IPCC Fourth Assessment Report: Climate Change 2007
Climate Change 2007: Working Group I: The Physical Science Basis

2.3.6 Ozone

In the TAR, separate estimates for RF due to changes in tropospheric and stratospheric ozone were given. Stratospheric ozone RF was derived from observations of ozone change from roughly 1979 to 1998. Tropospheric ozone RF was based on chemical model results employing changes in precursor hydrocarbons, CO and nitrogen oxides (NOx). Over the satellite era (since approximately 1980), stratospheric ozone trends have been primarily caused by the Montreal Protocol gases, and in Ramaswamy et al. (2001) the stratospheric ozone RF was implicitly attributed to these gases. Studies since then have investigated a number of possible causes of ozone change in the stratosphere and troposphere and the attribution of ozone trends to a given precursor is less clear. Nevertheless, stratospheric ozone and tropospheric ozone RFs are still treated separately in this report. However, the RFs are more associated with the vertical location of the ozone change than they are with the agent(s) responsible for the change. Stratospheric Ozone

The TAR reported that ozone depletion in the stratosphere had caused a negative RF of –0.15 W m–2 as a best estimate over the period since 1750. A number of recent reports have assessed changes in stratospheric ozone and the research into its causes, including Chapters 3 and 4 of the 2002 Scientific Assessment of Ozone Depletion (WMO, 2003) and Chapter 1 of IPCC/TEAP (2005). This section summarises the material from these reports and updates the key results using more recent research.

Global ozone amounts decreased between the late 1970s and early 1990s, with the lowest values occurring during 1992 to 1993 (roughly 6% below the 1964 to 1980 average), and slightly increasing values thereafter. Global ozone for the period 2000 to 2003 was approximately 4% below the 1964 to 1980 average values. Whether or not recently observed changes in ozone trends (Newchurch et al., 2003; Weatherhead and Andersen, 2006) are already indicative of recovery of the global ozone layer is not yet clear and requires more detailed attribution of the drivers of the changes (Steinbrecht et al., 2004a (see also comment and reply: Cunnold et al., 2004 and Steinbrecht et al., 2004b); Hadjinicolaou et al., 2005; Krizan and Lastovicka, 2005; Weatherhead and Andersen, 2006). The largest ozone changes since 1980 have occurred during the late winter and spring over Antarctica where average total column ozone in September and October is about 40 to 50% below pre-1980 values (WMO, 2003). Ozone decreases over the Arctic have been less severe than have those over the Antarctic, due to higher temperature in the lower stratosphere and thus fewer polar stratospheric clouds to cause the chemical destruction. Arctic stratospheric ozone levels are more variable due to interannual variability in chemical loss and transport.

The temporally and seasonally non-uniform nature of stratospheric ozone trends has important implications for the resulting RF. Global ozone decreases result primarily from changes in the lower stratospheric extratropics. Total column ozone changes over the mid-latitudes of the SH are significantly larger than over the mid-latitudes of the NH. Averaged over the period 2000 to 2003, SH values are 6% below pre-1980 values, while NH values are 3% lower. There is also significant seasonality in the NH ozone changes, with 4% decreases in winter to spring and 2% decreases in summer, while long-term SH changes are roughly 6% year round (WMO, 2003). Southern Hemisphere mid-latitude ozone shows significant decreases during the mid-1980s and essentially no response to the effects of the Mt. Pinatubo volcanic eruption in June 1991; both of these features remain unexplained. Pyle et al. (2005) and Chipperfield et al. (2003) assessed several studies that show that a substantial fraction (roughly 30%) of NH mid-latitude ozone trends are not directly attributable to anthropogenic chemistry, but are related to dynamical effects, such as tropopause height changes. These dynamical effects are likely to have contributed a larger fraction of the ozone RF in the NH mid-latitudes. The only study to assess this found that 50% of the RF related to stratospheric ozone changes between 20°N to 60°N over the period 1970 to 1997 is attributable to dynamics (Forster and Tourpali, 2001). These dynamical changes may well have an anthropogenic origin and could even be partly caused by stratospheric ozone changes themselves through lower stratospheric temperature changes (Chipperfield et al., 2003; Santer et al., 2004), but are not directly related to chemical ozone loss.

At the time of writing, no study has utilised ozone trend observations after 1998 to update the RF values presented in Ramaswamy et al. (2001). However, Hansen et al. (2005) repeated the RF calculation based on the same trend data set employed by studies assessed in Ramaswamy et al. (2001) and found an RF of roughly –0.06 W m–2. A considerably stronger RF of –0.2 ± 0.1 W m–2 previously estimated by the same group affected the Ramaswamy et al. (2001) assessment. The two other studies assessed in Ramaswamy et al. (2001), using similar trend data sets, found RFs of –0.01 W m–2 and –0.10 W m–2. Using the three estimates gives a revision of the observationally based RF for 1979 to 1998 to about –0.05 ± 0.05 W m–2.

Gauss et al. (2006) compared results from six chemical transport models that included changes in ozone precursors to simulate both the increase in the ozone in the troposphere and the ozone reduction in the stratosphere over the industrial era. The 1850 to 2000 annually averaged global mean stratospheric ozone column reduction for these models ranged between 14 and 29 Dobson units (DU). The overall pattern of the ozone changes from the models were similar but the magnitude of the ozone changes differed. The models showed a reduction in the ozone at high latitudes, ranging from around 20 to 40% in the SH and smaller changes in the NH. All models have a maximum ozone reduction around 15 km at high latitudes in the SH. Differences between the models were also found in the tropics, with some models simulating about a 10% increase in the lower stratosphere and other models simulating decreases. These differences were especially related to the altitude where the ozone trend switched from an increase in the troposphere to a decrease in the stratosphere, which ranged from close to the tropopause to around 27 km. Several studies have shown that ozone changes in the tropical lower stratosphere are very important for the magnitude and sign of the ozone RF (Ramaswamy et al., 2001). The resulting stratospheric ozone RF ranged between –0.12 and +0.07 W m–2. Note that the models with either a small negative or a positive RF also had a small increase in tropical lower stratospheric ozone, resulting from increases in tropospheric ozone precursors; most of this increase would have occurred before the time of stratospheric ozone destruction by the Montreal Protocol gases. These RF calculations also did not include any negative RF that may have resulted from stratospheric water vapour increases. It has been suggested (Shindell and Faluvegi, 2002) that stratospheric ozone during 1957 to 1975 was lower by about 7 DU relative to the first half of the 20th century as a result of possible stratospheric water vapour increases; however, these long-term increases in stratospheric water vapour are uncertain (see Sections 2.3.7 and 3.4).

The stratospheric ozone RF is assessed to be –0.05 ± 0.10 W m–2 between pre-industrial times and 2005. The best estimate is from the observationally based 1979 to 1998 RF of –0.05 ± 0.05 W m–2, with the uncertainty range increased to take into account ozone change prior to 1979, using the model results of Gauss et al. (2006) as a guide. Note that this estimate takes into account causes of stratospheric ozone change in addition to those due to the Montreal Protocol gases. The level of scientific understanding is medium, unchanged from the TAR (see Section 2.9, Table 2.11).