typical distributions of water...

C H A P T E R

4

Typical Distributions of WaterCharacteristics

4.1. INTRODUCTION

In this chapter, we describe the typical distri-butions of water properties such as temperature,salinity, oxygen, and nutrients. The propertieswere introduced in Chapter 3. Here we high-light distributions that are common, forinstance, to the Atlantic, Pacific, and IndianOceans, or to all subtropical regions, or to allthree equatorial regions, and so on. The over-view provides an essential framework for theheat and freshwater budgets of Chapter 5 andfor the detailed descriptions of properties andcirculation in each ocean basin presented in laterchapters. Summaries of some of the large-scalewater masses are included in Chapter 14.

Several central concepts are useful forstudying large-scale water properties. First,most water properties are initially set at the seasurface and are then modified within the oceanthrough a process called ventilation. Ventilationis the connection between the surface and theocean interior (similar to breathing). Second, theocean is vertically stratified in density, and flowwithin the ocean interior is primarily along isen-tropic (isopycnal) surfaces rather than acrossthem. That is, flow within the ocean interior isnearly adiabatic (without internal sources ofheat and freshwater). Third, as a result, waterproperties are helpful for identifying flow paths

from the surface into the interior, and for identi-fying forcing andmixing processes and locations.This is related to the usefulness of the concept ofwater masses, defined in the next section.

Most water characteristics have large andtypical variations in the vertical direction, whichencompasses an average of 5 km in the deepocean, whereas variations of similar magnitudein the horizontal occur over vastly greaterdistances. For instance, near the equator, thetemperature of the water may drop from 25�Cat the surface to 5�C at a depth of 1 km, but itmay be necessary to go 5000 km north or southfrom the equator to reach a latitude where thesurface temperature has fallen to 5�C. Theaverage vertical temperature gradient (changeof temperature per unit distance) in this case isabout 5000 times the horizontal one. However,the more gradual horizontal variations areimportant: the horizontal density differencesare associated with horizontal pressure differ-ences that drive the horizontal circulation,which is much stronger than the vertical circula-tion. To illustrate the three-dimensional distri-butions of water properties and velocities, weuse a number of one- and two-dimensionalrepresentations, such as profiles, vertical sec-tions, and horizontal maps.

Much of the geographic variation in proper-ties in the oceans and atmosphere occurs in

67Descriptive Physical Oceanography

� 2011. Lynne Talley, George Pickard, William Emery and James Swift.

Published by Elsevier Ltd. All rights reserved.

the north-south (meridional) direction. Proper-ties are often much more uniform in the east-west (zonal) direction. A principal exception tothe latter is the important zonal variation nearboundaries, especially on the west sides ofocean basins. In addition to the major oceanbasins, we also refer to general regions that aremainly distinguished by latitude ranges.

The equatorial region refers to the zone withinseveral degrees of the equator, while tropicalrefers to zones within the tropics (23�N or �Sof the equator). In the equatorial region, theeffect of the earth’s rotation on currents isminimal, leading to very distinctive currentsand water property distributions comparedwith other regions. Within the tropics, there isnet heating at the sea surface. The distinctionbetween equatorial and tropical is often signifi-cant, but when the two are to be lumpedtogether, they are referred to as the low latitudes.In contrast, the regions near the poles, north andsouth, are called the high latitudes. Subtropicalrefers to mid-latitude zones poleward of thetropics, characterized by atmospheric high pres-sure centers. Polar is used for the Arctic andAntarctic regions, where there is net coolingand usually sea ice formation. Subpolar refersto the region between the strictly polar condi-tions and those of the temperate mid-latitudes.The most marked seasonal changes takeplace in the temperate zones (approximately30e60�N or �S).

Throughout this chapter and in subsequentchapters we refer to the concept of a watermass, which is a body of water that has had itsproperties set by a single identifiable process.This process imprints properties that identifythe water mass as it is advected and mixedthrough the ocean. Most water masses areformed at the sea surface where their identi-fying characteristics are directly related tosurface forcing, but some water masses acquiretheir characteristics (e.g., an oxygen minimum)through subsurface processes that might bebiogeochemical as well as physical. Some water

masses are nearly global in extent while otherwater masses are confined to a region, such asa gyre in a specific ocean basin. Water masseshave been given names that are usually capital-ized. Some water masses have several names,simply because of the history of their study.Water type and source water type are usefulrelated concepts; a water type is a point in prop-erty space, usually defined by temperature andsalinity, and a source water type is the watertype at the source of the water mass (e.g.,Tomczak & Godfrey, 2003).

Each water mass is introduced in terms of(1) its identifying characteristic(s) and (2) theocean process that creates that specific charac-teristic. Descriptive physical oceanographersoften first identify an extremum or interestingcentral characteristic. They then seek to findthe process that created that characteristic.Once the process is identified, additionalinformation about the process is used to refinethat water mass’s definition, for example, thefull density range might be assigned to thewater mass. Information about the process andwater mass distribution assists in studying thecirculation.

The Mediterranean Water (MW) (Chapter 9)is an example of a water mass with a simpleidentifying characteristic. MW is a salinitymaximum layer in the North Atlantic at mid-depth (1000e2000 m) and a lateral salinitymaximum on any quasi-horizontal surfacecutting through the layer (e.g., Figure 6.4). Itssource is the saline outflow of water from theMediterranean through the Strait of Gibraltar.Its high salinity results from excess evapora-tion and internal dense water formation withinthe Mediterranean Sea (see the textbook Website, which contains supplementary materials,http://booksite.academicpress.com/DPO/ toview Section S8.10.2; “S” denotes supplementalmaterial). The MW density range within theNorth Atlantic is a function of both its highdensity at the Strait of Gibraltar and alsointense mixing with ambient (stratified) North

4. TYPICAL DISTRIBUTIONS OF WATER CHARACTERISTICS68

Atlantic water as it plunges down the conti-nental slope after it exits the Strait of Gibraltar.

Subtropical Mode Water (STMW) is anotherexample of a water mass with a simple verticalextremum; in this case its thickness (verticalhomogeneity) is compared with waters aboveand below it. A type of STMW is found ineach ocean’s subtropical gyre (Sections 9.8.2,10.9.1, and 11.8.1). STMW originates in a thicksurface winter mixed layer that is then advecteddown along isopycnals into the ocean interior.STMWretains its signature of relative thickness,just as the MW retains its signature of highsalinity. Slow mixing within the ocean interioreventually erodes these extrema, but theypersist far enough from their sources to beuseful tracers of flow.

Many other major world water masses areintroduced in this chapter. Detailed descriptionsof them and of their formation processes areprovided in the ocean basin chapters (9 through13), with a final summary in Chapter 14.

Taking into account the whole set of oceanproperties and information about water mas-ses, it is useful to think of the vertical structurein terms of four layers: upper, intermediate,deep, and bottom. The upper layer containsa surface mixed layer, thermocline and/orhalocline, pycnocline, and other structuresembedded in these (see descriptions withrespect to temperature and density in Sections4.2 and 4.4). The upper layer is in contactwith the atmosphere, either directly or throughbroad flow (relatively directly) into the upperocean through the subduction processdescribed in Sections 4.4.1 and 7.8.5. The inter-mediate, deep, and bottom layers are all belowthe pycnocline, or at most, embedded withinthe bottom of it. These layers are identified bywater masses that indicate surface origins,with respect to location and formation pro-cesses, and relative age.

Before describing some typical distributionsof each of the water properties, the followinginformation on ocean water temperatures

and salinities is given for orientation (seeFigure 3.1):

1. 75% of the total volume of the ocean waterhas a temperature between 0 and 6�C andsalinity between 34 and 35 psu,

2. 50% of the total volume of the oceans hasproperties between 1.3 and 3.8�C andbetween 34.6 and 34.7 psu,

3. The mean temperature of the world ocean is3.5�C and the mean salinity is 34.6 psu.

4.2. TEMPERATUREDISTRIBUTION OF THE OCEANS

The ocean and atmosphere interact at the seasurface. Surface forcing from the atmosphereand sun sets the overall pattern of sea surfacetemperature (SST) (Figure 4.1). High SST in thetropics is due to net heating, and low SST athigh latitudes is due to net cooling. Beyond thissimple meridional variation, the more complexfeatures of SST result from ocean circulationand spatial variations in atmospheric forcing.The ocean’s surface, which could include seaice, provides the forcing at the bottom of theatmosphere through various kinds of heatforcing and as a source of water vapor.

SST ranges from slightly more than 29�C inthe warmest regions of the tropics, to freezingtemperature (about �1.8�C; Figure 3.1) in ice-forming regions, with seasonal variations espe-cially apparent at middle to high latitudes.

Below the sea surface, we refer only to poten-tial temperature so that the pressure effect ontemperature is removed (Section 3.3 andFigure 3.3). The vertical potential temperaturestructure can usually be divided into threemajor zones (Figure 4.2): (1) the mixed layer,(2) the thermocline, and (3) the abyssal layer.This structure is typical of low and mid-latitudes with high SST. Relative to the four-layer structure introduced in Section 4.1, thefirst two zones are within the upper layer and

TEMPERATURE DISTRIBUTION OF THE OCEANS 69

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FIGURE 4.1 (a) Surface temperature (�C) of the oceans in winter (January, February, March north of the equator; July,August, September south of the equator) based on averaged (climatological) data from Levitus and Boyer (1994). (b) Satelliteinfrared sea surface temperature (�C; nighttime only), averaged to 50 km and 1 week, for January 3, 2008. White is sea ice.(See Figure S4.1 from the online supplementary material for this image and an image from July 3, 2008, both in color). Source:From NOAA NESDIS (2009).


the third temperature zone contains the inter-mediate, deep, and bottom layers.

In high latitudes where SST is low, this struc-ture differs, and can have a mixed layer,a vertical temperature minimum and under-lying maximum near the sea surface, and thenthe thermocline and abyssal layer.

The mixed layer (Section 4.2.2) is a surfacelayer of relatively well-mixed properties. Insummer in low latitudes, it can be very thin ornon-existent. In winter at middle to high lati-tudes, it can be hundreds of meters thick, andin isolated deep convection regions, the mixedlayer can be up to 2000 m thick. Mixed layersare mixed by both wind and surface buoyancyforcing (air-sea fluxes). The thermocline (Sections4.2.3 and 4.2.4) is a vertical zone of rapidtemperature decrease with a depth of roughly1000 m. In the abyssal layer, between the ther-mocline and ocean bottom, potential tempera-ture decreases slowly. At high latitudes,a near-surface temperature minimum (dichother-mal layer) is often found, a holdover from a cold

winter mixed layer that is “capped” withwarmer waters in other seasons (Figure 4.2c);the underlying temperature maximum (meso-thermal layer) results from advection of watersfrom somewhat warmer locations. This temper-ature structure is stable because there is strongsalinity stratification, with fresher water in thesurface layer.

Typical temperatures at subtropical latitudesare 20�C at the surface, 8�C at 500 m, 5�C at1000 m, and 1e2�C at 4000 m. All of these valuesand the actual shape of the temperature profileare a function of latitude, as shown by the threedifferent profiles in Figure 4.2.

There are some notable additions to this basicthree-layered structure. In all regions, springand summer warming produces a thin warmlayer overlying the winter’s mixed layer. Inthe western subtropical regions as well asother regions, there are often two thermoclineswith a less stratified (more isothermal) layer(thermostad) between them, all within the upper1000 m (Figure 4.2b). In some regions another

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FIGURE 4.2 Typical potential temperature (�C)/depth (m) profiles for the open ocean in (a) the tropical western NorthPacific (5�N), (b) the western and eastern subtropical North Pacific (24�N), and (c) the western subpolar North Pacific(47�N). Corresponding salinity profiles are shown in Figure 4.16.


mixed layer is found at the very bottom(“bottom boundary layer”) and can be up to100 m thick.

In many parts of the ocean, density is a strongfunction of temperature (Chapter 3), and has thesame layered structure as temperature; that is,an upper layer, a pycnocline with rapidlyincreasing density, and an abyssal zone. Salinityusually has a more complicated vertical struc-ture (Section 4.3). In regions of high precipita-tion and/or runoff (such as subpolar and highlatitude regions and parts of the tropics), salinitymay be more important than temperature insetting the vertical density structure, especiallyin the upper layer, since the water columnmust be vertically stable on average. A typicalvertical salinity profile in these regions includesa relatively fresh surface layer with a haloclineseparating the surface layer from the highersalinity water below. The higher underlyingsalinity is an indication of a sea-surface sourceof water in a less rainy area. On the otherhand, in the subtropics where the sea surfacesalinity is dominated by evaporation, surfacewater is usually more saline than the underlyingwater. Here temperature clearly dominates thevertical stability.

This three-layered structure is simpler thanour simplest description of overall watermass structure, for which at least four layersare usually required (Section 4.1). The abyssallayer, in terms of temperature, usually includesat least two or possibly three separate watermass layers: intermediate, deep, and bottomwaters. However, potential temperature is rela-tively low in all of these water mass-basedlayers, declining toward the bottom, and isnot a useful indicator of these water masslayers.

4.2.1. Surface Temperature

The temperature distribution at the surface ofthe open ocean is approximately zonal, with thecurves of constant temperature (isotherms)

running roughly east-west (Figure 4.1). Nearthe coast where the currents are diverted bythe boundaries, the isotherms may swing morenearly north and south. Also, along the easternboundaries of the oceans, surface temperaturesare often lower due to upwelling of subsurfacecool water, for example, along the west coastof North America in summer, causing theisotherms to trend equatorward. Upwellingalso causes lower surface temperatures in theeastern equatorial Pacific and Atlantic.

The open ocean SST, averaged over all longi-tudes and displayed as a function of latitude(Figure 4.3), decreases from as high as 28�Cjust north of the equator to nearly �1.8�C nearsea ice at high latitudes. This distribution corre-sponds closely with the input of short-waveradiation (mainly from the sun), which is high-est in the tropics and lowest at high latitudes(Section 5.4.3). The corresponding mean zonalsurface salinity and density are also shown.Salinity and density are discussed in Sections4.3 and 4.4. Density is dominated by tempera-ture. Salinity has subtropical maxima in boththe Northern and Southern Hemispheres anda minimum just north of the equator.

Because many satellites observe SSTand SST-related quantities, many different SST productsare available, providing daily and longer termaverage maps with higher spatial and tem-poral resolution than the climatology based onin situ data shown in Figure 4.1a. Global SSTbased on infrared imagery for one week inJanuary (boreal winter, austral summer) isshown in Figure 4.1b. (The equivalent imagefor July is included in the online supplementarymaterials as Figure S4.1.) The structures of oceancurrents, fronts, upwelling regions, eddies, andmeanders are more apparent in these nearlysynoptic SST images.

Non-zonal features of global SSTs that aremost apparent and important to note in Figure4.1 include the warm pool and the cold tongue.The warm pool is the warmest SST region,located in the western tropical Pacific, through


the Indonesian passages, and into the tropicalIndian Ocean. The cold tongue is the narrowtongue of colder water along the equator inboth the eastern Pacific and Atlantic. Thisforms due to upwelling of thermocline wateralong the equator. Because the thermocline isshallower in the eastern Pacific and Atlanticthan in the west, upwelling brings up colderwater in the east.

In each ocean, warm regions are centered inthe west, off the equator. Cooler waters cycleequatorward in the central and eastern parts ofeach ocean. These SST patterns reflect the

anticyclonic circulation of the subtropical gyres(clockwise in the Northern Hemisphere, coun-terclockwise in the Southern Hemisphere),which advects warm water away from thetropics and cooler water toward the equator.There are also regions of warmer water in theeastern tropical North Pacific and NorthAtlantic. These are found east of the subtropicalcirculation and north of the cold tongue;high temperatures are not suppressed byeither the anticyclonic circulation or equatorialupwelling.

In the subpolar North Pacific and NorthAtlantic, there is, again, evidence of the circula-tion in the SST pattern. Here the gyresare cyclonic (counterclockwise in the NorthernHemisphere). Warmer waters are advectednorthward in the eastern parts of these circula-tions (along the coast of British Columbia andalong northern Europe). Warmer water extendsfar to the north in the Atlantic toward the Arctic,along the Norwegian coast. Cold waters arefound in the western parts of these circulations,along the Kamchatka/Kuril region in the Pacificand Labrador/Newfoundland region in theAtlantic.

In the Southern Ocean, SST is not exactlyzonal. This reflects excursions in the AntarcticCircumpolar Current (ACC), which is also notzonal. Colder waters are farther north in theAtlantic and Indian Oceans and pushed south-ward in the Pacific (Section 13.4).

In the satellite SST maps (Figures 4.1b andS4.1 from the online supplementary material),eddy-scale (100e500 km) features are apparenteven with global maps, particularly where thecolor scaling provides large contrasts. Espe-cially visible are the large wavelike structuresin the equatorial regions; length scales of trop-ical waves are longer than at higher latitudesso they are better resolved in this map. Thewaves around the Pacific’s equatorial coldtongue are the Tropical Instability Waves(TIWs), with timescales of about a month(Section 10.7.6).

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FIGURE 4.3 Variation with latitude of surface (a)temperature, (b) salinity, and (c) density averaged for alloceans for winter. North of the equator: January, February,and March. South of the equator: July, August, andSeptember. Based on averaged (climatological) data fromLevitus and Boyer (1994) and Levitus et al. (1994b).


4.2.2. Upper Layer Temperature andMixed Layer

Within the ocean’s near-surface layer, prop-erties are sometimes very well mixed vertically,particularly at the end of the night (diurnalcycle) and in the cooling seasons (seasonalcycle). This is called the mixed layer. This layeris mixed by the wind and by buoyancy lossdue to net cooling or evaporation at the seasurface. It is unmixed by warming and precip-itation at the sea surface and by circulationswithin the mixed layer that move adjacentmixed waters of different properties overeach other. Processes that create and destroythe mixed layer are described in much greaterdetail in Section 7.4.1. Here we focus on theobserved structure and distribution of mixedlayers.

As a rule of thumb, wind-stirred mixedlayers do not extend much deeper than 100 or150 m and can reach this depth only at theend of winter. On the other hand, infrequentvigorous cooling or evaporation at the seasurface can cause the mixed layer to deepenlocally to several hundred meters, or evenbriefly in late winter to more than 1000 m inisolated deep convection locations. Mixedlayers in summer may be as thin as 1 or 2 m,overlying a set of remnant thin mixed layersfrom previous days with storms, and thickerremnant mixed layers from winter. Becausethe mixed layer is the surface layer thatconnects the ocean and atmosphere, andbecause sea-surface temperature is the mainway the ocean forces the atmosphere, observa-tions of the mixed layer and understandinghow it develops seasonally and on climate time-scales is important for modeling and under-standing climate.

A given vertical profile will not usuallyexhibit a thick, completely mixed layer ofuniform temperature, salinity, and density.Most often, there will be small steps, nearlydiscontinuities, in the profiles due to daily

restratification and remixing with layers slidingin from nearby. For a careful study of the mixedlayer, the investigator assigns the mixed layerdepth based on examination of every verticalprofile. However, for general use (e.g., with thegrowing profiling float data set, or for use inupper ocean property mapping for fisheries,climate prediction, or navigational use), it isnot feasible to examine each profile, and it isimportant to have consistent criteria for assign-ing the mixed layer depth. Functional defini-tions of mixed layer depth have beendeveloped, mostly based on finding a settemperature or density difference between thesurface observation and deeper observations;this is the so-called “threshold method.” Intropical and mid-latitudes, temperature-baseddefinitions are adequate, but at higher latitudes,it is common to find a subsurface temperaturemaximum lying underneath a low salinitysurface layer. Currently, the most commonlyused criterion is a density difference of sq ¼0.03 kg/m3 or temperature difference of 0.2�C,as used in the mixed layer maps shown inFigure 4.4a,b (deBoyer Montegut et al., 2004).Other treatments have employed larger thresh-olds (e.g., 0.8�C in Kara, Rochford, & Hurlburt,2003) or more detailed criteria that fit theobserved vertical profiles rather than relyingon a threshold (Holte & Talley, 2009). A globalmap of the maximum mixed layer depth, usingthe latter method, is shown in Figure 4.4c (Holte,Gilson, Talley, & Roemmich, 2011).

In all regions, winter mixed layers are muchthicker than summer mixed layers. The mainfeatures of the global winter mixed layer mapsare the thick mixed layers in the northern NorthAtlantic and in a nearly zonal band in theSouthern Ocean. These regions correspond tomaxima in anthropogenic carbon uptake(Sabine et al., 2004), so they have practical impli-cations for global climate. These thick wintermixed layers are the main source of ModeWaters, which are identified as relatively thicklayers in the upper ocean (Section 4.2.3).


Mixed layer development is affected by theamount of turbulence in the surface layer. Thisturbulence is generated by breaking surfaceand internal waves generated by the wind,decreasing with increasing depth. Mixed layerdevelopment can also be affected by Langmuircells, which are transient helical circulations(in the vertical plane) aligned parallel to thewind (Section 7.5.2). These create the “windrows” sometimes seen at the sea surface underthe wind, where the water is pushed together,or converges, in the Langmuir cells, which reachto about 50 m depth and 50 m width, and can

create turbulence that affects mixing in themixed layer.

Another dynamical phenomenon present inthe near-surface layer is the Ekman responseto wind forcing, which forces flow in theocean’s surface layer off to the right of thewind in the Northern Hemisphere (and tothe left in the Southern Hemisphere), becauseof the Coriolis force (Section 7.5.3). Turbulencein the surface layer acts like friction. In theNorthern Hemisphere, each thin layer withinthe surface layer pushes the one below it a littlemore to the right, and with a little smaller

FIGURE 4.4 Mixed layer depth in (a) January and (b) July, based on a temperature difference of 0.2�C from the near-surface temperature. Source: From deBoyer Montegut et al. (2004). (c) Averaged maximum mixed layer depth, using the 5deepest mixed layers in 1� � 1� bins from the Argo profiling float data set (2000e2009) and fitting the mixed layer structureas in Holte and Talley (2009). This figure can also be found in the color insert.


velocity than the layer above. This creates an“Ekman spiral” of decreasing velocities withincreasing depth. The whole spiral occurswithin the top 50 m of the ocean. If all of thevelocities are added together to calculate thetotal transport in the Ekman layer, the net effectis that this Ekman transport moves at exactlyright angles to the wind direction d to theright in the Northern Hemisphere and to theleft in the Southern Hemisphere. Ekman veloc-ities are small and do not generate turbulence.Thus they have no direct effect on mixed layerdevelopment and are affected by the upperlayer turbulence but not by the mixed layerstratification. The Ekman response is crucial,however, for conveying the effect of the windto the ocean, for development of the large-scaleand long timescale ocean circulation, as alsodescribed in Chapter 7.

4.2.3. Thermocline, Halocline, andPycnocline

Below the surface layer, which can be wellmixed or can include messy remnants of localmixing and unmixing, temperature begins todecrease rapidly with depth. This rapid dec-rease ceases after several hundred meters,with only small vertical changes in tempera-ture in the deep or abyssal layer that extendson down to the bottom. The region of highervertical temperature gradient (rate of decreaseof temperature with increasing depth) is calledthe thermocline. The thermocline is usuallya pycnocline (high vertical density gradient). Itis often hard to precisely define the depthlimits, particularly the lower limit, of the ther-mocline. However, in low and middle lati-tudes, a thermocline is always present atdepths between 200 and 1000 m. This isreferred to as the main or permanent thermo-cline. In polar and subpolar waters, where thesurface waters may be colder than the deepwaters, there is often no permanent thermo-cline, but there is usually a permanent halocline

(high vertical salinity gradient) and associatedpermanent pycnocline.

The continued existence of the thermoclineand pycnocline requires explanation. There aretwo complementary concepts, one based onvertical processes only, and the other based onhorizontal circulation of the waters that formthe thermocline away from where they outcropas mixed layers in winter. Both concepts areimportant and work together.

The vertical processes that affect the thermo-cline are downward transfer of heat from the seasurface and either upwelling or downwelling(these depend on the location in the ocean andon what creates the vertical motion). One mightexpect that as the upper waters are warmest,heat would be transferred downward by diffu-sion despite the inhibiting effect of the stabilityin the pycnocline/thermocline, and that thetemperature difference between the upper andlower layers would eventually disappear.However, the deeper cold waters are fed contin-uously from the sea surface at higher latitudes(deep and bottom water formation regions,mainly in the northernmost North Atlantic andGreenland Sea and in various regions aroundAntarctica). These deep inflows maintain thetemperature difference between the warmsurface waters and cold deep waters. The deepwaters upwell and warm up through down-ward diffusion of heat. If upwelling from thebottommost layers to near the surface occursthrough the whole ocean, the upward speedwould be 0.5e3.0 cm/day. Unfortunately thesespeeds are too small to accurately measurewith current instruments, so we are unable totest the hypothesis directly. The result of thedownward vertical diffusion of heat balancedby this persistent upwelling of the deepestcold waters results in an exponential verticalprofile of temperature (Munk, 1966), whichapproximates the shape of the permanentthermocline.

This simplified vertical model of the thermo-cline is depicted in Figure 4.5, which shows


a typical vertical temperature profile in theupper ocean containing the thermocline. Theresult of downward diffusion of heat is labeledas AvT

vz and the result of upward vertical advec-tion of colder, deeper water is labeled “wT”.(Equation (7.46) shows these two terms arethe vertical integrals of the vertical diffusionand vertical advection terms, assuming cons-tant eddy diffusivity A and constant verticalvelocity w. In this simplest of thermoclinemodels, it is assumed that downward diffusionof heat is entirely balanced by upward advec-tion.) If we assume that the difference betweenthese two terms is a constant, we have an equa-tion with an exponential solution for tempera-ture T, which in many cases approximates theshape of the thermocline. We can use similararguments relative to the vertical distributionof tracers like dissolved oxygen except thatsuch tracers can have both sources and sinkswithin the water column, ultimately resultingin subsurface maxima or minima.

A second, more horizontal, adiabatic andcomplementary process for maintaining thethermocline/pycnocline was suggested byIselin (1939) and further developed by Luyten,

Pedlosky, and Stommel (1983; Section 7.8.5). Ise-lin observed that the surface temperature-salinity relation along a long north-south swathin the North Atlantic strongly resembled the T-Srelation in the vertical (Figure 4.6). He hypothe-sized that the waters in the subtropical thermo-cline therefore originate as surface watersfarther to the north. As they move south, thecolder surface waters subduct beneath thewarmer surface waters to the south (usingthe term from Luyten et al., borrowed from platetectonics). Subduction of many layers builds upthe temperature, salinity, and density structureof the main pycnocline (thermocline) in thesubtropical gyre. This process is adiabatic, notrequiring any mixing or upwelling across iso-pycnals. Such one-dimensional diapycnalprocesses would then modify the thermoclinestructure, smoothing it out.

Double diffusion (Section 7.4.3) is anothervertical mixing process that might affect thethermocline (pycnocline). This process mightmodify the relation between temperature andsalinity within the pycnocline, smoothing theprofile that results from adiabatic subduction(Schmitt, 1981).

The main thermoclines/pycnoclines of theworld’s subtropical gyres are permanent fea-tures. The temperature-salinity relation in thethermocline of each subtropical gyre is shownin Figure 4.7. The main thermoclines are identi-fiable in temperature/salinity relations, andthey have a common formation history that issome combination of subduction and verticalupwelling/diffusion. Therefore, the waters inthe thermocline can be identified as a watermass. This is the first water mass that we intro-duce systematically, rather than as an example.The thermocline water mass is Central Water.Central Water differs from typical water massesbecause it has a large range of temperature,salinity and density.

So far, we have referred to the “main,” orpermanent, thermocline. There are also perma-nent, double thermoclines in some large but

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FIGURE 4.5 Vertical processes that can maintain thethermocline in a simplified one-dimensional model.


geographically restricted regions. For instance,two thermoclines are found in the SargassoSea just south of the Gulf Stream. A layer oflower vertical stratification separates the twothermoclines. The layer of lower stratification

is called a thermostad (or pycnostad for the equiv-alent density layer).

The thermostad/pycnostad is often given thewater mass name,ModeWater. This is the secondwater mass that we introduce. Mode Water is

FIGURE 4.6 Temperature-salinity along surface swaths in the North Atlantic (dots and squares), and in the vertical (solidcurves) at stations in the western North Atlantic (Sargasso Sea) and eastern North Atlantic. Source: From Iselin (1939).

FIGURE 4.7 Potential temperature-salinityrelation in the thermocline of each subtropicalgyre. These are the CentralWaters. R is the bestfit of a parameter associated with doublediffusive mixing (Section 7.4.3). Source: FromSchmitt (1981).


considered a water mass because it is identifiedby a particular characteristic (a verticalextremum in layer thickness), and because ithas a specific formation process (subduction ofthick mixed layers). The name “Mode Water”was introduced by Masuzawa (1969). Volumet-rically there is more water in a particulartemperature/salinity range than in the thermo-clines above and below it, so Mode Waterappears as a mode in the distribution of volumein temperature/salinity space.

In the region where the Mode Water outcropsas a thick mixed layer, the overlying thermoclineis actually a seasonal thermocline that disap-pears in late winter. After ModeWater subducts,its thermostad is embedded in the permanentthermocline, creating a double thermocline.

4.2.4. Temporal Variations ofTemperature in the Upper Layerand Thermocline

The temperature in theupper zoneand into thethermocline varies seasonally, particularly inmid-latitudes. In winter the surface temperatureis low, waves are large, and the mixed layer isdeep and may extend to the main thermocline.In summer the surface temperature rises, thewater becomes more stable, and a seasonal ther-mocline often develops in the upper layer.

The growth and decay of the seasonal ther-mocline is illustrated in Figure 4.8a usingmonthly mean temperature profiles fromMarch 1956 to January 1957 taken at OceanWeather Station P (“Papa”) in the northeastern(subpolar) North Pacific. FromMarch to August,the temperature gradually increases due toabsorption of solar energy. A mixed layer fromthe surface down to 30 m is evident all thetime. After August there is a net loss of heatand continued wind mixing; these erode awaythe seasonal thermocline until the isothermalcondition of March is approached again. Notethat March does not have the maximum heat

loss; rather, it is the last month of cooling beforeseasonal heating begins. Therefore total heatcontent is lowest in March. In tropical andsubtropical locations, the summer mixed layermay be even thinner.

These same data may be presented in alterna-tive forms; for instance, as a time series showing

FIGURE 4.8 Growth and decay of the seasonal thermo-cline at 50�N, 145�W in the eastern North Pacific as (a)vertical temperature profiles, (b) time series of isothermalcontours, and (c) a time series of temperatures at depthsshown.


the depths of the isotherms during the year(Figure 4.8b). (The original data include thealternate months, which were omitted fromFigure 4.8a to avoid crowding.) In Figure 4.8cthe temperatures are plotted at selected depths.The different forms in which the thermoclineappears in these three presentations should benoted. In Figure 4.8a, the permanent thermo-cline appears as a maximum gradient region inthe temperature/depth profiles. In Figure 4.8b,the thermocline appears as a crowding of theisotherms, which rises from about 50 m inMay to 30 m in August and then descends to100 m in January. In Figure 4.8c, the thermoclineappears as a wide separation of the 20- and 60-misobaths between May and October, andbetween the 60- and 100-m isobaths after thatas the thermocline descends.

At the highest latitudes, the surface tempera-tures are much lower than at lower latitudes,while the deep-water temperatures are littledifferent. As a consequence, the main

thermocline might not be present at high lati-tudes, and only a seasonal thermocline mightoccur. In high northern latitudes, there is oftena layer of cold water at 50e100 m (Figure 4.2c),with temperatures as low as �1.6�C, sand-wiched between the warmer surface anddeeper layers. As described at the beginning ofSection 4.2, this cold layer is referred to as thedichothermal layer. The warmer surface wateris often just seasonal, and the thermocline over-lying the dichothermal layer is thereforeseasonal.

Figure 4.9 shows the annual range of surfacetemperature over the globe. Annual variationsat the surface rise from 1e2�C at the equatorto between 5 and 10�C at 40� latitude in theopen ocean, then decrease toward the polarregions (due to the heat required in the meltingor freezing processes where sea ice occurs).Near the coast, larger annual variations(10e20�C) occur in sheltered areas and in thewestern subtropical regions of the Northern

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FIGURE 4.9 Annual range of sea surface temperature (�C), based on monthly climatological temperatures from theWorld Ocean Atlas (WOA05) (NODC, 2005a, 2009).


Hemisphere, where the Kuroshio and GulfStream are located and where surface heat lossis highest (Section 5.5, Figure 5.12). Theseannual variations in temperature decrease withdepth and are rarely perceptible below100e300 m. The maximum temperature at thesurface occurs at the end of the warming season,in August/September in the Northern Hemi-sphere, and the minimum at the end of the cool-ing season, during February/March. Below thesurface, the times of occurrence of the maximaand minima are delayed by as much as twomonths relative to the surface.

Diurnal variations of SST had been thought tobe small (<0.4�C) prior to satellite observations.Such measurements, verified by in situ observa-tions from a moored buoy in the Sargasso Seaover a period of two years (Stramma et al.,1986), have shown that diurnal variations to1�C are common with occasionally highervalues, up to 3e4�C. The larger diurnal varia-tions of 1�C or more are observed in conditionsof high insolation (solar radiation) and lowwind speed, and are generally limited to theupper few meters of water. Similar diurnal vari-ation has been observed elsewhere in the NorthAtlantic and in the Indian Ocean. In shelteredand shallow waters along the coast, values of2e3�C are common.

4.2.5. Deep-Water Temperature andPotential Temperature

Below the thermocline, the temperatureslowly decreases with increasing depth. (Thisvertical temperature change is much smallerthan through the thermocline.) In the deepestwaters, temperature can rise toward the bottom,almost entirely because the high pressure thatcompresses the water and raises its temperatureadiabatically (Section 3.3.3, Figure 3.3). To inter-pret variations in temperature, even in shallowwaters over a continental shelf as well as fromthe surface to thousands of meters, potentialtemperature (q) should always be used. Potentialtemperature reflects the original temperature ofthe water when it was near the sea surface.

An example of this difference between in situand potential temperature is shown in Table 4.1and in Figure 4.10 using data collected in 1976by the R/V T. Washington from the MarianaTrench (the deepest trench in the world ocean).While temperature (T) reaches a minimum atabout 4500 m and thereafter increases towardthe bottom, potential temperature is almostuniform. (Salinity also is almost uniformbetween 4500 m and the deepest observationas are potential densities relative to any refer-ence pressure.) Uniform properties from

TABLE 4.1 Comparison of in situ and Potential Temperatures and Potential Densities Relative to the Sea Surface(sq), 4000 dbar (s4) and 10,000 dbar (s10) in the Mariana Trench in the Western North Pacific

Depth (m) Salinity (psu) Temperature (�C) q (�C) sq (kg mL3) s4 (kg mL3) s10 (kg mL3)

1487 34.597 2.800 2.695 27.591 45.514 69.495

2590 34.660 1.730 1.544 27.734 45.777 69.903

3488 34.680 1.500 1.230 27.773 45.849 70.015

4685 34.697 1.431 1.028 27.800 45.898 70.090

5585 34.699 1.526 1.004 27.803 45.904 70.099

6484 34.599 1.658 1.005 27.803 45.904 70.099

9940 34.700 2.266 1.007 27.804 45.904 70.099

Data from R/V T. Washington, 1976.


4500e9940 m imply that the trench is filled withwater that passes over the sill into the trench,and that there is no other source of water. Theslight increase in potential temperature withdepth might be due to the weak geothermal hea-ting acting on this nearly stagnant thick layer.

It is not necessary to go to the deepest part ofthe ocean to see the important differencesbetween in situ temperature and potential

temperature. Through most of the deep ocean,there is a temperature minimum well abovethe ocean bottom, with higher temperature atthe bottom. However, potential temperaturedecreases to the ocean bottom almost every-where. This is because the densest waters thatfill the oceans are also the coldest, since salinityvariations are mostly too weak to control thedensity stratification in the deep waters. There

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FIGURE 4.10 Mariana Trench:(a) in situ temperature, T, andpotential temperature, q (�C); (b)salinity (psu); (c) potential densitysq (kg m�3) relative to the seasurface; and (d) potential densitys10 (kg m�3) relative to 10,000 dbar.


are some limited exceptions to the monotonicdecrease with depth: in localized regions ofdensest water formation, at some mid-oceanridges where geothermal heating slightlywarms waters right on the ridges, and in thecentral South Atlantic where there is significantvertical salinity variation at mid-depth(Figure 4.11b).

A global map of potential temperature at theocean bottom in the deep ocean (>3500 mdepth) is shown in Chapter 14 (Figure 14.14b).The bottom temperature distribution is mostlyset by the two sources of bottom water, fromthe Antarctic and the Nordic Seas. (Mid-oceanridges also result in bottom temperature varia-tions since they jut upward into warmerwaters.) Bottom waters of Antarctic origin arethe coldest; bottom temperatures are near thefreezing point near Antarctica, with tongues ofwater colder than 0�C extending northwardinto the deep basins of the Southern Hemi-sphere. Bottom waters of northern Atlanticorigin (which arise from overflows from theNordic Seas) are considerably warmer withtemperatures around 2�C.

4.2.6. Vertical Sections of PotentialTemperature

We now view potential temperature usingmeridional cross-sections through each of thethree oceans (Figures 4.11a, 4.12a, and 4.13a) toidentify common and typical features. Salinityand potential density sections are also shownto keep the vertical sections from each oceantogether. The salinity and density distributionsare described in Sections 4.3 and 4.4.

In all oceans, the warmest water is in theupper ocean with the highest temperatures inthe tropics. In the subtropics, the warm waterfills bowl-shaped regions. These bowls definethe upper ocean circulations, with westwardflow on the equatorward side of the bowlsand eastward flow on the poleward side ofthe bowls. Potential temperature decreases

downward through the thermocline into muchmore uniform, colder temperatures at depth.The coldest water is found at the surface athigh latitudes (and is vertically stable becauseof low salinity surface water). The coldest waterin these sections is in the Antarctic, since thenorthern ends of the sections do not extendinto the Arctic. In the Antarctic, the coldisotherms slope steeply downward between 60and 50�S. This marks the eastward flow of theACC (Chapter 13).

There are distinct differences in potentialtemperature distributions between the Northernand Southern Hemispheres. The cold surfacewaters are much more extensive in the south.Even the two bowls of higher temperature arenot symmetric; the southern bowl is more exten-sive than its northern counterpart. In the deeppart of the Atlantic, Pacific, and Indian Oceans,the coldest waters are in the south (in theAntarctic) and the potential temperatures areslightly higher in the north.

4.3. SALINITY DISTRIBUTION

The mean salinity of the world ocean is 34.6psu, based on integrating the climatologicaldata in Java Ocean Atlas (Osborne & Swift,2009; see the online supplementary materialslocated on the Web site for this text). There aresignificant differences between the oceanbasins. The Atlantic, and especially the NorthAtlantic, is the saltiest ocean and the Pacific isthe freshest (excluding the Arctic and SouthernOcean, which are both fresher than the Pacific).These basin differences are illustrated inFigure 4.14,which shows themean salinity alongwell-sampled hydrographic sections, averagedzonally, and from top to bottom of the ocean.

Salinity sections from south to north in eachocean are included in Figures 4.11, 4.12, and4.13. The following descriptions refer back tothese sections. It is apparent after comparingsalinity, potential temperature, and potential

SALINITY DISTRIBUTION 83

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FIGURE 4.11 (a) Potential temperature (�C), (b) salinity (psu), (c) potential density sq (top) and potential density s4 (bottom) (kg m�3), and (d)oxygen (mmol/kg) in the Atlantic Ocean at longitude 20� to 25�W. Data from the World Ocean Circulation Experiment. This figure can also be found inthe color insert.

4.TYPIC

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density sections for each ocean that the salinitydistribution is more complex than temperatureand density. While potential temperaturedecreases monotonically to the bottom in mostplaces, salinity has marked vertical structure;from the simplicity of the density field, it isapparent that it is dominated by potentialtemperature. Salinity therefore functions in partas a tracer of waters, even as it affects densityin a small way.

More detailed depictions of the globalsalinity distribution and seasonal changes are

available in the climatological (seasonallyaveraged) data set from Levitus, Burgett, andBoyer (1994b). They also showed the data usedas the basis for the climatologies. There arefar more observations (~90%) in the NorthernHemisphere than in the Southern Hemisphere(~10%), and far more observations in summerthan in winter (e.g., Figure 6.13). (This is alsotrue of temperature observations.) This samp-ling bias is rapidly being corrected in the upper1800 m by the global profiling float program(Argo) that began in the 2000s.

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FIGURE 4.12 (a) Potential temperature (�C), (b) salinity (psu), (c) potential density sq (top) and potential density s4

(bottom; kg m�3), and (d) oxygen (mmol/kg) in the Pacific Ocean at longitude 150�W. Data from the World Ocean CirculationExperiment. This figure can also be found in the color insert.


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FIGURE 4.13 (a) Potential temperature (�C), (b) salinity (psu), (c) potential density sq (top) and potential density s4 (bottom; kg m�3), and (d) oxygen(mmol/kg) in the Indian Ocean at longitude 95�E. Data from the World Ocean Circulation Experiment. This figure can also be found in the color insert.

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4.3.1. Surface Salinity

Surface salinity in the open ocean rangesfrom 33 to 37. Lower values occur locally nearcoasts where large rivers empty and in thepolar regions where the ice melts. Higher valuesoccur in regions of high evaporation, such asthe eastern Mediterranean (salinity of 39) andthe Red Sea (salinity of 41). On average, theNorth Atlantic is the most saline ocean at

the surface (35.5 psu), the South Atlantic andSouth Pacific are less so (about 35.2 psu),and the North Pacific is the least saline (34.2psu), which reflects the ocean basin differencesin salinity over the whole ocean depth(Figure 4.14).

The salinity distribution at the ocean’s surfaceis relatively zonal (Figure 4.15), although not asstrongly zonal as sea-surface temperature.UnlikeSST, which has a tropical maximum and polar

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FIGURE 4.14 Mean salinity,zonally averaged and from top tobottom, based on hydrographicsection data. The overall meansalinity is for just these sections anddoes not include the Arctic,Southern Ocean, or marginal seas.Source: From Talley (2008).

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Winter surface salinity

FIGURE 4.15 Surface salinity (psu) in winter (January, February, and March north of the equator; July, August, andSeptember south of the equator) based on averaged (climatological) data from Levitus et al. (1994b).


minima, salinity has a double-lobed structure,with maxima in the subtropics in both hemi-spheres and minima in the tropics and subpolarregions. This meridional variation is alsoapparent in the global zonal average of surfacesalinity (Figure 4.3b). In that figure, the salinitymaximum just north of 60�N (with correspond-ing density deviation) results from dominanceof the North Atlantic waters over North Pacificat these latitudes. This is a combination of geog-raphy and the higher overall salinity of the NorthAtlantic; as the North Pacific closes off at theselatitudes, the zonal average mainly includesmore saline North Atlantic waters, even thoughinternally the subpolar North Atlantic watersare fresher than its subtropical waters.

Surface salinity is set climatologically by theopposing effects of evaporation (increasing it)and precipitation, runoff, and ice melt (alldecreasing it), mostly captured by the map ofevaporation minus precipitation (Figure 5.4a).The meridional salinity maxima of Figures 4.3and 4.15 are in the trade wind regions andsubtropical high pressure regions where theannual evaporation (E) exceeds precipitation(P), so that (E�P) is positive. On the other hand,the surface temperature maximum is near theequator because the balance of energy into thesea has a single maximum there. Just north ofthe equator, precipitation is high and surfacesalinity is lower because of the IntertropicalConvergence Zone (ITCZ) in the atmosphere.

Generally the regions of high positive evapo-ration minus precipitation (E�P) are displacedto the east of the subtropical salinity maxima.This lateral displacement results from the circu-lation (advection) of the surface waters, so thatsalinity is highest at the downstream end ofthe flow of upper ocean waters through theevaporation maxima.

4.3.2. Upper Layer Salinity

The vertical salinity distribution (Figure 4.16and sections in Figures 4.11, 4.12, and 4.13) is

more complicated than the temperature distri-bution. In the upper ocean, in the tropics, andsubtropics and parts of the subpolar regions,temperature dominates the vertical stability(density profile). In the deep ocean, beneaththe pycnocline, temperature also dominatesover salinity. Therefore, warmer water (lowerdensity) is generally found in the upper layersand cooler water (higher density) in the deeperlayers. Salinity can have much more verticalstructure, ranging from low to high, withoutcreating vertical overturn. (In subpolar andhigh latitudes, where surface waters are quitefresh and also cold, salinity does dominate thevertical stability.) As a consequence of its lessimportant role in dictating the density structure,salinity is a more passive tracer than tempera-ture. Thus, salinity can often be used as amarkerof the flow directions of water masses (minimaor maxima).

In the subtropics, salinity is high near the seasurface due to subtropical net evaporation.Salinity decreases downward to a minimum inthe vertical at 600e1000 m. Below this, salinityincreases to a maximum, with the exact depthsof the vertical minimum andmaximum depend-ing on the ocean. In the Atlantic and IndianOceans, the salinity maximum is at depths of1500e2000 m. In the Pacific, the maximumsalinity is at the bottom.

In the tropics and southernmost part of thesubtropical gyres, salinity is often slightly lowerat the sea surface than in the main part of thesubtropics. Salinity increases to a sharp subsur-face maximum at depths of 100e200 m, close tothe top of the thermocline. This maximum arisesfrom the high salinity surface water in eachsubtropical gyre (Figures 4.7, 4.11b, 4.12b,4.13b, and 4.15). This high salinity watersubducts and flows equatorward and down-ward beneath the fresher, warmer tropicalsurface water, thus forming a salinity maximumlayer. This shallow salinity maximum is foundin the equatorward part of every subtropicalgyre, merging into the tropics. Because it has an


identifiable characteristic (salinity maximum)and common formation history (subductionfrom the high salinity surface water at mid-latitudes), it has acquired status as a watermass. Several names are used for this watermass. Our preference is Subtropical Underwater,following Worthington (1976). It is also referredto as “salinity maximum water.”

Low salinity layers also result from subduc-tion, in this case from the fresher but densernorthern outcrops of the subtropical gyres.Advection of these waters southward resultsin subduction and a low salinity layer that isfound around the eastern and into the southernside of the anticyclonic gyre. In the North andSouth Pacific, these are extensive features calledthe Shallow Salinity Minimum in each ocean(Reid, 1973). In the subpolar North Atlantic,there is a much less-extensive shallow salinityminimum associated with the subarctic front(part of the North Atlantic Current); it is calledSubarctic Intermediate Water.

In subpolar and high-latitude regions, withhigh precipitation, runoff, and seasonal icemelt, there is generally low salinity at thesea surface. The halocline, with a rapid down-ward increase of salinity, lies between thesurface low-salinity layer and the deeper,

saltier water. In such regions, the pycnoclineis often determined by the salinity distribu-tion rather than by temperature, whichremains relatively cold throughout the year,and may have only a weak thermocline oreven none at all. This condition, associatedwith runoff and precipitation, occurs throu-ghout the subpolar North Pacific. In the Arcticand Antarctic and other regions of sea iceformation, ice melt in spring creates a similarlyfreshened surface layer.

This low salinity surface layer in regionslike the subpolar North Pacific and aroundAntarctica permits a vertical temperatureminimum near the sea surface, with a warmerlayer below (the dichothermal and mesothermallayers, described in Section 4.2).

4.3.3. Intermediate Depth Salinity

At intermediate depths (around1000e1500m)in many regions of the world, there are hori-zontally extensive, vertically broad layers ofeither low salinity or high salinity. These areeasily identified in Figures 4.11, 4.12, and 4.13because of their vertical salinity extrema. In theNorth Pacific and Southern Hemisphere, thesalinity minimum layer is at about a depth of

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High latitude(subpolar N. Pac.)

(a) (b) (c)FIGURE 4.16 Typicalsalinity (psu) profiles forthe tropical, subtropical,and subpolar regions ofthe North Pacific. Corre-sponding temperatureprofiles are shown inFigure 4.2.


1000 m. The subpolar North Atlantic salinityminimum is at about a depth of 1500 m. Thelow salinity layers are located near the base ofthe pycnocline, with temperatures of 3e6�C.The two major intermediate-depth salinitymaximum layers are in the North Atlanticand northern Indian Ocean (not to be confusedwith the deeper salinity maximum associatedwith North Atlantic Deep Water; NADW). Theyare considerably warmer than the low salinityintermediate waters. The vertical salinityextrema reflect specific formation processes,described briefly here and in more detail in laterchapters. These layers are therefore labeled aswater masses and called the “intermediatewaters.”

A map of the locations of the major interme-diate water masses is provided in Chapter 14(Figure 14.13). Their low salinity and theirtemperature ranges indicate that they originateat the sea surface at subpolar latitudeswhere surface waters are relatively fresh, butwhere surface waters are warmer than freezing.TheNorth Pacific Intermediate Water (NPIW) orig-inates in the northwest Pacific and is foundthroughout the North Pacific. Labrador Sea Water(LSW) originates in the northwest Atlantic andis found through the North Atlantic. LSW isalso marked by high oxygen and chlorofluoro-carbons, and retains these signatures even as itloses its salinity minimum as it becomes partof the NADW in the tropical and South Atlantic.Antarctic Intermediate Water (AAIW) originatesin the Southern Ocean near South Americaand is found throughout the Southern Hemi-sphere and tropics. In these three ventilationregions, surface salinity is lower but density ishigher than the upper ocean and thermoclinewaters in the subtropics and tropics. Theventilated intermediate waters spread equator-ward and carry their low salinity signature withthem.

The two major salinity maximum interme-diate waters result from high salinity outflowsfrom the Mediterranean and Red Seas. The

source of these high salinity waters is surfaceinflow into these seas; high evaporation withinthe seas increases the salinity and coolingreduces their temperature, thus dense water isformed. When these saline, dense waters flowback into the open ocean, they are dense enoughto sink to mid-depths.

Other, more local, intermediate waters arealso identified by vertical salinity extrema. Forinstance, in the tropical Indian Ocean, a mid-depth salinity minimum originates from fresherPacific Ocean water that flows through the Indo-nesian Passages (Chapter 11). This intermediatesalinity minimum has been called IndonesianIntermediate Water or Banda Sea IntermediateWater (Rochford, 1961; Emery & Meincke,1986; Talley & Sprintall, 2005).

Each of these intermediate waters is dis-cussed in greater detail in the relevant oceanbasin chapter (9e13).

4.3.4. Deep-Water Salinity

The deep waters of the oceans exhibitsalinity variations that mark their origin. TheNorth Atlantic is the saltiest of all of the oceansat the sea surface, so dense waters formed inthe North Atlantic carry a signature of highsalinity as they move southward into theSouthern Hemisphere and then eastward andnorthward into the Indian and Pacific Oceans.This overall water mass is referred to as NorthAtlantic Deep Water. Dense waters formed inthe Antarctic are colder and denser than NorthAtlantic dense waters, so they are foundbeneath waters of North Atlantic origin. Thedense Antarctic waters are also fresher thanNorth Atlantic waters; their progress north-ward into the Atlantic can be tracked throughtheir lower salinity, where they are referred toas Antarctic Bottom Water (AABW). The verticaljuxtaposition of the salty NADW and fresherAABW is apparent in the Atlantic verticalsalinity section (Figure 4.11b). This NADW/AABW structure is also apparent in the


southern Indian Ocean since both NADW andAABW enter the Indian Ocean from the south(Figure 4.13b).

The northern Indian Ocean is tropical so nodense waters are formed there, but the highsalinity from the intermediate waters of theRed Sea penetrates and mixes quite deep,making northern Indian Ocean deep watersrelatively saline (Figure 4.13b). The NorthPacific does not form dense, abyssal watersbecause the sea surface in the subpolar NorthPacific is too fresh to allow formation of watersas dense as those from the Antarctic and NorthAtlantic. Therefore, the salinity structure in thedeep North Pacific is determined by the inflowof the mixture of Antarctic and North Atlanticdeep waters from the south; this mixture ismore saline than the local North Pacific watersso salinity increases monotonically to thebottom in the North Pacific (Figure 4.12b).

A global map of bottom salinity is shown inChapter 14 (Figure 14.14c). Globally, the salinityvariation in the deep waters is relatively small,with a range from 34.65 to 35.0 psu. Like bottomtemperatures, the bottom salinities reflect theAntarctic and Nordic Seas origins of the waters.The Antarctic bottom waters are freshest, withsalinities lower than 34.7 psu. The bottomwaters of Nordic Sea origin are the saltiest,with salinity up to 35.0 psu. Full interpretationof the bottom salinity map also requires consid-eration of the varying bottom depthd as ridgescut up into overlying deep waters d and ofdownward diffusion of properties from theoverlying deep waters, which are beyond thescope of this book.

Thus both the deep water temperature anddeep water salinity have small ranges. Thedeep water environment is relatively uniformin character compared with the upper oceanand thermocline and even the intermediatelayer. This relative uniformity is the result ofthe small number of distinct sources of densewaters, and the great distance and time thatthese waters travel, subjected to mixing with

each other and to downward diffusion fromlayers above them.

4.3.5. Temporal Variations of Salinity

Salinity variations at all timescales are lesswell documented than temperature variations,because temperature is more easily measured.Annual variations of surface salinity in theopen ocean are less than 0.5 psu. In regionsof marked annual variation in precipitationand runoff, such as the eastern North Pacificand the Bay of Bengal and near sea ice, thereare large seasonal salinity variations. Thesevariations are confined to the surface layersbecause in such regions the effect of reducedsalinity overrides the effect of temperature inreducing the seawater density. This keeps thelow salinity water in the surface layer. Diurnalvariations of salinity appear to be small, butagain this is a conclusion based on very fewobservations. Local rainstorms produce freshsurface waters even in the open ocean thatmix into the surrounding waters after severalweeks.

Temporal salinity variations at a given loca-tion can be large at large-scale fronts betweenwaters of different properties. These frontsare sometimes termed water mass boundaries.Temperature variations can also be quite largeacross these fronts. The fronts move about theirmean location, on weekly, seasonal, and longertimescales. Meandering of the fronts and crea-tion of eddies of the different types of waterscan cause large salinity and temperature varia-tions at a given location.

Interannual and long-term changes in large-scale salinity are observed and are part of thedocumentation of climate change. With theadvent of the global profiling float array, it isbecoming possible to document salinitychanges in all regions of the non-ice-coveredocean; significant patterns of surface salinitychange have already been detected (Hosoda,Suga, Shikama, & Mizuno, 2009; Durack &


Wijffels, 2010). Salinity variations in the NorthAtlantic and Nordic Seas are associated withchanges in mixed layer convection and withchanges in water mass formation in the Labra-dor and Greenland Seas (Chapters 9 and 12).LSW has dramatic decadal salinity variationsthat correspond with changes in its formation.See Figure S15.4 (Yashayaev, 2007) from theonline supplementary material. Decades longfreshening of the subpolar North Atlanticand Nordic Seas (see Chapter S15 from theonline supplementary material), followed bysalinification in recent years, has causedmuch interest in terms of NADW productionrates. Large-scale, coherent salinity changesover several decades have been documented(Boyer, Antonov, Levitus, & Locarnini, 2005;Durack & Wijffels, 2010) and can be associatedwith large-scale changes in precipitation andevaporation that might be related to overallwarming of the atmosphere (Bindoff et al.,2007).

4.3.6. Volumetric Distributionof Potential Temperature and Salinity

A classic (and typical) approach to looking atthe water mass structure is to display variousproperties as a function of each other; a moremodern statistical approach describes the watermasses in terms of all of their properties (seeSection 6.7). Potential temperature-salinity dia-grams are used throughout the basin chapters(9e13) to illustrate the water masses. A volu-metric q-S diagram from Worthington (1981) isintroduced as our first global summary ofwater properties (Figure 4.17). The method isdescribed in Section 6.7.2.

The underlying q-S in the upper panel(Figure 4.17a) shows three separate branchesstretching from low q-S to higher q-S; these arethe Central Waters of the pycnocline (as inFigure 4.7). The saltiest branch is the NorthAtlantic; the freshest branch is the North Pacific.The intermediate branch, with larger volumes,

is the three Southern Hemisphere basins (SouthAtlantic, South Pacific, and Indian). The impor-tance of the Southern Ocean connection bet-ween these latter three basins is immediatelyapparent, as the three have properties that aremore similar than the two Northern Hemi-sphere basins.

In the deep water (Figure 4.17b), the largestpeak is the Pacific Deep Water (or CommonWater); the large volume in a single q-S classindicates how well mixed this water mass is,which is a direct result of its great age (Section4.7). The coldest waters are the AABW, withthe single ridge again indicating SouthernOcean circumpolar connectivity. Above about0�C, the diagram splits into three branches dthe Pacific Ocean, Indian/Southern Ocean, andAtlantic Ocean, from freshest to saltiest. Thesalty Atlantic ridge has a long portion of highvolume, without a huge, single peak such as isfound in the Pacific. This reflects the multiplesources of NADW and its relatively young,unmixed character.

4.4. DENSITY DISTRIBUTION

Potential density must increase with depth ina system in equilibrium. To be more precise, thewater column must be statically stable, using thedefinition of static stability (Eq. 3.9) in Section3.5.6. This means that potential density, usinga local reference pressure, must increase withdepth. While potential temperature and salinitytogether determine density, individually they canhave maxima and minima in the vertical, as longas the water density increases with depth. Theonly exceptions to the monotonic increase occurat very short timescales, on the order of hours orless, which is the timescale for overturn. As soonas denserwater flows over lighterwater, or surfacelayer density is increased above that of the under-lying water, the water column becomes unstableand will overturn and mix, removing theinstability.


In Chapter 3, we discussed the use of differentreference pressures for reporting potentialdensity, or equivalently for use of an empiricallydefined type of density such as neutral density(Section 3.5.4). The potential density that is usedshould best approximate the local verticalstability and isentropic surfaces. Profiles of poten-tial density relative to both the sea surface and4000 dbar are used in constructing the potential

density sections of Figures 4.11, 4.12, and 4.13.When spatial variability in temperature andsalinity is very small, any typeofpotentialdensitywill increase monotonically with depth; anexample is the potential density relative to boththe sea surface and 10,000 dbar in the MarianaTrench (Figure 4.10). The North Pacific has littlevariation in temperature and salinity below thepycnocline, which is the vertical region of large

FIGURE 4.17 Potential tempera-ture-salinity-volume (q-S-V) dia-grams for (a) the whole watercolumn and (b) for waters colderthan 4�C. The shaded region in (a)corresponds to the figure in (b).Source: From Worthington (1981).

DENSITY DISTRIBUTION 93

density change. Therefore all potential densitychoices yield stable-appearing vertical profiles(Figure 4.20).

In the South Atlantic, on the other hand, thereare large-scale salinity inversions where thesaline NADW is layered between the fresherAAIW and AABW (Figure 4.11b). Here thedifferences in compressibility of warmer andcooler waters begin to matter. Figure 4.11cemphasizes the local potential density structure,which is decidedly stable in the vertical. To illus-trate the main drawback of using a surface refer-ence pressure for deep-water density, a verticalsection through the Atlantic Ocean of potentialdensity relative to the sea surface, sq, for thefull water column is shown in Figure 4.18a.There is a large-scale inversion of sq in the SouthAtlantic, most pronounced just south of theequator at a depth of about 3700 m. This is thebase of the high salinity NADW layer. Potentialtemperature contours are compressed below theNADW (Figure 4.11a). Potential density refer-enced to 4000 dbar, s4, hence locally referenced,has no inversion (Figure 4.11c).

Neutral density gN (Section 3.5.4; Jackett &McDougall, 1997) is commonly used to repre-sent the stable increase of “potential” densitywith depth.1 Like choosing appropriate locallyreferenced potential densities, neutral densityeliminates the apparent density inversions ofFigure 4.18a and also removes the need to usemultiple pressure reference levels such as inthe use of 0 and 4000 dbar references in Figures4.11c, 4.12c, and 4.13c. The neutral density gN

section for the Atlantic is clearly monotonic,with gN increasing from top to bottom. Thedeep contours resemble those of s4 (Figure 4.11c),and the distortions of sq in the region of theMediterranean salinity maxima at about 2000m in the North Atlantic are removed.

4.4.1. Density at the Sea Surface and inthe Upper Layer

The density of seawater at the ocean surfaceincreases from about sq ¼ 22 kg/m3 near theequator tosq¼ 26e 28 kg/m3 at 50e60� latitude,and beyond this it decreases slightly (Figures 4.3

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(a) (b)

Atlantic σθ

Atlantic γΝ

km

FIGURE 4.18 (a) Potential density sq (kg m�3) and (b) neutral density gN in the Atlantic Ocean at longitude 20� to 25�W.Compare with Figure 4.12c. Data from the World Ocean Circulation Experiment.

1 There continues to be energetic discussion of the most appropriate variable for density for constructing the most isentropic

surfaces in the sense of the direction of motion of water parcels and the directions of along-isopycnal and diapycnal

mixing.


and 4.19), due to lower salinity at higher lati-tudes. Surface densities at high latitudes in theAntarctic and North Atlantic are higher than inthe North Pacific even at the freezing point.North Pacific surface water must be less densesince its surface water is fresher.

In Figure 4.3, we see that the surface densityaveraged for all oceans follows surface temper-ature rather than surface salinity in the tropicsand mid-latitudes. At the highest northern andsouthern latitudes, poleward of 50�, surfacedensity follows salinity more than temperature,because temperature is close to the freezingpoint there, with little variation in latitude.

Surface density and the vertical stratificationdetermine the depth to which surface waterswill sink as they move away from their ventila-tion (“outcrop”) region. The combination ofsurface temperature and salinity for a givendensity also affects the sinking because of theireffect on compressibility, with warmer, saltier

water compressing less than colder, fresherwater at the same surface density. Thus thecolder, fresher parcel will become more denseand, consequently, deeper than the warmer,saltier parcel as they move into the ocean eventhough they start with the same surface density.See Section 3.5.4.

In late winter, surface waters reach their localdensity maximum as the cooling season draws toa close. (Cooling in many regions is also associ-ated with evaporation, so both temperature andsalinity may change together to create densewater, depending on the local amount of precip-itation.) Late winter density is associated withthe deepest mixed layers. As thewarming seasonbegins (March in the Northern Hemisphere,September in the Southern Hemisphere), thedensewintermixed layer is “capped” bywarmerwater at the surface. The capped winter watersmove (advect) away from the winter ventilationregion. If they move into a region where the

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Winter surface density

FIGURE 4.19 Surface density sq (kg m�3) in winter (January, February, and March north of the equator; July, August, andSeptember south of the equator) based on averaged (climatological) data from Levitus and Boyer (1994) and Levitus et al.(1994b).


winter surface waters are less dense, they sinkbeneath the local surface layers, and will not bereopened to the atmosphere during the nextwinter. This subduction process is a primarymechanism for moving surface waters into theocean interior (Luyten et al., 1983; Woods, 1985and Sections 4.2.3 and 7.8.5).

Longer timescale variations in surfacedensity can affect the amounts of intermediateand deep waters that form and the overall sizeof the regions impacted by them. During majorclimate changes associated with glacial/inter-glacial periods, surface density distributionsmust have been strongly altered, resulting invery different deep water distributions.

Winter mixed layer depths vary from tens ofmeters to hundreds of meters, depending onthe region (Figure 4.4). Because they haveusually been detected using temperaturecriteria, we discussed mixed layers in somedetail in Section 4.2.2. In the tropics, wintermixed layer depths may be less than 50 m.Winter mixed layer depths are greatest in thesubpolar North Atlantic, reaching more than

1000 m in the Labrador Sea, and in the SouthernHemisphere around the northern edge of themajor current that circles Antarctica at a latitudeof about 50�S, reaching up to about 500 mthickness.

4.4.2. Pycnocline

Like potential temperature, potential densitydoes not increase uniformly with depth(Figure 4.20). The vertical structure of densityis similar to that of potential temperature. Thereis usually a shallow upper layer of nearlyuniform density, then a layer where the densityincreases rapidly with depth, called the pycno-cline, analogous to the thermocline (Section4.2.3). Below this is the deep zone where thedensity increases more slowly with depth(Figures 4.10, 4.11, 4.12, 4.18, and 4.20). Thereis much smaller variation with latitude of thedeep-water density compared with upper oceandensity. As a consequence, in high latitudes,where the surface density rises to sq ¼ 27 kg/m3

or more, there is a smaller increase of density

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Potential density sq -1000 Potential density s4 -1000

FIGURE 4.20 Typical density/depth profiles for low and highlatitudes (North Pacific).


with depth than in the low latitudes, and thepycnocline is much weaker.

The double thermocline structure that occursin some broad regions (described in Section 4.2),is mirrored in density because of the strongdependence of ocean density on temperature.Layers of lower vertical density gradients arecalled pycnostads.

In all regions, there is a seasonal pycnoclinein the warm seasons. This results from seasonalwarming and/or ice melt, overlying theremnant of the winter mixed layer, which formsa pycnostad in non-winter seasons. A perma-nent double pycnocline, with a pycnostad lyingbetween the pycnoclines, is a common feature ofsubtropical regions. Mode Waters (Section 4.2.3)are pycnostads, and are best identified in termsof density stratification rather than temperaturestratification; that is, a minimum in verticalstability is the best identifier of a Mode Wateron a given vertical profile. Often Mode Watersand other water masses are tracked in terms oftheir potential vorticity (Eq. 3.11 and Section 7.7);the dominant contribution to potential vorticityin most of the ocean (except in strong currents)is proportional to the vertical stability. Potentialvorticity is a useful tracer because it is aconserved dynamical quantity in the absenceof mixing.

4.4.3. Depth Distribution of PotentialDensity

Potential density structure is simpler thanpotential temperature and salinity simplybecause the water column must be verticallystable. Potential density, appropriately defined,must increase downward. Below the pycno-cline, vertical potential density variations aremuch smaller, similar to potential temperaturestructure. There are no large-scale inversions indensity if the appropriate reference pressuresare used, as described in Section 3.6 and asseen in the vertical section through the AtlanticOcean (Figure 4.18 compared with Figure 4.11c

in Section 4.2.6). Horizontal variations inpotential density are associated with horizontalpressure gradients and therefore with large-scale currents (Section 7.6).

Potential density structure is displayedalong vertical sections through the north-southlength of each ocean (Figures 4.11, 4.12, and4.13). The main features are downward bowlsin the upper to intermediate ocean in thesubtropics, and a large upward slope towardthe southern (Antarctic) end of the sections.Below about 2000 m, the total range of potentialdensity is small, from about sq ¼ 27.6 to 27.9kg/m3 (or s4 ¼ 45.6 to 46.2 kg/m3, which ispotential density relative to 4000 dbar).

Because mixing is greatly inhibited by ver-tical stratification, there is a strong preferencefor flow and mixing to occur along isentropicsurfaces, which are approximately isopycnals(surfaces of constant potential density). In theupper ocean, surfaces of constant sq are useful.For instance, the processes that give oceanwaterstheir particular properties act almost exclusivelyat the surface, and the origin of even the deepestwatercanbe tracedbacktoaregionofformationatthe surface somewhere. Because deep oceanwater is of high density, it must form at high lati-tudes where cold, dense water is found at thesurface. After formation, it spreads down almostisopycnally (reference pressures should beadjusted to account for temperature dependenceof the compressibility). The sinking is combinedwith horizontal motion so that the water actuallymoves in a direction only slightly inclined to thehorizontal. Even in the regions of largest isopyc-nal slopes, for instance in the southern part ofFigures 4.11c, 4.12c, and 4.13c, the slopes are nomore than several kilometers down over severalhundred kilometers horizontally.

Even though there is large-scale structure inthe deep ocean’s salinity field (e.g., in theAtlantic in Figure 4.11b), temperature domi-nates the density structure in the deep ocean.Salinity is important for the density structurenear the sea surface at high latitudes where


precipitation or ice melt creates a low salinitysurface layer, such as in the Arctic, in the regionnext to Antarctica, and in the subpolar NorthPacific and coastal subpolar North Atlantic. Inshallow coastal waters, fjords, and estuaries,salinity is often the controlling factor in deter-mining density at all depths, while the tempera-ture variations are of secondary importance(e.g., Table 3.1).

In much of the ocean, the density profile withincreasing depth appears nearly exponential,asymptoting to a nearly constant value in thedeep ocean (Figure 4.20). However, in someregions, where deep waters from very differentsources are juxtaposed, there is a weak pycno-cline (higher density variation) between them.An obvious example is between the NADWand AABW in the South Atlantic, which iswhere we illustrated the necessity for local refer-encing of the potential density. Such regions aremost common in the subtropical SouthernHemisphere where dense Antarctic watersflow northward in a thick layer under slightlyless dense deep waters flowing southwardfrom the North Atlantic, Pacific, and IndianOceans (Figures 4.11 and 4.18).

4.5. DISSOLVED OXYGEN

Seawater contains dissolved gases, includingoxygen and carbon dioxide. Some of the tran-sient tracers are dissolved gases (Section 4.6).The ocean is an important part of the global(atmosphere/land/ocean) cycle of carbondioxide, which is a greenhouse gas. However,because of its complexity, we do not describethe ocean’s carbon chemistry in this book, andinstead refer readers to texts on biogeochemicalcycles (e.g., Broecker & Peng, 1982; Libes, 2009).

Dissolved oxygen content is used as animportant tracer of ocean circulation and anindicator of the time passed since a water parcelwas at the sea surface (ventilated) (Section 3.6).The range of oxygen values found in the sea is

from 0 to 350 mmol/kg (0 to 8 ml/L), but a largeproportion of values fall within the more limitedrange from 40 to 260 mmol/kg (0 to 6 ml/L). Theatmosphere is the main source of oxygen dis-solved in seawater. At the sea surface, the wateris usually close to saturation. Sometimes, in theupper 10e20 m, the water is supersaturatedwith oxygen, a by-product of photosynthesisby marine plants. Supersaturation also occursnear the sea surface if the water warms upthrough solar radiation that penetrates tens ofmeters into the ocean. (If the actual sea surfacebecomes warmer, it will lose its excess oxygento the atmosphere, so supersaturation is notfound right at the sea surface.) Sometimessurface waters are undersaturated; this occursrarely in winter if mixing of the surface layeris especially intense, entraining underlyingolder, less saturated waters. (The equilibrationtime of surface waters d time required torestore the waters to 100% saturation d isa few days to a few weeks and is a function ofwind speed and temperature.) Below thesurface layers, the oxygen saturation is lessthan 100% because oxygen is consumed byliving organisms and by the bacterial oxidationof detritus. Low values of dissolved oxygen inthe sea are often taken to indicate that the waterhas been away from the surface for a long time,the oxygen having been depleted by the biolog-ical and detrital chemical demands.

Figure 4.21 shows typical dissolved oxygenprofiles for the Atlantic and the Pacific for threelatitude zones. Figures 4.11d, 4.12d, and 4.13dshow oxygen along a south-north section foreach ocean. Common features of the AtlanticandPacificare (1)highoxygenclose to the surface,(2) an oxygen minimum at 500e2500 m, (3) rela-tively high values below 1500 m in the Atlantic(NADW), (4) low values in the North Pacificbeneath the surface layer, and (5) more similarsubsurface distributions in the southern lati-tudes in both oceans. Distributions in the IndianOcean are similar to those in the Pacific (southand tropics). The lower values in the deep water


of thePacific comparedwith theAtlantic indicatethat the Pacific water has been away from thesurface for a much longer time. In some regionsof extremely low oxygen, such as the Black Seaand the bottom of the Cariaco Trench (off Vene-zuela in the Caribbean), hydrogen sulfide ispresent, created from the reduction of sulfateion by bacteria. This indicates that the waterhas been stagnant there for a long time.

The oxygen minimum through the worldoceans at mid-depth, overlying higher oxygenat the bottom, results from at least several mech-anisms. One is that minimal circulation andmix-ing do not replace the oxygen consumed. Asecond is that the increase of density with depth(stability) allows biological detritus to accumu-late in this region, which increases the oxidationrate. A third is that the bottom waters in each ofthe oceans are relatively high in oxygen becauseof their surface source in the Antarctic. In thePacific and Indian Oceans, a three-layer struc-ture is obtained, with high oxygen at the surfacedecreasing through the pycnocline, an oxygenminimum layer in the intermediate and deepwater, and higher oxygen in the abyss. TheAtlantic has a four-layer structure because ofthe juxtaposition of high oxygen content in the

NADWonto this three-layer structure (the thicklayer of higher oxygen between 2000 and 4000dbar in Figure 4.11d corresponds with highsalinity in Figure 4.11b).

A pronounced vertical oxygen minimum isfound in the upper ocean in the tropical Atlantic(Figure 4.11d), eastern tropical Pacific(Figure 4.12d), and in the northern Indian Ocean(Figure 4.13d). These shallow oxygen minimaresult from very high biological productivityin the surface waters in these regions. Bacterialconsumption of the large amount of sinkingdetritus from these surface waters is large andconsumes almost all of the dissolved oxygenwithin the upper 300e400 m of the ocean.

The production and utilization of oxygen inthe sea are essentially biogeochemical matters(Section 3.6). Oxygen is a useful tracer, broadlyindicative of a water parcel’s age, but sinceit is non-conservative, it must be used carefully.

4.6. NUTRIENTS AND OTHERTRACERS

Other common water properties used as flowtracers or for identifying water masses include

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FIGURE 4.21 Profiles of dissolvedoxygen (mmol/kg) from the Atlantic(gray) and Pacific (black) Oceans.(a) 45�S, (b) 10�N, (c) 47�N. Datafrom the World Ocean CirculationExperiment.

NUTRIENTS AND OTHER TRACERS 99

the nutrients (phosphate, nitrate, nitrite, silicicacid, and ammonium), dissolved gases otherthan oxygen and carbon dioxide, and plankton,which are small organisms (both plant andanimal) that drift with the water. These watermass characteristics must be used with cautionbecause, like oxygen, they are non-conservative;in other words, they may be produced orconsumed within the water column. Otherchemical and radioactive tracers are nowmeasured widely (e.g. Broecker and Peng,1982). Here we concentrate on the main featuresof the nutrient distributions, which are ofinterest to marine biologists as well as physicaland chemical oceanographers.

Nutrient values are low in the upper fewhundredmeters with higher values in the deeperwater (Figure 4.22). In the North Pacific, thesedeeper distributions are in the form ofmid-depthmaxima, extending from north to south, withhighest values at 1000/2000 m for nitrate (NO3)and phosphate (PO4), and at 2000/3000 m forsilicate. Additionally, there are maxima in thesouth in the dense water formed in the Antarctic.In the Atlantic, the mid-depth low nutrienttongues extending from north to south are asso-ciated with the NADW (Section 9.8). Maximumvalues are found in the south and along thebottom in the dense waters formed in theAntarctic (AABW).

The low values of nutrients in the upperlayers result from utilization by phytoplanktonin the surface layer (euphotic zone, exposed tosunlight), while the increase in deeper watersis because of their release back to solution bybiological processes (respiration and nitrifica-tion, mostly microbial) during the decay ofdetrital material sinking from the upper layers.Therefore, nutrient distributions are approxi-mately mirror images of the oxygen distribu-tion. Phosphate and nitrate have similardistributions because they are governed byalmost the same biological cycle (see discus-sion of Redfield ratios in Section 3.6). (There-fore only nitrate sections are included in

Figure 4.22.) The dissolved silica (silicic acids)distribution is not as closely similar. Silica hasan additional source at the ocean bottom, asit can be dissolved into the seawater fromthe sediments or injected by hydrothermalsources.

Nutrient replenishment in the surface layersis strongly influenced by the physical processesof vertical diffusion, overturning, and upwell-ing. These bring nutrients from below theeuphotic zone up to the surface. The impact ofupwelling on surface nutrients is illustrated bythe nitrate distribution at the sea surface(Figure 4.23). Nitrate is nearly zero in thesubtropical regions where surface waters down-well (Chapters 9e11). Surface nitrate is non-zero(although small) where there is upwelling fromjust below the euphotic zone, which occurs insubtropical eastern boundary regions (Section7.9.1), along the equator, and in the subpolarregions. These are regions of high biologicalproductivity because of the nutrient supply tothe euphotic zone (see map of depth of theeuphotic zone in Figure 4.29).

At mid-depth, in the nutrient vertical maxi-ma, the Pacific nutrients are higher than Atlanticvalues by a factor of about two for phosphateand nitrate, and by a factor of three to ten for sili-cate. This is due to the much greater age of themid-depth and deep waters in the Pacific thanin the North Atlantic. The lower dissolvedoxygen values in the Pacific than in the Atlanticare attributed to the same cause.

Taken together, the oxygen and nutrientdistributions, along with salinity, provide ourprincipal identification for water masses belowthe pycnocline. The high oxygen, low nutrient,high salinity deep layer in the Atlantic Oceanis the NADW. The low oxygen, high nutrientlayer in the Pacific Ocean is the Pacific DeepWater, and the same layer in the Indian Oceanis the Indian Deep Water. The higher oxygen,lower nutrient, very cold bottom layer in alloceans is the AABW. When considering morecarefully the east-west distributions of


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FIGURE 4.22 Nitrate (mmol/kg) and dissolved silica (mmol/kg) for the Atlantic Ocean (a, b), the Pacific Ocean (c, d), andthe Indian Ocean (e, f). Note that the horizontal axes for each ocean differ. Data from the World Ocean CirculationExperiment. This figure can also be found in the color insert.

NUTRIENTS AND OTHER TRACERS 101

properties in the Southern Hemispheresubtropics, we can also distinguish betweenhigher oxygen, lower nutrient waters thatcome from the Antarctic compared with thelow oxygen, high nutrient Pacific and IndianDeep Waters. These Antarctic deep waters arenot as dense as AABW and are referred to asCircumpolar Deep Water. Different types ofCircumpolar Deep Water are described inChapter 13.

4.7. AGE, TURNOVER TIME, ANDVENTILATION RATE

Estimates of age and overturning rates ofocean waters assist in understanding the overalldistribution of temperature and salinity in theocean, the replenishment rate of nutrients inthe upper layers, and the exchange of gasesbetween the atmosphere and ocean. The

effectiveness and safety of the deep ocean asa dump for noxious materials depends on theturnover time of the deep waters. For theseand many other applications, it is useful to esti-mate timescales for ocean ventilation.

Age as applied to ocean water is the timesince a parcel of water was last at the seasurface, in contact with the atmosphere. Theventilation rate or production rate is the transportof water that leaves a surface formation siteand moves into the ocean interior. Turnovertime is the amount of time it takes to replenisha reservoir, such as an ocean basin or a layeror water mass in the ocean. The “reservoir”can also be construed in terms of a tracer ratherthan water particles (e.g., molecules of CO2, orzooplankton, etc.). Residence time is the timea particle spends in a reservoir.

The ages of waters can be estimated usingtracers (Section 3.6). Tracers that are biologicallyinert are more straightforward than those that

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FIGURE 4.23 Nitrate (mmol/L) at the sea surface, from the climatological data set of Conkright, Levitus, and Boyer(1994).


are biologically active. Anthropogenic transienttracers that have measured histories in the atmo-sphere are useful in the upper ocean and well-ventilated parts of the deep ocean. The penetra-tion of chlorofluorocarbons (CFCs; Pacific sectionshown in Figure 4.24a) and tritium (Pacific mapin Figure 4.25b) is evidence of recent ventilation;absence of these tracers is clear evidence of agethat is greater than 50e60 years.

Pairs of tracers whose concentration ratiochanges with time can be used to estimate age,including pairs of CFCs with different atmo-spheric time histories that result in a changingratio in surface sourcewaters (Figure 4.25a). Simi-larly, because tritium (3H) decays to 3He witha half-life of about 12 years, the 3H/3He pairreflects age (ignoring the smaller amounts ofnatural 3He injected in the deep ocean at themid-ocean ridges). This tracer ratio method isstraightforward only if the surrounding watersare free of the tracers because mixing betweenwaters with different ratios complicates the agecalculation. The tropical Pacific andNorth Pacificare ventilated only in the upper ocean, with nodeep water sources except far to the south, sothe CFC and tritium/3He pair “ages” are espe-cially useful for estimating water age there.

For the deep ocean, where water is too old tobe dated using anthropogenic tracers, and alsoas an alternate method of estimating age in thebetter ventilated parts of the ocean, naturaltracers such as oxygen, nutrients, and 14C(Figure 4.24b) are useful. Biological activityreduces oxygen and increases nutrient contentonce the water moves away from the seasurface. If the oxygen consumption rate ornutrient remineralization rates are known asa function of location and temperature, thenthe age of a water mass as it moves away fromits source can be estimated. As with anthropo-genic tracers, simplifying assumptions aboutmixing with waters of different oxygen andnutrient content are required.

Radiocarbon can be used for dating justas with terrestrial organic matter (Section 3.6).

14C is created in the atmosphere by cosmicrays and quickly becomes part of the atmo-spheric CO2. It enters the ocean with the CO2

that dissolves in surface water. When thesurface water is subducted or incorporated indeeper waters, the decay of 14C at a rate of 1%every 83 years results in increasingly negativevalues (deficits) along the pathways into thedeep ocean. The largest deficits globally arefound in the deep North Pacific, reflecting thegreat age of the waters there (Figure 4.24b).Use of 14C deficits to precisely date ocean wateris subject to caveats about mixing and alsocomplications due to local sinking of organicmatter from the surface and other sources of 14Cincluding nuclear testing. The gross estimateof ages of deep waters based directly on 14Cdeficits is 275 years for the Atlantic, 250 yearsfor the Indian, and 510 years for the Pacific. Itis easy to see these age estimates are biased sincethe oldest waters always mix with youngerwaters and vice versa. Thus the age of thedeep northern Pacific waters is likely higherthan their 14C age (around 1000 years), whilethe age of the deep northern Atlantic waters islower, as evidenced by invasion of CFCs to thebottom (Broecker et al., 2004).

The ventilation rate (production rate) of a watermass or layer can be defined in several differentways, which can lead to somewhat differentquantitative estimates. In all cases, the objectiveis to estimate the rate of injection of new waterinto a reservoir. One approach is to estimateproduction rate from the volume of the reservoirdivided by its age:

RP ¼ Volume=Age (4.1a)

which has units of transport (m3/sec). This isa straightforward concept, but difficult to imple-ment since the ocean is not composed of simpleboxes filled with waters of uniform age; there-fore somewhat complex calculations and simplemodels are used to obtain ventilation rates fromthe continuous distribution of ages. If Eq. (4.1a)

AGE, TURNOVER TIME, AND VENTILATION RATE 103

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FIGURE 4.24 (a) Chlorofluorocarbon content (CFC-11; pmol/kg) and (b) D14C (/mille) in the Pacific Ocean at 150�W. White areas in (a) indicateundetectable CFC-11. From the WOCE Pacific Ocean Atlas. Source: From Talley (2007).

4.TYPIC

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is written in terms of turnover time (Eq. 4.2below) instead of age, the rate that is obtainedcould differ since age and turnover time areusually not identical.

A related approach to estimating a ventilationrate using transient tracers starts with the totalamount (inventory) of the tracer and concentra-tion of the tracer at its source (sea surface). Forinstance, using CFC’s with an inventory ICFC(units of moles) and surface concentrationCsource (units of moles/kg), the ventilation rateis given by (Smethie & Fine, 2001):

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Since both the source concentration and inven-tory vary with time, this ventilation rate isobtained iteratively.

Ventilation rates, RP, are also estimated fromobservations of the transport of newly venti-lated waters very close to the source of the watermass. Farther from the source, quantitativewater mass identification techniques (Section6.7) can be used to estimate the portion ofobserved transport that can be attributed tothe source versus the portion due to mixingwith other waters. Indirect estimates are alsofrequently used, based on measuring the buoy-ancy forcing that results in ventilation withsimple or complex models to compute the venti-lation rate; an approach using airesea fluxes ofheat and freshwater within surface outcroppingregions of isopycnal layers was introduced byWalin (1982) and has been employed in numerouscalculations.

Turnover time is the time it takes to replenisha reservoir. If in reference to water rather thana tracer, it is equal to the volume V of the watermass or layer, in units of m3, divided by itsoutflow transport Rout in m3/sec. If in referenceto a tracer, it is the inventory of the tracer, inmoles, divided by the transport of the tracerout of the reservoir in mole/sec. Turnovertime, which has been defined generally for usein biogeochemistry, is written in terms ofthe exit flow because reservoirs are usuallywell-mixed, unlike the inflow sources, resultingin a simpler (proportional) relation betweenoutflow volume transport and turnover time.

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FIGURE 4.25 (a) Age (years) of Pacific Ocean waters onthe isopycnal surface 27.2 sq, using the ratio of chlorofluo-rocarbon-11 to chlorofluorocarbon-12. Source: From Fine,

Maillet, Sullivan, and Willey (2001). (b) Tritium concentrationat 500 m in the Pacific Ocean from the WOCE Pacific OceanAtlas. Source: From Talley (2007).

AGE, TURNOVER TIME, AND VENTILATION RATE 105

Turnover times for volume and for a tracer aregiven by:

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where vout and vin are velocities out of and into the reservoir, C is the concentration ofa tracer (e.g., in mmol/kg) and r is density.Eq. (4.2a) can also be written in terms ofmass rather than volume transport byincluding r in both the numerator anddenominator. The rightmost term with inflowvelocity in Eq. (4.2a) yields turnover time ifthe system is in steady state. In a steady state,Eq. (4.2b) could also be written in terms of theinflow.

Residence time is the time an individual waterparcel spends in a reservoir. The average resi-dence time is obtained by averaging over all ofthe water parcels passing through the reservoir.The average residence time is equal to the turn-over time if the system is in steady state. Theaverage residence time is twice the age if wateris moving steadily through the reservoir, sincethe age is the average of newest to oldest watersin the reservoir.

4.8. OPTICAL PROPERTIESOF SEAWATER

The transparency of the ocean depends onhow much suspended or living material is con-tained in it, as described in Section 3.8. If thewater is very transparent, then solar radiationpenetrates to greater depth than if there ismuch suspended material. Therefore opticalproperties of the surface waters affect surfacelayer heating, thereby affecting surface temper-ature and hence ocean-atmosphere interaction.Ocean color depends on suspended materials,

especially including chlorophyll-producingphytoplankton, so large-scale observations ofcolor and other optical properties can be usedto study biological productivity. Optical obser-vations of ocean color using satellite remotesensing are used routinely to quantify theamount of chlorophyll-a (green pigment;McClain, Hooker, Feldmand, & Bontempi,2006), and, more recently, particulate organiccarbon (POC; Gardner, Mishonov, & Richardson,2006; Stramski et al., 2008), and euphotic zonedepth (Lee et al., 2007).

Prior to invention of electronic opticaldevices, transparency was measured usinga Secchi disk (see Section S16.8 in the supple-mental materials on the text Web site for infor-mation about this instrument). This was doneby visually observing when the speciallypainted disk could no longer be seen from theship’s deck. An enormous number of Secchidisk depths (>120,000) were collected and arearchived at the U.S. National OceanographicData Center (Lewis, Kuring, & Yentsch, 1988).The majority of the values were for the northernoceans and taken in the summer. There are largeareas of the Southern Hemisphere open oceanwhere there are no values at all, but coastal areaswere generally well sampled. Large Secchidepths are found in the open oceans at lowand middle latitudes with lower values inhigher latitudes and along most coasts. The lat-itudinal variation is apparent in Figure 4.26,which shows averages of Secchi depths along180�� 20�W for the Pacific, and along 35�� 10�Wfor the Atlantic. Lewis et al. (1988) concludedthat the prime source of variability in the openocean is attenuating material in the water. Thesmaller Secchi depths correspond to higherchlorophyll-a values. The most marked featurein Figure 4.26 is the sharp decrease in Secchidepths beyond about 30� latitude, correspond-ing to higher productivity in the higher lati-tudes. The large Secchi depths in the Atlanticare in the Sargasso Sea, a region of notably lowbiological productivity. In a polynya in the


Weddell Sea in 1986, a Secchi disk was visible tofour observers at 79 m and disappeared at 80 m.This was claimed as a record: the Secchi depthcalculated for distilled water is 80m, so a greaterdepth is not possible. In coastal waters, values of2e10 m are common, and in silty waters nearrivers and in estuaries, values of less than 1 mare observed.

Modern in situ optical observations are madewith instruments that measure many differentaspects of seawater optical properties, whichare affected by suspended materials, includingsediments and plankton (Section 3.8; Figure 3.9).Fluorescence provides a measure of chlorophyllconcentration and therefore, phytoplankton.Within the water column, light transmission,beam attenuation, and fluorescence, amongother properties, are measured at differentwavelengths to quantify the amount andtype of suspended material (Gardner, 2009). Asan example, beam attenuation measured witha transmissometer, at a visible wavelength(660 nm), is shown for the equatorial Pacificand the eastern subpolar North Pacific (OceanWeather Station P or Papa; Figure 4.27). This

instrument provides its own light as it is low-ered through the water column, so the observa-tion is related to the local amount of scatteringand absorption by particles and absorption bywater, and not to the actual penetration ofsunlight. This particular beam attenuation canthen be related to the amount of POC, which ismeasured from actual samples of seawater.The high beam attenuation in the uppermostlayer indicates high POC.

Using ocean color remote sensing and in situobservations of chlorophyll-a and POC, algo-rithms have been developed to map the latterquantities. Chlorophyll-a maps are now stan-dard remote sensing products. Seasonal mapsof chlorophyll from remote sensing are shownin Figure 4.28. Notable features of the northernsummer chlorophyll distribution include verylow values throughout the subtropical gyres,high values in the equatorial regions and alongparts of the ACC, very high values in the highnorthern latitudes and Arctic, and high valuesin coastal regions. In austral summer, the highlatitude patterns reverse somewhat, withincreased chlorophyll along the margins of

FIGURE 4.26 Mean Secchi disk depths as functions of latitude in the Pacific and Atlantic Oceans. Source: From Lewis et al.(1988).

OPTICAL PROPERTIES OF SEAWATER 107

FIGURE 4.27 Profile of beam attenuation coefficient at 660 nm, from a transmissometer, converted to POC (solid line)and in situ measurements of POC (circles): (a) equatorial Pacific and (b) northeast Pacific at OWS Papa. Source: From Bishop

(1999).

FIGURE 4.28 Global images of chlorophyll derived from the Coastal Zone Color Scanner (CZCS). Global phytoplanktonconcentrations change seasonally, as revealed by these three-month “climatological” composites for all months betweenNovember 1978eJune 1986 during which the CZCS collected data: JanuaryeMarch (upper left), AprileJune (upper right),JulyeSeptember (lower left), and OctobereDecember (lower right). Note the “blooming” of phytoplankton over the entireNorth Atlantic with the advent of Northern Hemisphere spring, and seasonal increases in equatorial phytoplanktonconcentrations in both Atlantic and Pacific Oceans and off the western coasts of Africa and Peru. Figure 4.28 will also befound in the color insert. See Figure S4.2 from the online supplementary material for maps showing the similarity betweenparticulate organic carbon (POC) and chlorophyll. Source: From NASA (2009a).


Antarctica (now ice free) and reduced chloro-phyll in the high northern latitudes. The POCdistribution derived from ocean color is closelyrelated to the chlorophyll-a distribution (Gard-ner et al., 2006; see Figure S4.2 in the onlinesupplement).

The solar radiation that affects the upperocean is quantified as photosyntheticallyavailable radiation (PAR; Section 3.8.1), and ismapped routinely from ocean color sensors(NASA, 2009b). An example (Figure S4.3) isincluded in the online supplementary materials.(In the NASA images, 1 Einstein ¼ 1 mole ofphotons.) The reader is encouraged to visit theNASA Web site where images are continually

posted and where the large seasonal variabilityis readily apparent.

The euphotic zone depth (Figure 4.29), whichis defined as the depth of 1% light penetration,is also mapped from satellite color informationusing algorithms based on in situ observations(Lee et al., 2007). The euphotic zone depth isrelated to the historical Secchi disk depths(Figure 4.26 and Section S16.8 of the supple-mentary online materials); the features thatwere described previously for the zonally aver-aged Secchi depths are apparent in the satellite-based map.

Ocean color and derived products are map-ped at a resolution of 4e9 km (as in Figure 4.29).

FIGURE 4.29 Euphotic zone depth (m) from the Aqua MODIS satellite, 9 km resolution, monthly composite forSeptember 2007. (Black over oceans is cloud cover that could not be removed in the monthly composite.) See Figure S4.3from the online supplementary material for the related map of photosynthetically available radiation (PAR). This figure canalso be seen in the color insert. Source: From NASA (2009b).

OPTICAL PROPERTIES OF SEAWATER 109

Color then complements remotely sensed SSTdata of a similar spatial resolution. Chloro-phyll-a is somewhat independent of SST, sothe two products provide powerful informationabout local circulation (advection) and

upwelling (Simpson et al., 1986). The two fieldsare used extensively in studies of regional circu-lation and ecosystems. Examples of ocean colormaps to illustrate regional circulation areincluded throughout later chapters.


typical distributions of water...

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