The atmosphere

The major constituents of the present atmosphere are nitrogen, oxygen and water. For dry air at sea-level the nitrogen constitutes 78% by volume, oxygen 21% and minor constituents 1%, the major part of which is argon (0.93%)

and carbon dioxide (0.03%). The water content by volume ranges from close to zero at the poles to about 4% in tropical climates.

Typical variations of density, pressure and temperature, as functions of height, are shown in Fig. 2.27. Common names for the different layers are indicated. They are generally defined by turning points in the temperature profile ("pauses") or by the lower border of constant temperature regions if no sharp change in the sign of the temperature gradient takes place. Actual temperature profiles are not independent of latitude and season and do not exhibit sharp "turning points", as indicated in Fig. 2.28.

Estimates of the variation with height of the average abundance of some gaseous constituents of the atmosphere are shown in Fig. 2.29. The increasing abundance of ozone (O3) between 15 and 25 km height is instrumental in keeping the level of ultraviolet radiation down, thereby allowing life to develop on the Earth's surface (cf. Fig. 2.2). Above a height of about 80 km, many of the gases are ionised, and for this reason the part of the atmosphere above this level is sometimes referred to as the ionosphere. Charged particles at the height of 60 000 km are still largely following the motion of the Earth, with their motion being mainly determined by the magnetic field of the Earth. This layer is referred to as the magnetosphere.

Figure 2.27. Altitude dependence of temperature, pressure and density for the US standard atmosphere (U.S. Government, 1962).

Particles in the atmosphere

In addition to gaseous constituents, water and ice, the atmosphere contains particulate matter in varying quantity and composition. A number of mechanisms have been suggested which produce particles varying in size from 10-9 to 10-4 m. One is the ejection of particles from the tops of the oceans, in the form of sea salt sprays. Others involve the transformation of gases into particles. This can occur as a result of photochemical reactions or as a result of adsorption to water droplets (clouds, haze), chemical reactions depending on the presence of water and finally evaporation of the water, leaving behind a suspended particle rather than a gas.

Temperature (deg. K)

Temperature (deg. K)

Latitude (deg)

Figure 2.28a,b. Average temperature in the atmosphere as a function of height and latitude for January (above) and July (below) 1995, both at longitude zero (based on NCEP-NCAR, 1998).

Temperature (deg. K)

Figure 2.28a,b. Average temperature in the atmosphere as a function of height and latitude for January (above) and July (below) 1995, both at longitude zero (based on NCEP-NCAR, 1998).

Temperature (deg. K)

Latitude (deg)
Figure 2.29. Estimated mean concentrations of some gaseous constituents of the atmosphere, as functions of height. (Based on Almquist, 1974).

Figure 2.30. Trends of particle concentrations as functions of height, either for large particles only (left) or for all particles. The range of surface concentrations over different surfaces is indicated. "Background" refers to an extrapolated, unpolluted and clean surface, such as the Arctic ice covers. (Based on Craig, 1965; Wilson and Matthews, 1971).

Figure 2.30. Trends of particle concentrations as functions of height, either for large particles only (left) or for all particles. The range of surface concentrations over different surfaces is indicated. "Background" refers to an extrapolated, unpolluted and clean surface, such as the Arctic ice covers. (Based on Craig, 1965; Wilson and Matthews, 1971).

Dust particles from deserts or eroded soil also contribute to the particle concentration in the atmosphere, as does volcanic action. Man-made contributions are in the form of injected particles, from land being uncovered (deforestation, agricultural practice), industry and fuel burning, and in the form of gaseous emissions (SO2, H2S, NH3, etc.) which may later be transformed into particles by the above-mentioned processes or by catalytic transformations in the presence of heavy metal ions, for example.

Wilson and Matthews (1971) estimate that the yearly emission into or formation within the atmosphere of particles with a radius smaller than 2 X 10-5 m is in the range (0.9-2.6) X 1012 kg. Of this, (0.4-1.1) X 1012 kg is due to erosion, forest fires, sea salt sprays and volcanic debris, while (0.3-1.1) X 1012 kg is due to transformation of non-anthropogenic gaseous emissions into particles. Anthropogenic emissions are presently in the range (0.2-0.4) X 1012 kg, most of which is transformed from gaseous emissions.

Figure 2.31. Size distribution of particles in the atmosphere. The estimated distribution over land is shown separately. The distribution over oceans resembles that of the clean "background" distribution, except in the radius interval 10-6-10-5 m, where it approaches that for continental particles. (Based on Wilson and Matthews, 1971).

Figure 2.30 gives a very rough sketch of the height distribution of the particle number density. The distribution of large particles (radii in the range 10-7-2 X 10-5 m) is based on measurements at a given time and place (based on Craig, 1965), while the total particle counts are based on extrapolations of the size distribution to very small particles (see Fig. 2.31). It is clear that the total number of particles depends sensitively on the density of very small particles, which on the other hand do not contribute significantly to the total mass of particulate matter. The concentrations of rural and city particles in the right-hand part of Fig. 2.30 correspond to an extrapolation towards smaller radii of the size distributions in Fig. 2.31, such that the number of particles per m3 per decade of radius does not decrease rapidly towards zero.

Returning to mass units, the total mass of particles in the atmosphere can be estimated from the emission and formation rates, if the mean residence time is known. The removal processes active in the troposphere include dry deposition by sedimentation (important only for particles with a radius above 10-6 m), diffusion and impact on the Earth's surface, vegetation, etc., and wet deposition by precipitation. The physical processes involved in wet deposition are condensation of water vapour on particles, in-cloud adsorption of particles to water droplets (called rainout), or scavenging by falling rain (or ice) from higher-lying clouds (called washout). Typically, residence times for particles in the troposphere are in the range 70-500 h (Junge, 1963), leading to the estimate of a total particle content in the atmosphere of the order of 1014 kg.

Particles may also be removed by chemical or photochemical processes, transforming them into gases.

Figure 2.32. Transmittance of solar radiation in the direction of incidence, as measured in Hawaii near sunrise, during the periods before and after the volcanic eruption of Mt Agung in Bali. Other smaller volcanic eruptions in the following years are indicated by arrows. The curves have been smoothed over seasonal and other short-term variations, with the fluctuations being generally less than ±0.05 in the transmittance on clear days. (Based on Ellis and Pueschel, 1971).

Figure 2.32. Transmittance of solar radiation in the direction of incidence, as measured in Hawaii near sunrise, during the periods before and after the volcanic eruption of Mt Agung in Bali. Other smaller volcanic eruptions in the following years are indicated by arrows. The curves have been smoothed over seasonal and other short-term variations, with the fluctuations being generally less than ±0.05 in the transmittance on clear days. (Based on Ellis and Pueschel, 1971).

Particle residence times in the stratosphere are much longer than in the troposphere. Sulphates (e.g. from volcanic debris) dominate, and the particles are spread rather homogeneously over each hemisphere, no matter where they are being injected into the stratosphere. Residence times are of the order of years, and the removal seems to be by diffusion into the troposphere at latitudes around 55°. This picture is based on evidence from major volcanic eruption and from detonation of nuclear weapons of megaton size.

Both of these inject particulate matter into the stratosphere. The particles in the stratosphere can be detected from the ground, because they modify the transmission of solar radiation, particularly close to sunrise or sunset, where the path-length through the atmosphere is at its maximum. Figure 2.32 shows atmospheric transmittance at sunrise, measured in Hawaii before and after the major volcanic eruption of Mt. Agung on Bali in 1963. The arrows indicate other eruptions which have taken place near the Equator during the following years and which may have delayed the return of the transmittance towards its value from before 1963. Figure 2.33 shows the ground deposition of radioactive debris (fallout), at a latitude of 56°N, after the enforcement of the test ban in 1963, which greatly reduced but did not fully end nuclear tests in the atmosphere. It is concluded that the amount of radioactivity residing in the stratosphere is reduced to half each year without new injection. Observations of increased scattering (resulting in a shift towards longer wavelengths) following the very large volcanic eruption of Krakatoa in 1883 have been reported (sky reddening at sunrise and sunset). It is believed that the influence of such eruptions on the net radiation flux has had significant, although in the case of Krakatoa apparently transient, effects on the climate of the Earth.

Figure 2.33. Yearly fallout of 90Sr following the atmospheric bomb test ban in 1963, as measured in Denmark. (Based on Aarkrog, 1971).

Absorption and scattering in the atmosphere

Absorption and scattering of radiation take place throughout the atmosphere. The radiation may cause electrons in atoms to go from one orbital to another, or it may cause transition between vibrational or rotational levels in molecules. Generally, the fundamental energy differences associated with rotational excitations are smaller than those associated with vibrational excitations, which again are smaller than those associated with low-lying electron states. Only with a fairly small probability is the solar radiation capable of ionising atoms (expelling electrons) or dissociating molecules.

Since the wavelengths in the solar spectrum are much longer than the dimensions of atoms and molecules found in the atmosphere, the time-dependent fields of the atomic or molecular system can be approximated by the first few terms in a multipole expansion. In the dipole approximation, the probability of a transition between two levels i and f, accompanied by the emission or absorption of radiation with frequency equal to (or very close to)

Vf = lEf-EI/h, where Ef is the energy of level if), is proportional to the square of the dipole transition moment df (an integral over the initial and final state quantal wavefunctions times the position vector; see e.g. Merzbacker, 1970). The absorption probability (or "rate") is further proportional to the spectral intensity at the required frequency.

"Rate of absorption" Idfl2I(v¡f).

If the absorption and re-emission processes were sufficiently frequent, a state of thermal equilibrium would be reached, in the sense that the relative population of any two atomic or molecular levels would be given by the Boltzmann factor exp(-hVf/(kT)). If the radiation has a frequency distribution corresponding to the Planck law, the common temperature T will be that appearing in the Planck distribution for the radiation. In this way a single atom or molecule can be in equilibrium with radiation and can be ascribed the same temperature as that of the radiation.

On the other hand, the thermodynamic temperature is defined from the distribution of kinetic energy in the "external" motion of the atoms or molecules. This distribution is maintained by collisions between the atoms or molecules. If the frequency of collisions is comparable to that of interactions with the radiation, the kinetic and radiation temperatures will become similar (e.g. the radiative transfer section of the Sun).

Where collisions are much more frequent than interactions with the radiation field, the radiative energy absorbed by a given atomic or molecular transition will become shared among all the atoms or molecules, and if the corresponding emission spectrum can be described by the black-body law, the temperature entering should be the kinetic temperature, which need not be equal to the temperature of the radiation field (in fact, the kinetic temperature must be lower than the radiation temperature, if there are no other energy sources). Re-emission at frequencies corresponding to the same definite atomic or molecular transitions that caused absorption of radiation can be regarded as scattering of the incoming radiation.

In the Earth's atmosphere, the number of molecular collisions is generally large compared with the number of interactions with the solar radiation. Thus, the radiation which is not re-emitted as scattered light will be redistributed over a large number of thermal degrees of freedom, i.e. the absorbed light is transformed into heat. The population of various molecular levels will not be in equilibrium with the temperature of the solar radiation, and the spectrum of scattered light will not be of Planck form (cf. Fig. 2.13).

The scattering process may be viewed as the creation of a dipole field in the atom or molecule, as a result of the radiation field, and a subsequent emission of radiation, as is well known from an oscillating dipole. The angular distribution of scattered light is proportional to (1+cos2y), where y is the angle between incident and scattered light.

Absorption processes in different frequency regions

Even in the thermosphere, most ultraviolet radiation with wavelength below 1.8 X 10-7 m is already being absorbed by N2O and O2. Owing to its low density, the thermosphere exhibits a rapid temperature response to variations in solar intensity, such as those caused by sunspots. However, for the same reason, these temperature changes have little or no effect on the lowerlying layers of the atmosphere. Figure 2.34 shows the penetration depth of ultraviolet radiation, confirming that the shortest wavelengths are being stopped at a height of 100 km or more.

Figure 2.34. The altitude, as a function of wavelength, at which the intensity of solar radiation is reduced to l/e. The main absorbing agents are indicated for different regions of the ultraviolet part of the spectrum. (Based on Ratcliffe, 1960).

This figure also shows that most of the radiation in the wavelength interval 1.8 X 10-7 to 3 X 10-7m is absorbed in the mesosphere or the stratosphere. The agents of absorption are O2 and in particular O3. The rate of absorption by ozone peaks around 2.5 X 10-7m wavelength, and the amount of energy converted to heat is sufficient to explain the peak in the temperature profile at about 50 km (see Fig. 2.28).

The ozone concentration is not uniform as a function of geographical position, as indicated by Fig. 2.35. Ozone is formed in the stratosphere by photodissociation of molecular oxygen,

Figure 2.35a,b. Daily satellite measurements of ozone distributions for 1 January (above) and 1 April (below) 1997. The figures indicated are the volumes (in Dobson units, 10-5 m3 m-2) which the ozone would occupy at sea-level standard temperature and pressure, if all the ozone in a vertical column were put together. Areas obscured from the satellite are shown as white (NASA, 1998).

Figure 2.35a,b. Daily satellite measurements of ozone distributions for 1 January (above) and 1 April (below) 1997. The figures indicated are the volumes (in Dobson units, 10-5 m3 m-2) which the ozone would occupy at sea-level standard temperature and pressure, if all the ozone in a vertical column were put together. Areas obscured from the satellite are shown as white (NASA, 1998).

(Chapman, 1943). Natural removal processes include

(Hesstvedt, 1973).

Wavelength (I0~6m) 198 I 99 2 00

Detail f spectrum mI^/MHJV

o 1_I

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