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E. Linacre and B. Geerts |
3/’02 |
Climate changes on the scale of several decades to millenia are strongly controlled by surface and deep ocean currents. For instance, in Europe the Ice Age cooling was larger than the global mean, due to a southward shift of the westward flow in the south Atlantic ocean gyre. Part of this surface current is planed off by the northeastern corner of Brazil and then goes into the Mexican Gulf, becoming the Gulf Stream. The southward shift of the westward flow reduces the fraction forming the Gulf Stream, weakening it, so that it is less able to warm Europe. Another example is the change in deep-water formation off Greenland at 11,500 aBP, which caused sudden cooling in northwestern Europe (heralding a brief period known as the lower Dryas).
Deep ocean currents are usually buoyant (i.e. thermohaline, driven by differences of temperature and salinity) (1), though the upwelling of deep water is generally forced by surface winds. Deep water is formed when dense water at less than 0° C sinks near Greenland and Antarctica, and it wells up mainly along eastern boundaries of ocean basins around 30° latitude (Section 11.5). The 'age' of deep water is the time since the water mass was last at the surface. This age can be estimated fairly accurately from the chemical composition and isotope ratios. Along the east coast of the Americas the deep water has an age increasing from 4 years just south of Greenland, to 19 years off Florida, to 35 years off Rio de Janeiro (2). An important deep-ocean conveyor belt is shown in Fig 11.19 of the textbook, but there is no single ‘conveyor belt’ circulation; upwelling occurs in several areas. Deep-ocean circulations, and especially their long-term variations, are still poorly understood.
The most important deep water masses are the Subantarctic Mode Water (SAMW), the Antarctic Intermediate Water (AIW), both formed north of the east-bound Antarctic Circumpolar Current, and, thirdly, the Antarctic Bottom Water (ABW), formed on the Antarctic continental shelf (Fig. 11.18 in the book). The AIW and ABW masses penetrate into the northern hemisphere. More rapid climate changes over no more than a few years, or one or two decades, are governed by near-surface and intermediate-depth currents within the oceans, influenced by surface winds, and by SAMW and AIW (3).
Oceans store a large amount of heat, so that small changes in ocean currents can have a large effect on coastal and global climate. Currents carry enormous amounts of heat north and south. The magnitude of the heat fluxes in the various ocean basins is shown in Table 1.
|
meridional heat flux (1015 watts) |
at 30° S |
at 30° N |
|
|
Atlantic Ocean |
-0.4* |
1.2 |
*a negative heat flux is towards the equator |
|
Indian Ocean |
1.7 |
none |
|
|
Pacific Ocean |
-0.4* |
0.8 |
|
|
total |
0.9 |
2.0 |
|
Table 1: Meridional net heat flux (in 1015 watts) at 30° latitude in the oceans. (3)
A convergence zone occurs within the westerly Antarctic Circumpolar Current, known as the Antarctic Convergence Zone (Fig 11.18 in the book). The sea surface temperature has a large gradient along this zone, and therefore transient mesoscale eddies form along the boundary. There are also much larger, longer-lived SST anomalies. These anomalies travel at about 7 cm/s, taking 8-10 years to circle the pole, presumably influencing atmospheric conditions to the north as they go (3). They affect sea-level air pressure anomalies, as well as sea-ice coverage.
The interannual variation of climate in Australia, Indonesia and Southeast Asia is affected by the ENSO. But there are three other aspects of the ocean circulation that have an important impact on temperature and mainly rainfall anomalies in this region.
2.1 Indonesian ThroughflowRegional climates are influenced by sea-surface temperatures, which depend on ocean currents. Thus rainfall amounts in Australasia are affected by the amount of ‘Indonesian Throughflow’, the widely variable warm current of around 5 Sverdrups (Sv)* eastward on the north side of the island of Flores, then westward on each side of Timor, into the Indian Ocean. The SST there influences the temperature of west winds onto Australia.
2.2 Indonesia/Indian Ocean DipoleAny fluctuation of SST’s in the central Indian Ocean (around 13°S, 80°E) is often in opposition to the SST variation between northwest Australia and Indonesia, in other words there is a see-saw pattern or dipole. What is regarded as a warm polarity of this Indonesia/Indian Ocean Dipole (IIOD) means warmth near Indonesia (and therefore coolness in the central Indian Ocean), which leads to increased rainfall over northern and eastern Australia. This is particularly evident in the southern winter.
Cool polarity at the same time as an El Niño causes especially severe droughts in Australia, as in 1982 and 1994. Cool polarity is made more likely by ocean upwelling off the south coast of Java (due to local winds) and by a reduced Indonesian Throughflow (due to weaker Trade winds in the Pacific).
2.3 East Australian CurrentThe East Australian Current runs along the east coast of Australia. This current amounts to about 27 Sv to the south across 28°S, partly offset by 18 Sv northwards, further east in the Tasman Sea, around 160°E. However, the EAC varies between seasons and from year to year.
*1 Sv is a million cubic metres per second; the Antarctic Circumpolar Current flows at 130 Sv.
References
(1) Rintoul, S., G. Meyers, J. Church, S. Godfrey, M. Moore & B. Stanton 1997 Ocean Processes, climate and sea level. In ???? (ed.) 1997 Greenhouse: coping with climate change ( ??), 127-44.
(2) Gould, J. & J. Church 1996. Oceans and climate. Physics World (Dec), 33-7.
(3) Church, J. 1997. Observations of DecCen variability in the southern hemisphere oceans. Publ. series 11 (ICPO, World Climate Res. Prog., World Meteor. Organ.), 40-52.