U.S. Dept. of Commerce / NOAA / OAR / PMEL / Publications

The Tropical Ocean-Global Atmosphere observing system: A decade of progress

Michael J. McPhaden,1 Antonio J. Busalacchi,2 Robert Cheney,3 Jean-René Donguy,4 Kenneth S. Gage,5 David Halpern,6 Ming Ji,7 Paul Julian,8 Gary Meyers,9 Gary T. Mitchum,10 Pearn P. Niiler,11 Joel Picaut,12,13 Richard W. Reynolds,7 Neville Smith,14 and Kensuke Takeuchi15

1Pacific Marine Environmental Laboratory, NOAA, Seattle, Washington
2NASA Goddard Space Flight Center, Greenbelt, Maryland
3National Ocean Service, NOAA, Silver Spring, Maryland
4Institut Français de Recherche Scientifique pour le Développement en Coopération, Plouzane, France
5Aeronomy Laboratory, NOAA, Boulder, Colorado
6Jet Propulsion Laboratory, California Institute of Technology, Pasadena
7National Centers for Environmental Prediction, NOAA, Camp Springs, Maryland
8Suitland, Maryland
9Commonwealth Scientific and Industrial Research Organization, Tasmania, Australia
10Department of Marine Science, University of South Florida, Saint Petersburg
11Scripps Institution of Oceanography, La Jolla, California
12Institut Français de Recherche Scientifique pour le Développement on Coopération
13Now at NASA Goddard Space Flight Center, Greenbelt, Maryland
14Bureau of Meteorology Research Centre, Melbourne, Victoria, Australia
15Institute of Low Temperature Science, Hokkaido University, Sapporo, Japan

Journal of Geophysical Research, 103(C7), 14,169-14,240 (1998).
Copyright ©1998 by the American Geophysical Union. Further electronic distribution is not allowed.

.  Scientific Progress: Improved Description and Understanding

3.1  Long-Term Mean and Mean Seasonal Cycle

The long-term mean and mean seasonal cycle are crucial for understanding interannual variations in the coupled system. Background stratification, for example, affects the length scales, timescales, and phase speeds of planetary equatorial waves thought to be important in the ENSO cycle. Likewise, zonal asymmetries in the background state of the equatorial ocean due to mean trade wind forcing, e.g., the mean zonal slope of the equatorial thermocline and zonal SST gradient associated with it (shown schematically in Figure 1), establish conditions necessary for the growth of ENSO-related SST anomalies [e.g., Battisti and Hirst, 1989]. El Niño anomalies also tend to be phase locked to the seasonal cycle, with warmest El Niño SST anomalies often occurring in boreal winter in the equatorial cold tongue, when SST is seasonally at its coldest [Rasmusson and Carpenter, 1982]. Empirical and modeling studies have indicated that persistence and predictability of ENSO anomalies is seasonally modulated, being highest in boreal summer and winter and falling off through the boreal spring [Latif and Graham, 1992; Webster and Yang, 1992; Latif et al., 1994; Balmaseda et al., 1995]. Some theories also suggest that the mean seasonal cycle determines the basic periodicity and irregularity of the ENSO cycle via chaotic nonlinear self-interaction [e.g., Jin et al., 1994; Tziperman et al., 1994; Chang et al., 1995]. However, few, if any, coupled ocean general circulation models (GCMs) are capable of simulating both the mean seasonal cycle and interannual ENSO-like variability with equal degrees of veracity [Mechoso et al., 1995]. Finally, seasonal variations for some variables (e.g., SST in the eastern Pacific) are as large as, or larger than, ENSO-related interannual anomalies. Therefore, at minimum, one requires a clear definition of the climatological mean seasonal cycle for model validation and in order to accurately define interannual climate anomalies. Climatologies existed prior to TOGA, but in some cases, especially for subsurface oceanographic variables, they were of poor quality because of the sparsity of data on which they were based.

3.1.1  Long-term mean

Key features important in characterizing the coupled ocean-atmosphere system in the equatorial Pacific include the western Pacific warm pool with SSTs > 28°C and the equatorial cold tongue of the eastern and central equatorial Pacific (Figure 4). These structures, evident in all long-term mean SST climatologies, are modulated in intensity and areal coverage on seasonal, interannual, and decadal timescales. Understanding how these features relate to surface winds and subsurface ocean hydrodynamics is critical to understanding climate variability related to ENSO.

An example of the improved definition from the TOGA observing system of mean upper ocean temperature, surface dynamic height, and wind stress along the equator is shown in Figure 6. The mean temperature section, on the basis of all available TAO data between 2°N and 2°S, is similar to that presented by Kessler et al. [1996]. It shows the increase in SST from east to west, the warm pool of 28°C water in the upper 100 m of the western Pacific, the downward sloping thermocline in the upper 300 m, and the existence of a weakly stratified "thermostad" of 13°C water in the eastern Pacific [Stroup, 1969]. Situated in the middle of the highly stratified upper thermocline is the 20°C isotherm; for this reason this isotherm is often used as an index for the depth of the thermocline in the tropical Pacific. The mean surface dynamic height associated with the temperature field rises by 40 dynamic centimeters (dyn. cm) between 95°W and 170°E, after which it decreases slightly to the west. Zonal variations in dynamic height and thermocline depth along the equator are a response to steady easterly trade wind forcing in the eastern and central Pacific [McPhaden and Taft, 1988]; reversal of these gradients in the western Pacific is associated with local westerly winds [see also Wyrtki, 1984; Mangum et al., 1990; McPhaden et al., 1990a]. The zonal section in Figure 6 has many features in common with sections composited from different individual cruises prior to TOGA [e.g., Philander, 1973; Halpern, 1980] but is more representative of long-term mean conditions.

Figure 6: Zonal section of mean temperature averaged between 2°N and 2°S on the basis of available TAO time series data in 1980–1996. Also shown is the corresponding mean zonal wind stress (computed using a constant drag coefficient of 1.2 × 10-3) and dynamic height 0–500 dbar (computed using mean temperature/salinity relationships based on work by Levitus and Boyer [1994] and Levitus et al. [1994a]). Crosses indicate depths and longitudes where temperature data were available. An average at a particular location was computed only if a minimum of 2 years of data was available.

The mean thermal structure of the Pacific along quasi-meridionally oriented VOS XBT lines (Figure 7) also shows the downward slope of the thermocline toward the west in response to mean trade wind forcing. In addition, the meridional structure of ridges and troughs in the thermocline, which are related to major zonal currents [e.g., Donguy and Meyers, 1996a], is also clearly delineated. Evidence of trade-wind-driven equatorial upwelling (local minima in temperatures near the equator in the surface layer) is apparent in the central and eastern Pacific sections.

Figure 7: Mean temperature for the period 1985–1994 on four well-sampled XBT lines. Typically, 120 or more realizations of the quasi-synoptic temperature field were obtained during the decade for each section. The standard deviation of seasonal-to-interannual temperature variability during 1985–1994 from the Australian ocean thermal analysis system [Smith, 1995b] is indicated by shading. Westernmost section is at the top, easternmost at the bottom.

Methods to estimate the volume transport of the major equatorial currents from monthly, synoptic VOS XBT sections, as in Figure 7, were developed by Kessler and Taft [1987], Taft and Kessler [1991], Picaut and Tournier [1991], and Donguy and Meyers [1996a]. A comparison of transports from VOS XBT data to research vessel data (Table 5) shows that all of the geostrophic current transports can be reasonably well monitored by the VOS program. Differences between means based on research vessel and VOS data are of the order of only 7–20% (Tables 5a and 5b). The temporal variation inferred from research cruise data is highly correlated to the VOS estimates [Picaut and Tournier, 1991]. Although somewhat different methods were used to calculate XBT transports by Kessler and Taft [1987] and Picaut and Tournier [1991], the mean and standard deviation of transports over a 7-year period are only slightly different (Table 5c).

Table 5a. Mean Current Transports During the Hawaii-Tahiti Shuttle From March 1979 to June 1980

Table 5b. Mean Current Transports During the Line Islands Profiling Projects (LIPP) From March 1982 to June 1983

Table 5c. Mean Current Transports From January 1979 to June 1985

Drifter data allow for a definition of the surface circulation (combined Ekman and geostrophic components) across the entire basin, rather than just along prevailing shipping routes. The average velocity at 15-m depth from the drifter data for 1988–1994 (Figure 8) shows the persistent and well-documented surface current systems of the tropical Pacific: the North Equatorial Current (NEC), South Equatorial Current (SEC), North Equatorial Countercurrent (NECC), and a vestigial South Equatorial Countercurrent (SECC) (in the region 6°–10°S, 160°–176°E). The standard error of the velocity shows that the general circulation of the tropical Pacific is well defined everywhere, even to the extent that divergence and relative vorticity fields can be computed from this data with a high degree of confidence.

Figure 8: Mean surface layer (15 m) circulation in the tropical Pacific based on Surface Velocity Program drifter data for the period 1988–1994. The ellipse at the end of each vector is the 95% confidence interval.

Significant departures from the patterns that have been reported by ship drift charts, or from interpretation of the gradients of dynamic height as an index of the surface current, emerge from the drifter data. For example, dynamic height maps show that there should be a geostrophic flow toward the equator nearly everywhere, while drifter data indicate that there is a flow toward the pole nearly everywhere. Thus the meridional Ekman flows are strong enough not only to cancel the near-surface geostrophic currents but also to transport surface layer water in the opposite direction. Surface layer Ekman divergence near the equator in particular is important in determining the equatorial upwelling circulation [Wyrtki, 1981]. Also, compared to ship drift charts, the drifter data show a splitting and divergence of the South Equatorial Current between 110° and 136°W, with maxima in westward flow to the north and south of the equator.

3.1.2  Mean seasonal cycle

The seasonal cycle of SST in the equatorial Pacific has been well documented from COADS and other VOS-based analyses [e.g., Reynolds and Smith, 1995]. Warmest SSTs in the cold tongue occur in boreal spring, and coolest SSTs occur in boreal autumn. The amplitude of these annual period variations diminishes from east to west as the thermocline deepens (Figure 9); similarly, the timing of maximum temperatures occurs later in the boreal spring progressing from west to east [e.g., Horel, 1981; Enfield, 1986; Chao and Philander, 1991]. The westward progression of the annual cycle of SST along the equator in the Pacific is related to the westward progression in the zonal winds [Chang, 1994; Xie, 1994]. Annual variations in SST in turn set up atmospheric boundary layer pressure gradients which drive annual period zonal wind variations [Nigam and Chao, 1996].

Figure 9: Mean seasonal cycles of temperature and zonal velocity at four sites along the equator based on multiyear analyses (1980–1994 at 110°W, 1983–1994 at 140°W, 1988–1994 at 170°W, and 1986–1993 at 165°E). The 110°W, 140°W, and 165°E analyses are updated versions of those found in work by McPhaden and McCarty [1992] and McCarty and McPhaden [1993]. The 170°W analysis is based on data presented by Weisberg and Hayes [1995], extended through 1994.

Although solar forcing near the equator is predominantly at semiannual periods, SST in the equatorial cold tongue of the eastern and central Pacific is dominated by annual period variations because of the importance of ocean dynamics and the influence of land masses bordering the Pacific [Li and Philander, 1996]. Recent diagnostic studies and model results illustrate the complex mix of ocean processes in accounting for the amplitude and phase of seasonal SST variations in this region [Hayes et al., 1991b; Köberle and Philander, 1994; Chang, 1993, 1994; Chen et al., 1994a]. The shallow mean thermocline depth in the eastern Pacific, which is due to large-scale wind forcing (Figure 6), is important in facilitating upwelling and vertical mixing to cool the surface. Zonal advection associated with seasonally varying currents is also important, particularly in the central Pacific [Chen et al., 1994a; Minobe and Takeuchi, 1995]. Variations in surface heat fluxes (mainly solar irradiance and latent heat flux) are significant at all locations. These fluxes assume a dominant role as ocean dynamical processes diminish poleward away from the equator and in the western equatorial Pacific where the thermocline is deep. In this latter region the semiannual period in solar irradiance forcing leads to the dominant semiannual period in SST (Figure 9).

Studies using XBT and conductivity-temperature-depth (CTD) data have described the seasonal cycle of upper ocean thermal structure based on the dynamics of Ekman pumping and Rossby waves [Delcroix and Henin, 1989; Kessler, 1990; Kessler and McCreary, 1993]. Seasonal variations in transports of major currents have also been documented using XBT and tide gauge data by Taft and Kessler [1991], Picaut and Tournier [1991], and Donguy and Meyers [1996a]. Mitchum and Lukas [1990] used a set of sea level data lying along the North Equatorial Countercurrent trough to show that annual variations propagate to the west as a Rossby wave resonantly forced by westward propagating components in the wind field. Recent model simulations of the seasonal cycle, validated against TOGA observations [e.g., Minobe and Takeuchi, 1995], confirm the results of these empirical studies on the importance of wind stress forcing and equatorial wave processes.

Reverdin et al. [1994], developed a climatology of the surface currents in the tropical Pacific from TOGA drifter and mooring data. A notable aspect of the mean seasonal cycle along the equator is the "springtime reversal" of the normally westward flowing South Equatorial Current [Halpern, 1987b]. It is most evident in the eastern Pacific where, for example, eastward flow of over 30 cm s-1 occurs in April–May at 110°W (Figure 9). This reversal in flow propagates westward along the equator [McPhaden and Taft, 1988], as do zonal winds and SST [Horel, 1981; Lukas and Firing, 1985], with variations at 140° and 170°W lagging those farther to the east. The springtime reversal in the SEC had been known for nearly a century [Puls, 1895], though its magnitude was underestimated because of contamination of ship drift estimates by windage on ship's hulls [McPhaden et al., 1991]. Model simulations suggest that the springtime reversal results from the seasonal relaxation of the zonal component of trade winds, causing flow to accelerate eastward down the zonal pressure gradient [Chao and Philander, 1991; Yu et al., 1997].

The mean seasonal cycle of the Equatorial Undercurrent along the equator has been described in several reports [Halpern, 1987b; McPhaden and McCarty, 1992; McCarty and McPhaden, 1993; Weisberg and Hayes, 1995]. Juxtaposing seasonal analyses based on these studies (Figure 9) helps to highlight some of the important characteristics of variability on this timescale. The EUC, on average, is located in the upper thermocline and is therefore found at greater depths in the west than in the east. Zonal current variations are confined principally to above the Undercurrent core, with a maximum eastward flow in the thermocline occurring in boreal spring at all longitudes.

Recent analyses suggest that the seasonal cycle is nonstationary in the eastern equatorial Pacific [Gu et al., 1998]. Specifically, at 110°W the annual period in thermocline depth variations was much more pronounced in the 1990s than in the 1980s, presumably because of changes in the annual cycle of zonal wind forcing farther to the west. Interestingly, amplification of thermocline depth variations was not reflected in amplified annual SST variations at 110°W. The mean depth of the thermocline remained sufficiently shallow in the eastern Pacific that, consistent with the theories of Köberle and Philander [1994] and Xie [1994], the efficiency of ocean-atmosphere interactions and ocean dynamical processes to cool the surface would not have been significantly impacted.

3.2  ENSO Variability

Some of the hallmark manifestations of the ENSO cycle are illustrated in Plate 1, which shows time series of the Southern Oscillation Index (SOI) and of surface zonal wind stress anomalies and sea surface temperature anomalies along the equator. The period shown (1982–1995) encompasses the 1982–1983 El Niño and interannual variability during the TOGA decade (1985–1994). Each warm episode (1982–1983, 1986–1987, 1991–1992, 1993, and 1994–1995) is associated with negative SOI values and weaker than normal trade winds over about 60° of longitude in the central and western Pacific. In the case of the intense 1982–1983 El Niño the trade winds weakened progressively from west to east all the way across the basin. Conversely, the 1988–1989 cold La Niña event was associated with high SOI values and a strengthening of the trade winds over roughly 60° of longitude. Also noteworthy in Plate 1 is the persistence of warm SST anomalies near the date line and the occurrence of three distinct warm episodes in the eastern Pacific in concert with consistently low Southern Oscillation Index values between 1991 and 1995. Although it is known that the frequency and intensity of ENSO events are modulated on decadal and longer timescales [Gu and Philander, 1995], the duration of warm phase ENSO conditions over 5 calendar years is unparalleled in this century [Trenberth and Hoar, 1996].

Plate 1: Time-longitude plots of zonal pseudostress (in m2 s-2) and SST (in °C) between 2°N and 2°S along the equator from 1982–1995. Pseudostress time series are from the Florida State University (FSU) analyses [Stricherz et al., 1992], and the SST is from Reynolds and Smith [1994]. Also shown is the Southern Oscillation Index (SOI) for the same time period. The SOI, defined as the normalized difference in surface pressure between Tahiti, French Polynesia and Darwin, Australia is a measure of the strength of the trade winds, which have a component of flow from regions of high to low pressure in the tropical marine boundary layer. High SOI (large pressure difference) is associated with stronger than normal trade winds and La Niña conditions, and low SOI (smaller pressure difference) is associated with weaker than normal trade winds and El Niño conditions. All time series have been smoothed with a 5-month triangle filter (roughly equivalent to a seasonal average). The FSU pseudostress and Reynolds SST have also been smoothed zonally over 10° longitude.

The relationship between surface winds and SST for December 1994 (Figure 10) illustrates another important aspect of ENSO variability. Deep atmospheric convection typically occurs over the warmest SSTs in the tropical Pacific [e.g., Graham and Barnett, 1987]. Warmest SSTs (> 30°C) in December 1994 were situated just south of the equator near the date line in a region of strongly convergent surface winds and active deep atmospheric convection [Climate Analysis Center, 1994]. Converging winds act to sustain both deep convection (via moisture convergence) and warm SSTs (via ocean dynamics) [Philander et al., 1984]. These processes tend to locally reinforce one another, and representing them properly in coupled ocean-atmosphere models has been one of the challenges of ENSO modeling [e.g., Zebiak and Cane, 1987; Battisti, 1988; Battisti and Hirst, 1989; Schopf and Suarez, 1988].

Figure 10: Wind vectors and SSTs from the TAO array for December 1994. (top) Monthly means; (bottom) monthly anomalies from the COADS wind climatology and NCEP SST climatology (1950–1979). SSTs warmer than 29°C and colder than 27°C are shaded; SST anomalies >1°C and <-1°C are shaded.

An important oceanic feature of the ENSO cycle is the zonal redistribution of warm surface layer water masses [White et al., 1985; Donguy, 1987; Donguy et al., 1989; McPhaden et al., 1990a; McPhaden and Hayes, 1990b; Kessler and McPhaden, 1995a]. In the western Pacific the thermocline (as indicated by the depth of the 20°C isotherm) shoals 20–50 m in the latitude band 15°S to 20°N during El Niño, whereas in the eastern Pacific the thermocline deepens by a comparable amount but in a narrower band of latitudes than in the west. These thermocline depth variations, illustrated along the equator in Figure 11 for the 1991–1993 El Niño, are correlated with changes in the strength of major currents. The westward SEC weakens significantly during El Niño episodes, while in some events the NECC intensifies [Taft and Kessler, 1991; Kessler and McPhaden, 1995a]. Thus there is an anomalous eastward mass transport of warm water by the equatorial surface currents during the onset of warm events.

Figure 11: Time-longitude sections of anomalies in surface zonal winds (in m s-1), sea surface temperature (in °C), and 20°C isotherm depth (in meters) for January 1991 to December 1993. Analysis is based on 5-day averages between 2°N and 2°S of moored time series data from the TAO Array. Anomalies are relative to monthly climatologies cubic spline fitted to 5-day intervals (COADS winds, Reynolds and Smith [1995] SST, CTD/XBT 20°C depths). Shading indicates anomaly magnitudes > 2 m s-1, 1°C, and 20 m for winds, temperatures, and 20°C depths, respectively. Positive winds are westerly. Squares on the top abscissa indicate longitudes where data were available at the start of the time series, and squares on the bottom abscissa indicate where data were available at the end of the time series.

Changes in the zonal distribution of upper ocean heat content are reflected in sea level variations [e.g., Rebert et al., 1985; Delcroix and Gautier, 1987] because of the vertically coherent structure of the upper ocean thermal field on seasonal-to-interannual timescales. In other words, anomalously deep thermocline tends to be associated with anomalously high sea level and vice versa. Wyrtki [1984] described the sea surface height gradient along the equator during the 1982–1983 El Niño assuming that the long-term mean sea level at tide gauges along the equator was equal to the long-term surface dynamic height relative to a deep reference level. He showed that the normal upward slope of sea level from east to west (Figure 7) was sharply reduced and at times reversed in the eastern and central Pacific during 1982–1983. Reduction and reversal of the sea surface slope also occurred in the 19861987 and 1991–1992 El Niño events (Figure 12). Variations were weaker at these times than in 1982–1983 though, as expected from the weaker and less zonally extensive westerly wind anomalies along the equator (Plate 1). Conversely, during the 1988–1989 cold La Niña event the sea level slope along the equator intensified, in association with stronger than normal trade winds (Figure 12).

Figure 12: Zonal slope of sea surface height along the equator. Sea level anomalies from the 1975–1987 mean seasonal cycle were taken from seven locations near the equator: Rabaul (4°S, 152°E), Kapingamarangi (1°N, 155°E), Nauru (0.5°S, 167°E), Tarawa (1°N, 173°E), Kanton (3°S, 172°W), Christmas Island (2°N, 157°W), and the Galapagos Islands (0.5°S, 90°W). These anomalies were added to the mean dynamic topography difference (0–1000 dbar) computed from the Levitus and Boyer [1994] and Levitus et al. [1994a] temperature and salinity climatologies in order to calculate absolute heights. (top) Mean conditions during three warm events are shown as solid circles (June 1982 to May 1983), crosses (January to December 1987), and open circles (June 1991 to May 1992). The heavy solid line is the long-term mean conditions taken from the Levitus climatology. (bottom) Warm and cold conditions are contrasted by showing the difference (the vertical bars) of the mean sea level anomaly in 1988 (cold) minus the mean sea level anomaly in 1987 (warm).

Sea level slope along the equator is an index for the strength of the zonal pressure gradient, which is the driving force for the Equatorial Undercurrent [Philander and Pacanowski, 1980; McCreary, 1980; McPhaden, 1981]. Reduction and reversal of this sea level slope were associated with a significant weakening and disappearance of the EUC in the thermocline during the 1982–1983 El Niño [Firing et al., 1983; Halpern, 1987b] and the 1986–1987 El Niño [McPhaden et al., 1990a]. The EUC, though it did not disappear during the 1991–1993 El Niño, was greatly reduced in strength in the central Pacific for several months [Kessler and McPhaden, 1995a]. El Niño related reductions in Undercurrent strength have significant implications for the heat balance of the surface layer, since the Undercurrent is normally a source of cold water to feed equatorial upwelling [Bryden and Brady, 1985].

Near the equator, adjustment of the upper ocean heat and mass is strongly influenced by excitation and propagation of equatorial Kelvin and long Rossby waves, which are the primary mechanisms by which the winds communicate their influence to other parts of the ocean basin. The Kelvin waves most prominent in equatorial time series data are associated with forcing by westerly wind bursts and the atmospheric Madden and Julian Oscillation [Miller et al., 1988; McPhaden et al., 1988a; Kessler et al., 1995]. These waves are clearly evident in 20°C isotherm depth variations (e.g., Figure 11), as well as in time series of sea level, dynamic height, and zonal currents within 2° latitude of the equator. Using TAO data and Geosat-derived sea level data, Cheney et al. [1987], Miller et al. [1988], McPhaden et al. [1988a], McPhaden and Hayes [1990b], Delcroix et al. [1991, 1994], Johnson and McPhaden [1993a], and Picaut and Delcroix [1995] clearly documented equatorial Kelvin waves propagating eastward with first baroclinic mode phase speeds of 2–3 m s-1 prior to and during the 1986–1987 El Niño. Similarly, analysis of TAO data and TOPEX/POSEIDON sea level data indicated prominent oceanic variability due to equatorial Kelvin waves generated by wind forcing west of the date line during 1991–1995 [Busalacchi et al., 1994; Kessler et al., 1995; Boulanger and Menkes, 1995].

Weakening of the trade winds near the equator in the central and western Pacific at the onset of warm ENSO events leads to a pattern of upwelling favorable wind stress curl which elevates the thermocline locally at extraequatorial latitudes [e.g., Kessler, 1990]. Weakening of the trade winds also excites upwelling long Rossby waves [White et al., 1985, 1987; Kessler, 1990; Boulanger and Menkes, 1995; Kessler and McPhaden, 1995b], the fastest of which propagates westward at phase speeds of one third the Kelvin wave speed. The slower propagation speed of these waves compared to equatorial Kelvin waves implies that elevation of the thermocline in the west lags depression of the thermocline in the east by several months as evident in thermal field and sea level analyses (e.g., for 20°C along the equator between late 1991 to early 1992 in Figure 11). The Geosat analysis of Delcroix et al. [1991] and subsequent modeling study of du Penhoat et al. [1992] for the 1986–1987 El Niño suggest that, in addition to wind forcing, eastern boundary reflections of equatorial Kelvin waves can generate equatorial Rossby waves that affect the evolution of ENSO.

Empirical studies of the surface layer heat balance emphasize the complex mix of processes controlling SST variability on ENSO timescales. For example, the importance of remotely forced equatorial waves in mediating SST variability in the eastern and central Pacific can be inferred from Plate 1. Largest ENSO SST anomalies during 1980–1995 were located significantly to the east of the largest zonal wind anomalies; moreover, large SST anomalies were found in the far eastern Pacific where zonal wind anomalies were weak. Waves affect SST in the cold tongue region by inducing changes in thermocline depth which affect upwelling and vertical mixing rates [e.g., Hayes et al., 1991b; Kessler and McPhaden, 1995a, b]. Waves can also advect temperature fields meridionally and, more importantly, zonally along the equator. Wave- and current-induced zonal advection of the eastern edge of the warm pool produces large interannual SST anomalies in the central Pacific [McPhaden and Picaut, 1990; Picaut and Delcroix, 1995; Picaut et al., 1996].

Local air-sea heat exchanges are also important in the surface layer heat balance of the tropical Pacific on interannual time scales [Liu and Gautier, 1990; Hayes et al., 1991b; Kessler and McPhaden, 1995a]. The most strongly varying components of the surface energy balance are solar irradiance, which is modulated by changes in cloudiness, and latent heat flux which is modulated by changes in wind speed, SST, and relative humidity [Liu, 1988; Waliser et al., 1994]. East of the date line, where ocean dynamics are crucial for generating SST anomalies on interannual time scales, latent heat flux tends to increase with increasing SST, and therefore acts as a negative feedback on developing SST anomalies [Kessler and McPhaden, 1995a; Weisberg and Wang, 1997]. In the western Pacific warm pool, the thermocline is deep, mean horizontal SST gradients are weak, and ocean dynamical processes are less capable of generating large scale SST anomalies than further east. In this region air-sea turbulent heat exchange is an important generating mechanism for SST anomalies, through enhanced evaporation during periods of strong westerly winds [Meyers et al., 1986]. Variations in short wave radiation tend to damp developing SST anomalies throughout the tropical Pacific since high cloudiness, which reduces insolation, tends to occur over the warmest surface waters [Waliser et al., 1994].

Data from the TOGA observing system have been used to test various theories of El Niño and the ENSO cycle. An early theory espoused by Wyrtki [1975] suggested that prior to El Niño, the trade winds strengthened, and there was a increase in sea level (a proxy for heat content) in the western Pacific warm pool. When the trade winds weakened, the overcharged warm water pool would collapse and surge eastward in the form of a Kelvin wave to initiate a warm event. The importance of Kelvin waves in the development of El Niño has been confirmed by many studies. However, other aspects of Wyrtki's theory were undermined when prior to the 1982–1983 El Niño, the strongest of the century, there was no anomalous rise in sea level in the western Pacific or intensification of the easterly trades [Cane, 1984]. Similarly, prior to the equatorial warming in 1993, there was no buildup of heat content in the western Pacific warm pool or intensification of the easterlies [Kessler and McPhaden, 1995b].

Wyrtki [1985a] proposed another hypothesis, namely that warm water accumulated in the tropical Pacific prior to an El Niño on a zonally averaged basis between 15°N and 15°S. In this scenario, El Niño represents a mechanism whereby excess heat is purged to higher latitudes. Cane et al. [1986] interpreted the interannual oscillations in their coupled ocean-atmosphere model in terms of this mechanism. Springer et al. [1990], in a wind-forced ocean model simulation, found a buildup of heat content near the equator prior to the 1982–1983 El Niño as hypothesized by Wyrtki, but only between 5°N and 5°S. The difference in latitude bands over which the buildup was assumed to occur resulted from Wyrtki's use of tide gauge station data which had to be interpolated over great distances zonally beyond 5°N–5°S [Springer et al., 1990]. Miller and Cheney [1990], however, did not find a buildup at all prior to the 1986–1987 El Niño event using Geosat data. Thus Wyrtki's [1985a] mechanism, modified to a narrower band of longitudes, may be operative during some but not all El Niño events.

McCreary [1983] proposed a theory for ENSO in which the timescale between warm events was set by the slow westward propagation of long extraequatorial Rossby waves and their reflection off the western boundary as equatorial Kelvin waves. The reflected Kelvin waves would alter thermocline depths (and by proxy SST) in the eastern Pacific, thereby affecting the strength of the trade winds. In order to get a realistic 3–4-year periodicity for the ENSO cycle, Rossby waves with significant amplitudes at roughly 20° latitude from the equator were required. Using XBT data, Graham and White [1988] argued for the existence of extraequatorial Rossby waves along 12°N and 12°S and their reflection into equatorial Kelvin waves at the western boundary. However, Kessler [1990] offered alternative explanations for the observed variability along the equator in terms of direct wind forcing rather than Rossby wave reflection, and Kessler [1991] showed that only Rossby waves equatorward of about 8° latitude could reflect into equatorial Kelvin waves with significant amplitudes.

The delayed oscillator theory of ENSO [Battisti, 1988; Battisti and Hirst, 1989; Schopf and Suarez, 1988] also involves the reflection of Rossby waves into equatorial Kelvin waves at the western Pacific boundary. In contrast to McCreary's [1983] theory though, equatorial Rossby waves closely trapped to the equator, rather than extraequatorial Rossby waves at higher latitudes, are most relevant. Thermocline changes associated with reflected Kelvin waves lead to SST anomalies in the eastern Pacific cold tongue by altering upwelling rates. The SST anomalies affect the atmospheric convection and circulation, giving rise to local positive feedbacks that reinforce the SST and wind anomalies (e.g., Figure 10). The anomalous surface winds in turn excite equatorial oceanic waves of opposite sign to those that generated the original SST anomalies. The timescale for the ENSO cycle in this theory is set by the competition between the local positive feedbacks and delayed negative feedbacks associated with remotely forced equatorial waves and their western boundary wave reflections.

Tests of the delayed oscillator have focused primarily on the question of whether equatorial Rossby waves can reflect from the irregular and gappy coastal geometry of the western Pacific. Theories suggest coastal irregularities should not be a fundamental limitation to this reflection process [Clarke, 1991; du Penhoat and Cane, 1991]. However, although in principle western boundary reflections should work equally well to both initiate and terminate El Niño events, it appears that they are most effective in terminating events [Li and Clarke, 1994; Mantua and Battisti, 1994]. In this situation, reflection of an upwelling Rossby wave at the western boundary excites an upwelling equatorial Kelvin wave train which erodes the warm SST anomaly in the cold tongue, eventually leading to cool La Niña SST anomalies. Even so, not all warm events appear to be terminated by western boundary reflections. Boulanger and Menkes [1995], for example, found that wind-forced upwelling Kelvin waves, rather than boundary-reflected Kelvin waves, led to cooling along the equator in the eastern Pacific in late 1993. Also, Picaut and Delcroix [1995] argued that the 1986–1987 El Niño was terminated by Rossby waves emanating from the eastern boundary, rather than Kelvin waves emanating from the western boundary.

Few, if any, El Niño events of the TOGA decade appear to have been initiated by delayed oscillator physics. Through extended empirical orthogonal function (EOF) analysis of Geosat data during the 1986–1989 El Niño-La Niña cycle, White and Tai [1992] suggested that an equatorial Rossby wave reflected into an equatorial Kelvin wave at the western boundary, consistent with delayed oscillator theory. However, a detailed projection of Geosat sea level and derived surface currents on individual equatorial wave modes indicated very little evidence of first meridional Rossby wave reflection into Kelvin waves during this time [Delcroix et al., 1994]. Similarly, Kessler and McPhaden [1995b], using TAO and XBT data during 1988–1993, and Boulanger and Menkes [1995], using TAO and TOPEX/POSEIDON data during 1992–1993, found little evidence for the initiation of warm events via Rossby wave reflections at the western boundary. Boulanger and Fu [1996], using TOPEX/POSEIDON altimeter data and ERS-1 wind data, detected wind-forced downwelling equatorial Rossby waves that reflected into downwelling Kelvin waves prior to warming along the equator in middle to late 1994. They interpreted these reflections as evidence for delayed oscillator physics as a trigger for the 1994–1995 El Niño. In contrast, however, Goddard and Graham [1997] argued that this same 1994–1995 warm event in the NCEP reanalysis [Ji and Smith, 1995; see also section 4.4] was not initiated Rossby wave reflection at the western boundary, but rather direct wind forcing near the equator.

Another perspective of the ENSO cycle was proposed by Picaut and Delcroix [1995] and Picaut et al. [1996]. Using hypothetical drifters moved by current fields derived from Geosat and TOPEX/POSEIDON data, TAO mooring data and SVP drifter data, and three different classes of ocean models, these authors found that ENSO-related SST anomalies in the central western Pacific were primarily the result of zonal advection (Figure 13). Picaut and Delcroix [1995] and Picaut et al. [1997] argued that Rossby waves excited by eastern boundary reflections, in addition to the direct effects of wind forcing, were instrumental in generating these currents. Since the impacts of SST variations on the atmosphere are most pronounced in the central and western equatorial Pacific [Geisler et al., 1985], Picaut et al. [1997] argue for a revision of the delayed oscillator theory to provide more weight to oceanic processes affecting this region, including eastern boundary wave reflections. It is evident from this wide variety of theoretical, modeling, and empirical studies that, despite progress made during TOGA on understanding the ENSO cycle, there are many as-of-yet unresolved issues related to the coupled ocean-atmosphere interactions that require further investigation.

Figure 13: (left) Longitude-time distribution of 4°N–4°S averaged SST. Contour interval is 1°C, except for the 28.5°C isotherm. Superimposed as thick lines are the trajectories of two hypothetical drifters moved by 4°N–4°S averaged surface current anomalies derived from Geosat data (thick solid lines correspond to the total currents; thick dashed lines correspond to the Kelvin and Rossby wave contributions). (right) Longitude-time distribution of 4°N–4°S averaged surface current anomaly derived from Geosat. Contour interval is 10 cm s-1. Solid (dashed) lines denote eastward (westward) current anomalies. Thick solid and thick dashed lines are as in Figure 13 (left). From Picaut and Delcroix [1995].

 

3.3  Intraseasonal Kelvin Waves

The Kelvin waves most prominent in equatorial Pacific time series data have energy across a broad band of periods spanning roughly 40–120 days, with maximum energy concentrated near periods of 60–90 days. Sea level, thermocline depth, and zonal currents associated with these waves propagate eastward with 2–3 m s-1 phase speeds [Enfield, 1987; McPhaden and Taft, 1988; Johnson and McPhaden, 1993a, b]. Vertical structures suggest significant energy in both the first and second vertical modes [Kessler and McPhaden, 1995b], consistent with model simulations [e.g., Busalacchi and Cane, 1985; Giese and Harrison, 1990; Kindle and Phoebus, 1995]. There is also evidence that the wave structures are modified by wave-mean flow interactions [Johnson and McPhaden, 1993a, b]. Upon reaching the eastern boundary, the waves can be traced along the coasts of North and South America as coastal Kelvin waves [Spillane et al., 1987].

These Kelvin waves are forced primarily by surface zonal wind variations associated with westerly wind bursts and the Madden and Julian Oscillation in the western Pacific (Figure 11). The amplitude of the ocean wave response depends on the structure of the wind forcing, namely its temporal evolution, zonal fetch, and meridional structure [Knox, 1987; Harrison and Giese, 1991; Giese and Harrison, 1991]. In terms of frequency content, wave energy is concentrated at periods decidedly longer than the dominant 30–60-day period of the wind forcing itself [McPhaden and Taft, 1988]. Kessler et al. [1995] explain this "red shift" as the result of a scale selection process related to wind fetch, which favors excitation of the lower-frequency Kelvin waves in response to wind forcing in the intraseasonal band. Their results are analogous to Knox's [1987] analysis in the time domain, which indicated that an equatorial wind event of duration T and zonal fetch L, would lead to a Kelvin pulse of longer duration T +  L/c, where c is the zonal phase speed of the Kelvin wave.

Intraseasonal Kelvin waves affect SST in the equatorial Pacific in a variety of ways. They can warm SST by zonal advection in the equatorial cold tongue as documented for the 1986–1987 El Niño [Johnson and McPhaden, 1993a] and the 1991–1993 El Niño [Kessler and McPhaden, 1995a]. Downwelling Kelvin waves also depress the thermocline [McPhaden and Hayes, 1990b; Kessler et al., 1995], which can lead to surface warming by reducing the efficiency of local wind-driven upwelling to cool the surface. Lien et al. [1995] found that the passage of a downwelling Kelvin wave during the 1991–1992 El Niño led to a reduction in upper ocean turbulent mixing in the central equatorial Pacific, which would likewise favor the development of warm SST anomalies.

There is a notable relationship between enhanced intraseasonal variability and El Niño in both the ocean and atmosphere [e.g., Keen, 1982; Luther et al., 1983; Lau and Chan, 1986; Enfield, 1987; McPhaden and Hayes, 1990b; Kessler et al., 1995; Kindle and Phoebus, 1995]. During El Niño westerly wind bursts tend to be more prominent, deep convection associated with the Madden and Julian Oscillation tends to be stronger and extend farther eastward along the equator in the Pacific, and intraseasonal equatorial Kelvin waves tend to be of larger amplitude. These findings have led to suggestions that intraseasonal variability, rather than chaotic interactions of the seasonal cycle with itself (see section 3.1), may be responsible for the irregularity of the ENSO cycle [e.g., Zebiak, 1989].

Nonlinear interactions between the ocean and the atmosphere are necessary to couple intraseasonal variations to the ENSO cycle. Harrison and Schopf [1984] proposed a mechanism whereby zonal advection by short-period Kelvin waves could initiate low-frequency warming in the equatorial cold tongue of the eastern and central Pacific, and some coupled models bear out the potential for this mechanism to trigger an El Niño [Latif et al., 1988]. Likewise, Kessler et al. [1995] described how intraseasonal Kelvin waves can contribute to the slow eastward displacement of the western Pacific warm pool, which would favor the development of warm El Niño SST anomalies.

3.4  Local Response to Westerly Wind Burst Forcing

The importance of the local response to strong westerly wind burst forcing in the western Pacific warm pool was first highlighted by Lukas and Lindstrom [1991]. That study and related work ultimately contributed to the design and implementation of the TOGA Coupled Ocean Atmosphere Response Experiment (COARE) [Godfrey et al., this issue]. Westerly wind bursts typically occur during the westerly phase of the Madden and Julian Oscillation [Sui and Lau, 1992], during which surface westerlies may attain speeds of 5–10 m s-1. These wind events lead to dramatic zonal current reversals in time and depth in the upper 100–150 m of the water column [McPhaden et al., 1988a, 1992; Delcroix et al., 1993; Kuroda and McPhaden, 1993; Kutsuwada and Inaba, 1995; Ralph et al., 1997]. The surface flow accelerates eastward and can reach speeds of over 100 cm s-1 in the course of a week. The resultant jet may extend over 40° of longitude, with anomalous eastward transports of 50 Sverdrups (1 Sv = 106 m3 s-1) between 5°N and 5°S. Westerly wind bursts and the westerly phase of the Madden and Julian Oscillation are usually associated with a drop in SST due to increased latent heat flux and reduced insolation [McPhaden and Hayes, 1991; Weller and Anderson, 1996; Cronin and McPhaden, 1997]. Strong wind forced currents advecting fresher water from the west, in combination with enhanced precipitation, generally lead to a freshening of the surface layer near the equator in the warm pool region. These processes can lead to barrier layer formation [Sprintall and McPhaden, 1994; Roemmich et al., 1994; Anderson et al., 1996], which Lukas and Lindstrom [1991] hypothesized as important for understanding the evolution of ENSO warm events. Wind burst forcing also excites downwelling equatorial Kelvin waves which propagate into the eastern Pacific as discussed in the previous section.

3.5  Instability Waves

Tropical instability waves, first observed in the Pacific in satellite SST imagery [Legeckis, 1977], typically propagate westward with zonal wavelengths of 800–2000 km and periods of 20–30 days. They have been observed in ocean currents, temperatures, and salinity [Philander et al., 1985; Pullen et al., 1987; Halpern et al., 1988; McPhaden et al., 1990c; Kessler and McPhaden, 1995a; Qiao and Weisberg, 1995; McPhaden, 1996; Flament et al., 1996]. They are also detectable in Geosat and TOPEX/POSEIDON altimetry data [Perigaud, 1990; Giese et al., 1994; Busalacchi et al., 1994] despite the relatively coarse temporal resolution of altimeters compared to the basic frequency of the waves. Instability waves are seasonally and interannually modulated, being weakest during boreal spring and during the warm phase of ENSO. The waves derive their energy from the large-scale, seasonally varying zonal equatorial currents through shear instability [Philander, 1978; Cox, 1980; Philander et al., 1986; Luther and Johnson, 1990] and possibly through SST frontal instabilities [Yu et al., 1995]. As such, they are a significant source of drag on the South Equatorial Current and Equatorial Undercurrent, and they heat the cold tongue through large downgradient (i.e., equatorward) eddy heat transports [Hansen and Paul, 1984; Bryden and Brady, 1989]. The waves also affect the stability of the atmospheric boundary layer [Hayes et al., 1989b], the distribution of cloudiness [Deser et al., 1993], latent heat fluxes [Zhang and McPhaden, 1995], and the distribution of nutrients, pCO2, and other chemical species in the eastern equatorial Pacific [Feely et al., 1994]. Instability waves of similar character have been documented in the equatorial Atlantic, where they are evident during the boreal summer season [e.g., Weisberg and Weingartner, 1988; Musman, 1992]. They are a potentially significant source of aliased energy which, if unresolved (as in infrequently sampled shipboard data), add noise contamination to lower-frequency signals of climatic interest [Hayes and McPhaden, 1992; Kessler et al., 1996].

3.6  ENSO and the Indo-Pacific Throughflow

Wyrtki [1987] first attempted to monitor the variations of the throughflow by computing the large-scale pressure gradient between the western Pacific and eastern Indian Oceans. Davao in the Philippines was used for the western Pacific, and Darwin in Australia was used for the eastern Indian Ocean. He found that this difference was dominated by seasonal variations but that the two records were coherent at interannual timescales, resulting in a small difference on ENSO timescales. Later, Clarke [1991] modeled the reflection and transmission of large-scale, low-frequency waves at a gappy western Pacific boundary and found that the interannual sea level variations along northern Australia were in fact of Pacific origin. These results implied that the Davao-Darwin sea level difference was not an appropriate index for the throughflow at interannual timescales. Clarke and Liu [1993, 1994] argued that a better index of the throughflow would be based on differences between northern and southeastern Indian Ocean sea levels. Their index suggested that the throughflow increased during cold ENSO events and decreased during warm events.

Thermal structure associated with the Indonesian throughflow in the eastern Indian Ocean has marked interannual variations, which have been documented on a frequently repeated XBT line between Shark Bay (northwestern Australia) and Sunda Strait (Java) [Meyers, 1996]. The largest variations of dynamic height and depth of the thermocline are near the coast of Australia (Figure 14, left), and they are highly correlated to the ENSO signal in the western equatorial Pacific (Figure 14, right). The XBT observations are consistent with the study by Clarke and Liu [1994] and with their model of the generation of the variations by wind forcing with long timescales. The XBT measurements also document how the signal extends into the ocean interior and how it is related to variations on the coast of Indonesia. The observations and model consistently indicate that variations near the coast of western Australia are generated by winds over the equatorial Pacific, while variations near the coast of Indonesia are generated by winds over the equatorial Indian Ocean. The differences in vertically integrated dynamic height between the coasts of Australia and Indonesia are a measure of the transport of Indonesian throughflow. The estimated mean transport, based on XBT data, is 7 Sv [Meyers et al., 1995]. Consistent with the tide gauge measurements, the throughflow is weaker during El Niño, with a peak-to-trough amplitude on interannual timescales of transport in the upper 400 m of about 5 × 106 m3 s-1. What impact ENSO timescale variations in throughflow have on the climate of the Indian Ocean region is, however, unclear.

Figure 14: Joint empirical orthogonal functions (EOFs) of anomalies of SST, dynamic height (0–400 dbar) and depth of the 20° isotherm on a frequently repeated XBT line between Shark Bay (westernmost point of Australia) and Sunda Strait (western end of Java). (left) The first EOF (34% of the variance) shows the ENSO signal entering the Indian Ocean along the coast of Australia. (right) The temporal coefficients of the first EOF are highly correlated with the Southern Oscillation Index (SOI). From Meyers [1996].

 

3.7  ENSO and Global Oceanic Variability

Although the TOGA observing system focused primarily on the ENSO phenomenon in the tropical Pacific, satellite and some in situ measurement programs (e.g., VOS, tide gauges, and drifters) provided a global perspective on climate variations during the TOGA decade. In this section we briefly review studies of climate phenomena facilitated by measurements outside the tropical Pacific, with emphasis on variability related to ENSO.

Atmospheric teleconnections associated with the ENSO cycle affect oceanic variability in wide-ranging parts of the globe. Over the North Pacific Ocean, for example, the Aleutian Low becomes anomalously strong during the late fall and winter of an El Niño year. Associated with these changes in atmospheric pressure, the axis of the subtropical jet stream splits, one branch displaced southward, steering storms into the southwestern United States, and another branch displaced northward into the Pacific Northwest. Air-sea heat exchange is enhanced at midlatitudes by these changes in atmospheric circulation [Alexander, 1992], leading to cold open ocean SST anomalies during El Niño years [Wallace et al., this issue]. Along the west coast of the United States, on the other hand, anomalous alongshore southerly winds during El Niño can lead to reduced coastal upwelling, which contributes to warmer coastal SSTs and higher coastal sea level [Enfield and Allen, 1980; Ramp et al., 1997, and references therein].

In addition to this atmospheric teleconnection pathway between the tropical and midlatitude Pacific Ocean, equatorial oceanic Kelvin waves impinge on the eastern boundary, forcing poleward propagating coastal Kelvin waves in both hemispheres [Enfield and Allen, 1980; Chelton and Davis, 1982; Clarke, 1992; Clarke and Van Gorder, 1994; Ramp et al., 1997; Shaffer et al., 1997]. Roach et al. [1989] concluded that these signals dominate sea level variability as far north as San Francisco. These waves are particularly energetic at intraseasonal periods [Spillane et al., 1987]. Recently, Jacobs et al. [1994] found that Rossby wave signals forced at the eastern boundary by the passage of El Niño-related coastal Kelvin waves associated with the 1982–1983 El Niño could be detected in the central and western North Pacific a decade later. Jacobs et al. [1994] speculated that these Rossby waves contributed to the development of SST anomalies in the midlatitude North Pacific by rerouting the warm, normally eastward flowing Kuroshio Extension off Japan to a more northeasterly course in the early 1990s.

White and Peterson [1996] have recently detected a 4–5-year eastward propagating, zonal wave number two oscillation encircling the globe in the Antarctic Circumpolar Current. The wave is characterized by coherent oscillations in SST, sea level pressure, meridional winds, and sea ice extent. White and Peterson [1996] hypothesized that this wave may be related to forcing associated with El Niño through atmospheric teleconnections between the tropical Pacific and the Southern Ocean.

The tropical Atlantic is characterized by a prominent mean seasonal cycle in surface winds, sea level upper ocean currents, and temperatures [e.g., Carton and Katz, 1990; Reverdin et al., 1991a, b; Molinari and Johns, 1994; Katz et al., 1995b]. In addition, two important modes of interannual-to-decadal variability are evident around this seasonal cycle, one of which consists of warm events with variability concentrated near the equator [Philander, 1986; Houghton, 1991; Zebiak, 1993; Carton and Huang, 1994] and another of which consists of interhemispheric variations in tropical SST [Moura and Shukla, 1981; Servain, 1991; Houghton, 1991; Houghton and Tourre, 1992]. Dynamics intrinsic to the ocean-atmosphere-land system in the Atlantic basin are important in determining the variability associated with these low-frequency climate signals. However, ENSO teleconnections through the atmosphere influence their evolution as well, as discussed by Servain [1991], Delecluse et al. [1994], and Enfield and Mayer [1997].

Variability in the Indian Ocean is dominated by a pronounced seasonal cycle related to monsoon wind forcing [Rao et al., 1989; Molinari et al., 1990; Perigaud and Delecluse, 1992; Mizuno et al., 1995; Donguy and Meyers, 1995, 1996a; Meyers et al., 1995]. However, interannual anomalies on ENSO timescales are detectable as well [e.g., Perigaud and Delecluse, 1993; Tourre and White, 1995]. Tourre and White's [1995] simultaneous analysis of upper ocean thermal data in all three tropical ocean basins indicated what appeared to be a coherent eastward propagating interannual wave in upper ocean heat content near the equator. On the strength of this result they suggested the possibility of oceanic precursors to ENSO in the Indian Ocean thermal field, in addition to atmospheric precursors believed to be important in association with the monsoons [Webster and Yang, 1992]. Latif and Barnett [1995], on the other hand, argued that the Pacific forces the tropical Indian and Atlantic Oceans remotely through atmospheric teleconnections on ENSO timescales and that this forcing accounts for a significant percentage of the observed thermal variability described by Tourre and White [1995].

3.8  Salinity Variations

For the three tropical oceans, long-term averaged sea surface salinity (SSS) exhibits well-documented minima associated with the Intertropical Convergence Zones as well as relatively high salinities, mainly where evaporation significantly exceeds precipitation. Maximum seasonal SSS variations are found primarily in the Intertropical Convergence Zones and in the South Pacific Convergence Zone, in close relation to seasonal variations in rainfall [Delcroix and Henin, 1991; Dessier and Donguy, 1994; Donguy and Meyers, 1996b]. There is also notable ENSO-related SSS variability. During El Niño periods the SSS field west of about 150°W is characterized by fresher than average SSS within 8°N–8°S; conversely, saltier than average SSS is found poleward of 8° latitude [Delcroix and Henin, 1991; Delcroix et al., 1996]. There is also significant freshening of the surface layer in the eastern Pacific within 10° of the equator during El Niño, particularly east of 110°W [Ando and McPhaden, 1997]. SSS anomalies of reverse sign are observed during La Niña periods. In the equatorial band these interannual modifications in the salinity field result mainly from the combined effects of rainfall and horizontal salt advection, the latter process apparently dominating west of about 165°E [Picaut et al., 1996; Delcroix and Picaut, 1998; Ando and McPhaden, 1997; Henin et al., 1998].

Lukas and Lindstrom [1991] proposed that salinity variability of the upper ocean may be an important determinant in the evolution of ENSO. They hypothesized that in regions of heavy rainfall, thin surface mixed layers form which are isolat