NO Emissions in Tg NNOyr

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Figure 2.4. Sources of nitrogen oxides (NO J to the rrnpi»sphere (data from Wang et al., W8). See also color plate section,


Biomass Burning ;11 6)

Fossil Fuel (21.0)

Figure 2.4. Sources of nitrogen oxides (NO J to the rrnpi»sphere (data from Wang et al., W8). See also color plate section,


Ocean (<0,1) Lightning Stratosphere (0 1:

Biomass Burning ;11 6)

Marepnan 3amnineHHbw as ropes um npaeow of which takes place in the Tropics (Andreae, 1993), and the production of NOx by lightning, which is also abundant in the deep convective thunderstorms of the ITCZ. Because vegetation fires can occur only when the vegetation is dry enough to burn, they are most abundant in the dry season, when the trade wind inversion with its large-scale subsidence prevails over the part of the Tropics in question. Because this inversion prevents convection to heights of more than a few kilometers, it was initially thought that the linkage between dry conditions and subsidence more or less precluded the transport of pyrogenic ozone precursors to the middle and upper troposphere. Recent work has shown, however, that large amounts of smoke can get swept by low-level circulation, such as the trade winds, toward convergent regions over the continents or the ITGZ, and there become subject to deep convection (Andreae et aL, 1999; Chatfield et at., 1996; Thompson et aL, 1996). This transport pattern can explain the abundance of fire-related O3 and Oj precursors observed in the middle and upper troposphere by remote sensingand in situ measurements (Browell etal,, 1996; Connors etal., 1996; Olson etal., 1996). Figu re 2.5 shows the distribution of Oj over the tropical South Atlantic during September-October 1992 in comparison with results from earlier studies (DECAFE-88 in the Congo, Andreae et ah, 1992; Tropical Atlantic, Kirchhoff et aL, 1991) and the ozone climatology over the Pacific Ocean. These results show dramatically the impact that O3 from biomass burning can have on the entire tropospheric column.

Whether this impact will grow in the future depends both on climate change and on human factors. The amount of fuel available at a given place for burning is a function of ecological factors, such as soil fertility, precipitation, and temperature. It also depends on land use: whether the area has been burned previously, is used for grazing or agriculture, and so on. If climatic variations become more extreme, as climate models have suggested, we can expect a more frequent occurrence of drought years following very wet years. This would result in large amounts of fuel ready to burn in the fire season. Furthermore, in a warmer climate, fire frequency is likely to increase. That would reduce biomass

And Potassium Availability Soil
Figure 2.5. Impact of tropical biomass burning on the vertical distribution uf ozone over the Equatorial ocean regions. Sec also color plate section

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carbon storage by changing the age class structure of vegetation as well as cause increased emissions of ozone precursors.

Human activities are of central importance to the frequency and severity of biomass fires. If large parts of the humid Tropics are further deforested, they will transition from a biome essentially free of fires (the tropical rainforest) to biomes with much more frequent fires {grazing lands, agricultural lands, and wastelands). With a higher human population density, the frequency of ignition will also go up. And finally, the amount of biomass burned for cooking and domestic heating, already a major source of emissions in tropical countries, will increase further.

2.6 Aerosols: Complex Spatiotemporal Distributions and Radiative Interactions

Over the past decade, a growing amount of attention has been focused on the climatic effects of atmospheric aerosols. Initially stimulated by a discussion on the climate effects of natural and anthropogenic sulfate aerosols (Charlson et al., 1987; Charlson et al., 1992; Schwartz, 1988), research now spans almost all aerosol types and source mechanisms. Recent reviews can be found in a number of books and articles (Andreae, 1995; Charlson and Heintzenberg, 1995; Houghton et al., 1996; Shine and Forster, 1999). In this chapter, there is not enough space to provide an exhaustive review of all the recent exciting developments in this rapidly expanding field, and I limit the discussion to a few less frequently addressed issues: the role of mineral dust, the function of organic aerosols from biogenic precursors, the link between stratospheric ozone and biogenic sulfate, and the influence of aerosols on smog chemistry.

A few key points must be made to put the interactions among aerosols, climate, and biota into perspective. First, there is no clear distinction between anthropogenic and natural sources. Like Oj, aerosols form in the atmosphere from precursor substances, with the rates of production depending simultaneously on the concentrations of several precursor molecules, most of which could be either biogenic or anthropogenic. I luman perturbations can increase or decrease the yields of aerosols from natural precursors, often in surprising ways, as we discuss below.

Second, aerosols interact with climate in much more complex ways than do gaseous molecules. In addition to being able to absorb light (and thereby warm the atmosphere), aerosols can scatter light back into space or enhance the backscattering of light by clouds, something that cools the Earth. Aerosols can also reduce precipitation from clouds, and that enhances clouds' lifetime (a cooling effect). Or, if aerosols absorb radiation and warm an atmospheric layer, that may reduce cloud formation, something that would warm the Earth. Because particles in the atmosphere arc created and removed at time scales of days or less, they are very unevenly distributed and cannot be adequately represented by global means, as can the long-lived greenhouse gases. As a consequence of this complex interaction, aerosol effects on climate are usually calculated using three-dimensional climate models, which attempt to include the in homogeneous distribution of aerosol in time and space. The difficulty both of correctly representing the aerosol distributions in such models and of adequately representing and characterizing the

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atmospheric physics involved is reflected in the large differences between predictions of aerosol radiative forcing from different models, often as large as a factor of 2 or 3 (Shine and Forster, 1999). This compares with differences of about 7%-10% for the forcing estimates for the well-mixed greenhouse gases. Overall, the cooling effect due to aerosols is considered to be roughly about 50%-!00% of the warming effect of the greenhouse gases (Houghton et al., 19%; Shine and Forster, 1999),

Third, aerosols are chemically just as complex as are the gaseous constituents of the atmosphere. This point applies obviously to the organic aerosol, which makes up a substantial fraction of atmospheric particles, but is also true for other aerosol components. It is entirely unrealistic to treat aerosols as simple "pure" compounds, such as ammonium sulfate.

Finally, we must move away from treating aerosols as largely inert products of chemical processes, products that play no important "active" role in atmospheric chemistry. Recent work has shown that reactions in and on aerosols may play an important role in the halogen and sulfur budgets of the atmosphere (Andreae and Crutzen, 1997). Scattering and absorption of UV radiation by aerosols can influence "smog" chemistry in polluted atmospheres (Dickerson et al., 1997). Furthermore, modifications of the optical and chemical characteristics of clouds may have an effect on OH concentrations (Mauldin et al, 1997), In the following paragraphs I illustrate some of these issues.

At first glance, it may be surprising that human perturbations of the atmospheric aerosol load could be significant enough to perturb climate, given that only some 11% of the global aerosol emissions are estimated to come from anthropogenic sources. The production of soil dust and sea spray aerosol, on the other hand, accounts for about 80% of the global source strength ( Andreae, 1995). This view, which has been used to argue against a potential influence from anthropogenic aerosols on climate, has several flaws. First, it is not the source strength that is relevant to the climate effect, but rather it is the amounts present in the atmosphere at any given time, the atmospheric burden. Because seasalt aerosol and dust consist mostly of coarse particles, which are rapidly deposited, their share of the burden (68%) is substantially reduced as compared with other sources. Second, it appears that about half of the soil dust aerosol is mobilized as a result of human disturbance of soils and can therefore be considered "anthropogenic" ( fegen and Fung, 1994; Tegen and Fung, 1995; Tegen et al., 1996). When this factor is taken into account, we find that about half of the global aerosol burden is the result of human activity (Figure 2.6), or, in other words, that humans have approximately doubled the aerosol load of the atmosphere.

The case for a substantial effect of anthropogenic aerosols on climate becomes even stronger when we consider the way aerosols interact with the flux of radiation through the atmosphere. We distinguish two basic mechanisms: the scattering and absorption of radiation by the aerosol particles themselves ("direct effect") and the scattering of light by clouds, which can be modified by variations in the concentration of cloud condensation nuclei ("indirect effect"). Light scattering by aerosols is strongly size-dependent, with a maximum effect when t he size of the particle and the w avelength of the scattered light are of the same order. For this reason, the submicron sulfate, organic, and smoke aerosols from S(>2 emission and biomass burning have stronger radiative

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Emmision Plate

Figure 2,6, Global atmospheric aerosol burden resulting from various sources. See also color plate section.

Global burden for major aerosol types

■ Soil dust {mineral aerosol)

■ Volcanic dust

■ Biological debris

□ Sulfates from biogenic gases

□ Sulfates from volcanic S02

■ Organic matter from biogenic VOC

■ Oust from disturbed soils H Industrial dust, etc.

■ Biomass burning (w/o soot carbon) HS Nitrates from pollution NO*

■ Organic from anth ropogenic VOC

Figure 2,6, Global atmospheric aerosol burden resulting from various sources. See also color plate section.

effects than the more abundant soil dust and seasalt aerosols (Figure 2,7) (Andreae, 1995)* As a result, whereas the anthropogenic fraction of the aerosol burden is about 50%, the anthropogenic share of the radiative effect is higher, about 60%.

The fact that shortwave light absorption by dark aerosols (soot) and longwave absorption by silicate minerals act in a warming direction - the opposite of the cooling effect of light scattering - makes the assessment of the net climate effect very difficult. Current estimates suggest that the cooling effect predominates, with a global mean forcing of ^—0,4 W m but with an uncertainty of about ±0,8 W m J (Shine and Förster, 1999).

The indirect effect by means of enhancement of cloud albedo resulting from increased cloud condensation nuclei (CCN) concentrations is related to the number rather than the mass concentration of aerosols. Here again, a very rough estimate based on the burden of fine aerosols, which account for most of the CCN burden, suggests that the introduction of anthropogenic particles has more than doubled the amount of CCN in the atmosphere (Andreae, 1995). Because the effect of added CCN is very sensitive to the number of CCN already present and to the type of cloud into which the CCN are introduced (Twomey, 1977), knowledge of the spatiotemporal d istribution of CCN sources is critical to an assessment of their effect.

Because the strongest climate effects from increased CCN concentrations arc expected for clouds of intermediate optical thickness and low initial (natural) CCN concentrations, the regions of most concern used to be the large areas of marine stratus in the eastern parts of the ocean basins (Charlson et ah, 1987). Continental clouds were not thought to be very susceptible to the indirect effect, because it was assumed that continental air had high natural CCN levels. This assumption may have to change in light of recent observations of very low CCN concentrations over Amazonia in the wet season (Roberts et al., 1998). If these measurements prove representative of CCN levels over the tropical continents, they imply that deep convection and rain formation in these regions are occurring naturally at very low cloud droplet number concentrations (CDNC), resulting in very high precipitation probability. Given the low natural CCN concentrations* it would not require very high amounts of anthropogenic emissions to significantly increase CDNC, which would change the rainout efficiency ("overseeding") and could lead to significant changes in the regional water cycle. This could even influence the water vapor content of the tropical troposphere as well as the en-ergv transfer processes in the tropical lladley celL The complexity of ice formation in clouds and our sparse knowledge about identity and sources of ice-nucleating panicles in clouds further complicate an assessment of the human impact on tropical clouds (Baker, 1997).

So far, we have ignored interactions between human activities and the rate of production of "natural" aerosols. However, this may not be appropriate in a number of instances. Consider the production of sulfate aerosols from marine biogenic dimethyl sulfide (DMS), which has been proposed as the main source for CCN in pristine marine regions {Charlson et ah, 1987). An important step in this process is the production of new particles (nucleation) from the gaseous precursor, 11?SO4. The rate at which this occurs depends on the concentration of gaseous H2SO4 molecules, which in turn

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Figure 2,7. Extinction of visible radiation at the Rarth's surface resulting from plate section.

Radiative effect (at Earth surface) from major aerosol types

ISoii dust (mineral aerosol} I Sea salt I Volcanic dust I Biological debris

□ Sulfates Irorn biogenic gases

□ Sulfates Irom volcanic S02

■ Organic matter from biogenic VOC

□ Nitrates from biogenic NO, I Dust from disturbed soils

SB industrial dust, etc I Soot carbon

■ Sulfates Irom SOs

■ Biomass burning (w/o soot carbon) 3 Nitrates from pollution NO*

■ Organic from anthropogenic VOC

Figure 2,7. Extinction of visible radiation at the Rarth's surface resulting from plate section.

the various aerosol types. See also color depends on their rate of production from the reaction between SO? and OI I. Any perturbation of the atmosphere that changes the OH levels in the marine boundary layer may therefore interfere with the rate of new particle and CCN production and have an influence on climate.

One such perturbation may be the increased U V flux that reaches the Earth's surface as a result of the thinning of the stratospheric UV layer (Tang and Madronich, 1995). Toumi et al. (1994) proposed that stratospheric ozone loss from increasing halogen levels may have led to a 3% increase in OH concentration during the 1980s, something that would have resulted in an increase in the production rate of CCN. However, the gaps in our understanding of the mechanisms of CCN production and the climate effects of changing CCN are so large that the uncertainty in the predicted climate effect ranges from an insignificant value to one that would more than compensate for the increased greenhouse forcing over the same period. It may be interesting to explore what changes in the OH concentration in the marine atmosphere have occurred and may still occur as a result of increasing CH4, CO, and NOx levels, and what effects those changes may have on CCN concentrations and climate.

An even more striking example of an impact of anthropogenic activities on the production of "natural" aerosols is the oxidation of biogenic v olatile organic carbon (VOC) compounds, particularly terpenes, to low-volatility compounds, which condense into aerosol particles (Kanakidou, 1998). The reaction mechanisms and products depend on the chemical environment in which the oxidation reactions occur. In the presence of elevated levels of NOx, ozone is formed, and it reacts with terpencs very rapidly to form products of low volatility and high aerosol yield (Bowman ct aL, 1995; Hoff mann et al., 1997; Seinfeld and Pandis, 1998). At low NOx levels, on the other hand, ozone production is low and terpenes are oxidized predominantly via attack by OI 1, with lower aerosol yields. Recent estimates suggest that more than 1000 Tg C are emitted annually in the form of biogenic VOC, of which maybe 30% are potential aerosol precursor substances (Guenther et al., 1995). In view of the vast amounts of VOC emitted from vegetation, even small changes in the aerosol production efficiency result in major perturbations of the atmospheric aerosol budget (Andreae and Crutzen, 1997). Kanakidou (1998) has estimated that as much as 80% of the global organic aerosol production may be the indirect result of human impacts on atmospheric chemistry.

The final point 1 would like to address in this section is the influence of aerosols on gas phase chemistry. The potential of seasalt aerosols to act as a source of reactive gaseous halogen species in the marine boundary layer is now well recognized (Andreae and Crutzen, 1997; Keene et al., 1998; Sander and Crutzen, 1996; Sander et al., 1997; Yogt et al., 1996). This process can be responsible for significant rates of photochemical hydrocarbon oxidation, O3 destruction, and other reactions in the marine boundary layer (MBL) in addition to OH-bascd chemistry. Mineral dust in the atmosphere is also far from being an inert substance, It can act as a sink for acidic trace gases, such as SO2 and HNO3, and thereby interact with the sulfur and nitrogen cycles (Dentener et al., 1996; Li-Jones and Prospero, 1998; Talbot et al., 1986). Coatings with soluble substances, such as sulfate or nitrate, change the ability of mineral dust aerosols to nucleate cloud droplets (Levin et al., 1996).

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In addition to direct chemical interactions, aerosol particles can influence atmospheric chemistry by modifying the UV radiation field. Submicron particles scatter UV radiation very efficiently, and the aerosol loadings present in polluted environments can easily triple or even quadruple the turbidity of the atmosphere at UV wavelengths. This increases the effective path length of the photons in a smoggy environment, and it can thereby enhance the rates of photochemical reactions, including ozone formation (Dickerson et alM 1997). There may be even an analog of the "indirect" climate effect: As variations in CCN change cloud radiative properties, they may also influence the actinic flux in and above clouds and thereby influence the production of the OH radical (Madronich, 1987; Mauldin et aL, 1997).

2.7 Climate-Chemistry Feedbacks and the Arctic "Ozone Hole"

In the 1990s, ozone loss in the Arctic stratosphere during the polar sunrise period accelerated dramatically, peaking in 1997 with an ozone loss of ^100-120 Dobson unit (DU) (Muller et al., 1997; Newman et al., 1997). This decrease is comparable to the ozone loss over Antarctica in 1985, at the time when the Antarctic "ozone hole" was first reported (Farman et al., 1985). At the same time, temperatures in the polar vortex dropped to lower temperatures and remained there over longer periods than before. These observations are closely connected: For the development of the chemical reaction sequence that leads to rapid ozone loss, it is necessary that temperatures fall low enough for polar stratospheric clouds (PSC) to form. These clouds must persist long enough for the reactions that regenerate active chlorine species (Cb, CI, CIO) from the inactive reservoir species (CIONO2,1IC1), and for HNOi-rich PSC particles to settle out of the stratosphere. After ozone loss has taken place, the lesser amounts of Oj present result in less absorption of UV radiation in the stratosphere, and consequently in lower temperatures, which in turn promote ozone loss. Consequently, stratospheric temperature change, PSC] formation, and ozone depletion form a feedback system with positive gain, mutually reinforcing one another (Danilin et al, 1998; Portmann et al., 1996).

The situation is further complicated by the fact that the build-up of greenhouse gases in the atmosphere, while it warms the lower atmosphere, actually cools the stratosphere (Ramaswamy and Bowen, 1994). The cooling of the Arctic stratosphere during the past decade, which made rapid ozone loss in the polar region possible, can thus be attributed to three causes: the ozone loss itself through the feedback described in the preceding paragraph; the radiative forcing due to the greenhouse gases; or unrelated fluctuations in the climate system (or a combination of these factors). Which of these mechanisms dominates is crucial to the time scale at which we can anticipate the recovery from low ozone conditions over the Arctic. If the ozone-temperature feedback dominates, recovery will occur approximately at the time scale at which chlorine concentrations in the stratosphere return to levels below those that cause rapid ozone loss. If, on the other hand, cooling was mostly caused by the effect of CO? and other greenhouse gases, the temperature/ozone-loss feedback will persist much longer ("slow" recovery). This is because less stratospheric chlorine is required to cause rapid ozone loss at the lower temperatures that would prevail as long as greenhouse gas concentrations laiepnan, 3amnineHHbiw aBTopcKMM npaBOM

remain elevated. If, finally, the observed cooling were mostly caused by unrelated climate fluctuations, recovery would be unpredictable altogether, injection of large amounts of volcanic aerosol into the stratosphere, another unpredictable "external" forcing, would also act to accelerate O3 loss and delay recovery (Portmann et al, 1996).

Recent modeling studies suggest that all these mechanisms in fact contributed to the development of an Arctic ozone hole in the 1990s, with the largest effects resulting from the cooling driven by ozone loss interacting with a natural mode of variability (Graf et al., 1998; Shindell ct al., 1998), This mode links a strengthened polar night vortex with an enhanced North Atlantic oscillation (Graf et al., 1998). Shindell et al+ (1998) have proposed a mechanism by which anthropogenic climate change might be coupled to a strengthened polar vortex. They find that in their model simulations the changes in temperature and winds resulting from increased greenhouse gas concentrations alter the propagation of planetary waves. As a result, planetary waves break up the Arctic polar vortex less frequently, and that leads to significantly colder temperatures existing over longer periods of time in the Arctic stratosphere. These authors estimate that because of this effect the ozone loss over the Arctic by the vear 2020 will be double what w mf it would be without greenhouse gas increases, and that recovery from Arctic ozone depletion will be delayed by some 10-15 years.

2.8 Conclusion

This chapter examines the linkages and connections between human perturbations of the Earth System and its chemical, physical, and ecological characteristics. We have seen that it is usually not adequate to consider only the emission of trace gases and aerosols; it is essential also to consider the complex interconnections between any given perturbation and the overall Earth System.

I n a few cases, it is relatively easy to assess the impact of anthropogenic emissions on the atmosphere - for example, when the sources of a substance are industrial and its sinks are chemical reactions with first-order kinetics, Hut in most instances, the emission and removal of climatically active gases and aerosols depend on a multiplicity of human activities and ecological factors, including climate itself

When land use and agricultural practices change, the emissions of trace gases such as N^O, NO, and CH4 change in highly complex ways, which are extremely difficult to assess at the scales of interest. When land use change reaches such vast extent as in the deforestation of the Tropics, it may even cause changes in the climate system, including the hvdrologieal cycle. As a result of these chemical and physical perturbations, the chemical functioning of the atmosphere, and consequently the production rates and lifetimes of aerosols and greenhouse gases, will be modified. The most obvious example for such a mechanism is the large-scale change of trace gas inputs into the tropical troposphere, the vast photochemical reactor where most of the photooxidation of long-lived trace gases takes place. Because of the long time scales involved in ecological change, biogeochemical cycles and climate have a memory of past land use change, and, conversely, current land use change may have long-term consequences reaching far into the future.

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In some cases, the human perturbation consists of the release of a precursor compound (e.g., SO2K which is transformed in the atmosphere to a climatically active substance. In this example, the actual amount of radiativcly active sulfate aerosol produced is determined by a complex interplay of atmospheric transport processes, chemical processes in the gas phase, and interactions with other aerosol species. In other cases, such as the production of organic aerosols from biogenic VOCs or sulfate aerosols from DMS, aerosol yields can be modified by anthropogenic changes in atmospheric photooxidation

At longer time scales, we must consider feedback loops in which climate change results in different circulation patterns, changes in water availability at the surface, changes in water vapor content of the atmosphere, and so on. These factors in turn modify the sources, sinks, and atmospheric budgets of trace gases and aerosols, again affecting climate. A dramatic example of this kind of interaction is the coupling between changes in stratospheric temperatures and ozone depletion, something that has shown up over the Arctic during the last decade.

Understanding the complex interactions between tropospheric chemistry and global change presents a formidable scientific challenge. Exciting progress has been made in this area, especially over the past decade, by intensified cooperation between scientific disciplines, close interaction between observation and modeling, and broad international cooperation. However, in our excitement about new conceptual insights into the complexity of the Farth System's workings, we must not lose sight of the fact that the observational database for testing our concepts and models remains rather sparse. High priority must therefore be given to developing new tools and programs for the investigation of our changing planet.

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