11.1.1. Heat, freshwater, and buoyancy fluxes
As discussed in Chapter 4, atmospheric convection is triggered by warming at the surface. Vertical mass transport is confined to a few regions of strong updrafts driven by deep convection over the warmest oceans and land masses in the tropics, with broader areas of subsidence in between (see the atmospheric mean meridional circulation plotted in Fig. 5.21). In contrast to the atmosphere, the ocean is forced from above by air-sea fluxes. We therefore expect ocean convection to be most prevalent in the coldest regions, where the interior stratification is small, most likely at high latitudes in winter where surface density can increase through:
1. direct cooling, reducing temperature and hence increasing density.
2. brine rejection in ice formation, thus increasing salinity (and hence density) of the water immediately below the ice.
Whether a parcel of water sinks depends on its buoyancy anomaly (as described in Section 4.2.1), defined by Eq. 4-3, which we write out again here adopting our oceanographic notation:
where g is the acceleration due to gravity, and a - ao is the difference between the density of the parcel and its surroundings (see Eq. 9-5). As discussed in Section 9.1.3, the buoyancy of seawater at the surface depends on both the T and S distribution. To determine whether convection will occur we must therefore consider both the flux of heat and freshwater across the ocean surface, which induce T, S and hence buoyancy changes, as well as the ambient, pre-existing stratification of the water column.
The equations governing the evolution of T and S are:
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