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oh+ho2-h2o+o2

Saw These reactions are also of interest for tropo-sphertc chemistry (see other parts of this book J.

Saw These reactions are also of interest for tropo-sphertc chemistry (see other parts of this book J.

What actually happens is that NO, destroys ozone at high altitude but forms it at low altitude. As a result, whether there is a net increase or decrease in total ozone depends on the altitude at which NO, is injected. The transition from ozone formation to destruction appears to occur somewhat below 20 km, the precise level depending on details of the mathematical model (Johnston and Podolske, 1978). Thus the flight altitude of a supersonic transport is critical ; NO, from sources at the Earth's surface always seems to increase ozone (Turco et ai, 1978).

Some years ago Molina and Rowland (1974) demonstrated that chlorofluoro-methanes (see Section 3.3) are photochemically destroyed in the stratosphere by solar radiation with wavelength shorter than 0.23 fim to form chlorine atoms:

Thus, these substances, non-reactive under tropospheric conditions, give a species which reacts rather rapidly with 03 and the resultant CIO reacts with atomic oxygen to regenerate CI (Rowland, 1976)13:

The halogenated hydrocarbons reaching the stratosphere through the tropopause by slow diffusion or through the tropopause gaps are mainly CH3C1 (methyl chloride), CC14 (carbon tetrachloride), CFC13 (fluorotrichloromethane) and CF2C12 (dichlorodifluoromethane). While CH3C1 is predominantly of natural origin, the other compounds mentioned are man-made. Thus, we cannot exclude the possibility that man can modify the stratospheric photochemistry by the release of organochlorine species (see Chapter 6).

It is to be noted that the above concept was modified several times after the publication of the original paper of Molina and Rowland (1974). Thus it has been shown by Rowland, and his co-workers (see, e.g. Vupputuri, 1977) that the following reactions stabilize the ozone destruction by the CI-CIO cycle in the mid-stratosphere:

Obviously the reactions [3.29a] and [3.29b] reduce substantially the ozone destruction due to anthropogenic F-l 1 and F-12 emissions. However, Howard and Evenson (1977) found that the rate of the chemical reaction

was much faster than previous indirect measurements had indicated. This finding again increased the value of the ozone reduction obtained by different model calculations14 (see Chapter 6).

13 For further details see e.g.: Rundel and Stolarski (1976).

'■•This is due to the fact that the reaction [3.30] removes H02 which can react with CI: CI + H02-»HC1 + 02. Furthermore by [3.30] OH radicals are formed. Thus the reaction HC1 + 0H-»H20 + C1 is accelerated.

It can be seen from reactions [3.27] and [3.28] that CI and CIO are reformed by these processes. A temporary chemical sink for chlorine species is provided in the stratosphere by some reaction steps leading to the formation of hydrogen chloride. HC1 can be removed from stratospheric air, like nitric acid, by slow downward mixing.

As in the case of nitrogen species, the value of theoretical predictions can be evaluated by comparing the measured stratospheric concentrations of halocarbons with calculated profiles. The results of stratospheric halocarbon analyses have been recently reviewed by Volz et al. (1978). Their data show that halocarbon levels decrease rather rapidly in the stratosphere with increasing altitude. The rate of this decrease is comparable to the theoretical value.

However, the experience of the change in direction of the effect of NO v must make us cautious in concluding that we understand the chemistry of the stratosphere. The recent past has included the discovery that one key reaction rate, supposedly well established by previous research, was in error by several orders of magnitude. Findings that do not agree with the models are coming to light. The field of stratospheric research, especially with regard to possible anthropogenic effects on essential trace gases, is certain to challenge researchers for a number of years to come.

3.44 The distribution of total otoiw.

Vertical profile of the o/one concentration

Generally, spectrophotometric methods are used to measure the atmospheric ozone concentration. By means of suitable optical devices15 one detects the intensity of solar radiation in two narrow wavelength bands. In one of these bands ozone strongly absorbs the radiation while, in the other, 03 has little absorption. The ozone content of the air can be calculated by comparing the two values measured. In this way the total ozone quantity in an air column is determined. This parameter expressed in cm is termed the total ozone. It represents the thickness of the layer which the same amount of ozone would form if it were separated from the air and held at normal temperature and pressure.

Figure 10 represents the average distribution of the total ozone (labelled x) as a function of geographical latitude and month of the year. Data refer to the Northern Hemisphere (Khrguian, 1969 and 1973). The curves in Fig. 10 were constructed on the basis of measurements carried out between 1957 and 1964. It can be seen that at tropical latitudes where changes in solar radiation are insignificant the variation of total ozone can be neglected. With increasing latitude the amplitude of variation increases and the maximum in late winter or early spring becomes more and more pronounced. The minimum can be found at mid-latitudes during autumn. It follows

" E.g. Dobson spectrophotometers.

from this distribution picture that the latitudinal gradient is greater in the early spring than during other seasons.

On the basis of photochemical considerations discussed in the last section we would expect a very different ozone distribution. For this reason it can be concluded that the pattern given by Fig. 10 can be explained only by the effect of atmospheric motions. This explanation is supported, together with other evidence by the fact that the results of individual measurements are also very much influenced by the atmospheric circulation.

Annual changes in the amount of total ozone (*) in the Northern Hemisphere as a function of geographical latitude (Khrguian, 1973). (By courtesy of Gidrometizdat)

Annual changes in the amount of total ozone (*) in the Northern Hemisphere as a function of geographical latitude (Khrguian, 1973). (By courtesy of Gidrometizdat)

Khrguian's (1969) calculations based on data obtained in 1958 and 1959 show that on average x=0.298 cm in the Northern Hemisphere. The corresponding value for the Southern Hemisphere is equal to 0.307 cm. Based on these figures he calculated that the total 03 mass in the atmosphere is 3360 x 106 t. Junge (1962) calculated the total ozone quantity in the atmospheric reservoir over the Northern Hemisphere on the basis of another data set. He found a value of 1750 x 106 t. The two estimates are in good agreement.

The vertical profile of the ozone concentration can also be determined by spectrophotometric observations16. However, more accurate distributions are measured by ozone sondes lifted by balloons into the upper layers of the

16 This procedure is called the "Umkehr" method (Götz, 1951).

atmosphere. These ozone sondes generally work on optical or chemical principles. The results obtained by chemical sondes (identifying ozone by iodometric methods) over the Northern Hemisphere have recently been compiled by Dutsch (1978). For levels above 10 mb and for the Southern Hemisphere he also used data gained by the Umkehr procedure. His average profiles for the Northern Hemisphere are reproduced in Fig. 11 as a function of latitude and season. One can see that, especially in winter and spring, there is a strong poleward gradient in the lower stratosphere. Above about 30 mb the poleward gradient becomes negative. Comparing Figs. 10 and 11 we can say that the late winter or early spring maximum in the value of total ozone observed over higher latitudes is caused by the huge quantity of ozone in the lower stratosphere at this time.

200nb mb

200nb

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