On the Variability of Winds, Sea Surface Temperature, and Surface Layer Heat Content in the Western Equatorial Pacific

Michael J. McPhaden and Stanley P. Hayes

NOAA/Pacific Marine Environmental Laboratory, Seattle, Washington

Journal of Geophysical Research, 96, supplement, 3331-3342 (1991)
This paper is not subject to U.S. copyright. Published in 1991 by the American Geophysical Union.

7. Summary and Discussion

The foregoing analysis has focused on a diagnosis of the surface layer heat balance using equatorial mooring data along 165°E during 1986 and 1987. Several processes related to wind forcing were examined, using statistical relationships between winds, temperature, and currents. We found that among the processes considered, evaporative cooling was the most important, accounting for heat content changes of typically 35 W m in the upper 75 m. Moreover, these evaporative heat fluxes converged nonuniformly in the surface layer, giving rise to temperature variations of about 0.1°C in the upper 10 m, but only about 0.02°C between 10 and 75 m. Entrainment from the thermocline due to wind work was an order of magnitude smaller. In addition, the affects of lateral advection were not significant, in spite of large (50-100 cm s) velocity fluctuations in the surface layer. This is because the horizontal temperature gradients were weak near 0°, 165°E during the time period of our analysis.

The inferred nonuniform convergence of evaporative heat flux in the surface layer, the weak temperature variations at 50 m relative to those at the surface (for example, Figure 7), and the weakness of wind work entrainment from the thermocline indicate that some process or combination of processes significantly limits the vertical extent to which heat can be mixed in the western Pacific warm pool. One possibility is the formation of a salt-stratified barrier layer, as suggested by Lukas and Lindstrom [this issue] and Godfrey and Lindstrom [1989]. Although we do not have salinity time series from the period of our analysis to test this hypothesis quantitatively, CTD data along 165°E show the existence of shallow, 0- to 10-m thick lenses of low-salinity water, with surface values up to 1 psu (practical salinity unit) lower than normal. These may be related to enhanced precipitation associated with an eastward displacement of deep convection across 165°E during the 1986-1987 ENSO. It has yet to be determined, however, whether the haloclines associated with these buoyant lenses are sufficiently strong to trap heat in a shallow layer near the surface.

Another possibility is that formation of the diurnal thermocline may be involved in limiting the depth penetration of heat on longer time scales. Daytime radiative heating leads to typical diurnal temperature variations of O(0.1°C) in the upper 10 m of the western equatorial Pacific [Taft and McPhaden, 1988]. The amplitude of the diurnal cycle is related to wind speed in qualitatively the same manner in which daily averaged SST is related to wind speed, that is, the diurnal cycle is weak when wind speed is high (>6 m s) and strong (>1°C) during periods of calm winds. Wind speed was greater than 6 m s for 22% of the time during the period from December 1986 to October 1987; however, most of the time the winds were not strong enough to erode the diurnal thermocline at the mooring site. One-dimensional mixed layer models indicate that inclusion of the diurnal heating cycle quenches convection during the day, which leads to shallower mixed layer depths than would be otherwise expected without the diurnal cycle because of nonlinearities in the turbulent energy balance [Woods and Barkmann, 1986]. Moreover, in the eastern equatorial Pacific Imawaki et al. [1988] have shown that the diurnal cycle modulates the turbulent heat flux out of the surface layer on longer time scales. Most of the turbulence and mixing there occurs at night in association with free convection [Peters et al., 1988]. Hence if formation of the diurnal thermocline were inhibited, nighttime convection could start sooner, last longer, be more intense, and penetrate to greater depths, resulting in a larger net downward heat flux. Conversely, if the winds were weak during the day and the diurnal thermocline strong, nighttime convection could be delayed, of shorter duration, weaker, and less vertically penetrative.

Lukas and Lindstrom [this issue] have suggested that repeated westerly wind burst forcing associated with ENSO may lead to long-period, interannual cooling of the western Pacific warm pool through wind work-generated entrainment mixing with the thermocline. As noted above, entrainment does not appear to be important in general along 165°E during 1987 at the height of the ENSO. Moreover, visually and by least squares analysis, there is no significant SST cooling trend over the record length shown in Figure 7, even though several 5-10 m s westerly wind events occur during this time period. In addition, the mean SST for December 1986 to October 1987 was not significantly colder than climatology (Figure 1); specifically, SST was 29.2°C (plus or minus one standard error of 0.1°C) at 0°, 165°E compared to a climatological mean over the same time period of 29.0°C [Reynolds, 1988]. Thus while we have shown that short-period westerly wind fluctuations lead to local SST variability on the same time scale, there is no evidence from our data that they lead to local low-frequency SST cooling associated with ENSO. Whether this result applies to other locations in the western equatorial Pacific as well is not clear.

Our conclusions with regard to advection and evaporation rely primarily on inference, because we do not have specific humidity to calculate latent heat flux, nor do we have accurate enough information on horizontal temperature gradients to estimate advection explicitly. In addition, our estimate of 5 m as a typical entrainment heat flux based on wind work is only approximate, because we do not have the data to incorporate into our estimate the affects of other sources and sinks of turbulent energy (for example, buoyancy forcing due to insolation and precipitation; vertical shear instability, etc.). Because of these and other data limitations, we have been able to account for only about 10-25% of the observed SST and heat content variance in terms of evaporative cooling, based on correlations with wind speed of 0.3-0.5 (for example, Figures 8 and 13). Shortwave radiation is likely to account for a large percentage of the unexplained temperature variance in the surface layer. Insolation at the equator under clear-sky conditions ranges between about 300 and 350 W m. Using Reed's [1977] flux formula, changes in cloudiness of only 10% would lead to changes in the net shortwave radiation at the sea surface of about 20 W m. Changes in cloudiness of this magnitude are likely to occur on daily to monthly time scales, given that normally the skies are 60-70% cloud-covered owing to very active deep convection in the western equatorial Pacific.

In conclusion, we note that the SST changes observed in our mooring data are only O(0.1°C). This is an order of magnitude smaller than those seen in the eastern equatorial Pacific, where O(1°C) fluctuations on daily to interannual time scales are common [e.g., McPhaden and Hayes, 1990]. It is encouraging that in spite of these small temperature signals and our limited data set, we have been able to draw inferences about physical processes at work in the surface layer heat balance. We emphasize that these inferences are specific to the time period and location of this study. Entrainment from the thermocline, for example, could be more important during periods of strong upwelling favorable easterly winds. Likewise, lateral advection could be important during periods when horizontal SST gradients are pronounced, or on longer time scales (e.g. interannual) than those considered in this study. Our analysis emphasizes the need, therefore, for longer and more complete data sets in future investigations of air-sea interaction in the western equatorial Pacific.

Acknowledgments. We would like to thank Stuart Godfrey of CSIRO, Hobart, Australia, for helpful comments on an earlier version of this manuscript. We would also like to acknowledge M. McCarty and H.P. Freitag for their assistance in carrying out the analyses. This work was supported by the U.S. Tropical Ocean-Global Atmosphere (TOGA) Project Office. Contribution 1166 from the Pacific Marine Environmental Laboratory, National Oceanic and Atmospheric Administration.

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