Greenhouse: Coping with Climate Change

By CSIRO PUBLISHING

Consideration of climate change deals increasingly with impacts and responses, and therefore involves a wide range of technical issues and a diverse community of experts. One of the challenges faced is that of ensuring effective communication between these different areas of expertise. For example, climate change studies require new types of collaboration between carbon cycle modellers and economists, and between meteorologists and coastal geomorphologists. Furthermore, there is a need to distil balanced assessments ranging across many disciplines for the benefit of all policymakers.Greenhouse: Coping with Climate Change brings together the contributions of many experts to the climate change debate.

This book is a landmark publication summarising our understanding of climate change issues as they affect Oceania. It contains review papers that report on the status of knowledge, methodologies and developments; and a selection of focused papers that expand on specific issues and present significant new developments of wide general interest and relevance to the region.

• Language: English
• Publisher: CSIRO PUBLISHING
• Release date: Jan 1, 1996
• ISBN: 9780643105737

Book preview

SCIENCE

THE CHANGING COMPOSITION OF THE ATMOSPHERE

Martin R. Manning,¹ Graeme I. Pearman,² David M. Etheridge,²

Paul J. Fraser,² David C. Lowe¹ and L. Paul Steele²

¹  National Institute of Water and Atmospheric Research, Lower Hutt, New Zealand

²  CSIRO Division of Atmospheric Research, Aspendale, VIC 3195, Australia

Abstract

This paper reviews changes that have occurred in the atmospheric levels of major greenhouse gases and consequent changes to heating of the lower atmosphere through the enhanced greenhouse effect. The naturally occurring greenhouse gases carbon dioxide, methane, and nitrous oxide have increased by about 27%, 140% and 10%, respectively, over the last 300 years and over the last decade have increased at an average rate of 0.4, 0.7 and 0.3% y–1, respectively. The rapid increases in these gases that occurred after the 18th century have been due to changes in anthropogenic sources and sinks. Direct measurements in the atmosphere, and studies of source and sink processes, are now producing detailed budgets for these gases.

Several industrially produced greenhouse gases, such as chlorofluorocarbons (CFCs), are entirely anthropogenic in origin. Until recently CFCs were increasing rapidly in the atmosphere, however, emissions of CFC-11 and CFC-12 have been reduced because of adherence to the Montreal Protocol. Increases in their atmospheric concentrations have thus slowed and are expected to stabilize and then decline before the end of the century. Hydrochlorofluorocarbons (HCFCs), used as substitutes for CFCs, and some other man-made greenhouse gases are increasing rapidly in the atmosphere.

Ozone is a greenhouse gas which is produced by chemical reactions in the atmosphere. Emissions of CFCs have decreased ozone in the stratosphere, whereas emissions of nitrogen oxides and volatile organic compounds have increased ozone in the troposphere, at least in the northern hemisphere. Global atmospheric chemistry models indicate that changes in ozone over the last 200 years have probably contributed more to the enhanced greenhouse effect than changes in nitrous oxide or CFCs.

INTRODUCTION

Concerns about anthropogenic climate change arise because of observed changes in the composition of the atmosphere. Here we review current understanding of changes in greenhouse gases over the last 1000 years and the increase in heating of the atmosphere that this has caused. The other major change in atmospheric composition is an increase in airborne particulates (aerosols) which act to cool the atmosphere. Aerosols are discussed in more detail by Ayers and Boers (1996).

The major greenhouse gases are carbon dioxide (CO2), methane (CH4), nitrous oxide (N2O), chlorofluorocarbons (CFCs), and ozone (O3). Although these gases constitute less than 0.04% of dry air, they play a major role in climate because they absorb radiant heat emitted from the surface of the Earth. Atmospheric water vapour also makes a large contribution to the natural greenhouse effect.

Concerns about future climate change are based on predictions of how the climate system will respond to further increases in greenhouse gas concentrations. A consistent basis for such predictions must start with a scenario for future emissions of greenhouse gases and develop estimates of how atmospheric concentrations will change. This paper reviews the information available on changes in greenhouse gas concentrations and the present understanding of their sources and sinks, which underpin models used to predict future greenhouse gas concentrations.

Ice cores provide a reliable source of information regarding past atmospheric composition. Polar ice contains air bubbles which are formed at depths of about 70–100 m as the snow metamorphoses into solid ice. These bubbles generally store air intact within the ice. This air is representative of the ‘background’ contemporary atmosphere, and can be found with ages ranging from decades to hundreds of thousands of years. Recent improvements to the ice-core records have come from better dating, better air-age resolution, more comprehensive coverage of the industrial period and the late-Holocene, more precise air-extraction and analysis techniques, and more accurate calibration. These records of past atmospheric composition indicate the range of natural variability in greenhouse gases and clearly show the net effect of anthropogenic changes over the last 200 years.

The predominant greenhouse gases have residence times in the atmosphere of 10 years or more. This is long compared to the time of about 1 year required for gases to become mixed throughout the atmosphere. As a result they are fairly evenly dispersed and it is possible to use observations at representative sites to assess the total amount present in the atmosphere, and to track the net rate at which they are being added or removed.

Current understanding of the addition and removal budgets for the major greenhouse gases comes from two complementary approaches. First, precise measurements of atmospheric concentrations at a network of sites around the world (see Figure 1) enable estimates of the net changes in the global atmospheric inventory for a gas species. Such data also provide information on the spatial and temporal distribution of the fluxes to and from the atmosphere. Second, direct studies of the processes that release or remove a gas are extrapolated to estimate total fluxes for the planet. Agreement between these approaches is needed to form a basis for robust global budget estimates.

Greenhouse gases, like most chemical species, occur in several different isotopic forms that have the same chemical behaviour but have slightly different masses. Many of the processes that release or remove gases from the atmosphere differentiate slightly between the different isotopic forms. For example, photosynthetic uptake of atmospheric CO2 favours lighter molecules over heavier ones. Measurements of the ratios of the different isotopic species of a gas, and how these vary, thus provide additional information on release and removal processes.

Figure 1. Some of the key sites used for regular measurement of atmospheric composition in background air around the world are marked with a shaded circle.

CARBON DIOXIDE

Carbon dioxide plays a fundamental role in biological and geological processes and, because it is a dominant greenhouse gas, links these with climate.

The record of atmospheric CO2 for the last millennium is shown in Figure 2. The results are selected from Antarctic ice cores only (Neftel et al., 1985; Stauffer and Oeschger, 1985; Friedli et al., 1986; Etheridge et al., 1988; Siegenthaler et al., 1988). Some earlier records have not been included because of their lesser precision. Greenland records have not been included because of concerns that Greenland ice may not store CO2 reliably (Delmas, 1993).

The results in Figure 2 come from ice cores from 4 different sites, using 3 different air extraction and analysis techniques, and demonstrate a high degree of consistency between the measurements. Further confirmation that ice-core air is representative of the past atmosphere is the overlap of the ice-core CO2 concentrations with those from direct atmospheric measurements of the last several decades (e.g. at South Pole, Keeling, 1991).

The record shows that CO2 concentrations in the pre-industrial Holocene ranged from 276 to 286 ppmv, i.e. the current global mean concentration of 358 ppmv is about 27% higher than the mean pre-industrial level. There is also evidence of smaller, natural CO2 variations, including an increase of about 10 ppmv in the 13th century and a decrease around 1550–1800 AD, coinciding with a period of cooler temperatures often called the ‘Little Ice Age’ (Etheridge et al., 1995; Barnola et al., 1995). Although Antarctic ice cores can now be measured for CO2 with an accuracy of around 1 ppmv, the possibility of finding the inter-hemispheric difference in CO2 in pre-industrial times may be restricted by poorer precision of the Greenland ice-core results.

Direct measurements of CO2 concentrations in clean air have been made at an increasing number of sites since 1958. Concentrations are reported for most of the sites shown in Figure 1 and are available through the World Meteorological Organization Greenhouse Gases Data Centre or from the groups concerned. Data from sites operated by the US National Oceanic and Atmospheric Administration’s Climate Monitoring and Diagnostics Laboratory are shown in Figure 3 as a function of latitude for the period 1980–1993 (Conway et al., 1994).

Continuous records of CO2 concentration show two clear features: the steady increase of annual averaged concentration; and a seasonal cycle which varies with latitude. Isotopic measurements confirm that the large seasonal cycles observed in the northern hemisphere are due to CO2 exchange with the terrestrial biosphere. Concentrations are driven lower in the spring and summer by photosynthetic uptake of carbon during plant growth and subsequent respiration and decay of plant material leads to winter maxima.

Figure 2. Concentrations of atmospheric CO2 from ice core records and direct atmospheric measurements for the period 1000 AD to present Data from five different ice cores are shown with symbols as noted in the legend. The solid curve denotes in situ measurements at the South Pole.

The annual growth rate of atmospheric CO2 is quite variable. Conway et al. (1994) cite growth rates varying from 0.6 to 2.5 ppmv y–1 during the 1981–1992 period with an average of 1.43 ppmv y–1. These authors and others have shown evidence for a relationship between inter-annual variations in CO2 growth rates and the El Niño–Southern Oscillation (ENSO), although the mechanisms are not well understood. During 1992 and 1993, low CO2 growth rates have coincided with a prolonged ENSO anomaly and cooling of sea surface temperatures due to the eruption of Mt. Pinatubo. Thus, as 1992 climatic conditions were unusual, the CO2 growth rate of 1.51 ppmv y–1 over the 1981–1991 period (Conway et al., 1994) is probably the best indication of current average growth rates.

The steady increase in global annual average CO2 concentration implies that source fluxes exceed removal (or sink) fluxes. In addition, higher concentrations in the northern hemisphere indicate that excess sources are higher there than in the southern hemisphere. Ice core and direct measurements show that, over the last 200 years, the increase in atmospheric CO2 over pre-industrial levels has increased at the same rate as estimated cumulative emissions from fossil fuels and deforestation. For example, the atmospheric CO2 increase between 1860 and 1989 was equivalent to the accumulation of about 40% of estimated anthropogenic emissions over that period (Siegenthaler and Sarmiento, 1993). There is some evidence that the fraction of CO2 emissions accumulating in the atmosphere is increasing, as the atmospheric increase between 1980 and 1989 corresponded to about 46% of cumulative anthropogenic emissions (Siegenthaler and Sarmiento, 1993).

Figure 3. The global distribution of atmospheric CO2 concentrations as a function of latitude and time for the period 1981–1993 from measurements in the NOAA/CMDL global air sampling network (courtesy T. Conway, CMDL).

This correspondence in growth rates provides strong circumstantial evidence that increasing CO2 is due to anthropogenic emissions. Further evidence comes from isotopic measurements. The ratio of the carbon isotopes ¹³C and ¹²C is different in atmospheric CO2, CO2 derived from biospheric sources (including fossil fuels), and from oceanic sources. Both recent atmospheric measurements and measurements of air retrieved from ice cores show that the ¹³C/¹²C ratio in atmospheric CO2 has decreased over the last 200 years as would be expected if all the additional CO2 was from terrestrial vegetation and fossil fuels (Keeling et al., 1989).

Several authors have compiled information on CO2 sources and sinks taking account of both atmospheric observations and studies of emission processes. The most recent such review (Schimel et al., 1994) was carried out as part of an Intergovernmental Panel on Climate Change (IPCC) assessment, and produced the estimates of average exchanges between major carbon reservoirs for the decade 1980–1989 (Table 1).

The term labeled ‘additional uptake’ in this carbon cycle budget is derived by difference from the other terms which are estimated more directly. This sink of atmospheric CO2 is consistent with some estimates of the CO2 fertilization effect which causes plants to grow more rapidly in atmospheres containing higher CO2 concentrations. Deposition of nitrogen compounds from anthropogenic emissions may also have stimulated plant growth and additional carbon storage (Hudson et al., 1994).

There is evidence from analyses of the spatial distribution of CO2 concentrations for uptake of CO2 in the mid-latitudes of the northern hemisphere (Enting and Mansbridge, 1989). This can be attributed to a terrestrial vegetation sink by comparing the total sink with that expected for the oceans (Tans et al., 1990) and by analysis of variations in ¹³C/¹²C isotopic ratios (Ciais et al., 1994).

The ability of ¹³C/¹²C to differentiate between biospheric and oceanic exchange of CO2 with the atmosphere has also been used to analyse inter-annual variations in these (Francey et al., 1995; Keeling et al., 1995). It is clear that small fluctuations in either oceanic or biospheric exchange (driven, for example. by climate variations) can produce inter-annual variations in the net effect of these natural sources or sinks which can explain observed variability in atmospheric CO2 growth rates.

METHANE

The record of methane (CH4) concentration over the last millennium has recently been improved by results from 3 ice cores, shown in Figure 4. The Antarctic ice core DE08 provides air from the period 1841–1978 with high age resolution (Etheridge et al., 1992). An approximate doubling of CH4 concentration over the industrial period is evident, consistent with previous ice-core studies. The CH4 growth rate stabilized during 1920–1945 and possibly during the 1970s. These previous events appear similar to a recent slowing of growth rate during the 1980s (Steele et al., 1992). The most recent ice core measurements of CH4 concentration overlap the direct atmospheric measurements in the 1970s and confirms that ice core air correctly reflects the past atmosphere.

Table 1. Estimated average sources and sinks of atmospheric CO2 for the 1980s and the related change in the atmospheric carbon inventory (Schimel et al., 1994). (1 Pg C = 10¹⁵g Carbon)

Figure 4. Concentrations of atmospheric CH4 from ice core records and direct atmospheric measurements for the period 1000 AD to present. Data from 5 different ice cores are shown with symbols as noted in the legend. The solid curve denotes in situ measurements at Cape Grim, Tasmania.

Figure 5. The global distribution of atmospheric CH4 concentrations as a function of latitude and time for the period 1983–1994 from measurements in the NOAA/CMDL global air sampling network (courtesy E. Dlugokencky, CMDL).

In Greenland ice CH4 is stored more reliably than CO2. Blunier et al. (1993) used two techniques of air extraction and analysis on an ice core from Summit, Greenland, to produce a record covering the last 1000 years. The pre-industrial level (700 ppbv) was found to be similar to that of other studies. Variations in concentration of about 70 ppbv were detected before 1500 AD, which the authors attributed to changed oxidizing capacity of the atmosphere, climatic impacts on wetlands and possible early changes in agricultural sources.

Nakazawa et al. (1993) analysed ice cores from Greenland and Antarctica to compare atmospheric CH4 concentrations in each hemisphere from pre-industrial times to about 40 years ago. The pre-industrial inter-hemispheric difference was found to be 55 ± 20 ppbv, about one-third of the present day difference, with higher concentrations in the northern hemisphere suggesting stronger natural CH4 sources there.

Methane is currently measured at many of the sites shown in Figure 1. Figure 5 shows the variation of CH4 in clean air as a function of latitude and time, based on sites operated by the US National Oceanic and Atmospheric Administration’s Climate Monitoring and Diagnostics Laboratory for the period 1983–1994. As in the case of CO2, an increasing trend and seasonal cycle are evident at all locations.

During the 1983–92 period, CH4 accumulated in the atmosphere at a globally averaged rate of 11.1 ± 0.2 ppbv y–1, a relative rate of increase of 0.66% y–1 (Dlugokencky et al., 1994b). However, there was a downward trend in CH4 growth rates from 13.5 ppbv y–1 in 1983 to 9.3 ppbv y–1 in 1991. The ice-core data suggest that this slowing of the atmospheric CH4 growth rate has been occurring for about 20 years, with a maximum in the growth rate in the early 1970s (Etheridge et al., 1992).

Most sources of CH4 can be classified into biogenic (rice paddies, wetlands, ruminant animals, and decomposition of organic material in landfills); fossil fuel, related to coal and gas production and use; and biomass burning (e.g. Fung et al., 1991). The last two are essentially anthropogenic whereas around 40% of biogenic sources appear to be natural in origin.

Biogenic CH4 is produced by bacterial communities which derive energy from organic materials such as acetates and release CH4 as a waste product. Fossil fuel emissions of CH4 are associated with release of gas trapped in ore bodies during coal extraction and leakage or venting during the recovery of oil and natural gas. Biomass burning, much of it in connection with tropical agricultural practices, leads to emission of CH4 from incomplete combustion.

The ¹³C/¹²C isotope ratio of CH4 differs between source types. Biogenic sources have ¹³C/¹²C ratios at least 1% less than atmospheric CH4, while fossil fuel and biomass burning sources have ratios about 1% greater. The isotope ratio of atmospheric CH4 is determined by the relative combination of different sources and so its measurement places constraints on the actual source mix.

The weakly radioactive ¹⁴C isotope occurs naturally in biogenic sources of CH4 but, as ¹⁴C decays with a half life of 5730 years, it is absent in fossil CH4 sources. Wahlen et al. (1989), Manning et al. (1990), and Quay et al. (1991) have used ¹⁴C measurements to show that approximately 20% of current atmospheric CH4 is derived from fossil sources.

The principal removal mechanism for atmospheric CH4 is a slow oxidation process in the atmosphere producing CO2. A lesser removal mechanism, accounting for about 10% of the total removal, is through consumption by methanotrophic bacterial communities in soils. Both removal processes favour the lighter ¹²CH4 isotopic species and cause a relative enrichment of the heavier ¹³CH4 in the atmosphere. The best estimates of methane removal rates come from analysis of atmospheric methyl chloroform levels (Prinn et al., 1995). Methyl chloroform is entirely anthropogenic in origin, so its sources are relatively well known, and its removal is by atmospheric chemistry closely related to that for CH4.

The most systematic attempt to determine the global CH4 budget to date is that by Fung et al. (1991). Prather et al. (1995) reviewed the global CH4 budget in a recent IPCC assessment, but do not attempt to simultaneously meet all the constraints imposed by atmospheric observations. Subak et al. (1993) have compiled estimates of anthropogenic emissions using methodologies developed for reporting national greenhouse gas emission inventories and provide a detailed analysis for the year 1988.

Table 2 compares estimates of the major sources of CH4 for the 1980s from these three works. During this period sources exceeded sinks by about 30 Tg CH4 y–1 (1 Tg CH4 = 10¹² g of CH4). The total amount emitted to the atmosphere each year is known to within about 20%, but the partitioning among the known CH4 sources is less well known.

Table 2. Estimates of major source and sink processes for atmospheric CH4 in the 1980s given in Tg CH4 y–1

In 1992 there was a dramatic drop in the CH4 growth rate followed by a recovery in 1993. Dlugokencky et al. (1994a) report an increase in global average CH4 values of only 4.7 ppbv during the calendar year 1992. Analysis of the spatial and temporal variations in growth rates shows that these were lowest in late 1992 and in the high northern latitudes. Some Arctic sites had lower concentrations during 1992 than in the corresponding month of 1991. More recently, the December 1993 global average CH4 concentration has been estimated as about 7 ppbv above that of December 1992 (E. Dlugokencky, pers. comm.). This suggests that the 1992 anomaly is a feature separate from the overall decline in growth rates.

Isotopic ¹³C measurements (Lowe et al., 1994) in the southern hemisphere show a shift towards isotopically lighter CH4 during 1992–93 (Figure 6). This indicates that the low growth rates in this period are not due to an increase in the removal rate, as that would have increased atmospheric ¹³C/¹²C ratios. A reduction in isotopically heavy CH4 sources, such as those associated with fossil fuels and biomass burning, is indicated.

Dlugokencky et al. (1994a) suggested that changes in the former Soviet Union may have led to a significant reduction in coal and gas production and distribution. However, Fridman et al. (1994) have reported total CH4 emissions from Russia for the period 1990–1993 giving a 9% reduction over that period. This would account for only a small part of the observed decrease in growth rate.

An alternative hypothesis (Lowe et al., 1994; Lassey et al., 1996) is that drought conditions since 1991 in Southern Africa and America may have resulted in a reduction of plant material subjected to agricultural burning and hence to lower CH4 emissions from this source. Some additional support for lower biomass burning in 1992 comes from the weaker seasonal cycle in ¹³CH4 apparent in Figure 6, which is presumed to be due to an influx of isotopically heavy CH4 to the southern hemisphere following burning in the tropical dry season of July–September.

Figure 6. Measurements of ¹³C in atmospheric CH4 in the southern hemisphere showing a seasonal cycle and trend towards isotopically lighter CH4 in 1992. The ¹³C/¹²C ratio is given by δ¹³C which is the relative difference in that ratio between the sample and a standard reference material exressed in (10 = 1 ). Small circles with error bars denote data from Baring Head, New Zealand, and intermediate sized circles denote data from Scott Base, Antarctica, Error bars on the latter are similar in size to those for Baring Head. The large circles connected by a line, and the open squares denote measurements of two surveillance standards used to check against drift in the measurements.

NITROUS OXIDE

Nitrous oxide (N2O) is a naturally occurring greenhouse gas which is released from soils and water as a product of nitrification and de-nitrification processes. Although it is only present in the atmosphere in low concentrations it has a long lifetime of about 120 years and is a strong absorber of infra-red radiation making it a significant greenhouse gas.

Analysis of ice-core records shows that atmospheric N2O increased by about 9%, from about 285 ppbv in the 1000 years before industrialization to 310 ppbv in 1990 (e.g. Etheridge et al., 1988; Zardini et al., 1989).

Results from two Byrd Station ice cores and a Dye 3 ice core from Greenland (Leuenberger and Siegenthaler, 1992) were obtained by mass spectrometer rather than the previous gas chromatographic techniques. The Dye 3 results cover the period 1781–1947 and suggest a pre-industrial N2O level of about 260 ppbv, 10–25 ppbv lower than other records. Comparisons with modern air indicate that the mass spectrometer technique is unlikely to be the cause of the discrepancy. The N2O concentrations in the Byrd core samples extend back to the early Holocene and are closer to the levels found in previous records.

New N2O results from an Antarctic ice core have also been presented by Machida et al. (1994). Pre-industrial concentrations, measured by gas chromatography, were found to be in the range 273–280 ppbv. This is midway between the previously mentioned results.

Terrestrial ecosystem sources of N2O are difficult to characterize because they are highly variable in space and time. Denitrification processes occur in anaerobic conditions, including anaerobic microsites within aerobic soils, and generally result from bacteria using nitrate as a source of oxygen when molecular oxygen is unavailable. Production rates are dependent on both current levels and the recent history of soil oxygen, moisture and temperature. Forest soils, particularly wet soils in the tropics, are the dominant natural source (Keller et al., 1986). Grasslands and savannahs are important secondary sources (Sanhueza et al., 1990). Some estimates of oceanic sources (e.g. Codispoti et al. 1992) suggest that these may also be significant.

Prior to 1988 it was generally believed that industrial combustion processes comprised a large anthropogenic source of N2O. This was shown to be incorrect when it was discovered that N2O was being generated as an artifact of sampling procedures (Linak et al. 1990). It is now believed that there is a large number of anthropogenic sources including use of fertilizers, decomposition of animal wastes, production of industrial chemicals, biomass burning and land-use change (e.g. Khalil and Rasmussen, 1992).

The recent atmospheric record (Prather et al., 1995) shows a steady increase of about 0.2 % y–l (0.6 ppbv y–l) over the last 15 years. Growth rates appear to have been slightly higher than average during the late 1980s. This long-term trend implies that sources exceed sinks by about 3.9 Tg N y–1. Table 3 compares estimates of major sources from Khalil and Rasmussen (1992) and the recent IPCC assessment of trace gas budgets (Prather et al., 1995). Note the small discrepancy between the total of identified sources and the sum of estimated atmospheric removal and accumulation.

Table 3. Estimated sources and sinks of atmospheric N2O (Tg N y–l)

CHLOROCARBONS

Anthropogenic chlorocarbon species have contributed significantly to the growing levels of chlorine in the atmosphere, and hence to global stratospheric ozone depletion. Chlorocarbons include the CFCs CFC-11 (CCl3F), CFC-12 (CCl2F2), CFC-113 (CCl2FCCIF2), HCFC-22 (CHClF2), and methyl chloroform (CH3CCl3) and carbon tetrachloride (CCl4). They are powerful greenhouse gases because of their relatively long atmospheric lifetimes (5–100 years) and strong absorption of infrared radiation.

Because CFC-11, -12, -113, HCFC-22 and CH3CCl3 do not occur naturally their global emissions can be tracked from industry production figures. These are shown in Figure 7 (Fisher et al., 1994). The emissions of all species except HCFC-22 are now in decline due to the apparent success of the Montreal Protocol. Under the Protocol, emissions of HCFCs are allowed to increase rapidly over the next few years to in excess of 500 Gg y–1 (1 Gg = 10⁹ g) if required, because this species is an interim substitute for CFCs. However, HCFC emissions are to be reduced to about 280 Gg y–1 by 2004 with substantial further reductions over the following decade.

Figure 7. The global emissions (kt) of, and Monstreal protocol (Copenhagen 1992 Amendments) target (…) for, CFC-11, -12, -113, HCFC-22, CCl4 and CH3CCl3 (Simmonds et al., 1988; Fisher et al., 1994). Global CCl4 emissions (1986–1992) are assumed to be 12% of the sum of global CFC-11 and -12 emissions.

Emissions of CFC-11 and -12, largely from foam and refrigeration losses, peaked in 1988 at about 350 and 460 Gg and by 1992 had declined by about 40 and 30%, respectively. CFC-113 emissions, largely from its use as a solvent in the electronics industry, peaked in 1989 at 270 Gg and by 1992 had declined by 35%. Emissions of CH3CCl3, a metal cleaning solvent, peaked in 1990 at 720 Gg and by 1993 had declined by 60%, whereas emissions of HCFC-22 continue to grow monotonically, having reached about 215 Gg by 1992.

Emissions of CFC, CCl4 and CH3CCl3 appear to be on target with respect to the requirements of the Montreal Protocol, although some applications of these chemicals are more easily switched to ozone-benign alternatives than others and, therefore, targets may become increasingly difficult to achieve. Final emission targets for the developed world are close to zero while developing countries have a 10 year ‘period of grace’ to reduce emissions.

Concentration measurements of CFCs, CCl4 and CH3CCl3 have been made by in situ gas chromatography in the background atmosphere at sites around the world including Cape Grim, Tasmania, and by analysis of air collected since 1978 and stored in the Cape Grim air archive (Simmonds et al., 1988; Prinn et al., 1992; Fraser and Derek, 1994; Fraser et al., 1994a; Cunnold et al., 1994). Measurements of HCFC-22 using gas chromatography–mass spectrometry techniques have been carried out at the University of East Anglia on air from the Cape Grim air archive, as well as on flask air samples collected at Cape Grim (Montzka et al., 1993; Fraser et al., 1994b; Oram et al., 1994).

Data for the CFCs, CCl4 and CH3CCl3 are shown in Figure 8. The 1993 annual mean concentrations and recent growth rates are shown in Table 4. The HCFC-22 growth rates are only approximate, and the true CFC-12 growth rates for 1992–93 and 1993–94 may be larger and smaller, respectively, than indicated in Table 4, because of calibration problems in 1993, which are still being investigated.

Concentrations of CFC-12 and -113 in the atmosphere are still growing, but at rates that are now about 50% lower than the long-term growth rates observed before the introduction of the Montreal Protocol. CFC-11 and CCl4 have stopped growing, reaching maximum levels of about 262 and 132 pptv in early 1994 and 1990, respectively. Maximum CH3CCl3 levels of about 145 pptv were observed in mid-1992 and are now falling at 4–5% y–1. The HCFC-22 growth rate is accelerating, currently being 50% higher than prior to the Montreal Protocol.

Table 4. Chlorocarbon concentrations (pptv) and growth rates (pptv y–1) observed at Cape Grim, Tasmania (Fraser and Derek, 1994; Oram et al., 1994). Recent data are still provisional and subject to revision.

Figure 8. Monthly mean CFC, CCl4 and CH3CCl³ concentrations observed at Cape Grim, Tasmania. Data affected by local and regional pollution episodes are not included in the monthly means (Simmonds et al., 1988; Prinn et al., 1992; Fraser and Derek, 1994; Fraser et at., 1994a; Cunnold et al., 1994; P. Fraser, unpublished data).

The growth rate patterns exhibited by the CFCs, HCFC-22 and CH³CCl³ are broadly consistent with the emissions shown in Figure 8 (Cunnold et al., 1994; Fraser et al., 1994b; Oram et al., 1994; Prinn et al., 1992, 1995). There are no reliable emission data for CCl4 (Simmonds et al., 1988), but the almost stationary concentrations observed over the past 3–4 years suggest that the industrial emissions are static and probably close in magnitude to the atmospheric sink for CCl4.

OZONE

Ozone (O3) is a reactive oxygen species formed by chemical reactions in the atmosphere. It is generally short lived (hours to weeks depending on location) and its levels are the result of a dynamic equilibrium between generation and removal processes. Several different ozone regimes exist in the atmosphere depending on the concentrations of chemical precursors and the levels of photochemically active radiation (Volz-Thomas and Ridley, 1994).

In the stratosphere, at altitudes above 12 km, UV radiation is the dominant factor. This causes photo-dissociation of both molecular oxygen (O2) and ozone (O3) producing highly reactive atomic oxygen (O) which then recombines with the parent species. The result is a dynamic mix of O, O2 and O3. About 90% of the total atmospheric ozone is in the so called ‘ozone layer’ in the stratosphere. Absorption of UV radiation there prevents most UV radiation reaching the surface of the Earth where it can cause biological damage. Over the last few decades large increases in chlorine levels in the stratosphere, due to the photo-dissociation of CFCs, have perturbed natural ozone chemistry and introduced catalytic removal processes which have depleted stratospheric ozone levels.

Exchange of air between the stratosphere and the troposphere brings ozone rich air to lower altitudes. However, other processes remove and generate ozone directly in the troposphere. Tropospheric ozone chemistry is highly dependent on levels of the reactive oxides of nitrogen, NO and NO2 collectively referred to as NOx. The ‘clean remote troposphere’, e.g. marine air in the southern hemisphere, is characterized by low NOx levels. In such circumstances, photodissociation of ozone and reactions with hydroxyl and peroxy radicals lead to net removal of ozone (Ayers et al., 1992). Ozone is also removed at soil and water surfaces (Galbally and Roy, 1980).

In areas affected by industrial pollution, elevated levels of NOx and non-methane hydrocarbons cause other chemical reactions, sometimes referred to as smog chemistry, which result in the formation of ozone. The most widespread effect is that, at higher NOx levels, peroxy radicals react with NO rather than ozone and catalytic generation of ozone occurs. Through most of the troposphere NOx is the rate-limiting ozone precursor.

Stratospheric ozone levels are inferred from total column measurements based on observing absorption of solar or lunar radiation at the surface of the Earth, and from satellite, balloon or aircraft-borne instruments. Extensive analysis has been carried out of recent trends in stratospheric ozone (Albritton et al., 1994) showing that overall levels have declined by a few percent. In addition, the Antarctic ozone hole and large depletions over the Arctic are known to be due to the chemical effects of excess chlorine and stratospheric aerosols at very low temperatures (Albritton et al., 1994).

Tropospheric ozone has been measured at a few surface sites, mainly in the northern hemisphere. Because ozone concentrations are highly variable in space and time, available data do not enable global tropospheric ozone budgets to be calculated with confidence. Some analyses (Oltmans and Levy, 1994; Albritton et al., 1994) indicate that tropospheric ozone has increased in regions which are subject to increasing industrial pollution. Longer term measurements suggest that tropospheric ozone may have more than doubled in Europe over the last 80 years (Marenco et al., 1994). On the other hand, recent reductions in pollutant emissions, particularly in Europe and North America, are expected to decrease tropospheric ozone levels in those regions.

Considerable progress has been made over the last decade in constructing global-scale models of atmospheric chemistry. These currently suggest that changes in ozone precursors since pre-industrial time should be responsible for about a doubling of tropospheric ozone in the northern hemisphere (e.g. Hauglustaine et al., 1994).

RADIATIVE FORCING

Increases in greenhouse gas levels lead to an enhanced greenhouse effect and more heating of the Earth’s surface and lower atmosphere. Direct solar heating of the Earth’s surface is about 240 W m–2, which would cause an average surface temperature of about –18°C. However, the insulating effect of greenhouse gases and water vapour, i.e. the re-radiation of the energy they absorb, is equivalent to an additional heating of about 150 W m–2. This raises the surface temperature by about 33°C (Ackerman, 1992).

Radiative forcing is defined as the change in the net balance of radiation at the top of the troposphere due to a change in atmospheric composition. It is a measure of the perturbation to the Earth’s heat budget given as a global average in W m–2. The radiative forcing due to a change in the concentration of a greenhouse gas is calculated from a detailed knowledge of the radiative absorption properties of the gas. These calculations generally use 1-dimensional radiative convective models of the atmosphere which keep other relevant properties of the atmosphere fixed.

A doubling of CO2 concentrations would lead to a radiative forcing of 4.4 W m–2. On the basis of simple energy balance considerations this would lead directly to an increase in surface temperature of about 1.2°C (Ackerman, 1992). However, an increased surface temperature would lead to higher atmospheric water vapour levels and hence to further absorption of infra-red radiation. When such feedback effects are calculated, using general circulation models of the atmosphere and interactions with other components of the climate system, it is found that doubling of CO2 concentrations leads to an increase in global average surface temperature of between 1.5 and 4.5°C (Gates et al., 1992).

The increase in greenhouse gases that has occurred from pre-industrial times to the present is calculated to have increased surface heating by about 2.7 W m–2 (Shine et al., 1995). Simultaneous changes in aerosol concentrations are expected to have reduced heating of the lower atmosphere (Charlson et al., 1991). The role of aerosols is covered in more detail by Ayers and Boers (1996).

In considering strategies to mitigate climate change due to greenhouse gas emissions, a useful construct is the Global Warming Potential (GWP). This compares the radiative forcing resulting from emission of 1 kg of a gas to that from emission of 1 kg of CO2, integrated over a specified future time period. Albritton et al. (1995), in a recent IPCC assessment, have summarized recent estimates of GWPs for most greenhouse gases.

Global Warming Potentials take into account the differing atmospheric lifetimes and abilities of various gases to absorb radiation. The direct component of the GWP is based on the absorption of radiation over the lifetime of the gas. For several gases, an indirect component arises from additional effects on the atmospheric radiation balance, particularly through consequent changes to the levels of other greenhouse gases. Thus increases in CH4 tend to decrease the oxidizing power of the atmosphere and hence the removal rate of CH4 and several other greenhouse gases. In addition, the interaction of CH4 and NOx leads to the formation of ozone which, as noted above, is a significant greenhouse gas. These chemical effects mean that emissions of CH4 cause more radiative forcing than that caused by the direct effect.

The GWP concept strictly applies only to gases whose emissions become fairly evenly dispersed around the globe. This does not apply to short-lived gases such as NOx. As already noted, NOx is the rate-limiting precursor for ozone formation, and increases in NOx emissions lead to increased ozone and radiative forcing (Jacob et al., 1993). However, a GWP is not defined for NOx because its short lived nature makes any global-scale change in concentration difficult to characterize. Given that ozone produced by the interaction of CH4 and NOx is included in the CH4 GWP, there is clearly some inconsistency in the present methodology for assessing radiative forcing due to trace gas emissions.

Indirect GWP effects have so far been limited to atmospheric chemistry and do not include biospheric or other feedbacks on climate. A notable exception is that an estimate of the ‘CO2 fertilisation’ effect on the terrestrial biosphere is inherently included in calculating CO2 removal from the atmosphere. Similar fertilisation effects due to emissions of nitrogen species are not included. Thus the GWP concept can be criticised as neither an entirely consistent nor a comprehensive approach to determining the relative impacts of different greenhouse gases. However, at present there appears to be no better alternative for comparing greenhouse gas emissions.

Direct and total GWPs, including indirect effects, taken from the 1994 WMO-UNEP review of stratospheric ozone (Solomon et al., 1994) and the IPCC 1994 report (Albritton et al., 1995) are shown in Table 5 and compared with earlier IPCC estimates. Some of these GWPs, particularly that for CH4, are expected to be reduced when a recent revision of atmospheric oxidation rates (Prinn et al., 1995) is incorporated.

The direct component of chlorocarbon GWPs is positive and large due to their high infrared absorption properties. Indirect effects of chlorocarbon emissions include the cooling effect of depletion of stratospheric ozone, and increases in tropospheric oxidation rates due to enhanced UV radiation. Hauglustaine et al., (1994) have suggested that further feedback effects of ozone depletion on the radiation balance need to be taken into account. In general, however, indirect effects of ozone-depleting gases reduce and can even change the sign of their direct GWPs.

Despite the large GWPs of many greenhouse gases, the greatest changes in greenhouse warming are due to emissions of CO2. This is because of the much larger amounts of CO2 being emitted and accumulating in the atmosphere. Figure 9 shows changes in radiative forcing due to the accumulated change in various greenhouse gases since the pre-industrial era. These are based on the IPCC 1994 assessment (Shine et al., 1995) which gives forcing due to individual greenhouse gases as 1.56 W m–2 for CO2, 0.47 W m–2 for CH4, 0.14 W m–2 for N²O, 0.06 for CFC-11, and 0.14 for CFC-12. Radiative forcing due to other gases was estimated at about 0.08 W m–2.

Table 5. Global Warming Potentials (direct and net) for the major greenhouse gases referenced to CO2 over a 100-year time horizon. Adapted from Albritton et al. (1995) and Solomon et al. (1994). Uncertainties in the direct GWPs are typically 35%.

Figure 9. Relative contributions of various gases to the global net radiative forcing at the tropopause due to changes in concentrations from pre-industrial times to the present (1990). Note that the CFC contribution does not include effects of HCFC-22, CCl4 and CH3CCl3 emissions, but does include the indirect effects of ozone depletion (adapted from Shine et al., 1995; and Hauglustaine et al., 1994).

Figure 9 also shows IPCC 1994 estimates of other climate forcing changes since the pre-industrial era (Shine et al., 1995). In particular, it is estimated that changes in tropospheric ozone have caused a forcing of between 0.2 and 0.6 W m–2. Significant negative forcing (i.e. cooling) is attributed to the effects of increasing aerosol levels. The direct effect of aerosol scattering of incoming solar radiation is estimated to have caused a negative forcing of between –0.25 and –0.9 W m–2. Furthermore, increases in aerosol levels can affect the reflectivity of clouds (the indirect effect) and an additional, but highly uncertain, negative forcing due to this is estimated to be in the range 0 to –1.5 W m–2. The simultaneous effects of change in solar output over the period are also uncertain, but estimated to lie in the range 0.1 to 0.5 W m–2.

DISCUSSION AND ATMOSPHERIC STABILIZATION

Increases in anthropogenic sources of greenhouse gases has led to significant increases in their atmospheric concentration. Records of atmospheric concentrations from air bubbles trapped in ice cores, which overlap and agree with recent direct measurements, provide a detailed picture of major changes in atmospheric composition since pre-industrial times.

Through an international network of atmospheric observation sites and measurement of variations in concentrations and isotopic ratios, much is known about the sources and global scale budgets of the major greenhouse gases. Current long-term trends based on observations and current budget estimates are fairly well defined. However, recent anomalies in the growth rates of CO2 and CH4 show that we still do not have sufficient understanding to predict inter-annual variations in growth rates.

The increase in greenhouse gases, including ozone changes caused by other pollutants, is estimated to have added about 2.7 W m–2 to the greenhouse effect warming of the Earth (Shine et al., 1995). At the same time, increased aerosols are expected to have caused a cooling effect. The magnitude of this is uncertain with estimates ranging from –0.3 to –3.0 W m–2.

Stabilization of atmospheric greenhouse gas levels is an expressed goal of the Framework Convention on Climate Change (FCCC). Although the FCCC does not cover CFCs, their growth in the atmosphere has stopped, or is expected to shortly, in response to greatly curtailed emissions under the Montreal Protocol. Once emissions of CFCs are less than their atmospheric loss by photolysis, their concentrations will start to decline with 50–100 year time constants.

Current CH4 sources are only slightly larger than their sinks. A reduction of about 10% in anthropogenic sources would lead to stabilization at concentrations close to present levels. Recent fluctuations in the CH4 growth rate suggest that extreme year to year variations in sources are of this magnitude. The budget for N2O is not well understood, but a reduction of about 30% in anthropogenic sources would probably stabilize this gas near present levels. Alternatively, if current emissions are kept constant the atmospheric concentration is expected to stabilize at levels about 30% higher than present ones.

The removal of CO2 from the atmosphere occurs by redistribution of carbon to other reservoirs of the global carbon cycle. A discussion of global carbon cycle models and their role in estimating future climate is given by Schimel et al. (1994). Stabilizing CO2 concentrations at various levels has been studied in some detail as part of a recent IPCC assessment (Schimel et al., 1995). Although the details of these studies are beyond the scope of this paper, it is clear that stabilization of CO2 at about twice pre-industrial levels would require eventual reduction of CO2 emissions to less than 30% of current emissions.

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CLIMATE, CLOUDS AND THE SULFUR CYCLE

Greg P. Ayers and Reinout Boers

CSIRO Division of Atmospheric Research, Aspendale,

VIC 3195, Australia

Abstract

One of the major constraints on present efforts to model the climatic effects of increases in greenhouse gas concentration over the last few centuries is uncertainty about the role played by atmospheric aerosol, especially sulfate particles, and by carbon-containing and other atmospheric particles. Recent analyses suggest that in the northern hemisphere the combined ‘direct’ and ‘indirect’ (mediated through cloud reflectivity) effects of anthropogenic sulfate aerosols may have had a cooling effect approximately equal to the warming effect of increased atmospheric carbon dioxide concentrations since pre-industrial times. Clearly this possibility is of immense significance. In this paper, current understanding of the ‘direct’ and ‘indirect’ roles of atmospheric sulfate particles and clouds is reviewed, and other second-order climatic effects via aerosol-induced modification of cloud microphysical properties are discussed.

INTRODUCTION

Aerosol particles