Atmosfera, Tempo e Clima

Atmosfera, Tempo e Clima

(Parte 6 de 9)

The light gases (hydrogen and helium especially) might be expected to become more abundant in the upper atmosphere, but large-scale turbulent mixing of the atmosphere prevents such diffusive separation up to at least 100 km above the surface. The height variations that do occur are related to the source locations of the two major non-permanent gases – water vapour and ozone. Since both absorb some solar and terrestrial radiation, the heat budget and vertical temperature structure of the atmosphere are affected considerably by the distribution of these two gases.

Water vapour comprises up to 4 per cent of the atmosphere by volume (about 3 per cent by weight) near the surface, but only 3 to 6 ppmv (parts per million by volume) above 10 to 12 km. It is supplied to the atmosphere by evaporation from surface water or by transpiration from plants and is transferred upwards by atmospheric turbulence. Turbulence is most effective

Figure 2.1Atmospheric particles. (A) Mass distribution, together with a depiction of the surface–atmosphere processes that create and modify atmospheric aerosols, illustrating the three size modes. Aitken nuclei are solid and liquid particles that act as condensation nuclei and capture ions, thus playing a role in cloud electrification. (B) Distribution of surface area per unit volume.

Sources: (A) After Glenn E. Shaw, University of Alaska, Geophysics Institute. (B) After Slinn (1983).

below about 10 or 15 km and as the maximum possible water vapour density of cold air is very low anyway (see B.2, this chapter), there is little water vapour in the upper layers of the atmosphere.

Ozone (O3) is concentrated mainly between 15 and 35 km. The upper layers of the atmosphere are irradiated by ultraviolet radiation from the sun (see C.1, this chapter), which causes the breakup of oxygen molecules

at altitudes above 30 km (i.e. O2→O O). These separated atoms (O + O) may then combine individu- ally with other oxygen molecules to create ozone, as illustrated by the simple photochemical scheme:

where M represents the energy and momentum balance provided by collision with a third atom or molecule; this Chapman cycle is shown schematically in Figure 2.2A. Such three-body collisions are rare at 80 to 100 km because of the very low density of the atmosphere, while below about 35 km most of the incoming ultraviolet radiation has already been absorbed at higher levels. Therefore ozone is formed mainly between

30 and 60 km, where collisions between O and O2are more likely. Ozone itself is unstable; its abundance is determined by three different photochemical interactions. Above 40 km odd oxygen is destroyed primarily by a cycle involving molecular oxygen; between 20 and 40 km NOxcycles are dominant; while below 20 km a hydrogen–oxygen radical (HO2) is responsible. Additional important cycles involve

chlorine (ClO) and bromine (BrO) chains at various altitudes. Collisions with monatomic oxygen may recreate oxygen (see Figure 2.2B), but ozone is destroyed mainly through cycles involving catalytic reactions, some of which are photochemical associated with longer wavelength ultraviolet radiation (2.3 to 2.9 µm). The destruction of ozone involves a recombination with atomic oxygen, causing a net loss of the odd oxygen. This takes place through the catalytic effect of a radical such as OH (hydroxyl):

The odd hydrogen atoms and OH result from the dissociation of water vapour, molecular hydrogen and methane (CH4).

Stratospheric ozone is similarly destroyed in the presence of nitrogen oxides (NOx, i.e. NO2and NO) and chlorine radicals (Cl, ClO). The source gas of the NOx is nitrous oxide (N2O), which is produced by combustion and fertilizer use, while chlorofluorocarbons

(CFCs), manufactured for ‘freon’, give rise to the chlorines. These source gases are transported up to the stratosphere from the surface and are converted by oxidation into NOx, and by UV photodecomposition into chlorine radicals, respectively.

The chlorine chain involves:

and

Both reactions result in a conversion of O3to O2and the removal of all odd oxygens. Another cycle may involve an interaction of the oxides of chlorine and bromine (Br). It appears that the increases of Cl and Br species during the years 1970 to 1990 are sufficient to explain the observed decrease of stratospheric ozone over Antarctica (see Box 2.1). A mechanism that may enhance the catalytic process involves polar stratospheric clouds. These can form readily during the austral

Figure 2.2Schematic illustrations of (A) the Chapman cycle of ozone formation and (B) ozone destruction. X is any ozonedestroying species (e.g. H, OH, NO, CR, Br).

Source: After Hales (1996), from Bulletin of the American Meteorological Society, by permission of the American Meteorological Society.

spring (October), when temperatures decrease to 185 to 195 K, permitting the formation of particles of nitric acid (HNO3) ice and water ice. It is apparent, however, that anthropogenic sources of the trace gases are the primary factor in the ozone decline. Conditions in the Arctic are somewhat different as the stratosphere is warmer and there is more mixing of air from lower latitudes. Nevertheless, ozone decreases are now observed in the boreal spring in the Arctic stratosphere.

The constant metamorphosis of oxygen to ozone and from ozone back to oxygen involves a very complex set of photochemical processes, which tend to maintain an approximate equilibrium above about 40 km. However, the ozone mixing ratio is at its maximum at about 35 km, whereas maximum ozone concentration (see Note 1) occurs lower down, between 20 and 25 km in low latitudes and between 10 and 20 km in high latitudes. This is the result of a circulation mechanism transporting ozone downward to levels where its destruction is less likely, allowing an accumulation of the gas to occur. Despite the importance of the ozone layer, it is essential to realize that if the atmosphere were compressed to sealevel (at normal sea-level temperature and pressure) ozone would contribute only about 3 m to the total atmospheric thickness of 8 km (Figure 2.3).

6Variations with latitude and season

Variations of atmospheric composition with latitude and season are particularly important in the case of water vapour and stratospheric ozone.

Ozone content is low over the equator and high in subpolar latitudes in spring (see Figure 2.3). If the distribution were solely the result of photochemical

Ozone measurements were first made in the 1930s. Two properties are of interest: (i) the total ozone in an atmospheric column. This is measured with the Dobson spectrophotometer by comparing the solar radiation at a wavelength where ozone absorption occurs with that in another wavelength where such effects are absent; (i) the vertical distribution of ozone. This can be measured by chemical soundings of the stratosphere, or calculated at the surface using the Umkehrmethod; here the effect of solar elevation angle on the scattering of solar radiation is measured. Ozone measurements, begun in the Antarctic during the International Geophysical Year, 1957–58, showed a regular annual cycle with an austral spring (October–November) peak as ozone-rich air from mid-latitudes was transported poleward as the winter polar vortex in the stratosphere broke down. Values declined seasonally from around 450 Dobson units (DU) in spring to about 300 DU in summer and continued about this level through the autumn and winter. Scientists of the British Antarctic Survey noted a different pattern at Halley Base beginning in the 1970s. In spring, with the return of sunlight, values began to decrease steadily between about 12 and 20 km altitude. Also in the 1970s, satellite sounders began mapping the spatial distribution of ozone over the polar regions. These revealed that low values formed a central core and the term “Antarctic ozone hole” came into use. Since the mid-1970s, values start decreasing in late winter and reach minima of around 100 DU in the austral spring.

Using a boundary of 220 DU (corresponding to a thin, 2.2-m ozone layer, if all the gas were brought to sea level temperature and pressure), the extent of the Antarctic ozone hole at the end of September averaged 21 million km2, during 190–9. This expanded to cover 27 million km2by early September in 199 and 2000.

In the Arctic, temperatures in the stratosphere are not as low as over the Antarctic, but in recent years ozone depletion has been large when temperatures fall well below normal in the winter stratosphere. In February 1996, for example, column totals averaging 330 DU for the Arctic vortex were recorded compared with 360 DU, or higher, in other years. A series of mini-holes was observed over Greenland, the northern North Atlantic and northern Europe with an absolute low over Greenland below 180 DU. An extensive ozone hole is less likely to develop in the Arctic because the more dynamic stratospheric circulation, compared with the Antarctic, transports ozone poleward from mid-latitudes.

box 2.1 significant 20th-c. advance

processes, the maximum would occur in June near the equator, so the anomalous pattern must result from a poleward transport of ozone. Apparently, ozone moves from higher levels (30 to 40 km) in low latitudes towards lower levels (20 to 25 km) in high latitudes during the winter months. Here the ozone is stored during the polar night, giving rise to an ozone-rich layer in early spring under natural conditions. It is this feature that has been disrupted by the stratospheric ozone ‘hole’ that now forms each spring in the Antarctic and in some recent years in the Arctic also (see Box 2.1). The type of circulation responsible for this transfer is not yet known with certainty, although it does not seem to be a simple, direct one.

The water vapour content of the atmosphere is related closely to air temperature (see B.2, this chapter, and Chapter 4B and C) and is therefore greatest in summer and in low latitudes. There are, however, obvious exceptions to this generalization, such as the tropical desert areas of the world.

The carbon dioxide content of the air (currently averaging 372 parts per million (ppm)) has a large seasonal range in higher latitudes in the northern hemisphere associated with photosynthesis and decay in the biosphere. At 50°N, the concentration ranges from about 365 ppm in late summer to 378 ppm in spring. The low summer values are related to the assimilation of

CO2by the cold polar seas. Over the year, a small net transfer of CO2from low to high altitudes takes place to maintain an equilibrium content in the air.

The quantities of carbon dioxide, other greenhouse gases and particles in the atmosphere undergo long-term variations that may play an important role in the earth’s radiation budget. Measurements of atmospheric trace gases show increases in nearly all of them since the Industrial Revolution began (Table 2.3). The burning of fossil fuels is the primary source of these increasing trace gas concentrations. Heating, transportation and industrial activities generate almost 5 1020J/year of energy. Oil and natural gas consumption account for 60 per cent of global energy and coal about 25 per cent. Natural gas is almost 90 per cent methane

(CH4), whereas the burning of coal and oil releases not only CO2but also odd nitrogen (NOx), sulphur and carbon monoxide (CO). Other factors relating to agricultural practices (land clearance, farming, paddy cultivation and cattle raising) also contribute to modifying the atmospheric composition. The concentrations and sources of the most important greenhouse gases are considered in turn.

Carbon dioxide(CO2). The major reservoirs of carbon are in limestone sediments and fossil fuels. The atmosphere contains just over 775 1012kg of carbon

(C), corresponding to a CO2concentration of 370 ppm

(Figure 2.4). The major fluxes of CO2are a result of solution/dissolution in the ocean and photosynthesis/ respiration and decomposition by biota. The average

Figure 2.3Variation of total ozone with latitude and season in Dobson units (milliatmosphere centimeters) for two time intervals: (top) 1964–1980 and (bottom) 1984–1993. Values over 350 units are stippled.

Source: From Bojkov and Fioletov (1995). From Journal of Geophysical Research100 (D), Fig. 15, p. 16, 548. Courtesy of American Geophysical Union.

time for a CO2molecule to be dissolved in the ocean or taken up by plants is about four years. Photosynthetic activity leading to primary production on land involves 50 1012kg of carbon annually, representing 7 per cent of atmospheric carbon; this accounts for the annual oscillation in CO2observed in the northern hemisphere due to its extensive land biosphere.

The oceans play a key role in the global carbon cycle.

Photosynthesis by phytoplankton generates organic compounds of aqueous carbon dioxide. Eventually, some of the biogenic matter sinks into deeper water, where it undergoes decomposition and oxidation back into carbon dioxide. This process transfers carbon dioxide from the surface water and sequesters it in the ocean deep water. As a consequence, atmospheric concentrations of CO2can be maintained at a lower level than otherwise. This mechanism is known as a ‘biologic pump’; long-term changes in its operation may have caused the rise in atmospheric CO2at the end of the last glaciation. Ocean biomass productivity is limited by the availability of nutrients and by light. Hence, unlike the land biosphere, increasing CO2levels will not necessarily affect ocean productivity; inputs of ferti- lizers in river runoff may be a more significant factor. In the oceans, the carbon dioxide ultimately goes to produce carbonate of lime, partly in the form of shells and the skeletons of marine creatures. On land, the dead matter becomes humus, which may subsequently

Table 2.3Anthropogenically induced changes in concentration of atmospheric trace gases.

GasConcentrationAnnual Increase Sources (%) 1850* 2000 1990s

Carbon dioxide280 ppm370 ppm0.4Fossil fuels Methane800 ppbv1750ppbv0.3Rice paddies, cows, wetlands

Nitrous oxide280 ppbv316 ppbv0.25Microbiological activity, fertilizer, fossil fuel

CFC–1 0 0.27 ppbv ≈ 0 Freon† HCFC–20.1 ppbv5CFC substitute Ozone ? 10–50 ppbv ≈ 0 Photochemical reactions (troposphere)

Notes:*Pre-industrial levels are derived primarily from measurements in ice cores where air bubbles are trapped as snow accumulates on polar ice sheets. †Production began in the 1930s.

Source: Updated from Schimel et al. (1996), in Houghton et al. (1996).

Figure 2.4Global carbon reservoirs (gigatonnes of carbon (GtC): where 1 Gt = 109metric tons = 1012 kg) and gross annual fluxes (GtC yr–1). Numbers emboldened in the reservoirs suggest the net annual accumulation due to anthropogenic causes.

Source: Based on Sundquist, Trabalka, Bolin and Siegenthaler; after Houghton et al. (1990 and 2001).

form a fossil fuel. These transfers within the oceans and lithosphere involve very long timescales compared with exchanges involving the atmosphere.

As Figure 2.4 shows, the exchanges between the atmosphere and the other reservoirs are more or less balanced. Yet this balance is not an absolute one; between AD1750 and 2001 the concentration of atmos- pheric CO2is estimated to have increased by 32 per cent, from 280 to 370 ppm (Figure 2.5). Half of this increase has taken place since the mid-1960s; currently, atmos- pheric CO2levels are increasing by 1.5 ppmv per year. The primary net source is fossil fuel combustion, now accounting for 6.5 1012kg C/year. Tropical deforestation and fires may contribute a further 2 1012 kg C/year; the figure is still uncertain. Fires destroy only above-ground biomass, and a large fraction of the carbon is stored as charcoal in the soil. The consumption of fossil fuels should actually have produced an increase almost twice as great as is observed. Uptake and dissolution in the oceans and the terrestrial biosphere account primarily for the difference.

(Parte 6 de 9)

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