University of Colorado
Boulder, Colorado USA
Historically, the interannual anomalies associated with ENSO have been considered independently from the mean state and the annual cycle, although it was recognized that there was some relationship between the phase of ENSO and the phase of the annual cycle (cf. Rasmusson and Carpenter, 1982). While it was recognized that latent heat transport is important in the mean and seasonal climatology of the atmosphere over the Pacific, anomalies of diabatic heating were considered to be controlled more-or-less directly by SST anomalies. (The model of Zebiak and Cane  showed that moisture convergence feedback results in a somewhat more complex relationship with maximum heating of the atmosphere possibly displaced from the maximum SST anomaly.) Most important was the assumption that SST anomalies could be simply related to the wind anomalies, through anomalous horizontal advection and through grossly simplified models of vertical mixing.
The purpose of this paper is to show that a complete consideration of the hydrological cycle leads to the conclusion that the mean state of the western Pacific warm pool depends on the existence of variability (transients, intraseasonal oscillations, and ENSO), as suggested by Lukas (1990). The assumption of a simple parametric dependence of interannual variability on the mean and seasonally-varying climate will not work in the modeling of the year-to-year variations of the warm pool.
The cycling of water between liquid and vapor phases results in the transfer of heat from the oceans to the atmosphere. These phase changes result in highly nonlinear behavior of the coupled ocean-atmosphere system. This can easily be seen in the form of the bulk parameterization of evaporation:
where E is evaporation, Ce is a turbulent exchange coefficient, qo is the saturation specific humidity at the temperature of the sea surface, qA is the near-surface atmospheric specific humidity, and U is the magnitude of the near-surface wind. In general, Ce depends on the atmospheric stability, which depends on the flow, the air-sea temperature difference, and the lapse rate. In particular, at low wind speeds, Ce is strongly dependent on U (Liu et al., 1979; Fairall et al., 1995). There are both local and large-scale relationships between Ce, qo, and U, and evaporation is very obviously one of the strong sources of nonlinearity introduced by the hydrological cycle.
Evaporation causes a decrease of buoyancy of the surface waters by cooling and increased salinity, resulting in convection in the upper ocean. Precipitation provides buoyancy to the upper ocean, reducing vertical mixing (Lukas and Lindstrom, 1991) especially nocturnal convection (Anderson et al., 1995). This adds additional potential feedbacks into the coupled system (Figure 1).
Moisture is transported by winds and ocean currents, which are forced by horizontal buoyancy gradients. Diabatic heating is responsible for a major fraction of the buoyancy gradients, but the contributions from the distribution of water vapor in the atmosphere and salinity in the ocean are not negligible. Water vapor is buoyant relative to dry air, but most of the buoyancy flux from the ocean to the atmosphere is potential buoyancy associated with the latent heat flux. Thus the regions of free convection in the ocean and atmosphere can be separated by great distances. This is the situation in the Pacific Ocean, with significant convection in the central subtropical oceanic gyres of both hemispheres, and deep atmospheric convection over the warm pool (Figure 2).
In order to compensate the moisture convergence over the warm pool, and the net freshwater flux, the ocean circulation supports a salt convergence into the warm pool within the upper thermocline circulation. This is the oceanic counterpart to the Walker and Hadley circulations of the atmosphere. Subduction of the saline waters created by the trade winds occurs as these waters are carried by the wind-driven ocean circulation into the western equatorial Pacific (Figure 3 and Figure 4; see also McCreary and Lu  and Shinoda and Lukas ). The resulting strong mean vertical salinity gradient in the warm pool inhibits vertical mixing, concentrating the heat from insolation in the near-surface layer of the warm pool where it can be most easily extracted by the atmosphere through turbulent heat fluxes on the short time and space scales which were the focus of the TOGA Coupled Ocean-Atmosphere Response Experiment (COARE). An imbalance between the net freshwater flux into the warm pool and the convergence of salt below will result in a modification of vertical mixing which feeds back onto the heat budget of the warm pool.
Of the 3-5 m annual average rainfall over this region, roughly 3050% is imported, yielding a net freshwater flux of 12 myr¯ ¹ (Figure 3; Oberhuber, 1988). Thus, a significant recycling occurs over the warm pool, with the heat being supplied from the ocean. Results from COARE measurements indicate that the latent heat flux was between 90 and 150 W m¯ ² during the COARE IOP, with the lower end of the range during light winds typical of the warm pool (S. Anderson, personal communication). This heat was provided by the 160240 W m¯ ² of penetrating shortwave radiation. (Other flux terms reduced the net flux into the ocean to 1520 W m¯ ².) Thus, much of the diurnal heating of the upper ocean by penetrating solar radiation is released back to atmosphere in the form of latent heat. On what time scale does this occur? Because of nonlinearities in the coupling, it happens on a spectrum of time scales, including ENSO.
The largest modulation of incoming solar radiation is obviously the diurnal solar cycle, while the variability of evaporation is strongest on the time scale of the transients and intraseasonal oscillations noted above. This suggests an oceanic mechanism which is sequestering a fraction of the diurnal heating and making it available to the atmosphere on these longer time scales. This mechanism is the penetration of solar radiation below the depth of the nocturnal mixed layer, which in the warm pool is largely controlled by the freshwater flux (Anderson et al., 1995). Siegel et al. (1995) measured typical values of 23 W m¯ ² passing through 30 m depth. Until surface winds are strong enough to mix below this depth, this region of the water column will continue to warm, and observations of frequent temperature inversions near this depth are consistent with this mechanism. Further support for this rectification of the diurnal cycle in the upper ocean comes from the observation that during the earliest stage of the December 1992 westerly wind burst, entrainment of deeper waters into the mixed layer appeared to contribute to a heating of the mixed layer (more than offset by latent heat loss, however.)
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