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On the Coupling of Wind Stress and Sea Surface Temperature 15 APRIL 2006 SAMELSON ET AL. 1557 On the Coupling of Wind Stress and Sea Surface Temperature R. M. SAMELSON,E.D.SKYLLINGSTAD,D.B.CHELTON,S.K.ESBENSEN,L.W.O’NEILL, AND N. THUM College of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, Oregon (Manuscript received 25 January 2005, in final form 5 August 2005) ABSTRACT A simple quasi-equilibrium analytical model is used to explore hypotheses related to observed spatial correlations between sea surface temperatures and wind stress on horizontal scales of 50–500 km. It is argued that a plausible contributor to the observed correlations is the approximate linear relationship between the surface wind stress and stress boundary layer depth under conditions in which the stress boundary layer has come into approximate equilibrium with steady free-atmospheric forcing. Warmer sea surface temperature is associated with deeper boundary layers and stronger wind stress, while colder temperature is associated with shallower boundary layers and weaker wind stress. Two interpretations of a previous hypothesis involving the downward mixing of horizontal momentum are discussed, and it is argued that neither is appropriate for the warm-to-cold transition or quasi-equilibrium conditions, while one may be appropriate for the cold-to-warm transition. Solutions of a turbulent large-eddy simulation numerical ␥ model illustrate some of the processes represented in the analytical model. A dimensionless ratio ␶A is introduced to measure the relative influence of lateral momentum advection and local surface stress on the ␥ Ͻ boundary layer wind profile. It is argued that when ␶A 1, and under conditions in which the thermo- dynamically induced lateral pressure gradients are small, the boundary layer depth effect will dominate lateral advection and control the surface stress. 1. Introduction horizontal momentum from the faster flow above to the slower near-surface flow. Other possible mechanisms Satellite observations of sea surface temperature include flows driven by thermodynamically induced (SST) and wind stress have recently revealed systematic pressure gradients (e.g., Lindzen and Nigam 1987; Wai correlations between warm SST anomalies and en- and Stage 1989; Small et al. 2003). hanced wind stress on horizontal scales of 50–500 km in Here, a new hypothesis for this observed correlation both the Tropics and the extratropics (e.g., Liu et al. is advanced for the case of quasi-equilibrium condi- 2000; Chelton et al. 2001; Hashizume et al. 2001; Polito tions, and the viability of the downward-mixing mecha- et al. 2001; O’Neill et al. 2003; see also the reviews in nism is reconsidered. A simple analytical model of the Chelton et al. 2004; Xie 2004). One possible mechanism marine atmospheric boundary layer is used to illustrate that has been invoked in several studies (e.g., Sweet et the concepts. The main goal is to clarify the effect of al. 1981; Wallace et al. 1989; de Szoeke and Bretherton variations in the intensity of turbulent mixing within the 2004) as a possible source of these correlations involves boundary layer on near-surface winds under quasi- changes in the stability of the boundary layer, which are equilibrium conditions by considering a model that re- presumed to control the “downward mixing” of mo- tains a basic representation of the essential physics, but mentum from a low-level jet to the lower boundary is simple enough to yield explicit analytical solutions layer. In this scenario, it is argued that the enhanced whose dependence on parameters can be readily iden- winds over warm water are maintained by enhanced tified and interpreted. The discussion of numerical downward fluxes, induced by convective instability, of model results by Wai and Stage (1989) anticipates some aspects of the present hypothesis, but focuses on the pressure gradient mechanism and on variability with smaller horizontal scales. The present study does not Corresponding author address: R. M. Samelson, College of Oce- anic and Atmospheric Sciences, Oregon State University, 104 address the pressure gradient mechanism or the force COAS Admin. Bldg., Corvallis, OR 97331-5503. balances in the complex transition region in the imme- E-mail: [email protected] diate vicinity of SST fronts. © 2006 American Meteorological Society 1558 JOURNAL OF CLIMATE VOLUME 19 2. Quasi-equilibrium models a. Nonrotating integral momentum balance Consider the downstream momentum balance in a turbulent marine atmospheric boundary layer with flow in, say, the meridional direction. Define z ϭ h, the top of the boundary layer, as the level at which the (kine- matic) turbulent stress ␶ ϭϪϽ␷ЈwЈϾ vanishes, where ␷Ј and wЈ, respectively, are meridional and vertical fluc- tuating turbulent velocities. The depth h might be esti- mated from physical balances (e.g., by Monin– Obukhov theory in the stable case), but for the present purposes it is adequate to specify it directly. Assume a ϭ ␳Ϫ1 ␳ fixed meridional pressure gradient G 0 py, where 0 is a constant reference density, and (for simplicity; note also that this follows a standard assumption in fluid FIG. 1. Dimensionless magnitude U/Ug (thick solid line) and boundary layer theory) take G to be independent of zonal (solid) and meridional (dashed) components of horizontal depth in the boundary layer. Neglect rotation, time de- wind in the slab boundary layer model vs dimensionless ratio pendence, and horizontal variations, and consider only rs /(hf ) of rotational to frictional time scales. The dimensionless the vertical divergence of the vertical turbulent stress. pressure gradient is scaled to one and oriented meridionally (G ϭ 1). For fixed pressure gradient G, drag coefficient r , and Coriolis Numerical solutions discussed below (section 4c) ad- s parameter f, wind speed and stress rsU increase with increasing dress the general quasi-equilibrium case, with rotation, boundary layer depth h. stratification, nonlinearity, and time dependence. As discussed below, those solutions support the qualitative picture illustrated here through the consideration of ever, because this direction depends on the wind and simple analytical models. turbulent stress profiles as well as the pressure gradient, With these assumptions, the meridional momentum these profiles must be determined before a similar de- equation in the boundary layer reduces to a balance pendence of surface stress on boundary layer depth can between pressure gradient and stress divergence, be inferred. A standard, analytically accessible boundary layer ϭϪ ϩ ␶ր ͑ ͒ 0 G d dz. 2.1 model that provides these profiles and illustrates this Integrate vertically over the boundary layer and divide dependence in the presence of rotation is a slab mixed- by the boundary layer depth h to obtain the vertically layer model, in which the horizontal velocities are as- averaged momentum balance sumed independent of height in the boundary layer. With a linear surface stress parameterization, the cor- 0 ϭϪG Ϫ ␶ րh, ͑2.2͒ s responding momentum equations are where ␶ ϭ ␶(z ϭ 0) is the surface stress. The balance s Ϫf␷ ϭϪr uրh, ͑2.3͒ (2.2) is between the pressure gradient force and the s ϭϪ Ϫ ␷ր ͑ ͒ mean vertical divergence of the vertical turbulent fu G rs h. 2.4 stress. From this relation, it is immediately evident that, Here f is the Coriolis parameter, G and h are as before, ␶ under these assumptions, the surface stress s is inde- and rs is a linearized surface drag coefficient. The so- pendent of the internal profile of the turbulent stress, lution is that is, of the form of ␶(z) for z Ͼ 0. For fixed pressure ␶ ϭ ͑ ␷͒ ϭ ͑ 2 ϩ 2ր 2͒Ϫ1 ͑Ϫ Ϫ ր ͒ ͑ ͒ gradient force G, however, s increases linearly with the u u, f rs h G f, rs h , 2.5 boundary layer depth h, because the stress vanishes by with wind speed magnitude assumption at the top of the stress boundary layer, and ϭ ͑ 2 ϩ ␷2͒1ր2 ϭ ͑ 2 ϩ 2ր 2͒Ϫ1ր2 ͑ ͒ the stress divergence is distributed over a deeper layer. U u f rs h G. 2.6 For fixed f, r , and G, the magnitudes U and r U of the b. Rotational effects s s wind speed and the surface stress, respectively, both When rotation is included, an integral relation of the increase monotonically with h toward the geostrophic form (2.2) may still be derived for the downwind mo- limit (Fig. 1). If h changes so that the ratio of the rota- mentum balance, where the downwind direction is de- tional time scale 1/f to the frictional time scale h/rs var- fined in terms of the vertically integrated flow. How- ies between 0 and 2, the surface stress changes by a 15 APRIL 2006 SAMELSON ET AL. 1559 factor of 2. Thus, in this model, the presence of rotation does not qualitatively alter the strong dependence of surface stress on boundary layer depth as suggested by the nonrotating balance (2.2). A second, standard boundary layer model that in- cludes rotation is the classical Ekman layer (e.g., Ped- losky 1987), for which a similar proportionality be- tween stress and boundary layer depth also arises for fixed pressure gradient. In that case, the surface stress is ␶ ϭ ␦ ͑ ␪ ␪ ͒ ͑ ͒ s G cos s, sin s , 2.7 where ␦ is an Ekman depth, proportional to the square FIG. 2. Schematic of the model marine atmospheric boundary ␪ ϭ ␪ Ϯ ␲ ϭ root of the eddy diffusivity, and s G /4 is the layer depth near an SST front. The depth z h is the level where hemispherically dependent direction of the surface the stress vanishes, which in the figure is deeper on the warm side ␪ of the SST front than on the cold side. The horizontal and vertical stress, with G now the magnitude and G the direction Ϯ␲ scales are several hundred kilometers and several hundred of the pressure gradient.
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