2. The global methane budget and the role of methane in climate forcing
2.4 Methane and the hydroxyl radical
Atmospheric oxidation by OH in the troposphere is the dominant removal mechanism for methane. The atmospheric concentration of OH will, therefore, be the dominant factor controlling the lifetime of methane. The primary source of hydroxyl in the troposphere results from the photolysis (the absorption of a photon of solar radiation, hv, with sufficient energy to break the molecule) of ozone producing an electronically excited atomic oxygen,
O3 + hv O2 + O(1D) Eq. 2.5 The excited atomic oxygen will predominantly react with molecular oxygen (O2) or nitrogen (N2) to relinquish the extra energy it carries, however a small fraction will react with water vapor to produce hydroxyl:
O(1D) + H2O 2OH Eq. 2.6 Note that only the excited form of atomic oxygen can react with water vapor (H2O) to produce OH. Thus the production of hydroxyl is favored by high levels of incident solar radiation and higher concentrations of water vapor making the tropical lower troposphere a globally important region for OH.
In addition to the primary production of OH from water vapor, OH can be recycled. The reaction of OH with methane, seen above, produces methyl peroxy radicals which may then react with nitrogen monoxide (NO) by:
NO + CH3O2 (+O2) CH2 + HO2 + NO2 Eq. 2.7 Note that this reaction implicitly includes a second, separate reaction (the addition of O2) that happens almost instantaneously after the first. The hydroperoxyl radical (HO2), produced above can then be recycled back to OH by further reaction with NO,
NO + HO2 NO2 + OH Eq. 2.8
Alternatively, and increasingly important when concentrations of NO are low, the HO2 can react with itself by,
HO2 + HO2 H2O2 + O2 Eq. 2.9 The hydrogen peroxide (H2O2) produced is relatively long-lived to further chemical reactions (on the order of days). It is, however, quite soluble meaning H2O2 is readily removed from the atmosphere by either wet or dry deposition processes, with the consequence that the OH is not recycled. In addition to self-reaction, HO2 may also react with the methyl peroxy radical (CH3O2) and generate products that similarly lead to an inefficient recycling of OH. The balance between the self-reaction of peroxy radicals and the recycling of OH through reaction with NO illustrates the important role the concentration of NO plays in affecting the concentration of OH in the atmosphere.
In addition to reaction with methane, OH can also react with a wide variety of organic compounds found in the atmosphere;
it is referred to as the ‘detergent’ of the atmosphere for the role it plays in initiating atmospheric oxidation and the eventual removal of compounds from the atmosphere. Of note, OH may react with carbon monoxide (CO) by:
OH + CO (+O2) CO2 + HO2 Eq. 2.10 Carbon monoxide is one of the intermediate products of methane oxidation and is also emitted into the atmosphere in
Loss process Magnitude of sink, Tg CH4/y Lifetime to sink, years
Tropospheric OH 390–490 10.1–12.6
Tropospheric chlorine 15–40 125–330
Uptake in dry soils 25–40 125–200
Chemical loss in stratosphere 30–40 125–165
Total loss 460–610 8.1–10.7
Table 2.2 Summary of methane loss processes expressed as the loss rate per year and as atmospheric lifetime. The atmospheric burden used in the calculation of lifetime is 4932 Tg CH4 (Prather et al. 2012). The estimate of loss to tropospheric OH is from Prather et al. (2012) with their stated uncertainties. Other estimates are as described in the early part of Sect. 2.3.
Box 2.1 Relationship between atmospheric lifetime and the response of atmospheric concentration to changes in emissions Control of short-lived climate forcers is of interest for climate
change mitigation because their atmospheric concentration, and hence the radiative forcing of the constituents (gas or particle) on climate, responds rapidly to changes in emissions due to their short atmospheric lifetimes. To illustrate this point, results from a simple box model are used to investigate the response of the atmospheric concentration to changes in emissions for two gases with different atmospheric lifetimes.
The first is methane, with a lifetime that is sufficiently long (9.1 years) that it has only small spatial variations in concentration, relative to the global average, due to local or regional processes. Because methane is well-mixed, the response of the global average methane concentration to changes in emissions can be estimated by assuming that the atmosphere (the troposphere) behaves like a single well-mixed box. The response of methane will be compared with that of nitrous oxide (N2O), a long-lived greenhouse gas with a lifetime of ~ 130 years (Prather et al. 2012; SPARC 2013). Note that for methane, the feedback of changes in the concentration of methane on the methane lifetime has been included using an assumed 0.25% decrease in the loss rate for a 1% increase in concentration (Myhre et al. 2013), with the 9.1 year lifetime specified for the concentration of methane in 2010. Although there is a similar, though weaker, feedback of concentration on lifetime for N2O it has not been considered here – a constant lifetime of 131 years has been used.
Figure 2.3 presents the results from the box model simulations of the global average concentration of N2O and methane as a function of time. The emissions for both species increase with time to maintain a constant 0.25% per year increase in concentration up to 2010, close to the currently observed rate of increase for both methane and N2O. At 2010, three scenarios are explored: a case where emissions remain constant at the 2010 rate; an instantaneous decrease of the 2010 emission rate by 10%; and an instantaneous decrease of the 2010 emission rate by 20%. Methane establishes a new steady-state concentration
reflecting the change in emissions after approximately 40 years, with much of the adjustment occurring within 20 years of the change. In contrast, the N2O concentration continues to change (increase) past 2100. The different time scales to establish a new steady state concentration in balance with the constant emissions is one aspect of the different atmospheric lifetimes.
A second aspect of the different behavior is the level at which the new steady-state is established. For the case with constant emissions after 2010, methane stabilizes at a concentration approximately 3% higher than the 2010 concentration while the concentration of N2O at 2100 is 16% higher than at 2010 and still increasing. The initial 0.25% per year increase in concentration represents the same relative increase in atmospheric burden for each species. Although the much smaller fractional loss of N2O each year, due to the much longer atmospheric lifetime, means that the annual increase in N2O is much greater relative to removal compared with methane. As a result, to establish a balance between constant emissions after 2010 and atmospheric removal the N2O loss must increase by a larger relative amount than for CH4. Since the annual loss is proportional to the concentration, the N2O concentration must increase more than the concentration of methane to establish a new steady state.
The results show that for species with atmospheric lifetimes similar to that of methane (i.e. years to one or two decades), the atmospheric concentration will stabilize rapidly after any increase in emissions stops and decreases in emissions will rapidly be reflected in decreased atmospheric concentrations, with the atmospheric lifetime being indicative of the timescales for the concentration to adjust. For long-lived species, the period of time over which atmospheric concentrations adjust to changes in emissions will be longer. Additionally, the change in emissions required to stabilize increasing concentrations will be greater than for short-lived species, assuming the same relative rates of increase in atmospheric concentration.
Constant 10% decrease 20% decrease
N2O-like species CH4-like tracer
1980 2000 2020 2040 2060 2080 2100
0.7 0.7
0.8 0.9 1.0 1.1 1.2
1980 2000 2020 2040 2060 2080 2100
0.8 0.9 1.0 1.1 1.2
Relative concentration Relative concentration
Fig. 2.3 Evolution of the global average concentration of two chemical species with time calculated using a simple one-box model of the atmosphere.
The left-hand panel shows the evolution of a nitrous oxide (N2O)-like species with an assumed lifetime of 131 years and the right-hand panel shows the evolution of a methane (CH4)-like tracer with a lifetime of 9.1 years. The methane tracer includes the effect of changes in concentration on its own lifetime, where the total loss decreases by 0.25% for a 1% increase in concentration. Both tracers increase at a rate of 0.25% per year until 2010, at which point the emissions are assumed to either remain constant, instantaneously decrease by 10% or instantaneously decrease by 20%.
The concentrations of both species are normalized to their value at the beginning of 2010.
large quantities by incomplete combustion processes, including industrial processes, internal combustion engines and biomass burning (Duncan et al. 2007).
The broad outline of OH production and cycling in the atmosphere given above illustrates how the primary production of OH is sensitive to ozone, solar radiation and water vapor concentrations.
The reactions also show that increasing concentrations of methane and CO will act to depress the concentration of OH, while increased concentrations of NO will generally have the effect of increasing OH through more efficient recycling. As discussed in Box 2.2, increased concentrations of NO will also lead to the photochemical production of ozone, which will further increase OH concentration since the products of ozone photolysis participate in the primary production mechanism of OH. These general aspects of atmospheric chemistry are important to understand how climate change and emissions may influence methane lifetime.
2.4.1
Observation-based estimates of hydroxyl
While challenging, the concentration of OH can be directly measured in the atmosphere (e.g. Wennberg et al. 1995).
However, due to the high reactivity of OH it is extremely short-lived, with a lifetime on the order of seconds (Lelieveld et al.
2004), and the atmospheric concentration is therefore strongly dependent on local rates of production and destruction, which are highly variable in space and time. For example, the concentration of OH has been found to vary by a factor of two between measurements made under a cloud and measurements made beside the cloud, with the variation matching the observed variation in the rate of ozone photolysis (Mauldin et al. 2001).
Therefore, it is not possible to extrapolate measurements of OH to derive a global average concentration that would be useful for constraining methane losses.
Observations of methyl chloroform (MCF), a chemical with no known natural sources, that depletes stratospheric ozone and whose use was subsequently phased-out under the Montreal Protocol, has provided an opportunity to estimate the global abundance of OH. Like methane, the dominant removal mechanism of MCF from the atmosphere is through reaction with OH and when the continued production and use of MCF was discontinued, atmospheric abundances began to decrease. Surface observations, shown in Fig. 2.4, demonstrate the exponential decay of the MCF concentration with the rate Box 2.2 The photochemical production of ozone
Methane also has important impacts on climate through the role it plays in the photochemical production of ozone (O3) in the troposphere and lower stratosphere (Crutzen and Zimmermann 1991). During sunlit conditions, the chemical species nitrogen monoxide (NO) and nitrogen dioxide (NO2) rapidly pass through a cycle given by the following two reactions:
NO + O3 NO2 + O2 NO2 + hv (+O2) NO + O3
Of note, the sum of these two reactions does not result in any change of concentration for any of the participating species. Where NO2 is produced by the reaction of NO with ozone, the subsequent photolysis of NO2 merely regenerates ozone. However, NO2 may be produced by the reaction of NO with HO2 and with other peroxy radicals;
in the case of methane, CH3O2. In this case, the subsequent photolysis of NO2 will result in the net production of ozone.
This sequence of reactions gives rise to the photochemical production of ozone due to the atmospheric oxidation of organic compounds in the presence of nitrogen oxides (NOX; due to the rapid interconversion, NOX is frequently used to refer to the sum of NO and NO2).
Although the discussion of organic compounds that undergo atmospheric oxidation is limited here to methane, there is a tremendous variety of organic compounds present in the atmosphere from both anthropogenic and natural sources (e.g. Houweling et al. 1998). The details of the chemistry vary from compound to compound, sometimes in important ways, but the underlying process by which the atmospheric oxidation of organic compounds interact with NOX to photochemically produce ozone is similar. It is usual to differentiate between methane and the group of much shorter-lived organic compounds by referring to the latter as non-methane volatile organic compounds (nmVOCs).
Over industrialized regions of the world, the combination of intense sunshine, warm temperatures and high concentrations of NOX and organic compounds from anthropogenic and natural sources can result in the rapid photochemical generation of ozone in the atmosphere and give rise to large concentrations of ozone near the surface, creating air quality problems (Crutzen 1974). While more broadly, emissions of anthropogenic NOX and organic compounds, including methane, affect the concentration of ozone throughout the troposphere. Atmospheric oxidation of methane in the lowermost stratosphere can also lead to photochemical production of ozone in that region, although the NOX is likely to be of stratospheric origin (Portmann and Solomon 2007; Fleming et al. 2011).
Fig. 2.4 Observed hemispheric monthly average concentration of methyl chloroform (MCF). The hemispheric average concentrations are derived from flask samples taken approximately weekly at nine remote surface sites operated by the US National Oceanic and Atmospheric Administration (NOAA)/Global Monitoring Division following the method described by Montzka et al. (2011).
1992 1996 2000 2004 2008 2012 2016
0 20 40 60 80 100 120 140 160
MCF in air, parts per trillion
Northern hemisphere Southern hemisphere
of decay reflecting the atmospheric lifetime. Until ~1997 a significant inter-hemispheric gradient was observed, with higher concentrations in the northern hemisphere reflecting the predominance of emissions in that hemisphere. As emissions decreased to levels that were insignificant compared to the burden and removal processes, the inter-hemispheric difference became small and the global average removal rate (the exponential first-order loss rate) stabilized at around -0.181 ± 0.005/y (Montzka et al. 2011). Accounting for other (minor) loss processes, loss in the stratosphere and uptake by the oceans, it is possible to derive the MCF loss rate to OH in the troposphere from the observed rate of decrease. The MCF loss to OH can then be scaled to that of methane using the relative reaction rates of MCF and methane with OH. In this way, estimates of methane loss to OH (440 ± 52 Tg CH4/y) with relatively small estimated uncertainties have been calculated (Prather et al. 2012).
The observed decay rate of MCF can also be used to provide an estimate of the amount of interannual variability in the global average concentration of OH. Earlier estimates suggested interannual variability in OH of 7–9%, with maximum year-to-year changes on the order of 20% (Prinn et al. 2005). It is now believed that these earlier estimates of large interannual variability were complicated by the fact that ongoing emissions of MCF were still considerable. Estimates derived from MCF observations for the post-1998 period show an interannual variability of less than 3% (Montzka et al. 2011).
2.4.2
Photochemical modelling estimates of hydroxyl
Global chemical models also provide estimates of present-day OH concentrations. These models include a representation of emissions, transport, chemistry and removal of a suite of chemical compounds that captures the important chemical processes of the troposphere. A suite of current-generation models that participated in the Atmospheric Chemistry Climate Model Intercomparison Project (ACCMIP) estimated the present-day chemical lifetime of methane to reaction with tropospheric OH to be 9.3 ± 0.9 years, which corresponds to a methane sink of 530 ± 50 Tg CH4/y (Voulgarakis et al. 2013).
While an earlier multi-model study, based on a substantially different set of global models and conducted for the Task Force on Hemispheric Transport of Air Pollution (TF-HTAP), estimated a tropospheric OH methane sink of 480 ± 80 Tg CH4/y (Fiore et al. 2009). Both of these estimates were calculated as the mean of the participating models, with the ranges given as one standard deviation of the individual model estimates. The model estimates can be compared with the methane loss of 440 ± 52 Tg CH4/y derived from the observed decay of MCF.
While some individual models calculate a tropospheric OH sink that is in agreement with the observationally-derived estimate, and in fact the multi-model mean from the TF-HTAP study falls within the stated uncertainty of the observational estimate, models in general tend to overestimate methane loss to OH.
Factors contributing to the range of model estimates include uncertainties in the chemistry (Dillon and Crowley 2008;
Lelieveld et al. 2008; Fuchs et al. 2013) and the representation of non-methane hydrocarbons in the models (Voulgarakis et al.
2013). Differences in temperature, water vapor, stratospheric
ozone column or the concentration of ozone in the troposphere have also been shown to play a role (Holmes et al. 2013; Naik et al. 2013). The effect of clouds on photolysis has been shown to have only minor effects on methane lifetime on a global scale (Voulgarakis et al. 2009), although the general treatment of photolysis appears to be an important factor driving the spread in simulated present-day methane lifetime between models (Voulgarakis et al. 2013).
2.4.3
Long-term changes in hydroxyl
While model estimates of present-day methane loss to OH remain uncertain, they still provide critical insight into how tropospheric OH, and methane loss by reaction with OH, may have changed in the past and how it may change in the future.
As the magnitude of the methane sink to OH is proportional to the methane abundance, for longer-term changes in the methane budget it is necessary to analyze changes in the methane lifetime, introduced in Sect. 2.3. For changes from preindustrial to present-day conditions, the suite of ACCMIP models analyzed by Naik et al. (2013) suggest a slight decrease in the methane lifetime to tropospheric OH between 1850 and 2000, from 10.1 to 9.7 years. The small change in methane loss from 1850 to present-day is the result of a balance of positive and negative influences on OH. Changes that tended to enhance the OH concentration between 1850 and 2000 include increased water vapor and temperature (owing to climate change), increased tropospheric ozone, decreased stratospheric ozone (which allows more ultraviolet radiation into the troposphere) and increased concentrations of NO (which act to recycle OH). Factors that depressed the concentration of OH between 1850 and 2000 include the factor of 2.2 increase in the concentration of methane and higher concentrations of CO from anthropogenic emissions, both of which react directly with OH (Naik et al. 2013). Note that although the multi-model mean showed a slight decrease in the methane lifetime to tropospheric OH, individual models showed changes that ranged between an increase of one year and a decrease of one year. The range of changes in methane lifetime found for the ACCMIP models is similar to the range of changes found in earlier studies, with individual studies showing both small increases in methane lifetime (Wang and Jacob 1998; Wild and Palmer 2008) and small decreases (Berntsen et al. 1997). It is worth noting, however, that these earlier studies did not account for changes in climate, which would have affected OH through changes in temperature and water vapor.
Projections of how the methane lifetime may change to 2100 were also made as part of ACCMIP. As for the historical period, generally small changes in methane lifetime were projected between 2000 and 2100 with individual models showing changes that ranged from small increases to small decreases for each of the four future scenarios investigated – the four Representative Concentration Pathways, or RCPs, specified for the Fifth Assessment of the Intergovernmental Panel on Climate Change. For the three RCPs with the lowest radiative forcing (RCP2.6, RCP4.5, RCP6.0), the models generally showed small decreases in the methane lifetime to OH loss of less than one year for RCP4.5 and decreases of less than one-half year for RCP2.6 and RCP6.0. The clearest signal in the
Projections of how the methane lifetime may change to 2100 were also made as part of ACCMIP. As for the historical period, generally small changes in methane lifetime were projected between 2000 and 2100 with individual models showing changes that ranged from small increases to small decreases for each of the four future scenarios investigated – the four Representative Concentration Pathways, or RCPs, specified for the Fifth Assessment of the Intergovernmental Panel on Climate Change. For the three RCPs with the lowest radiative forcing (RCP2.6, RCP4.5, RCP6.0), the models generally showed small decreases in the methane lifetime to OH loss of less than one year for RCP4.5 and decreases of less than one-half year for RCP2.6 and RCP6.0. The clearest signal in the