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 NO_x, 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 NO_x, 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, HO_x radicals (OH and HO₂) together with NO_x are key catalysts in the production of tropospheric O₃ (see Section 4.2.3.3). 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 NO_x 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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