CIESIN Reproduced, with permission, from: Madronich, S. 1993. Tropospheric photochemistry and its response to UV changes. In The role of the stratosphere in global change. Vol. 18. NATO-ASI Series, ed. M-L. Chanin, 437-61. Amsterdam: Springer-Verlag.


Sasha Madronich, National Center for Atmospheric Research Boulder, Colorado USA

ABSTRACT. The photochemistry of the troposphere is highly non-linear, and may be changing due to emissions of gases related to human activities. Increases in tropospheric ultraviolet (UV) radiation, due to stratospheric ozone depletion, may also perturb the troposphere.


Numerous environmental concerns arise from tropospheric chemistry. On the urban scale (ca. 50 x 50 km[2]), intense tropospheric pollution can have direct effects on human health through noxious gases such as carbon monoxide (CO), nitrogen oxides (NOX = NO + NO2, with NO = nitric oxide and NO2 = nitrogen dioxide), organic compounds including methane (CH4) and non-methane hydrocarbons (NMHC) and their partly oxidized intermediates such as peroxy acetyl nitrate (CH3-CO-OONO2, or PAN), and airborne particulates. This pollution is not confined to urban areas, however, and can spread over regional scales (ca. 500 x 500 km[2]), resulting in acid rain through the formation of compounds like nitric acid (HNO3), sulfuric acid (H2SO4) and organic acids (R-CO-OH, where R is an organic radical group). Vegetation damage can result also from exposure to high levels of oxidizing gases such as O3 and hydrogen peroxide (H2O2).

The concern that air pollution may grow to the global scale (4pi 6371[2] km[2]) is real. There is good evidence that many tropospheric gases, active either chemically or radiatively, have grown in abundance since the pre-industrial era. Among the chemically reactive species, long term increases in CH4 and tropospheric O3 are well known (IPCC, 1990). For NOX, current athropogenic emissions are thought to be comparable to or larger than global natural sources (IPCC, 1990; WMO, 1991). The first major global concern is that an increase in these gases, some of which are strong infrared absorbers, may change the radiative balance of the atmosphere, possibly resulting in climate change. The combined radiative effect of all of these gases is about double that of carbon dioxide (CO2) alone. Additional climate perturbations may come from aerosols, formed when some compounds such as sulfur dioxide (SO2) and dimethyl sulfide (CH3-S-CH3, or DMS) react in the troposphere to form H2SO4, which can lead to the formation of sulfate aerosols. These aerosols can absorb and scatter light, and, if appropriate in size, can also function as cloud condensation nuclei (CCN), changing the size distribution of cloud droplets and affecting cloud reflectivities.

The second major global issue is the tropospheric "oxidizing" or "self-cleaning" capacity. Many compounds in the atmosphere are removed by reaction with the hydroxyl radical (OH) (Levy, 1971; Thompson, 1992 and refs. therein). OH is not available in infinite supply, and can be lost or at least reduced, if emissions associated with human activities exceed certain values. Should this happen, the lifetime of many compounds will increase, and ultimately their atmospheric concentrations will rise to higher levels. Tropospheric OH is made mostly from tropospheric O3, and both are sensitive to ultraviolet (UV) radiation, and therefore to the overhead ozone abundance. Thus, one consequence of stratospheric ozone depletion is the alteration of tropospheric chemistry. Ultimately, the chemistries of the troposphere and the stratosphere must be viewed as fully coupled, because the troposphere is also the first chemical filter for surface-emitted compounds on their way to the stratosphere.

Tropospheric chemistry is non-linear, it involves a large number of compounds emitted at the surface, and is complicated by interactions between different phases including gas, liquid, aerosol, and various surfaces. Large uncertainties remain that go beyond simple computing resource limitations. Among all the complications, a few of the basic features do appear to be understood, particularly with regards to the roles of OH, O3, CH4, and NOx, and are discussed below.


Ozone is made in the troposphere by the same reaction responsible for its formation in the stratosphere: the addition of ground state oxygen atoms O([3]P) to molecular oxygen O2, assisted by any third body M to ensure simultaneous momentum and energy conservation,


However, the source of the O atoms is different than in the stratosphere where the O atoms are made by the photodissociation of O2 at UV wavelengths lamba < 240 nm. In the troposphere, only UV radiation with lamba > 290 nm is available, due to essentially complete absorption of shorter wavelengths by O2 and O3 above the tropopause. Thus, an alternate source of O atoms is necessary to explain tropospheric ozone formation. Exhaustive search of plausible tropospheric reactions reveals only a few possibilities which yield O atoms, and among these only one appears to be important: the photodissociation of NO2,


This reaction however cannot be the complete story. While the O([3]P) produced by NO2 photolysis does indeed form ozone, the NO produced in the same reaction reacts rapidly with ozone,


Thus, the net ozone production is strongly limited. This is illustrated in Figure 1. Starting with an initial NO2 amount [NO2]o and zero nitric oxide and ozone, the ozone concentration at time t is


and is therefore always smaller than the initial NO2. In practice, however, observations show nearly always [O3] >> [NOx], so that an alternate way of producing the ozone is indicated. The alternative must still use NO2 photolysis as the primary source of O3, but the NO2 must be created without the loss of ozone that occurred in reaction 3. The necessary reactions are supplied by the oxidation of CO and various hydrocarbons.

The simple case of methane oxidation is illustrated in Table 1. Reactions 1-3 form ozone as discussed above. The real starting point is reaction 4, the UV photolysis of O3 leading to the formation of excited oxygen atoms, O([1]D). If these atoms are quenched (i .e., collisionally transfer their excitation and return to the ground state), as occurs with about 95% probability, ozone is simply generated through reaction 1 and no net change occurs. With some smaller probability, the O([1]D) atoms can encounter a water molecule (H2O) and produce two OH radicals. The OH radicals then initiate a sequence of reactions which oxidize methane first to formaldehyde (CH2O) (reactions 7-11), then to CO and ultimately CO2 (reactions 12-15). In the presence of NOx, the complete oxidation of CH4 to CO2 can lead to a total of five NO-->NO2 conversions, and also regenerates 2 OH radicals necessary to initiate the oxidation of other CH4 molecules. Each NO-->NO2 conversion which occurs without loss of ozone is effectively a net ozone production step.

The methane oxidation illustrated above is clearly a powerful agent for NO-->NO2 conversions, and may be expected to lead to extremely large tropospheric ozone amounts. However, other chemical reactions are in competition with the reactions listed in Table 1. Some of these may destroy ozone directly, i.e., the HOx (= OH + HO2) catalytic cycle,


Another important ozone loss is reaction 6, since it represents the net loss of ozone dissociated in reaction 4. Also, the effectiveness of the oxidation of CO and CH4 can be greatly reduced by termination reactions in which two radicals are lost,


and by reactions in which both radicals and NOx are lost,


The products H2O2, CH3O2H, and HNO3 should be viewed as reservoirs, rather than terminal products, since they can lead to radical (or NOx) regeneration, e.g.,


if they are not removed from the atmosphere by rainout or surface deposition.

The effects of CO and hydrocarbon oxidation on ozone production may be summarized as


where R-OO is intended to represent organic peroxy radicals (roughly analogous to CH3O2) derived from non-methane hydrocarbons. In all cases, the ozone production is dependent on the NO levels. Surprisingly small values of NO can offset the ozone losses. Considering for simplicity only reactions 11 and 16, ozone production is greater than loss when


which is about 5 ppt (1ppt = 10[-12] molar ratio = 2.69x10[7] molec cm[-3] at STP) given the values of k16 and k11, and typical tropospheric ozone values. The full dependence of ozone production and loss on NOx levels is shown in Figure 2. The importance of NOx is clear, and will be discussed further in the next section.


Tropospheric nitrogen oxides originate primarily from the heating of air to temperatures where the Zeldovich mechanism becomes operative,


These temperatures are reached during most combustion processes (including fossil fuel and biomass burning) and lightning. Additional NOx sources may be associated with bacterial processes in soils.

Once in the atmosphere, NO and NO2 partake in many chemical reactions. Some of these are simple NO - NO2 interconversions, while others are actual NOx sinks. In particular, the reaction


removes NOx quickly, with about 1 day lifetime for typical mid-latitude conditions. The short NOx lifetime has one important implication: If the sources of NOx are not geographically uniform (which is certainly the case), the global NOx distributions will be highly variable, being very sensitive to both chemical and meteorological (transport) processes. Typical NOx values observed in different non-urban regions of the northern hemisphere are shown in Figure 3. NOx levels are seen to span about 3 orders of magnitude, and can be on either side of the ozone net -production threshold shown in Figure 2.

While most of the reactive nitrogen is emitted as NO or NO2, field data suggest that atmospheric reactive nitrogen may be present also in different forms, such as HNO3, PAN, particulate nitrate (NO3-), and organic nitrates, among others. The ratio of NOx to the total reactive nitrogen (NOy, measured by instruments which are sensitive to the total reactive nitrogen rather than to individual components), is shown in Figure 4. The general trend is that in relatively polluted regions, NOx is a high fraction of NOy, but decreases toward more remote regions. This is consistent with the NOx short lifetime and with conversions to other nitrogen species during several days of atmospheric transport. However, the budget of the NOy species is still incomplete. This is shown in Figure 5, which compares simultaneous but independent measurements of total reactive nitrogen (NOy) with the individually measured species NO, NO2, HNO3, PAN, and NO3-. Clearly there is a shortfall in the NOy which cannot be explained by any measured species. The magnitude of the shortfall can be greater than the NOx amounts, a disturbing fact in view of the potential role of even small amounts of NOx in global ozone production. The identity of the "mystery" NOy species is still uncertain.


The simplest hydrocarbon, CH4, is believed to be of considerable importance on the global scale. But on urban and regional scales, it is really the non-methane hydrocarbons (see Table 2) that dominate the chemistry of O3 and NOx. Their rate constants for reaction with OH are several orders of magnitude faster than that for the OH + CH4 reaction, and their emission rates are at least as large as that of CH4 on the global scale, and much larger on some local scales (WMO 1991). The chemistry of the NMHC proceeds along the general lines described for methane in the previous sections, but is more complex due to the much larger number of intermediate, reservoir, species that can be formed. A few of these intermediates are shown in Table 3.

The NMHCs confound tropospheric chemistry in several ways. First, near their sources, their strong reactivity may have an immediate effect on local O3, OH, and NOx levels. Second, because of their short lifetimes, the geographical distribution of the NMHCs will be highly variable and determined by complex couplings between chemistry and transport. Third, and perhaps most important for global considerations, the NMHCs are not converted immediately to their ultimate oxidation products CO2 and H2O; instead, intermediate compounds are produced after the initial OH attack, and may persist much longer in the atmosphere than the initial NMHC, possibly being transported over global distances. It has been suggested, for example, that these intermediate compounds may tie up at least some of the missing NOy species discussed in the previous section. These organic compounds, which may collectively be called Cy by analogy with NOy, may pervade the troposphere but, as of today, remain largely undetected.


The reaction with OH radicals is one of the main pathways for the removal of many compounds from the troposphere. Changes in the global amount of OH may have an important impact on the concentration of these compounds. In turn, rising levels of pollutants may change the concentrations of OH, thus affecting the self-cleaning (oxidizing) capacity of the troposphere.

Consider, specifically, a compound X whose atmospheric burden is determined by the balance between its emission fluxes (Fx), and its sinks which are inversely proportional to its atmospheric lifetime taux. Its atmospheric concentration, [X], can rise for two reasons: (a) in response to increased fluxes, and (b) in response to an increase in taux, which for many gases is given by the reaction with OH radicals. If the OH concentration, [OH], is in turn dependent on [X], a potentially powerful chemical feedback is possible: increases in Fx may increase the atmospheric concentrations of X, which in turn decrease [OH] and increase the lifetime of X, leading to further increases in [X]. The response of [OH] to current and future emissions is thus a major issue in tropospheric chemistry.

Current OH levels are illustrated in Figure 6, together with some other species of interest to the oxidizing capacity. OH is produced primarily from photolysis reactions, and is therefore nearly negligible at night, while noon-time values reach about 4x10[6] molec cm[-3] for the conditions shown in the Figure. Table 4 gives a detailed breakdown of the chemical reactions which contribute to the formation and destruction of OH in the remote free troposphere. It is seen that (for the 24-hr average), primary photolytic production accounts for about 60% of the OH radical production, while conversion of HO2 accounts for the remainder. OH loss is dominated by reaction with CO (49%) and CH4 (18%), with another 12% arising from intermediates formed by methane oxidation (CH3OOH, CH2O). The dependence of OH on NOx levels is fairly small for the low-NOx conditions shown in Table 4 (5-20 ppt NOx), but increases in more polluted regions. This is illustrated in Figure 7, which shows both OH and HO2 radicals as a function of NOx concentrations. If NOx is increased, OH is seen to first increases due to the NO + HO2 --> NO2 + OH reaction, but ultimately falls due to the radical termination reaction OH+NO2+M-->HNO3+M.

The possibility of serious OH changes due to anthropogenic emissions is illustrated in Figure 8. Here, reasonably high steady state OH concentrations (~10[6] molec cm[-3]) are found for low CO and NOx fluxes. But as FCO and FNO are increased beyond certain threshold values, [OH] can decrease catastrophically, although some specific combinations of high NO and CO inputs can still result in acceptable oxidizing capacity.

The chemical system is clearly complex, and may even exhibit some multiple solutions, that is, different possible atmospheric composition states for identical input conditions. This is illustrated in Figure 9, where the high-NO collapse of Figure 8 is shown in cross sectional detail. Starting at PO, increases in FNO lead first to more OH (to P1), but subsequent increases in FNO lead to much lower OH levels (to P2). Suppose now that, after the environmental seriousness of the OH collapse is realized, attempts are made to restore oxidizing capacity by reducing the FNO Reversing the FNO trend will not result in the same OH solution until the point P3 is reached, where FNO is considerably lower than the FNO which was capable of sustaining the highest OH levels at P1.

A most important question relates to how today's fluxes compare with the threshold fluxes at which the oxidizing capacity fails. Rough estimates suggest that we are within an order of magnitude of the critical values, but such estimates should be viewed in the context of large uncertainties in both the estimation of fluxes and in the model calculations of the global oxidizing capacity.


Tropospheric UV levels are kept well below the extraterrestrial values by the filtering action of stratospheric ozone. All other factors being taken as constant, stratospheric ozone depletion leads to a direct increase in tropospheric UV radiation, and therefore in the photolysis rates for various tropospheric photo-active species. Some of the UV-sensitive tropospheric chemical processes include reactions 2, 4, 12, 23, and 24, among others.

The lifetime of a molecules X with respect to photodissociation is given by the inverse of the photolysis rate coefficient (J, also termed the photodissociation frequency),


where lamba is wavelength (nm), F(lamba) is the spectral actinic flux (quanta cm[-2] s[-1] nm[-1]), sigmaX(lamba) is the absorption cross section (cm[2] molec[-1]) and phiXi(lamba) is the quantum yield (molec quantum[-1]) for the dissociation of X into one of several possible product sets, say channel i. Changes in photolysis rates are likely to have a profound impact on various aspect of tropospheric composition, including the formation and destruction of O3 and OH radicals, and the lifetime of various reduced species.

The calculated variation of tropospheric O3 with changing stratospheric O3 levels (and therefore UV levels) is shown in Figure 10. The net effect on ozone production depends, once again, on NOx levels. At low NOx, the UV increase reduces surface O3 levels, while at high NOx the net effect is an increase in O3. The high-NOx regime is of particular interest in cities where photochemical smog and ozone formation are of concern, because higher UV levels will make it more difficult to attain air quality standards, as illustrated in Table 5. The low-NOx ozone decrease may already have been observed, as shown in Figure 11, under the extreme conditions of UV increase experienced under the Antarctic ozone hole.

OH concentrations are generally expected to increase with increasing UV levels, as shown in Figure 12. The calculated relative increases in [OH] are greatest at low NOx, (about 50% increase for a 20% ozone column reduction), with decreasing sensitivity at high NOx. An analysis of recent total ozone column data from 1979 to 1989 (TOMS version 6, as reported by Stolarski et al. 1991 ) suggests an increase in the J value for the reaction


which is the primary OH source over most of the troposphere. The globally averaged J trend is about 4% per decade, with the latitudinal distribution shown in Table 6. Because this trend translates almost directly into OH increase at low NOx, the globally averaged OH concentration should also have increased by about 4% per decade. This OH increase has significant implications for the global concentrations of methane, which had been increasing by about 1% per year until the early 1980s, but recent observations indicate a slower increase of about 0.6% per year. This slowing of the methane trend is consistent with the J trend estimated from the TOMS ozone data. While other factors may have been contributing to the changing CH4 trends (i.e., changes in CH4 emissions, and factors which affect OH such as NOx, etc.), the UV-induced component is obviously significant.


It is very difficult to predict what the chemical state of the troposphere will be a few decades from now. In addition to the non-linearities which distinguish the basic HOX - NOX - CO - CH4 system, numerous other tropospheric processes occur. The troposphere is rich and highly varied in its composition, with numerous different compounds of carbon, sulfur, nitrogen, and the halogens. Chemical transformations may occur not only in the gas phase, but also in cloud and rain water, and on the surface of aerosols. Perhaps most importantly, the troposphere is strongly coupled to the other "spheres." One stratospheric coupling is via UV radiation, as discussed in the previous section, yet another is the change in transport of gases across the tropopause (in both directions) if the chemical composition of either the troposphere or the stratosphere is altered. The troposphere is also strongly coupled to the earth's surface which can be source and sink for tropospheric gases. The biosphere plays a central role here, because many of the emissions are related to natural ecosystems and to human activities (here, too, UV can play a role if it can induce significant changes in biological activity on land and ocean). Uncertainties abound. At best, some of these processes are beginning to be understood in one direction, (for example, the effect of isoprene emissions on tropospheric chemistry) but very little is known on the reciprocity of the coupling. Future global atmospheric change, whether it is climate, or UV irradiation, or exposure to altered tropospheric air, will most likely change how the biosphere affects the troposphere, completing a feedback which today is beyond our predictive skills.


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