The Location of Diapycnal Mixing and the Meridional Overturning Circulation

3578 VOLUME 32J O U R N A L O F P H Y S I C A L O C E A N O G R A P H Y

q 2002 American Meteorological Society

The Location of Diapycnal Mixing and the Meridional Overturning Circulation

JEFFERY R. SCOTT

Program in Atmospheres, Oceans, and Climate, Massachusetts Institute of Technology, Cambridge, Massachusetts

JOCHEM MAROTZKE

School of Ocean and Earth Science, Southampton Oceanography Centre, Southampton, United Kingdom

(Manuscript received 1 May 2001, in final form 14 June 2002)

ABSTRACT

The large-scale consequences of diapycnal mixing location are explored using an idealized three-dimensionalmodel of buoyancy-forced flow in a single hemisphere. Diapycnal mixing is most effective in supporting a strongmeridional overturning circulation (MOC) if mixing occurs in regions of strong stratification, that is, in the low-latitude thermocline where diffusion causes strong vertical buoyancy fluxes. Where stratification is weak, suchas at high latitudes, diapycnal mixing plays little role in determining MOC strength, consistent with weakdiffusive buoyancy fluxes at these latitudes. Boundary mixing is more efficient than interior mixing at drivingthe MOC; with interior mixing the planetary vorticity constraint inhibits the communication of interior watermass properties and the eastern boundary. Mixing below the thermocline affects the abyssal stratification andupwelling profile but does not contribute significantly to the meridional flow through the thermocline or theocean’s meridional heat transport. The abyssal heat budget is dominated by the downward mass transport ofbuoyant water versus the spread of denser water tied to the properties of deep convection, with mixing of minorimportance. These results are in contrast to the widespread expectation that the observed enhanced abyssalmixing can maintain the MOC; rather, they suggest that enhanced boundary mixing in the thermocline needsto be identified in observations.

1. Introduction

Microstructure and tracer release measurements ofdiapycnal mixing in the ocean (Polzin et al. 1997; Led-well et al. 1993, 2000) show that mixing is stronglylocalized, with diffusivities exceeding 1024 m2 s21

above rough bottom topography and an order of mag-nitude less above smooth abyssal plains and in the ther-mocline. The energy for the diapycnal mixing is thoughtto come from winds and tides (Munk and Wunsch 1998,hereafter MW) and perhaps geothermal sources (Huang1999). The strength of the meridional overturning cir-culation (MOC) in models depends strongly on thechoice of the vertical or diapycnal diffusivity (Bryan1987; Colin de Verdiere 1988; Zhang et al. 1999).Hence, diapycnal mixing plays a crucial role in the dy-namics of the MOC and of climate because the MOCis an important transport agent of properties relevant forclimate.

To date, there have been only a few model studiesthat specifically examine the dynamical consequencesof mixing location. Using an idealized single-hemi-

Corresponding author address: Dr. Jeffery R. Scott, MIT, Room54-1711, Cambridge, MA 02139.E-mail: [email protected]

sphere ocean general circulation model, Cummins et al.(1990, hereafter CHG) parameterized vertical diffusivityas a function of the buoyancy frequency, effectivelyincreasing mixing at depth, particularly below the ther-mocline. Cummins (1991) examined the results of sev-eral additional runs with specified increased mixing be-low the thermocline. Using a similar model withoutwind forcing, Marotzke (1997, hereafter M97) imposedmixing only along the boundaries and applied the resultsas a foundation for a self-contained theory predictingthe strength of the MOC. Samelson (1998) applied lo-calized mixing on the eastern boundary to an idealizedwind- and buoyancy-forced single hemisphere, plane-tary geostrophic model. Hasumi and Suginohara (1999)investigated the effects of enhanced mixing over to-pography in a global model, and Marotzke and Klinger(2000) analyzed the effects of equatorially asymmetricvertical mixing.

Using a single-hemisphere model of the ocean, weexplore the effect of spatially varying diapycnal mixing.We juxtapose various extreme scenarios, in that mixingis concentrated entirely at low or high latitudes; at thewestern boundary, the eastern boundary, or the interior;in or below the thermocline. To our knowledge, theMOC’s sensitivity to mixing in so clearly identifiableregimes of the ocean has never been investigated. Our

DECEMBER 2002 3579S C O T T A N D M A R O T Z K E

TABLE 1. Summary of numerical parameters.

Parameter Value

Basin width, lengthBasin depthHorizontal, vertical viscosityHorizontal, isopycnal diffusivityIsopycnal thickness diffusionLongitude, latitude grid spacingNumber of levelsTemperature, salinity restoringtimescale

Time step, momentumTime step, tracers

608, 6484500 m3.3 3 104 m2 s21, 1022 m2 s21

0 m2 s21, 103 m2 s21

103 m2 s21

1.8758, 283030 days

30 min12 h

numerical results lead us to revisit the advective–dif-fusive balance in the abyss, the deep-ocean heat budget,and how planetary vorticity conservation helps us tounderstand the MOC. Our configuration is very highlyidealized (like M97, we omit wind forcing), but wemaintain that the single hemisphere ocean is an appro-priate configuration to represent the fundamental dy-namics of the MOC.

This paper is organized as follows. In section 2, webriefly describe the ocean model. The ramifications ofmixing that is highly localized in the horizontal are ex-amined in section 3. In section 4 we present experimentswith depth-dependent diapycnal mixing and examine thedeep ocean heat balance and the structure of the over-turning cell. We end with a discussion and conclusionsin section 5 and section 6, respectively.

2. Model description

We employ the z-coordinate, primitive equation mod-el MOM 2 (beta version 2.0), as described in Paca-nowski (1996). Default parameters are listed in Table1. The ocean configuration and forcing are identical tothat in M97: the domain is a 608 wide single hemispheresector, ranging from the equator to 648, with a constantdepth of 4500 m; temperature and salinity are forcedusing an identical, zonally uniform cosine profile, withpeak-to-peak amplitudes of 278C and 1.5 psu and a 30-day relaxation time constant. Given these identical forc-ing profiles, the model can be thought of as being forcedby buoyancy. Higher-order effects from the nonlinearequation of state are included in the model but are notthought to be important in the results presented here.Therefore, we do not distinguish between temperatureand buoyancy. For simplicity, no wind stress is imposed.Our vertical grid spacing ranges from 50 m at the surfaceto 250 m at the lowest level.

As in M97, diapycnal mixing is imposed in the col-umns adjacent to the north, south, east, and west side-walls, and is set to zero elsewhere; this is thought tomimic the effect of enhanced mixing due to a slopinglateral boundary. Diapycnal mixing at the equator is asurrogate for the global integral of mixing throughoutthe rest of the world’s oceans. Although there is evi-

dence of enhanced mixing at the equator (Gregg 1987;Peters et al. 1988) and the dynamics there are uniquedue to the vanishing of the Coriolis parameter (Gill1982), our choice of the equator for our southern wallis for practical reasons, as cross-equatorial flow involvescomplicated dynamics (Marotzke and Klinger 2000) thatare beyond the scope of our investigation. To insure thatequatorial dynamics were not important in our findings,we spun up a run with the southern boundary mixingmoved one grid cell northward (i.e., removed from theequator), with only minor differences resulting.

Mixing is implemented along isopycnals, using theRedi (1982) isoneutral diffusion tensor. We employedthe MOM 2 ‘‘full tensor’’ option, keeping all terms inthe diffusion tensor. In locations of steeply sloping is-opycnals, use of this option requires only mild rescalingof the isoneutral diffusion coefficient, as compared tothat using the approximate form of the tensor (Griffieset al. 1998; see also Pacanowski 1996). All runs use theGent–McWilliams (1990) parameterization to representthe effect of mesoscale eddies on isopycnals. Our runsuse the (original) advective-flux implementation of theGent–McWilliams parameterization, as is implementedin MOM 2.

All model runs were integrated to equilibrium, as de-fined by a basin-averaged surface heat flux of 5 3 1023

W m22 or less, where practical, and/or when overturningis discernibly within 0.1 Sv (Sv [ 106 m3 s21) of itsfinal value. The control run was integrated to equilib-rium from an isothermal, motionless ocean. All otherexperiments were started from the control run or fromthe equilibrated state of another experiment.

A summary of the numerical experiments is presentedin Table 2. Diapycnal mixing does not vary with depthin experiments A–K; in these runs, we aim to understandthe effect of mixing location (and degree of localization)in the horizontal. Our control run, experiment A, is verysimilar to the run described at length in M97. Exceptfor experiments G–K, which employ a background dif-fusivity of 0.1 3 1024 m2 s21 for numerical purposes,diffusivity is set to zero where not prescribed in Table2. In experiments L–O, diapycnal mixing occurs onlyalong the boundaries, as in our control run, but variesin the vertical, as plotted in Fig. 1.

For computation efficiency, experiments F–K wererun at half resolution, that is, 3.758 3 48, 16 verticallevels. In these runs, the diffusivity was adjusted so thatthe area-integrated diapycnal mixing approximatelymatched that of the control run. In addition, horizontalviscosity was changed to resolve the Munk boundarylayer at the new zonal grid spacing. In section 4, severalof the experiments were run with high vertical resolution(90 evenly spaced levels) in order to minimize any ad-verse numerical effects and to allow for a smootherrepresentation of stratification. In several direct com-parisons (not shown), model results did not differ sig-nificantly between similarly configured runs at differentvertical and/or horizontal resolution.


TABLE 2. Summary of numerical experiments. Mixing in expts A–K is constant with depth.

Expt Mixing location k (m2 s21)

ABCDEFGHIJK

N, S, E, W boundariesUniformN, S, E, W boundaries, double-widthS boundary; E, W boundaries, 0–368NS boundary onlyN, S boundaries; 7.58 wide midbasin meridional strip3.758 3 48 patch along western boundary, various latitudes3.758 3 48 patch along eastern boundary, various latitudes3.758 3 48 patch, located at midbasin, various latitudesThree 3.758 3 48 patches, located along the S boundary3.758 3 48 patch, centered at 1.8758E, 28N; rapid restoring

10 3 1024

1.15 3 1024

5 3 1024

10 3 1024

10 3 1024

5 3 1024

80 3 1024

80 3 1024

80 3 1024

27 3 1024

80 3 1024

Upper k Abyssal k

LMNO

Weak thermocline mixing; N, S, E, W boundariesWeak deep mixing; N, S, E, W boundariesStrong deep mixing; N, S, E, W boundariesAlternate strong deep mixing; N, S, E, W boundaries

see Fig. 110 3 1024

10 3 1024

10 3 1024

10 3 1024

see Fig. 1see Fig. 1see Fig. 1

FIG. 1. Vertical profile of boundary diapycnal diffusivity for weakthermocline mixing (expt L), weak deep mixing (expt M), and ourtwo strong deep mixing experiments (expts N and O).

In order to minimize numerical ‘‘wiggles’’ resultingfrom zero diapycnal mixing in the ocean interior, weemployed the flux-corrected transport advectionscheme, a nonlinear compromise between upstream andcentered differences (Gerdes et al. 1991). In effect, thisscheme minimizes numerical noise through the intro-duction of some diapycnal mixing. We maintain thisscheme’s supplementary mixing is inconsequential here,based on trial experiments using MOM’s other advectionschemes and from the results of runs with very weakboundary mixing. The reader is referred to M97 for adiscussion of other numerical issues involved with theboundary mixing implementation.

Before proceeding, some additional comments are inorder regarding boundary layers, several of which can-not be resolved by our model. Reasonable treatment ofwestern boundary currents is possible, albeit using an

unrealistic viscosity coefficient due to our coarse gridspacing. As noted in Huck et al. (1999), however, theparameterization of the lateral boundary conditions caninfluence the large-scale circulation. Here, two notablefeatures of our solution—narrow upwelling along theeastern and western boundary and deep downwelling inthe northeast corner—are enhanced by (or perhaps evencaused by) by our use of no-slip side boundaries withLaplacian momentum dissipation. The so-called Veroniseffect (Veronis 1975) whereby a Cartesian implemen-tation of diffusion is thought to effect spurious hori-zontal mixing in the western boundary, producing up-welling, does not occur here given our use of the iso-pycnal mixing scheme. Huck et al. (1999) argue thatthe lateral boundary parameterization induces upwell-ing, which in turn causes the Veronis effect rather thanvice versa. Huck et al. also showed that the model so-lution differs when a linear frictional closure schemefor tangential velocity is introduced into the vorticityequation, such as proposed in Winton (1993). It is notclear which boundary parameterization is superior forthe purpose of modeling the real ocean, given a lack ofobservational evidence and our limited understandingof eddy dissipation processes (see Huck et al. for a morecomplete discussion). For practical reasons, we mustassume that the results described here reflect upon thefundamental thermodynamics of the model’s large-scalecirculation rather than local features specific to the mod-el implementation. The effect of other omissions, par-ticularly the absence of wind stresses and topography,is addressed in our discussion section.

3. Horizontal location of mixing

a. Boundary versus uniform mixing

In the model runs of M97, boundary flows set up aneast–west temperature gradient that, through thermalwind balance, supports a MOC. We have repeated theM97 boundary mixing control run here (expt A); minor


FIG. 2. Boundary mixing run (expt A), k 5 10 3 1024 m2 s21. (a)Temperature (contour levels as indicated) and flow along the westernwall; (b) temperature and flow along the eastern wall. Vertical andhorizontal velocity scales are shown for reference.

FIG. 3. Meridional overturning streamfunction (contours) and zon-ally averaged temperature (shading) for (a) boundary mixing run (exptA), k 5 10 3 1024 m2 s21 and (b) uniform mixing run (expt B), k5 1.15 3 1024 m2 s21. In this and all subsequent plots of overturningstreamfunction and zonally averaged temperature, overturning con-tours interval is 1 Sv; isotherms are at 0.05, 0.1, 0.2, 0.4, and 0.8 3DT (DT 5 278C); flow is oriented clockwise around overturningmaximum.

differences are due to our use of improved resolutionand the isopycnal mixing scheme. As in M97, upwellingoccurs all along the west wall (Fig. 2a), advecting densedeep waters into the thermocline. On the east wall (Fig.2b), however, the vertical flow pattern is more compli-cated. Upwelling occurs at depth, but surface flowdownwells to increasing depths at high latitudes. Sincethe downwelling surface water is relatively warm, theeastern wall is less dense than the western wall, pro-viding the necessary shear for zonally integrated south-ward flow at depth and northward flow in the upperocean.

The meridional mass transport streamfunction for thecontrol run is shown in Fig. 3a. Over 5 Sv, or almosthalf of the net mass transport, upwells adjacent to the

equator where diapycnal mixing is concentrated. To ex-amine the effect of the boundary mixing parameteri-zation, we equilibrated a run with uniform diapycnalmixing diffusivity of 1.15 3 1024 m2 s21 (expt B), thatis, that which produces an area-integrated diffusivityequivalent to that of the control run. Detailed analysesof a similarly configured uniform mixing run, albeit withcruder numerics, are presented in Colin de Verdiere(1988). Superficially, the MOC of the uniform mixingrun (Fig. 3b) shows little difference from the boundarymixing run. The maximum of the overturning streamfunction is 1.7 Sv less in the uniform case, reducing theocean’s northward peak heat transport from 0.55 to 0.47PW (PW 5 1015 watts). With mixing spread out moreevenly over the low latitudes, a much smaller proportionof the upward mass transport flows adjacent to the south-ern boundary. As would be suggested by our discussionin the previous section, strong vertical flows are presentalong the east and west boundaries, even in the uniformmixing case. Some of this flow recirculates zonally with-out contributing to the MOC [see Bryan (1987) for adiagnosis of the meridionally averaged circulation and


FIG. 4. Two-column boundary mixing run (expt C), k 5 5 3 1024 m2 s21. (a) Temperature and flow along the meridional plane through258E; (b) temperature and flow along the zonal plane through 138N; (c) temperature and flow along the zonal plane through 358N; (d)temperature and flow along the zonal plane through 558N.

its scaling behavior with vertical diffusivity], whilesome of western upwelling moves northward along is-opycnals. As such, it is not immediately clear what com-ponent of these boundary flows is diapycnal.

To determine which boundary flows are directly in-duced by diapycnal mixing, presumably through ad-vective–diffusive balance, and which are a largely aconsequence of lateral boundaries, we ran an experimentwhere we expanded the region of diapycnal mixing totwo boundary grid columns around the model sidewalls(expt C). For consistency, we decreased the magnitudeof k by 50%. In addition, we resolved the Munk bound-ary layer across two zonal grid points. As with the uni-form mixing run, imposing mixing away from theboundaries results in a decreased maximum in the over-

turning streamfunction (not shown), although the re-duction here is only 0.3 Sv. Near the equator, both me-ridional grid columns with mixing show strong up-welling, as shown in Fig. 4a, while interior flows arevery weak in the meridional–vertical plane. In low lat-itudes, there are two columns of strong upwelling atboth the eastern and western boundaries (Fig. 4b). Thisresult is consistent with Colin de Verdiere (1988), whodiagnosed that the primary balance in low latitudes wasbetween diffusive heating and cold upwelling. Similarly,Samelson (1998) observed strong upwelling in low lat-itudes along the eastern boundary, where his mixingwas concentrated. Although some upwelling is evidentin two columns at the boundaries at midlatitudes (Fig.4c), the magnitude is much larger in the columns di-


FIG. 5. Meridional overturning streamfunction (contours) and zon-ally averaged temperature (shading) for (a) low-latitude boundarymixing run (expt D), k 5 10 3 1024 m2 s21 between 0 and 368 lat,otherwise diffusivity is set to zero; and (b) equatorial mixing run(expt E), k 5 10 3 1024 m2 s21 between 08 and 28 lat, otherwisediffusivity is set to zero.

rectly adjacent to the boundary. At high latitudes, nearlyall upwelling in the west occurs adjacent to the bound-ary, as shown in Fig. 4d; in the east, some upwellingis apparent in both columns, although there is a largedisparity in the velocities, as in the midlatitude section.These results suggest that the mixing at low latitudes,in effect, drives local upwelling through vertical ad-vective–diffusive balance. At middle and high latitudes,where stratification is generally weak, a large percentageof the vertical flow at the east–west boundaries is theresult of mass convergence and subsequently a com-ponent of the vertical flow is oriented along isopycnals.

b. Low-latitude mixing and midbasin mixing

Motivated by these results, we wish to examinewhether mid- and high-latitude mixing plays any sig-nificant role in the dynamics of the modeled MOC. Fig-ure 5a shows the meridional overturning streamfunctiongiven boundary mixing from the equator to 368N, withno diapycnal mixing to the north (expt D). As comparedwith the control run (Fig. 3a), the center of the over-turning cell is several hundred meters higher in the water

column, but there are no apparent differences in thezonally averaged temperature profile and the differencein overturning maximum is only 0.3 Sv. In addition,there are only slight differences in the east and westwall boundary layer flows in mid and high latitudes (notshown).

If we further concentrate all mixing at the equator(expt E), the maximum in overturning drops by 38% to7.9 Sv (Fig. 5b), as we have reduced the area of mixingfrom the previous low-latitude mixing experiment by50%. Thus, the subtropical mixing on the east and westwalls does contribute in driving the MOC. This resultis consistent with significant upwelling into the ther-mocline along these latitudes, as suggested by the over-turning pattern in runs that include mixing in the sub-tropics (e.g., Figs. 3a and 5a).

As mentioned in the introduction, the results fromrecent microstructure measurements suggest elevatedmixing in the water column above the Mid-AtlanticRidge. To this end, we ran a variation on our controlrun where we moved the mixing on the eastern andwestern boundaries to two adjacent meridional stripsdown the middle of the ocean, preserving the area-in-tegrated diffusivity (expt F). This experiment produceda similar pattern of overturning as the control run (notshown). However, consistent with other runs that im-posed interior mixing, the overturning cell was slightlyweaker, comparable in magnitude with the uniform mix-ing experiment.

c. Highly localized mixing

The previous results suggest that the MOC cell de-pends critically on the meridional distribution of dia-pycnal mixing, with the zonal distribution being lessimportant. In this subsection we take these experimentsto their logical extreme, localizing mixing to a singlegrid column. Our motivation here is to facilitate a moredetailed examination of the dynamical significance ofinterior versus boundary mixing (we do not claim thatthese extreme scenarios are representative of the realocean). The diapycnal diffusivity in the ‘‘mixing col-umn’’ was chosen so that the area-weighted diffusivityin latitudes 08–368N matched that of the control run.However, to control numerical difficulties we also addeda background diffusivity of 0.1 3 1024 m2 s21, the valuetypically assumed for the ‘‘pelagic diffusivity’’ (MW).In a run with uniform diffusivity set at this backgroundvalue, the overturning streamfunction maximum is 2.3Sv, considerably weaker than that observed in this setof experiments.

We equilibrated runs with the mixing column at threezonal locations—the western boundary (expt G), at theeastern boundary (expt H), and at midbasin (expt I)—and at latitudes ranging from 28N to 508N. Consistentwith the previous results, nearly all the MOC’s up-welling occurs where mixing is located. Figures 6a–cshow the temperature and flow in the bottom layer, ther-


FIG. 6. Plan view of the circulation and temperature (contours)with the mixing column located along the western boundary at 188N(expt G): (a) lowest model level (depth 4245 m), (b) thermoclinelevel (depth 700 m), and (c) uppermost model level (depth 25 m).Velocity scales are shown for reference.

mocline, and surface layer, respectively, for a mixingcolumn located at 188N along the western boundary.The plots show deep convergence and upper level di-vergence in the mixing column; note that in this con-figuration, the deep western boundary current is effec-tively short-circuited by the upwelling in the mixingcolumn. Figures 7 and 8 show a similar low-level con-vergence and upper-level divergence in the mixing col-umn when it is situated at the eastern boundary or oceaninterior, respectively.

Figure 9 shows the maximum in overturning stream-function for the three series as a function of the mixingcolumn latitude. As the mixing location moves farthernorth, the MOC decreases in intensity, most noticeablywhen the mixing is located in the interior. Conversely,the circulation remains strongest when the mixing islocated on the eastern boundary.

As a starting point in explaining Fig. 9, let us examine

the scaling behavior of the MOC. A simple predictivescaling relationship for the overturning circulation wasfirst presented by Bryan and Cox (1967) (see also We-lander 1971). The following balances are employed:

vertical advective–diffusive balance:

kW ; (1)

D

continuity:

V W; (2)

Dx D

thermal wind balance:

V ga DT; , (3)

D f Lx

where V and W are horizontal and vertical velocity


FIG. 7. Plan view of the circulation and temperature (contours)with the mixing column located along the eastern boundary at 188N(expt H): (a) lowest model level (depth 4245 m) and (b) thermoclinelevel (depth 700 m).

FIG. 8. Plan view of the circulation and temperature (contours)with the mixing column located at midbasin (328E) and 188N (exptI): (a) lowest model level (depth 4245 m) and (b) thermocline level(depth 700 m).

scales, respectively, D is a vertical length scale [albeitsomewhat ambiguously defined; see Scott (2000) for amore complete discussion of how this contributes toproblems in using (1)–(3) to predict scaling behavior],and k is a vertical (or diapycnal) diffusivity. Tradition-ally, DT is taken as the equator-to-pole temperature gra-dient, and Dx and Lx are taken to be identical, a hori-zontal length scale. Marotzke (1997) and Park and Bry-an (2000) suggest that the equator-to-pole temperaturegradient scales as the east–west temperature gradient; itis the latter that is directly related to the meridionalvelocity and thus meridional overturning. Here, we con-sider Dx as a horizontal length scale over which up-welling occurs, corresponding to the generation of hor-izontal velocity by convergence of vertical velocity; Lx

is a horizontal length scale over which the east–westtemperature gradient occurs (presumably at midlati-tudes, near the maximum in overturning). As noted inM97, the relevant east–west temperature gradient occursacross the western boundary current.

By eliminating W from (1) and (2) and eliminatingV using (3), the following scales for V and D are ob-tained:

1/3 1/32kDx(gaDT ) kDxf LxV ; , D ; . (4)2 2 1 2[ ]f L gaDTx

The meridional overturning C can in turn be estimatedas ;VDL, suggesting a (kDx)2/3 dependence. In otherwords, both the depth scale and the maximum of over-turning scale as a function of the area-integrated mixing


FIG. 9. Maximum in overturning streamfunction for highly local-ized mixing expts G, H, I, and J. The abscissa reflects the meridionallocation of the mixing column with k 5 160 3 1024 m2 s21, otherwisethe diffusivity is set to a background value of k 5 0.1 3 1024 m2

s21 (these runs were done using 3.758 3 48 resolution). The threeseries show results for different zonal locations of the mixing column,i.e., adjacent to the western wall, adjacent to the eastern wall and atmidbasin (expts G, H, and I, respectively). The solid star indicatesthe result when mixing is equally divided into three equatorial col-umns at these zonal locations (expt J).

(see also Samelson 1998), consistent with our resultshere (Dx encapsulates both the meridional scale and thezonal scale of the mixing area). Note that neither thebasin width nor the basin’s meridional scale (i.e., thedistance between the tropical mixing and high latitudes)enters the scaling relationship. This was confirmed in atest run where we decreased the zonal width of the basin(with mixing in a single column) and a second test runwith ‘‘tropical’’ mixing (and tropical temperatures) lo-cated at 208N rather than along the equator. In both thesetest runs, the maximum in overturning was littlechanged. Thus, changes in either zonal or meridionallength scales associated with our localized mixing runparameterization do not explain the behavior exhibitedin Fig. 9.

Upon further scrutiny, there are three competing ef-fects that are important here, which we address sepa-rately. Although the first effect is captured in (4), thelatter two involve dynamics that the scaling relation isincapable of representing.

1) SURFACE BOUNDARY CONDITIONS

An examination of the surface temperature immedi-ately suggests why all of the highly localized mixingexperiments plotted in Fig. 9 produce weaker overturn-ing than the control mixing run: where upwelling occurs,

the surface is quite cold, approximately 108C colder thanneighboring grid points, so the diffusion of heat into thethermocline is far less efficient than in the less localizedmixing runs [i.e., decreasing DT in (4)]. In the midbasin,equatorial mixing run the surface anomaly in temper-ature is about 18C less than when the mixing is at theequatorial western or eastern boundary, consistent withits slightly stronger overturning circulation. To furthertest this hypothesis, we ran two additional experiments.First, we divided the mixing evenly between three equa-torial columns located in the west, east, and midbasin(expt J). As indicated by the solid star in Fig. 9, theoverturning circulation was 2 Sv stronger, consistentwith the steady-state temperature at these three pointsbeing much closer to the restoring profile than in thesingle column mixing runs. Second, we changed thesurface restoring time constant from 30 days to 2 days(expt K). With the mixing column located in the south-west corner, the circulation increased to match that ofthe control run (but since the mixing column experi-ments employ the weak pelagic background mixing, wecaution that the nearly exact agreement is not quite as‘‘clean’’ as this result might suggest).

As the mixing column moves north, the surface re-storing temperature decreases, providing a simple ex-planation for the noted decrease in overturning strength.The dashed line in Fig. 9 is a plot of the observed modelsurface temperature for the eastern mixing series as afunction of mixing column latitude, taken to the two-thirds power and then normalized to coincide with themaximum overturning at the equator (the decrease insurface temperature with latitude is similar for the west-ern and midbasin mixing series). This power law is cho-sen as the approximate functional dependence of max-imum overturning on the high to low latitude temper-ature difference [Scott 2000; note this differs somewhatfrom that predicted by (4)]. We see that with westernboundary mixing, the fall off with latitude is close tothat predicted by the decrease in surface temperature.However, at high latitudes the eastern boundary mixingseries is stronger than predicted while the interior mix-ing series is weaker, suggesting that additional factorsplay a role in determining the overturning strength asthe mixing latitude is varied.

2) INTERIOR VERSUS BOUNDARY MIXING

Dynamical considerations suggest a different behav-ior of the interior localized mixing runs, compared tothe boundary mixing runs; in this section, we go throughthis—fairly involved—chain of reasoning. In all cases,the mixing causes upwelling at the mixing locations;this upwelling must be fed by converging horizontalflow, ultimately by southward flow emanating from thedeep-sinking locations. But the upwelling and meridi-onal flow are also linked dynamically: Assuming a geo-strophic ocean interior, the planetary geostrophic vor-ticity equation


]wby 5 f (5)

]z

implies that any deep convergence (producing upwell-ing and hence vortex stretching) must be balanced bynorthward flow. This connection exists for upwelling inthe basin interior but not for upwelling at the side walls,where (5) would not be expected to be a good approx-imation (Stommel and Arons 1960; Spall 2000). Thefollowing considerations apply in a scenario that has thesame amount of water upwelling in the interior as wouldupwell at the same latitude but at a wall.

When mixing is highly localized in the ocean interior,the magnitude of ]w/]z is quite large, producing a strongrecirculation in both the abyss and in the thermocline,as shown in Figs. 8a and 8b, respectively [see Pedlosky(1996, 405–409) for an analytical treatment of a local-ized abyssal sink; see also Spall (2000)]. At both depths,the recirculation is associated with a warm anomalyextending westward from the mixing column. In thethermocline, vortex compression produces a recircula-tion of opposite sense as the flow at depth, consistentwith anomalous warmth above the level of no motion.The upper level recirculation is supported geostrophi-cally by the diffusive heat flux ‘‘trapped’’ between themixing column and the western wall. The westwardpropagation of the anomaly is due to the b effect, thedynamics of which are described in Stommel (1982) fora geothermally driven warm anomaly. Notice that theanomaly is not advected to the east wall, where it couldinstead support the MOC.

To support a given rate of upwelling at different lat-itudes, the intensity of the warm anomaly scales as f 2/b, from (5) and the requirement that the anomaly be inthermal wind balance with y at the upwelling location.As f 2/b increases monotonically with latitude, an in-creasingly strong anomaly would be required, whichhowever cannot be created. Thus, only weaker upwell-ing can be supported, and the difference between themidbasin and boundary mixing series plotted in Fig. 9grows sharply with mixing latitude.

3) EASTERN VERSUS WESTERN BOUNDARY MIXING

From thermal wind considerations alone one mightexpect diffusive warming on the eastern boundary tosupport a strong MOC, whereas it is not clear howwarming on the western boundary can support even aweak MOC. In reality, the dynamics are more compli-cated than suggested by this argument. If the low-lati-tude thermocline is heated diffusively, whether in theeast or west, a meridional temperature gradient existsat midlatitudes in concert with strong zonal flow andsubsequent downwelling on the eastern boundary. Thisdownwelling is the main mechanism that warms theeastern boundary, thus helping to provide the shear nec-essary for the MOC. Note that the high-latitude flowand temperature structure is similar in the thermocline

whether mixing is located in the west (Fig. 6b), east(Fig. 7b), or midbasin (Fig. 8b).

Instead, it is at lower latitudes where the distinctionbetween the eastern and western localized mixing runsis more apparent. Notice, first, that the MOC is confinedto the north of the mixing latitude. When mixing occurson the eastern wall, the deep western boundary currentturns and flows eastward across the basin (Fig. 7a) atthe latitude of the mixing column. The low-level con-vergence into the eastern boundary causes upwellingfrom the abyss and divergence in the thermocline, pro-ducing opposite (westward) flow across the basin (Fig.7b). In order to support this flow geostrophically, it mustbe colder in the Tropics, as this depth is above the levelof no motion. It must remain warm northward of themixing latitude, or else geostrophic eastward flow wouldoccur; note the large tongue of water with temperaturegreater than 38C.

In contrast, when mixing occurs only at the westernboundary, significant zonal flow does not occur at lowlatitudes (Fig. 6b). At this depth the temperature in theTropics is a full degree higher than with localized east-ern mixing, yielding the result that the tropical ther-mocline is actually deeper in the run with smaller over-turning. This may seem surprising, as it has long beenassumed that the meridional overturning scales as thethermocline depth, but the caveat here is that the over-turning does not extend to these tropical latitudes.

Building on these differences, we now address whythe aforementioned disparity in the eastern and westernmixing series increases with mixing latitude (Fig. 9).As the mixing column moves northward and f increases,weaker flow is in thermal wind balance with a similardensity gradient. With mixing on the western boundary,it becomes increasingly difficult for geostrophic currentsto advect the diffusive warming over to the easternboundary at midlatitudes. Conversely, as mixing on theeastern boundary moves northward it approaches thesite of warm water injection, forming a cohesive warmanomaly that supports the dominant circulation—an an-ticyclonic gyre above a deep cyclonic gyre—even asthe mixing column moves quite far to the north.

Finally, we return to our comparison of the boundarymixing control run with the uniform mixing run. Aswith localized interior mixing, uniform mixing leads tohorizontal recirculation (as evidenced by stronger west-ern boundary currents) supported by a warm anomaly(or, equivalently, less anomalous cooling) on the westernside of the basin interior, consistent with the modestreduction in the overturning maximum.

The wavelike pattern in Fig. 6b along the westernboundary to the north of the mixing latitude is, to alesser extent, present to the south of the mixing latitudewhen the mixing is located on the east (as shown in Fig.7b). Since this behavior does not ostensibly affect theconclusions presented here, we did not investigate fur-ther whether we were, in fact, observing stationaryRossby waves or spurious ‘‘wiggles’’ related to contrast


FIG. 10. Meridional overturning streamfunction (contours) and zon-ally averaged temperature (shading) for (a) weak deep mixing (exptM) and (b) strong deep mixing (expt N).

between the intense local and weak background diffu-sivities.

The maximum in meridional heat flux also decreasessharply as the mixing column moves northward (notshown), as might be expected given the decrease inoverturning strength. However, there is less disparity inthe three series as compared to that shown in Fig. 9.Although the overturning is weaker with midbasin mix-ing, a small ‘‘eddy’’ contribution to the meridional heatflux (i.e., due to deviations from the zonal mean) ispositive at the latitude of maximum flux, whereas thecontribution is negative for both other series, particu-larly with mixing located on the eastern boundary. Wealso note that the latitude of maximum meridional heatflux, approximately 208N for our standard boundarymixing run, moves to the north if the mixing column ismoved northward of this location. The reader is referredto Samelson (1998) for a more in-depth discussion ofthe effect of localized mixing on the meridional heattransport.

4. Depth-dependent mixing

a. Numerical results

The preceding section investigated the response ofthe MOC to the horizontal location of mixing. Now weturn to the dependence on where in depth mixing occurs.We performed four additional runs, with vertical profilesof mixing given in Fig. 1: a weak thermocline mixingcase (expt L), with our standard boundary mixing below1000 m and exponentially decaying boundary diffusiv-ity toward the surface; a weak deep mixing case (exptM), retaining the standard boundary mixing in the top1000 m but exponentially decreasing the boundary dif-fusivity below; a strong deep mixing case (expt N)where the boundary mixing is exponentially increasedbelow 1000 m; and a second strong deep mixing sce-nario (expt O), with a sharp increase below the ther-mocline but nearly constant diffusivity in the abyss.Note that our choice of 1000 m for the thermocline depthwas determined a posteriori, based on results from thecontrol experiment.

When mixing is decreased in the thermocline (exptL), the thermocline depth decreases and the maximumin overturning strength is reduced to 7.0 Sv (not shown).The model’s meridional heat flux is also considerablyweaker. The ramifications of this result, particularly inthe context of thermodynamic considerations, are con-sidered in our discussion section.

The overturning streamfunction for the weak deepmixing case (expt M) is shown in Fig. 10a. The max-imum in overturning strength decreases by only 0.2 Svas compared with the control run (Fig. 3a), and thereis no apparent change in the zonally averaged ther-mocline. However, upwelling at depth along the equatoris much weaker so that upwelling through the abyss is

more evenly distributed meridionally rather than con-centrated where mixing (albeit weak) is parameterized.

In contrast, the exponentially increasing strong deepmixing run (expt N) gives rise to vigorous upwelling atthe equator (Fig. 10b), producing a deep secondary max-imum in overturning. Approximately 3 Sv upwells nearthe equator at depth but subsequently downwells in thesubtropics. In as much as the equatorial upwellingthrough the thermocline here is similar to that in thecontrol and weak deep mixing runs, the overturningmaximum is increased by less than 1 Sv. The maximummeridional heat transport in the three runs varies by lessthan 0.01 PW, again suggesting that any changes in thedeep circulation do not affect the flow through the sur-face layer or thermocline.

Our experiment N is similar to those discussed inCHG, although their parameterization of vertical dif-fusivity as a function of N21 also implies increased mix-ing with depth within the thermocline. Cummins (1991)performed several experiments with varied mixing atdepth, although his profiles of vertical diffusivity ex-hibited a rather sharp increase at the base of the ther-mocline, in contrast to our slow exponential increase inour experiment N. In both CHG and Cummins, the in-crease in deep abyssal diffusivity was an order of mag-


FIG. 11. (a) Vertical potential temperature structure in low latitudesalong the boundary, as exemplified by this plot of T(z) 2 T(2H) anddT/dz at 08E, 08N (from expt A). (b) Vertical velocity adjacent to theequatorial boundary, as shown for the westmost and eastmost gridpoints.

nitude less than that here. Despite these differences inthe implementation of deep mixing, all results concurthat the meridional heat flux is not sensitive to deepmixing. In contrast with our study, however, the CHGand Cummins results suggest that the maximum in over-turning streamfunction can, in fact, be quite sensitiveto deep mixing. To examine this disparity, experimentO has a more sharply increased diffusivity near the baseof the thermocline (i.e., similar to the Cummins profile,although our increase was somewhat less steep than thatindicated by his Fig. 1). In experiment O, our model’sMOC exhibited a more significant increase than in ex-periment N, again with only a minor change in the me-ridional heat flux. Nevertheless, our maximum in over-turning is less sensitive to deep mixing than in CHGand Cummins. Note that our secondary meridional cellis much stronger than that shown in CHG’s Fig. 4a,which we attribute to our boundary mixing implemen-tation. With strong equatorial mixing the secondary cellis quite distant from the high-latitude maximum in over-turning, allowing for less superposition of the deep cir-culation and the large-scale buoyancy-driven overturn-ing. It is also possible that their use of a Cartesian mix-ing scheme contributes to their sensitivity. For example,the maximum of the MOC is deeper in M97’s Cartesianboundary mixing run than in our isopycnal mixing run.With a deeper maximum, more superposition with thissecondary deep circulation is possible.

b. Abyssal heat balance

That deep diffusivity plays so little role in setting theoverall strength of the MOC is surprising, given theimportance that has been attributed to abyssal mixing(Munk 1966; Polzin et al. 1997; MW). To understandhow the steady-state dynamics are affected by deep mix-ing, it is useful to more carefully examine the model’sdeep stratification and abyssal heat budget. Figure 11ashows a comparison of T(z) 2 T(2H) with ]T/]z forour ‘‘control’’ boundary mixing run. Both quantities areshown as measured at the western equatorial boundary(the behavior is qualitatively similar throughout theTropics). It is readily shown that in the Tropics, whereupwelling is ‘‘induced’’ by prescribed mixing, the mod-el’s stratification and w are consistent with one-dimen-sional advective–diffusive balance (see also Samelson1998), with the surface and bottom temperatures asboundary conditions; in fact, the resulting solution isnearly an exact match with the model’s temperature pro-file. Thus, horizontal advection (i.e., due to sloping is-opycnals) does not play any significant role in settingstratification in the tropics, except at the bottom bound-ary where w(z) vanishes.

The abyssal heat budget as illustrated in Fig. 12.Downwelling water in the northeast corner is relativelybuoyant (Marotzke and Scott 1999), producing a warmanomaly with associated cyclonic flow in the deepocean. Some of the flow immediately turns and upwells

along the eastern boundary; as illustrated in Fig. 2b, thestrongest upwelling occurs adjacent to downwelling,tending to cool the eastern boundary higher in the watercolumn (note the cold anomaly between 308–408N onthe eastern boundary in Figs. 6b, 7b, and 8b). Most ofthe flow however continues westward across the basin,passing near deep convection. In these model runs deepconvection reaches the bottom in the northwestern cor-ner, which is relatively stagnant (and therefore cold)because the upper western boundary current separatesfrom the ‘‘coast’’ between 408N and 508N. Deep con-vection is not a source of deep mass flux, so there isno divergence of flow to spread the water mass prop-erties of the convectively mixed column. Because the


FIG. 12. Illustration (plan view) of abyssal flow and heat budget.Flow downwells (as indicated by the circled ‘‘3’’) in the northeastcorner at temperature TD and its water mass properties are subse-quently modified as it flows adjacent to deep convection in the north-west. Flow then proceeds along the western boundary; some flowupwells (as indicated by the circled ‘‘●’’) in the western boundary,some flow reaches the equator and upwells, and some flow movesacross the basin before upwelling in the east. Downward diffusionof heat along the boundaries is balanced by upwelling so that theflow is warmed slightly in low latitudes, upwelling at average tem-perature TU.

site of deep convection is anomalously cold, the geo-strophic flow is around rather than through this region,so horizontal advection is also ineffective at conveyingits water mass properties. Rather, this cold water isspread through the deep ocean by mesoscale eddy trans-fers, here represented by the Gent and McWilliams(1990) parameterization. One way to quantify this effectis via a ‘‘bolus’’ downwelling at the deep convectionsite and a bolus upwelling in the warmer sectors of thedeep ocean. Thus, both deep downwelling and deep con-vection play a role in determining abyssal water massproperties (see also Marotzke and Scott 1999; Huck etal. 1999).

The relatively minor role played by abyssal mixingin setting MOC strength through the thermocline standsin marked contrast to other recent discussions (e.g.,MW; Ledwell et al. 2000), so a careful analysis of theabyssal heat budget is warranted. To reconcile the abys-sal heat budget, consider the signs of the relevant fluxterms. Let us first consider the mixing processes. Insteady state, convective mixing causes a heat loss, anddiffusion produces a heat gain. In our control boundarymixing run, however, the contribution from both of thesesources is small. Given that deep flow is around ratherthan through the site of deep convection, scant heat isconvectively mixed out of the abyss. The heat gain fromdiffusion is also small due to the model’s weak strati-fication in the abyss.

The vertical advective heat gain in the abyss is pro-portional to the following expression:

wT dA 5 w T 1 w T , (6)E E D D E U U

downwelling upwelling

where the U and D subscripts refer to upwelling anddownwelling, respectively. This term seemingly couldbe positive or negative, but we suggest it must yield aheat gain (i.e., TD . TU, and by continuity 2wD 5 wU),unless there are other deep diabatic sources/sinks suchas geothermal heating (Scott et al. 2001). The advectiveheat gain is largely balanced by the remaining term inthe budget, namely a heat loss from the parameterizationof mesoscale eddies along isopycnals, which leads to abolus heat transport opposite that of the advective trans-port.

For a more quantitative justification of these results,we made some rough calculations of the heat fluxes intothe upper and lower aybss. In the lower abyss (bottom1000 m), the diffusive heat flux rcpk]zT is of order 0.1TW, whereas the larger balance is between heating viaadvection and cooling through bolus transport, withterms on the order of 1 TW. Thus, despite the verticalbalance in the Tropics (i.e., where mixing occurs) beingadvective–diffusive, the diffusive heat flux is a lowerorder term in the full balance. In the upper abyss, thestratification increases by an order of magnitude (andtherefore the diffusive heat transport increases), but boththe advective heating and bolus transport cooling alsoincrease, due to both an increased magnitude of flowand an increased disparity between TU and TD. In thethermocline, the balance among these heat fluxes is quitedifferent, particularly the top 500 m. Here, the diffusiveflux is quite large, approximately 0.5 PW (i.e., the mag-nitude of the equator-to-pole heat flux), whereas bothadvection and bolus transport contribute a upward flux,but less than 0.1 PW. Presumably, the balancing upwardheat flux is by convection in (deep) mixed layers.

The insensitivity of the overturning maximum to deepmixing, as would be suggested by our numerical results,is consistent with our observation that diffusive heatingis not a dominant term in the abyssal heat budget. Onthe other hand, the differences in flow through the abyssas shown in Figs. 10a and 10b imply that the strengthof deep mixing affects the density structure of the deepocean. Figures 13a and 14a show the temperature andstratification (]T/]z) at the west and east sides of theequator, respectively, for the depth-dependent mixingruns. With strong deep mixing, the bottom water isslightly warmer in low latitudes, that is, the tail end ofthe abyssal flow pathway, consistent with a larger dif-fusive heat flux to the bottom. Similarly, weaker dif-fusive fluxes leads to colder bottom water at low lati-tudes. The temperature at the deep convection site,which is linked to the coldest surface temperature, islittle changed in both cases (not shown).

Surprisingly, the low-latitude abyss is less stratifiedin both the weak and strong mixing cases. With weakdeep mixing, less heat is diffused downward, whereas


FIG. 13. (a) Vertical potential temperature structure and (b) verticalvelocity at the western extreme of the equator for the control boundarymixing (expt A), weak deep mixing (expt M), and strong deep mixingrun (expt N).

FIG. 14. (a) Vertical potential temperature structure and (b) verticalvelocity at the eastern extreme of the equator for the control boundarymixing (expt A), weak deep mixing (expt M), and strong deep mixingrun (expt N).

in the strong mixing case diffusion is so efficient atmixing heat that the temperature gradient is degraded.This latter result is in contrast with CHG and Cummins(1991), where stronger stratification with increased deepmixing was observed. However, when we scaled backthe increase in deep mixing so that the area-weighteddiffusivity at each vertical level was more similar to thatin Cummins (expt O), we too observed increased strat-ification at depth (not shown). More specifically, thedeep ocean stratification doubled, although stratificationin the 1000 m below the thermocline was weaker. Wespeculate that some ‘‘optimum’’ profile of k could leadto a maximum stratification at depth, although furtherresearch along these lines is beyond the scope of thispaper.

As suggested by the plots of meridional overturningstreamfunction, upwelling at the equator varies consid-erably between deep mixing runs (Figs. 13b and 14b).A weaker diffusive flux requires less upwelling forsteady-state balance, and therefore it is no longer nec-essary for such a large percentage of abyssal upwellingto occur at the equator. Conversely, in the strong deepmixing case the larger equatorial diffusive heat fluxmust be balanced by strong upwelling. Without suffi-cient mixing in the thermocline, however, this upwellingessentially ‘‘detrains’’ from the larger cell, flowing hor-izontally and downward away from the equator.

Our analysis of the abyssal heat budget also suggestsan explanation for the differences in the behavior on theeast and west (cf. Fig. 13 vs Fig. 14). Because the abys-


FIG. 15. Potential temperature difference between the eastern andwestern boundary for (a) weak deep mixing (expt M) and (b) strongdeep mixing run (expt N). The difference is negative in shaded areas.

sal flow reaches the western boundary soon after down-welling, diffusion has had less time to alter the strati-fication. Thus, the stratification in the west is less af-fected by the magnitude of deep mixing. In contrast, inthe east the ocean is virtually unstratified in the bottom2000 m in the weak mixing run, and stratification alsofalls off more sharply (as compared with the controlrun) in the strong mixing case.

Differences in the upwelling profile are consistentwith thermal wind balance of the zonally averaged deepoverturning circulation. More specifically, note that thewestern boundary upwelling in the weak deep mixingrun is sharply reduced. As a result, some heat penetratesdownward on the west (mixing may be weak, but is stillnonzero), so the nearly unstratified eastern boundary iscolder than the west throughout much of the abyss (Fig.15a). The resulting east–west temperature difference issuch that zonally averaged southward flow increasesfrom the bottom upward, consistent with the overturningpattern observed in Fig. 10a. In the strong deep mixingrun, upwelling on the eastern boundary peaks higher inthe water column as compared to the west, which inturn produces a dipole pattern east–west temperaturedifference at depth (Fig. 15b). The warmer easternboundary near the bottom is necessary to support the

shear required for the deep equatorial overturning cell,but the east must also be colder near the base of thethermocline in order to attenuate the northward flowassociated with the top of this cell.

5. Discussion

We have presented a series of numerical experimentsthat explore the large-scale consequences of mixing lo-cation. Our single-hemisphere model is highly idealized,lacking wind forcing and topography, although we sub-mit that our results provide context for speculation aboutthe dynamics of the real ocean.

Given the different processes thought to play a rolein the steady state balance of the MOC—convection,rotation, and buoyancy forcing—it is not directly ap-parent why the strength of the MOC is a function ofthe magnitude and distribution of diapycnal diffusivity.According to ‘‘Sandstrom’s theorem’’ (Sandstrom1908), given surface heating at a higher geopotentialthan cooling (neglecting the smaller geothermal heatfluxes at the ocean floor), the steady-state ocean cir-culation should, for all intents and purposes, be mo-tionless except in a thin upper layer. In other words,given heating in the Tropics, the ocean should not beable to operate as a heat engine, extracting energy fromthe surface buoyancy forcing to maintain a strong MOC.However, Jeffreys (1925) argued that turbulent mixingcould effectively lower the geopotential of heating, lead-ing to horizontal temperature gradients at depth whichwould in turn lead to a vigorous circulation [see Colinde Verdiere (1993), MW, and Huang (1999) for a morethorough discussion of Sandstrom’s theorem and thecontroversy surrounding its application to the ocean].

In the model, we find that diapycnal (nonconvective)mixing at mid and high latitudes is not critical in orderto generate a MOC, given that surface temperaturesthere are relatively low and hence diffusive heat fluxesare much weaker than in the Tropics. Mixing at lowlatitudes, where the surface temperature is high, is moreefficient at diffusing heat beneath the mixed layer andhence more effective at driving the MOC. This resultsuggests that thermodynamic considerations of theocean circulation are indeed fundamental: The strengthof the MOC is a direct function of the surface heat inputthat diffuses into the thermocline. Diapycnal mixing inthe tropical thermocline communicates the surfacebuoyancy fluxes into the interior, which in turn increasespotential energy. In conjunction with convective mixingat high latitudes, the penetration of heat leads to hori-zontal temperature gradients beneath the surface. In geo-strophic balance, these temperature gradients generatestrong zonal flows into the eastern boundary that sub-sequently downwell, leading to an east–west tempera-ture difference that provides the vertical shear necessaryfor the MOC (Zhang et al. 1992; Colin de Verdiere 1993;Marotzke 1997).

Comparison of boundary mixing with interior mixing


indicates that planetary geostrophic vorticity balance inthe interior, the cornerstone of the Stommel and Arons(1960) theory for the abyssal component of the large-scale overturning circulation, actually hinders the pro-cess leading to east–west density differences, which cansupport the MOC through thermal wind shear. As re-quired by conservation of planetary vorticity, vortexstretching or compression in the ocean interior is ac-companied by meridional flow. Here, we find that vortexcompression in the thermocline effectively restricts thecommunication of warm waters to the eastern boundary,and therefore interior mixing is less effective at drivinga vigorous MOC than boundary mixing. With boundarymixing, where frictional effects enter the vorticity bal-ance, heat penetration into the thermocline leads tostronger zonal flows that downwell to greater depth atthe eastern boundary.

Our model results suggest that deep downwelling wa-ters are relatively buoyant, and to lowest order the abys-sal heat balance is between advective transport (pro-ducing a heat gain) and cooling through the effect ofmesoscale eddies along isopycnals, here represented bythe Gent and McWilliams (1990) parameterization.Thus, the abyssal water mass properties are an averageof the properties tied to deep convection and that ofdeep mass injection. The magnitude of deep mixing canaffect the bottom water temperature, however. With veryweak mixing, the abyss is colder and nearly homoge-neous, and flow upwells quasi-adiabatically into thethermocline. Stronger mixing produces a warmer abyssand dictates that upwelling in the abyss occurs diabat-ically, where mixing is located. An interesting corollaryis that the temperature at the base of the thermoclineappears to be set by the temperature of the downwellingwater (see Fig. 2b).

Observed stratification of the real abyss might suggestthe presence of enhanced deep mixing there. Dissipationof tides is thought to produce elevated mixing at depthnear rough bottom topography (Polzin et al. 1997), andalso several 100 m above, presumably through upwardinternal wave propagation. An idea expressed in MWis that diapycnal mixing in the abyss, resulting from theenergy input of winds and tides, is fundamentally nec-essary to return deep waters back into the thermocline.Our results suggest that this enhanced deep mixing byitself is insufficient to support a strong MOC and heattransport, and has little effect on the strength of thecirculation through the thermocline; rather, elevatedmixing must be found at thermocline depths. A corollaryof this result is that abyssal mixing is not necessary fora vigorous deep circulation (i.e., mixing in the ther-mocline is sufficient). Recent microstructure measure-ments off Cape Hatteras indicate strong mixing at ther-mocline depths above rough bottom topography, al-though of very limited spatial extent and primarily con-fined to depths below 500 m (K. Polzin 2002, personalcommunication). Similar measurements in the GulfStream, above gently sloping terrain, suggest only

slightly elevated levels of mixing. Only very recentlyhave significant areas of elevated mixing in the ther-mocline been located, in the salt-fingering area of thewestern subtropical North Atlantic (R. Schmitt 2002,personal communication).

In the real ocean, mixing at depth may play an ad-ditional role, which is not addressed here, namely itscapacity to homogenize water masses of different origin(e.g., North Atlantic Deep Water and Antarctic BottomWater). Our results show that the deep circulation pat-tern induced by deep mixing (or lack thereof ) is con-fined below the thermocline and does not transport anysignificant amount of heat, and therefore does not playa significant role in determining the oceanic meridionalheat transport. Although the bottom-water temperatureis affected by the deep mixing, we suggest that this hasonly minor impact on the meridional heat flux, in con-trast with the argument put forth in Cummins et al.(1991) to explain this insensitivity. The maximum inthe overturning streamfunction may be affected by deepmixing through superposition of the deep circulationwith the surface-forced overturning, depending on thevertical profile of mixing in the abyss.

In part due to the lack of observed thermocline mix-ing, other ideas regarding the importance of the South-ern Hemisphere in driving North Atlantic Deep Waterproduction have recently been gaining favor. One pos-sibility is that the winds over the Antarctic CircumpolarChannel (ACC) lead to enhanced mixing there, as dis-cussed in Wunsch (1998). Using a general circulationmodel of an idealized ocean basin, Marotzke and Klin-ger (2000) found increased cross-equatorial transportwith enhanced mixing in the Southern Hemisphere, con-sistent with our results that show upwelling occurswhere mixing is located. However, if mixing is con-centrated in the latitude band of the ACC, our modelresults suggest that this would not be effective in drivinga strong MOC, as the cool surface temperatures wouldlead to weak diffusive heat flux penetrating into thethermocline.

A second variation on the role of the Southern Oceanis that the wind stress over the ACC produces a ‘‘DrakePassage effect’’ whereby these winds induce a north-ward Ekman transport that can only return southwardgeostrophically below the sill of Drake Passage, hencerequiring its transformation into North Atlantic DeepWater (Toggweiler and Samuels 1995). The flow is sub-sequently returned to the surface though Ekman suctionin Drake Passage. In coarse-resolution general circu-lation models, a nearly linearly relation between north-ern deep-water production and Southern Hemispherewinds has been observed (McDermott 1996), perhapsobviating the need for any significant sources of mixingin order to produce a vigorous overturning circulation(Toggweiler and Samuels 1998). Again, we caution thedirect applicability of our idealized model to the realocean, in as much as we do not have any wind-inducedupwelling, nor does our model do justice when it comes


to representing the complexity of the oceans’ water massproperties. Although model results show that the pres-ence of the channel can lead to increased thermoclinedepth (e.g., Gill and Bryan 1971; Vallis 2000), the ther-modynamic considerations we have examined suggestthat it must be investigated how this mechanism leadsto the east–west temperature differences necessary tosupport a pole-to-pole overturning cell in thermal windbalance.

Our standard boundary mixing is parameterized invertical columns many kilometers wide, which providesan equal area of mixing at all depths. In the real ocean,the boundaries are more horizontal than vertical, andmixing likely occurs over a significantly reduced lengthscale normal to the boundary. Moreover, the effect ofdiapycnal mixing on sloping boundary has been shownto have dynamical consequences (Garrett 1991, 2001;Thompson and Johnson 1996), which may influence thelarge-scale circulation. The presence of sloping bound-aries at high latitudes affects the volume of deep masstransport by requiring deep convection to occur in theopen ocean (Spall and Pickart 2001), which may alterour depiction of the abyssal heat balance.

6. Conclusions

The main conclusions from our idealized model ex-periments are as follows.

1) Boundary mixing is more efficient than interior mix-ing in causing a strong MOC.

2) The MOC strength through the thermocline, and theassociated heat transport, are mainly determined bythermocline mixing at low latitudes, where the ver-tical temperature gradient is strong. In contrast, high-latitude and deep mixing play lesser roles.

3) Mixing plays a minor role in the deep-ocean heatbudget.

Acknowledgments. The authors thank Carl Wunsch,Kerry Emanuel, Barry Klinger, Peter Stone, and HuaRu for discussions and comments on an earlier versionof this manuscript. The anonymous reviewers signifi-cantly contributed to improving the paper. JS was sup-ported by the MIT Joint Program on the Science andPolicy of Global Change and by the U.S. Departmentof Energy’s Office of Biological and Environmental Re-search Grant DE-FG02-93ER61677. JM was supportedby NSF Grant 9810800.

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The Location of Diapycnal Mixing and the Meridional Overturning Circulation

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