7.4.6 Stratospheric Ozone and Climate
From about 1980 to the mid-1990s a negative trend in globally averaged total ozone occurred, due primarily to an increase in Cl and bromine loading (Montzka et al., 1999). A reduction in halogen loading appears to have occurred recently (Montzka et al., 2003) as well as the beginning of ozone recovery (e.g., Newchurch et al., 2003; Huck et al., 2005; Reinsel et al., 2005; Yang et al., 2005). Evidence suggests that a sustainable recovery of ozone is not expected before the end of the current decade (e.g., Steinbrecht et al., 2004; Dameris et al., 2006). Atmospheric concentrations of LLGHGs have increased (see Chapter 2) and are expected to continue to increase, with consequences for the ozone layer. This section assesses current understanding of interactions and feedbacks between stratospheric ozone and climate. More detailed discussions can be found in recent reports (European Commission, 2003; IPCC/TEAP, 2005).
Stratospheric ozone is affected by climate change through changes in dynamics and in the chemical composition of the troposphere and stratosphere. An increase in the concentrations of LLGHGs, especially CO2, cools the stratosphere, allowing the possibility of more PSCs, and alters the ozone distribution (Rosenlof et al., 2001; Rosenfield et al., 2002; Randel et al., 2004, 2006; Fueglistaler and Haynes, 2005). With the possible exception of the polar lower stratosphere, a decrease in temperature reduces ozone depletion leading to higher ozone column amounts and a positive correction to the LLGHG-induced radiative cooling of the stratosphere. Moreover, ozone itself is a greenhouse gas and absorbs UV radiation in the stratosphere. Absorption of UV radiation provides the heating responsible for the observed temperature increase with height above the tropopause. Changes in stratospheric temperatures, induced by changes in ozone or LLGHG concentration, alter the Brewer-Dobson circulation (Butchart and Scaife, 2001; Butchart et al., 2006), controlling the rate at which long-lived molecules, such as LLGHGs, CFCs, HCFCs and halogens are transported from the troposphere to various levels in the stratosphere. Furthermore, increases in the Brewer-Dobson circulation increase temperatures adiabatically in the polar regions and decrease temperatures adiabatically in the tropics.
Climate is affected by changes in stratospheric ozone, which radiates infrared radiation down to the troposphere. For a given percentage change in the vertical structure of ozone, the largest dependence of the radiative forcing is in the upper troposphere and ozone layer regions (e.g., TAR, Figure 6.1). Past ozone depletion has induced surface cooling (Chapter 2). The observed decrease in stratospheric ozone and the resultant increase in UV irradiance (e.g., Zerefos et al., 1998; McKenzie et al., 1999) have affected the biosphere and biogenic emissions (Larsen, 2005). Such UV radiation increases lead to an enhanced OH production, reducing the lifetime of CH4 and influencing tropospheric ozone, both important greenhouse gases (European Commission, 2003). In addition to global mean equilibrium surface temperature changes, local surface temperature changes have been identified by Gillett and Thompson (2003) as a result of ozone loss from the lower stratosphere. Observational (e.g., Baldwin and Dunkerton, 1999, 2001; Thompson et al., 2005) and modelling (Polvani and Kushner, 2002; Norton, 2003; Song and Robinson, 2004; Thompson et al., 2005) evidence exists for month-to-month changes to the stratospheric flow feedback to the troposphere, affecting its circulation. Model results show that trends in the SH stratosphere can affect high-latitude surface climate (Gillett and Thompson, 2003).