separated western boundary current provided the source of warm, saline water required to initiate the anomaly development. Advection set the oscillation time scale, which was given by the length of time it took a particle to be advected from the mid-ocean region, between the subpolar and subtropical gyres, to the eastern boundary and then, as subsurface flow, toward the polar boundary. In more complicated geometry, they argued, one would expect this time scale to be slightly longer since the advective paths would no longer be along straight lines.
Figure 3, from Weaver and Sarachik (1991b), illustrates the meridional overturning stream function throughout the course of one particular oscillation in a single, Southern Hemisphere basin. During the oscillation the poleward heat transport changed by as much as a factor of 3 at certain latitudes. For example, at 26°S, during the most intense stage of the oscillation (Figure 3d) 0.29 petawatts (1 petawatt = 1015 W) of heat was being transported poleward, whereas during the weakest phase (Figure 3h) this was reduced to only 0.11 petawatts. The changes in heat transport corresponded directly to changes in the heat lost to the overlying atmosphere at high latitudes, since the ocean stored little heat during the oscillation. If such internal variability were to exist in the real ocean, it would evidently have a profound effect on global climate.
Weaver et al. (1991) attempted to understand why Marotzke (1989, 1990, 1991) and Marotzke and Willebrand (1991) almost never saw such spontaneous internal decadal oceanic variability (although one of the two-basin experiments described in Marotzke (1990) indicated the presence of decadal-to-interdecadal-scale oscillations in the region of the Antarctic Circumpolar Current), whereas Weaver and Sarachik (1991a,b) almost always found such variability. Through a systematic analysis of the forcing fields used in these works, they concluded that the presence of decadal-to-interdecadal variability was linked to an area of negative P - E at middle to high latitudes, and fresh-water gain further north. The meridional gradients in the fresh-water flux forcing field also had to be sufficiently strong that the system was in a haline-dominant regime. Figure 4a illustrates the basin-averaged surface heat flux over the course of the integration of one of the single-hemisphere experiments of Weaver et al. (1991), while Figure 4b shows the power spectral density of this curve over the first 8,214 years of integration. What is readily evident from this figure is that the dominant variability is in the decadal band, with associated basin-averaged surface heat-flux anomalies of between + 6 and - 10 W m-2 over one oscillation. Although the integration shown in Figure 4a eventually reached a steady state, Weaver et al. (1993) showed that the inclusion of a stochastic component in the fresh-water flux forcing field continually excited the decadal variability, with the result that no equilibrium was ever reached.
Weaver et al. (1991) suggested that when the thermoha-
line circulation was weak, it slowly passed through the region with negative P - E, and hence the surface waters became more saline. A warm, saline surface anomaly then developed through convection, and this anomaly was advected to the eastern boundary by the mean flow, from which it was convected to the deeper ocean (as in Weaver and Sarachik, 1991b). This led to the subsequent generation of the reverse cell seen in Figures 3a and 3b and Figures 3i and 3j, which in turn caused the thermohaline circulation to intensify. The intensified thermohaline circulation passed rapidly through the evaporative region; hence, the surface waters did not become as saline. Deep water then formed at high latitudes until high-latitude freshening dominated and the thermohaline circulation slowed down. This whole process repeated itself, with the time scale of the oscillation