As discussed in Sections 1.3.2 and 4.5, the moisture distribution in the atmosphere is strongly controlled by the temperature distribution; the atmosphere is moist near the surface in the tropics where it is very warm and drier aloft and in polar latitudes where it is cold. As shown in Fig. 5.15, the specific humidity, defined in Eq. 4-23, reaches a maximum (of around 18gkg-1) at the surface near the equator and decreases

The control by temperature of the specific humidity distribution can be seen more directly by comparing Fig. 5.15 with Fig. 5.16, which shows q*, the specific humidity at saturation given by Eq. 4-24 with es given by Eq. 1-4. We see that q has the same spatial form as q* but never reaches saturation even at the surface. As discussed in Section 4.5, U, the relative humidity defined in Eq. 4-25, is the ratio of q in Fig. 5.15 to q* in Fig. 5.16. Zonal mean relative humidity, shown in Fig. 5.17, is (on average) 70-80% everywhere near the ground. The reason for the decrease of relative humidity with altitude is a little more subtle. Vertical transport of water vapor is effected mostly by convection, which (as we have seen) lifts the air to saturation. It may therefore seem odd that even the relative humidity decreases significantly with height through the troposphere. To understand this, we need to think about the entire circulation of a convective system, and not just the updraft. Consider Fig. 5.18.

The updraft in a convective cloud—the part considered in the parcel stability argument of Section 4.5.2—is rather narrow. Of course, the air must return and does so

Zonal-Average Specific Humidity (g/kg)

FIGURE 5.15. Zonally averaged specific humidity q, Eq. 4-23, in gkg 1 under annual mean conditions. Note that almost all the water vapor in the atmosphere is found where T > 0°C (see Fig. 5.7).
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