CHAPTER S8 - Scripps Institution of Oceanographyltalley/DPO/supplementary/ch_s8.pdf · CHAPTER S8 Gravity Waves, Tides, and Coastal Oceanography: Supplementary Materials This web-based

C H A P T E R

S8

Gravity Waves, Tides, and CoastalOceanography: Supplementary Materials

This web-based content continues the topicsof Chapter 8, covering several different aspectsof coastal oceanography (river runoff, estuaries,and coral reefs in Sections S8.7, S8.8, and S8.9),followed by an extended discussion of adjacentseas (Section S8.10), including the Mediterra-nean, Black, Baltic, and North Seas from theAtlantic; the Bering, Okhotsk, Japan (East),Yellow, East China, South China Seas, and Gulfof California from the Pacific Ocean; and theRed Sea and Persian Gulf from the Indian Ocean.

Figure numbering in this portion of Chapter 8continues from the last figure of the print textbut with “S” denoting online material, thusstarting with Figure S8.16.

S8.7. WATER PROPERTIES INCOASTAL REGIONS: RIVER

RUNOFF

River runoff affects coastal regions. It reducesthe salinity of the surface layer and even of thedeeper water if there is sufficient vertical mixing.It often carries a large amount of suspended sedi-ment, as seen for the Mississippi River outflowand the outflows from the Himalayas into theBayofBengal, including theGangesRiver (FigureS8.16). Generally, river runoff has a pronounced

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seasonal variation, resulting in much largerseasonal fluctuations of salinity in coastal watersthan in the open ocean. In a coastal region whereprecipitation occurs chiefly as rain, the seasonalsalinity variation will closely follow the localprecipitation pattern. In regions where rivers arefed by meltwater from snowfields or glaciers,the river runoff increases in the summer tomany times the winter rate and causes a corre-sponding decrease of salinity that lags the snow-fall by several months.

Fresh river water flowing out over saltieropen ocean water creates a strong halocline,with high vertical stability. This can inhibit mix-ing with water below the halocline. In the warmseasons, this can result in higher temperaturesin the surface layer. In winter in high northernlatitudes, the halocline permits the surface layerto cool to below the temperature of the waterbeneath the halocline, producing a temperatureinversion. Ice thus tends to form first in coastalwaters, as “fast ice” (Section 3.9.1; the shallow-ness of the coastal region also contributes). Inregions of multiyear ice such as the Arctic, thenew coastal ice spreads seaward until it contactsthe first-year ice spreading shoreward from themultiyear pack ice.

Since river water frequently carries sus-pended sediment (Figure S8.16), coastal waters

FIGURE S8.16 (a) Sediments in the Ganges River plumein the northern Bay of Bengal. The Himalayas are the line ofsnow-covered mountains across the top of the image; thewhites in the center right are clouds. (b) Mississippi Riverestuary. Images from the Moderate Resolution ImagingSpectroradiometer (MODIS). Source: From NASA Goddard

Earth Sciences, (2007c, 2008).

S8. GRAVITY WAVES, TIDES, AND COASTAL OCEANOGRAPHY: SUPPLEMENTARY MATERIALS2

often have low optical transparency (Section3.8.1). Sometimes this sediment is carried inthe surface low-salinity layer for some distancewhile the deeper, more saline water remainsclear. The deposition of this sediment causesshoaling and consequent hazards to navigation.Frequently the location of the deposition isinfluenced by the salinity distribution becauseincreases in salinity can cause flocculation ofthe sediment and rapid settling.

The effect of runoff, especially from largerivers, can often be traced a long way from thecoast, both by reduced salinity and by the sedi-ment in the water. Some examples of major influ-ences on the open ocean include the Amazonand Congo Rivers in the tropical Atlantic, inthe northeast Pacific from the many rivers flow-ing off the North American continent, and in theBay of Bengal from the large rivers whose runoffis strongly influenced by the monsoon. Lowsalinity from these river sources can be seeneven in the global surface salinity map (Figure4.16). The net freshwater input from these sour-ces is an important part of the ocean’s freshwaterbudget as apparent in the list of outflows of themajor rivers of the world in order of volumetransport in Dai and Trenberth (2002). Runoff iscomparable to open ocean precipitation andevaporation because it deposits net precipitationover land into the ocean.

The result of the sediment deposition fromrunoff into the Bay of Bengal is apparent evenin open ocean bottom topography, as a smoothsediment fan spreading down to 5000 m depth,across the equator in the Indian Ocean, evidentin the bathymetry in Figure 4.13 (Curray,Emmel, & Moore, 2003).

S8.8. ESTUARIES

An estuary, in the strictest definition, isformed at the mouth of a river, where the rivermeets the sea (Dyer, 1997). Cameron and Pritch-ard (1963) defined an estuary as “a semi-enclosed

ESTUARIES 3

coastal body of water having a free connection tothe open sea and within which the sea-water ismeasurably diluted with fresh water derivingfrom land drainage.” They restrict the definitionto coastal features and exclude large bodies ofwater such as the Baltic Sea. The river water,which enters the estuary, mixes to some extentwith the salt water therein and eventually flowsout to the open sea in the upper layer. The mix-ing processes are mainly due to tides and thewind. A corresponding inflow of seawater takesplace below the upper layer. The inflow andoutflow are dynamically associated so that whilean increase in river flow tends to reduce thesalinity of the estuary water, it also causes anincreased inflow of seawater, which tends toincrease the salinity. Thus an approximatesteady state prevails.

The defining characteristic of estuarine circu-lation is that inflow is denser than outflow,which is diluted relative to the inflow. Some-times this concept is applied heuristically tomuch larger bodies of water, such as the BlackSea, or even the Indian and Pacific Oceans, butthe study of estuarine circulation is defined tobe within the confined coastal regions.

Extended descriptions of estuaries and estua-rine circulation can be found in the texts by Dyer(1997), Officer (1976), and Hardisty (2007).Compilations of papers on specific estuaries ortypes of estuaries have appeared over the years.Beardsley and Boicourt (1981) summarized theMiddle Atlantic Bight and Gulf of Maine;Farmer and Freeland (1983) reviewed the phys-ical oceanography of fjord estuaries; papers inNeilson, Kuo, & Brubaker (1989) summarizedthen-contemporary ideas about estuaries; andthe Estuarine and Coastal Science Associationperiodically produces compilations of papers.

S8.8.1. Types of Estuaries

There are many types of estuaries and manytypes of flow in estuaries. A classification systemis useful as an introduction, but inevitably results

in oversimplification (Pritchard, 1989). Estuariesare classified in termsofboth their shapeand theirstratification. They can also be classified in termsof tidal and wind forcing. The inland end of anestuary is called the head and the seaward endthe mouth. “Positive” estuaries have a river orrivers emptying into them, usually at the head.

In terms of geology, three specific types ofestuary are recognized: the coastal plain(drowned river valleys), the deep basin (e.g.,fjords), and the bar-built estuary; there are alsotypes that do not fit in these categories (Dyer,1997; Pritchard, 1989). The first is the result ofland subsidence or a rise of sea level that floodsa river valley; North American examples are theSt. Lawrence River valley and Chesapeake Bay.Typical examples of the deep basin are the fjordsof Norway, Greenland, Canada, South America,and New Zealand. Most of these have a sill orregion toward the seaward end, which is shal-lower than both the main basin of the fjordand the sea outside, so it restricts the exchangeof deep water. The third type is the narrowchannel between the shore and a bar, whichhas built up close to shore through sedimenta-tion or wave action.

In terms of stratification and salinity struc-ture, estuaries have been classified based onthe distribution of water properties as (a) verti-cally mixed, (b) slightly stratified, (c) highlystratified, and (d) salt wedge estuaries (FigureS8.17; Dyer, 1997; Pritchard, 1989). The stratifica-tion is due to salinity, because density in estu-aries is determined mainly by salinity ratherthan by temperature. The classification systemis not rigid. In the left-hand column of FigureS8.17, the salinity distributions are shown asvertical profiles at each of four stations betweenthe head and the mouth of the estuary (see sche-matic plan view at the top). The right-handcolumn shows simplified longitudinal sectionsof salinity from head to mouth for the full depthof the estuary. In most estuaries, unlike the sche-matics in Figure S8.17, the bottom depth is shal-lowest at the head.

FIGURE S8.17 Typical salinity/depth profiles (left) and longitudinal salinity sections (right) in different types of estu-aries: (a) vertically mixed, (b) slightly stratified, (c) highly stratified, and (d) salt wedge.


The vertically mixed estuary (Figure S8.17a) isgenerally shallow and the water is mixed verti-cally making it homogeneous from the surfaceto the bottom at any particular place along theestuary. The salinity increases with distancealong the estuary from head to mouth. The river

water in this type of estuary flows toward themouth while the salt may be considered to prog-ress from the sea toward the head by eddy diffu-sion at all depths. In the right-hand figure, thevertical isohalines indicate the homogeneity ofthe water at each location while the straight

ESTUARIES 5

arrows indicate that the direction of net flow ofthe water is seaward at all depths. (The circulararrows symbolize the mixing taking place at alldepths.) The Severn River in England is anexample of a vertically mixed estuary.

In the slightly stratified estuary (Figure S8.17b),which is usually also shallow, the salinityincreases from head to mouth at all depths.The water is essentially in two layers with theupper layer a little less saline than the deeperone at each position along the estuary, witha mixing layer between them (symbolized bythe circular arrows in Figure S8.17b). In thistype of estuary there is a net seaward (outward)flow in the upper layer and a net inward flow inthe deeper layer as shown by the straight arrowsin the vertical salinity section. In addition tothese flows at both levels there is the verticalmixing of both fresh and salt water giving riseto the longitudinal variation of salinity in bothlayers. The James River in Chesapeake Bay isan example of this type of estuary.

In the highly stratified estuary (Figure S8.17c),of which fjords are typical, the upper layerincreases in salinity from near zero in the riverat the head to a value close to that of the outsidesea at the mouth. The deep water, however, is ofalmost uniform salinity from head to mouth.Again there is a net outflow in the upper layerand inflow in the deeper water as shown bythe straight arrows in the salinity section. Inthese estuaries there is a very strong haloclinebetween the upper water and the deep water,particularly at the head where vertical salinitygradients of 10 to 20 psu per meter may occurin summer during the period of greatest riverrunoff. There is vertical mixing, but this resultspredominantly in an upward movement of saltwater from below into the upper layer, withlittle downward movement of fresh water. Oneexplanation for this almost unidirectional mix-ing is that internal waves are generated by thevelocity shear between the upper low salinitylayer and the deeper more saline water, andthat the tops of these waves break and throw

off a “spray” of saline water into the upper layerinto which it mixes. There is much less breakingat the bottom of the internal waves and there-fore no spray of fresh water downward intothe saline water.

For the salt wedge estuary (Fig S8.17d), thelongitudinal section indicates the reason for itsname. The saline water intrudes from the seaas a wedge below the river water. This situationis typical of rivers of large volume transportsuch as the Fraser or Mississippi Rivers (FigureS8.16b). It should be noted that the section inFigure S8.17 is exaggerated in the vertical direc-tion; the salt wedge really has a much smallerangle than shown, with almost horizontalisohalines.

The salt wedge estuary has features incommon with the stratified estuaries. There isa horizontal gradient of salinity at the bottom asin a slightly stratified estuary and a pronouncedvertical salinity gradient as in a highly stratifiedestuary. The distinction is in the lack of salinewater at the surface until it reaches the sea at themouth of the estuary, because of the large riverflow. In this type of estuary the salt wedgemigrates up and down the estuary as the tidefloods andebbs, sometimesbyseveral kilometers.

In terms of mixing, Stommel (reported byPritchard, 1989) suggested that estuaries be clas-sified in terms of tidal and wind forcing, whichare the main modes of mixing in estuaries. Thetides that are important in estuaries are almostalways co-oscillation tides (Section 8.6.2). Thisclassification has evolved to consider the tidalrange and effect of friction on tides in the estuary(Dyer, 1997). Both the winds and tides havetemporal and spatial variations. The stratifica-tion in an estuary can therefore vary significantlywith time as well as location in the estuary.

S8.8.2. Estuarine Circulation

In an estuary, the flow is out to the ocean inthe upper layer and into the estuary in thebottom layer. In stratified estuaries, the depth


of the halocline (thickness of the upper, lowsalinity layer) remains substantially constantfrom head to mouth of an estuary for a givenriver runoff. If the estuary width does notchange much, then the depth remains constant,which means that the cross-sectional area of theupper layer outflow remains the same while itsvolume transport increases because of theentrainment of salt water from below. Conse-quently, the speed of the outflowing surfacelayer markedly increases along the estuaryfrom head to mouth. The increase in volumeand speed can be considerable, with the outflowat the mouth as much as 10 to 30 times thevolume flow of the river. In his classical studyof Alberni Inlet d a typical, highly stratified,fjord-type estuary in British Columbia d Tully(1949) demonstrated the above features. Healso showed that the depth of the upper layerdecreased as the river runoff increased up toa critical value and thereafter increased asrunoff increased.

Estuarine circulation depends on severalfactors: the sill depth, river runoff rate, and thecharacter of the outside water density distribu-tion. Tides and mixing also impact the circula-tion. If the sill is so shallow that it penetratesinto the low-salinity, out-flowing upper layer,the full estuarine circulation cannot developand the subsurface inflow of saline water doesnot occur regularly. As a result, the deep wateris not exchanged regularly and tends to becomestagnant. This situation occurs in some of thesmaller Norwegian fjords, but is by no meanstypical of deep basin estuaries. Most of thefjords in Norway, as well as on the west coastsof North and South America and New Zealand,have sills that are deeper than the upper layer.Therefore the estuarine circulation is developedsufficiently to affect continual renewal of thedeep water and stagnation does not occur (Pick-ard, 1961; Pickard & Stanton, 1980). The rate ofrenewal is proportional to the circulation, whichis proportional to the river runoff. Fjord estu-aries with small river runoff show more

evidence of limited circulation in the form oflow oxygen values than those with large runoff.The depth of the sill has little effect as long as itis greater than the depth of the low-salinity, out-flowing upper layer.

The other major factor influencing theexchange of the deep basin water is seasonalvariation in the density structure of the outsideseawater. Although the downward mixing offresh water in an estuary is small, it does occurto some extent. Therefore the salinity, and hencethe density of the basin water, tends to decreaseslowly. If a change then occurs in the outsidewater such that the density outside becomesgreater than that inside at similar levels abovethe sill depth, then there will be an inflow ofwater from the sea. The inflowing water is likelyto sink, although not necessarily to the bottom,in the estuary basin and displace upward andoutward some of the previously resident water.In this way the basin water becomes refreshed.In deep-sill estuaries this refreshment mayoccur annually, but in shallow-sill estuaries itmay occur only at intervals of many years; thedisturbance to the biological regime may becataclysmic on these occasions (by displacingupward into the biotic zone the low-oxygenwater from the bottom). This type of basin-water replacement has been well documentedfor some Norwegian fjords (with very shallowsills), but it should not be considered character-istic of all fjord estuaries.

The previous remarks only briefly describesome of the salient characteristics of stratifiedestuaries; the property distributions in FigureS8.17 are smoothed and schematic. Real distri-butions show fine and mesoscale structure anddetailed features, some general and some local.In particular, because the density structure isdetermined largely by the salinity distribution,temperature maxima and minima are quitecommon in the water column. Mixing betweenfresh and salt water is largely governed by tidalmovements and the effects of internal waves.The circulation that was just reviewed for

CORAL REEFS 7

stratified estuaries is greatly modulated by thestrong tidal currents in the estuaries. This briefdescription also neglects the horizontal vari-ability and horizontal circulation in estuaries.

Estuarine characteristics and processes areobserved in ocean areas as well as near the coast.In the northeast Pacific and in the Bay of Bengal,where there is considerable river runoff, thedensity of the upper layer is controlled by thesalinity rather than by temperature as is usuallythe case in the open ocean. The upper, low-salinity layer of perhaps 100 m depth in thenortheast Pacific is much less dense than thedeeper, more saline water and the stability inthe halocline between them inhibits mixing.Consequently, the summer input of heat is trap-ped in the surface layer and a marked seasonalthermocline develops as shown in Figure 4.8.

S8.8.3. Flushing Time of Estuaries

The time that it takes to replace the fresh-water within an estuary through river dischargeis called the flushing time. This is important forwater quality within estuaries. The flushingtime has significant temporal variation, espe-cially since river flows have strong variability.Following Dyer (1997), the flushing time is thefreshwater volume (VF in units of m3) dividedby the river discharge (R, in units of m3/sec).Both the freshwater volume and the riverdischarge can be time dependent. Using obser-vations of the average salinity <S> within theestuary compared with the seawater salinity Sooutside the estuary, the freshwater fraction canbe estimated as F ¼ (So � <S>)/So. The fresh-water volume is the total volume, V, multipliedby the freshwater fraction. The flushing time isthen

tF ¼ VF=R ¼ FV=R: (S8.10)

Flushing times range from several days toa year. Dyer lists flushing times for several estu-aries: Narragansett Bay, Massachusetts e 12 to40 days depending on river flow; Mersey e 5.3

days; Bay of Fundy e 76 days; and the SevernEstuary e 100 to 300 days depending on riverflow.

Observing the salinity at all locations in theestuary at all times is unrealistic, so variousapproximate methods are used to determinethe flushing time. Dyer (1997) is a good sourcefor these different methods.

S8.9. CORAL REEFS

The physical oceanography of coral reefs wasof particular interest to George Pickard, the orig-inal author of this text. He published severalpapers and a book on the Great Barrier Reef in1977 (Pickard, Donguy, Henin, & Rougerie,1977). Therefore we retain this section ofChapter 8, but it has not been updated. Manypapers and some books have been publishedon the physical oceanography of coral reefssince the previous edition, and Wolanski (2001)and Monismith (2007) are suggested as startingpoints.

Prior to about 1970, most physical oceanog-raphy in coral reef areas had been carried outas ancillary to biological or chemical studies.Studies of the dynamics started in the late1970s with most studies carried out in the GreatBarrier Reef of Australia. Through the 1980s, ofthe 200 publications on the physical oceanog-raphy of coral reef regions, 70% refer to the GreatBarrier Reef and only about 30% to other reefareas. Because of observed degradation of coralreefs worldwide, coral reef research expandedgreatly in the 1990s, with numerous reefs nowthe focus of careful, ongoing studies d manywith physical oceanographic components.

S8.9.1. Topography of Coral Reefs

Coral reefs are features of many coastalregions between the tropics, along the conti-nental shelves, around islands, and also on thetops of shoals and seamounts in the open ocean

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with little or no emergent land in their vicinity,such as atolls. Reefs act as complex barriers toflow in their neighborhood. Living coral cannotwithstand exposure above water so it onlyoccurs below low-tide level; it also requires lightfor growth so it generally cannot survive belowabout 50 m depth.

Along a land boundary, two types of reef arerecognized: (1) the fringing reef, which extendsout from the shore and (2) the barrier reef, whichis located away from the shore with a relativelyreef-free region (lagoon) between it and theshore. Despite its name, a barrier reef is rarelycontinuous for very long distances; instead itconsists of a series of reefs with gaps betweenthem. Water exchange with the ocean can takeplace through the gaps as well as over the reefs.In some cases the outer reefs may be long(parallel to the coast) and narrow; in others,the “barrier” may consist of a number of indi-vidual reefs dotted over a relatively wide band(50 km or more) parallel to the shore, with irreg-ular passages between them connecting thelagoon with the open ocean outside. Open oceanreefs, based on shoals or seamounts, usuallyextend around the shoals; the “lagoon” is thebody of shallow water within the reef perimeter,which usually has gaps that permit some directexchange with the ocean. Atolls are reefs wheresufficient material has collected on parts of thereef to raise the level a few meters above sealevel and on which shrubs and trees may grow.

S8.9.2. Water Properties in Coral Reefs

Extensive and long-term measurements ofwater properties have been made in the GreatBarrier Reef; in the southwest lagoon of NewCaledonia; and in the Hawaiian, Floridian, andCaribbean reefs. The annual variation of watertemperature is generally approximately sinu-soidal with the maximum in the local summerand closely correlated with the air temperature.The annual range of temperature variationdecreases toward the equator.

Salinity variations are less regular than thoseof temperature. Near land, decreases occur dueto local precipitation and river runoff associatedwith monsoons, whereas for atolls only precipi-tation is effective. Increases of salinity are due toevaporation. An increase from an oceanic valueof 35.7 psu near the pass into Canton Islandlagoon to 39.5 psu at the back of the lagoonapproximately 15 km away was recorded, butthis is probably an extreme example.

Water depths in lagoons and around reefs aregenerally small, less than 50e100 m, and thewater is usually unstratified, for example,well-mixed vertically due to turbulence fromwind-wave effects and the rough character ofthe bottom over the reefs. Some stratificationin the upper 10e20 m occurs near river mouthsduring periods of heavy runoff, but even thenthe variations in the vertical are generally lessthan 1�C in temperature and 1psu in salinity.Along the Great Barrier Reef, intrusions ofcooler, more saline water can be evident nearthe bottom of passes through the outer reefswith Dt ¼ -5�C and DS ¼ +1 psu relative to theupper layer reef-area waters. In the very shallowwater over fringing reefs, diel (day-night) varia-tions of as much as 10 to 12�C have beenobserved and attributed to solar heating duringthe day and radiant cooling at night. It is prob-able that the relatively reef-free lagoon betweenthe shore and off-lying barrier reefs occursbecause coral is intolerant both of the freshwater and silt carried in by rivers.

S8.9.3. Currents in Coral Reefs

To describe the currents we divide them intothree classes: drift or long-period (periods ofweeks or more), weather band (periods ofdays), and tidal (periods of hours). Drift currentsare generated by steady wind stress (e.g., thetrade winds) or long-shore pressure gradients.The oceanic equatorial currents generated bythe trade winds cause flow over the mid-oceanreefs, such as at Bikini. In the central Great

CORAL REEFS 9

Barrier Reef, a 25-year time series of currentmeasurements showed equatorward currents of20 cm/sec during the south-east Trade Windseason,while at other times therewas a polewardcurrent of about 30 cm/sec attributable to thedownward slope to the south associated withthe southward flow of the East AustralianCurrent outside the Great Barrier Reef.

Weather band currents associated with conti-nental shelf waves have also been documentedfor the central and southern Great Barrier Reef.They are a consequence of fluctuations of windstress as weather systems move eastward withtheir centers over the southern part of Australia.The weather systems have periods of 10 to 20days and speeds of some 500 km/day (equator-ward). As the resulting ocean currents havea vertical range of only 10 to 30 cm (near theshore and diminishing to zero outside thereef), they are not evident to the eye and canonly be identified by analysis of tide or currentrecords. Water particle speeds are approxi-mately 20e40 cm/sec or 17e35 km/day. Thiscan result in long-shore displacements of waterof 100 to 200 km that can be very significant intransporting pollutants or plankton over suchdistances in the reef area.

Tidal-flows through reef passes of approxi-mately 200 cm/sec are common; speeds as highas 370 cm/sec (13 km/h) at Aldabra Atoll havebeen recorded. It should be noted that althoughsuch tidal speeds through passes can be large,the inflowing water then spreads out in thelagoon and the distance of penetration of oceanwater during the flood (which only lasts about6 hours) may be only a few kilometers. This issmall compared to the diameter of many atolllagoons and therefore tidal flows may onlyhave a limited effect on water exchange andflushing. Also, it has been observed that thereis often little mixing between the intruding oceanwater and the resident lagoon water.

The term “reef flat” refers to extensive areas ofcoral of relatively uniform height; the waterdepths over them are generally only a few

meters. Flow over them may be due to driftcurrents, shelf waves, and tidal currents as wellas to the local wind stress. Note that as the waterin these areas is very shallow, bottom friction ismore important than Coriolis force so thewind-driven flow is downwind, not to the leftor right of the wind direction as in deeper water(see Section 7.5.3). Tidal currents over the reefflats may have speeds of 100 cm/sec or more.These flows (over the surrounding reef) willnot necessarily go into a lagoon during the floodor out during the ebb. For instance, in the GreatBarrier Reef, the tide wave approaches from thenortheast so that the tidal flow in the area duringthe flood is to the southwest and can be overa surrounding reef into its lagoon on the north-east side but out of the lagoon on the southwestside at the same time. Wave-overtopping alsocontributes to the water in a reef lagoon. Thisoccurs when ocean waves or swells break onthe outside of a reef generating a slope acrossthe reef, which causes flow across it. This compo-nent can contribute as much transport acrossa reef as the other mechanisms combined.

S8.9.4. Circulation in Lagoons

The circulation in individual reef lagoons maybe forced by some or all of the current mecha-nisms described previously. In the extensiveopen lagoon between the shore and the barrierreef in theGreat Barrier Reef, steadydrift currentsdue to wind stress and pressure gradients, andperiodic currents due to tides and continentalshelfwaves, all contribute to thewatercirculation.In the New Caledonia lagoon, about 20 � 80 kmlong inside a narrow barrier reef, the tide appearsto be the main contributor to water circulationduring light wind conditions. During strongsoutheast trade winds the wind stress superim-poses a general northwestmotion over the barrierreef and across the lagoon.

Within theGreat Barrier Reef, individual reefsform partial obstacles to the general flow d“partial” because, except at very low water,


someflowoccurs continually over the coral reefs.Eddies often form downstream of reefs, particu-larly during flows associatedwith tidal currents.Such eddies can increase mixing on a small scale(tens to hundreds of meters) as well as formclosed volumes in which plankton can be heldfor hours or longer. Shedding of eddies behindreefs probablydoes not occur very often, becauseflow speeds are not great enough.

For the roughly circular lagoons within indi-vidual reefs in the Great Barrier Reef, again thedrift and periodic currents contribute to thecirculation together with inflow due to wave-overtopping. Studies within atoll lagoons havedemonstrated that in addition to inflow due toocean currents, tides, and wave-overtopping,wind stress causes downwind flow in the upperlayer at speeds of about 3% of the wind speedwhile a compensating upwind flow developsin the deeper water.

Residence times for water within lagoonscover a wide range. For lagoons of 2e10 kmdiameter in the Great Barrier Reef, times of 0.5to 4 days have been estimated, for Bikini Atoll40 to 80 days, and for very shallow lagoonssuch as at Fanning Atoll (18 km long but onlya few meters deep) periods of up to 11 monthswere estimated. Rougerie (1986) estimated resi-dence times in the New Caledonia lagoon as 2 to28 days, depending on the particular area andthe runoff, wind, and tide characteristics.

S8.10. ADJACENT SEAS

S8.10.1. General Inflow and OutflowCharacteristics

Adjacent seas affect the open ocean’s stratifi-cation and circulation through water masstransformation. Transformation can be fromdense to light water (as in the Black Sea andBaltic, where fresh water is added within thesea, greatly reducing the salinity, Chapter 9),from light to dense (as in the Nordic Seas and

the Mediterranean Sea, where there is largecooling and evaporation, Chapter 9), or due tovigorous mixing (as in the Indonesian passages,Chapter 11).

Exchanges between basins (inflow andoutflow) can be principally separated in thevertical or in the horizontal directions (FigureS8.18). Adjacent seas that are separated froma larger basin by a narrow strait usually havevertically stratified exchange, with inflow inone layer and outflow in a layer of a differentdensity. The inflow and outflow “layers” areseparated by an interfacial layer within whichvigorous vertical mixing modifies both theinflow and outflow. The Mediterranean, Black,and Baltic Seas, discussed in the followingsections, have this type of exchange. Otherexamples include the Red Sea and PersianGulf in the Indian Ocean (Section 11.6 andSection S8.10.7 below). For the Mediterranean,Red Sea and Persian Gulf, inflow is in the upperlayer and outflow in the lower layer, witha density increase within the sea. In the Blackand Baltic Seas, inflow is in the lower layerand outflow is in the upper layer, with a densitydecrease within the sea.

When the exchanges between basins canoccur over a much broader region than justa narrow strait, the inflow/outflow geometrycan be more horizontal, with water of onedensity entering in one region and transformedwater exiting in another (Figure S8.18b). Theinflow and outflow might adjoin each other, oroccur through separate passages. The CaribbeanSea and Gulf of Mexico (Section 9.3.1) are of thistype as are Fram Strait between Greenland andSpitsbergen (Chapter 12); the Indonesianpassages (Section 11.5); and the Bering,Okhotsk, and Japan Seas (Sections S8.10.5 andS8.10.6).

Many exchanges are a mixture of horizontaland vertical. The exchange between the NorthAtlantic andNorth Sea (Section S8.10.4) is mostlyhorizontal, with inflow through Dover Strait andnear the Shetland Islands and outflow along the

FIGURE S8.18 Schematic diagram of various types of circulation in seas adjacent to the oceans. (a) Vertically separatedinflow and outflow, typified by the Mediterranean and Red Seas (surface inflow and subsurface outflow with transformationwithin the sea), Baltic Sea and Hudson Bay (subsurface inflow and surface outflow, with transformation within the sea), andBlack Sea (subsurface inflow and surface ouflow, with no deep ventilation within the sea). (b) Horizontally separated inflowand outflow, typified by the Arctic Ocean and Caribbean Sea and also by the North Pacific marginal seas (surface inflow andsurface outflow, usually through a different strait from the inflow), and by the Nordic Seas, Labrador Sea/Baffin Bay, andalso the Persian Gulf (surface inflow and both surface and subsurface outflow).

ADJACENT SEAS 11

Norwegian coast. But there is also inflowof salinewater within the Norwegian Trench, beneath thefresher outflow along the Norwegian coast. FortheNordic Seasexchangewith theNorthAtlantic,the exchange is mostly horizontal, with generally

less dense Atlantic Water (AW) entering theNordic Seas along the eastern boundary in theNorwegian Atlantic Current, and denser outflowoccurring across each of the threemain deep sills.However, the easternmost of these outflows, over


the Faroe-Shetland Ridge, is beneath the north-ward inflow of AW.

S8.10.2. Mediterranean Sea

The Mediterranean Sea (Figure S8.19) isa nearly enclosed marginal sea in the eastern

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FIGURE S8.19 Mediterranean Sea. (a) Surface circulation scLetage (2005) and Robinson et al. (1991). Blue shows dense flow dwhich are roughly indicated with purple. Etopo2 topographyoverturning circulation (surface, intermediate, and deep layers)can flow over the sill.

North Atlantic, connected to the Atlanticthrough the Strait of Gibraltar, which has a silldepth of 284 m. The maximum depths withinthe sea are about 3400 m in the western basinand 4200 m in the eastern, separated by theStrait of Sicily, which is 430 m deep. The Medi-terranean is connected to the Black Sea in the

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; curved arrows indicate that only the lighter part of the layer

ADJACENT SEAS 13

northeast through the Dardanelles andBosporus.

The tidal range in the Mediterranean issmall, decreasing from 0.8 m at Gibraltar inthe west to 0.4 m at Port Said in the east (to aslow as 0.2 m along the French coast in thenorth). Sea level decreases in a northeasterlydirection by about 0.7 m from the African coastto the Aegean Sea.

The Mediterranean Sea is the prototype ofa “negative” basin in terms of water balance(Section 5.3.1), with evaporation exceedingprecipitation and runoff. There is also net cool-ing within the sea. Therefore, the outflow fromthe Mediterranean is denser (saltier and cooler)than the inflow. The sea is well ventilated to itsbottom as a result. The Mediterranean Sea hasa profound impact on North Atlantic waterproperties because of the high salinity anddensity of its outflow. While the Mediterraneancontributes only about one-third of the net evap-oration of the Atlantic Ocean, its cooling of thesaline surface waters allows them to sink todepth in the Atlantic, which is important forthe entire North Atlantic Deep Water formationprocess.

S8.10.2.1 Exchange at the Strait of Gibraltar

The net exchange through the Strait ofGibraltar is small, on the order of 0.7 Sv (Bryden,Candela, & Kinder, 1994). The salinity differencebetween the Atlantic inflow in the surface layerand Mediterranean outflow below is large: 2.3psu (from 36.1 to 38.4 psu). The inflow temper-ature is strongly seasonal, but averages around15e16�C, hence a potential density of sq ¼ 26.6to 26.8 kg/m3. The temperature of the outflowis around 13.3 �C, so the potential density ofthe outflow is about 28.95 kg/m3 (Figure 9.23b).

The flow through the Strait of Gibraltar ishydraulically controlled and modulated bytides (Armi & Farmer, 1988). Minimum silldepth is at the Camarinal Sill, while theminimum horizontal constriction is farther tothe east, at Tarifa Narrows; both affect the

outflow. The interface depth between the AWand outflowing Mediterranean Water is about100 m at the sill, and slopes downward to almost250 m depth on the Atlantic side (Bray, Ochoa, &Kinder, 1995). This downward slope is typical ofhydraulically supercritical flow. The interfacealso slopes upward toward the northern sideof the strait because of the Coriolis force, sothe saltiest, densest water is banked to the north(Figure 9.23a).

The saline Mediterranean Water at the Straitof Gibraltar is one of the densest water massesin the world ocean; it is denser than the variousNordic Seas Overflow Waters (NSOW) at theiroverflow sills. However, instead of sinking tothe bottom of the North Atlantic like theNSOW, the Mediterranean Water equilibratesat about 1200 m depth due to the difference inthe stratification of the entrained waters forthese two overflows (Price & Baringer, 1994;Figure S7.6b). The warmth of the Mediterraneanoutflow also means that it compresses less thanNSOW as both descend to high pressure; interms of potential density referenced to 4000dbar, the NSOW is actually denser than theMediterranean outflow.

S8.10.2.2 Circulation of the MediterraneanSea

The horizontal and vertical circulations in theMediterranean Sea are strongly affected by thebasin geometry, which is separated into westernand eastern basins by the Strait of Sicily andwhich has a saddle depth of about 430 m (FigureS8.19). The general sense of mean circulation inthe Mediterranean Sea is cyclonic. Surface waterenters from the North Atlantic through the Straitof Gibraltar. Within the Alboran Sea close to thestrait, the circulation is anticyclonic (Alborangyre), but then becomes cyclonic as the AWflows into the Algerian Basin, following theNorth African coastline. This eastward coastalflow is called the Algerian Current. The flowsplits at the Strait of Sicily into a branch thatcontinues eastward through the strait and


a northward branch feeding the cyclonicwestern Mediterranean circulation. The meanwestward flow along the northern side can becalled the Northern Current (Millot, 1991).

In the eastern Mediterranean, the Adriaticand Aegean circulations are each cyclonic.Circulation in the Levantine Basin is morecomplex, with a quasi-permanent, meanderingeastward jet in mid-basin (the mid-Mediterra-nean or mid-Levantine jet; e.g., Robinson et al.,1991 and Ozsoy et al., 1991). South of the jet,there are quasi-permanent, anticyclonic, sub-basin-scale gyres. North of the jet, there arecyclonic, sub-basin-scale gyres. Of the severalcommonly occurring gyres, the cyclonic Rhodesgyre between Crete and Cyprus is singled outhere, as it is the site of Levantine IntermediateWater (LIW) formation (see later in this section).

The Mediterranean circulation is markedlytime dependent. In the western Mediterranean,the Algerian Current regularly spawns large(100e200 km) anticyclonic mesoscale features,while the Northern Current region is alsoeddy-rich, although the features are not ascoherent (Millot, 1991). In the easternMediterra-nean, the sub-basin-scale anticyclonic andcyclonic gyres are strongly time dependent(Ozsoy et al., 1991; Robinson et al., 1991).

The mean horizontal circulation in the Medi-terranean at both intermediate and deep levelsis cyclonic, similar to the surface circulation(Millot & Taupier-Letage, 2005).

Transports of the main currents in the Medi-terranean are on the order of 1 to 3 Sv. That is,the mean circulation is on the order of theexchange through the Strait of Gibraltar, anddoes not reach the strength of the open NorthAtlantic currents.

The vertical circulation of the Mediterraneanis best described in terms of its water masses(Figure S8.19b and Section S8.10.2.3). AW flowsin through the Strait of Gibraltar and circulatesat the surface through the Mediterranean. Inthe eastern Mediterranean, LIW is formed inthe vicinity of the Rhodes Gyre and then

spreads cyclonically at mid-depth (200e600 m).LIW is the source of the deep waters in boththe eastern and western Mediterranean. LIWand the lighter part of the Eastern Mediterra-nean Deep Water flow back to the westthrough the Strait of Sicily beneath the east-ward flow of AW. This joins the Western Medi-terranean Deep Water (WMDW). Outflowthrough the Strait of Gibraltar originates inthe Northern Current along the coast of Spainand is composed of LIW and the upper part ofthe WMDW.

S8.10.2.3 Properties and Water MassesWithin the Mediterranean Sea

The Mediterranean is a saline, warm, well-ventilated basin. Due to prevailing dry north-west winds and frequent sunny days, there isa large excess (about 100 cm/year) of evapora-tion over precipitation in the eastern part ofthe Mediterranean Sea. The high temperaturesand salinities are surpassed only in the RedSea. Salinity ranges from 36.1 psu, in theentering AW, to 39.1 at the surface in the easternMediterranean. Bottom temperatures are above12.5�C even at 4000 m (Wust, 1961) and havebecome warmer (Klein et al., 1999). Bottomwater densities exceed sq ¼ 29.25 kg/m3 dueto the high salinity and deep oxygen exceeds200 mmol/kg. These properties differ greatlyfrom those at the same depth in the adjacentNorth Atlantic, which are 2.4�C, 34.9 psu, andsq ¼ 27.8 kg/m3.

The nomenclature for Mediterranean watermasses has varied. An official list of namesand acronyms is maintained by CIESM (2001).We limit our discussion to four primary watermasses: AW in the surface layer; LIW in theintermediate layer; and two dense waters,WMDW and the denser Eastern MediterraneanDeep Water (EMDW). LIW is formed in thenorthern Levantine Basin, off the south coastof Turkey near the island of Rhodes. The DeepWaters are formed at the northern edges of thebasins, chiefly in the Gulf of Lions in the western

ADJACENT SEAS 15

basin (WMDW) and in the southern Adriaticand in the Aegean (EMDW).

LIW is recognized throughout the Mediterra-nean by a subsurface vertical maximum ofsalinity between 200 and 600 m (Figure S8.20;Wust, 1961). At formation, LIW salinity isgreater than 39.1 psu and its temperature isaround 15�C. After passing westward through

FIGURE S8.20 (a) Salinity at the vertical salinity maximumdinal salinity section in winter to show the LIW.

the Strait of Sicily its core becomes colder (13.5�C), fresher (38.5), slightly less dense, and some-what deeper than in the eastern Mediterranean.

LIW formation was observed in the cyclonicRhodes Gyre in early 1995 (Malanotte-Rizzoliet al., 2003). Deep convection to 900 m occurredin this gyre (Figure S8.21), acting as a classicconvective chimney (Section 7.10). The dense

characterizing LIW. Source: From Wust (1961). (b) Longitu-

(a)

(b) (c)

FIGURE S8.21 Eastern Mediterranean convection. Formation of dense water in the Levantine Basin in January 1995.(a) Surface dynamic height relative to 800 dbar. The cyclonic low centered at 28�30’E, 35�N is the Rhodes gyre. (b) Salinityand (c) potential density sq along section A. LIW is the salinity maximum at 50e100 m. The convective chimney in the centeris forming a deep water. Source: From Malanotte-Rizzoli et al. (2003).


ADJACENT SEAS 17

water in the chimney in that year was a newlyidentified deep water, rather than LIW. NewLIW was the shallow salinity maximum on theoutside of the gyre (right and left at 50e100 mdepth in Figure S8.21b,c). LIW spread muchfarther into the eastern Mediterranean at theend of that winter than did the new deep water,which was trapped in the Rhodes Gyre.

The Deep Waters originate at several loca-tions along the northern coast through wintercooling of the widespread high salinity LIW. Inthe Adriatic and Aegean, cold winter outbreakswith intense winds (the “Bora”) create EMDWfrom LIW. The Adriatic has historically beenthe site of the densest formation (Schlitzer etal., 1991). However, changes in winter condi-tions shifted the densest water production tothe Aegean in the 1980s, and then back to theAdriatic in the 1990s (Klein et al., 2000). Thepresence of EMDW affects the properties ofnew LIW in the eastern Mediterranean.

In the western Mediterranean, dense waterformation contributing to WMDW occursprimarily in the Gulf of Lions in a sub-basin-scale cyclonic gyre (Figure S8.22), in responseto cold, dry winter winds (the “Mistral”). Thefirst observations of the classic stages of deepconvection (preconditioning, convective mixingand spreading; Section 7.10.1) were made herein 1969 (MEDOC Group, 1970; Sankey, 1973).Deep convection has occurred reliably in thisregion in many other years within a cyclonicdome of about 100 km scale. Dense water prop-erties here are around 12.8�C, 38.45 psu, andsq ¼ 29.1 kg/m3 (Marshall & Schott, 1999).Thus WMDW is not as dense as EMDW, butthe Strait of Sicily blocks the densest EMDWfrom flowing into the western Mediterranean.Thus the deep water of the western Mediterra-nean is a mixture of local Gulf of Lions densewaters and the shallower part of EMDW.

Water mass properties and formation rates inthe Mediterranean are demonstrably affected bythe North Atlantic Oscillation and by the warm-ing and drying trends of global climate change

(Section S15.6). Because the sea is relativelysmall, these changes affect the relative balanceand properties of the different deep and inter-mediate waters. The net effect is observed inchanges of salinity and temperature of theoutflow at the Strait of Gibraltar. This hasaffected Mediterranean Water properties withinthe North Atlantic (Potter & Lozier, 2004).

S8.10.3. Black Sea

The Black Sea is an almost completely iso-lated marginal sea with a maximum depthof over 2200 m. It is connected to the north-eastern Mediterranean Sea through the narrowBosphorus and Dardanelles, which have depthsof only 33 and 70 m, respectively (Figure S8.23).The small Sea of Marmara lies between the twostraits. Inflow from the Mediterranean is moresaline and denser than the fresh outflow fromthe Black Sea. The Black Sea thus representsa classic estuarine circulation with inflow atthe bottom and outflow at the surface in thestraits; it is a “positive sea” because it hasa net input of freshwater (Section 5.3.2). TheBlack Sea is also a prototypical anoxic sea withno dissolved oxygen below the pycnoclinebecause of the very long residence time of itsdeep water. Black Sea physical oceanographywas reviewed by Ozsoy and Unluata (1998)and Oguz et al. (2006).

The surface circulation of the Black Sea iscyclonic overall, with a cyclonic gyre in eachof the west and east basins, including cycloniceddies (Figure S8.23a and Oguz et al., 2006). ARim Current circulates around the exterior,roughly following the continental shelf break.Its maximum velocity is 50e100 cm/sec atthe surface. Inshore of the Rim Current isa series of anticyclonic eddies or small gyresthat connect the coastal regions with thecyclonic circulation (light contours in FigureS8.23a). The whole is dominated by time-dependent eddies and seasonal changes incirculation.

FIGURE S8.22 Western Mediterranean convection. (a) Deep convection region in the Gulf of Lions for three winters.Circulation (arrows) and isopycnal depth (m) showing cyclonic doming. (b) Potential density through the convection regionin February 1992. Source: From Marshall and Schott (1999).


ADJACENT SEAS 19

The narrowness and shallowness of thepassages between the Black Sea and Mediterra-nean result in high current speeds and verticalshear. The consequent turbulence causesvertical mixing between the inflow and outflowlayers. Therefore the surface water that leavesthe Black Sea with a salinity of about 17 psu rea-ches the Mediterranean with its salinityincreased to about 30 psu, while the salinity of38.5 psu of incoming subsurface Mediterranean

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in the Black Sea, 1988. Adapted from Murray et al. (1989). (c) Ovet al. (2006).

water is reduced to about 34 psu by the time thatit reaches the Black Sea.

Exchange with the Mediterranean includesinflow on the order of 300 km3 yr�1 (300 km3

yr�1 is equal to 9.5 � 103 m3sec�1, hence 0.0095Sv) and outflow of the order of 600 km3 yr�1

(Oguz et al., 2006). The higher outflow is dueto net freshwater input. Evaporation and precip-itation within the Black Sea are each on the orderof 300 km3 yr�1 each, so they are nearly

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FIGURE S8.23 (Continued).


balanced. Inflow from the large rivers, includingthe Danube, Dniester, and Dnieper in the north-west, is also on the order of 300 km3 yr�1. Thuswithout the river inflow, salinity in the Black Seawould be nearly neutrally balanced. In Section5.3.2, we calculated a residence time for theBlack Sea of 1000 to 2000 years.

As a result of the net input of freshwater andthe long residence time, the Black Sea is one ofthe world’s major brackish (low salinity) seas.

A halocline between 50 and 100 m separatesthe fresher surface layer from the subsurfacewater column (Figure S8.23b). The sharp halo-cline/pycnocline separates the upper low-salinity, oxygenated (oxic) water from thedeeper oxygen-free (anoxic) water.

A subsurface temperatureminimum of<8 �Cnear 100 m is called the Cold Intermediate Layer. Itis most likely a remnant of the winter surfacemixed layer. Because the surface layer is so

ADJACENT SEAS 21

fresh, this density structure is stable; it is similarto the subpolar structures found in the northernNorth Pacific and parts of the Southern Ocean.The Deep Water in the Black Sea has a salinityof 22.3 psu and potential temperature of 8.9 �C.

The salinity structure and isolation of thedeep Black Sea produce interesting doublediffusive and deep geothermal convectivefeatures. The overall vertical structure is diffu-sive, with colder, fresh water in the Cold Inter-mediate Layer overlying warmer, saltier waterbelow the pycnocline (Ozsoy & Unluata, 1998;Kelley et al., 2003). Inflowing salty Mediterra-nean water enters the pycnocline and below inintrusions that are double diffusive. Weakgeothermal heating at the bottom of the BlackSea, along with the long residence time,creates a very thick (>450 m) convectivebottom layer, rivaled only in the deep Arctic(Section 12.5.3; Timmermans, Garrett, &Carmack, 2003).

Over geological time, the Black Sea hasvaried from being fresh (as recently as 7000years ago), to moderately saline. As global sealevels rose and fell during glacials and intergla-cials and river outlets changed location, theexchange between the Black Sea and Aegeanmay have reversed direction and even ceasedbecause the Bosphorus and Dardanelles are soshallow. There is an ongoing paleoclimatedebate about whether overflow in through theBosphorus 5600 years ago created a sea levelrise of tens of meters and massive floodingaround the Black Sea or milder changes (Giosan,Filip, & Constatinescu, 2009).

The considerable river runoff into the BlackSea has decreased by 15% in the last severaldecades due to the diversion of the river waterfor agricultural purposes. Observations in 1988compared with 1969 (Figure S8.23b) showedhigher salinity by about 0.1 psu, ascribed tothe change in runoff (Murray et al., 1989). Thelower temperature in the surface layer in FigureS8.23b is due to a month of observation in thetwo years.

S8.10.4. Baltic and North Seas

The North Sea is the semi-enclosed, shallow,continental shelf sea of about 100 m depthbetween the British Isles, Norway, and Europe;it is connected to the open North Atlanticthrough a broad region between Scotland andNorway at 61�e62�N and through Dover Strait(Figure S8.24). The Baltic Sea is the nearlyenclosed sea east of Denmark. The Baltic isconnected to the North Sea at its southwestend through a complex of passages with a silldepth of 18 m, leading to the Kattegat andthe North Sea. The Kattegat is the small seabetween Denmark and Sweden. The Baltic,which includes the Gulf of Bothnia to the northand the Gulf of Finland to the east, is thelargest area of brackish (nearly fresh) water inthe ocean system. It has irregular bottomtopography, with a mean depth of 57 m, anda number of basins of which the deepest is459 m deep.

The physical oceanography of both seas wasreviewed in Rodhe (1998) and Rodhe, Tett, andWulff, (2006). Since 1992, the Baltic Sea hasbeen the focus of an intensive hydrological cyclestudy called “Baltex” (the Baltic Sea Experi-ment); as an outcome of the study, Lepparantaand Myrberg (2009) provided a thorough over-view of Baltic Sea physical oceanography.

Surface salinity clearly illustrates the connec-tion of the North Sea to the open ocean and themuch greater isolation of the Baltic (FigureS8.24b). Surface salinity in the North Sea is closeto oceanic values, with a tongue of high salinity(>35) entering from the north. Through the Kat-tegat and into the Baltic, there is an enormousdecrease, with salinity in the southern Balticbetween 7 and 8 psu, dropping to less than 2psu in the northernmost Gulf of Bothnia andeasternmost Gulf of Finland.

The North Sea circulation is cyclonic, withmost water entering and leaving across thecontinental shelf break in the north. Theexchange transport is about 2 Sv. Properties


and circulation along the western side arestrongly modulated by tides, which verticallymix the inflowing waters, with some dilutionof salinity due to river inflow. Inflow from theBaltic through the Kattegat introduces muchlower salinity waters into the North Sea; thenorthward flow along the coast of Norwaythat feeds the outflow is strongly stratified insalinity as a result. Its circulation is estuarinebecause of this net freshwater input.

The Baltic Sea, like the North Sea, has an estu-arine circulation (Figure S8.24c) with the upper

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FIGURE S8.24 North Sea and Baltic Sea. (a) Surface circuLepparanta and Myrberg (2009) with Etopo2 topography from NOAKattegat. (b) Surface salinity in August in the Baltic and NorthBaltic. Source: From Winsor, Rodhe, and Omstedt. (2001).

layer outflow in the Kattegat having a salinityof 20 psu and the bottom layer inflow havinga salinity of 30e34 psu. Large-scale meteorolog-ical conditions can override the estuarine circu-lation and result in full depth inflow or outflowat times.

The mean circulation in the Baltic Sea is weakand cyclonic with mean surface currents ofabout 5 cm/sec and no major stable features.The complicated geography of this sea createsa complex deeper circulation. However, thecirculation is extremely time dependent and

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FIGURE S8.24 (Continued).

ADJACENT SEAS 23

strongly coupled to the wind (through theEkman layer) because of the shallowness ofthe sea; currents can reach 50 cm/sec in theopen sea and up to 100 cm/sec in straits duringstorms (Lepparanta & Myrberg, 2009).

The very low salinity of the Baltic Sea resultsprimarily from river runoff and a long resi-dence time of waters within the Baltic of morethan 30 years (see Section 4.7). Evaporationand precipitation are estimated to be nearly


equal at about 47 cm/year. The freshwaterbudget is always positive, even on a monthlylevel, because of the river inflows, which areequivalent to 123 cm of water over the wholearea of the Baltic Sea. The bulk is due to theNeva River into the Gulf of Finland, equivalentto 400 cm/year over that gulf, and 170 cm/yearto the Gulf of Bothnia, with significant year-to-year variations.

The Baltic is basically a two-layer system,with a well-mixed upper layer (in terms ofsalinity) that is 30e 50 m deep in the south,increasing to 60e70 m in the central Baltic. Theupper layer temperature in summer is over 10�C with a thermocline at 15e20 m, which issignificantly shallower than the halocline.Surface layer salinities are 6e 8 psu while thedeeper waters are usually 10e 13 psu, some-times exceeding 17 psu in the south when largeinflows occur. These bursts of inflow typicallyoccur no more than once a year, with the stron-gest events sometimes separated by a decade ormore (Jakobsen, 1995). Surface salinity in thegulfs is even lower at 2e7 psu. Dissolvedoxygen reaches 100% saturation in the surfacelayers, but is relatively low in the deep water,with variations on a decadal timescale relatedto variations of inflows from the Kattegat.Anoxic conditions occur in many of the deeperBaltic basins where the residence times areseveral years.

Starting in January, sea ice forms in the northand east gulfs (Gulfs of Bothnia and Finland)along the coast. The ice often extends to mid-gulf but is less extensive in the central Baltic.

S8.10.5. Subtropical North PacificMarginal Seas

Along the western boundary of the tropicaland subtropical Pacific lies a set of marginalseas that interact differently with the openNorth Pacific circulation (Figure 10.1). Fromsouth to north, these are the South China Sea,the East China Sea, the Yellow Sea, and the

Japan (or East) Sea. Farther to the north lie thesubpolar Okhotsk and Bering Seas (SectionS8.10.6). Wind forcing for the southern part ofthis region is strongly monsoonal, with the sea-sonality weakening to the north. These seas areconnected through shallow straits or are locatedon the continental shelf; the net transportexchanges between them and with the Kur-oshio, which lies in deep water to the east, there-fore are limited and of the order of 2 Sv.

The South China Sea circulation is highlyseasonal, driven by the Asian monsoon. Inflowfrom the Pacific occurs throughout the yearthrough Luzon Strait between Luzon and Tai-wan; these are Kuroshio waters. Exit is north-ward through Taiwan Strait, between Taiwanand the continent, also throughout the year.The exchange rate is approximately 2 Sv (Xueet al., 2004). Within the South China Sea, thesummer circulation is mostly anticyclonic,driven by southwesterly winds, while in winterit is mostly cyclonic, driven by northeasterlywinds (Hu, Kawamura, Hong, & Qi, 2000). InFigure 10.1, only the cyclonic winter circulationis depicted. The western boundary currentalong Malaysia, Vietnam, Hainan, and southernChina reverses from southward in winter tonorthward in summer. However, in the north,the South China Sea Warm Current flows north-ward through Taiwan Strait throughout theyear, with maximum transport in summerwhen the full western boundary currentcomplex is northward.

The East China Sea and Yellow Sea constitutethe broad continental shelf region that lies eastof China, north of Taiwan, and west and southof Korea. The eastern edge of the continentalshelf is the effective western boundary for theNorth Pacific’s circulation, hence for the Kur-oshio, which flows northward along the conti-nental slope. Flow enters the East China Seafrom the South China Sea through Taiwan Straitin the TaiwanWarm Current. Exit is to the northinto the Japan Sea through Tsushima Strait(Korea Strait) in the Tsushima Warm Current.

ADJACENT SEAS 25

Within the Yellow Sea, the circulation is cyclonicand much stronger in winter than in summerdue to the strong northerly monsoonal windforcing (Naimie, Blain, & Lynch, 2001). Thisoverall region also absorbs the major freshwateroutput from the Changjiang River (YangtzeRiver). Cross-shelf exchange with the Kuroshiomodifies the properties of the East China Seawaters. The Bohai Sea, which is the gulf northof the Yellow Sea, forms sea ice in winter; thisis the southernmost ice-covered region in theNorthern Hemisphere.

The Japan Sea (East Sea) lies between Asia andJapan. It is deep with bottom depths exceeding3000 m, but it is connected to the North Pacificand Okhotsk Sea only through shallow straits.Water enters the Japan Sea from the souththrough Tsushima Strait (140 m deep). Thesource of this warm, saline subtropical water isthe East China Sea with some possible inputfrom an onshore branch of the Kuroshio. Thenet transport into the Japan Sea is estimated ata little less than 2 Sv (Teague et al., 2006). Waterexits from the Japan Sea mainly through Tsu-garu Strait, between Honshu and Hokkaido(130 m deep). There is also small but importanttransport into the Okhotsk Sea through SoyaStrait between Hokkaido and Sakhalin, andthrough the very shallow Tatar Strait far to thenorth, between Siberia and Sakhalin.

Within the Japan Sea, there are typicalsubtropical and subpolar circulations drivenby Ekman downwelling in the south andupwelling in the north, and separated by a zonalsubarctic front that is similar to the North Pacif-ic’s subarctic front. The northward subtropicalwestern boundary current is the East KoreanWarm Current. The subpolar western boundarycurrent is the Primorye (or Liman) Currentwhere it flows along the coast of Russia andthe North Korean Cold Current where the flowintrudes southward along the Korean coast.

The Japan Sea circulation deviates froma typical open ocean gyre system because ofits vigorous eastern boundary current, the

Tsushima Warm Current, which flows north-ward along the coast of Honshu. This resultsfrom the “island effect,” which is related tothe wind forcing of the entire North Pacificcirculation east of Japan with open straits onthe southern and northern sides of the island(this is outside the scope of this text).

A principal role of the Japan Sea in the NorthPacific circulation is to carry warm, salinesubtropical water northward west of Japan,(cool and freshen it), and then expel the still-saline water north of the Kuroshio’s separationpoint. This impacts details of formation of thesalinity minimum of North Pacific IntermediateWater east of Japan (Section 10.9.2; review inTalley et al., 2006).

S8.10.6. Bering and Okhotsk Seas

The Bering and Okhotsk Seas are separatedfrom the North Pacific by the long Aleutianand Kuril Island chains. The North Pacific’scyclonic subpolar circulation partially loopsthrough these adjacent seas. The two seas areintrinsically part of the North Pacific’s circula-tion, but the island chains create leaky barriersthat partially support boundary currents anda large amount of mixing in the island passagesdue to tides. Both seas have sea ice formationand brine rejection processes in the winter thatcreate denser shelf waters. Because the OkhotskSea has a salty external source of water from theJapan (East) Sea, through Soya Strait, its brinerejection process produces denser water thanin the Bering Sea. The Bering Sea’s special roleis as a small conduit of Pacific waters to theAtlantic Ocean.

In the Bering Sea, cyclonic circulation entersfrom the Alaskan Stream beginning with theeasternmost passages through the Aleutians;the principal deep inlet straits are AmchitkaPass at about the date line (1155 m), and NearStrait at 170�E (2000 m), just west of Attu Island(Stabeno & Reed, 1995). Most of the exit flowis through Kamchatka Strait (4420 m depth)


between Kamchatka and the KomadorskiyIslands with a smaller amount exiting to theArctic through the shallow Bering Strait. TheBering Sea is well known for its vigorous eddyfield that obscures much of the mean circulationin synoptic data (Reed, 1995). Flow aroundgroups of islands in the Aleutian chain is anticy-clonic and tidally driven.

Within the Bering Sea, the inflow from thecentral and eastern straits proceeds cyclonically,with the principal northwestward flowfollowing the continental shelf topography asthe Bering Slope Current (Kinder, Coachman, &Galt, 1975). When it encounters the Kamchatkaboundary, the Bering Slope Current splits. Thesouthward flow along the coast of Kamchatkais the East Kamchatka Current (EKC), whichbecomes the principal western boundarycurrent for the North Pacific’s subpolar gyre.This is joined by water circulating in the deeperbasin of the Bering Sea, and exits to the NorthPacific following the Kamchatka coast. TheEKC outflow transport through KamchatkaStrait is 6e12 Sv (Stabeno & Reed, 1995).

The flow of water northward through BeringStrait to the Arctic Ocean is one of the principalpathways in the global overturning circulation,despite the shallowness of the strait (~50 m)and its small transport (~0.8 Sv; Roach et al.,1995). This is the only northern connectionbetween the Pacific and Atlantic. The flow isrelatively fresh (~32.5 psu) relative to globalmean salinities and is thus one of the freshwaterexits from the Pacific (Wijffels, Schmitt, Bryden,& Stigebrandt, 1992; Talley, 2008). The waterflowing into Bering Strait comes from a westernboundary flow, the Anadyr Current, which isthe northward branch of the Bering SlopeCurrent, and a warmer eastern boundary flow,the Alaskan Coastal Current, which is fed bycyclonic circulation around the broad BeringSea shelf (Woodgate & Aagaard, 2005).

The Okhotsk Sea is connected to the NorthPacific through the Kuril Island chain. It is thesource of the densest water in the North Pacific,

which contributes to the North Pacific Interme-diate Water. The predominantly cyclonic circu-lation of the Okhotsk Sea enters through thenorthernmost strait close to the southern endof the Kamchatka peninsula (KruzenshternStrait, ~1400 m depth), and through Bussol’Strait, which lies in the center of the Kurilsand is the deepest passage (~2300 m depth).Net outflow is mainly through Bussol’ Strait,which, like the other straits, has bidirectionalflow associated with anticyclonic flow aroundeach island. The net transport in and out of theOkhotsk Sea is approximately 3e4 Sv (Glady-shev et al., 2003. This water is greatly modifiedwithin the Okhotsk Sea.

Within the Okhotsk Sea, the cyclonic circula-tion flows northward along the western side ofKamchatka as the West Kamchatka Current, andwestward along the broad continental shelveson the northern Siberian boundary, where seaice formation produces especially dense shelfwaters (Section 10.9.2). The Amur River injectsfresh water in the northwest. The complexthen moves southward along the east coast ofSakhalin, as the East Sakhalin Current, which isa typical narrow western boundary current.Waters from the East Sakhalin Current enterthe region south of Sakhalin, join an anticycloniccirculation there, and then head for Bussol’Strait. They are joined by eastward flow in theSoya Current along the northern coast of Hok-kaido that enters the Okhotsk Sea from theJapan Sea through Soya Strait.

S8.10.7. Red Sea and Persian Gulf

Geographically, the Red Sea, west of theArabian Peninsula, is a rift valley, resultingfrom the separation of Africa and the ArabianPeninsula, which is closed at the north andopens to the Gulf of Aden, Arabian Sea, andthe Indian Ocean at the south through thenarrow strait of the Bab el Mandeb (or Bab alMandab; see Ross, 1983 in Ketchum). Thedepth averages 560 m, with maximum values

FIGURE S8.25 Schematic circulations: (a) Red Sea and(b) Persian Gulf. Source: From Johns et al. (1999).

ADJACENT SEAS 27

of 2900 m and a sill of about 110 m depth at theBab el Mandeb in the south.

In contrast, the Persian orArabianGulf, east ofthe Arabian Peninsula, is shallow with amaximum depth of 105 m and average depth of35 m (Swift & Bower, 2002). The Persian Gulf isconnected to the Arabian Sea through the Straitof Hormuz (86 m deep) and the Gulf of Oman.

This brief introduction to these marginal seasis well complemented by the greater detail inTomczak and Godfrey (1994).

A major aspect of the northwestern ArabianSea is the high evaporation rate of approxi-mately 100e 200 cm/year, while precipitationaverages about 7e10 cm/year. There are nomajor rivers flowing into the Red Sea. The Tigrisand Euphrates Rivers drain into the PersianGulf, but their freshwater contribution is farsmaller than the net evaporation.

The water structure in the Red Sea consists ofa shallow upper layer and a thick deep layerseparated by a thermocline/halocline at about200 m depth. At the surface, the temperaturein summer (June eSeptember) is 26�e30�C,and in winter (October eMay) it is 24e28�C.Below the thermocline the deep layer is nearlyisothermal, at 21.6e21.8�C. The Red Sea is themost saline large body of ocean water, withsurface layer values of 38e 40 psu (with highervalues to 42.5 psu in the north) and deep watervalues of 40.5e40.6 psu. The deep water isformed by winter cooling in the north. Thesurface layer is saturated with dissolvedoxygen, but the absolute values are low becauseof the high temperature (less than 175 mmol/kg).There is an oxygen minimum of 20e60 mmol/kgat 400 m below the thermocline/halocline,whereas the deep water below this has a contentof 80e90 mmol/kg.

A schematic Red Sea mean circulation is pre-sented in Figure S8.25a. The central Red Seaocean surface pressure (sea level) is dominatedby two highs with flanking lows in the northand south. Intermediate water forms in thenorthern low. All of the mean boundary


currents flow to the north, reflecting the netthermohaline overturn in the Red Sea.

The Red Sea circulation’s seasonal variation isrelated to thewinds (Johns et al., 1999; Sofianos&Johns, 2003). In summer (Southwest Monsoon)the winds are to the south over the whole Sea(Figure S8.26); the surface flow is southwardwith outflow through the Bab el Mandeb, whilethere is a subsurface inflow to the north andweak outflow at the bottom through that strait.In the winter (Northeast Monsoon) the windsover the southern half of the Red Sea change to

FIGURE S8.26 Wind stress (dyn/cm2) for the Red Sea and(Southwest Monsoon). Sources: From Johns et al. (1999); see also

northward; there is a northward surface flowover the whole of the Red Sea and a subsurfacesouthward flow with outflow through the Babel Mandeb. The residence time for the upperlayer has been estimated at 6 years and for thedeep water at 200 years.

Hot brine pools are found in some of the deep-est parts of the Red Sea (Karbe, 1987). Very hightemperatures of 58�C and salinities of 320 psu aredue to hydrothermal activity. The heat flow upthrough the bottom is much greater than theworld average of 4 � 10�2 W/m2. The high

Persian Gulf: (a) January (Northeast Monsoon) and (b) JulySofianos and Johns (2003).

REFERENCES 29

salinity value is not directly comparable to oceanwater salinities because the chemical constitutionof these brines is quite different. They havea much higher content of metal ions. (Forcomparison, a saturated solution of sodium chlo-ride in water has a salinity value in the oceano-graphic sense of about 270.) The favoredexplanation for the origin of the chemical constit-uents is that this is interstitial water from sedi-ments or solutions in water of crystallizationfrom solid materials in the sea bottom, whichare released by heating from below and forcedout through cracks into the deep basins of theRed Sea.

Circulation and formation of hypersalinewater in winter in the Persian Gulf was brieflysummarized in Section 11.6, based on Johns etal. (2003) and Swift and Bower (2003). Thecyclonic circulation, exchange through theStraits of Hormuz, and formation region in thesouth of the high salinity water mass are illus-trated in Figure S8.25b. Winter surface salinityin the southern region exceeds 42 psu. Becausethe Persian Gulf is so shallow, bottom salinitylargely mirrors surface salinity. Inflows offresher waters from the Gulf of Oman in thesoutheast and from the Tigris and Euphratesrivers in the northwest bracket the high salinityregion along the coast of the Arabian Peninsula.Outflow of the dense water occurs throughoutthe year, with only a weak seasonal signal, ata mean salinity of 39.5 psu (Johns et al., 2003).Part of the outflow occurs in the surface layerin the southern part of the Straits of Hormuz.The surface layer transport has a strong seasonalcycle.

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