Water vapour in the lower stratosphere is a very effective greenhouse gas. Baseline levels of stratospheric H2O are controlled by the temperature of the tropical tropopause, a parameter that changes with climate (Moyer et al., 1996; Rosenlof et al., 1997; Dessler, 1998; Mote et al., 1998). The oxidation of CH4 is a source of mid-stratospheric H2O and currently causes its abundance to increase from about 3 ppm at the tropopause to about 6 ppm in the upper stratosphere. In addition, future direct injections of H2O from high-flying aircraft may add H2O to the lower stratosphere (Penner et al., 1999). Oltmans and Hofmann (1995) report statistically significant increases in lower strato-spheric H2O above Boulder, Colorado between 1981 and 1994. The vertical profile and amplitude of these changes do not correspond quantitatively with that expected from the recognised anthropogenic sources (CH4 oxidation). Analyses of satellite and ground-based measurements (Nedoluha et al., 1998; Michelsen et al., 2000) find increases in upper stratospheric H2O from 1985 to 1997, but at rates (>1%/yr) that exceed those from identified anthropogenic sources (i.e., aviation and methane increases) and that obviously could not have been maintained over many decades. In principle such a temporary trend could be caused by a warming tropopause, but a recent analysis indicates instead a cooling tropopause (Simmons et al., 1999). It is important to resolve these apparent discrepancies; since, without a physical basis for this recent trend, no recommendation can be made here for projecting changes in lower stratospheric H2O over the 21st century.
The hydroxyl radical (OH) is the primary cleansing agent of the lower atmosphere, in particular, it provides the dominant sink for CH4 and HFCs as well as the pollutants NOx, CO and VOC. Once formed, tropospheric OH reacts with CH4 or CO within a second. The local abundance of OH is controlled by the local abundances of NOx, CO, VOC, CH4, O3, and H2O as well as the intensity of solar UV; and thus it varies greatly with time of day, season, and geographic location.
The primary source of tropospheric OH is a pair of reactions that start with the photodissociation of O3 by solar UV.
O3 + h
O(1D) + O2
O(1D) + H2O OH + OH
Although in polluted regions and in the upper troposphere, photodissociation of other trace gases such as peroxides, acetone and formaldehyde (Singh et al., 1995; Arnold et al., 1997) may provide the dominant source (e.g., Folkins et al., 1997; Prather and Jacob, 1997; Wennberg et al., 1998; Müller and Brasseur, 1999). OH reacts with many atmospheric trace gases, in most cases as the first and rate-determining step of a reaction chain that leads to more or less complete oxidation of the compound. These chains often lead to formation of an HO2 radical, which then reacts with O3 or NO to recycle back to OH. Tropospheric OH is lost through reactions with other radicals, e.g., the reaction with HO2 to form H2O or with NO2 to form HNO3. In addition to providing the primary loss for CH4 and other pollutants, HOx radicals (OH and HO2) together with NOx are key catalysts in the production of tropospheric O3 (see Section 18.104.22.168). The sources and sinks of OH involve most of the fast photochemistry of the troposphere.
Pre-industrial OH is likely to have been different than today, but because of the counteracting effects of lower CO and CH4 (increasing OH) and reduced NOx and O3 (decreasing OH), there is no consensus on the magnitude of this change (e.g., Wang and Jacob, 1998). Trends in the current OH burden appear to be <1%/yr. Separate analyses of the CH3CCl3 observations for the period 1978 to 1994 report two different but overlapping trends in global OH: no trend within the uncertainty range (Prinn et al., 1995), and 0.5 ± 0.6%/yr (Krol et al., 1998). Based on the OxComp workshop, the SRES projected emissions would lead to future changes in tropospheric OH that ranging from +5% to -20% (see Section 4.4).
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