Atmosfera, Tempo e Clima

Atmosfera, Tempo e Clima

(Parte 7 de 9)

Carbon dioxide has a significant impact on global temperature through its absorption and re-emission of radiation from the earth and atmosphere (see Chapter 3C). Calculations suggest that the increase from 320 ppm in the 1960s to 370 ppm (AD2001) raised the mean surface air temperature by 0.5°C (in the absence of other factors).

Research on deep ice cores taken from Antarctica has allowed changes in past atmospheric composition to be calculated by extracting air bubbles trapped in the old ice. This shows large natural variations in CO2 concentration over the ice age cycles (Figure 2.6). These variations of up to 100 ppm were contemporaneous with temperature changes that are estimated to be about 10°C. These long-term variations in carbon dioxide and climate are discussed further in Chapter 13.

Methane(CH4) concentration (1750 ppbv) is more than double the pre-industrial level (750 ppbv). It increased by about 4 to 5 ppbv annually in the 1990s but this dropped to zero in 1999 to 2000 (Figure 2.7). Methane has an atmospheric lifetime of about nine years and is responsible for some 18 per cent of the greenhouse effect. Cattle populations have increased by 5 per cent per year over thirty years and paddy rice area by 7 per cent per year, although it is uncertain whether these account quantitatively for the annual increase of 120 ppbv in methane over the past decade. Table 2.4, showing the mean annual release and consumption, indicates the uncertainties in our knowledge of its sources and sinks.

Nitrous oxide(N2O), which is relatively inert, orig-

Siple ice CoreEstimates of Callendar Machta

Mauna Loa Observations

Projected trends:


TION (ppmv)

Figure 2.5Estimated carbon dioxide concentration: since 1800 from air bubbles in an Antarctic ice core, early measurements from 1860 to 1960; observations at Mauna Loa, Hawaii, since 1957; and projected trends for this century.

Source: After Keeling, Callendar, Machta, Broecker and others. Note: (A) and (B) indicate different scenarios of global fossil fuel use (IPCC, 2001).

inates primarily from microbial activity (nitrification) in soils and in the oceans (4 to 8 109kg N/year), with about 1.0 109kg N/year from industrial processes. Other major anthropogenic sources are nitrogen fertil- izers and biomass burning. The concentration of N2O has increased from a pre-industrial level of about 285 ppbv to 316 ppbv (in clean air). Its increase began around 1940 and is now about 0.8 ppbv per year (Figure

2.8A). The major sink of N2O is in the stratosphere, where it is oxidized into NOx.

Chlorofluorocarbons (CF2Cl2 and CFCl3), better known as ‘freons’ CFC-1 and CFC-12, respectively, were first produced in the 1930s and now have a total atmospheric burden of 1010kg. They increased at 4 to 5 per cent per year up to 1990, but CFC-1 is declining slowly and CFC-12 is nearly static as a result of the

Figure 2.7Methane concentration (parts per million by volume) in air bubbles trapped in ice dating back to 1000 years BP obtained from ice cores in Greenland and Antarctica and the global average for AD 2000 (X).

Source: Data from Rasmussen and Khalil, Craig and Chou, and Robbins; adapted from Bolin et al. (eds) The Greenhouse Effect, Climatic Change, and Ecosystems(SCOPE 29). Copyright ©1986. Reprinted by permission of John Wiley & Sons, Inc.

Thousands of years (BP)

CO 2 (ppmv)


Figure 2.6Changes in atmospheric CO2(ppmv: parts per million by volume) and estimates of the resulting global temperature deviations from the present value obtained from air trapped in ice bubbles in cores at Vostok, Antarctica.

Source: Our Future World, Natural Environment Research Council (NERC) (1989).

Table 2.4Mean annual release and consumption of CH4 (Tg).

Mean Range

A Release

Natural wetlands 115 100–200 Rice paddies 110 25–170 Enteric fermentation 8065–110 (mammals)

B Consumption

Soils 30 15–30 Reaction with OH500400–600 Totalc. 530

Source: Tetlow-Smith 1995.

Montreal Protocol agreements to curtail production and use substitutes (see Figure 2.8B). Although their concentration is <1 ppbv, CFCs account for nearly 10 per cent of the greenhouse effect. They have a residence time of 5 to 130 years in the atmosphere. However, while the replacement of CFCs by hydrohalocarbons (HCFCs) can reduce significantly the depletion of stratospheric ozone, HCFCs still have a large greenhouse potential.

Ozone(O3) is distributed very unevenly with height and latitude (see Figure 2.3) as a result of the complex photochemistry involved in its production (A.2, this chapter). Since the late 1970s, dramatic declines in springtime total ozone have been detected over high southern latitudes. The normal increase in stratospheric ozone associated with increasing solar radiation in spring apparently failed to develop. Observations in Antarctica show a decrease in total ozone in September to October from 320 Dobson units (DU) (10–3cm at standard atmospheric temperature and pressure) in the 1960s to around 100 in the 1990s. Satellite measurements of stratospheric ozone (Figure 2.9) illustrate the presence of an ‘ozone hole’ over the south polar region (see Box 2.2). Similar reductions are also evident in the Arctic and at lower latitudes. Between 1979 and 1986, there was a 30 per cent decrease in ozone at 30 to 40-km altitude between latitudes 20 and 50°N and S (Figure 2.10); along with this there has been an increase in ozone in the lowest 10 km as a result of anthropogenic activities. Tropospheric ozone represents about 34 DU compared with 25 pre-industrially. These changes in the vertical distribution of ozone concentration are likely to lead to changes in atmospheric heating (Chapter 2C), with implications for future climate trends (see Chapter 13). The global mean column total decreased from 306 DU for 1964 to 1980 to 297 for 1984 to 1993 (see Figure 2.3). The decline over the past twenty-five years has exceeded 7 per cent in middle and high latitudes.

The effects of reduced stratospheric ozone are particularly important for their potential biological damage to living cells and human skin. It is estimated that a 1 per cent reduction in total ozone will increase ultraviolet-B radiation by 2 per cent, for example, and ultraviolet radiation at 0.30 µm is a thousand times more damaging to the skin than at 0.3 µm (see Chapter 3A). The ozone decrease would also be greater in higher latitudes. However, the mean latitudinal and altitudinal gradients of radiation imply that the effects of a 2 per cent UV-B increase in mid-latitudes could be offset by moving poleward 60 km or 100 m lower in altitude! Recent polar observations suggest dramatic changes. Stratospheric ozone totals in the 1990s over Palmer Station, Antarctica (65°S), now maintain low levels from September until early December, instead of recovering in November. Hence, the altitude of the sun has been higher and the incoming radiation much greater than in previous years, especially at wavelengths ≤0.30 µm. However, the possible effects of increased UV radiation on biota remain to be determined.

Aerosolloading may change due to natural and human-induced processes. Atmospheric particle con-

Figure 2.8Concentration of: (A) nitrous oxide, N2O (left scale), which has increased since the mid-eighteenth century and especially since 1950; and of (B) CFC-1 since 1950 (right scale). Both in parts per billion by volume (ppbv).

Source: After Houghton et al. (1990 and 2001).

Jul Aug Sep Oct Nov Dec

Column Ozone (DU)

Figure 2.9Total ozone measurements from ozonesondes over South Pole for 1967 to 1971, 1989, and 2001, showing deepening of the Antarctic ozone hole.

Source: Based on Climate Monitoring and Diagnostics Laboratory, NOAA.

centration derived from volcanic dust is extremely irregular (see Figure 2.1), but individual volcanic emissions are rapidly diffused geographically. As shown in Figure 2.12, a strong westerly wind circulation carried the El Chichón dust cloud at an average velocity of

20 m s–1so that it encircled the globe in less than three weeks. The spread of the Krakatoa dust in 1883 was more rapid and extensive due to the greater amount of fine dust that was blasted into the stratosphere. In June 1991, the eruption of Mount Pinatubo in the

% change per decade km MAM

Figure 2.10Changes in stratospheric ozone content (per cent per decade) during March to May and September to November 1978 to 1997 over Europe (composite of Belsk, Poland, Arosa, Switzerland and Observatoire de Haute Provence, France) based on umkehr measurements.

Source: Adapted from Bojkov et al. (2002), Meteorology and Atmospheric Physics, 79, p. 148, Fig. 14a.

K uwae Billy Mitchell


Aw u

Long Island Laki

Unknown T ambora


Krakatau 1902 eventsKatmai


Chichon Calibrated Optical Depth


Figure 2.11Record of volcanic eruptions in the GISP 2 ice core and calibrated visible optical depth for AD1300 to 2000, together with the names of major volcanic eruptions. Note that the record reflects eruptions in the northern hemisphere and equatorial region only; optical depth estimates depend on the latitude and the technique used for calibration.

Source: Updated after Zielinski et al. (1995), Journal of Geophysical Research100 (D), courtesy of the American Geophysical Union, p. 20, 950, Fig. 6.

Philippines injected twenty megatons of SO2into the stratosphere. However, only about twelve eruptions have produced measurable dust veils in the past 120 years. They occurred mainly between 1883 and 1912, and 1982 and 1992. In contrast, the contribution of manmade particles (particularly sulphates and soil) has been progressively increasing, and now accounts for about 30 per cent of the total.

The overall effect of aerosols on the lower atmosphere is uncertain; urban pollutants generally warm the atmosphere through absorption and reduce solar radiation reaching the surface (see Chapter 3C). Aerosols may lower the planetary albedo above a high-albedo desert or snow surface but increase it over an ocean surface. Thus the global role of tropospheric aerosols is difficult to evaluate, although many authorities now consider it to be one of cooling. Volcanic eruptions, which inject dust and sulphur dioxide high into the stratosphere, are known to cause a small deficit in surface heating with a global effect of –0.1°to –0.2°C, but the effect is short-lived, lasting only a year or so after the event (see Box 13.3). In addition, unless the eruption is in low latitudes, the dust and sulphate aerosols remain in one hemisphere and do not cross the equator.

Atmospheric gases obey a few simple laws in response to changes in pressure and temperature. The first, Boyle’s Law, states that, at a constant temperature, the volume (V) of a mass of gas varies inversely as its pressure (P), i.e.

k1 P= ––

(k1is a constant). The second, Charles’s Law, states that, at a constant pressure, volume varies directly with absolute temperature (T) measured in degrees Kelvin (see Note 2):

V= k2T

These laws imply that the three qualities of pressure, temperature and volume are completely interdependent, such that any change in one of them will cause a compensating change to occur in one, or both, of the remainder. The gas laws may be combined to give the following relationship:

PV= RmT where m= mass of air, and R= a gas constant for dry air (287 J kg–1K–1) (see Note 3). If mand Tare held fixed, we obtain Boyle’s Law; if mand Pare held fixed, we obtain Charles’s Law. Since it is convenient to use density, ρ(= mass/volume), rather than volume when studying the atmosphere, we can rewrite the equation in the form known as the equation of state:

Thus, at any given pressure, an increase in temperature causes a decrease in density, and vice versa.

1Total pressure

Air is highly compressible, such that its lower layers are much more dense than those above. Fifty per cent of the total mass of air is found below 5 km (see Figure 2.13), and the average density decreases from about 1.2 kg m–3at the surface to 0.7 kg m–3at 5000 m (approximately 16,0 ft), close to the extreme limit of human habitation.

Pressure is measured as a force per unit area. A force of 105newtons acting on 1 m2corresponds to the Pascal (Pa) which is the Système International (SI) unit of pressure. Meteorologists still commonly use the millibar (mb) unit; 1 millibar = 102Pa (or 1 hPa; h = hecto) (see Appendix 2). Pressure readings are made with a mercury barometer, which in effect measures the height of the column of mercury that the atmosphere is able to support in a vertical glass tube. The closed upper end of the tube has a vacuum space and its open lower end is immersed in a cistern of mercury. By exerting pressure downward on the surface of mercury in the cistern, the atmosphere is able to support a mercury column in the tube of about 760 m (29.9 in or approximately 1013 mb). The weight of air on a surface at sea-level is about 10,0 kg per square metre.

Pressures are standardized in three ways. The readings from a mercury barometer are adjusted to correspond to those for a standard temperature of 0°C (to allow for the thermal expansion of mercury); they are referred to a standard gravity value of 9.81 ms–2at 45° latitude (to allow for the slight latitudinal variation in g from 9.78 ms–2at the equator to 9.83 ms–2at the poles);

and they are calculated for mean sea-level to eliminate the effect of station elevation. This third correction is the most significant, because near sea-level pressure decreases with height by about 1 mb per 8 m. A fictitious temperature between the station and sea-level has to be assumed and in mountain areas this commonly causes bias in the calculated mean sea-level pressure (see Note 4).

The mean sea-level pressure (p0) can be estimated from the total mass of the atmosphere (M, the mean acceleration due to gravity (g0) and the mean earth radius (R):

where the denominator is the surface area of a spherical earth. Substituting appropriate values into this

106m), we find p0= 105kg ms–2= 105Nm–2, or 105 Pa. Hence the mean sea-level pressure is approxi- mately 105Pa or 1000 mb. The global mean value is 1013.25 mb. On average, nitrogen contributes about 760 mb, oxygen 240 mb and water vapour 10 mb. In other words, each gas exerts a partial pressure independent of the others.

Atmospheric pressure, depending as it does on the weight of the overlying atmosphere, decreases logarithmically with height. This relationship is expressed by the hydrostatic equation:

∂p –– = –gρ ∂z

Figure 2.12The spread of volcanic material in the atmosphere following major eruptions. (A) Approximate distributions of observed optical sky phenomena associated with the spread of Krakatoa volcanic dust between the eruption of 26 August and 30 November 1883. (B) The spread of the volcanic dust cloud following the main eruption of the El Chichón volcano in Mexico on 3 April 1982. Distributions on 5, 15 and 25 April are shown.

Sources: Russell and Archibald (1888), Simkin and Fiske (1983), Rampino and Self (1984), Robock and Matson (1983). (A) by permission of the Smithsonian Institution; (B) by permission of Scientific American Inc.

(Parte 7 de 9)