Other GHGs. IPCC Climate Change 2007: The Physical Science Basis

Transcription

1 Other GHGs IPCC Climate Change 2007: The Physical Science Basis 1

2 Atmospheric Chemistry and other long-lived GHG during the industrial period The radiative forcing of climate during the period due to CO 2 was about 1.5 Wm -2. The total forcing by all gases was ~3 Wm -2 (not including water vapor). Methane, tropospheric ozone, the halogenated hydrocarbons, and nitrous oxide (N 2 O) make up the remainder. Methane. The rapid increase in methane over the last 100 years has produced a climate forcing of ~0.5 W m -2, or 1/3 of that due to the increase in CO 2. Methane is produced biologically at Earth's surface by organisms in anoxic environments (such as wet lands, deep soils, landfills, etc.) IPCC estimates 600 Tg CH 4 yr -1 are produced of which anthropogenic sources are thought to contribute ~1/2 through agriculture, fossil fuel use and waste disposal. Sources and Sinks of CH 4 Evans, New Phytologist,

3 Biological Production of CH 4 CH 4 produced by methanogenic bacteria: grow only in low O 2 environments fermentation of cellulose and other organic material swamps, marshes, rice paddy fields rumina of cows and sheep. Warneck, Chemistry of the Natural Atmosphere, 2000 Methane is destroyed primarily in the troposphere (90% of the loss) when it is oxidized by the hydroxyl radical (OH) - ESE/Ch/Ge 171: OH + CH 4 CH 3 + H 2 O OH is the premier oxidant in Earth's atmosphere. It is formed in the daytime via gas phase photochemistry; its major source is: O 3 + hν O ( 1 D) + O 2 O ( 1 D) + H 2 O OH + OH hν represents a photon of wavelength < 315 nm. O ( 1 D) is the first electronically excited oxygen atom. In the troposphere, OH has an average mixing ratio of ~ 1 x 10 6 molecule cm -3. At 273 K, the rate coefficient for the reaction of OH with methane is about 3.5 x cm 3 molecule -1 s -1. Thus: τ CH4 = [CH 4 ]/(k OH+CH4 [OH][CH 4 ]) = 1/( k OH+CH4 [OH]) = 1/(2.8 x 10 8 s -1 ) = 9 years. Because the reaction of OH with CH 4 is an important sink of tropospheric OH, the lifetime of methane is not truly independent of the concentration of methane. As the concentration of methane increases, the lifetime of methane also increases - a positive feedback. 3

4 As illustrated in the IPCC CH 4 figure, the growth rate of atmospheric methane has been highly variable. During the 1980s, the trend was fairly constant at 10 ppbv /yr. During the 1990s the rate has oscillated with some years with essentially no trend (e.g , ). There is no consensus view on what causes the variability in the methane growth rate. Some research has pointed to variability in the sources driven by changes in the atmospheric hydrological cycle; others have pointed to possible variability in [OH] driven by changes in stratospheric ozone. Given the lack of understanding of recent trends, prediction of future methane concentration remains highly uncertain. Atmospheric Time Series of CH 4 Recent Data (ppm) (ppb) ppm (a) Solid line shows globally averaged CH 4 dry air mole fractions; dashed line is a deseasonalized trend curve fitted to the global averages. (b) Instantaneous growth rate for globally averaged atmospheric CH 4 (solid line; dashed lines are ±1σ [Steele et al., 1992]). The growth rate is the time-derivative of the dashed line in Figure 1a. Circles are annual increases, calculated from the trend line in Figure 1a as the increase from January 1 in one year to January 1 in the next. (c) Residuals from a function fitted to zonal averages for CH 4 (solid line), CO (dotted line), and MOPITT CO (circles) for polar northern latitudes (53.1 N to 90 N). (d) Same as Figure 1c, but for the tropics (17.5 S to 17.5 N). Dlugokencky, E. J., et al. (2009), Observational constraints on recent increases in the atmospheric CH 4 burden, Geophys. Res. Lett., 36, L18803, doi: /2009GL

5 The growth in atmospheric methane also contributes to moistening of the stratosphere. Air entering the stratosphere is desiccated to 3-4 ppmv by the very cold temperatures present at the tropical tropopause ( K). As air remains in the stratosphere the methane is oxidized to CO 2 and H 2 O and the moisture increases to nearly 8 ppmv in the 'oldest' air (i.e. where [CH 4 ] 0). Because H 2 O in the lower stratosphere is such a good GHG (it is cold and has high extinction), this source of H 2 O must be considered in future climate predictions. IPCC Climate Change 2007: The Physical Science Basis 5

6 Tropospheric Ozone. Ozone is a very efficient green house gas and the estimated change in tropospheric ozone drives a warming nearly as large as methane. In the stratosphere, ozone losses driven by both dynamical and chemical changes have produced a cooling at the surface. The losses of stratospheric ozone are widely expected to be reversed as a result of regulation of the release of long-lived halogenated compounds. Ozone, O 3, in the troposphere is an important source of oxidant via formation of the hydroxyl radical. It is also an important ingredient of smog and contributes to poor urban (and regional) air quality. Ozone in the troposphere is highly variable reflecting its variable (and often short) lifetime, as well as variability of the sources. Ozone is often near 0 in the marine boundary layer far from continents and can exceed 100 ppbv in the upper troposphere and downwind of urban centers. Presently ozone mixing ratio averages ~ 50 ppbv. A small fraction of the tropospheric source of ozone is from the stratosphere (10%) while a much larger source is derived from in situ photochemistry. Hydrocarbons + NO x + sunlight (hν) produce significant amounts of ozone: e.g.: OH + CO + O 2 HO 2 + CO 2 HO 2 + NO NO 2 + OH NO 2 + hν (350 nm) NO + O O + O 2 O 3 Net: CO + 2 O 2 CO 2 + O 3. In urban areas, hydrocarbons of much higher reactivity with OH (e.g octane) are responsible for ozone production. As a result of increased hydrocarbon and NOx emissions, it is estimated by the IPCC that tropospheric ozone has increased from 25 ppbv in preindustrial times to ~ 50 ppbv today. This estimate is highly uncertain. Unlike other long lived gases, no record of ozone exists from ice cores. There are some early measurements of ozone in Europe with iodine cells that suggest a large increase. Since 1970 when good records of ozone began, there has been a small, but statistically nonsignificant trend observed in mid-tropospheric ozone (IPCC Figure 4.8) Mid-tropospheric O3 abundance (ppb) in northern mid-latitudes (36 N-59 N) for the years 1970 to Observations between 630 and 400 hpa are averaged from nine ozone sonde stations (four in North America, three in Europe, two in Japan). Values are derived from the residuals of the trend fit with the trend added back to allow for discontinuities in the instruments. Monthly data (points) are shown with a smoothed 12-month-running mean (line) 6

7 CFC-11 and CFC-12 (and the other halogenated solvents such as CCl 4 ) are being replaced by other compounds that have much shorter lifetimes. These compounds typically have at least one hydrogen that can be removed by reaction with OH (similar to the chemistry of methane described above). As a result a much smaller atmospheric burden results. Figure 4.3 from the IPCC 2001 shows the recent history of these compounds. These compounds are also regulated by the Montreal Protocol and it is expected that they will also be phased out in the coming years. Halocarbons. Halocarbons have been manufactured for a myriad of uses. CFC-11 (CFCl 3 ) and CFC-12 (CF 2 Cl 2 ) were used extensively as refrigerants and solvents. Manufacture of these compounds is now highly regulated by the Montreal Protocol (and its amendments). These two compounds with a combined concentration of less than 1 ppbv contribute 0.25 Wm -2 forcing! Remember that the change in CO 2 (~ 80 ppmv = 80,000 times larger) contributes a forcing only 6 times that of these two halocarbons. Note also that the concentration and forcing are shown on the same plot and the relationship is linear. What does that tell us about the optical depth of these compounds in the atmosphere? The forcing by these compounds peaked in the late 1990s and will produce a negative forcing of 0.25 Wm -2 over the next 100 years as they are cleansed from the atmosphere. The lifetime of these compounds is very long because the only mechanism to destroy them is photolysis in the middle and upper stratosphere (leading to other worries - ozone depletion by halogens). 7

8 Nitrous oxide. Nitrous oxide is a very interesting trace gas. It is produced biologically at the surface as a leakage of fixed nitrogen within the nitrogen cycling. Figure 4.2 of the IPCC shows the history of N 2 O from ice core and the recent instrumental record. The concentration of N 2 O was very stable over the last 1000 years. During the industrial period its concentration has been growing steadily at ~ 0.25% yr -1. It is thought that the increase reflects the application of large amounts of fixed nitrogen (fertilizer) to soils. In tropical regions, fertility is often phosphorous rather than nitrogen limited and large emissions of N 2 O have been observed following application of chemical fertilizers. It is very unclear whether the trends in this gas can be stopped by changes in agricultural practices. Nitrous oxide is lost (almost exclusively) in the stratosphere when it is photolysed or reacts with O( 1 D). A small amount of this latter loss produces NO. This is very important for stratospheric chemistry as ozone production via photolysis of molecular O 2 : O 2 + hν O + O; O + O 2 O 3 is to first order balanced by loss via: NO + O 3 NO 2 + O O + NO 2 NO + O 2 Increases in nitrous oxide are thus expected to lead to decreases in stratospheric ozone. Evolution of anthropogenic CO 2 sources and sinks between 1765 and S Khatiwala et al. Nature 462, (2009) doi: /nature