h�1), only a small fraction of the sea surface iscovered by stage B whitecaps (0.04 or 4%), and aneven smaller fraction of that surface is covered bystage A whitecaps (0.002 or 0.2%). Yet the totalarea of all the world’s oceans is very great(3.61�1014 m2), and as a consequence the total areaof the global ocean covered by whitecaps at anyinstant is considerable. If a wind speed of 7 m s�1 istaken as a representative value, then at any instantsome 7.0�1010 m2, i.e. some 70 000 km2, of stageA whitecap area is present on the surface of theglobal ocean. Following from this, and includingsuch additional information as the terminal risevelocity of bubbles, it can be deduced that some7.2�1011 m2, i.e. some 720 000 km2 of individualbubble surface area are destroyed each second inall the stage A whitecaps present on the surfaceof all the oceans, and an equal area of bubblesurface is being generated in the same interval.The vast amount of bubble surface area destroyedeach second on the surface of all the world’soceans, and the great volume of water (some2.5�1011 m3) swept by all the bubbles that burston the sea surface each second, have profoundimplications for the global rate of air}sea exchangeof moisture, heat and gases. An additional prelimi-nary calculation following along these lines, sug-gests that all the bubbles breaking on the sea surfaceeach year collect some 2 Gt of carbon during theirrise to the ocean surface.

See also

Heat and Momentum Fluxes at the Sea Surface.Wave Generation by Wind.

Further ReadingAndreas EL, Edson JB, Monahan EC, Rouault MP and

Smith SD (1995) The spray contribution to net evapor-ation from the sea: review of recent progress. Bound-ary-Layer Meteorology 72: 3}52.

Blanchard DC (1963) The electriRcation of the atmo-sphere by particles from bubbles in the sea. Progress inOceanography 1: 73}202.

Bortkovskii RS (1987) Air}Sea Exchange of Heat andMoisture During Storms, revised English edition.Dordrecht: D. Reidel [Kluwer].

Liss PS and Duce RA (eds) (1997) The Sea Surface andGlobal Change. Cambridge: Cambridge UniversityPress.

Monahan EC and Lu M (1990) Acoustically relevantbubble assemblages and their dependence on meteoro-logical parameters. IEEE Journal of Oceanic Engineer-ing 15: 340}349.

Monahan EC and MacNiocaill G (eds) (1986) OceanicWhitecaps, and Their Role in Air}Sea Exchange Pro-cesses. Dordrecht: D. Reidel [Kluwer].

Monahan EC and O’Muircheartiaigh IG (1980) Optimalpower-law description of oceanic whitecap coveragedependence on wind speed. Journal of PhysicalOceanography 10: 2094}2099.

Monahan EC and O’Muircheartaigh IG (1986) White-caps and the passive remote sensing of the ocean sur-face. International Journal of Remote Sensing 7:627}642.

Monahan EC and Van Patten MA (eds) (1989) Climateand Health Implications of Bubble-Mediated Sea}AirExchange. Groton: Connecticut Sea Grant CollegeProgram.

Thorpe SA (1982) On the clouds of bubbles formed bybreaking wind waves in deep water, and their role inair}sea gas transfer. Philosophical Transactions of theRoyal Society [London] A304: 155}210.

WIND AND BUOYANCY-FORCED UPPER OCEAN

M. F. Cronin, NOAA Pacific Marine EnvironmentalLaboratory, WA, USAJ. Sprintall, University of California San Diego,La Jolla, CA, USA

Copyright ^ 2001 Academic Press

doi:10.1006/rwos.2001.0157

Introduction

Forcing from winds, heating and cooling, and rain-fall and evaporation, have a profound inSuenceon the distribution of mass and momentum in theocean. Although the effects from this wind and

buoyancy forcing are ultimately felt throughout theentire ocean, the most immediate impact is in thesurface mixed layer, the site of the active air}seaexchanges. The mixed layer is warmed by sunshineand cooled by radiation emitted from the surfaceand by latent heat loss due to evaporation (Figure1). The mixed layer also tends to be cooled bysensible heat loss since the surface air temperature isgenerally cooler than the ocean surface. Evaporationand precipitation change the mixed layer salinity.These salinity and temperature changes deRne theocean’s surface buoyancy. As the surface loses buoy-ancy, the surface can become denser than the sub-surface waters, causing convective overturning and

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WIND AND BUOYANCY-FORCED UPPER OCEAN 3217

Figure 1 Wind and buoyancy forces acting on the upper ocean mixed layer. Solar radiation (Qsw), net longwave radiation (Q lw),latent heat flux (Q lat), and sensible heat flux (Q sen) combine to form the net surface heat flux (Q0). Qpen is the solar radiationpenetrating the base of the mixed layer.

mixing to occur. Wind forcing can also cause sur-face overturning and mixing, as well as localizedoverturning at the base of the mixed layer throughsheared-Sow instability. This wind- and buoyancy-generated turbulence causes the surface water to bewell mixed and vertically uniform in temperature,salinity, and density. Furthermore, the turbulencecan entrain deeper water into the surface mixedlayer, causing the surface temperature and salinityto change and the layer of well-mixed, verticallyuniform water to thicken. Wind forcing can also setup oceanic currents and cause changes in the mixedlayer temperature and salinity through horizontaland vertical advection.

Although the ocean is forced by the atmosphere,the atmosphere can also respond to ocean surfaceconditions, particularly sea surface temperature(SST). Direct thermal circulation, in which moist airrises over warm SSTs and descends over cool SSTs,is most prevalent in the tropics. The resulting atmo-spheric circulation cells inSuence the patterns ofcloud, rain, and winds that combine to form thewind and buoyancy forcing for the ocean. Thus, theoceans and atmosphere form a coupled system,

where it is sometimes difRcult to distinguish forcingfrom response. Because water has a heat capacityand density nearly three orders of magnitude largerthan air, the ocean has thermal and mechanicalinertia relative to the atmosphere. The ocean thusacts as a memory for the coupled ocean}atmospheresystem.

We begin with a discussion of air}sea interactionthrough surface heat Suxes, moisture Suxes, andwind forcing. The primary external force driving theocean}atmosphere system is radiative warming fromthe sun. Because of the fundamental importance ofsolar radiation, the surface wind and buoyancy forc-ing is illustrated here with two examples of theseasonal cycle. The Rrst case describes the seasonalcycle in the north PaciRc, and can be considereda classic example of a one-dimensional (involvingonly vertical processes) ocean response to wind andbuoyancy forcing. In the second example, the sea-sonal cycle of the eastern tropical PaciRc, the atmo-sphere and ocean are coupled, so that wind andbuoyancy forcing lead to a sequence of events thatmake cause and effect difRcult to determine. Theimpact of wind and buoyancy forcing on the surface

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3218 WIND AND BUOYANCY-FORCED UPPER OCEAN

mixed layer and the deeper ocean is summarized inthe conclusion.

Air^Sea Interaction

Surface Heat Flux

As shown in Figure 1, the net surface heat Suxentering the ocean (Q0) includes solar radiation(Qsw), net long-wave radiation (Qlw), latent heatSux due to evaporation (Qlat), and sensible heatSux due to air and water having different surfacetemperatures (Qsen):

Q0"Qsw#Qlw#Qlat#Qsen [1]

The Earth’s seasons are largely deRned by theannual cycle in the net surface heat Sux associatedwith the astronomical orientation of the Earth rela-tive to the Sun. The Earth’s tilt causes solar radi-ation to strike the winter hemisphere more obliquelythan the summer hemisphere. As the Earthorbits the sun, winter shifts to summer and summershifts to winter, with the sun directly overheadat the equator twice per year, in March and againin September. Thus, one might expect the seasonalcycle in the tropics to be semiannual, rather thanannual. However, as discussed later, in someparts of the equatorial oceans, the annual cycledominates due to coupled ocean}atmosphere}landinteractions.

Solar radiation entering the Earth’s atmosphere isabsorbed, scattered, and reSected by water in bothits liquid and vapor forms. Consequently, theamount of solar radiation which crosses the oceansurface, Qsw, also depends on the cloud structures.The amount of solar radiation absorbed by theocean mixed layer depends on the transmissionproperties of the light in the water and can beestimated as the difference between the solar radi-ation entering the surface and the solar radiationpenetrating through the base of the mixed layer (i.e.,Qsw!Qpen in Figure 1).

The Earth’s surface also radiates energy at longerwavelengths similar to a black-body (i.e., propor-tional to the fourth power of the surface temper-ature in units kelvin). Long-wave radiation emittedfrom the ocean surface can reSect against cloudsand become downwelling long-wave radiation to bereabsorbed by the ocean. Thus net long-wave radi-ation, Qlw, is the combination of the outgoing andincoming long-wave radiation and tends to cool theocean.

The ocean and atmosphere also exchange heat viaconduction (‘sensible’ heat Sux). When the ocean

and atmosphere have different surface temperatures,sensible heat Sux will act to reduce the temperaturegradient. Thus when the ocean is warmer than theair (which is nearly always the case), sensible heatSux will tend to cool the ocean and warm theatmosphere. Likewise, the vapor pressure at theair}sea interface is saturated with water while theair just above the interface typically has relativehumidity less than 100%. Thus, moisture tends toevaporate from the ocean and in doing so, the oceanloses heat at a rate of:

Qlat"!L(�fwE) [2]

where Qlat is the latent heat Sux, L is the latent heatof evaporation, �fw is the freshwater density, andE is the rate of evaporation. Qlat has units W m�2,and (�fw E) has units kg s�1 m�2. This evaporatedmoisture can then condense in the atmosphere toform clouds, releasing heat to the atmosphere andaffecting the large-scale wind patterns.

Because sensible and latent heat loss are turbulentprocesses, they also depend on wind speed relativeto the ocean surface, S. Using similarity arguments,the latent (Qlat) and sensible (Qsen) heat Suxes canbe expressed in terms of ‘bulk’ properties at andnear the ocean surface:

Qlat"�aL C�S(qa!qs) [3]

Qsen"�acpaC�S(Ta!Ts) [4]

where �a is the air density, cpa is the speciRc heat ofair, C

�and C

�are the transfer coefRcients of latent

and sensible heat Sux, qs is the saturated speciRchumidity at Ts , the sea surface temperature, andqa and Ta are, respectively, the speciRc humidity andtemperature of the air at a few meters above theair}sea interface. The sign convention used here isthat a negative Sux tends to cool the ocean surface.The transfer coefRcients, C

�and C

�, depend upon

the wind speed and stability properties of the atmo-spheric boundary layer, making estimations of theheat Suxes quite difRcult. Most algorithms estimatethe turbulent heat Suxes iteratively, using Rrst esti-mates of the heat Suxes to compute the transfercoefRcients. Further, the dependence of heat Sux onwind speed and SST causes the system to be coupledsince the heat Suxes can change the wind speed andSST.

Figure 2 shows the climatological net surface heatSux, Q0, and SST for the entire globe. Several pat-terns are evident. (Note that the spatial structure ofthe climatological latent heat Sux can be inferred

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Figure 2 Mean climatologies of (A) net surface heat flux (Wm�2) and, (B) sea surface temperature (3C), and surface winds(m s�1). Climatologies provided by da Silva et al. (1994). A positive net surface heat flux acts to warm the ocean.

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Figure 3 Mean climatologies of (A) evaporation, and (B) precipitation from da Silva et al. (1994). Both have units mm day�1 andshare the scale shown on the right.

from the climatological evaporation shown in Figure3A.) In general, the tropics are heated more than thepoles causing warmer SST in the tropics and coolerSST at the poles. Also, there are signiRcant zonalasymmetries in both the net surface heat Sux andSST. The largest ocean surface heat losses occurover the western boundary currents. In these re-gions, latent and sensible heat loss are enhanced dueto the strong winds which are cool and dry as theyblow off the continent and over the warm watercarried poleward by the western boundary currents.In contrast, the ocean’s latent and sensible heat lossare reduced in the eastern boundary region wheremarine winds blow over the cool water. Conse-quently, the eastern boundary is a region where theocean gains heat from the atmosphere. These spatialpatterns exemplify the rich variability in theocean}atmosphere climate system that occurs ona variety of spatial and temporal scales. In particu-lar, seasonal conditions can often be quite differentfrom mean climatology. The seasonal warming andcooling in the north PaciRc and eastern equatorialPaciRc are discussed later.

Thermal and Haline Buoyancy Fluxes

Since the density of sea water depends on temper-ature and salinity, air}sea heat and moisture Suxescan change the surface density making the watercolumn more or less buoyant. SpeciRcally, the netsurface heat Sux (Q0), rate of evaporation (E) andprecipitation (P) can be expressed as a buoyancySux (B0), as:

B0"!g�Q0/(�cp)#g�(E!P)S0 [5]

where g is gravity, � is ocean density, cp is speciRcheat of water, S0 is surface salinity, � is the effectivethermal expansion coefRcient (!��1R�/RT), and� is the effective haline contraction coefRcient(��1R�/RS). Q0 has units W m�2, E and P have unitsm s�1, and B0 has units m2 s�3. A negative (i.e.,downward) buoyancy Sux, due to either surfacewarming or precipitation, tends to make the oceansurface more buoyant. Conversely, a positive buoy-ancy Sux, due to either surface cooling or evapor-ation, tends to make the ocean surface less buoyant.As the water column loses buoyancy, it can become

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3220 WIND AND BUOYANCY-FORCED UPPER OCEAN

convectively unstable with heavy water lying overlighter water. Turbulence, generated by the ensuingconvective overturning, can then cause deeper, gen-erally cooler, water to be entrained and mixed intothe surface mixed layer (Figure 1). Thus entrain-ment mixing typically causes the SST to cool andthe mixed layer to deepen. As discussed in the nextsection, entrainment mixing can also be generatedby wind forcing.

Figure 3 shows the climatological evaporationand precipitation Relds. Note that in terms of buoy-ancy, a 20 Wm�2 heat Sux is approximately equiva-lent to a 5 mm day�1 rain rate. Thus, in someregions of the world oceans the freshwater Sux termin eqn [5] dominates the buoyancy Sux, and henceis a major factor in the mixed layer thermo-dynamics. For example, in the tropical regions,heavy precipitation can result in a surface-trappedfreshwater pool that forms a shallower mixed layerwithin a deeper nearly isothermal layer. The differ-ence between the shallower mixed layer of uniformdensity and the deeper isothermal layer is referred toas a salinity-stratiRed barrier layer. As the namesuggests, a barrier layer can effectively limit turbu-lent mixing of heat between the ocean surface andthe deeper thermocline since entrainment mixing ofthe barrier layer water into the mixed layer resultsin no heat Sux, if the barrier layer water is the sametemperature as the mixed layer. In regions withparticularly strong freshwater stratiRcation, solarradiation can warm the waters below the mixedlayer, causing the temperature stratiRcation to beinverted (cool water above warmer water). In thiscase, entrainment mixing produces warming ratherthan cooling within the mixed layer.

In subpolar latitudes, freshwater Suxes can alsodominate the surface layer buoyancy proRle. Duringthe winter season, atmospheric cooling of the ocean,and stronger wind mixing leaves the water columnisothermal to great depths. Then, wintertime iceformation extracts freshwater from the surfacelayer, leaving a saltier brine that further increasesthe surface density, decreases the buoyancy, andenhances the deep convection. This process can leadto deep water formation as the cold and salty densewater sinks and spreads horizontally, forcing thedeep, slow thermohaline circulation. Conversely, insummer when the ice shelf and icebergs melt, freshwater is released, and the density in the surface layeris reduced so that the resultant stable halocline(pycnocline) inhibits the sinking of water.

Wind Forcing

The inSuence of the winds on the ocean circulationand mass Reld cannot be overstated. Wind blowing

over the ocean surface causes a tangential stress(‘wind stress’) at the interface which acts as avertical Sux of horizontal momentum. Similar to theair}sea Suxes of heat and moisture, this air}sea Suxof horizontal momentum, �0, can be expressed interms of bulk properties as:

�0"�aC�S(ua!us) [6]

where �a is the air density, C�

is the drag coefR-cient, S is the wind speed relative to the ocean, and(ua!us) is the surface wind relative to the oceansurface Sow. The units of the surface wind stress areN m�2. If the upper ocean has a neutral or stablestratiRcation, then the force imparted by the windstress will generate a momentum Sux that ismaximum at the surface and decreases with depththrough the water column. Consequently, a verticalshear in the velocity develops in the upper oceanmixed layer. The mechanisms by which the mo-mentum Sux extends below the interface are notwell understood. Some of the wind stress goes intogenerating ocean surface waves. However, most ofthe wave momentum later becomes available forgenerating currents through wave breaking, andwave}wave and wave}current interactions. Forexample, wave}current interactions associated withLangmuir circulation can set up large coherent vor-tices which carry momentum to near the base of themixed layer. As with convective overturning, windstirring can entrain cooler thermocline water intothe mixed layer, producing a colder and deepermixed layer. Likewise, current shear generatedby the winds can cause ‘Kelvin Helmholtz’ shearinstabilities that further mix properties within andat the base of the mixed layer.

Over timescales longer than roughly a day, theEarth’s spinning tends to cause a rotation of thevertical Sux of momentum. From the non-inertialperspective of an observer on the rotating Earth, thetendency to rotate appears as a force, referred to asthe Coriolis force. Consequently, the wind-forcedsurface layer transport (‘Ekman transport’) tends tobe perpendicular to the wind stress. Because theprojection of the Earth’s axis onto the local verticalaxis (direction in which gravity acts) changes sign atthe equator, the Ekman transport is to the right ofthe wind stress in the Northern Hemisphere and tothe left of the wind stress in the Southern Hemi-sphere. Convergence and divergence of this Ekmantransport leads to vertical motion which can deformthe thermocline and set the subsurface waters inmotion. In this way, meridional variations in theprevailing zonal wind stress drive the steady, large-scale ocean gyres.

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Figure 4 Seasonal climatologies at the ocean weather stationPAPA in the north Pacific. (A) Wind speed, (B) net surface heatflux, and (C) upper ocean temperature. The bold line representsthe base of the ocean mixed layer defined as the depth wherethe temperature is 0.53C cooler than the surface temperature.Wind speed and net surface heat flux climatologies are from daSilva et al. (1994).

The inSuence of Ekman upwelling on SST can beseen along the eastern boundary of the ocean basinsand along the equator (Figure 2B). Equatorwardwinds along the eastern boundaries of the PaciRcand Atlantic Oceans cause an offshore-directedEkman transport. Mass conservation requires thatthis water be replaced with upwelled water, waterthat is generally cooler than the surface waters out-side the upwelling zone. Likewise, in the tropics,prevailing easterly trade winds cause poleward Ek-man transport. At the equator, this poleward Sowresults in substantial surface divergence and upwell-ing. As with the eastern boundary, equatorial up-welling results in relatively cold SSTs (Figure 2B).Because of the geometry of the continents, thethermal equator favors the Northern Hemisphereand is generally found several degrees of latitudenorth of the equator.

In the tropics, winds tend to Sow from cool SSTsto warm SSTs, where deep atmospheric convectioncan occur. Thus, surface wind convergence in theinter-tropical convergence zone (ITCZ) is associatedwith the thermal equator, north of the equator. Therelationship between the SST gradient and windsaccounts for an important coupling mechanism inthe tropics.

The Seasonal Cycle

The North Paci\c: A One-dimensional OceanResponse to Wind and Buoyancy Forcing

From 1949 through 1981, a ship (Ocean StationPAPA) was stationed in the North PaciRc at503N 1453W with the primary mission of takingroutine ocean and atmosphere measurements.The seasonal climatology observed at this site(Figure 4) illustrates a classic near-one-dimensionalocean response to wind and buoyancy forcing.A one-dimensional response implies that only thevertical structure of the ocean is changed by theforcing.

During springtime, layers of warmer and lighterwater are formed in the upper surface in responseto the increasing solar warming. By summer, thisheating has built a stable (buoyant), shallowseasonal pycnocline that traps the warm surfacewaters. In fall, storms are more frequent andnet cooling sets in. By winter, the surface layeris mixed by wind stirring and convective overturn-ing. The summer pycnocline is eroded and themixed layer deepens to the top of the permanentpycnocline.

To Rrst approximation, horizontal advection doesnot seem to be important in the seasonal heat

budget. The progression appears to be consistentwith a surface heat budget described by:

RT/Rt"Q0/(�cpH) [7]

where RT/Rt is the local time rate of change ofthe mixed layer temperature, and H is the mixedlayer depth. Since only vertical processes (e.g.,turbulent mixing and surface forcing) affectthe depth and temperature of the mixed layer,the heat budget can be considered one-dimen-sional.

A similar one-dimensional progression occurs inresponse to the diurnal cycle of buoyancy forcingassociated with daytime heating and nighttime cool-ing. Mixed layer depths can vary from just a fewmeters thick during daytime to several tens ofmeters thick during nighttime. Daytime and night-time SST can sometimes differ by over one degreeCelsius. However, not all regions of the ocean havesuch an idealized mixed layer seasonal cycle. Oursecond example shows a more complicated seasonalcycle in which the tropical atmosphere and oceanare coupled.

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Figure 5 Seasonal climatologies of the tropical Pacific sea surface temperature (3C) and wind (m s�1) for (A) February}March}April, (B) May}June}July, (C) August}September}October, and (D) November}December}January. Winds are from da Silvaet al. (1994). SSTs are from Reynolds and Smith (1994).

The Eastern Equatorial Paci\c: CoupledOcean^Atmosphere Variability

Because there is no Coriolis turning at the equator,water and air Sow are particularly susceptible tohorizontal convergence and divergence. Smallchanges in the winds patterns can cause large vari-ations in oceanic upwelling, resulting in signiRcantchanges in SST and consequently in the atmosphericheating patterns. This ocean and atmosphere coup-ling thus causes initial changes to the system toperpetuate further changes.

At the equator, the sun is overhead twice peryear: in March and again in September. Thereforeone might expect a semiannual cycle in the mixedlayer properties. Although this is indeed found insome parts of the equatorial oceans (e.g., in thewestern equatorial PaciRc), in the eastern equatorialPaciRc the annual cycle dominates. During thewarm season (February}April), the solar equinoxcauses a maximum in insolation, equatorial SST iswarm and the meridional SST gradient is weak.

Consequently, the ITCZ is near the equator, andoften a double ITCZ is observed that is symmetricabout the equator. The weak winds associated withthe ITCZ cause a reduction in latent heat loss, windstirring, and upwelling, all of which lead to furtherwarming of the equatorial SSTs. Thus the warm SSTand surface heating are mutually reinforcing.

Beginning in about April}May, SSTs begin to coolin the far eastern equatorial PaciRc, perhaps in re-sponse to southerly winds associated with the conti-nental monsoon. The cooler SSTs on the equatorcause an increased meridional SST gradient thatintensiRes the southerly winds and the SST coolingin the far eastern PaciRc. As the meridional SSTgradient increases, the ITCZ begins to migratenorthward. Likewise, the cool SST anomaly in thefar east sets up a zonal SST gradient along theequator that intensiRes the zonal trade winds to thewest of the cool anomaly. These enhanced tradewinds then produce SST cooling (through increasedupwelling, wind stirring and latent heat loss) thatspreads westward (Figure 5).

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By September, the equatorial cold-tongue is fullyformed. Stratus clouds, which tend to form overthe very cool SSTs in the tropical PaciRc, causea reduction in solar radiation, despite the equinoc-tial increase. The large meridional gradient in SSTassociated with the fully formed cold tongue causesthe ITCZ to be at its northernmost latitude. Afterthe cold tongue is fully formed, the reduced zonalSST gradient within the cold tongue causes the tradewinds to weaken there, leading to reduced SST cool-ing along the equator. Finally, by February,the increased solar radiation associated with theapproaching vernal equinox causes the equatorialSSTs to warm and the cold tongue to disappear,bringing the coupled system back to the warm sea-son conditions.

Conclusion

Because the ocean mixed layer responds so rapidlyto surface-generated turbulence through wind andbuoyancy forced processes, the surface mixed layercan often be modeled successfully using one-dimen-sional (vertical processes only) physics. Surface heat-ing and cooling cause the ocean surface to warmand cool; evaporation and precipitation cause theocean surface to become saltier and fresher. Stabiliz-ing buoyancy forcing, whether from net surfaceheating or precipitation, stratiRes the surface andisolates it from the deeper waters. Whereas windstirring and destabilizing buoyancy forcing generatesurface turbulence that cause the surface propertiesto mix with deeper water. Eventually, however,one-dimensional models drift away from observa-tions, particularly in regions with strong ocean}atmosphere coupling and oceanic current structures.The effects of horizontal advection are explicitly notincluded in one-dimensional models. Likewise, verti-cal advection depends on horizontal convergencesand divergences and therefore is not truly a one-dimensional process. Finally, wind and buoyancyforcing can themselves depend on the horizontalSST patterns, blurring the distinction between forc-ing and response.

Although the mixed layer is the principal regionof wind and buoyancy forcing, ultimately the effectsare felt throughout the world’s oceans. Both thewind-driven motion below the mixed layer and thethermohaline motion in the relatively more quiesc-ent deeper ocean originate through forcing in the

surface layer that causes an adjustment in the massReld (i.e., density proRle). In addition, buoyancy andwind-forcing in the upper ocean deRne the propertycharacteristics for all the individual major watermasses found in the world oceans. On a globalscale, there is surprisingly little mixing betweenwater masses once they acquire the characteristicproperties at their formation region and are verti-cally subducted or convected from the active surfacelayer. As these subducted water masses circulatethrough the global oceans and later outcrop, theycan contain the memory of their origins at the sur-face through their water mass properties and thuscan potentially induce decadal and centennial modesof variability in the ocean}atmosphere climatesystem.

See also

Breaking Waves and Near-surface Turbulence. Ek-man Transport and Pumping. Langmuir Circulationand Instability. Paci\c Ocean Equatorial Currents.Penetrating Shortwave Radiation. ThermohalineCirculation. Upper Ocean Heat and FreshwaterBudgets. Upper Ocean Responses to Strong Forc-ing Events. Upper Ocean Time and Space Variabil-ity. Upper Ocean Vertical Structure. Wind DrivenCirculation. Water Types and Water Masses.

Further Readingda Silva AM, Young CC and Levitus S (1994) Atlas of

Surface Marine Data 1994, vol 1. Algorithms and Pro-cedures, NOAA Atlas NESDIS 6. Washington: USDepartment of Commerce.

Kraus EB and Businger JA (1994) Atmosphere}OceanInteraction, Oxford Monographs on Geology andGeophysics, 2nd edn. New York: Oxford UniversityPress.

Niiler PP and Kraus EB (1977) One-dimensional modelsof the upper ocean. In: Kraus EB (ed.) Modellingand Prediction of the Upper Layers of the Ocean,pp. 143}172. New York: Pergamon.

Philander SG (1990) El NinJ o, La NinJ a, and the SouthernOscillation. San Diego: Academic Press.

Price JF, Weller RA and Pinkel R (1986) Diurnal cycling:observations and models of the upper ocean responseto diurnal heating, cooling, and wind mixing. Journalof Geophysical Research 91: 8411}8427.

Reynolds RW and Smith TM (1994) Improved global seasurface temperature analysis using optimum inter-polation. Journal of Climate 7: 929}948.

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3224 WIND AND BUOYANCY-FORCED UPPER OCEAN