Latent and sensible heat
The ocean also transfers heat to the atmosphere by
the turbulent processes of conduction (sensible
heat flux) and evaporation (latent heat flux). The
sensible heat flux (Curry and Webster, 1999) is a
function of the temperature difference between the
ocean surface and the atmosphere immediately
above and is also limited by the rate at which heat
can be transferred to the air–sea interface by
processes within the atmospheric boundary layer.
Latent heat flux (Curry and Webster, 1999)
results from evaporation at the ocean surface. As
water changes from liquid in the ocean to vapour in
the atmosphere, the ocean loses energy in the form
of the latent heat of vaporization and the atmosphere gains heat when the vapour condenses. Since
water vapour pressure depends on temperature, sea
surface temperature is an important parameter in
determining the latent heat flux. As with the sensible heat flux, characteristics of the atmospheric
boundary layer (water vapour content, wind speed,
atmospheric stability) control the latent heat flux.
Net heat flux
The upper atmosphere receives more solar radiation at low than at high latitudes but the outgoing
radiation is more spatially uniform (Bryden and
Imawaki, Fig. 6.1.2). This requires a redistribution
of energy by the atmosphere and ocean. As a
result, there is a net heat flux into the ocean in the
equatorial regions and from the oceans at high latitudes. On the western side of ocean basins, where
swift western boundary currents carry warm water
polewards beneath a cooler atmosphere, there is
also a net heat loss by the ocean (Bryden and
Imawaki, Fig. 6.1.4). As with the surface wind
stress, quantitative descriptions of the air–sea heat
flux have been made using ship observations,
operational atmospheric analyses and satellite
observations. A comprehensive review of the available products is given by WGASF (2000). The
determination of heat (and fresh-water) transports,
which is a central objective of WOCE, will provide
constraints on the net air–sea heat (and freshwater) fluxes. Since these determinations will
require the availability of many key data sets, they
will only be able to be carried out in the later
stages of the WOCE project. Examples of earlier
determinations from both single zonal hydrographic sections and from global inversions are
given in Chapter 6.1.
1.2.2.3 Water exchange and surface salinity
Evaporation and precipitation move both heat and
fresh water between the ocean and the atmosphere. The rate of evaporation is directly proportional to the oceans’ latent heat loss and is
controlled by the ocean–atmosphere temperature
difference. Therefore the ocean influences evaporation through its surface temperature.
The difference between evaporation and precipitation (the evaporation minus precipitation (E9P)
flux, Wijffels, Fig. 6.2.1) changes the upper ocean
salinity. The highest rates of precipitation occur in
the tropics with secondary maxima at latitudes of
40 to 60° (Wijffels, Fig. 6.2.3). The main evaporative regions of the ocean are at about 20°N and S
(Wijffels, Fig. 6.2.2). Together with the addition of
fresh water through river discharge, spring runoff
from snow melt on land and through the melting
of sea ice, this results in a low-salinity surface
ocean in the equatorial and high-latitude regions
and high salinity near 20°N and S (Gordon,
Fig. 4.7.1b).
Neither evaporation nor precipitation is directly
affected by the surface salinity. With no direct
feedback, it is therefore a challenge to run coupled
atmosphere–ocean models that maintain upper
ocean salinity fields within climatological limits.
1.2.2.4 Particle exchanges
The formation of marine clouds and hence of precipitation depends on the presence of small particles to serve as condensation nuclei. Some of these
nuclei are provided by aerosols originating from
land-based natural and anthropogenic sources.
However, most are thought to arise from the ocean
when breaking waves release tiny subdroplets of
seawater into the atmospheric boundary layer. The
water content of these droplets rapidly evaporates,
leaving salt crystals that are carried aloft where
they can serve as condensation nuclei (Hudson
et al., 1998). In addition, some species of oceanic
phytoplankton produce organic sulphur compounds in the upper water column which are then
exchanged with the atmosphere. In the atmosphere
they are converted into small particles of dimethylsulphide, DMS, which also serve as condensation
nuclei. This exchange provides a potential climate
feedback if changes in ocean productivity lead
to changes in cloud cover and thus to changes in
the solar radiation reaching the ocean (Charlson
et al., 1987).
1.2 Ocean Processes and Climate Phenomena
15
Clarke, Church and Gould
The ocean also transfers heat to the atmosphere by
the turbulent processes of conduction (sensible
heat flux) and evaporation (latent heat flux). The
sensible heat flux (Curry and Webster, 1999) is a
function of the temperature difference between the
ocean surface and the atmosphere immediately
above and is also limited by the rate at which heat
can be transferred to the air–sea interface by
processes within the atmospheric boundary layer.
Latent heat flux (Curry and Webster, 1999)
results from evaporation at the ocean surface. As
water changes from liquid in the ocean to vapour in
the atmosphere, the ocean loses energy in the form
of the latent heat of vaporization and the atmosphere gains heat when the vapour condenses. Since
water vapour pressure depends on temperature, sea
surface temperature is an important parameter in
determining the latent heat flux. As with the sensible heat flux, characteristics of the atmospheric
boundary layer (water vapour content, wind speed,
atmospheric stability) control the latent heat flux.
Net heat flux
The upper atmosphere receives more solar radiation at low than at high latitudes but the outgoing
radiation is more spatially uniform (Bryden and
Imawaki, Fig. 6.1.2). This requires a redistribution
of energy by the atmosphere and ocean. As a
result, there is a net heat flux into the ocean in the
equatorial regions and from the oceans at high latitudes. On the western side of ocean basins, where
swift western boundary currents carry warm water
polewards beneath a cooler atmosphere, there is
also a net heat loss by the ocean (Bryden and
Imawaki, Fig. 6.1.4). As with the surface wind
stress, quantitative descriptions of the air–sea heat
flux have been made using ship observations,
operational atmospheric analyses and satellite
observations. A comprehensive review of the available products is given by WGASF (2000). The
determination of heat (and fresh-water) transports,
which is a central objective of WOCE, will provide
constraints on the net air–sea heat (and freshwater) fluxes. Since these determinations will
require the availability of many key data sets, they
will only be able to be carried out in the later
stages of the WOCE project. Examples of earlier
determinations from both single zonal hydrographic sections and from global inversions are
given in Chapter 6.1.
1.2.2.3 Water exchange and surface salinity
Evaporation and precipitation move both heat and
fresh water between the ocean and the atmosphere. The rate of evaporation is directly proportional to the oceans’ latent heat loss and is
controlled by the ocean–atmosphere temperature
difference. Therefore the ocean influences evaporation through its surface temperature.
The difference between evaporation and precipitation (the evaporation minus precipitation (E9P)
flux, Wijffels, Fig. 6.2.1) changes the upper ocean
salinity. The highest rates of precipitation occur in
the tropics with secondary maxima at latitudes of
40 to 60° (Wijffels, Fig. 6.2.3). The main evaporative regions of the ocean are at about 20°N and S
(Wijffels, Fig. 6.2.2). Together with the addition of
fresh water through river discharge, spring runoff
from snow melt on land and through the melting
of sea ice, this results in a low-salinity surface
ocean in the equatorial and high-latitude regions
and high salinity near 20°N and S (Gordon,
Fig. 4.7.1b).
Neither evaporation nor precipitation is directly
affected by the surface salinity. With no direct
feedback, it is therefore a challenge to run coupled
atmosphere–ocean models that maintain upper
ocean salinity fields within climatological limits.
1.2.2.4 Particle exchanges
The formation of marine clouds and hence of precipitation depends on the presence of small particles to serve as condensation nuclei. Some of these
nuclei are provided by aerosols originating from
land-based natural and anthropogenic sources.
However, most are thought to arise from the ocean
when breaking waves release tiny subdroplets of
seawater into the atmospheric boundary layer. The
water content of these droplets rapidly evaporates,
leaving salt crystals that are carried aloft where
they can serve as condensation nuclei (Hudson
et al., 1998). In addition, some species of oceanic
phytoplankton produce organic sulphur compounds in the upper water column which are then
exchanged with the atmosphere. In the atmosphere
they are converted into small particles of dimethylsulphide, DMS, which also serve as condensation
nuclei. This exchange provides a potential climate
feedback if changes in ocean productivity lead
to changes in cloud cover and thus to changes in
the solar radiation reaching the ocean (Charlson
et al., 1987).
1.2 Ocean Processes and Climate Phenomena
15
Clarke, Church and Gould