Influence of Indonesian Throughflow

Significant water exchange between the Pacific and Indian Oceans occurs in the Indonesian Archipelago, where the Indonesian Throughflow allows an outflow of upper-ocean water estimated to be 5-15 Sv (Fine, 1985; Gordon, 1986; Godfrey, 1989; Hirst and Godfrey, 1993).

It is balanced by an inflow of intermediate waters from the Antarctic Circumpolar Current (ACC;

Reid, 1997). We define the Pacific Interocean Circulation (IOC) to be all the pathways that connect this southern inflow to the Indonesian outflow. Defined in this way, the IOC is the Pacific branch of the hypothesized global “conveyor belt” (Gordon, 1986).

Influences of the Pacific IOC in the equatorial ocean have been discussed in several recent modelling studies. For example, Semtner and Chervin (1992) noted that 4-10 Sv of deep IOC water mixes upward into the equatorial thermocline in their solution. Blanke and Raynaud (1996) reported that IOC water flows in their model EUC, and as a result ²₃ of the extratropical water in the EUC is of southern-hemisphere origin. Shriver and Hurlburt (1996) outline the IOC pathways present in a solution to their 6-layer GCM. In all of the pathways, IOC water enters the Pacific in the southeastern basin, and eventually moves to the equator near the western boundary. In one of the pathways, IOC water within layer 3 is then allowed to flow directly out of the basin in the Throughflow. In another, IOC water within layer 2 flows eastward in the EUC, upwells into layer 1 in the eastern, equatorial basin, flows into the northern subtropical ocean where it subducts back into layer 2, and only then leaves the basin in the Throughflow.

Lu et al. (1998) simulated the IOC by specifying an inflow through the southern boundary of their basin and an outflow through an idealized representation of the Indonesian passages north of the equator. They found pathways similar to those in the Shriver and Hurlburt (1996) solution, and an additional one involving upwelling in the subpolar region (as discussed above).

They also concluded that ²₃ of the EUC extratropical water was of southern-hemisphere origin because of the IOC.

The IOC thus forces a north-south asymmetry to the Pacific equatorial circulation. It intensifies the amount of southern-hemisphere water that upwells along the equator, and strengthens the southern STC relative to the northern one. As such, the IOC plays an analogous role to the deep thermocline cell in the Atlantic.

3. Dynamics

3.1 Subtropical intrusions: A key property of many (if not all) of the above studies is the existence of a Shadow-Zone characteristic (Luyten et al., 19xx) that extends equatorward from a point at the eastern boundary of the basin x_e and at the ventilation (or “subduction cutoff”) latitude y_d. A typical value for y_d is 15-20° north and south of the equator. The layer-2 flow to the east of this characteristic is at rest (i.e., the Shadow Zone). More interesting for our

purposes, though, is that subtropical water “intrudes” into the tropical ocean west of the streamline. The intrusion does not occur in a linear model, resulting from nonlinearities in the divergence term of the continuity equation. Although strictly incorrect (due to the failure of geostrophy), the characteristic can be extended to the equator. It intersects at a distance L from the eastern boundary given by

L= 1 2

′ g ₁₂

x (H₂+2H₁) (1)

where g ₁₂′ is the reduced gravity coefficient between layers 1 and 2, H₁ and H₁ are the

thicknesses of layers 1 and 2 at the eastern boundary of the basin, and ^x is the average zonal wind stress along L. Equation (1) provides insight into whether an interior pathway exists: If L is larger than the width of the basin, then the Shadow-Zone characteristic intersects the

western boundary before intersecting the equator and there is no interior pathway. Note that L depends on the stratification parameters g ₁₂′ and H₂, and that an interior pathway is more possible when these parameters are smaller. This implies that an interior pathway will exist at shallower levels in the real ocean, consistent with models and observations. Also note that L is smaller when ^x is larger, a consequence of the system then having a larger near-equatorial interface tilt, and the system therefore being more nonlinear.

3.2 Strength: The basic reason for the existence of STCs is that there is a net divergence of upper-layer water from the tropical ocean: This poleward volume transport in layer 1 must be compensated for by an equatorward volume transport in layer 2. McCreary and Lu (1994) derived a simple expression for the upper-layer divergence in their analytic model [see their equation (17)], showing that it was determined mostly by the Ekman transport across the tropical boundaries

E(y_d)= ^x

x_w x_e

∫

where y_d is the subduction cutoff latitude north or south of the equator, and x_w is the longitude of the western boundary.

According to (2), STC strength is not determined by midlatitude wind-stress curl, as might be expected. It is true that an increase in midlatitude wind curl increases subduction by a

proportionate amount, but the additional subducted water can recirculate within the Subtropical Gyres, and need not flow into the tropical ocean. It will do so only if the wind-curl anomaly also acts to increase the Ekman divergence out of the tropics.

4. Variability

The work summarized above only discusses steady-state solutions for the STCs and their dynamics. There have been almost no studies done to understand the processes that govern STC variability. Variability of the STCs has been proposed to be important for climate variability in two different ways. In one first scenario anomalous temperatures are advected by the mean STC (labelled v T ′ ), and in the other the mean temperature field is advected by variations in STC strength (v T ).′

Gu and Philander (1997) first proposed the former idea ( v T ′ ), hypothesizing that midlatitude SST anomalies generated by surface heat fluxes could be subducted into the subsurface branch of the STCs, advected to the equator, and upwelled there to affect the Pacific cold tongue. A number of modelling studies have explored this idea. They suggest that midlatitude SST anomalies are weakened by a variety of processes by the time they reach the equator, and hence that their effect on equatorial SST anomalies is quite small (e.g., Xie and Nonaka, 2000).

Kleeman et al. (1999) first proposed the latter idea (v T ), showing that changes in the wind′ stress along y_d altered the strength of the North Pacific STC in their coupled,

ocean-atmosphere model, leading to changes in the size and strength of the cold tongue. Klinger et al. (2000) used the Kleeman et al. (1999) ocean model to investigate wind-forced STC variability in greater detail, forcing the model with idealized wind patches of zonal wind that were either switched-on or oscillatory. Among other things, they find that significant equatorial SST anomalies are generated by changes in the zonal wind field along y_d, generally consistent with the measure of STC strength in (2). In further support of the idea, Nonaka (2000; see abstract) analyzed the output of a GCM forced by interannual winds, showing that at decadal time scales tropical SST anomalies are closely correlated with variability in the STC heat transport whereas the relationship did not exist at all at interannual time scales.