M. O. Andreae
Biogeochemistry Department
Max Planck Institute for Chemistry
P.O. Box 3060
D-55020 Mainz, Germany
 
 

Abstract

Human activities are changing the composition of the atmosphere not only directly through the emission of trace gases and aerosols, but also indirectly through perturbations in the physical, chemical, and ecological characteristics of the Earth System, which in turn influence the rates of production and loss of atmospheric constituents.

The impact of direct anthropogenic emissions on the atmosphere is often relatively easy to assess, especially if they are tied to major industrial activities, where accurate and detailed records are kept for economic reasons. Classical examples are the release of chlorofluorocarbons and the emission of CO₂ from fossil fuel combustion. There are, however, also cases where it is much more difficult to obtain accurate emission estimates. An example is biomass burning, for which no economic incentive for record keeping exists, and which takes on many forms, each with a different emission profile.

A more complex case exists where human activities release a precursor compound, which is transformed in the atmosphere to a climatically active substance. This can be illustrated using the example of SO₂, from which sulfate aerosol can be formed. The actual amount of radiatively active sulfate aerosol produced, however, is determined by a complex interplay of atmospheric transport processes, chemical processes in the gas phase, and interactions with other aerosol species.

Some of the most important anthropogenic modifications of the atmosphere, however, are the indirect results of human-caused changes in the functioning of the Earth System. For example, when land use and agricultural practices change, the emissions of trace gases such as N₂O, NO, and CH₄ change in highly complex ways, which are extremely difficult to assess at the scales of interest. An even higher level of complexity is encountered, when human activities affect the atmospheric levels of some species, which in turn changes the chemical functioning of the atmosphere, and consequently the production rates and lifetimes of aerosol and greenhouse gases. An example for such a mechanism may be the large-scale change of trace gas inputs into the vast photochemical reactor of the tropical troposphere, where most of the photooxidation of long-lived trace gases takes place.

Finally, we must consider feedback loops, where global change begets global change. Climate change, caused by upsetting the Earth’s radiative balance, results in different circulation patterns, changes in water availability at the surface, water vapor content of the atmosphere, etc., all of which modify the atmospheric budgets of trace gases and aerosols. Thinning of the stratospheric ozone layer yields a higher UV flux into the troposphere, thereby accelerating photooxidation processes.

Understanding the complex interactions between tropospheric chemistry and global change presents a formidable scientific challenge, which can only be addressed by close cooperation between scientific disciplines, tight interaction between observation and modeling, and broad international cooperation.

Introduction

In the public mind, "global change" has become almost synonymous with "global warming" or "climate change", a narrow reduction of the original meaning. While there is no doubt that the possibility of climate change is of great concern to the Earth’s population, we must not forget that we are living in a period when almost all components of the Earth system are undergoing change. The chemical composition of the atmosphere is being perturbed at a vast scale by human activities. The terrestrial biota are modified by land use change, biomass burning, deforestation, and species extinction. Marine life is impacted by overfishing, eutrophication and pollution. There is a tendency to see these issues as independent environmental problems, each grabbing the public’s attention for some time, and each demanding a specific solution.

This approach obscures the fact that all these phenomena are occurring simultaneously, and within the same "Earth system". As a result, they interact with one another, reinforcing or damping each other, or changing each other's temporal evolution. This view is reflected in the "Bretherton" diagram (Figure 1),

Figure 1: The "Bretherton" diagram. This version of the diagram has been produced by the Earth System Science Education Program, Universities Space Research Association,  Whitelaw, Wisconsin (http://www.usra.edu/esse/essonline/whatis.html)

which shows the complex linkages between human activities, physical climate system and biogeochemical cycles.

It is especially important to examine the Earth system for possible feedbacks, which amplify the effect of perturbations. It is already well established, that increasing temperatures result in changes in ice albedo, atmospheric water vapor content and cloudiness, which in turn act to increase temperature. If additional positive feedbacks exist, they would add to known feedbacks, and, due to the extremely non-linear behavior at higher gains, could have disproportionately large effects (Lashof et al., 1997).

In this paper, I will explore some of these interactions between human activities, atmospheric chemistry, climate, and ecology, using selected examples or case studies. I will proceed from the (relatively) simple to the more complex, keeping in mind that exploring any of the issues addressed here in its full depth and complexity is well beyond the frame of an overview paper such as this.
 
 

1. The "Simplest" Case: Anthropogenic Halogenated Hydrocarbons

The clearest evidence of a global change in the Earth system is the changing composition of the atmosphere, particularly the increasing concentrations of some long-lived trace gases emitted by industrial activities. This was first documented for CO₂ by the long-term measurements made by C. D. Keeling at Mauna Loa Station on Hawaii (Bacastow et al., 1985; Keeling et al., 1995), and subsequently for numerous other trace gases such as methane (CH₄), nitrous oxide (N₂O), carbon monoxide (CO), and numerous halogenated hydrocarbons (Houghton et al., 1996, and references therein).

The latter class of compounds includes species such as the chlorofluorocarbons (CFCs, Freons), methyl chloroform (trichloroethane), and the partially halogenated chlorofluorocarbons (HCFCs), which are exclusively man-made and have no natural sources. Since they are produced industrially, often by relatively few manufacturers, accurate records exist on the amounts and times at which they were produced and released into the environment. Most anthropogenic halogenated hydrocarbons have no significant biological sinks and are resistant to hydrolysis in aquatic systems, so that their only major sink is photochemical breakdown in the atmosphere. For most substances, this sink follows first order kinetics, i.e., its rate is a linear function of the trace gas concentration.

As a result of the well-characterized and simple source and sink functions of these gases, their concentration in the atmosphere as a function of space and time coordinates can be relatively easily understood and modeled. Figure 2 illustrates this behavior with the example of the temporal record of methyl chloroform, a substance that has no known natural sources, and is removed almost exclusively by reaction with tropospheric OH. The use of methyl chloroform is severely restricted by the Montreal Protocol, and consequently its production declined sharply around 1990. This resulted in a reversal of its atmospheric trend in 1991, from an average increase of 4.5±0.1 %/year to a decline of about 14 %/year in 1995/96. This behavior can be described in an atmospheric model and used to deduce both its weighted-mean atmospheric lifetime (4.8 years), and the global-mean OH concentration (Montzka et al., 1996; Prinn et al., 1995).

Figure 2: Temporal evolution of the atmospheric mixing ration of methyl chloroform (CH₃CCl₃) in the troposphere over the Northern Hemisphere (from Kurylo et al. (1999)).

While we have used these compounds as examples of the simplest case, with minimal feedback processes, we should note that even here some complications might arise. If any of these substances could change the stratospheric ozone density to such a degree that it would significantly alter the UV flux into the troposphere, and thus the tropospheric OH abundance, it could influence its own lifetime. While it has been argued that significant changes in tropospheric OH have not occurred over the time scale represented in Figure 2 (Prinn et al., 1995), this issue is not without controversy (Krol et al., 1998). This issue will be discussed in more detail below.
 
 

2. A More Complex Case: CO₂

In a way, CO₂ represents the opposite extreme from the previous case: It is subject to strong and complex biogeochemical interactions, and its anthropogenic sources are but minor perturbations on the natural fluxes. The annual fluxes of CO₂ in and out of the terrestrial and marine biota make up some 150 Pg C per year (1 Pg = 10¹⁵ g), while the emission from fossil fuel combustion and cement manufacturing accounts for "only" about 6 Pg year^-1. Yet, it is this small increment added to the large biogenic fluxes of CO₂ that is responsible for most of the growth of its concentration in the atmosphere and for about half of the greenhouse gas effect. The rest of the atmospheric CO₂ increase is due to the effect of tropical deforestation, which moves carbon out of the "long-lived terrestrial biomass" reservoir into the atmosphere.

In order to understand and predict the atmospheric abundance of CO₂, we need a thorough understanding of all the complex biogeochemical interactions that control its transfer between the Earth’s compartments, including the deep and shallow oceans, the marine and terrestrial biota, the sediments and soils, and so on. There are a large number of known feedbacks between climate and the cycles of carbon and the "nutrient" elements (N, P, and S), and quite likely an even greater number that are still unexplored (see, for example, Lashof et al., 1997). Consequently, CO₂ is probably the most "interesting" trace gas to a biogeochemist. To an atmospheric chemist, however, it is quite "boring", since it does not undergo any relevant chemical reactions in the atmosphere. I will, therefore, not address the global carbon cycle in any detail in this paper, but restrict myself to these few short remarks.

It may be worthwhile, however, to already emphasized one point here, to which we will come back to several times in the following sections: the importance of the tropics in understanding global change. The tropics are the part of the globe with the most rapidly growing population, the most dramatic industrial expansion, and the most rapid and pervasive change in land use and land cover. The tropics contain also the largest standing stocks of terrestrial vegetation, and have the highest rates of photosynthesis and respiration (Houghton and Skole, 1990; Raich and Potter, 1995). It is therefore likely that changes in tropical land use will have a profound impact on the global carbon cycle in future decades (Houghton et al., 1998).
 
 

3. Trace Gases With Very Complex Source And Sink Patterns: CH₄, N₂O

After CO₂, methane is the most important greenhouse gas, and there is unequivocal evidence that its atmospheric concentration is increasing due to human activities. In an effort to understand this increase, a large effort has gone into elucidating the budget of this trace gas.

Figure 3 shows the sources and sinks of methane in the form of pie diagrams. It is evident from this figure that a large variety of sources contribute to the atmospheric methane budget, and that most of these sources are of comparable order of magnitude. The sources which are under human "control" (fossil fuel burning and handling, animal husbandry, rice agriculture, biomass burning, and waste disposal) make up almost three-quarters of the methane sources. In contrast to the halogenated hydrocarbons, however, which are produced industrially, most methane is produced biologically through microbial fermentation (with the exception of the fossil fuel source). Therefore, methane production is dominated by the harnessing of biological processes through human activity. For example, methane production from rice agriculture closely resembles that from natural wetlands, and the emissions from cattle are similar to those from the large herds of wild ruminants that grazed on natural grasslands.

The diversity of sources of methane, and their biological nature results in a complex source pattern, which makes assembling an accurate source inventory quite difficult. There are a considerable number of different rice farming practices, for example, all of which result in substantially different amounts of methane emitted per amount of rice produced, or per hectare farmed. On the other hand, this diversity provides a chance for mitigation by selecting practices that result in low methane emission factors. As a result, considerable effort has gone into refining our understanding of the processes that control methane emission from the various source types, and into producing more accurate source estimates.

A very different picture prevails on the sink side of the methane budget. The term labeled "storage" represents the amount of methane accumulating annually in the atmosphere, a value that can be relatively easily and accurately obtained from the existing measurement network. It appears in the budget as a "sink", in the sense of being a reservoir where some of the methane introduced into the atmosphere ends up, at least for some years. The true sinks, i.e., the processes removing methane from the atmosphere, are dominated by one process, the tropospheric oxidation of methane through its reaction with the hydroxyl radical, OH. This sink is so large, that even a relatively small uncertainty in its magnitude, some 10-20%, is as large a many of the individual methane sources.

It may seem, at first glance, that this sink is rather well known and predictable. The global mean OH concentration can be derived from analysis of the distribution in time and space of compounds such as methyl chloroform, as we have discussed above. The reaction rate constant of methane with OH, and the global concentration distribution of methane are also well known. Using present-day measured CH₄ and model-predicted OH concentrations, we obtain a photochemical sink that is of the right size to balance the methane budget. Yet, how well can we extrapolate this knowledge back into the past, or, more importantly, forward into the future?

Model calculations suggest that tropospheric OH concentrations have decreased by some 25% since the industrial revolution (Crutzen, 1995a). This value is, however, highly uncertain, and other models have predicted much smaller or much larger changes (Thompson, 1992). Furthermore, there is an ongoing controversy about changes in tropospheric OH over the last few decades (Krol et al., 1998; Prinn et al., 1995). The outcome of this discussion has important implications for our understanding of the current methane budget. If the methane lifetime has not been significantly affected by likely changes in OH, the decreasing rate at which methane accumulates in the atmosphere would be consistent with methane sources having reached a plateau, and the atmospheric methane concentration approaching a new steady state (Dlugokencky et al., 1998). If, on the other hand, OH had been growing significantly over the past decade, we would have to assume that methane sources are currently increasing as well.

To look into the future, we must remember that most methane oxidation takes place in the tropics, due to the high concentrations of OH resulting from high amounts of water vapor and high UV flux in that region (Andreae and Crutzen, 1997). The tropics are also the world’s most rapidly changing region. Deforestation of the Amazon Basin, and subsequent agricultural and industrial development is likely to substantially change the amounts of hydrocarbons and nitrogen oxides released into the tropical atmosphere, resulting in elevated ozone over the region. At the same time, deforestation would change the regional water balance including the atmospheric water vapor content. Since ozone and water vapor are the precursors of the OH radical, these changes must be expected to have a pronounced influence on OH and consequently on the lifetimes of CH₄, CO, and all the other atmospheric trace gases which are being removed by reaction with OH. Changes in cloudiness resulting from a perturbation of the hydrological cycle in the humid tropics would also impact the OH distribution because of the radiative and chemical effects of clouds on this radical (Mauldin et al., 1997).

The challenge of predicting future CH₄ levels is further complicated by the fact that atmospheric CH₄, CO, and OH are part of a coupled chemical reaction scheme, with complex, nonlinear behavior resulting from simple perturbations (Prather, 1996). Adding CO to this system actually leads to increased CH₄ concentrations, and the couplings and feedbacks in the system result in effects that take longer to decay than the lifetimes of the individual molecules involved. The nonlinearities in this system increase with the concentrations of methane (and CO present in the atmosphere), and at methane source fluxes around three times the present size, runaway growth of methane could occur.

When the possibilities of additional feedbacks with climate and biota are considered, even more complex feedbacks can be expected. The warming resulting from the greenhouse effect may release additional methane from wetlands due to enhanced microbial activity at elevated temperatures, at least initially (Cao et al., 1998; Chapman and Thurlow, 1996; Christensen and Cox, 1995; Lashof et al., 1997; Oechel and Vourlitis, 1994). At longer time scales, effects of changing water table levels and soil moisture content resulting from climate change may reverse the direction of this feedback. Global warming may also liberate methane locked into clathrates in continental slope sediments and permafrost (Harvey and Huang, 1995). This effect may in part be counteracted by stabilization of clathrates due to increased pressure resulting from rising sea levels (Gornitz and Fung, 1994). The overall impact of this feedback process is difficult to assess due to great uncertainties about the amounts of CH₄ present in clathrates, but is thought to be important mostly in the more distant future (beyond the 21^st century) and for high climate sensitivities.

In the previous section, we have already pointed out that gases that are not themselves greenhouse gases may have a climatic effect because they change the rates of production or destruction of greenhouse gases. In this sense, we can attribute a climate forcing and greenhouse warming potential to gases such as CO, which has no significant radiative effect of its own. This is because adding CO to the atmosphere increases the lifetime and abundance of methane, results in the production of ozone, and, following oxidation, adds some CO₂ to the atmosphere. When these effects were simulated in a photochemical model, the cumulative radiative forcing due to CO emissions exceeded at shorter times scales (<15 years) that due to anthropogenic N₂O, one of the important greenhouse gases (Daniel and Solomon, 1998).

Tropospheric ozone, a gas that has no direct emission sources, is the third most important greenhouse gas after CO₂ and CH₄ (Houghton et al., 1996; Portmann et al., 1997; Roelofs et al., 1997; Shine and Forster, 1999; van Dorland et al., 1997). Since ozone has a chemical lifetime in the troposphere that is of the same order as the timescales of many atmospheric transport processes (days to weeks), its temporal and spatial distribution is highly inhomogeneous. In the absence of vertically resolved and globally representative data sets on O₃ concentrations, the climatic effect due to this gas must therefore be estimated based on model calculations. The chemical precursors of ozone are hydrocarbons (including methane and NMHC), CO, and the oxides of nitrogen, NO_x. The latter play an especially important role in the ozone budget, since their abundance determines if the photochemical oxidation of hydrocarbons and CO results in net O₃ production or destruction (Crutzen, 1995b; National Research Council (U.S.) Committee on Tropospheric Ozone Formation and Measurement, 1991).

In many regions of the Earth, especially on the continents, biogenic NMHC emissions are relatively abundant (Fehsenfeld et al., 1992; Guenther et al., 1995), and, in the absence of strong NO_x emissions, their photooxidation results in net O₃ destruction. When NO_x emissions in these regions increase due to development, or because deforestation lets NO_x from soil microbial production escape more readily into the troposphere, the system can switch to net O₃ production, strongly enhancing ozone levels (Keller et al., 1991). This is especially critical in the tropics where O₃ can be entrained into the intertropical convergence zone (ITCZ) and transported by deep convection into the upper troposphere, where it has the strongest climatic effect. Modeling studies suggest that input of pollutants into convective regions may have strong effects on O₃ levels in the free troposphere (Ellis et al., 1996).

Figure 4: Sources of nitrogen oxides (NO_x) to the troposphere (data from Wang et al., (1998))

Figure 4 shows the sources of nitrogen oxides to the troposphere. Of particular importance to the tropical atmosphere are the emissions from biomass burning, most of which takes place in the tropics (Andreae, 1993), and the production of NO_x 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 amount of smoke can get swept by low-level circulation, e.g., the trade winds, towards convergent regions over the continents or the ITCZ, and there become subject to deep convection (Andreae et al., 1999; Chatfield et al., 1996; Thompson et al., 1996). This transport pattern can explain the abundance of fire-related O₃ and O₃-precursors observed in the middle and upper troposphere by remote sensing and in-situ measurements (Browell et al., 1996; Connors et al., 1996; Olson et al., 1996). Figure 5 shows the distribution of O₃ over the tropical South Atlantic during September-October 1992 in comparison with results from earlier studies (DECAFE-88 in the Congo (Andreae et al., 1992); Tropical Atlantic (Kirchhoff et al., 1991)) and the ozone climatology over the Pacific Ocean. These results show dramatically the impact 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, e.g., soil fertility, precipitation, and temperature. It also depends on land use, i.e., if 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, which would reduce biomass carbon storage by changing the age class structure of vegetation, as well as causing 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 go up as well. And finally, the amount of biomass burned for cooking and domestic heating, already a major source of emissions in tropical countries, will increase further.
 
 

5. Aerosols: Complex Spatiotemporal Distributions and Radiative Interactions

Over the last 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 practically 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 paper, there is not enough space to provide an exhaustive review of all the recent exciting developments in this rapidly expanding field, and I will 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 between aerosols, climate, and biota into perspective. First, there is no clear distinction between anthropogenic and natural sources. Just like O₃, 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. Human perturbations can increase or decrease the yields of aerosols from natural precursors, often in surprising ways, as we will discuss below.

Second, aerosols interact with climate in much more complex ways than 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, which cools the Earth. Aerosols can also reduce precipitation from clouds, which enhances their lifetime (a cooling effect). Or, if they absorb radiation and warm an atmospheric layer, that may reduce cloud formation, which would warm the Earth. Since particles in the atmosphere are created and removed at timescales of days or less, they are very unevenly distributed, and can not be adequately represented by global means, like 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 inhomogeneous distribution of aerosol in time and space. The difficulty of correctly representing both the aerosol distributions in such models, and of adequately representing and characterizing the 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 two or three (Shine and Forster, 1999). This compares to 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 50-100% of the warming effect of the greenhouse gases (Houghton et al., 1996; Shine and Forster, 1999).

Third, aerosols are chemically just as complex as the gaseous constituents of the atmosphere are. 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 have to move away from treating aerosols as largely inert products of chemical processes, which 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 will illustrate some of these issues.

Figure 6: Global atmospheric aerosol burden resulting from various sources

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 the amounts present in the atmosphere at any given time, the atmospheric burden. Since seasalt aerosol and dust consist mostly of coarse particles, which are rapidly deposited, their share of the burden (68%) is substantially reduced as compared to 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" (Tegen 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 6), or, in other words, that humans have approximately doubled the aerosol load of the atmosphere.

Figure 7: Extinction of visible radiation at the Earth's surface resulting from the various aerosol types.

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 the size of the particle and the wavelength of the scattered light are of the same order. For this reason, the submicron sulfate, organic, and smoke aerosols from SO₂ emission and biomass burning have stronger radiative effects than the more abundant soil dust and sea salt aerosols (Figure 7) (Andreae, 1995). As a result, while the anthropogenic fraction of the aerosol burden is about 50%, the anthropogenic share of the radiative effect is higher, about 60%.

The fact that short-wave light absorption by dark aerosols (soot) and long-wave absorption by silicate minerals acts in a warming direction, opposite to 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 ca. -0.4 W m^-2, but with an uncertainty of about ±0.8 W m^-2 (Shine and Forster, 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, suggest 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 on the type of cloud into which the CCN are introduced (Twomey, 1977), knowledge of the spatiotemporal distribution of CCN sources is critical to an assessment of their effect.

Since the strongest climate effects from increased CCN concentrations are 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 al., 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 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 would imply that deep convection and rain formation in these regions is 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 ("over-seeding") and could lead to significant changes in the regional water cycle. This could even influence the water vapor content of the tropical troposphere, and the energy transfer processes in the tropical Hadley cell. The complexity of ice formation in clouds and our sparse knowledge about identity and sources of ice-nucleating particles in clouds further complicates 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. This may, however, 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 region (Charlson et al., 1987). An important step in this process is the production of new particles (nucleation) from the gaseous precursor, H₂SO₄. The rate at which this occurs depends on the concentration of gaseous H₂SO₄ molecules, which in turn depends on their rate of production from the reaction between SO₂ and OH. 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 UV flux which 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 over the period of the 1980s, which 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 which would more than compensate 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 NO_x levels, and what effects that may have on CCN concentrations and climate.

An even more striking example for an impact of anthropogenic activities on the production of "natural" aerosols is the oxidation of biogenic volatile organic carbon compounds (VOC), 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 NO_x, ozone is formed, which reacts with terpenes very rapidly to form products of low volatility and high aerosol yield (Bowman et al., 1995; Hoffmann et al., 1997; Seinfeld and Pandis, 1998). At low NO_x levels, on the other hand, ozone production is low and terpenes are oxidized predominantly via attack by OH, with lower aerosol yields. Recent estimates suggest that over 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 I 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; Vogt et al., 1996). This process can be responsible for significant rates of photochemical hydrocarbon oxidation, O₃ destruction, and other reactions in the MBL in addition to OH-based 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, will change the ability of mineral dust aerosols to nucleate cloud droplets (Levin et al., 1996).

In addition to direct chemical interactions, aerosol particles can also 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 can thereby enhance the rates of photochemical reactions, including ozone formation (Dickerson et al., 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).
 
 

6. 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 ca 100-120 DU (Müller 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 have to persist long enough for the reactions which regenerate active chlorine species (Cl₂, Cl, ClO) from the inactive reservoir species (ClONO₂, HCl), and for HNO₃-rich PSC particles to settle out of the stratosphere. Once ozone loss has taken place, the lesser amounts of O₃ present result in less absorption of UV radiation in the stratosphere, and consequently to 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 last 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 previous 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 timescale 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 timescale 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 CO₂ 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 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 et 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, which leads to significantly colder temperatures existing over longer periods of time in the Arctic stratosphere. They estimate that due to this effect the ozone loss over the Arctic by the year 2020 will be double what it would be without greenhouse gas increases, and that recovery from Arctic ozone depletion will be delayed by some 10-15 years.

7. Conclusion

In the preceding sections, we have examined 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 just the emission of trace gases and aerosols, but that it is essential to consider the complex interconnections between any given perturbation and the overall Earth System.

There are a few cases, where the impact of anthropogenic emissions on the atmosphere is relatively easy to assess, for example those where the sources of a substance are industrial and its sinks are chemical reactions with first-order kinetics. But in most instances, the emission and removal of climatically active gases and aerosols depends 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 N₂O, 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 hydrological cycle. As a result of these chemical and physical perturbations, the chemical functioning of the atmosphere will be modified, and consequently the production rates and lifetimes of aerosols and greenhouse gases. 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 times scales involved in ecological change, biogeochemical cycles and climate have a memory of past land use change, and, conversely, present land use change may have long-term consequences reaching far into the future.

In some cases, the human perturbation consists of the release of a precursor compound (e.g., SO₂), which is transformed in the atmosphere to a climatically active substance. In this example, the actual amount of radiatively 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 processes.

At longer timescales, we must consider feedback loops where climate change results in different circulation patterns, changes in water availability at the surface, water vapor content of the atmosphere, etc. These factors in turn modify the sources, sinks and atmospheric budgets of trace gases and aerosols, again affecting climate. A dramatic example for this kind of interaction is the coupling between changes in stratospheric temperatures and ozone depletion, which 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 last decade by intensified cooperation between scientific disciplines, close interaction between observation and modeling, and broad international cooperation. We must, however, over our excitement with new conceptual insights into the complexity of the Earth System’s workings not lose sight of the fact that the observational data base for testing our concepts and models remains still 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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