Other GHGs
IPCC Climate Change 2007: The Physical Science Basis
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Atmospheric Chemistry and other long-lived GHG
during the industrial period 1750-2000
The radiative forcing of climate during the period 1750-2000 due to
CO2 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 (N2O) 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
CO2. 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 CH4 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 CH4
Evans, New Phytologist, 2007.
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Biological Production of CH4
CH4 produced by “methanogenic” bacteria:
• grow only in low O2 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 + CH4 → CH3 + H2O
OH is the premier oxidant in Earth's atmosphere. It is formed in the
daytime via gas phase photochemistry; its major source is:
O3 + hν →O (1D) + O2
O (1D) + H2O → OH + OH
hν represents a photon of wavelength < 315 nm. O (1D) is the first
electronically excited oxygen atom. In the troposphere, OH has an
average mixing ratio of ~ 1 x 106 molecule cm-3. At 273 K, the rate
coefficient for the reaction of OH with methane is about 3.5 x 10-15 cm3
molecule-1 s-1. Thus:
τCH4 = [CH4]/(kOH+CH4 [OH][CH4]) = 1/( kOH+CH4 [OH]) = 1/(2.8 x 108 s-1)
= 9 years.
Because the reaction of OH with CH4 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.
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As illustrated in the
IPCC CH4 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. 1992-1993,
1999-2000).
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.
(ppm)
Atmospheric Time Series of CH4 – Recent Data
(ppb)
1.775 ppm
(a) Solid line shows globally averaged
CH4 dry air mole fractions; dashed line is
a deseasonalized trend curve fitted to the
global averages. (b) Instantaneous growth
rate for globally averaged atmospheric
CH4 (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 CH4 (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 CH4 burden, Geophys.
Res. Lett., 36, L18803, doi:
10.1029/2009GL039780.
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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
(190-200 K). As air remains in the stratosphere the
methane is oxidized to CO2 and H2O and the moisture
increases to nearly 8 ppmv in the 'oldest' air (i.e.
where [CH4] → 0). Because H2O in the lower
stratosphere is such a good GHG (it is cold and has
high extinction), this source of H2O must be
considered in future climate predictions.
IPCC Climate Change 2007: The Physical Science Basis
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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, O3, 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 + NOx + sunlight (hν) produce
significant amounts of ozone: e.g.:
OH + CO + O2 → HO2 + CO2
HO2 + NO → NO2 + OH
NO2 + hν (350 nm) → NO + O
O + O2 → O 3
Net: CO + 2 O2 → CO2 + O3.
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 1996. 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)
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CFC-11 and CFC-12 (and the
other halogenated solvents
such as CCl4) 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 (CFCl3) and
CFC-12 (CF2Cl2) 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 CO2 (~ 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).
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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 N2O
from ice core and the recent
instrumental record.
The concentration of N2O 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 N2O 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(1D). A small amount of this latter loss produces NO. This is very
important for stratospheric chemistry as ozone production via photolysis of molecular
O2:
O2 + hν →O + O; O + O2 →O3
is to first order balanced by loss via:
NO + O3 → NO2 + O
O + NO2 → NO + O2
Increases in nitrous oxide are thus expected to lead to decreases in stratospheric ozone.
Evolution of anthropogenic CO2 sources and sinks between 1765 and 2005.
S Khatiwala et al. Nature 462, 346-349 (2009) doi:10.1038/nature08526
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