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TROPOSPHERIC CHEMISTRY AND COMPOSITION | Oxidizing Capacity
Oxidizing Capacity DH Ehhalt, F Rohrer, and A Wahner, Forschungszentrum Jülich, Jülich, Germany Ó 2015 Elsevier Ltd. All rights reserved. This article is a revision of the previous edition article by D H Ehhalt, A Wahner, volume 6, pp 2415–2424, Ó 2003, Elsevier Ltd.
Synopsis The oxidizing capacity of the troposphere relies on oxidants such as O3, NO3, and OH. Of these, OH is by far the most important. The supply of all the oxidants is limited and thus the oxidizing capacity is finite. The factors that control OH, O3, and NO3 are reviewed and examples of their tropospheric distributions are presented. The resulting lifetimes of the most important tropospheric trace gases are listed. The possible change of the oxidant distributions under the impact of anthropogenic emissions is indicated. A quantitative measure of the oxidation capacity is derived from the total loss rate of all trace gases due to OH.
Introduction The term ‘oxidizing capacity’ is widely used, but loosely defined. Its attraction is that it conveys in a succinct manner, two major features of atmospheric chemistry: namely, that the atmosphere actively oxidizes gaseous trace constituents and pollutants, and thereby initiates their removal from the atmosphere; and also, that this capacity is finite. The first feature is not unexpected; after all molecular oxygen, O2, is a major constituent of the Earth’s atmosphere. The second is somewhat surprising – considering that O2 is present at about 21%, whereas the most abundant oxidizable trace gases are present at ppm level. The explanation is, of course, that O2 does not directly react with those molecules. Tropospheric oxidation is rather initiated by a number of oxidants, above all by the hydroxyl radical, OH, but also by ozone, O3, the nitrate radical, NO3, and to a lesser extent by chlorine (Cl) and bromine (Br) atoms. Ground state atomic oxygen, O(3P), plays only a very minor role. In the liquid phase, for example in cloud droplets, hydrogen peroxide also acts as an oxidant. All these molecules and radicals are generated within the troposphere, and the rates of the respective generation processes limit their supply and thus the oxidizing capacity of the troposphere. In the following, we briefly review the factors that control the tropospheric concentrations of the major oxidants, OH, O3, and NO3. We present examples of their tropospheric distributions, and indicate how these may have changed under the impact of anthropogenic emissions. From the superposition of the various loss processes, we derive the atmospheric lifetimes for the most important tropospheric trace gases. Finally, based on all this information we present a measure of the tropospheric oxidizing capacity through an approximate but simple expression.
The Tropospheric Chemistries of OH and O3 The tropospheric chemistries of OH and O3 are very closely interlinked. In fact, the primary production of OH is based on the photolysis of O3. At wavelength below 340 nm, this photolysis yields electronically excited oxygen atoms, O(1D) (eqn [I]): O3 þ hn/Oð1 DÞ þ O2
l 340 nm
Encyclopedia of Atmospheric Sciences 2nd Edition, Volume 6
[I]
Most of the O(1D) atoms produced in reaction [I] are quenched to ground state atomic oxygen, O(3P), in collisions with atmospheric nitrogen and oxygen molecules (eqns [II] and [III]), but a fraction collides with water vapor molecules, to form the very reactive hydroxyl radical, OH (eqn [IV]). O(1D) þ N2 / O(3P) þ N2 O(1D) þ O2 / O(3P) þ O2 O(1D) þ H2O / OH þ OH
[II] [III] [IV]
H2O is ubiquitous in the troposphere. For instance, in surface air at midlatitudes, H2O is present at a volume mixing ratio of about 1%. There about 10% of the O(1D) generate OH, because reaction [IV] has a rate constant of approximately a factor of 10 higher than the quenching reactions [II] and [III]. The conversion of O3 to OH via the reactions [I] and [IV] at the same time forms a major sink of tropospheric O3, whereas the O(3P) formed by [II] and [III] returns to O3 by recombination with O2 (eqn [V]). Oð3 PÞ þ O2 þ M/O3 þ M ðM ¼ N2 ; O2 Þ
[V]
The same holds for the O(3P) formed directly in the photolysis of O3 at wavelengths longer than 320 nm (eqn [VI]). O3 þ hn/Oð3 PÞ þ O2
l > 340 nm
[VI]
The generation of O(3P) mainly by photolysis (eqn [VI]) but also by reactions [II] and [III] and others is balanced by the loss due to recombination. During daytime this results in a steady-state concentration of a few thousand O(3P) per cm3. Given the relatively low reactivity of O(3P), this concentration is too small to be of significance for the oxidation of trace gases in the troposphere. The fact that the O atoms generated in reaction [VI] interact rapidly with one of the major constituents of air limits their possible importance. Hydroxyl radicals, on the other hand, do not react with any of the major constituents of air. Rather, they have the ability to initiate chain reactions in an O2-containing atmosphere. When reacting with trace gas molecules, OH is not consumed, but is regenerated in catalytic cycles. In this way, relatively large OH concentrations, up to 107 cm3, are maintained in the sunlit troposphere despite the high reactivity of OH toward most pollutants and other trace gases. These two properties, high
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reactivity and relatively high concentrations, make OH the most important oxidizing agent in the troposphere.
OH Reaction with Molecules
whose photolysis leads to formation of HOx and thus introduces a degree of autocatalysis. Parallel to reaction [IX], HO2 reacts directly with O3 to return OH (eqn [XII]). HO2 þ O3 / OH þ 2O2
[XII]
The simplest example of such a reaction is given by the atmospheric oxidation of carbon monoxide. The reaction of CO with OH immediately forms the stable end product CO2 (eqn [VII]).
Thus even without NO there is a certain measure of recycling. But in this case the net result of the reaction chain, [VII], [VIII], and [XII], is a destruction of O3 (eqn [XIII]).
CO þ OH / CO2 þ H
CO þ O3 / CO2 þ O2
[VII]
[XIII]
The reaction also forms a hydrogen atom, which is very reactive. The hydrogen atom rapidly combines with O2 to form a hydroperoxy radical, HO2 (eqn [VIII]).
Together with reaction [XIV], reaction [XII] forms the second important loss mechanism of tropospheric O3, again closely tied to the HOx chemistry.
H þ O2 þ M / HO2 þ M
OH þ O3 / HO2 þ O2
[VIII]
The addition of a hydrogen atom to O2 weakens the bond between the oxygen atoms, and HO2 reacts much more readily than O2. In particular, HO2 oxidizes nitric oxide, NO (eqn [IX]). NO þ HO2 / NO2 þ OH
[IX]
In the planetary boundary layer over the industrialized and highly populated continents, where daytime concentrations of NO exceed 0.1 ppb, reaction [IX] is by far the fastest HO2 reaction. Most important for our present argument is the fact that the OH radical consumed in reaction [VII] is regenerated in reaction [IX]. This is generally true: virtually all reactions of OH with molecular species lead to chain reactions that recycle OH. Reaction [IX] also highlights one of the roles of the nitrogen oxides for the OH budget. In the form of NO they quickly return the less reactive HO2 back to the highly reactive OH, thus increasing the OH concentration. The nitrogen dioxide molecule, NO2, generated in reaction [IX] photolyzes readily in the near-UV region and therefore contributes to tropospheric photochemistry (eqn [X]). NO2 þ hy / NO þ O(3P)
[X]
In the sunlit atmosphere, the lifetime of NO2 against photolysis is a few minutes. The resulting O atom immediately combines with O2 to form O3 (reaction [V]). This NO-mediated process consisting of reactions [IX], [X], and [V], in which an O atom from a peroxy radical is passed along to O3, is the major ozone formation mechanism in the troposphere. It illustrates the fact that in an NO-containing atmosphere, trace gas removal is invariably linked with the production of O3 or other photo oxidants. Supplemented by reactions [X] and [V], the hydroxyl radical reactions [VII]–[IX] combine to the net reaction shown in eqn [XI]. CO þ 2O2 þ hy / CO2 þ O3
[XI]
It does not consume OH, HO2, NO, and NO2, hence the cycle consisting of the reactions [VII]–[X] and [V] can be completed repeatedly before that chain reaction is interrupted by termination reactions. By this and other chain reactions, OH and HO2 are interconverted within a matter of seconds. Therefore, both are often lumped together as HOx. The oxidation of more complex molecules, such as hydrocarbons, leads to other peroxy radicals, RO2, which also convert NO to NO2 and augment the generation of O3. In addition, intermediates such as formaldehyde (HCHO) are formed,
[XIV]
Destruction of HOx To produce a net loss of HOx and terminate the reaction chains, HOx radicals have to react with other radicals. Eqns [XV] and [XVI] are responsible for HOx loss in clean air. OH þ HO2 / H2O þ O2 HO2 þ HO2 / H2O2 þ O2
[XV] [XVI]
The addition of OH to NO2, forming a nitric acid molecule, HNO3 (eqn [XVII]), is the dominant HOx loss reaction in the polluted atmosphere. OH þ NO2 þ M / HNO3 þ M
[XVII]
At the same time, reaction [XVII] provides the major loss mechanism for NOx. It also illustrates the tendency of atmospheric oxidation to produce acidic end products, which are eventually removed from the atmosphere by rainout.
The OH Balance Equation The various production and destruction reactions of OH can be combined in a local OH budget equation of the type shown as eqn [1], where POH and DOH stand for the local production and destruction terms of OH. d ½OH ¼ POH DOH z0 dt
[1]
In the case of a very short-lived species like OH, contributions from transport, i.e., the flux divergence, are small compared to POH and DOH, and can be neglected. The same holds for the temporal change, d[OH]/dt. POH consists of two, systematically different, terms: the primary production of OH (POH) through the photolysis of precursors such as O3 or H2O2 and the OH produced from recycling HO2 (POH). By far the largest contribution to POH on a global scale comes from the photolysis of O3, which can be derived from reactions [I]–[IV] to give eqn [2]. POH ¼
2 ½O3 $ ½H2 O $ J1 $ k4 þ/ k2 ½N2 þ k3 ½O2 þ k4 ½H2 O
[2]
J1 is the photolysis frequency of reaction [I], and the ki refers to the rate constants of reactions [II]–[IV]. Smaller contributions,
such as that from the photolysis of H2O2, are indicated by the ellipsis ‘.’. POH is determined by reactions [IX] and [XII] as eqn [3]. [3]
DOH, the destruction of OH, is given by eqn [4]. DOH ¼ ½OH $ ðk7 ½CO þ k14 ½O3 þ / þ k17 ½NO2 þ k15 ½HO2 þ /Þ
OH P(O3)
2.0 1.5
4
1.0 2 0.5
[4]
¼ ½OH$s1 OH
0
A complete chemical mechanism would contain contributions to the production and destruction of OH other than those listed above. Again, these are indicated by the dotted ellipsis ‘.’. These additional terms would, however, fall in the same categories distinguished here. The first line in eqn [4] contains the reactions of OH with molecules, i.e., those that eventually generate HO2; the second line contains the reactions of OH with radicals that lead to a net HOx loss. The whole term in parenthesis in eqn [4] represents the pseudo-first-order reaction frequency or inverse lifetime, s1 OH , of an OH radical. We now define the recycling ratio r ¼ POH/POH. This is the number of HO2 radicals converted to OH divided by the number of OH generated directly. We note that the precursors for direct OH production, mainly O3 but also H2O2, are also produced by OH chemistry, although not necessarily at the same location, and, thus, contain an element of recycling. Because of their long lifetime their recycling is slow and incomplete and their local concentration is largely determined by transport. We further note that POH also includes a slow component, namely HO2 derived from aldehyde photolysis, besides the fast recycling of HO2 generated immediately after the OH attack. Nonetheless, for most of the troposphere, r also reasonably approximates the chain length, that is, the number of times an OH is recycled via HO2 before it is removed by a radical–radical reaction. Introducing sOH and r, eqn [1] takes the simple form of eqn [5a] or eqn [5b]. ½OH ¼ 0 sOH
[5a]
½OH ¼ POH $sOH ð1 þ rÞ
[5b]
POH ð1 þ rÞ
OH (106 cm–3)
POH ¼ ½HO2 $ðk9 ½NO þ k12 ½O3 Þ
2.5 6
or
Equation [5b] demonstrates nicely that the ability of OH for chain reaction and slow recycling enhances its concentration by a factor (1 þ r) over that of a nonchain-reacting radical, whose steady-state concentration would be given merely by the product of its production and lifetime, P $ s. Tropospheric values for r vary between 0.3 and 10 depending on NOx in a manner similar to that of OH shown in Figure 1. We note that the factors r and sOH still depend on the HOx concentration. Thus, eqn [5b] by itself is not sufficient for calculating the OH concentration. It helps, however, to categorize the action of the various reactions on OH. The nitrogen oxides, NOx ¼ NO þ NO2, for example, act on OH in two different ways. In the form of NO, they accelerate the recycling from HO2 to OH (reaction [IX]; cf eqn [3]) and thus enhance the concentration of OH; in the form of NO2 they enhance the net loss of OH via reaction [XVII], and thus decrease sOH and
0.01
0.0 0.10
1.00
245
Net O3 production rate (ppbv h –1)
Tropospheric Chemistry and Composition j Oxidizing Capacity
10.00
NOx (ppbv)
Figure 1 Dependence of the OH concentration and net O3 production on NOx calculated with a steady-state model for remote, rural conditions. Adapted from Ehhalt, D.H., 1999b. Photooxidation of trace gases in the troposphere. Physical Chemistry Chemical Physics 1, 5401–5408.
therefore the concentration of OH (cf eqn [4]). The opposing actions of the two processes lead to a highly nonlinear dependence of the OH concentration on NOx (see Figure 1). At low NOx, where recycling dominates, the addition of NOx leads to an increase in OH concentration and thus to an increase in photochemical activity; at high NOx, reaction [XVII] eventually becomes the dominant loss path of OH and further addition of NOx reduces the OH concentration. The position of the maximum with respect to NOx depends to some extent on the local mix of reactive trace gases. Except in polluted areas, however, tropospheric chemistry operates in NOx regimes defined by the left flank of the OH curve in Figure 1; that is, an increase in global NOx concentrations will lead to an increase in global OH. The dependence of OH on other parameters is more monotonic, though still not linear. For the NOx regime of the global troposphere, an increase in CO or other oxidizable molecules will lead to a decrease in OH (cf reaction [VII]; eqn [4]). But a change in H2O will lead to a change of equal sign in OH (cf eqn [2]). Solar radiation acts through several channels: through POH via the photolysis of O3, but also through r and sOH, because photolysis of NO2 shifts the [NO]/[NO2] ratio. These factors all conspire to give a nearly linear relationship between tropospheric OH and solar UV radiation. Figure 1 also shows the dependence on NOx of the net formation rate of O3 given by d ½O3 =dt ¼ PO3 DO3 , where the local O3 production, PO3 , is maintained by reaction cycles such as that consisting of [VII]–[X] and [V]. The O3 destruction, DO3 , is given by reactions such as [XII], [XIV], or [I] combined with [IV]. In the clean environment considered in Figure 1, the net O3 production rate is relatively small; even at the maximum it amounts to no more than 2.3 ppb h1. The profile of the net O3 formation rate is similar to that of the OH concentration, i.e., to the photochemical activity. At low NOx concentrations, however, it begins to deviate significantly, because there the O3-destroying reactions, such as reaction [I] combined with reaction [IV], remain active, whereas the O3-producing reactions, such as the sequence [VII]–[X], decrease with decreasing NOx. Eventually, below an NOx concentration of about 0.07 ppb, photochemistry leads to a net destruction of O3.
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Chemistry of NO3
NO3 þ NO2 # N2O5
The nitrate radical, NO3, is formed primarily by the relatively slow reaction [XVIII]. NO2 þ O3 / NO3 þ O2
[XVIII]
However, in daylight NO3 is photolyzed within tens of seconds via reactions [XIXa] and [XIXb]. NO3 þ hy / NO2 þ O(3P) NO3 þ hy / NO þ O2
[XIXa] [XIXb]
In addition, NO3 reacts rapidly with NO (eqn [XX]) that is always present during daytime owing to photolysis of NO2 (reaction [X]) such that during daytime NO3 remains at very low concentrations. NO3 þ NO / NO2 þ NO2
[XXI]
This reaction is important especially at low temperatures since the thermal decomposition of N2O5 to NO2 and NO3 is strongly temperature dependent. N2O5 reacts with condensed water to form HNO3. It thus reacts rapidly upon collision with the wet surfaces such as cloud droplets or aerosol particles at relative humidities above 60%. In nights with high relative humidities, this reaction path together with the direct uptake of NO3 provides an efficient loss of NOx. Finally, NO3 reacts with organic molecules in a manner similar to OH but for most molecules much more slowly than the corresponding reactions with OH. Nevertheless, in polluted regions it can provide a significant loss for organic molecules.
[XX]
During nighttime, however, NO3 can build up to significant concentrations because the remaining loss reactions are much slower. These are (1) the heterogeneous uptake of NO3 by moist aerosol surfaces, fog, or cloud droplets; and (2) the heterogeneous uptake of N2O5, which is formed from combination of NO3 with NO2 (eqn [XXI]).
Global Distributions of OH, O3, and NO3 The tropospheric chemistry outlined above maintains global distributions of OH, O3, and NO3 that vary with latitude, longitude, daytime, and season. For reference, Figure 2 presents the mean zonal distributions of OH, O3, and NO3 for January
January
July
200
12.0
OH 400
7.1
600
4.2
800
2.0 0
200
12.0
400
7.1
O3
600
4.2
800
2.0
1000
0
200 400
Altitude (km)
Pressure (hPa)
1000
12.0
NO3
7.1
600
4.2
800
2.0
1000 S –60 –30 0
30 60 N
S –60 –30 0
30 60 N
0
Latitude Figure 2 Mean zonal distributions of OH, O3, and NO3 for January and July based on 3D chemical transport models. OH distribution (top panels, contour lines 105 OH cm3) adapted from Spivakovsky, C.M., Logan, J.A., Montzka, S.A., et al., 2000. Three-dimensional climatological distribution of tropospheric OH: update and evaluation. J. Geophys. Res. 105, 8931–8980. O3 distribution (middle panels, contour lines ppbv) adapted from Wang, Y., Jacob, D.J., Logan, J.A., 1998. Global simulation of tropospheric O3-NOx-hydrocarbon chemistry. Journal of Geophysical Research 103, 10757–10767. NO3 distribution (bottom panels, contour lines pptv), private communication Hauglustaine, D., 2000.
Tropospheric Chemistry and Composition j Oxidizing Capacity and July. They are based on 3D chemical transport models, since measurements are too sparse to allow a representative and consistent reconstruction of the average concentration fields. Nevertheless, the concentration fields shown are expected to closely approach reality. For O3, the measured seasonal variations, as well as height profiles, show good agreement with those modeled for the same sites. For OH, the globally averaged concentration agrees well with the value of 1.0 106 cm3 derived empirically from the budget of methyl chloroform, CH3CCl3. The main oxidant, OH, shows a strong latitudinal variation at all altitudes, with a broad maximum in the tropics centered at the latitude of maximum solar radiation. It also varies with altitude with a broad maximum around 4 km at all latitudes. In addition, OH, as well as the other oxidants, O3 and NO3, exhibits higher concentrations in the Northern Hemisphere whose continents provide most of the global emissions, whether natural or artificial. For the same reason, the longitude by latitude presentation (Figure 3) also indicates higher concentrations of OH over the continents. The chemistry outlined above also predicts that the increase of the anthropogenic emissions of NOx, CO, and CH4 that accompanied the industrialization and population growth of the last century must have caused significant changes in the concentrations of the tropospheric oxidants. There is indeed
experimental evidence mainly from historical records over Europe that the concentration of O3 has increased by about a factor of two over that period. Figure 4 presents the modelcalculated change in the zonally and annually averaged distribution of O3 and OH from preindustrial to present time. For O3 the models predict a general increase of about 60%, which is, however, lower in the Southern Hemisphere (þ50% in the case presented) than in the Northern Hemisphere (þ80%). The model-calculated globally averaged OH decreased by about 10%, a relatively small value. However, as Figure 4 indicates, the changes are quite unevenly distributed, with large zones of substantial positive change. The OH concentration over the continents increased strongly, whereas it decreased over the oceans, mapping the nonuniform distribution of NOx, which is short lived and mostly emitted over the continents. In fact, the global distribution of OH, as well as its trend, strongly depends on the assumed spatial and temporal distribution of the NOx emissions. The most recent analysis of the CH3CCl3 budget suggests a relatively constant mean global OH concentration over the past 20 years with a small interannual variability of 2.1 1.8% per year, but no significant long-term trend. The difference between the global change for OH and O3 is not unexpected. For O3 the increased emissions of CO and NOx reinforce each other, as more CO leads to higher formation of
200
N Pressure (hPa)
60 30 Latitude
247
0 –30
400
600
800
–60 S W
1000 –90
–120
–60
0
60
120
–60
–30
E
EQ
30
60
90
30
60
90
Latitude
Longitude N
200
Pressure (hPa)
60
Latitude
30 0 –30
400
600
800
–60 S W
–120
–60
0
60
120
E
Longitude
Figure 3 Global distribution of OH for January (top panel) and July (bottom panel) at 700 mb (contour lines 105 OH cm3). Adapted from Spivakovsky, C.M., Logan, J.A., Montzka, S.A., et al., 2000. Threedimensional climatological distribution of tropospheric OH: update and evaluation. Journal of Geophysical Research 105, 8931–8980.
1000 –90
–60
–30
EQ Latitude
Figure 4 Calculated relative change in the zonally and annually averaged distribution of O3 (top panel) and OH (bottom panel) from preindustrial to present times (contour lines in %). Adapted from Wang, Y., Jacob, D.J., 1998. Anthropogenic forcing on tropospheric ozone and OH since preindustrial times. Journal of Geophysical Research 103, 31123–31135.
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HO2, which in turn is efficiently converted to O3 by the enhanced NO. For OH they counteract each other – increased NOx induces a larger and increased CO a lower OH concentration. For NO3 one would also expect a global increase in the concentration field since preindustrial times, since both the NO2 and O3 concentrations have increased globally. However, no hindcasts have been published for NO3. Obviously, as long as the emissions of NOx, CO, and CH4 continue to grow, further changes in the global distributions of OH, O3, and NO3 have to be expected. For example, for the future emission scenario A1F1 from IPCC 2001, which presents one of the worst cases and assumes global emissions of NOx, CO, VOC, and CH4 of 243, 193, 198, and 128%, respectively, of the values in the year 2000, model calculations predict a decrease in the mean global OH of 14% and an increase in O3 of 47% by the year 2100.
The Global Loss Rate The removal of a given trace gas from the troposphere is usually generated by several different processes. In addition to the tropospheric oxidation by OH, O3, NO3, and other radicals, direct photolysis, dry deposition at the Earth’s surface, and export to the stratosphere may play a role. For a gas Xi, the global tropospheric loss rate, Li, is given by the sum of all the individual loss rates integrated over the volume of the troposphere and averaged over a full year. Correspondingly, the global tropospheric lifetime of a trace gas, sI, is defined as its tropospheric burden, Mi, divided by its global tropospheric loss rate. The contribution of the major oxidants to the global loss rate of Xi is given by eqn [6], where the brackets denote the local concentrations in units of cm3, the ki are the respective rate constants, which generally depend on temperature and thus also on location and season, v stands for tropospheric volume, t for time, and a ¼ 1 year. Li ¼
1 a
Z
ki;OH $½OH þ ki;O3 ½O3 þ ki;NO3 $½NO3 $½Xi dv$dt
v;t
¼ Mi s1 i [6]
Examples of the rate constants are given in Table 1 for a few selected organic trace molecules. Clearly, the rate constants vary over many orders of magnitude, reactions with OH having the highest and those with O3 the lowest rate constants. We also note that the ratios kOH : kO3 : kNO3 vary substantially between the various trace gases. Thus the relative contributions of the global OH, O3, and NO3 fields to the global oxidation rate vary greatly with the trace gas considered. Generally, however, tropospheric oxidation is dominated by reaction with OH. This is emphasized by Table 2, which lists the trace gases with the highest turnover in the atmosphere. The most important three trace gases, namely CO, CH4, and H2, are not attacked by O3 or NO3. In fact, significant direct oxidation by O3 is limited to the terpenes and NOx; and that by NO3 to the terpenes and dimethyl sulfide. In terms of global tropospheric removal, direct attack of O3 and of NO3 adds up to a loss rate of 3.75 Tmol per year, i.e., about 2% of the total loss rate of 180 Tmol per year given by the sum over the individual trace gas losses in Table 2 and the oxidation of their secondary products. This does not preclude that oxidation by NO3 and O3 becomes important on a local scale, or for certain molecules even on a global scale, but its impact on the total turnover of oxidizable molecules is small. Similarly, the tropospheric role of Cl and Br atoms appears small. Based on the global budgets of ethane and tetrachloroethene, the average Cl atom concentration in the Northern Hemisphere has been estimated to less than 1000 cm3 and in the Southern Hemisphere less than 2000 cm3. This is three orders of magnitude smaller than the globally averaged OH concentrations. This difference is so large that it cannot be compensated by the generally higher rate constants of hydrocarbons with Cl atoms. Finally, the role of H2O2 is pretty much limited to the oxidation of SO2 in the liquid phase (Table 2) and its overall impact is on the order of 1% of the total loss rate of 180 Tmol per year. As a consequence, the tropospheric oxidizing capacity is generally viewed as synonymous with the OH chemistry in the troposphere. Moreover, at least implicitly, the mean global OH abundance has been, and occasionally still is, used as a quantitative measure for the oxidizing capacity. Note that according to this measure, the oxidizing capacity is predicted not to have changed much over the last 100 years. At first glance this choice appears plausible, because OH determines the loss rates of many
Table 1 Rate constants of reactions of various organic trace molecules with OH, O3, and NO3. The rate constants are given for 298 K in cm3 molecule1 s1. The respective lifetimes are calculated on the basis of following concentrations: OH: 1 106 cm3, O3: 30 ppbv, NO3: 1 pptv Rate constants (at 298 K)
Organic kOH n-Butane Ethene Propene Trans-2-butene Isoprene a-Pinene Formaldehyde CH3CHO (CH3)2S
2.4 8.6 2.9 6.4 1.0 5.3 8.5 1.5 5.0
1012 1012 1011 1011 1010 1011 1012 1011 1012
Lifetimes
kO 3
kNO 3
<1022 1.6 1018 1.0 1017 1.9 1016 1.3 1017 8.7 1017 – – <1018
4.6 2.1 9.5 3.9 7.0 6.2 5.6 2.7 1.1
1017 1016 1015 1013 1013 1012 1016 1015 1012
sOH
sO 3
sNO 3
4.8 days 32 h 9.6 h 4.3 h 2.8 h 5.2 h 33 h 19 h 56 h
– 9.6 days 37 h 1.9 h 28 h 4.1 h – – >15 days
28 years 6 years 49 days 28 h 16 h 1.8 h 2.3 years 170 days 10 h
Atkinson, R., 1997. Gas-phase tropospheric chemistry of volatile organic compounds: 1. Alkanes and alkenes. J. Phys. Chem. Ref. Data 26, 215–290; Atkinson, R., Baulch, D.L., Cox, R.A., et al., 2006. Evaluated kinetic and photochemical data for atmospheric chemistry: volume II – gas phase reactions of organic species. Atmospheric Chemistry and Physics 6, 3625–4055.
Tropospheric Chemistry and Composition j Oxidizing Capacity Table 2
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Global turnover of the major tropospheric trace gases and attribution to the major oxidants
Trace gas
Global loss rate (Tmol per year)
Global lifetime
Removal (%) OH
CO H2 CH4 Isoprene SO2 NOx Terpenes C2H6 N2O (CH3)2S
1.5 months 2 years 8 years Hoursd Daysd 0.3–5 daysd Hoursd 2 months 120 years Daysd
100 38 36 8 5 3 1 0.7 0.6 0.5
85 25 90 80 30 50 20 80 – 70f
a
O3b
NO3c
Other
– – – 7 – 40 25 – – –
– – – 13 – – 55 – – 30f
15 (Soil uptake) 75 (Soil uptake) 10 (Soil uptake; stratos.) – 70 (Hetero. in clouds) 10 (Soil uptake) – 20 (Cl reaction)e 100 (Stratosphere) –
Using a mean global OH concentration of 1 106 cm3. Using a mean global O3 concentration of 30 ppbv. c Using a mean global NO3 concentration of 1 pptv. d Order of magnitude or range of local lifetimes. These are too short and variable to make a global lifetime meaningful. e Upper limit using a mean global Cl concentration of 1 103 cm3. f 3D model; Isaksen, private communication. Due to a strong anticorrelation induced by the oceanic emission of (CH3)2S and continental emission of NO2 removal by NO3 is less than that calculated from the global means (footnote a), (footnote c) which would assign 85% of the (CH3)2S loss to NO3. Adapted and expanded from Ehhalt, D.H., 1999a. Chapter 2: Gas phase chemistry of the troposphere. In: Zellner, R. (Guest ed.), Baumgärtel, H., Grünbein, W., Hensel, F. (eds.), Global Aspects of Atmospheric Chemistry, Topics in Physical Chemistry, vol. 6. Steinkopff, Darmstadt, pp. 21–109. a
b
trace gases (see Table 2), such that the knowledge of global OH allows a straightforward estimate of the mean global lifetimes of the individual trace gases. However, taken by itself, the OH concentration says nothing about the rates of the photochemical system and thus nothing about the absolute strength of the oxidizing capacity. A regional example illustrates the point: The OH concentrations in the lower and upper troposphere are about equal (Figure 2). Nevertheless, the loss rate of CH4 in the upper troposphere is more than a factor of 10 lower than in the lower troposphere, because of the lower temperature and lower CH4 concentration of the former. In addition, this measure provides little immediate insight into the parameters controlling the oxidizing capacity of the troposphere. Therefore, more recently, another more appropriate measure for the oxidizing capacity has been introduced. This consists of the total annual oxidative loss of trace gases in the troposphere, Lt (eqn [7], with Li defined by eqn [6]). X Lt ¼ Li [7] i
Restricting, as above, the losses considered to those faciliP tated by OH, the integrand for i Li becomes identical to DOH (eqn [4]), the local loss rate of OH, with one exception: the term k15 [HO2][OH] does not appear among the Li. This term, however, contributes only about 1% to the sum defining DOH, so that this difference can be safely neglected. Thus we obtain eqn [8], where LOH is the annual mean global loss rate of OH. Z 1 Lt y DOH dvdt ¼ LOH [8] a v;t
Using eqn [5b], we can rewrite eqn [8] as eqn [9]. Z Z 1 1 POH dvdt Lt yLOH ¼ POH ð1 þ rÞdvdt ¼ ð1 þ rÞ a a v;t
¼ ð1 þ rÞPOH;g [9]
The total oxidative loss rate of trace gases is given by the total loss rate of OH, which in turn is identical to the annual global primary production rate of OH, POH,g, multiplied by an appropriately averaged feedback factor ð1 þ rÞ. In other words, this measure for the oxidizing capacity is identical to the total global supply of OH. This definition has several advantages: 1. It quantifies the total trace gas loss initiated by OH in the troposphere. 2. It indicates directly the essential dependences of the oxidizing capacity on its controlling parameters. Since POH is essentially given by the photolysis of O3, i.e., POH J1[H2O][O3] (see eqn [2]), it is clear that POH,g is dominated by the tropospheric fields of O3, H2O, and solar UV radiation. All of these are influenced to some extent by human activities – for example, O3 by tropospheric chemistry acting on anthropogenic emissions, H2O by a warming climate, and solar UV by the loss of stratospheric O3. Clearly r is also influenced via tropospheric chemistry, but see point (4) below. 3. This measure also suggests that the oxidizing capacity of the troposphere depends via O3 and r on the emissions of NOx and volatile organic compounds and thus may respond with some resilience to the load placed on it by anthropogenic emissions, for example, by increasing the burden of tropospheric O3. 4. It directly indicates that the oxidizing capacity will always remain finite, since both, POH,g and r remain finite for any foreseeable emission scenario. For the present troposphere, POH,g and r can be estimated from 3D model calculations to about 108 Tmol per year and 0.74, respectively; the corresponding values for the preindustrial troposphere were 71.5 Tmol per year and 0.62. For very high loads of certain pollutants, r can become negligibly small. Thus, if the global emission flux of NOx was to exceed POH,g, the global OH abundance would be successively titrated away by reaction [XVII], while the NOx burden would build up until checked by other losses. It has been pointed out that such
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events of overwhelmed oxidation capacity can be observed regionally over areas with high NOx emissions. There, during wintertime when the flux of solar UV is low, the levels of oxidants such as O3 and H2O2 are greatly reduced in favor of a buildup of the primary pollutants, NOx and hydrocarbons, quite in contrast to summertime conditions. In such a case, the oxidizing capacity is limited to POH,g. Generally, however, emissions of NOx and oxidizable molecules have increased together causing an increase in r. The new measure, Lt, predicts a substantial increase of the tropospheric oxidizing capacity, about 60%, over the past 100 years, mainly because of the increase of global O3 by about 60%, but also due to an increasing r as a consequence of the global increase in NOx. The same conclusion can be reached directly from eqn [8], since the predicted global abundance of OH decreased by only 10%, whereas the predicted concentrations of the major molecules CH4 and CO increased by a factor of 2–3, greatly enhancing the total global loss rate of trace gases. The model-derived value of Lt for the oxidizing capacity of the present troposphere is 188 Tmol per year. This value includes the oxidation of secondary products such as HCHO. It is uncertain, because the actual emissions of the various tropospheric trace gases are uncertain and some of the minor trace gases may not be treated in the models at all. Moreover, the degradation pathways of some hydrocarbons are still not fully understood and may include OH recycling processes, not considered so far. Such seems to be the case for isoprene. Thus, the above value for Lt may rather represent a lower limit for the current oxidizing capacity of the troposphere.
See also: Ozone Depletion and Related Topics: Photochemistry of Ozone. Tropospheric Chemistry and Composition: Aliphatic Hydrocarbons; Aromatic Hydrocarbons; Hydroxyl Radical; Peroxyacetyl Nitrate; Sulfur Chemistry, Organic; Volatile Organic Compounds Overview: Anthropogenic.
Further Reading Ehhalt, D.H., 1999a. Chapter 2: Gas phase chemistry of the troposphere. In: Zellner, R. (Guest ed.), Baumgärtel, H., Grünbein, W., Hensel, F. (eds.), Global Aspects of Atmospheric Chemistry, Topics in Physical Chemistry, vol. 6. Steinkopff, Darmstadt, pp. 21–109. Ehhalt, D.H., 1999b. Photooxidation of trace gases in the troposphere. Physical Chemistry Chemical Physics 1, 5401–5408. Kleinmann, L.I., 1991. Seasonal dependence of boundary layer peroxide concentrations: the low and high NOx regimes. Journal of Geophysical Research 96, 20721–20733. Lelieveld, J., Dentener, F.J., Peters, W., Krol, M.C., 2004. On the role of hydroxyl radicals in the self-cleansing capacity of the troposphere. Atmospheric Chemistry and Physics 4, 2337–2344. SRef_ID: 1680-7324/acp/2004-4-2337. Lelieveld, J., Butler, T.M., Crowley, J.N., et al., 2008. Atmospheric oxidation capacity sustained by a tropical forest. Nature 452, 737–740. http://dx.doi.org/10.1038/ nature06870. Montzka, S.A., Krol, M., Dlugokencky, E., et al., 2011. Small interannual variability of global atmospheric hydroxyl. Science 331, 67–69. http://dx.doi.org/10.1126/ science.1197640. Prinn, R.G., 2003. The cleansing capacity of the atmosphere. Annual Review of Environmental Resource 28, 29–57. http://dx.doi.org/10.1146/annurev. energy.28.011503.163425.
Synopsis The oxidizing capacity of the troposphere relies on oxidants such as O3, NO3, and OH. Of these, OH is by far the most important. The supply of all the oxidants is limited and thus the oxidizing capacity is finite. The factors that control OH, O3, and NO3 are reviewed and examples of their tropospheric distributions are presented. The resulting lifetimes of the most important tropospheric trace gases are listed. The possible change of the oxidant distributions under the impact of anthropogenic emissions is indicated. A quantitative measure of the oxidation capacity is derived from the total loss rate of all trace gases due to OH.
Introduction The term ‘oxidizing capacity’ is widely used, but loosely defined. Its attraction is that it conveys in a succinct manner, two major features of atmospheric chemistry: namely, that the atmosphere actively oxidizes gaseous trace constituents and pollutants, and thereby initiates their removal from the atmosphere; and also, that this capacity is finite. The first feature is not unexpected; after all molecular oxygen, O2, is a major constituent of the Earth’s atmosphere. The second is somewhat surprising – considering that O2 is present at about 21%, whereas the most abundant oxidizable trace gases are present at ppm level. The explanation is, of course, that O2 does not directly react with those molecules. Tropospheric oxidation is rather initiated by a number of oxidants, above all by the hydroxyl radical, OH, but also by ozone, O3, the nitrate radical, NO3, and to a lesser extent by chlorine (Cl) and bromine (Br) atoms. Ground state atomic oxygen, O(3P), plays only a very minor role. In the liquid phase, for example in cloud droplets, hydrogen peroxide also acts as an oxidant. All these molecules and radicals are generated within the troposphere, and the rates of the respective generation processes limit their supply and thus the oxidizing capacity of the troposphere. In the following, we briefly review the factors that control the tropospheric concentrations of the major oxidants, OH, O3, and NO3. We present examples of their tropospheric distributions, and indicate how these may have changed under the impact of anthropogenic emissions. From the superposition of the various loss processes, we derive the atmospheric lifetimes for the most important tropospheric trace gases. Finally, based on all this information we present a measure of the tropospheric oxidizing capacity through an approximate but simple expression.
The Tropospheric Chemistries of OH and O3 The tropospheric chemistries of OH and O3 are very closely interlinked. In fact, the primary production of OH is based on the photolysis of O3. At wavelength below 340 nm, this photolysis yields electronically excited oxygen atoms, O(1D) (eqn [I]): O3 þ hn/Oð1 DÞ þ O2
l 340 nm
Encyclopedia of Atmospheric Sciences 2nd Edition, Volume 6
[I]
Most of the O(1D) atoms produced in reaction [I] are quenched to ground state atomic oxygen, O(3P), in collisions with atmospheric nitrogen and oxygen molecules (eqns [II] and [III]), but a fraction collides with water vapor molecules, to form the very reactive hydroxyl radical, OH (eqn [IV]). O(1D) þ N2 / O(3P) þ N2 O(1D) þ O2 / O(3P) þ O2 O(1D) þ H2O / OH þ OH
[II] [III] [IV]
H2O is ubiquitous in the troposphere. For instance, in surface air at midlatitudes, H2O is present at a volume mixing ratio of about 1%. There about 10% of the O(1D) generate OH, because reaction [IV] has a rate constant of approximately a factor of 10 higher than the quenching reactions [II] and [III]. The conversion of O3 to OH via the reactions [I] and [IV] at the same time forms a major sink of tropospheric O3, whereas the O(3P) formed by [II] and [III] returns to O3 by recombination with O2 (eqn [V]). Oð3 PÞ þ O2 þ M/O3 þ M ðM ¼ N2 ; O2 Þ
[V]
The same holds for the O(3P) formed directly in the photolysis of O3 at wavelengths longer than 320 nm (eqn [VI]). O3 þ hn/Oð3 PÞ þ O2
l > 340 nm
[VI]
The generation of O(3P) mainly by photolysis (eqn [VI]) but also by reactions [II] and [III] and others is balanced by the loss due to recombination. During daytime this results in a steady-state concentration of a few thousand O(3P) per cm3. Given the relatively low reactivity of O(3P), this concentration is too small to be of significance for the oxidation of trace gases in the troposphere. The fact that the O atoms generated in reaction [VI] interact rapidly with one of the major constituents of air limits their possible importance. Hydroxyl radicals, on the other hand, do not react with any of the major constituents of air. Rather, they have the ability to initiate chain reactions in an O2-containing atmosphere. When reacting with trace gas molecules, OH is not consumed, but is regenerated in catalytic cycles. In this way, relatively large OH concentrations, up to 107 cm3, are maintained in the sunlit troposphere despite the high reactivity of OH toward most pollutants and other trace gases. These two properties, high
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reactivity and relatively high concentrations, make OH the most important oxidizing agent in the troposphere.
OH Reaction with Molecules
whose photolysis leads to formation of HOx and thus introduces a degree of autocatalysis. Parallel to reaction [IX], HO2 reacts directly with O3 to return OH (eqn [XII]). HO2 þ O3 / OH þ 2O2
[XII]
The simplest example of such a reaction is given by the atmospheric oxidation of carbon monoxide. The reaction of CO with OH immediately forms the stable end product CO2 (eqn [VII]).
Thus even without NO there is a certain measure of recycling. But in this case the net result of the reaction chain, [VII], [VIII], and [XII], is a destruction of O3 (eqn [XIII]).
CO þ OH / CO2 þ H
CO þ O3 / CO2 þ O2
[VII]
[XIII]
The reaction also forms a hydrogen atom, which is very reactive. The hydrogen atom rapidly combines with O2 to form a hydroperoxy radical, HO2 (eqn [VIII]).
Together with reaction [XIV], reaction [XII] forms the second important loss mechanism of tropospheric O3, again closely tied to the HOx chemistry.
H þ O2 þ M / HO2 þ M
OH þ O3 / HO2 þ O2
[VIII]
The addition of a hydrogen atom to O2 weakens the bond between the oxygen atoms, and HO2 reacts much more readily than O2. In particular, HO2 oxidizes nitric oxide, NO (eqn [IX]). NO þ HO2 / NO2 þ OH
[IX]
In the planetary boundary layer over the industrialized and highly populated continents, where daytime concentrations of NO exceed 0.1 ppb, reaction [IX] is by far the fastest HO2 reaction. Most important for our present argument is the fact that the OH radical consumed in reaction [VII] is regenerated in reaction [IX]. This is generally true: virtually all reactions of OH with molecular species lead to chain reactions that recycle OH. Reaction [IX] also highlights one of the roles of the nitrogen oxides for the OH budget. In the form of NO they quickly return the less reactive HO2 back to the highly reactive OH, thus increasing the OH concentration. The nitrogen dioxide molecule, NO2, generated in reaction [IX] photolyzes readily in the near-UV region and therefore contributes to tropospheric photochemistry (eqn [X]). NO2 þ hy / NO þ O(3P)
[X]
In the sunlit atmosphere, the lifetime of NO2 against photolysis is a few minutes. The resulting O atom immediately combines with O2 to form O3 (reaction [V]). This NO-mediated process consisting of reactions [IX], [X], and [V], in which an O atom from a peroxy radical is passed along to O3, is the major ozone formation mechanism in the troposphere. It illustrates the fact that in an NO-containing atmosphere, trace gas removal is invariably linked with the production of O3 or other photo oxidants. Supplemented by reactions [X] and [V], the hydroxyl radical reactions [VII]–[IX] combine to the net reaction shown in eqn [XI]. CO þ 2O2 þ hy / CO2 þ O3
[XI]
It does not consume OH, HO2, NO, and NO2, hence the cycle consisting of the reactions [VII]–[X] and [V] can be completed repeatedly before that chain reaction is interrupted by termination reactions. By this and other chain reactions, OH and HO2 are interconverted within a matter of seconds. Therefore, both are often lumped together as HOx. The oxidation of more complex molecules, such as hydrocarbons, leads to other peroxy radicals, RO2, which also convert NO to NO2 and augment the generation of O3. In addition, intermediates such as formaldehyde (HCHO) are formed,
[XIV]
Destruction of HOx To produce a net loss of HOx and terminate the reaction chains, HOx radicals have to react with other radicals. Eqns [XV] and [XVI] are responsible for HOx loss in clean air. OH þ HO2 / H2O þ O2 HO2 þ HO2 / H2O2 þ O2
[XV] [XVI]
The addition of OH to NO2, forming a nitric acid molecule, HNO3 (eqn [XVII]), is the dominant HOx loss reaction in the polluted atmosphere. OH þ NO2 þ M / HNO3 þ M
[XVII]
At the same time, reaction [XVII] provides the major loss mechanism for NOx. It also illustrates the tendency of atmospheric oxidation to produce acidic end products, which are eventually removed from the atmosphere by rainout.
The OH Balance Equation The various production and destruction reactions of OH can be combined in a local OH budget equation of the type shown as eqn [1], where POH and DOH stand for the local production and destruction terms of OH. d ½OH ¼ POH DOH z0 dt
[1]
In the case of a very short-lived species like OH, contributions from transport, i.e., the flux divergence, are small compared to POH and DOH, and can be neglected. The same holds for the temporal change, d[OH]/dt. POH consists of two, systematically different, terms: the primary production of OH (POH) through the photolysis of precursors such as O3 or H2O2 and the OH produced from recycling HO2 (POH). By far the largest contribution to POH on a global scale comes from the photolysis of O3, which can be derived from reactions [I]–[IV] to give eqn [2]. POH ¼
2 ½O3 $ ½H2 O $ J1 $ k4 þ/ k2 ½N2 þ k3 ½O2 þ k4 ½H2 O
[2]
J1 is the photolysis frequency of reaction [I], and the ki refers to the rate constants of reactions [II]–[IV]. Smaller contributions,
such as that from the photolysis of H2O2, are indicated by the ellipsis ‘.’. POH is determined by reactions [IX] and [XII] as eqn [3]. [3]
DOH, the destruction of OH, is given by eqn [4]. DOH ¼ ½OH $ ðk7 ½CO þ k14 ½O3 þ / þ k17 ½NO2 þ k15 ½HO2 þ /Þ
OH P(O3)
2.0 1.5
4
1.0 2 0.5
[4]
¼ ½OH$s1 OH
0
A complete chemical mechanism would contain contributions to the production and destruction of OH other than those listed above. Again, these are indicated by the dotted ellipsis ‘.’. These additional terms would, however, fall in the same categories distinguished here. The first line in eqn [4] contains the reactions of OH with molecules, i.e., those that eventually generate HO2; the second line contains the reactions of OH with radicals that lead to a net HOx loss. The whole term in parenthesis in eqn [4] represents the pseudo-first-order reaction frequency or inverse lifetime, s1 OH , of an OH radical. We now define the recycling ratio r ¼ POH/POH. This is the number of HO2 radicals converted to OH divided by the number of OH generated directly. We note that the precursors for direct OH production, mainly O3 but also H2O2, are also produced by OH chemistry, although not necessarily at the same location, and, thus, contain an element of recycling. Because of their long lifetime their recycling is slow and incomplete and their local concentration is largely determined by transport. We further note that POH also includes a slow component, namely HO2 derived from aldehyde photolysis, besides the fast recycling of HO2 generated immediately after the OH attack. Nonetheless, for most of the troposphere, r also reasonably approximates the chain length, that is, the number of times an OH is recycled via HO2 before it is removed by a radical–radical reaction. Introducing sOH and r, eqn [1] takes the simple form of eqn [5a] or eqn [5b]. ½OH ¼ 0 sOH
[5a]
½OH ¼ POH $sOH ð1 þ rÞ
[5b]
POH ð1 þ rÞ
OH (106 cm–3)
POH ¼ ½HO2 $ðk9 ½NO þ k12 ½O3 Þ
2.5 6
or
Equation [5b] demonstrates nicely that the ability of OH for chain reaction and slow recycling enhances its concentration by a factor (1 þ r) over that of a nonchain-reacting radical, whose steady-state concentration would be given merely by the product of its production and lifetime, P $ s. Tropospheric values for r vary between 0.3 and 10 depending on NOx in a manner similar to that of OH shown in Figure 1. We note that the factors r and sOH still depend on the HOx concentration. Thus, eqn [5b] by itself is not sufficient for calculating the OH concentration. It helps, however, to categorize the action of the various reactions on OH. The nitrogen oxides, NOx ¼ NO þ NO2, for example, act on OH in two different ways. In the form of NO, they accelerate the recycling from HO2 to OH (reaction [IX]; cf eqn [3]) and thus enhance the concentration of OH; in the form of NO2 they enhance the net loss of OH via reaction [XVII], and thus decrease sOH and
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10.00
NOx (ppbv)
Figure 1 Dependence of the OH concentration and net O3 production on NOx calculated with a steady-state model for remote, rural conditions. Adapted from Ehhalt, D.H., 1999b. Photooxidation of trace gases in the troposphere. Physical Chemistry Chemical Physics 1, 5401–5408.
therefore the concentration of OH (cf eqn [4]). The opposing actions of the two processes lead to a highly nonlinear dependence of the OH concentration on NOx (see Figure 1). At low NOx, where recycling dominates, the addition of NOx leads to an increase in OH concentration and thus to an increase in photochemical activity; at high NOx, reaction [XVII] eventually becomes the dominant loss path of OH and further addition of NOx reduces the OH concentration. The position of the maximum with respect to NOx depends to some extent on the local mix of reactive trace gases. Except in polluted areas, however, tropospheric chemistry operates in NOx regimes defined by the left flank of the OH curve in Figure 1; that is, an increase in global NOx concentrations will lead to an increase in global OH. The dependence of OH on other parameters is more monotonic, though still not linear. For the NOx regime of the global troposphere, an increase in CO or other oxidizable molecules will lead to a decrease in OH (cf reaction [VII]; eqn [4]). But a change in H2O will lead to a change of equal sign in OH (cf eqn [2]). Solar radiation acts through several channels: through POH via the photolysis of O3, but also through r and sOH, because photolysis of NO2 shifts the [NO]/[NO2] ratio. These factors all conspire to give a nearly linear relationship between tropospheric OH and solar UV radiation. Figure 1 also shows the dependence on NOx of the net formation rate of O3 given by d ½O3 =dt ¼ PO3 DO3 , where the local O3 production, PO3 , is maintained by reaction cycles such as that consisting of [VII]–[X] and [V]. The O3 destruction, DO3 , is given by reactions such as [XII], [XIV], or [I] combined with [IV]. In the clean environment considered in Figure 1, the net O3 production rate is relatively small; even at the maximum it amounts to no more than 2.3 ppb h1. The profile of the net O3 formation rate is similar to that of the OH concentration, i.e., to the photochemical activity. At low NOx concentrations, however, it begins to deviate significantly, because there the O3-destroying reactions, such as reaction [I] combined with reaction [IV], remain active, whereas the O3-producing reactions, such as the sequence [VII]–[X], decrease with decreasing NOx. Eventually, below an NOx concentration of about 0.07 ppb, photochemistry leads to a net destruction of O3.
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Chemistry of NO3
NO3 þ NO2 # N2O5
The nitrate radical, NO3, is formed primarily by the relatively slow reaction [XVIII]. NO2 þ O3 / NO3 þ O2
[XVIII]
However, in daylight NO3 is photolyzed within tens of seconds via reactions [XIXa] and [XIXb]. NO3 þ hy / NO2 þ O(3P) NO3 þ hy / NO þ O2
[XIXa] [XIXb]
In addition, NO3 reacts rapidly with NO (eqn [XX]) that is always present during daytime owing to photolysis of NO2 (reaction [X]) such that during daytime NO3 remains at very low concentrations. NO3 þ NO / NO2 þ NO2
[XXI]
This reaction is important especially at low temperatures since the thermal decomposition of N2O5 to NO2 and NO3 is strongly temperature dependent. N2O5 reacts with condensed water to form HNO3. It thus reacts rapidly upon collision with the wet surfaces such as cloud droplets or aerosol particles at relative humidities above 60%. In nights with high relative humidities, this reaction path together with the direct uptake of NO3 provides an efficient loss of NOx. Finally, NO3 reacts with organic molecules in a manner similar to OH but for most molecules much more slowly than the corresponding reactions with OH. Nevertheless, in polluted regions it can provide a significant loss for organic molecules.
[XX]
During nighttime, however, NO3 can build up to significant concentrations because the remaining loss reactions are much slower. These are (1) the heterogeneous uptake of NO3 by moist aerosol surfaces, fog, or cloud droplets; and (2) the heterogeneous uptake of N2O5, which is formed from combination of NO3 with NO2 (eqn [XXI]).
Global Distributions of OH, O3, and NO3 The tropospheric chemistry outlined above maintains global distributions of OH, O3, and NO3 that vary with latitude, longitude, daytime, and season. For reference, Figure 2 presents the mean zonal distributions of OH, O3, and NO3 for January
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Latitude Figure 2 Mean zonal distributions of OH, O3, and NO3 for January and July based on 3D chemical transport models. OH distribution (top panels, contour lines 105 OH cm3) adapted from Spivakovsky, C.M., Logan, J.A., Montzka, S.A., et al., 2000. Three-dimensional climatological distribution of tropospheric OH: update and evaluation. J. Geophys. Res. 105, 8931–8980. O3 distribution (middle panels, contour lines ppbv) adapted from Wang, Y., Jacob, D.J., Logan, J.A., 1998. Global simulation of tropospheric O3-NOx-hydrocarbon chemistry. Journal of Geophysical Research 103, 10757–10767. NO3 distribution (bottom panels, contour lines pptv), private communication Hauglustaine, D., 2000.
Tropospheric Chemistry and Composition j Oxidizing Capacity and July. They are based on 3D chemical transport models, since measurements are too sparse to allow a representative and consistent reconstruction of the average concentration fields. Nevertheless, the concentration fields shown are expected to closely approach reality. For O3, the measured seasonal variations, as well as height profiles, show good agreement with those modeled for the same sites. For OH, the globally averaged concentration agrees well with the value of 1.0 106 cm3 derived empirically from the budget of methyl chloroform, CH3CCl3. The main oxidant, OH, shows a strong latitudinal variation at all altitudes, with a broad maximum in the tropics centered at the latitude of maximum solar radiation. It also varies with altitude with a broad maximum around 4 km at all latitudes. In addition, OH, as well as the other oxidants, O3 and NO3, exhibits higher concentrations in the Northern Hemisphere whose continents provide most of the global emissions, whether natural or artificial. For the same reason, the longitude by latitude presentation (Figure 3) also indicates higher concentrations of OH over the continents. The chemistry outlined above also predicts that the increase of the anthropogenic emissions of NOx, CO, and CH4 that accompanied the industrialization and population growth of the last century must have caused significant changes in the concentrations of the tropospheric oxidants. There is indeed
experimental evidence mainly from historical records over Europe that the concentration of O3 has increased by about a factor of two over that period. Figure 4 presents the modelcalculated change in the zonally and annually averaged distribution of O3 and OH from preindustrial to present time. For O3 the models predict a general increase of about 60%, which is, however, lower in the Southern Hemisphere (þ50% in the case presented) than in the Northern Hemisphere (þ80%). The model-calculated globally averaged OH decreased by about 10%, a relatively small value. However, as Figure 4 indicates, the changes are quite unevenly distributed, with large zones of substantial positive change. The OH concentration over the continents increased strongly, whereas it decreased over the oceans, mapping the nonuniform distribution of NOx, which is short lived and mostly emitted over the continents. In fact, the global distribution of OH, as well as its trend, strongly depends on the assumed spatial and temporal distribution of the NOx emissions. The most recent analysis of the CH3CCl3 budget suggests a relatively constant mean global OH concentration over the past 20 years with a small interannual variability of 2.1 1.8% per year, but no significant long-term trend. The difference between the global change for OH and O3 is not unexpected. For O3 the increased emissions of CO and NOx reinforce each other, as more CO leads to higher formation of
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Figure 3 Global distribution of OH for January (top panel) and July (bottom panel) at 700 mb (contour lines 105 OH cm3). Adapted from Spivakovsky, C.M., Logan, J.A., Montzka, S.A., et al., 2000. Threedimensional climatological distribution of tropospheric OH: update and evaluation. Journal of Geophysical Research 105, 8931–8980.
1000 –90
–60
–30
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Figure 4 Calculated relative change in the zonally and annually averaged distribution of O3 (top panel) and OH (bottom panel) from preindustrial to present times (contour lines in %). Adapted from Wang, Y., Jacob, D.J., 1998. Anthropogenic forcing on tropospheric ozone and OH since preindustrial times. Journal of Geophysical Research 103, 31123–31135.
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HO2, which in turn is efficiently converted to O3 by the enhanced NO. For OH they counteract each other – increased NOx induces a larger and increased CO a lower OH concentration. For NO3 one would also expect a global increase in the concentration field since preindustrial times, since both the NO2 and O3 concentrations have increased globally. However, no hindcasts have been published for NO3. Obviously, as long as the emissions of NOx, CO, and CH4 continue to grow, further changes in the global distributions of OH, O3, and NO3 have to be expected. For example, for the future emission scenario A1F1 from IPCC 2001, which presents one of the worst cases and assumes global emissions of NOx, CO, VOC, and CH4 of 243, 193, 198, and 128%, respectively, of the values in the year 2000, model calculations predict a decrease in the mean global OH of 14% and an increase in O3 of 47% by the year 2100.
The Global Loss Rate The removal of a given trace gas from the troposphere is usually generated by several different processes. In addition to the tropospheric oxidation by OH, O3, NO3, and other radicals, direct photolysis, dry deposition at the Earth’s surface, and export to the stratosphere may play a role. For a gas Xi, the global tropospheric loss rate, Li, is given by the sum of all the individual loss rates integrated over the volume of the troposphere and averaged over a full year. Correspondingly, the global tropospheric lifetime of a trace gas, sI, is defined as its tropospheric burden, Mi, divided by its global tropospheric loss rate. The contribution of the major oxidants to the global loss rate of Xi is given by eqn [6], where the brackets denote the local concentrations in units of cm3, the ki are the respective rate constants, which generally depend on temperature and thus also on location and season, v stands for tropospheric volume, t for time, and a ¼ 1 year. Li ¼
1 a
Z
ki;OH $½OH þ ki;O3 ½O3 þ ki;NO3 $½NO3 $½Xi dv$dt
v;t
¼ Mi s1 i [6]
Examples of the rate constants are given in Table 1 for a few selected organic trace molecules. Clearly, the rate constants vary over many orders of magnitude, reactions with OH having the highest and those with O3 the lowest rate constants. We also note that the ratios kOH : kO3 : kNO3 vary substantially between the various trace gases. Thus the relative contributions of the global OH, O3, and NO3 fields to the global oxidation rate vary greatly with the trace gas considered. Generally, however, tropospheric oxidation is dominated by reaction with OH. This is emphasized by Table 2, which lists the trace gases with the highest turnover in the atmosphere. The most important three trace gases, namely CO, CH4, and H2, are not attacked by O3 or NO3. In fact, significant direct oxidation by O3 is limited to the terpenes and NOx; and that by NO3 to the terpenes and dimethyl sulfide. In terms of global tropospheric removal, direct attack of O3 and of NO3 adds up to a loss rate of 3.75 Tmol per year, i.e., about 2% of the total loss rate of 180 Tmol per year given by the sum over the individual trace gas losses in Table 2 and the oxidation of their secondary products. This does not preclude that oxidation by NO3 and O3 becomes important on a local scale, or for certain molecules even on a global scale, but its impact on the total turnover of oxidizable molecules is small. Similarly, the tropospheric role of Cl and Br atoms appears small. Based on the global budgets of ethane and tetrachloroethene, the average Cl atom concentration in the Northern Hemisphere has been estimated to less than 1000 cm3 and in the Southern Hemisphere less than 2000 cm3. This is three orders of magnitude smaller than the globally averaged OH concentrations. This difference is so large that it cannot be compensated by the generally higher rate constants of hydrocarbons with Cl atoms. Finally, the role of H2O2 is pretty much limited to the oxidation of SO2 in the liquid phase (Table 2) and its overall impact is on the order of 1% of the total loss rate of 180 Tmol per year. As a consequence, the tropospheric oxidizing capacity is generally viewed as synonymous with the OH chemistry in the troposphere. Moreover, at least implicitly, the mean global OH abundance has been, and occasionally still is, used as a quantitative measure for the oxidizing capacity. Note that according to this measure, the oxidizing capacity is predicted not to have changed much over the last 100 years. At first glance this choice appears plausible, because OH determines the loss rates of many
Table 1 Rate constants of reactions of various organic trace molecules with OH, O3, and NO3. The rate constants are given for 298 K in cm3 molecule1 s1. The respective lifetimes are calculated on the basis of following concentrations: OH: 1 106 cm3, O3: 30 ppbv, NO3: 1 pptv Rate constants (at 298 K)
Organic kOH n-Butane Ethene Propene Trans-2-butene Isoprene a-Pinene Formaldehyde CH3CHO (CH3)2S
2.4 8.6 2.9 6.4 1.0 5.3 8.5 1.5 5.0
1012 1012 1011 1011 1010 1011 1012 1011 1012
Lifetimes
kO 3
kNO 3
<1022 1.6 1018 1.0 1017 1.9 1016 1.3 1017 8.7 1017 – – <1018
4.6 2.1 9.5 3.9 7.0 6.2 5.6 2.7 1.1
1017 1016 1015 1013 1013 1012 1016 1015 1012
sOH
sO 3
sNO 3
4.8 days 32 h 9.6 h 4.3 h 2.8 h 5.2 h 33 h 19 h 56 h
– 9.6 days 37 h 1.9 h 28 h 4.1 h – – >15 days
28 years 6 years 49 days 28 h 16 h 1.8 h 2.3 years 170 days 10 h
Atkinson, R., 1997. Gas-phase tropospheric chemistry of volatile organic compounds: 1. Alkanes and alkenes. J. Phys. Chem. Ref. Data 26, 215–290; Atkinson, R., Baulch, D.L., Cox, R.A., et al., 2006. Evaluated kinetic and photochemical data for atmospheric chemistry: volume II – gas phase reactions of organic species. Atmospheric Chemistry and Physics 6, 3625–4055.
Tropospheric Chemistry and Composition j Oxidizing Capacity Table 2
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Global turnover of the major tropospheric trace gases and attribution to the major oxidants
Trace gas
Global loss rate (Tmol per year)
Global lifetime
Removal (%) OH
CO H2 CH4 Isoprene SO2 NOx Terpenes C2H6 N2O (CH3)2S
1.5 months 2 years 8 years Hoursd Daysd 0.3–5 daysd Hoursd 2 months 120 years Daysd
100 38 36 8 5 3 1 0.7 0.6 0.5
85 25 90 80 30 50 20 80 – 70f
a
O3b
NO3c
Other
– – – 7 – 40 25 – – –
– – – 13 – – 55 – – 30f
15 (Soil uptake) 75 (Soil uptake) 10 (Soil uptake; stratos.) – 70 (Hetero. in clouds) 10 (Soil uptake) – 20 (Cl reaction)e 100 (Stratosphere) –
Using a mean global OH concentration of 1 106 cm3. Using a mean global O3 concentration of 30 ppbv. c Using a mean global NO3 concentration of 1 pptv. d Order of magnitude or range of local lifetimes. These are too short and variable to make a global lifetime meaningful. e Upper limit using a mean global Cl concentration of 1 103 cm3. f 3D model; Isaksen, private communication. Due to a strong anticorrelation induced by the oceanic emission of (CH3)2S and continental emission of NO2 removal by NO3 is less than that calculated from the global means (footnote a), (footnote c) which would assign 85% of the (CH3)2S loss to NO3. Adapted and expanded from Ehhalt, D.H., 1999a. Chapter 2: Gas phase chemistry of the troposphere. In: Zellner, R. (Guest ed.), Baumgärtel, H., Grünbein, W., Hensel, F. (eds.), Global Aspects of Atmospheric Chemistry, Topics in Physical Chemistry, vol. 6. Steinkopff, Darmstadt, pp. 21–109. a
b
trace gases (see Table 2), such that the knowledge of global OH allows a straightforward estimate of the mean global lifetimes of the individual trace gases. However, taken by itself, the OH concentration says nothing about the rates of the photochemical system and thus nothing about the absolute strength of the oxidizing capacity. A regional example illustrates the point: The OH concentrations in the lower and upper troposphere are about equal (Figure 2). Nevertheless, the loss rate of CH4 in the upper troposphere is more than a factor of 10 lower than in the lower troposphere, because of the lower temperature and lower CH4 concentration of the former. In addition, this measure provides little immediate insight into the parameters controlling the oxidizing capacity of the troposphere. Therefore, more recently, another more appropriate measure for the oxidizing capacity has been introduced. This consists of the total annual oxidative loss of trace gases in the troposphere, Lt (eqn [7], with Li defined by eqn [6]). X Lt ¼ Li [7] i
Restricting, as above, the losses considered to those faciliP tated by OH, the integrand for i Li becomes identical to DOH (eqn [4]), the local loss rate of OH, with one exception: the term k15 [HO2][OH] does not appear among the Li. This term, however, contributes only about 1% to the sum defining DOH, so that this difference can be safely neglected. Thus we obtain eqn [8], where LOH is the annual mean global loss rate of OH. Z 1 Lt y DOH dvdt ¼ LOH [8] a v;t
Using eqn [5b], we can rewrite eqn [8] as eqn [9]. Z Z 1 1 POH dvdt Lt yLOH ¼ POH ð1 þ rÞdvdt ¼ ð1 þ rÞ a a v;t
¼ ð1 þ rÞPOH;g [9]
The total oxidative loss rate of trace gases is given by the total loss rate of OH, which in turn is identical to the annual global primary production rate of OH, POH,g, multiplied by an appropriately averaged feedback factor ð1 þ rÞ. In other words, this measure for the oxidizing capacity is identical to the total global supply of OH. This definition has several advantages: 1. It quantifies the total trace gas loss initiated by OH in the troposphere. 2. It indicates directly the essential dependences of the oxidizing capacity on its controlling parameters. Since POH is essentially given by the photolysis of O3, i.e., POH J1[H2O][O3] (see eqn [2]), it is clear that POH,g is dominated by the tropospheric fields of O3, H2O, and solar UV radiation. All of these are influenced to some extent by human activities – for example, O3 by tropospheric chemistry acting on anthropogenic emissions, H2O by a warming climate, and solar UV by the loss of stratospheric O3. Clearly r is also influenced via tropospheric chemistry, but see point (4) below. 3. This measure also suggests that the oxidizing capacity of the troposphere depends via O3 and r on the emissions of NOx and volatile organic compounds and thus may respond with some resilience to the load placed on it by anthropogenic emissions, for example, by increasing the burden of tropospheric O3. 4. It directly indicates that the oxidizing capacity will always remain finite, since both, POH,g and r remain finite for any foreseeable emission scenario. For the present troposphere, POH,g and r can be estimated from 3D model calculations to about 108 Tmol per year and 0.74, respectively; the corresponding values for the preindustrial troposphere were 71.5 Tmol per year and 0.62. For very high loads of certain pollutants, r can become negligibly small. Thus, if the global emission flux of NOx was to exceed POH,g, the global OH abundance would be successively titrated away by reaction [XVII], while the NOx burden would build up until checked by other losses. It has been pointed out that such
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events of overwhelmed oxidation capacity can be observed regionally over areas with high NOx emissions. There, during wintertime when the flux of solar UV is low, the levels of oxidants such as O3 and H2O2 are greatly reduced in favor of a buildup of the primary pollutants, NOx and hydrocarbons, quite in contrast to summertime conditions. In such a case, the oxidizing capacity is limited to POH,g. Generally, however, emissions of NOx and oxidizable molecules have increased together causing an increase in r. The new measure, Lt, predicts a substantial increase of the tropospheric oxidizing capacity, about 60%, over the past 100 years, mainly because of the increase of global O3 by about 60%, but also due to an increasing r as a consequence of the global increase in NOx. The same conclusion can be reached directly from eqn [8], since the predicted global abundance of OH decreased by only 10%, whereas the predicted concentrations of the major molecules CH4 and CO increased by a factor of 2–3, greatly enhancing the total global loss rate of trace gases. The model-derived value of Lt for the oxidizing capacity of the present troposphere is 188 Tmol per year. This value includes the oxidation of secondary products such as HCHO. It is uncertain, because the actual emissions of the various tropospheric trace gases are uncertain and some of the minor trace gases may not be treated in the models at all. Moreover, the degradation pathways of some hydrocarbons are still not fully understood and may include OH recycling processes, not considered so far. Such seems to be the case for isoprene. Thus, the above value for Lt may rather represent a lower limit for the current oxidizing capacity of the troposphere.
See also: Ozone Depletion and Related Topics: Photochemistry of Ozone. Tropospheric Chemistry and Composition: Aliphatic Hydrocarbons; Aromatic Hydrocarbons; Hydroxyl Radical; Peroxyacetyl Nitrate; Sulfur Chemistry, Organic; Volatile Organic Compounds Overview: Anthropogenic.
Further Reading Ehhalt, D.H., 1999a. Chapter 2: Gas phase chemistry of the troposphere. In: Zellner, R. (Guest ed.), Baumgärtel, H., Grünbein, W., Hensel, F. (eds.), Global Aspects of Atmospheric Chemistry, Topics in Physical Chemistry, vol. 6. Steinkopff, Darmstadt, pp. 21–109. Ehhalt, D.H., 1999b. Photooxidation of trace gases in the troposphere. Physical Chemistry Chemical Physics 1, 5401–5408. Kleinmann, L.I., 1991. Seasonal dependence of boundary layer peroxide concentrations: the low and high NOx regimes. Journal of Geophysical Research 96, 20721–20733. Lelieveld, J., Dentener, F.J., Peters, W., Krol, M.C., 2004. On the role of hydroxyl radicals in the self-cleansing capacity of the troposphere. Atmospheric Chemistry and Physics 4, 2337–2344. SRef_ID: 1680-7324/acp/2004-4-2337. Lelieveld, J., Butler, T.M., Crowley, J.N., et al., 2008. Atmospheric oxidation capacity sustained by a tropical forest. Nature 452, 737–740. http://dx.doi.org/10.1038/ nature06870. Montzka, S.A., Krol, M., Dlugokencky, E., et al., 2011. Small interannual variability of global atmospheric hydroxyl. Science 331, 67–69. http://dx.doi.org/10.1126/ science.1197640. Prinn, R.G., 2003. The cleansing capacity of the atmosphere. Annual Review of Environmental Resource 28, 29–57. http://dx.doi.org/10.1146/annurev. energy.28.011503.163425.