184.108.40.206 Consequences for Lifetimes
220.127.116.11.1 Lifetime definition
The global instantaneous atmospheric lifetime of a trace gas in the atmosphere is obtained by integrating the loss frequency l over the atmospheric domain considered. The integral must be weighted by the distribution of the trace gas on which the sink processes act. Considering a distribution of the trace gas C(x,y,z,t), a global instantaneous lifetime derived from the budget can be defined as:
τglobal = ∫ C dv / ∫ C l dv (7.4)
where dv is an atmospheric volume element. This expression can be averaged over one year to determine the global and annual mean lifetime. The global atmospheric lifetime (also called ‘burden lifetime’ or ‘turnover lifetime’) characterises the time required to turn over the global atmospheric burden.
The global atmospheric lifetime characterises the time to achieve an e-fold decrease of the global atmospheric burden. Unfortunately τglobal is a constant only in very limited circumstances. In the case where the loss rate depends on the burden, the perturbation or pulse decay lifetime (τpert) is introduced (see Velders et al., 2005). The perturbation lifetime is used to determine how a one-time pulse emission may decay as a function of time as needed for the calculation of Global Warming Potentials (GWPs). The perturbation lifetime can be distinctly different from the global atmospheric lifetime. For example, if the CH4 abundance increases above its present-day value due to a one-time emission, the time it takes for CH4 to decay back to its background value is longer than its global unperturbed atmospheric lifetime. This delay occurs because the added CH4 will cause a suppression of OH, in turn increasing the background CH4. Such feedbacks cause the decay time of a perturbation (τpert) to differ from the global atmospheric lifetime (τglobal). In the limit of small perturbations, the relation between the perturbation lifetime of a gas and its global atmospheric lifetime can be derived from a simple budget relationship as τpert = τglobal / (1– f), where the sensitivity coefficient f = dln(τglobal) / dln(B). Prather et al. (2001) estimated the feedback of CH4 to tropospheric OH and its lifetime and determined a sensitivity coefficient f = 0.28, giving a ratio τpert / τglobal of 1.4. Stevenson et al. (2006), from 25 CTMs, calculate an ensemble mean and 1 standard deviation uncertainty in present-day CH4 global lifetime τglobal of 8.7 ± 1.3 years, which is the AR4 updated value. The corresponding perturbation lifetime that should be used in the GWP calculation is 12 ± 1.8 years.
Perturbation lifetimes can be estimated from global models by simulating the injection of a pulse of gas and tracking the decay of the added amount. The pulse of added CO, HCFCs or hydrocarbons, by causing the concentration of OH to decrease and thus the lifetime of CH4 to increase temporarily, causes a buildup of CH4 while the added burden of the gas persists. Thus, changes in the emissions of short-lived gases can generate long-lived perturbations as shown in global models (Derwent et al., 2001; Wild et al., 2001; Collins et al., 2002). Changes in tropospheric ozone accompany the CH4 decay on a 12-year time scale as an inherent component of this mode, a key example of chemical coupling in the troposphere. Any chemically reactive gas, whether a greenhouse gas or not, will produce some level of indirect greenhouse effect through its impact on atmospheric chemistry.
18.104.22.168.2 Changes in lifetime
Since OH is the primary oxidant in the atmosphere of many greenhouse gases including CH4 and hydrogenated halogen species, changes in OH will directly affect their lifetime in the atmosphere and hence their impact on the climate system. Recent studies show that interannual variations in the chemical removal of CH4 by OH have an important impact on the variability of the CH4 growth rate (Johnson et al., 2002; Warwick et al., 2002; J. Wang et al., 2004). Variations in CH4 oxidation by OH contribute to a significant fraction of the observed variations in the annual accumulation rate of CH4 in the atmosphere. In particular, the 1992 to 1993 anomaly in the CH4 growth rate can be explained by fluctuations in OH and wetland emissions after the eruption of Mt. Pinatubo (J. Wang et al., 2004). CH4 variability simulated by Johnson et al. (2002), resulting only from OH sink processes, also indicates that the ENSO cycle is the largest component of that variability. These findings are consistent with the variability of global OH reconstructed by Prinn et al. (2005), Manning and Keeling (2006) and Bousquet et al. (2005), which is strongly affected by large-scale wildfires as in 1997 to 1998, by El Niño events and by the Mt. Pinatubo eruption.
The effect of climate change on tropospheric chemistry has been investigated in several studies. In most cases, the future CH4 lifetime increases when emissions increase and climate change is ignored (Brasseur et al., 1998; Stevenson et al., 2000; Hauglustaine and Brasseur, 2001; Prather et al., 2001; Hauglustaine et al., 2005). This reflects the fact that increased levels of CH4 and CO depress OH, reducing the CH4 sink. However, climate warming increases the temperature-dependent CH4 oxidation rate coefficient (Johnson et al., 1999), and increases in water vapour and NOx concentrations tend to increase OH. In most cases, these effects partly offset or exceed the CH4 lifetime increase due to emissions. As a consequence, the future CH4 lifetime calculated by Brasseur et al. (1998), Stevenson et al. (2000) and Hauglustaine et al. (2005) remains relatively constant (within a few percent) over the 21st century. In their transient simulation over the period 1990 to 2100, Johnson et al. (2001) find a dominant effect of climate change on OH in the free troposphere so that the global CH4 lifetime declines from about 9 years in 1990 to about 8.3 years by 2025 but does not change significantly thereafter. Hence the evolution of the CH4 lifetime depends on the relative timing of NOx and hydrocarbon emission changes in the emission scenarios, causing the calculated CH4 increase in 2100 to be reduced by 27% when climate change is considered. Stevenson et al. (2006) reach a similar conclusion about the relatively constant CH4 lifetime. As a result of future changes in emissions, the CH4 steady-state lifetime simulated by 25 state-of-the-art CTMs increases by 2.7 ± 2.3% in 2030 from an ensemble mean of 8.7 ± 1.3 years for the present day (mean ± 1 standard deviation) for a current legislation scenario of future emissions of ozone precursors. Under the 2030 warmer climate scenario, the lifetime is reduced by 4.0 ± 1.8%: the total effect of both emission and climate changes reduces the CH4 lifetime by only 1.3%.