A ko T [M

and f is a constant that for atmospheric conditions a value of about 0.6 adequately fits the reaction data. There are simpler limiting forms for kf depending whether or not we have low or high pressure. At low pressures kf — ko(T) = ko(300)

If the number density [M] of the catalytic third body is in units of molecules cm~3 then the units of ko are cm6

molecules 2 s 1

cm3 molecules 1 s 1. Usually, ko(300), ka 2006, IUPAC 2006, NIST 2006).

while the units of kare (300), n and m are tabulated (JPL

7.8 Ozone photochemistry

7.8.1 The Chapman mechanism

In 1930 Chapman proposed a mechanism for the production of stratospheric ozone. The construction of the photoelectric spectrophotometer for measuring the amount (column density) of atmospheric ozone by Dobson (1931) and Chapman's mechanism played a key role in initiating research on stratospheric ozone. Chapman proposed that an ozone layer should form in the stratosphere through the photolysis of O2 resulting in the production of atomic oxygen both in the ground state O(3P) and in the excited state O(1D), eqn (7.53) and eqn (7.54), which could recombine with molecular oxygen to produce ozone. As we have seen in §7.6.3, O(1D) is formed above 60 km altitude while O forms also below 60 km with diffusion transporting these atoms to lower altitudes, as can be seen in Fig. 7.10. Above about 110 km atomic oxygen becomes the second most abundant species after molecular nitrogen. The reaction scheme

03 + hv O + O2, (7.93) O3 + hv — O(1D) + O2, (7.94)

10 " 10 " 10 " 10 " 10"10 10^ 10-8 10 7 10-6 10-5 10"* 10 3 10 2 101 10e

Volume mixing ratio

10 " 10 " 10 " 10 " 10"10 10^ 10-8 10 7 10-6 10-5 10"* 10 3 10 2 101 10e

Volume mixing ratio

FlG. 7.10. Global mean vertical profiles of oxygen and nitrogen species involved in O3 photochemistry.

involves only oxygen species with the deactivation of O(1D) primarily by molecular nitrogen and oxygen (Appendix B, Table B.4). These reactions produce an ozone layer between 20 and 40 km with a peak concentration at 30 km of about 1.5x1013 molecules cm~3 that is a factor 4 greater then the global mean observed stratospheric peak value of about 3.5x1012 at 25 km. Note that the peak in ozone mixing ratio occurs at about 35 km with a value close to 10 ppmv, as shown in Fig. 7.10. The discrepancy arises from the absence of important cat-alytical destruction cycles involving NO and OH, and related species NO2 and HO2.

7.8.2 N2O and NOx photochemistry

The oxides of nitrogen, especially NO and NO2 referred to as NOx, play an important role in reducing stratospheric ozone levels and explain the difference between the ozone levels predicted by the Chapman mechanism and observations

O3 + O ^ 2O2, O3 + O(1D) ^ 2O2, O3 + O(1D) ^ O2 + 2O,

(Crutzen 1970, 1971, McElroy and McConnell 1971). The destruction mechanism is the following catalytic cycle

The source of NO is primarily nitrous oxide N2O from biological processes in soil and in the oceans with an annual emission rate of about 25 Tg (Table 7.4). Nitrous oxide is a very stable molecule within the troposphere with a residence time of 10-20 years but is oxidized by O(XD) to NO in the stratosphere, and pho-tolyzed to N2. NO is also produced by oxidation of N2 by the high temperatures of jet engines.

In Fig. 7.10 is shown the vertical distribution of N2O where it can be seen that its mixing ratio is constant at 314 ppbv until about 20 km altitude. Other important reactions involving nitrogen-containing species are given in Appendix B, Tables B.4 and B.6, amongst which are NO3, HNO3, HNO4, N2O5, referred to as nitrogen reservoir species as they do not directly destroy ozone. The NOK species plus the reservoir species together are referred to as the NOy chemical family.

Daytime-nighttime effects become important for the NOy family as NO3 and N2O5 are formed at night since NO3 is rapidly destroyed by sunlight. Also, HNO3 is formed by reaction between OH and NO2 during nighttime as OH is formed in the night by the reaction of O(XD) with H2O up to an altitude of 60 km. We note that HNO3 (nitric acid) is highly soluble and is removed from the atmosphere by rain.

Another significant source of atmospheric NO arises from the reaction of excited nitrogen atoms, N(2D) with molecular oxygen above 100 km altitude. Nitric oxide is produced primarily via the following chain of reactions (Appendix B, Table B.5)

Thus, the above chain of reactions is initiated by the photolysis of molecular oxygen and the dissociative ionization of molecular nitrogen below 37.5 nm. Another

Ionosphere Composition
FIg. 7.11. Global mean vertical profiles of ionic species that control mesospheric NO+, together with N(2D).

important production mechanism for N(2 D) is the dissociation of molecular nitrogen by energetic electrons (about 10 eV) that are produced by solar soft X-rays (2-7 nm) and extreme ultraviolet radiation (EUV) in the spectral region 7-20 nm. We note that the solar flux in these spectral regions varies significantly over the solar cycle (see §5.2.3) and so does thermospheric NO. Electron precipitation in the auroral polar region is an additional source of energetic electrons. The NO plays an important role in the energy balance of the thermosphere via cooling to space and if transported to the mesosphere it can contribute to the catalytic destruction of ozone. Global mean vertical profiles of ionic species, which control mesospheric NO+ together with N(2D), are shown in Fig. 7.11. The dissociation of molecular nitrogen, oxygen, and the hydrogen flux control the variation of the atmospheric molecular weight with altitude in the thermosphere and lead to the formation of the heterosphere, where the atmospheric composition is no longer uniform, as shown by the rapid decrease in the atmospheric molecular weight above about 100 km in Fig. 7.12.

7.8.3 Water vapour and HOx photochemistry

The water-vapour mixing ratio is determined by ocean evaporation (see §8.9), the tropospheric air temperature and by diffusion and transport to the stratosphere.

fig. 7.12. Model vertical profiles of number density, pressure, temperature and mean molecular weight.

It is also produced by methane oxidation in the stratosphere and its mixing ratio there is fairly constant, ranging between 3 and 6 ppmv (see Fig. 7.13). In the stratosphere it is destroyed by O(1D) during nighttime via the reaction

while above 70 km H2O is readily destroyed by solar Lyman-a radiation (see Fig. 7.9 and Lewis et al. 1983). Photolysis of water vapour by Lyman-a can result in the production of active and inactive hydrogen species, with respect to ozone destruction. As we saw earlier, the third channel is the dominant with branching ratio a = 0.78 (§7.6.5) so the direct production of H2 from the photodissociation of water vapour is a small fraction of the process of Lyman-a photodissociation. The role of the HOœ species (H, OH and HO2) in destroying ozone was investigated in detail by Allen et al. (1984) who stressed the key role played by reactions that convert the active-hydrogen HOœ species, to the inactive-hydrogen species H2 and H2 O.

The hydroxyl radical can efficiently destroy ozone in the stratosphere and mesosphere via the catalytic cycle

Volume Mixing Ratio

flg. 7.13. Global mean vertical profiles of key species involved in water vapour and CO2-initiated photochemistry.

In the mesosphere, atomic hydrogen plays an important role in the destruction of ozone via the reaction

resulting in a deep minimum in mesopause ozone. The depth of this minimum can range from a mixing ratio of 0.1 ppmv to 0.01 ppmv. This depends on the reaction of H with HO2 and on the branching ratio a for H2O photolysis. The reaction of H with HO2 has three reaction channels (Appendix B, Table B.6)

where channel a that leads to the production of the active hydrogen species OH is dominant (branching ratio 0.9). It is the strength of channel c that limits the role of water vapour photolysis in ozone destruction through the conversion of

flG. 7.14. Model global mean diurnally averaged ozone profiles with and without channels a, b, and c of the reaction H + HO2, with a = 1. (Vardavas et al. 1998)

atomic hydrogen to inactive molecular hydrogen. In Fig. 7.14 is shown the effect of including sequentially channels a, b and c, for a photolysis branching ratio a =1.

7.8.4 Chlorine and ClOx photochemistry

Chlorine is naturally emitted into the atmosphere from the oceans and biomass burning in the form of methyl chloride, CH3Cl. Significant anthropogenic sources are carbon tetrachloride, CCl4, methylchloroform, CH3CCl3 and chlorofluoro-carbons such as CFCl3 (CFC-11) and CF2Cl2 (CFC-12), amongst others. CFCs were introduced into the atmosphere after 1960 through their industrial use, refrigeration and in sprays.

In 1974 Molina and Rowland warned of the potential role of CFCs in the destruction of stratospheric ozone associated with their increasing atmospheric concentrations. As can be seen in Table 7.5, the total Cl mixing ratio trebled in about 40 years from about 1 ppbv pre-1960. Molina and Rowland showed that the catalytic destruction of ozone by ClOx (Cl and ClO) would have a major effect on stratospheric ozone. This and the discovery of a substantial reduction

Table 7.5 Approximate surface chlorine species mixing ratios (ppbv) and total Cl from pre-1960 to 1998. (WMO 1999, IPCC 2001)

Species

pre-1960

1980

1990

1998

CFC13(CFC-11)

0.00

0.17

0.24

0.27

CF2 Cl2(CFC-12)

0.00

0.29

0.50

0.53

CH3 Cl

0.62

0.62

0.62

0.62

CH3 CC13

0.00

0.10

0.10

0.07

CC14

0.10

0.10

0.10

0.10

Total Cl

1.02

2.41

3.04

3.10

in the ozone column density in Antarctica (ozone hole) led to the Montreal Protocol (1987) that has resulted in the control and reduction in CFC emissions (IPCC 2001). Their suggested replacements by the ozone-inert hydrochlorofluo-rocarbons (HCFCs) and hydrofluorocarbons (HFCs), however, have introduced a new potential source of greenhouse warming.

CFC molecules are stable in the troposphere but undergo photolysis in the stratosphere by solar ultraviolet radiation (see Appendix B) releasing chlorine atoms

The chlorine source species diffuse slowly into the stratosphere (Fig. 7.15) where they are destroyed and have tropospheric lifetimes ranging between 2 and 100 years. The ClOK family can destroy ozone via the catalytic cycle

The catalytic cycle is limited by the conversion of ClOK to chlorine reservoir species such as HCl, HNO3, ClONO2 and HOCl via the reactions (Appendix B, Table B.7)

so there is coupling between chlorine and methane chemistry, and chlorine and NOk chemistry. The reservoir species HCl and HNO3 are soluble in water and so can be removed from the troposphere by rain, while ClONO2 can be converted to Cl through heterogeneous chemistry (gas and solid/liquid phase reactions) taking place on polar stratospheric clouds (PSCs).

flg. 7.15. Global mean vertical profiles of chlorine species based on pre-1960 sources of chlorine.

7.8.5 Polar stratospheric clouds

The lifetime of HCl is weeks and that of ClONO2 a day. HCl is destroyed by OH, which we saw is a nighttime species in the stratosphere, while ClONO2 is destroyed by photolysis. Hence, during nighttime conditions these reservoir species are protected from decomposition back to ozone-active components. Thus, within the polar winter there should be a buildup of these reservoir species. However, the very low temperatures (typically 195 K) of the lower stratosphere during the Antarctic winter allows stratospheric clouds made of ice crystals to form at such low pressures. On ice crystals (containing water and HCl), the following heterogeneous reactions take place that decompose the chlorine reservoir species

ClONO2(gas) + HCl(solid) ^ Cl2(gas) + HNO3(solid), (7.129) ClONO2(gas) + H2O(solid) ^ HOCl(gas) + HNO3(solid). (7.130)

Molecular chlorine and HOCl are easily photolysed to produce Cl during the polar spring (daytime) and hence significant ozone depletion. The HNO3 can be removed from the atmosphere through coagulation and particle sedimentation with the result that NO2 decreases and formation of the reservoir species ClONO2 is limited.

The discovery of the decrease in the total ozone column in Antarctica by the British Antarctic Survey (WMO 1995), by about a factor of 2 between 1955 and 1985, led eventually to the correlation of the presence of high ClO concentrations with Antarctic ozone depletion. This led to the understanding of the role of the ClO dimer (ClOOCl) and BrO in ozone depletion when ClO concentrations are maintained high by the Antartic vortex (wind circulation about the pole) that limits flow into the region during winter.

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