EFFECTS OF INCREASED SOLAR
ULTRAVIOLET RADIATION ON TROPOSPHERIC
COMPOSITION AND AIR QUALITY
X. Tang (China) and S. Madronich (USA)
Reductions of stratospheric ozone and the concomitant increases of UV-B radiation penetrating to the lower atmosphere result in higher photodissociation rates of key trace gases that control the chemical reactivity of the troposphere. This can increase both production and destruction of tropospheric ozone (O3) and related oxidants such as hydrogen peroxide (H2O2), which are known to have adverse effects on human health, terrestrial plants, and outdoor materials. Changes in the atmospheric concentrations of the hydroxyl radical (OH) may change the atmospheric lifetimes of climatically important gases such as methane (CH4) and the CFC substitutes.
Trends in the photodissociation rate coefficient of tropospheric O3, of about +0.32+/-0.04 percent per year in the northern hemisphere and +0.40+/-0.05 percent per year in the southern hemisphere, have been estimated from satellite measurements of the ozone column between 1979 and 1992. The corresponding model-calculated changes in tropospheric chemical composition are non-linear and sensitive to the prevailing levels of nitrogen oxides (NOx). In polluted regions (high NOx), tropospheric O3 is expected to increase, reaching potentially harmfully concentrations earlier in the day, and leading to more frequent exceedance of oxidant standards for air quality in urban areas where O3 levels are routinely near such air quality thresholds. In more pristine regions (lower NOx), O3 increases can be lower or even negative. Other oxidants, such as H2O2 and OH, are projected to increase for both polluted and pristine regions. Changes to H2O2 concentrations may have some impact on the geographical distribution of acid precipitation. Rural regions may become more urban-like and the percentage of areas with remote tropospheric conditions may decline.
Increases in OH concentrations cause a nearly proportionate decrease in the steady state tropospheric concentrations of CH4 and CFC substitutes such as the HCFCs and HFCs. Thus, the measured reductions in the ozone column (TOMS, 1979-92) are likely to have moderated CH4 increases over the past decade, and may account for about 1/3 of the slowing of the global CH4 trends.
Increased tropospheric reactivity could also lead to changes in the production of particulates such as cloud condensation nuclei, from the oxidation and subsequent nucleation of sulfur of both anthropogenic and natural origin (e.g., carbonyl sulfide and dimethylsulfide). While these processes are still not fully understood, they exemplify the possibility of complex feedbacks between stratospheric ozone reductions, tropospheric chemistry, and climate change.
Ultraviolet-B radiation (UV-B, 280-315 nm) is one of the key environmental factors controlling the chemistry of the lower atmosphere (the troposphere). UV-B is sufficiently energetic to break the bonds of gases such as ozone (O3), nitrogen dioxide (NO2), formaldehyde (HCHO), hydrogen peroxide (H2O2) and nitric acid (HNO3), producing highly reactive atomic and molecular radical species (O, H, OH, HO2, etc.). These photo-fragments effectively drive the tropospheric oxidation cycles, and thus determine the lifetimes and abundance of many other atmospheric chemical species.
UV-B enhancements (resulting from stratospheric O3 reductions, see Chapter 1) are expected to affect tropospheric oxidant concentrations on many geographical scales. Tropospheric O3 is a major constituent of the photochemical smog encountered in many polluted urban environments, and is known to have adverse effects on human health [Lippmann, 1991] and outdoor materials [Andrady, 1993]. On broader suburban and rural scales, oxidants including O3 and H2O2 cause damage to vegetation and play a role in the acidification of rain [Penkett et al., 1979; Moller, 1989; NAPAP, 1991]. Global-scale changes in tropospheric O3 are of concern both because O3 is a major greenhouse gas, and because it is the chemical precursor of the OH radical, the principal cleaning (oxidizing) agent of the global troposphere. Changes in OH concentrations affect directly the rate of removal of many other atmospheric gases involved in climate and ozone chemistry, including methane (CH4) and the hydrogen-containing CFC substitutes (e.g., HFCs, HCFCs) [WMO, 1994].
The rates of some chemical reactions in the troposphere depend directly on the amount of available UV-B [Leighton, 1961]. The reaction rate coefficient, or J value, is given by the expression
where 
is the wavelength (nm),
F() is the spectral actinic flux
(quanta cm-2 s-1 nm-1), 
()
is the molecular absorption cross section (cm2 molec-1),
and 
()
is the photodissociation quantum yield (molec quanta-1).
The spectral actinic flux in the UV-B region is a strong function of the ozone column. The most sensitive J values for key reactions are listed in Table 1 of Chapter 1. Comparable sensitivities have been calculated by Madronich and Granier [1994] and Fuglestvedt et al. [1994], though for different locations and times. Madronich and Granier [1994] calculated the trends in J1, the O3 photodissociation rate coefficient, associated with the 1979-1992 TOMS O3 column data record. Figure 1 shows that the global J1 has increased by about 0.36+/-0.04 percent per year, with slightly higher values in the southern hemisphere (0.40+/-0.05 percent per year) than in the northern hemisphere (0.32+/-0.05 percent per year).
Fig. 6.1. Changes in mid-tropospheric rate coefficients for the photodissociation reaction O3 + hv --> O(1D) + O2, computed using ozone column measurements from the Nimbus 7 satellite (TOMS, version 6 data). Values are given as monthly deviations from the corresponding 1979-1992 averages. Thick solid curve is the area-weighted global average, thin solid line and dotted line are respectively the southern hemisphere and northern hemisphere area-weighted averages. Values next to legend are linear trends, and their corresponding uncertainties, expressed as percent per year relative to the 1979 intercept. From Madronich and Granier [1994].
The photochemical production of tropospheric O3 is due almost exclusively to the photodissociation of NO2,
(<
420 nm) --> NO + O(3P)
followed rapidly by recombination of the atom with molecular oxygen,
Ozone loss occurs when its photodissociation,
(<
320 nm) --> O(1D) + O2
is followed by reaction of O(1D), usually with water vapor:
These last two reaction are also the main source of tropospheric OH radicals [Levy, 1971]. Additional loss of O3 may occur through the catalytic cycle
as well as other reactions. The net ozone production/destruction depends sensitively on the amounts and partitioning of the NOx (= NO + NO2) and HOx (= OH + HO2) species, which are in turn determined by complex sequences of reactions involving carbon monoxide (CO) and hydrocarbons, as well as other tropospheric compounds. Detailed discussions of these reaction sequences are beyond the scope here (but see for example [Finlayson-Pitts and Pitts, 1986]), although some of the non-linearities that they induce may already be evident from the well-known reactions
The HOx species are ultimately removed by termination reactions, leading to the formation of oxidants such as hydrogen peroxide, e.g.
Consideration of the above reactions under scenarios of enhanced UV-B radiation shows that both increased production and increased destruction of tropospheric O3 are possible. The sign and magnitude of the net effect are complex functions of the concentrations of various compounds (especially O3, H2O, NOx, CO, and hydrocarbons), and may well be different for different chemical regimes (e.g., polluted vs. pristine).
Estimates of the tropospheric response to UV changes are still a matter of active research and the results appear to be sensitive to model formulation. Liu and Trainer [1988] used a simple zero-dimensional chemistry model to derive the changes in O3 and OH as a function of prescribed NOx concentrations. They found that a 20 percent loss of stratospheric ozone at northern mid-latitudes could reduce tropospheric O3 by 10-35 percent for NOx levels between 0.01 to 0.10 ppbv (parts per billion by volume), while O3 increases substantially when NOx levels are higher than 0.1-0.2 ppbv (e.g., in many rural and most urban regions). Different results were obtained by Thompson et al., [1991] using a one-dimensional model for different chemical regimes, with decreases in tropospheric O3 concentrations for all regions considered, including relatively polluted areas.
Modeling studies also show that OH concentrations respond to changes in J1 in a complex way. For a one percent increase in J1, the OH concentration increased by about 0.9 percent at low NOx and by 0.3 percent at high NOx [Liu and Trainer, 1988]; by 0.3 percent for marine low latitudes and 0.7 percent for urban mid-latitude regions [Thompson et al. 1991]; and by 0.5-1.0 percent when averaged globally and annually based on a recent three-dimensional model [Fuglestvedt et al., 1994].
H2O2 levels are projected to increase globally [Thompson et al., 1989, 1991; Fuglestvedt et al., 1994]. For scenarios with higher NOx, significantly larger H2O2 (and HO2) levels were calculated for rural and continental conditions, while for the remote regions only a slight increase was obtained.
Thus, it may be concluded from current models that enhanced UV will increase the OH, HO2 and H2O2 concentrations, but the exact amount of the increase remains unclear, with results apparently being sensitive to model formulation. Model-calculated changes in tropospheric O3 appear to be very sensitive to the NOx levels of specific regions. Some O3 is also transported downward from the stratosphere, so that decreases of stratospheric O3 could result in lower O3 input from the stratosphere to the troposphere, but quantitative assessments of this effect have not been made. Measurements by Schnell et al. [1991] show that Antarctic surface O3 decreased by 17 percent over 1976-89 during spring and summer, and were attributed to increased UV-B levels associated with the O3 hole, rather than reduced downward transport from the stratosphere.
Many environmental concerns related to urban and regional air quality are likely to be affected by increased oxidant concentrations. Urban and surrounding areas routinely experience high NOx and volatile organic carbon concentrations due to anthropogenic emissions, which leads to very reactive chemistry and generates high levels of ozone and other oxidants through the photo- chemical smog formation process. Some of these oxidants, such as formaldehyde, HCHO, can produce odd-hydrogen radicals upon absorption of UV-B [Finlayson-Pitts and Pitts, 1986], providing additional chemical channels that could contribute to increased urban reactivity with future increases in UV-B. Increased UV-B radiation could produce even higher levels of urban ozone, as well as potentially harmful concentrations of ozone earlier in the day, and nearer to emission sources and population centers [Gery et al., 1987; Whitten and Gery, 1986]. It was suggested by Gery et al. [1987] that, if UV levels are increased, significantly more stringent regulation of volatile organic emissions will be required in urban locations to maintain photochemical oxidants below air quality standards. De Leeuw and Van Rheineck Leyssius [1991] have similarly estimated that, given constant emissions, air quality violations will be much more frequent in European urban areas if UV levels increase.
Because increased UV-B would make the lower troposphere more reactive, some rural regions may become more urban-like and the percentage of areas with remote tropospheric conditions may decline; this will change the global distribution of polluted areas [Gery, 1993]. On the global scale, the increased tropospheric reactivity caused by UV-B radiation could lead to increases in global aerosol production, providing additional condensation nuclei to the lower stratosphere where heterogeneous ozone destruction processes could become more important in non-polar regions [Gery, 1993].
H2O2 is the most important oxidant of sulfur(IV) in the aqueous phase when the pH is less than about 4.5 [Penkett et al., 1979] and has attracted increasing attention over the past decade, because it was identified as one of the dominant trace species in polar ice [Neftel et al., 1984] and its concentrations have increased by 50 percent over the past 200 years, with most of the increase occurring in the past 20 years [Sigg et al., 1991]. Potential changes to H2O2 concentrations would be interesting with respect to the formation of acidic precipitation. The predicted substantial increase in H2O2 concentrations over large geographical areas caused by increased UV-B levels in troposphere [Gery et al., 1987; Fuglestvedt et al., 1994] may have some impact on the geographical distribution of acid precipitation. However, no quantitative estimates of these effects are yet available.
UV-related increases in tropospheric OH may affect the concentration of other important gases. Methane is removed from the atmosphere primarily via the reaction
with a lifetime of about 10 years. Pre-industrial atmospheric CH4 concentrations were near 600 ppbv, but increased emissions (most likely due to human activities) have brought the current value near 1,800 ppbv. Recent measurements of the rate of CH4 increases show that the trend has slowed in the past decade from about 14 ppbv per year to about 9 ppbv per year. It has been proposed that the slowing of the CH4 trend may be partly due to the increased OH resulting from tropospheric UV increases [Madronich and Granier, 1992, 1994; Fuglestvedt et al. , 1994]. Increase in OH concentrations will decrease the steady state CH4 concentration by nearly the same proportion. The time-dependent response of CH4 depends on the detailed scenario for OH increase. Madronich and Grainer [1992, 1994] used the 1979-92 O3 column measurements (Nimbus 7 TOMS), to calculate the global increase of J1 (see Figure 6.1) and assumed a corresponding linear OH increases at a rate of 0.36 percent per year beginning in 1979. In this scenario, in 1993 the CH4 is reduced by about 35 ppbv from the value it would have achieved without the OH increases, which may be compared with a measured atmospheric increase of about 150 ppbv over the same time [Khalil et al., 1993; Steele et al., 1992]. Three-dimensional model results obtained by Fuglestvedt et al. [1994] are roughly comparable, suggesting that about 1/3 of the slowing is due to the increased UV levels.
As mentioned in previous chapters, both natural terrestrial and aquatic ecosystems are likely to be affected by UV radiation changes. For example, increased penetration of UV radiation into the oceans is expected to affect both the viability of living organisms, and the photochemical processing of non-living organic and inorganic matter. Thus, changes may be expected in the ocean-to-atmosphere emissions of various gases including carbon monoxide (CO), carbonyl sulfide (COS), dimethyl sulfide (DMS), and other carbon and sulfur compounds [Najjar et al., 1994]. Some of these species are believed to be the primary source of natural sulfate aerosols, which have roles in the destruction of ozone, climate change, and acidification of precipitation. For example, aerosols formed by DMS oxidation are believed to influence the ability of marine clouds to reflect incoming solar radiation, and may thus influence climate [Charlson et al., 1987]. Another example is COS, which is believed to be a major source of sulfur atoms to the stratosphere and of the resulting natural stratospheric sulfate aerosol layer. Current theories of stratospheric O3 chemistry suggest that aerosols may play a role in the destruction of O3, so that UV-induced increases in emissions of, e.g., COS, may constitute a yet unquantified positive feedback on stratospheric O3 depletion.
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