In order to study the effect of the atmosphere on the outcoming longwave radiation, it is convenient to subdivide it into a ``clean dry'' component, water vapor, and aerosols (water droplets, as well as ice crystals, salt grains & dust particles, which serve as condensation seeds for water).
Table 10.1 gives the standard composition of the ``clean dry'' air in the troposphere (i.e at altitudes km). Except for CO, whose abundance at ground level may vary between day and night by up to a factor of 2, this composition is remarkably homogeneous and constant. Other trace components, most of which are unstable (SO, O, NO, CO, ...) have abundances (in volume) that never exceeds 10.
The abundance of water is highly variable, but hardly exceeds 1% in mass, even locally. Most of the water in the air is in the form of vapor. Even inside the clouds, precipitation and turbulence insure that the mass of water droplets per cm seldom equals that of water vapor. In addition the water vapor mixing ratios above 15-20 km are under 10 ppmv making it just another trace gas there.
Ozone and to much less extent molecular Oxygen are responsible for most of the absorption of the solar radiation in the UV, especially between 180 and 290 nm, thanks to these processes: (Hartley band), (Schuman-Runge band), and (Herzberg band),
In the visible, the air is fairly transparent except for scattering by aerosols, mostly water droplets, ice crystals and dust particles. In the infrared, HO, CO (around 15 m) and O (around 10m) are very efficient absorbers of the solar and ground radiation, to the extent that they prevent ground-based observations in large regions of the electromagnetic spectrum.
By clear weather, the atmospheric absorption at millimeter and submillimeter wavelengths is dominated by rotational and fine structure lines of molecules in their ground electronic and low vibrational states.
The strongest molecular rotational resonances appear in polar molecules (HO and O being the most important of such molecules in the atmosphere) and are of the type E1 (electric dipole transitions). Intrinsically weaker M1 (magnetic dipole) transitions are of considerable practical importance in the atmosphere due to the high abundance of O. We will see that collision induced E2 (electric quadrupole) absorption involving N and O is measurable in the atmosphere. The different atmospheric hydrometeors (water droplets, snow, graupel, hale, ice cristals,...) scatter and absorb following different patterns across the longwave (radio to submm) spectral region. All the mechanisms involved in the radiative transfer of longwave radiation in the atmosphere will be described in this chapter.
A mixture of ideal gases behaves like an ideal gas:
Dry air is a mixture of N, O, ... molecules. It behaves indeed very much like an ideal gas: J/mol-deg (vs 8.3149 for an ideal gas), = 1.404 (vs 1.400 for ideal rigid molecules).
Wet air (without clouds) is a mixture of dry air + HO molecules. It is customary to denote by the partial pressure of water vapor, that of dry air, and the wet air pressure. The specific heats of water vapor are not that different from those of ideal gases: ; , vs and for a rigid asymmetric top.
Then, Dalton's law yields:
Then, for the adiabatic expansion of a wet air bubble, one has:
At large scales, the air pressure and density depend essentially on the massive and slowly varying dry component and are well described by hydrostatic equilibrium. The air temperature, as we have seen, depends significantly on the abundance and distribution of water, CO (and O for the stratosphere).
For larger altitudes, from Eq.10.7 and , then:
Although the above equations represent fairly well the density and pressure throughout the troposphere, the temperature distribution can depart significantly from the above linear variation near the ground. This ground heats up faster than the transparent air during the day, and cools off more rapidly during the night. The temperature gradient at low altitudes (up to 1-2 km) can be thus steeper or shallower than described by (Kkm). Occasionally, it can be inverted, the temperature increasing with altitude. The inversion layer usually stops briskly at 1 or 2 km altitude and the normal temperature gradient resumes above. Inversion layers are common during the night over bare land. They can also be caused by hot winds blowing from the sea.
The local temperature gradient determines stability of the air to vertical motions. A rising bubble of wet air expands adiabatically as long as the water vapor it contains does not condense. Expanding, it cools almost as an ideal gas with:
The scale height of water, , which results from a fast evaporation/condensation process, is small (typically km) compared to the equilibrium scale height km. At km, the altitude of the Plateau de Bure, the water vapor column density (or , ``amount of precipitable water'', when expressed in g.cm, or cm of water) is normally reduced by a factor of 3-4, with respect to sea level. This factor, as we have seen, is strongly modulated by the local temperature gradient. is lower in the presence of a low altitude inversion layer which reduces the vertical turbulence and traps most of the water well below the observatory.