166
R. Stuhlmann
7.5 Effects of Clouds on the Atmospheric Radiation
Budget
As shown in many general circulation studies (e.g.: Johnson et al.. 1985; Oort and Peixoto.
1983). it is the vertical and horizontal distribution of the total diabatic heating of the atmosphere which determines the energy conversion and the dynamic structure of the atmosphere.
The full vertical profile of radiative flux divergence is also needed to determine the effect of radiation on the generation of Available Potential Energy (Lorenz. 1995; Stuhlmann and Smith.
1988). Thus, to understand how cloudiness will affect the general circulation and therefore
weather and climate. it is necessary to estimate the change in the three-dimensional structure
of the diabatic heating field as forced by a change in cloudiness.
The radiative heating of the total atmosphere and also the effect of clouds on that heating can
be derived from a combination of the reults for the top of the atmosphere and the surface.
Results on the shortwave cloud forcing of the atmosphere. C Fsw(A), are given by Laszlo
and Pinker (1993). Rieland and Stuhlmann (1993) and Rossow and Zhang (1995). Figure 7.12
presents the annual average shortwave atmospheric cloud forcing, C Fsw(A). determined for the
Meteosat field of view by Rieland and Stuhlmann (1993). The figure shows that the shortwave
atmospheric cloud forcing turns out to be about one order of magnitude less than that at the
top and falls inside a range between a warming of C Fsw(A) > 20 Wm- 2 and a cooling of
CFsw(A) < -5 Wm- 2 • A short-wave cooling of the total atmospheric column is found for
clouds within the tropics having high cloud tops. Here. the increase in backscattering of solar
radiation by these clouds, which takes place in the upper tropospheric levels. strongly reduces
the atmospheric short-wave absorption in comparison to a clear-sky situation. Contrary to this
fact, clouds with tops at lower levels cause a small increase in shortwave absorption.
Because the interaction of clouds with shortwave radiation can cool or heat the total atmosperic
column, CFsw(A) is not highly correlated with the mean cloud cover (R = 0.37). The largest
atmospheric shortwave heating of C Fsw( A) > 20 Wm -2. for example, is found for the Sahara.
an area with only a very low amount of cloudiness, but a very highly reflective surface. In
contrast. for the Atlantic ocean within the same latitudinal belt, a region with a high amount
of cloudiness. but a very small surface albedo, only a small shortwave heating of CFsw(A) <
10 Wm- 2 is found. As a result, it is concluded that the influence of clouds on CFsw(A) is
not determined by the cloud properties alone but in combination with the spectral surface
reflectance properties. An annual and regional average for the Meteosat field of view, ± 60°
longitude and latitude, of a warming of C Fsw(A) = 5 Wm- 2 is given by Rieland and Stuhlmann
(1993), which corresponds to a 1.4% increase in atmospheric solar absorption due to cloudiness.
Similar to the above. the longwave cloud forcing of the atmosphere. C FLW(A), depends on
the vertical distribution of the clouds within the atmosphere. The zonally averaged CFLW(A),
as calculated by Gupta et al. (1993) for January 1986 and July 1985. and presented in Figure
7.13. show a strong warming of the atmosphere in the tropics. the regions where deep convective
clouds are present. On the contrary, for the subtropics and midlatitudes of both hemispheres.
the regions with prevailing middle and low level clouds. a moderate cooling is derived. The
annual global average of CFLW(A) is calculated to be a cooling of -3.4 Wm- 2 (Gupta et
al., 1993). For hemispheric averages Gupta et al. found a strong seasonal change, which could
change, for the northern hemisphere example, from a warming of 4.7 Wm- 2 in July to a cooling
of -7.3 Wm- 2 in January. These changes are associated with the shift of the warm band in
the tropics across the equator. The cooling, which is found outside the tropics. undergoes only
minor seasonal variations.
R. Stuhlmann
7.5 Effects of Clouds on the Atmospheric Radiation
Budget
As shown in many general circulation studies (e.g.: Johnson et al.. 1985; Oort and Peixoto.
1983). it is the vertical and horizontal distribution of the total diabatic heating of the atmosphere which determines the energy conversion and the dynamic structure of the atmosphere.
The full vertical profile of radiative flux divergence is also needed to determine the effect of radiation on the generation of Available Potential Energy (Lorenz. 1995; Stuhlmann and Smith.
1988). Thus, to understand how cloudiness will affect the general circulation and therefore
weather and climate. it is necessary to estimate the change in the three-dimensional structure
of the diabatic heating field as forced by a change in cloudiness.
The radiative heating of the total atmosphere and also the effect of clouds on that heating can
be derived from a combination of the reults for the top of the atmosphere and the surface.
Results on the shortwave cloud forcing of the atmosphere. C Fsw(A), are given by Laszlo
and Pinker (1993). Rieland and Stuhlmann (1993) and Rossow and Zhang (1995). Figure 7.12
presents the annual average shortwave atmospheric cloud forcing, C Fsw(A). determined for the
Meteosat field of view by Rieland and Stuhlmann (1993). The figure shows that the shortwave
atmospheric cloud forcing turns out to be about one order of magnitude less than that at the
top and falls inside a range between a warming of C Fsw(A) > 20 Wm- 2 and a cooling of
CFsw(A) < -5 Wm- 2 • A short-wave cooling of the total atmospheric column is found for
clouds within the tropics having high cloud tops. Here. the increase in backscattering of solar
radiation by these clouds, which takes place in the upper tropospheric levels. strongly reduces
the atmospheric short-wave absorption in comparison to a clear-sky situation. Contrary to this
fact, clouds with tops at lower levels cause a small increase in shortwave absorption.
Because the interaction of clouds with shortwave radiation can cool or heat the total atmosperic
column, CFsw(A) is not highly correlated with the mean cloud cover (R = 0.37). The largest
atmospheric shortwave heating of C Fsw( A) > 20 Wm -2. for example, is found for the Sahara.
an area with only a very low amount of cloudiness, but a very highly reflective surface. In
contrast. for the Atlantic ocean within the same latitudinal belt, a region with a high amount
of cloudiness. but a very small surface albedo, only a small shortwave heating of CFsw(A) <
10 Wm- 2 is found. As a result, it is concluded that the influence of clouds on CFsw(A) is
not determined by the cloud properties alone but in combination with the spectral surface
reflectance properties. An annual and regional average for the Meteosat field of view, ± 60°
longitude and latitude, of a warming of C Fsw(A) = 5 Wm- 2 is given by Rieland and Stuhlmann
(1993), which corresponds to a 1.4% increase in atmospheric solar absorption due to cloudiness.
Similar to the above. the longwave cloud forcing of the atmosphere. C FLW(A), depends on
the vertical distribution of the clouds within the atmosphere. The zonally averaged CFLW(A),
as calculated by Gupta et al. (1993) for January 1986 and July 1985. and presented in Figure
7.13. show a strong warming of the atmosphere in the tropics. the regions where deep convective
clouds are present. On the contrary, for the subtropics and midlatitudes of both hemispheres.
the regions with prevailing middle and low level clouds. a moderate cooling is derived. The
annual global average of CFLW(A) is calculated to be a cooling of -3.4 Wm- 2 (Gupta et
al., 1993). For hemispheric averages Gupta et al. found a strong seasonal change, which could
change, for the northern hemisphere example, from a warming of 4.7 Wm- 2 in July to a cooling
of -7.3 Wm- 2 in January. These changes are associated with the shift of the warm band in
the tropics across the equator. The cooling, which is found outside the tropics. undergoes only
minor seasonal variations.