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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties 2 State of Scientific Understanding In this chapter the state of understanding of radiative forcing from individual agents is reviewed. Over the past 15 years, the Intergovernmental Panel on Climate Change (IPCC) has produced assessments in which at least one chapter has been devoted to a thorough review of current understanding about radiative forcings. The discussions here summarize the findings of the IPCC’s Third Assessment Report (IPCC, 2001) and scientific advances since it was published. The Third Assessment Report (IPCC, 2001) includes a summary figure of the global and annual mean radiative forcings from 1750 to 2000 due to a range of perturbations (Figure 2-1), including the well-mixed greenhouse gases, ozone, aerosols, aviation effects on clouds, land use, and the Sun. The largest positive forcing (warming) since 1750 is associated with the increase of the well-mixed greenhouse gases (carbon dioxide [CO2]; nitrous oxide [N2O]; methane [CH4]; and chlorofluorocarbons [CFCs]) and amounts to 2.4 W m−2. The greatest uncertainty in Figure 2-1 is associated with the direct and indirect radiative effects of aerosols. If the actual negative forcing from aerosols were at the high end (most negative) of the uncertainty range, then it would have offset essentially all of the positive forcing due to greenhouse gases (see also Boucher and Haywood, 2001). According to the IPCC definition, applied to the data in Figure 2-1, “The radiative forcing of the surface-troposphere system due to the perturbation in or the introduction of an agent is the change in net irradiance at the tropopause after allowing for stratospheric temperatures to readjust to radiative equilibrium, but with the surface and tropospheric temperatures and state held fixed at the unperturbed values.” This definition of forcing is
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties FIGURE 2-1 Estimated radiative forcings since preindustrial times for the Earth and troposphere system (TOA radiative forcing with adjusted stratospheric temperatures). The height of the rectangular bar denotes a central or best estimate of the forcing, while each vertical line is an estimate of the uncertainty range associated with the forcing, guided by the spread in the published record and physical understanding, and with no statistical connotation. Each forcing agent is associated with a level of scientific understanding, which is based on an assessment of the nature of assumptions involved, the uncertainties prevailing about the processes that govern the forcing, and the resulting confidence in the numerical values of the estimate. On the vertical axis, the direction of expected surface temperature change due to each radiative forcing is indicated by the labels “warming” and “cooling.” SOURCE: IPCC (2001). restricted to changes in the radiation balance of the Earth-troposphere system imposed by external factors, with no changes in stratospheric dynamics, without any surface and tropospheric feedbacks in operation, and with no dynamically induced changes in the amount and distribution of atmospheric water. A somewhat broader perspective is applied in this chapter to include, in particular, volcanic aerosols, the effects of land-use changes and aerosols on precipitation, and the radiative forcing due to changes in ocean color.
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties Figure 2-1 has been an effective way to portray the relative magnitudes of different radiative forcings, the associated scientific uncertainties, and an assessment of the current level of understanding. It has been used widely in the scientific and policy communities. However, it has some important limitations, including the following: The figure does not provide information about the timescales over which each of the forcings is active. For example, the greenhouse gases in the first bar (CO2, CH4, N2O, and halocarbons) remain in the atmosphere for decades or longer, whereas the various aerosols persist for days to weeks. The figure shows globally averaged forcings and therefore does not provide information about regional variation in forcing or vertical partitioning of forcing. The figure does not provide information about other climate effects of each forcing agent, such as impacts on the hydrological cycle. The figure gives the impression that one can simply sum the bars to determine an overall or net radiative forcing; however, such a calculation does not give a reasonable description of the cumulative effect of all the forcings. The uncertainty ranges are generally estimated from the range of published values and cannot be readily combined to determine a cumulative uncertainty. The figure does not consistently indicate the forcing associated with specific sources (e.g., coal, gas, agricultural practices). The figure omits nonradiative forcings as discussed in this report. Although it would be unrealistic to expect a single figure to fully portray all of these aspects of radiative forcings, there are clearly opportunities to improve upon Figure 2-1 and to introduce new figures that address these limitations in the next IPCC report. WELL-MIXED GREENHOUSE GASES The radiative forcing due to CO2, CH4, N2O, and various halocarbons is due to absorption of infrared (IR) radiation. It is well characterized and well understood. These gases remain in the atmosphere long enough to be well mixed; thus, their abundances are well known and have little spatial variability. Their concentrations have increased substantially since preindustrial times (see Table 2-1), and they are the greatest contributors to total anthropogenic radiative forcing. As shown in Figure 2-1, the IPCC estimate of the radiative forcing due to well-mixed greenhouse gases is +2.43 W m−2 from 1750 to 1998 (present), comprising CO2 (1.46 W m−2),
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties TABLE 2-1 Well-Mixed Greenhouse Gases CO2 CH4 N2O CFC-11 HFC-23 CF4 Preindustrial concentration ~280 ppm ~700 ppb ~270 ppb 0 0 40 ppt Concentration in 1998 365 ppm 1745 ppb 314 ppb 268 ppt 14 ppt 80 ppt Rate of concentration changea 1.5 ppm/yrb 7.0 ppb/yrb 0.8 ppb/yr −1.4 ppt/yr 0.55 ppt/yr 1 ppt/yr Atmospheric lifetime 5 to 200 yrc 12 yrd 114 yrd 45 yr 260 yr >50,000 yr NOTE: CF4 = perfluoromethane; CFC-11 = chlorofluorocarbon-11; HFC-23 = hydrofluorocarbon-23; ppm = parts per million; ppb = parts per billion; ppt = parts per trillion. a Rate is calculated over the period 1990 to 1999. b Rate fluctuated between 0.9 and 2.8 ppm yr-1 for CO2 and between 0 and 13 ppb yr-1 for CH4 over the period 1990 to 1999. c No single lifetime can be defined for CO2 because of coupling with surface reservoirs. d This lifetime has been defined as an “adjustment time” that takes into account the indirect effect of the gas on its own residence time. SOURCE: IPCC (2001).
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties CH4 (0.48 W m−2), N2O (0.15 W m−2), and halocarbons (0.34 W m−2) (IPCC, 2001). The estimated uncertainty associated with this forcing is 10 percent, with that for CO2 and N2O being less and that for the other gases being greater. The estimated uncertainty for halocarbons is 10-15 percent for those molecules that have been studied in detail and is not well characterized for other halocarbons. Recent research on well-mixed greenhouse gases focuses on refining the models used to do radiative transfer calculations (e.g., Evans and Puckrin, 1999), considering the small temporal and spatial variations in concentrations which can lead to errors up to about 5-10 percent (Forster et al., 1997; Myhre and Stordal, 1997; Freckleton et al., 1998), and accounting for the extent to which clouds reduce radiative forcing (e.g., Myhre and Stordal, 1997). The IPCC estimate for CH4 forcing includes an observation-based estimate for both the direct forcing of CH4 and the indirect forcing due to changes in the hydroxyl radical (OH) and tropospheric ozone (O3) resulting from methane oxidation. The oxidation of CH4 leads to a net loss of OH in the atmosphere, thereby lengthening the CH4 lifetime. It is estimated that this indirect effect of CH4 increases its radiative forcing by 25-35 percent over the direct CH4 forcing (Lelieveld and Crutzen, 1992; Brühl, 1993; Lelieveld et al., 1993, 1998; Hauglustaine et al., 1994; Fuglestvedt et al., 1996). The oxidation of CH4 also leads to the formation of tropospheric ozone, indirectly increasing the CH4 forcing by 30-40 percent through the greenhouse effect of the additional tropospheric O3. In the stratosphere, oxidation of CH4 is a source of water vapor. In situ measurements of water vapor in the lower stratosphere indicate an increase of about 1 percent per year for 1954-2000 (Rosenlof et al., 2001), whereas satellite measurements of water vapor in the stratosphere in the 1990s showed no steady rate of change (Randel et al., 1996). A 1 percent annual increase in stratospheric water vapor would be associated with an estimated radiative forcing of 0.2 W m−2 since 1980 (Forster and Shine, 1999). The oxidation of CH4 can explain only a fraction of such a water vapor increase. TROPOSPHERIC AND STRATOSPHERIC OZONE Atmospheric ozone modifies the radiative budget of the Earth system by absorbing radiation both in the IR and in the ultraviolet (UV). It acts both as a radiative forcing agent and as a climate feedback. Ozone is produced and destroyed by solar UV radiation and by chemical reactions involving natural and anthropogenic gases. Changes in ozone driven by anthropogenic emissions represent a forcing. However, ozone concentrations also respond to changes in temperature and UV radiation, transport patterns, and natural emissions from lightning and vegetation; these responses represent climate feedbacks. In what follows, tropospheric and
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties stratospheric ozone are discussed separately because they are produced by different mechanisms and have very different radiative implications. Tropospheric Ozone Only 10 percent of atmospheric ozone resides in the troposphere, but this small fraction is of particular importance for climate forcing. Tropospheric ozone directly affects the radiative budget of the troposphere, and pressure broadening allows ozone absorption lines in the troposphere to extend into otherwise optically thin regions of the spectrum. Ozone is produced in the troposphere by photochemical oxidation of volatile organic compounds (VOCs) and carbon monoxide (CO) in the presence of nitrogen oxides (NOx = NO + NO2). Anthropogenic emissions of these precursors have caused large increases in tropospheric ozone over the past century. The increase is estimated to be 50-100 percent globally according to current global three dimensional chemical transport models (CTMs), and the resulting radiative forcing is in the range 0.2-0.5 W m−2 (IPCC, 2001). An important caveat is that the CTMs are unable to reproduce the low ozone concentrations observed in the late nineteenth century and early twentieth century, suggesting (if the observations are correct) that they overestimate the natural source of ozone. Model calculations constrained with the historical observations indicate a larger forcing from anthropogenic tropospheric ozone, up to 0.8 W m−2 (Mickley et al., 2001; Shindell and Faluvegi, 2002). Beyond this direct radiative forcing effect, tropospheric ozone also has an indirect effect as the primary precursor of the OH radical. Increasing ozone causes tropospheric OH to increase, thus decreasing the lifetime of methane and facilitating aerosol nucleation (both negative forcings). Assessing this indirect effect is complicated because the increase in ozone is driven by emissions of its precursors, which themselves have intrinsic effects on OH. Increasing NOx thus causes OH to increase, while increasing CO and VOCs cause OH to decrease. According to the current generation of CTMs, OH concentrations have decreased by about 10 percent over the past century (Wang and Jacob, 1998). Another indirect effect of tropospheric ozone is to cool the stratosphere (Joshi et al., 2003; Mickley et al., 2004), affecting stratospheric ozone and polar stratospheric cloud (PSC) levels. The greatest uncertainty in quantifying the direct radiative forcing from tropospheric ozone lies in reconstructing its concentration field in the past and projecting it into the future. The inability of current models to reproduce ozone observations from the early twentieth century could reflect calibration problems in the observations, as well as model errors in the estimates of natural sources. CTMs also have problems in simulating the
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties well-calibrated ozone trends over the past 30 years (Fusco and Logan, 2003), implying that fundamental problems remain in our understanding of tropospheric ozone chemistry. Uncertainty in quantifying the indirect radiative effect of ozone is related mainly to the complexity of factors controlling OH concentrations (Lawrence et al., 2001). Uncertainties in predicting the climatic response to changes in tropospheric ozone are also large and require further investigation using general circulation models (GCMs). Stratospheric Ozone Depletion of stratospheric ozone over the past 30 years has caused both a positive radiative forcing at the Earth’s surface (due to increased UV penetration) and a negative forcing (due to reduced IR emission from the stratosphere to the troposphere). The consensus from current radiative models constrained by observed ozone trends is that the net forcing is negative and of magnitude −0.10 ± 0.05 W m−2. Forster and Tourpali (2001) argue that about half of this forcing is due to an increase in tropopause heights and thus should not be considered a forcing but rather a feedback. The main indirect radiative effects of stratospheric ozone depletion are (1) increased UV penetration to the troposphere, increasing tropospheric OH concentrations and hence decreasing the lifetime of methane (IPCC, 2001), and (2) changes in stratospheric water vapor. Several GCM studies have examined the climate response to changes in stratospheric ozone. Shindell et al. (1999) finds that changes in the upper stratosphere elicit far greater surface climate response than changes in the lower stratosphere. Stuber et al. (2001) find that changes in stratospheric ozone have a greater effect per unit forcing than changes in CO2, largely because of feedbacks associated with stratospheric water vapor. The greatest uncertainty in quantifying radiative forcing from past changes in stratospheric ozone is the vertical distribution of the ozone trend in relation to temperature, since the magnitude of the forcing depends crucially on temperature (IPCC, 2001). Another critical issue is to better quantify indirect radiative forcings, particularly the effect on stratospheric water vapor, which could double the effective forcing according to Stuber et al. (2001). DIRECT EFFECT OF AEROSOLS Aerosol particles both scatter and absorb radiation, representing a direct radiative forcing; scattering generally dominates (except for black carbon particles) so that the net effect is of cooling. Global models have demonstrated the important role of sulfate aerosols in providing the cooling effect missing in past models of the atmospheric radiation balance (Kiehl et
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties al., 1995). The average global mean aerosol direct forcing from fossil fuel combustion and biomass burning is in the range from −0.2 to −2.0 W m−2 (IPCC, 2001). This large range results from uncertainties in aerosol sources, composition, and properties used in different models. Recent advances in modeling and measurements have provided important constraints on th direct effect of aerosols on radiation (Ramanathan et al., 2001a; Russell et al., 1999; Conant et al., 2003). Critical gaps, discussed further below, relate to spatial heterogeneity of the aerosol distribution, which results from the short lifetime (a few days to a week) against wet deposition; chemical composition, especially the organic fraction; mixing state and behavior (hygroscopicity, density, reactivity, and acidity); and optical properties associated with mixing and morphology (refractive index, shape, solid inclusions). The chemical composition of particles is in general not well known. The mixing state and relative humidity history of sulfate-nitrate-ammonium aerosols have important implications for their water content and hence their direct radiative effect (Martin et al., 2004). Uncertainties are particularly large for the 20 to 70 percent of particle mass that consists of organic compounds (NARSTO, 2003). Measurement of organic components is inherently difficult for three reasons: (1) small sample sizes and analytical difficulties, (2) the complexity of mixtures, and (3) artifacts in sampling procedures. The ideal approach for characterizing organic mass in aerosol particles would identify, molecule by molecule, the composition of each individual particle. Such instrumentation is unlikely to become available in the near future. In the meantime, it will be important to use partial information from traditional and new approaches, including evolved gas analysis techniques, gas chromatography, time-of-flight and chemical ionization mass spectrometry, Fourier transform infrared spectroscopy, and near-edge X-ray absorption fine structure (Morrical et al., 1998; Schauer et al., 1999; Huebert and Charlson, 2000; Russell et al., 2002; Bahreini et al., 2003; Russell, 2003). Radiative and climate models generally assume that aerosols are “externally mixed,” that is, individual particles are made up of a single component (Koch, 2001; Cooke et al., 2002). Actual aerosol particles are multicomponent mixtures for which properties, such as water uptake, differ from those expected from the simple addition of components because of nonlinear interactions between components. The consequence for radiative forcing is that water uptake by particles is not predicted accurately. The presence of organic compounds has two competing effects on particle hy-
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties groscopicity: (1) it reduces the mass of water taken up, and (2) it initiates water uptake at a lower relative humidity (called the deliquescence relative humidity) (Ming and Russell, 2002). The reduction in water uptake results from the typically low solubility of organic compounds. Particles containing organic compounds will grow less as a function of relative humidity, meaning that models that use the properties of sulfate aerosol will overestimate the direct radiative forcing. Many GCM aerosol schemes tend to omit organic particles or underestimate their size at typical boundary layer humidity (~80 percent). Chung and Seinfeld (2002) estimate the effect of organic carbon on radiative forcing of the climate to be between −0.09 and −0.21 W m−2, with the range of uncertainty driven by the role of water uptake by organic aerosols. The combination of the role of organic carbon and its water uptake with the externally or internally mixed states of other components results in a direct aerosol forcing range of −0.86 to −1.26 W m−2, or an uncertainty of ±50 percent (Chung and Seinfeld, 2002). In addition to water uptake, uncertainties in aerosol lifetime and optical properties contribute to the range of uncertainty. In addition to these general difficulties in describing the direct radiative forcing from aerosols, specific uncertainties relate to (1) light-absorbing black carbon, and (2) categorization of aerosol types in modeling. These are discussed below. Black Carbon Individual aerosol particles may contain light-absorbing carbon-containing compounds referred to collectively as “black” carbon (BC) or equivalently as soot. In addition to elemental carbon, BC frequently includes low-volatility solid or liquid organic compounds, typically composed of long hydrocarbon chains with high molecular weights (Marley et al., 2001). The presence of trace amounts of BC (as little as 5 to 10 percent of the total mass in anthropogenic aerosols) can result in large atmospheric solar absorption. This absorption can be enhanced when BC is embedded in refractive particles (Chylek et al., 1996; Fuller et al., 1999). Current understanding of the global emission of BC is uncertain by factors of two or more (Cooke et al., 2002; Bond et al., 2004). Biomass burning and fuel combustion are the two main contributors. BC has been detected in remote oceanic regions, implying hemispheric-wide dispersal. The direct forcing due to carbonaceous aerosols can be separated into three components: A portion of the direct solar beam is scattered back to space, which leads to a reduction in solar radiation reaching the surface. This reduction
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties manifests as increased reflection at the top of the atmosphere (TOA), i.e., a negative radiative forcing (cooling). A portion of the direct solar beam is absorbed by the aerosol, and this atmospheric absorption leads to further reduction in solar radiation reaching the surface. As shown later, this shielding of the surface by BC is the dominant absorption term for anthropogenic aerosols with as little as 10 percent of BC. This absorption leads to a positive radiative forcing of the atmosphere and a negative radiative forcing of the surface. The upward diffuse beam from scattered radiation is absorbed by the BC aerosol, reducing the solar radiation that escapes to space and resulting in a positive radiative forcing for the surface-atmosphere column. This effect could be large in cloudy skies if BC lies above low clouds (Haywood and Ramaswamy, 1998). The TOA forcing reported in IPCC and other global warming studies is the sum of processes (1) and (2). Process (3), albeit the largest in terms of magnitude, does not contribute significantly to TOA forcing, since it adds solar radiation to the atmosphere and reduces surface solar heating by the same magnitude. The net effect of BC is to increase the radiative heating of the atmosphere and decrease the radiative heating of the surface. The TOA radiative forcing is the sum of the surface and the atmospheric forcing. At the TOA the BC effect opposes the cooling effect of sulfates and organics, while at the surface all aerosols lead to reduction of solar radiation. Thus, aerosol-induced changes at the surface can far exceed those at the TOA. The direct effect of BC aerosol in the atmosphere has important implications for “global dimming.” Black carbon emissions may have increased by a factor of two to four during the last 50 years (Novakov et al., 2003). Given such large increases in BC emissions and the large impact of BC on reducing surface solar radiation, large decreases in surface solar radiation should be observed downwind of major sources of BC. Long-term negative trends in surface solar irradiances have been observed by surface radiometers worldwide (Ohmura et al., 1998; Stanhill and Cohen, 2001; Liepert, 2002). The reported trends in the annual mean irradiance vary from −5 percent (10 W m−2) between 1958 and 1985 for all land stations to about −1 to −3 percent per decade for the last four decades over many of the 1500 stations in the global datasets. The decreases are so large that there is skepticism about the measurements. Trends in surface radiometer all-sky observations are subject to large uncertainties due to difficulties in maintaining accurate calibration for routine surface observations in remote locations and measurement errors inherent in broadband radiometric measurements (Dutton et al., 2001). This global dimming is thought to be caused by both absorbing aerosols and increases in cloud cover (Liepert, 2002). Most of the radiometer
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties observations are over land, and we need to understand whether they are affected by local urban haze. Anthropogenic absorbing aerosols, by themselves, can reduce land-averaged solar radiation by about 3 to 5 W m−2 (Ramanathan et al., 1995; Jacobson, 2002). Such large reductions in surface solar radiation have implications for the hydrological cycle, since roughly 70 percent of the absorbed solar radiation is balanced by the latent heat flux of evaporation. This global dimming may also be related to changes in the ratio of direct and diffuse solar irradiance received by vegetation. As documented by Gu et al. (2003), the increase in diffuse irradiance for the two years following the eruption of Mount Pinatubo in 1991 resulted in a 23 percent increase of noontime photosynthesis of a deciduous forest in 1992 under cloudless conditions. The increased diffuse irradiance permits a greater penetration of photosynthetically active sunlight into the canopy. However, if there is a sufficient reduction of total solar irradiance received at the ground, vegetation growth could be stunted. Chameides et al. (1999), for example, reported on reductions in crop yield in China due to reduced total solar irradiance from pollution aerosols. Krakauer and Randerson (2003) concluded from tree ring data that with respect to volcanic eruptions, the beneficial effect of aerosol light scattering for high northern latitudes appears to be offset entirely by the deleterious effect of eruption-induced climate change. Using field observations, Niyogi et al. (2004) found that increased aerosol loading led to increases in carbon assimilation for forests and crops and decreases for grasslands. The effect on carbon assimilation was larger with aerosols than with clouds, since clouds reduced the total solar irradiance more than the aerosols did. Deposition of BC aerosols over snow-covered areas can result in changes to the surface albedo (Chylek et al., 1983). Further reductions in albedo occur due to the enhanced melting that accompanies the heating of absorbing soot particles in snow. Chylek et al. (1983) estimate this enhancement to be up to a factor of ten in the rate of melting. Recent model results indicate radiative forcings of +0.3 W m−2 in the Northern Hemisphere associated with albedo effects of soot on snow and ice (Hansen and Nazarenko, 2004). Model Discretization of Aerosols Typically the behavior of aerosol particles in the atmosphere has been described in models by discretization of both size and chemical composition. The continuous particle size spectrum is described by a limited number of modes, moments, or sections (Seinfeld and Pandis, 1998). For many problems, such as the evolution of marine aerosol, the computational simplicity of a few modes is sufficient to characterize changes in the particle
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties (Lean et al., 2002). Critical examination of a broader distribution of stars and examination of their “solar-likeness” appear to contradict the initial findings that noncycling Sun-like stars undergo Maunder-type episodes with reduced overall brightness. Likewise, an initial study of (uncalibrated) solar images failed to find evidence of a varying brightness component in the past century (Foukal and Milano, 2001). The most recent studies therefore raise the possibility that long-term solar irradiance variations may be limited to 11-year cycles. Solar forcing estimates based on changes in total solar irradiance are only approximations of the actual forcing. This is because of the wavelength dependence of both the magnitude and the variability of the solar spectrum and of atmospheric absorption that differentially attenuates the spectrum. Forcing estimates based on total solar irradiance assume that energy at all wavelengths reaches the Earth’s surface (or at least the troposphere) and that radiation at all wavelengths changes by the same amount. However the solar spectrum has less flux, but varies more, at shorter wavelengths. Furthermore, the Earth’s stratosphere absorbs solar irradiance at wavelengths less than 310 nm. The energy in the solar spectrum at wavelengths from 200 to 300 nm (15.3 W m−2) changes by about 1 percent during the 11-year cycle and accounts for 13 percent of the corresponding total irradiance cycle (Lean et al., 1997). Variations in solar ultraviolet irradiance alter the production and destruction of ozone, thereby influencing stratospheric temperature, dynamics, and chemistry. The subsequent coupling of the stratosphere with the troposphere (via radiative and dynamical pathways) is considered to produce indirect climate forcing by solar irradiance. Knowledge of variations in solar spectral irradiance is much poorer than for total solar irradiance. Observations have been made primarily in the ultraviolet spectrum, for only one decade, at wavelengths less than 400 nm. Estimates of variations in the visible and infrared regions have thus far relied on models of the wavelength dependence of the competing sunspot and faculae influences. Only the recently launched SORCE spacecraft has the capability to measure the solar spectral irradiance with the needed long-term precision. Preliminary data already raise questions about the modeled infrared spectrum variability (Fontenla et al., 2003). Whereas current understanding is that faculae are dark in the IR spectrum, SORCE observes increased IR irradiance when faculae are present on the Sun. Ionization and Production of Cloud Condensation Nuclei Galactic cosmic rays have one billionth of the total solar irradiance energy, but can reach the troposphere where they produce ions that may serve as nuclei for cloud condensation, with subsequent climatic impacts.
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties Because the heliosphere influences their transport, cosmic rays exhibit fluctuations that mirror solar activity. When solar activity is high, the more complex magnetic configuration of the heliosphere and the solar wind that flows through it from the Sun to the Earth reduce the cosmic ray flux. There is a close inverse correspondence with solar activity of products of the collision of cosmic ray particles with particles in the Earth’s atmosphere (smaller pions, muons), which ground-based neutron monitors have been measuring since the 1950s. The approximate 15 percent modulation of cosmic ray flux by solar activity produces an energy change less than one millionth of the energy change in the 0.1 percent total solar irradiance cycle. Cosmic rays also interact with air nuclei to produce isotopes such as 14C in tree-rings and 10Be in ice cores. Fluctuations in the 14C and 10Be records (which are superimposed on the larger variations associated with changes in the Earth’s magnetic field) are believed to reflect primarily changes in long-term solar activity (Stuiver, 1965; Beer et al., 1990), although climate effects cannot be ruled out (Lal, 1988). During the last 100 years, the 10Be record suggests a 15 percent overall decline in cosmic ray flux. By altering the population of cloud condensation nuclei and hence microphysical cloud properties (droplet number and concentration), cosmic rays may induce processes analogous to the indirect effect of tropospheric aerosols (Carslaw et al., 2002). Since the plasma produced by cosmic ray ionization in the troposphere is part of an electric circuit that extends from the Earth’s surface to the ionosphere, cosmic rays may also affect thunderstorm electrification (Carslaw et al., 2002). Analysis of cloud cover data reveals decadal variations apparently related to solar-modulated galactic cosmic ray fluxes (Svensmark and Friis-Christensen, 1997), but because solar activity modulates both cosmic ray fluxes and solar irradiance, it is difficult to distinguish which forcing mechanism is responsible for such empirical evidence. The evidence can readily be reinterpreted as association of solar irradiance and cloud cover (Udelhofen and Cess, 2001; Kristjánsson et al., 2002). Using 16.5 years of cloud cover data from the International Satellite Cloud Climatology Project (ISCCP), Kristjánsson et al. (2002) found that low cloud cover correlates better and more consistently with total solar irradiance than with galactic cosmic rays. The data suggest that solar irradiance variations are amplified by interactions with sea surface temperature, which in turn interacts with low cloud cover. In another study, Udelhofen and Cess (2001) found a high coherence between cloud cover inferred from ground-based observations and solar variability over the United States from 1900 to 1987 (but of opposite phase to that found in the ISCCP low clouds). Using cloud coverage simulated by a climate model, they found cloud cover variations in phase with solar
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties variability but not with the galactic cosmic ray flux. They suggest that the cloud variability is affected by a modulation of the atmospheric circulation resulting from variations in the solar-UV-ozone-induced heating of the atmosphere. Orbital Variation Earth’s distance from the Sun does not remain constant at 1 AU. Rather, the eccentricity of Earth’s orbit (currently 0.0167) and the tilt of its axis relative to the orbital plane result in continual changes in the amount and distribution of solar electromagnetic radiation that the Earth receives. In modern times this variation is ±3.5 percent during the year, with maximum energy and minimum distance in January. Indirect Effects Through the Stratosphere Of the Sun’s mean total radiative output of 1365 W m−2, 15 W m−2 (~1 percent) of energy is in the ultraviolet spectrum and does not reach the Earth’s surface (e.g., Lean et al., 1997). This energy is deposited in the stratosphere, where it drives ozone formation (and also destruction). Although unavailable for direct forcing of climate, it may induce indirect climate effects as a result of radiative and dynamical coupling of the stratosphere and troposphere. The regional pattern of such indirect climate forcing is likely quite different from the effects of direct surface heating by solar radiation. The effect of solar cycle UV irradiance changes on stratospheric ozone are now relatively well established as a result of extensive space-based datasets that span more than two solar activity cycles (McCormack et al., 1997). As Figure 2-7 illustrates, the 11-year cycle of ~1 percent peak-to-peak amplitude in middle UV radiation is associated with a 2 to 3 percent modulation of global total atmospheric ozone. The solar UV-induced ozone effects vary with geographic location and altitude, and appear to induce a significant tropopause response (Hood, 2003). As with tropospheric climate, solar-induced ozone changes occur simultaneously with other natural and anthropogenic effects that must be understood and quantified in order to isolate the solar component (Jackman et al., 1996; Geller and Smyshlyaev, 2002). Most evident is a long-term downward trend in total ozone concentrations associated with increasing concentrations of CFCs. The 11-year solar cycle is superimposed on this trend (Figure 2-7), as are the influences of volcanic aerosols (which warm the stratosphere while cooling the surface), greenhouse gas increases (which
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties cool the stratosphere while warming the surface), and internal variability modes (in particular the quasi-biennial oscillation). Energetic particles (1 to 100 MeV) produced during eruptive solar events can also produce significant episodic ozone depletion, primarily at higher latitudes (where the particles preferentially enter the Earth’s atmosphere) and for relatively short periods (days). Ozone depletion arises from the odd nitrogen chemical destruction cycle that the particles initiate (Jackman et al., 2001). These depletion events, whose frequency and strength vary with solar activity, are superimposed on the more sustained solar UV radiation-induced ozone changes that occur during the 11-year solar cycle. The extent to which solar UV radiation and energetic particle effects FIGURE 2-7 Top panel: global total atmospheric ozone observed by the TOMS satellite (McCormack et al., 1997). Bottom panel: solar ultraviolet irradiance observed at 200-295 nm (Lean et al., 1997).
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties have indirect climatic impacts depends on the coupling of the stratosphere with the troposphere. Both radiative and dynamical couplings are surmised (Figure 2-8). Since ozone absorbs electromagnetic radiation in the UV, visible, and IR spectral regions, changes in ozone concentration can affect Earth’s radiative balance by altering both incoming solar radiation and outgoing terrestrial radiation. Simulations of this effect (Lacis et al., 1990) show (Figure 2-8) that the net change of surface temperature depends on FIGURE 2-8 Change in surface temperature resulting from a change in ozone concentration as a function of altitude. SOURCE: adapted from Lacis et al. (1990).
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties the altitude of the ozone change; increases below 29 km produce surface warming, and increases above 29 km produce surface cooling. Model simulations suggest that such radiative coupling effects can alter the strength of the Hadley cell circulation, with attendant effects on, for example, Atlantic hurricane flows (Haigh, 2003). Solar-induced indirect effects on climate may also involve altered modes of variability. Model simulations and analyses of patterns of variability suggest that the Arctic Oscillation (AO), or Northern Annular Mode (NAM), and its subset the North Atlantic Oscillation (NAO) propagate from the stratosphere to the troposphere (Baldwin and Dunkerton, 1999). Radiative forcings that impact the stratosphere could alter this coupling. Contemporary observations suggest that the NAM manifests itself primarily in the North Atlantic sector, as the NAO, during solar cycle minima, and extends more uniformly over all longitudes, as the AO, during solar maxima (Kodera, 2002). The effect of the Sun on the NAM may further depend on the phase of the quasi-biennial oscillation (QBO) in stratospheric equatorial winds (Ruzmaikin and Feynman, 2002). Reduced solar activity in the Maunder Minimum may have produced a negative NAO phase (compared with the current positive phase), based on empirical analysis of historical surface temperature fields and model simulations (Shindell et al., 2001b). Additional evidence that the phase of the QBO changes with the solar cycle (Salby and Callaghan, 2000) underscores the complicated, multifaceted nature of indirect solar effects on climate. VOLCANIC ERUPTIONS Emissions from volcanic eruptions have multiple effects on climate as listed in Table 2-3 (Robock, 2002). A number of studies have evaluated the role of volcanic forcing in climate change during the twentieth and earlier centuries (Free and Robock, 1999; Crowley, 2000; Bertrand et al., 2002; Bauer et al., 2003). These studies suggest that volcanic forcing is the dominant source of natural global radiative forcing over the past millennium. The greater prevalence of explosive volcanic activity during both the early and the late twentieth century and the dearth of eruptions over the interval from 1915 to 1960 represents a significant natural radiative forcing of twentieth century climate (e.g., Crowley, 2000). Similarly, the longer-term volcanic radiative forcing has been associated with a significant long-term forced cooling from A.D. 1000 to A.D. 1900 resulting from a general increase in explosive volcanic activity in later centuries (Crowley, 2000; Bertrand et al., 2002; Bauer et al., 2003; Crowley et al., 2003; Hegerl et al., 2003). Some spatially resolved simulations of volcanic forcing indicate a large continental summer cooling but a tendency for a dynamically induced, offsetting winter warming (Stenchikov et al., 2002; Shindell et al.,
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties TABLE 2-3 Effects of Large Explosive Volcanoes on Weather and Climate Effect and mechanism Begins Duration Reduction of diurnal cycle Immediately 1-4 days Blockage of shortwave and emission of longwave radiation Reduced tropical precipitation 1-3 months 3-6 months Blockage of shortwave radiation, reduced evaporation Summer cooling of Northern Hemisphere tropics and subtropics 1-3 months 1-2 years Blockage of shortwave radiation Reduced Sahel precipitation 1-3 months 1-2 years Blockage of shortwave radiation, reduced land temperature, reduced evaporation Stratospheric warming 1-3 months 1-2 years Stratospheric absorption of shortwave and longwave radiation Winter warming of Northern Hemisphere continents 6-18 months 1 or 2 winters Stratospheric absorption of shortwave and longwave radiation, dynamics Global cooling Immediately 1-3 years Blockage of shortwave radiation Global cooling from multiple eruptions Blockage of shortwave radiation Immediately Up to decades Ozone depletion, enhanced UV radiation Dilution, heterogeneous chemistry on aerosols 1 day 1-2 years SOURCE: Robock (2000). 2003). This result contrasts with the response to solar forcing, for which the dynamical and radiative responses appear to reinforce constructively. Past histories of radiative forcing by explosive volcanic activity are typically constructed from sulfate aerosols contained in annual ice core layers (e.g., Robock, 2000). Spikes of sulfate in the ice core records reflect volcanic injection to the lower stratosphere, where the lifetime is a year or longer, allowing transport to polar regions and eventual deposition after subsidence to the troposphere (Robock and Free, 1995). The longer the residence time of the aerosol in the lower stratosphere, the greater is the associated negative shortwave radiative forcing of the surface through the
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties reflection of radiation back to space. Assumptions must be made regarding the relationship between the sulfate aerosol deposited at the surface and the extent and duration of a significant stratospheric dust veil. These assumptions are highly uncertain and can be only partially tested for a few recent eruptions (e.g., Stenchikov et al., 1998). Greater concentrations of trapped sulfates are typically indicative of larger eruptions, although the proximity of the source region to the ice core may be a complicating factor. Explosive tropical eruptions are more likely to impart a significant global radiative forcing because they provide an opportunity for the aerosol to spread throughout the global lower stratosphere. An eruption is assumed to have occurred in the tropics if its aerosols are recorded in ice cores at both poles. Other indices of past volcanic activity have also been developed in past work. These include the Volcanic Explosivity Index, or VEI, which is based on qualitative volcanological information and should therefore be used with caution in studies seeking quantitative estimates of climate response (Robock and Free, 1995), and the Dust Veil Index (DVI; see, e.g., Robock, 2000). Some authors argue that the use of climate information in some DVI estimates leads to a potential circularity in using this index to diagnose climate response. Tree-ring reconstructions of continental summer temperature variations have also been used to estimate past volcanic forcing histories (Briffa et al., 1998), although a similar circularity obviously exists if the associated volcanic histories are used to diagnose the climate response to volcanic forcing. Zielinski (2000) and Robock (2000) provide excellent reviews and critiques of various indices of past explosive volcanic activity. Ice core volcanic radiative forcing estimates have been developed for the past century to the past couple of millennia by numerous researchers (Robock and Free, 1995, 1996; Robertson et al., 1998; Robock, 2000; Crowley, 2000; Ammann et al., 2003; Crowley et al., 2003). The choice of ice cores used to define the volcanic forcing chronology leads to some significant differences among these different estimates. Some of the estimates assume that tropical eruptions dominate the annual global mean radiative forcing (e.g., Free and Robock, 1999; Crowley, 2000), whereas other reconstructions seek to take into account the influence of the latitudinal and seasonal characteristics of the eruptions (Robertson et al., 1998; Ammann et al., 2003). OCEAN COLOR Recognized biologically related surface forcings associated with the oceans include ocean color as well as biogenic aerosol emissions (e.g., dimethyl sulfide). Ocean color refers to the radiance backscattered at the air-sea interface. It is determined by water molecules (the blue wavelengths in particular), phytoplankton and detrital particles, and nonbiogenic sedi-
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties ments in coastal waters (Yoder et al., 2001). Bacteria, viruses, colloids, and small bubbles are also possible contributors. Shell et al. (2003) used a global climate model to assess the role of ocean color in the sea surface temperature and other aspects of the climate system. They found that phytoplankton warm the surface by about 0.05°C on a global average basis. They also found that the large-scale atmospheric circulation is significantly affected by regional alterations of ocean color. These results suggest that the radiative effects of phytoplankton should not be overlooked in studies of climate change. Frouin and Iacobellis (2002) also determined that absorption of sunlight by phytoplankton must be included in the global radiation budget. They estimated that, compared to pure seawater, the globally and annually averaged outgoing radiative flux is decreased by 0.25 W m−2 due to ocean phytoplankton. In coastal and high-latitude regions, the forcing can reach around 1.5 W m−2. They also found that the amount absorbed was species dependent. TELECONNECTIONS AND RADIATIVE FORCING Linkages between weather or climate changes occurring in widely separated regions of the globe are referred to as teleconnections. The extent to which regionally concentrated radiative forcing can affect climate via teleconnections is a matter of current research. Determining the importance of regional forcings, such as those from aerosols or land-use change, requires an understanding of the role of teleconnections that can lead forcings in one region to have effects on other regions far away. Teleconnections are most commonly thought of with respect to the transport of energy by atmospheric waves (Tsonis, 2001). For example, regional and global weather patterns have been associated with sea surface temperature anomalies (e.g., Hoerling and Kumar, 2003). Radiative and nonradiative forcing due to regional land-use change can also result in large differences in atmospheric circulation patterns at large distances from the landscape disturbance. For example, land-use change can alter deep cumulonimbus patterns, which affect atmospheric circulation in distant regions (Chase et al., 2000a). Avissar and Werth (2005) found that deforestation of tropical regions, through teleconnections similar to those produced during El Niño events, has a significant impact on the rainfall of other regions. In particular, they found that the U.S. Midwest is the continental region the most negatively affected by the deforestation of Amazonia and Central Africa during spring and summer, when rainfall decrease could severely damage agricultural productivity in that region. These results are summarized in Figure 2-9. Avissar and Werth (2005) conclude that tropical deforestation considerably
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties FIGURE 2-9 Annual cycle of precipitation (mm day−1) in continental regions particularly affected by the deforestation of Amazonia (red), Central Africa (green), and Southeast Asia (blue). The blue curves represent the mean monthly precipitation before massive deforestation started in tropical regions (i.e., the “control” case). The red curves indicate the corresponding precipitation following tropical deforestation. The size and location of the color-coded areas corresponding to the deforested regions are at scale. Color-coded ellipses indicate the regions in which tropical forest (in green on the 1-km resolution land-cover map used for the background) was replaced with a mixture of shrubs and grassland. SOURCE: Avissar and Werth (2005). alters the sensible and latent heat released into the atmosphere and the associated change of pressure distribution modifies the zones of atmospheric convergence and divergence, which shift the typical pattern of the Polar Jet Stream and the precipitation that it engenders. Radiative forcing by aerosols has also been associated with teleconnected responses in distant locations. For example, a GCM simulation by Chung and Ramanathan (2003) shows that absorbing aerosols over South Asia and the North Indian Ocean can cause subsidence motions over most of the tropics, which would have a drying effect (Figure 2-10).
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Radiative Forcing of Climate Change: Expanding the Concept and Addressing Uncertainties FIGURE 2-10 Velocity potential for the lower troposphere (850 hPa or about 1.5 km) in the lower panel and for the upper troposphere (200 hPa) in the upper panel. The solar heating by absorbing aerosols, mainly due to black carbon, is concentrated over South Asia and the North Indian Ocean (i.e., over the red shaded regions in the lower panel). The red region in the lower panel shows areas of convergence of air or alternately rising motions in response to the solar heating of the lower atmosphere by black carbon. The convergence at the lower levels is followed by divergence (air flowing out of the region) at the upper levels (the blue shaded region in the upper troposphere) over the source region. SOURCE: Adapted from Chung and Ramanathan (2003).
Representative terms from entire chapter: