wiebe & Smith (1977) suggested, there is close coupling between phytoplankton DOC excretion and bacterial utilization, because DOC accumulation may be only a small fraction of the total flux between phytoplankton and bacteria. Smith et al. (1977) specifically discounted any rapid and complete removal of phytoplankton exudate by bacteria. A similar approach of relating diurnal changes in DOC with heterotrophic activity was used by Sieburth, Johnson, Burney & Lavoie (1977); they found that, in regions of intense microbial activity, there was a difference of 13% in DOC concentration and 32% in total carbohydrate between day and night samples. This result implies that bacterial utilization of algal DOC must be greater than Smith et al. (1977) found but much less than the rates reported by Wiebe & Smith (1977). Clearly, there are two questions which future studies must answer; how important is phytoplankton exudation of DOC to the total pool of DOC available for bacterial utilization and, are the reported differences in the coupling of phytoplankton exudation and bacterial uptake due to experimental artefacts or are there regional differences in the proportion of algal excretion which is directly utilized by bacteria Recently, Fuhrman, Ammerman & Azam (1980) studied bacterial production in a coastal euphotic zone and found that bacterial growth rate was related more to the standing stock of phytoplankton than to their production. It may be that bacteria were utilizing organic compounds produced by zooplankton predation rather than using DOC excreted during photosynthesis. If this result is found in other regions, then zooplankton grazing may play a dominant rôle in supplying organic matter for bacterial growth.

TROPHIC RÔLE OF MARINE BACTERIA The classic view of the marine food web is of three or four trophic levels from phytoplankton to fish (Ryther, 1969) but it is only recently that the place of marine bacteria in this scheme has been considered. A considerable stimulus was provided by the work of Russian workers who claimed that bacterial production was often greater than primary production. Sorokin (1971a, b, 1973, 1977) reported that the biomass and production of bacteria in the tropical Pacific considerably exceeded primary production. Apart from one report of bacterial production being less than primary production in a winter diatom bloom under ice in the Sea of Japan (Sorokin & Konovalova, 1973), Sorokin’s results consistently suggested that bacterial, heterotrophic production was greater than phytoplankton, autotrophic production. The obvious question is, where does


107

the organic matter come from to fuel this bacterial production? Sorokin (1971b) suggested the source was advection of deep-water from the more productive regions of the ocean into the area of the tropical Pacific that he was studying. Banse (1974) criticized this hypothesis because it could not be reconciled with the observed geographical distribution of DOC. Banse also reasoned that bacterial production could not be greater than 10% of phytoplankton production and that Sorokin’s estimates of bacterial production must be too high by at least an order of magnitude. The procedure used by the Russian workers to estimate heterotrophic production is open to criticism. The technique, first proposed by Romanenko (1964), is based on the fixation of 14CO2 by heterotrophic processes and assumes a constant relationship between dark 14CO2 fixation and total heterotrophic assimilation. Such dark CO2 fixation, however, is not strictly analagous to autotrophic carbon reduction because there is evidence (Overbeck, 1972) that the CO2 is fixed in anaplerotic reactions; these reactions are involved in replenishing intermediates of the TCA cycle which are diverted into biosynthesis (Kornberg, 1966) and this CO2 fixation does not represent de novo synthesis of biomass. The amount of carbon fixed in anaplerotic reactions varies considerably with substrate utilization and it would be improbable if there were a constant relationship between dark CO2 fixation and bacterial production. Other problems, such as the certain dark CO2 fixation by algae and the inconstancy of the ratio of dark CO2 fixation to total CO2 fixation have been discussed by Overbeck & Daley (1973) who cautioned against using the Romanenko technique. Yet, it would be premature to dismiss out of hand the Russian workers’ results on the basis of a doubtful technique. Other workers have also presented evidence that bacterial production is much greater than has previously been assumed. Joiris (1977) reported that heterotrophic consumption of organic matter in the southern North Sea was ten times greater than primary production. In considering all the possible sources of error in his methods, Joiris concluded that his estimate of heterotrophic respiration was unlikely to be an under-estimate but he had less confidence in the measurements of primary production, which were based on changes in dissolved oxygen concentration. Sieburth et al. (1976) suggested that 30–40% of the plankton biomass in the north Atlantic was composed of bacteria and that their estimates of bacterial production agreed well with Sorokin (1971b). This bacterial production could not be sustained by the estimates of primary production which are obtained using the standard 14C technique (Sieburth, 1977). It is generally accepted that the 14C method may be under-estimating phytoplankton production (see review by Peterson, 1980) but it is hardly credible that a technique which has been so widely used should under-estimate primary production by an order of magnitude as these estimates of bacterial production imply. Sieburth (1977) thought that the bacterial production estimates were more reliable because they were based on a number of different techniques, whereas primary production is measured almost exclusively by the 14C method. None of the methods used to measure bacterial production is, however, free from error and, perhaps, an error of two or three in both heterotrophic and autotrophic estimates may remove these apparent discrepancies between bacterial and phytoplankton production. It may also be that the above arguments are misleading because they relate bacterial production to primary production, implying that bacteria are competing directly with


582

zooplankton for the organic matter produced by phytoplankton. This need not be the case and an alternative hypothesis which depends on the efficiency with which zooplankton utilize ingested phytoplankton carbon has already been discussed (Fig. 1, p. 70). Such a process, coupled with phytoplankton production which we have previously speculated may not be completely utilized by zooplankton because of temporal mismatch in biomass, could result in a large proportion of the production passing to the DOC- and POC-pools. The end result is that a significant proportion of the total primary production (although much less than 100%) could be utilized by bacteria, without depriving zooplankton of any food. GRAZING ON BACTERIA More abundant bacterial biomass and greater bacterial production inevitably raises the question of what happens to the bacteria and is there a component of the food web which can exert a significant grazing pressure on marine bacteria? Boyd (1976) stated that few copepods are capable of filtering particles hollow sphere > ring > cap > disc which produces a full range variation in settling of about a factor of 2 for the same sphere radius (particle volume equals the equivalent sphere volume). The effect is that the equivalent sphere radius over-estimates settling. An aid to the reduction in errors is the estimation of particle density by S.E.M.-E.D.X-R. analysis and the calculation of viscosity for given depths based on hydrographic data because viscosity varies by a factor of 2 from warm surface water to deep cold water. Chase (1979) reported that the presence of dissolved organic matter in sea water and organic coatings on particles increased the settling velocity of particles in the range 5– 500 µm (based on the mean of two characteristic chords). The greatest effect was observed for small particles where variation of about an order of magnitude was noted, reducing to a factor of 2 for large particles. It is interesting to note that Brun-Cottan (1976) stated that Stokes’ settling does not apply to particles

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