Equatorial westerly (easterly) wind anomalies, phase‐locked to the seasonal cycle, typically ‘propagate’ from the eastern equatorial Indian Ocean and into the western Pacific immediately before an El Niño (La Niña). A space‐time integration of this Indo‐Pacific signal yields an index τ that, for 11 out of 12 months of the calendar year, leads the El Niño index NINO3.4 with a correlation of 0.5 or greater for at least some lead in the range 10–15 months. Cross‐validated hindcasts suggest that a linear combination of this atmospheric index and the ocean indices NINO3.4 and (t), the anomalous equatorial Pacific upper ocean heat content, is an excellent predictor of El Niño. It can predict across the nearest spring persistence barrier, but not the one after that. The present El Niño should die over the spring, leaving near neutral conditions for the rest of 2003.

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Discharge and recharge of the warm water volume (WWV) above the 20°C isotherm in an equatorial Pacific Ocean box extending across the Pacific from 156°E to the eastern ocean boundary between latitudes 5°S and 5°N are key variables in ENSO dynamics. A formula linking WWV anomalies, zonally integrated wind stress curl anomalies along the northern and southern edges of the box, and flow into the western end of the box is derived and tested using monthly data since 1993. Consistent with previous work, a WWV balance can only be achieved if the 20°C isotherm surface is not a material surface; that is, warm water can pass through it. For example, during El Niño, part of the WWV anomaly entering the box is cooled so that it is less than 20°C and therefore passes out of the bottom of the box, the 20°C isotherm surface. The observations suggest that the anomalous volume passing through the 20°C isotherm is approximately the same as TЈ W , the anomalous WWV entering the western end of the box. Therefore the observed WWV anomaly can be regarded as being driven by the anomalous wind stress curl along the northern and southern edges of the box. The curl anomaly changes the WWV both by divergent meridional flow at the edges of the box and vortex stretching; that is, the Sverdrup balance does not hold in the upper ocean. A typical amplitude for the rate of change of WWV for the 5°S-5°N box is 9.6 Sv (Sv ϵ 10 6 m 3 s Ϫ1 ). The wind stress curl anomaly and the transport anomaly into the western end of the box are highly correlated with the El Niño index Niño-3.4 [the average sea surface temperature anomaly (SSTA) over the region 5°S-5°N, 170°-120°W] and Niño-3.4 leads minus the WWV anomaly by one-quarter of a cycle. Based on the preceding results, a simple discharge/recharge coupled ENSO model is derived. Only water warmer than about 27.5°-28°C can give rise to deep atmospheric convection, so, unlike past discharge/recharge oscillator models, the west-central rather than eastern equatorial SSTAs are emphasized. The model consists of two variables: TЈ, the SSTA averaged over the region of strong ENSO air-sea interaction in the west-central Pacific equatorial strip 5°S-5°N, 156°E-140°W and DЈ, the 20°C isotherm depth anomaly averaged over the same region. As in the observations, TЈ lags DЈ by one-quarter of a cycle; that is, ‫ץ‬TЈ/‫ץ‬t ϭ DЈ for some positive constant . Physically, when DЈ Ͼ 0, the thermocline is deeper and warmer water is entrained through the base of the mixed layer, the anomalous heat flux causing ‫ץ‬TЈ/‫ץ‬t Ͼ 0. Also, when DЈ Ͼ 0, the eastward current anomaly is greater than zero and warm water is advected into the region, again causing ‫ץ‬TЈ/‫ץ‬t Ͼ 0. Opposite effects occur for DЈ Ͻ 0. A second relationship between TЈ and DЈ results because the water is warm enough that TЈ causes deep atmospheric convection anomalies that drive the wind stress curl anomalies that change the heat storage ‫ץ‬DЈ/‫ץ‬t. The atmosphere responds essentially instantly to the TЈ forcing and the curl causes a discharge of WWV during El Ni...

Many years of simultaneous hourly buoy wind and directional wave spectra data in the Gulf of Mexico and the Pacific were used to estimate Stokes drift and u * e w where u * = (magnitude of the local windstress/water density) 1/2 and e w is a unit vector in the direction of the local wind. Stokes drift and u * e w were strongly vectorally correlated, the two vectors on average being within a few degrees of one another. This result remained valid even when there was evidence of remotely forced swell. Extension of the observed wave spectra above 0.35 Hz to the u * -dependent wave breaking frequency shows that typically the e-folding scale of the Stokes drift with depth is less than 1.8 m, much smaller than the Ekman layer e-folding scale. Therefore, there is negligible induced Eulerian cancellation of the Stokes drift, and the surface particle movement is governed by the Eulerian velocity + |u Stokes |e w . Taking into account wave spreading, |u Stokes | typically ranges from about 3 to 13 cm/s. Thus, the Stokes drift, which can be estimated directly from the wind stress, is an order one contributor to the surface transport of particles.Plain Language Summary Although crucial for the movement of oil spills, red tide, fish eggs and larvae, and floating garbage, much still has to be learned about net particle movement in the top 1 or 2 m of the ocean. George Gabriel Stokes showed mathematically in 1847 that ocean surface waves may affect the net movement of particles at the ocean surface, but later it was shown that because we live on a rotating Earth, the net particle movement in the direction of the waves (the "Stokes drift") might be canceled by another opposite flow. In this paper we demonstrate that because of the ocean turbulence generated by the wind, the main part of the Stokes drift in the top 2 m of the ocean is not canceled by an opposing flow. Furthermore, analysis of simultaneous hourly wind and wave measurements for many years at Christmas Island in the equatorial Pacific, ocean station Papa in the north Pacific, and 10 stations in the Gulf of Mexico shows that Stokes drift is strongly related to the local wind and is in the direction of the wind. Stokes drift is therefore not primarily due to remotely driven swell; rather, it is mainly due to the much shorter waves that the local wind generates. By taking into account when the short waves break, it is shown how Stokes drift can be approximately estimated directly from the local wind.It has been recognized (see, e.g., Tamura et al., 2012;Webb & Fox-Kemper, 2011) that since Stokes drift is proportional to frequency cubed, its magnitude depends strongly on the high-frequency tail of the spectrum. However, so far, it has been unclear how to calculate the Stokes drift, because a justification of how to limit the high-frequency contribution has not been given. In section 5 we offer a physical argument, backed by observations, for how this might be done.But before doing this we discuss two other ways that the inclusion of the high-frequency tail is crucial to ...