Convec,on, cloud and radia,on Convection redistributes the thermal energy yielding (globally-averaged), a mean lapse rate of ~ -6.5 o C/km. Radiative processes tend to produce a more negative temperature gradient (cooling). Nonradiative upward heat transfer opposes such cooling. The smaller lapse rate, -6.5 o C/km, reflects both the importance of convection (and water condensation in particular) and the fact that the atmosphere is, on mean, stable. 1
Hartmann 2
In Fig. 3.17 from Hartmann the importance of CO 2, H 2 O, and O 3, are illustrated for a clear sky calculation. 3
Clouds & Radia,on Clouds are the Achilles heal of climate modeling. The optical properties of clouds vary substantially with both the amount of water in the cloud, the size and shape of the cloud particles / droplets, and their distribution in space. Water droplets and ice interact with both long and short wave radiation. In figure 3.13 of Hartmann, the dependence of cloud albedo and cloud absorption on both the liquid water content and SZA is shown for a highly idealized cloud (uniform and horizontally infinite). Note that the albedo saturates for liquid water paths of ~ 1000 g m -2. These are dark clouds and any additional radiation scattered within the cloud is unlikely to make it back out the top. Absorption of solar radiation decreases with SZA because of the increasing albedo. Radia,ve Proper,es of Cloud 4
In figure 3.14 of Hartmann, the calculated albedo is shown for a cloud with the same liquid water content by varying cloud droplet size. Scattering, which to first order depends on surface area, increases with decreasing cloud droplet size. In addition, the cloud lifetime generally also increases with decreasing cloud droplet size. Clouds absorb terrestrial radiation very effectively (why?). Figure 3.15 of Hartmann shows the emissivity of water and ice clouds as a function of liquid water content. At 20 g m -2 clouds become opaque to longwave radiation. Cirrus (wispy clouds at high altitude) are sometimes not optically deep in the IR. Note that cloud albedo continues to increase long after they have become opaque in the IR. What does this mean for the greenhouse effect of clouds? Why does a low cloud cool the surface while a thin high cloud warms the surface? 5
Clouds of sufficient thickness are essentially perfect absorbers of the thermal IR (3-100 um). At the same time, clouds act to increase the albedo. Thus clouds have both cooling and warming properties (a cloudy day is cooler than a clear day; a cloudy night can be much warmer than a clear night). In figure 3.19 of Hartmann, a simple radiative-convective model is used to illustrate the influence of clouds. Remember that in general as water content grows, clouds become essentially perfect blackbodies in the IR long before they become perfect reflectors of solar radiation. As a result, clouds that form at lower altitude, where more water is available, tend to have higher albedos. In addition, the longwave forcing by clouds depends on their temperature. A warm cloud near the surface looks just like a warm surface; a cold high cloud emits much less IR. To be a bit more quantitative, we can write down the radiation balance at the top of the atmosphere for clear and cloudy conditions. This energy balance is the difference between the absorbed solar radiation and the outgoing longwave radiation (OLR): In our simple slab atmosphere, R TOA = 0. Here, however, we would like to calculate the difference in R TOA between clear and cloudy atmosphere. If we know the albedo for both clear and cloudy conditions, the difference in Q is: ΔQ abs = - (S o /4) Δα p. 6
On the longwave side, let s assume for the moment that the cloud is a perfect blackbody and that the transmission above cloud height is perfect. This approximation is clearly not perfect, but for clouds above ~ 5 km, it is not terrible. With this approximation, ΔF (infinity) = σt 4 ct F clear(infinity) With these simplifications, ΔR TOA = - (S o /4) Δα p - σt 4 ct + F clear(infinity) Thus, the change in net radiation depends on the contrast between the clear and cloudy albedo and the temperature (and thus height) of the cloud. If the cloud top is at just the right height, the change in albedo will cancel the change in longwave forcing and the surface temperature will not change. For ΔR TOA = 0 we can solve for the cloud temperature (and height if we assume Γ = -6.5 o C/km and a surface temperature). The result is shown in Figure 3.20. Note how large these values are (both + and - ). For comparison, the radiative forcing due to change in CO 2 from 1750 is 1.5 W m -2. 7