Physical Chemistry of Climate Metrics - Chemical Reviews (ACS

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Physical Chemistry of Climate Metrics A. R. Ravishankara,† Y. Rudich,*,‡ and D. J. Wuebbles§ †

Departments of Chemistry and Atmospheric Science, Colorado State University, 1872 Campus Delivery, Fort Collins, Colorado 80523, United States ‡ Department of Earth and Planetary Sciences, Weizmann Institute, Rehovot 76100, Israel § Department of Atmospheric Sciences, University of Illinois, 105 South Gregory Street, Urbana, Illinois 61801, United States provided enough references to such material so that the interested readers can find the details and the nuances of the issues discussed here. In particular, we refer to an excellent review on radiative forcing and radiative efficiency by Hodnebrog et al.1 for some of the details mentioned here. We start with a very brief review of the basic concepts and quickly move to the physical chemistry that underlines the climate metrics. Then we describe the metrics and their uses and limitations. The greenhouse effect is a concept that has been known for almost two centuries.2 The potential impact of CO2 from fossil CONTENTS fuel combustion on surface temperature has also been known 1. Introduction A for more than a century.3 Thus, climate change due to 2. Physical Chemistry Considerations C anthropogenic activity is not a new idea. However, it has risen 2.1. Greenhouse Gases C in prominence in the past few decades and is now considered 2.1.1. Atmospheric Infrared Window C among the most important environmental issues of the 21st 2.1.2. Infrared Cross Sections D century. As such, understanding climate change and its effects 2.2. Atmospheric Lifetime D on humanity and our planet requires rigorous physical science 2.3. Aerosols E tools to understand its past and future behavior as well as to 2.3.1. Aerosol Distribution and Properties E project future changes in the face of actions taken by society. 2.3.2. Aerosol-Cloud Interactions and InfluThe energy flow in the Earth system involves energy input ence on Climate F from the sun, part of which is intercepted before reaching the 2.3.3. Absorption and Scattering by Particles G surface by complex exchange of energy within the atmosphere, 3. Metrics for Changes in Climate I reflection of the light by the surface, absorption at the surface, 3.1. Radiative Forcing J emission of infrared radiation from the surface, and eventually 3.1.1. Radiative Forcing by Greenhouse Gases K emission of energy back to space. Figure 1 shows the 3.1.2. Radiative Forcing by Aerosols K complexities and various pathways for the flow of energy in 3.1.3. Radiative Forcing and Its Connection to the atmosphere. Climate Change K The sun is the major (essentially the only) source of energy 3.2. Global Warming Potentials (GWPs) M to the Earth’s surface and the atmosphere. The solar spectrum 3.2.1. CO2 Equivalent N is roughly represented by a blackbody radiation characteristic of 3.3. Global Temperature Change Potentials ∼6000 K.5 It spans all wavelengths. However, the portion of the (GTPs) O spectrum that carries the most energy extends from the far3.4. Recent Evaluations of GWPs and GTPs Q ultraviolet (UV) to far-infrared. The atmosphere absorbs 3.5. Uses and Abuses of Climate Metrics R almost all the UV radiation at wavelengths shorter than roughly 4. Future Needs and Conclusions R 290 nm by the time sunlight reaches the surface. Most of the Author Information R energy reaching the surface is in the visible region; the Corresponding Author R maximum of the radiant energy distribution is at approximately Notes R 500 nm. The fraction of sunlight that reaches the surface Biographies S depends on clouds, absorption and scattering by aerosols and Acknowledgments S molecules above the surface, and by the solar zenith angle. References T Further, a part of the incoming radiation is reflected by the surface, and that fraction depends on the surface characteristics. The remaining energy is absorbed and heats the surface. The 1. INTRODUCTION This paper is written with chemists in mind, and therefore, concepts are explained in physical chemistry terms. In the process, some of the important atmospheric dynamics and radiation transfer concepts have been simplified. We have © XXXX American Chemical Society

Special Issue: 2015 Chemistry in Climate Received: January 5, 2015

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Figure 2. Simplified schematic of the greenhouse effect. Incoming energy from the sun is shown as the blue arrow. While some radiation is reflected to space due to the Earth’s albedo (not shown in the graphic), a significant fraction of the radiation from the sun passes through the atmosphere and warms the surface and the lower atmosphere (see Figure 1). Then, infrared radiation from the surface (and the lower atmosphere) characteristic of a temperature between ∼200 and 320 K is emitted and is shown by the orange arrows. Some of the radiation escapes the Earth back to space and cools the planet. In the absence of greenhouse gases (left diagram), the radiation from the surface will equal the incoming energy when the system has reached equilibrium. The presence of infrared absorbing gases (right diagram) intercepts some of the infrared radiation and re-emits it at the local temperature, with the net effect that some of the energy is trapped as shown by the thick orange arrow and heats in the Earth atmosphere system. However, the warmer lower atmosphere and the surface emits radiation that is at slightly shorter wavelengths (blueshifted) from the one in the absence of the greenhouse gases. As a result, the surface and the lower atmosphere will be warmer. Eventually, the radiation from the surface and the warmed atmosphere/surface reaches a new equilibrium state, and again, the incoming radiation is equal to the outgoing radiation. However, the surface and the lower atmosphere will be warmer as a result.

Figure 1. Complex energy flow in the earth system. Key points to note include the following: the incoming solar energy and the outgoing energy are in the form of both visible and infrared radiation. The left side of the diagram shows the interactions with the incoming radiation, and the right side shows the effects on the outgoing IR radiation. The effects of greenhouse gases (right side) and aerosols (center) on the radiation balance are shown. In this paper, we focus on the effects of human activities on changes in atmospheric abundances of aerosols and greenhouse gases, and the resulting forcing on the climate system, and not on other effects such as changes in ocean color, ice/snow cover, vegetation, and land use. These factors also influence atmospheric concentrations of greenhouse gases as well as the abundances and properties of aerosols. (Reproduced with permission from ref 4. Copyright 2013 Cambridge University Press.)

brium because the greenhouse gases are increasing. Even if emissions of greenhouse gases were to stop, it would take hundreds of years for the Earth’s surface and lower atmospheric temperature to reach equilibrium, but the majority of the response would occur in about 20−30 years. The long time scales are because of the response time for the oceans and the large ice masses. In addition to greenhouse gases, condensed matter aerosols, and clouds, also influence the radiation balance. Aerosols are condensed matter suspended in air; however, cloud droplets or ice particles are not considered aerosols. The primary effect of aerosols is to scatter and/or absorb incoming solar radiation. If the aerosols were pure scattering particles, they would reduce the amount of incoming solar energy reaching the surface and would thus lead to cooling of the surface. If the aerosol particles absorbed incoming solar radiation, they would heat the atmosphere. They would still reduce radiation reaching the surface, but the overall effect on the surface and lower atmosphere would be warming. There are no aerosol particles that only scatter or only absorb radiation throughout the entire spectrum. Therefore, the net effects of aerosol particles depend on their relative scattering and absorption capabilities.4,8 In addition to their direct interactions with incoming solar radiations, aerosol particles alter other characteristics of the atmosphere such as cloudiness, light scattering properties of clouds, and precipitation. These secondary influences of aerosols are large and are collectively called aerosol indirect effects or aerosol−cloud interactions. Changes in the abundance of greenhouse gases and aerosols create a forcing on climate, called radiative forcing (see below), which is balanced by a change in the Earth’s surface temperature. In addition, it has other impacts such as changes

surface then emits back radiation characteristic of 200−320 K, according to the Stefan−Boltzmann law, and this occurs mostly in the infrared (IR) region (see below). Without greenhouse gases (GHGs) in the atmosphere, the IR radiation emitted by the surface would escape to space and cool the planet (Figure 2). At equilibrium, the incoming energy would be equal to the outgoing energy. Under these conditions, the calculated mean surface temperature of the Earth would be about 250 K.6 The so-called “greenhouse effect” is the trapping of infrared radiation emitted by the surface and the lower atmosphere by infrared active chemicals in the atmosphere. Greenhouse gases in the atmosphere absorb much of the infrared radiation emitted by Earth (both from the surface and the lower atmosphere). This absorption leads to heating of the lower atmosphere, the surface, and the oceans. When there is a constant amount of greenhouse gases in the atmosphere and when the system reaches equilibrium, the outgoing energy again equals incoming energy, but the surface and the lower atmosphere remain warmer than what it would be without the greenhouse gases. In other words, because of the presence of atmospheric greenhouse gases, the surface temperature is higher than what would be the case if there were no greenhouse gases. The most important natural greenhouse gases are water vapor and carbon dioxide; without either one of them the Earth would be a frozen planet.6 Other planets with infrared absorbing gases in their atmospheres also exhibit a greenhouse effect.7 Increase in the atmospheric burden of the greenhouse gases due to human activity leads to further increases in surface temperature. Currently, Earth’s atmosphere is not in equiliB

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scatters and absorbs light. We need to know the atmospheric lifetime to connect the amounts in the atmosphere to emissions (as described below). To a first approximation, most of the input regarding the properties of the GHGs and aerosols that force the climate are physicochemical in nature, and they are derived from laboratory and/or field studies. This is the primary role played by physical chemistry in determining the metrics for climate change.

in evaporation and precipitation, increase in sea level, melting of glaciers, etc. Nonetheless, changes in globally averaged surface temperature are commonly used as a proxy for climate change because it is one of the primary impacts and it is better understood than the changes in other atmospheric impacts. Climate metrics are tools used for quantifying impacts of specific emissions from human activity on Earth’s climate; they are especially useful in comparing two or more emissions for their impacts. These metrics aggregate and generalize information about different emissions and then place them on a common scale to simplify comparison of impacts. Metrics such as the ozone depletion potentials, ODPs, have proven to be valuable for making and negotiating policy for stratospheric ozone-layer depletion.9 Similar climate metrics include radiative forcing (RF) and global warming potentials (GWPs, defined below), and they are being used in climate policy and management related activities. These metrics have all been used in past assessments of ozone and climate, including major international assessments of the science.4,9 Newer metrics, such as global temperature change potentials (GTPs),10 are being considered for use in policymaking considerations.

2.1. Greenhouse Gases

The main point to consider when assessing their effect on climate is the influence that molecules emitted into the atmosphere have on the energy balance. Therefore, the properties of the greenhouse gases of interest are (1) the superposition of the thermal infrared spectrum of the molecule of interest with the emission spectrum of the surface and the lower atmosphere characterized by a blackbody temperature of 200−320 K, i.e., ∼4 to 100 μm or 2000 to 100 cm−1; (2) the absorption cross sections of the molecule of interest that determine how much of the radiation may be absorbed; and (3) the atmospheric lifetime of the molecule that determines how long and where the molecule emitted to the atmosphere will remain and accumulate in the atmosphere, and thus be available to absorb IR radiation. These three parameters are derived mostly from laboratory studies coupled with atmospheric chemistry modeling calculations. They are briefly discussed below. 2.1.1. Atmospheric Infrared Window. We are interested in the change in the absorption of infrared radiation in the atmosphere due to the presence of a greenhouse gas emitted into the atmosphere. The radiation emitted by the surface and the lower atmosphere is characteristic of a blackbody between ∼200 and ∼320 K (see Figure 4). This figure also shows the wavelength distribution of the radiation at the top of the atmosphere that escapes to space as measured by a satellite instrument looking at Earth. If there were no IR absorbers in the atmosphere, the spectrum would be that of a blackbody at a temperature characteristic of the surface (and lower atmosphere) that emits the IR radiation. That spectrum, of course, would depend on the location, time of the year, etc. For example, it is obvious that the emissions from an ice-covered surface would be at a lower temperature than those from a hot desert. The key point to note is that there are some regions of the spectrum that are essentially opaque because of the presence of strong absorbers at high concentrations in the present atmosphere. Water vapor and CO2 absorb a large fraction of the radiation between 25 and 14 μm, near the peak of the IR emission curves where most of the outgoing energy is present. Therefore, further addition of CO2 and H2O is less effective than addition of molecules that absorb in the window regions. Consequently, increased absorption due to the emission of CO2 above what is already present does not lead to a proportionate response. In the “atmospheric window” regions, in contrast, the extent of absorption increases linearly with the increase in the concentration of a greenhouse gas. The absorptions by molecules already present in the atmosphere complicate the calculation of the effectiveness of the increase in the concentration of a given greenhouse gas. Most potent greenhouse gases (other than CO2 and water vapor) absorb in the atmospheric window region. Therefore, knowledge of the IR spectrum of a greenhouse gas in this region is of great importance. A high resolution IR spectrum of a gas is needed to

2. PHYSICAL CHEMISTRY CONSIDERATIONS Many physical factors affect the determination of the climate metrics. A simple flowchart for the calculation of climate metrics is shown in Figure 3 below. This flowchart is inspired by a detailed diagram for greenhouse gases given by Hodnebrog et al.1 The figure shows that the key input for calculating the effect of a species on the radiation balance is its amount and distribution in the atmosphere along with how the species

Figure 3. Chart of the flow of information for the calculation of climate change metrics. Optical properties of greenhouse gases and aerosols are a critical input and are primarily derived from laboratory studies or field observations. In addition, the atmospheric lifetime (derived from laboratory data and atmospheric modeling) is a primary input that connects emissions to atmospheric abundances. This article focuses on the red arrows that are used to derive the climate metrics. The topics discussed in this article are shown in blue in the graphic. In the case of long-lived (well-mixed) greenhouse gases, the current radiative forcing is calculated from atmospheric observations rather than from a derivation of their concentration via modeling studies. C

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elaborate since the incoming and outgoing radiations are calculated at multiple vertical levels at a given geographical location for specified conditions of sun angle, overhead amounts of absorbers and scatters, time of the year, and other parameters such as surface albedo. These codes account for absorption, scattering, and reflection. Many different computer codes have been developed. A commonly used code for calculating incoming solar radiation is the NCAR TUV calculator.12 There are multiple codes for calculating outgoing IR. For example see Forster et al.13 and Clough and Iacono.14 The IR cross sections for many of the molecules of interest are already in IR databases, with the HITRAN database15 being one of the most commonly used resources. In some cases, the spectrum is calculated using known (assessed and tabulated) molecular parameters of the greenhouse gas. This database is updated frequently and is used not only for calculation of climate metrics but also for satellite retrievals. The infrared cross sections of most greenhouse gases are known and accessible. When a new molecule is suggested for commercial use, assessment of its climatic effects requires measurement of its infrared cross sections. Lastly, IR absorption line positions and spectra can often be calculated or estimated from knowledge of the molecular structure. Therefore, rough estimates of the potential contributions of a greenhouse gas can be made even without laboratory measurements. However, accurate values require measurements of IR spectra, absorption strengths, and linebroadening parameters.

Figure 4. Wavelength distribution of radiation from Earth as measured in space. Spectrum of the IR radiation escaping from earth measured from space via the NASA IRIS D spectrometer aboard the Nimbus-4 satellite. It shows the spectrum measured above the hot Sahara desert. Black body curves for 220 to 320 K in 20 K intervals are superimposed on the IR spectrum. Clearly, the measured spectrum does not fit any one specific blackbody curve, but is close to that for 320 K blackbody for the peak values. Attenuation of IR radiation in certain regions shown in the graphic is due to strong absorptions by H2O, CO2, and O3, which are identified in the graphic. In addition, there are many other IR features in the spectrum that are less obvious but can be identified by careful analyses. The regions between ∼10 and ∼12.5 μm and between ∼8 and ∼9 μm, which have minimal attenuation, are referred to as window regions; these are marked in the graphic as blue shaded regions. Molecules that absorb in these regions tend to be potent greenhouse gases. Reproduced with permission from ref 11. Copyright 1972 The American Geophysical Union.

2.2. Atmospheric Lifetime

Atmospheric lifetime is the key parameter for determining how much of an emission accumulates in the atmosphere and the rate of decay of the radiative forcing from a pulsed injection of a greenhouse gas (see Figure 5). Lifetimes of precursor gases are the determining factors for calculating the concentrations of other greenhouse gases produced in the atmosphere, in

calculate the overlap of its absorption features with those of other absorbers in the atmosphere. The effectiveness of a new added absorber in attenuating the outgoing radiation depends on the specific wavelength and widths of its absorption features, which can change with temperature and altitude. Since the pressure varies with altitude, knowledge of the location of the absorbing molecule in the atmosphere is also important to account for pressure-broadening for each IR feature. Similarly, the temperature, which affects the partition function of the greenhouse gas, has to be taken into account in calculating its absorption at a given location. Therefore, ideally, one needs the absorption spectra at various temperatures and pressures of the atmosphere. 2.1.2. Infrared Cross Sections. The IR absorption cross sections of greenhouse gases are usually measured in the laboratory. In general, measurements of the IR spectra and absorption cross sections for individual ro-vibronic lines are straightforward since all the gases of interest are stable molecules. Usually, standard IR spectral measurement methods are used to determine the cross sections. Cross sections at as high a resolution as feasible are needed so that the data can be binned or integrated to any needed resolution. (In some cases, one needs to include the cross sections at high resolutions.) In most atmospheric calculations, integrated band strengths are used instead of the cross sections for the individual ro-vibronic line.1 If high resolution cross sections are available, they can be degraded to the resolution needed for the specific radiative transfer calculations. Radiative transfer calculations are often

Figure 5. Decay of the abundance of a pulsed injection of chemical according to first order kinetics. The simplest definition of an atmospheric lifetime of a molecule is the time taken for the removal from the atmosphere of 63% (1/e) of a pulsed emission of a gas into the atmosphere. In atmospheric chemistry, other concepts such as turnover time are used in place of a lifetime.16 (Also see Burkholder et al.17 in this issue.) However, for the climate metrics under consideration here, simple lifetime as noted here is used. D

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removal rate. In practice, the concentrations of the reactants are calculated from numerical atmospheric chemistry-transport models. In some cases, it is also important to consider situations when the concentration of the molecule of interest affects the concentration of the free radical responsible for its removal. An example is the lifetime of CH4 (due mostly to its reaction with OH), where the atmospheric concentration of OH radicals responsible for CH4 removal is affected by CH4 itself!22 The key parameter needed for calculating the photolytic loss rate is the so-called “J-value”, which is the first order rate coefficient for photolytic removal of the chemical. This is also calculated using atmospheric models coupled to radiative transfer calculations or simply using models such as the TUV model noted above. The primary quantities needed for the calculation are UV−vis absorption cross sections and quantum yields for the removal of the molecule following the photon absorption.23 These quantities are also obtained via laboratory studies. Calculation of the removal by heterogeneous processes is somewhat more complex. All of these have been described and reviewed in the past; see articles by Ravishankara and Lovejoy,17 Kurylo and Orkin,16,24 and Kolb et al.25 Therefore, we will not describe them further in this paper. Heterogeneous reactions also lead to loss of a molecule from the atmosphere. They include nonreactive uptake of the molecule on/into condensed phase and reactive uptake of that which transforms the molecules. In reality, whether an uptake is reactive or nonreactive could depend on how long the particle exists in the atmosphere. Even a dissolved gas that is in equilibrium with the gas phase (taken up reversibly) could be slowly lost via reaction with another constituent in the condensed phase. If the reaction of the molecule is with water, the uptake can lead to rapid removal; thus, hydrolysis of a chemical is always a consideration. Such reactions can affect the gas phase concentration of species (such as uptake of soluble species by aqueous droplets and aerosols) or the chemical compositions of the participating particles. Burkholder et al.17 give more details on the processes and formulation of the calculation of the removal rates. Again, compilation and evaluations of necessary data for calculating heterogeneous reactions are available; they are updated periodically to provide up-to-date information.21 We have described above processes that determine the atmospheric lifetimes of a species emitted into and/or produced in the atmosphere. All these processes occur simultaneously in the atmosphere, albeit at different rates, and the overall lifetime has contributions from each of them. Necessarily, the dominating processes will be one or another of the above loss pathways; accurate quantification of the dominant loss process is a major pursuit in this research area. In general, these lifetimes are calculated using numerical atmospheric models.

particular ozone. Ozone is not emitted into the atmosphere but made and destroyed (partly) in the atmosphere. The calculation of atmospheric lifetime requires consideration of all removal processes such as reactions with atmospheric free radicals and oxidants (e.g., OH, NO3, Cl, O3), heterogeneous uptake into condensed matter, photolysis by solar radiation, and deposition to the surface. The lifetime is needed in the equations for the climate metrics discussed below. From a physical chemistry perspective, the concept of lifetime is easiest to define when the loss process for the molecule is first order in its concentration. The removal of short-lived (less than a year) chemical species cannot be characterized by a single lifetime. Lifetimes of short-lived chemicals depend on where and when they are emitted, and their forcing will be spatially inhomogeneous.9b,18 For example, n-propyl bromide, used as a solvent, is highly reactive, and its lifetime varies by a factor of 3 from around 10 days to 30 days depending on the location of the emissions.18 If a chemical were even shorter lived, there can be even a wider range of lifetimes depending on the location of the emissions. Short-lived chemicals do not accumulate much in the atmosphere, and their abundances will be small, unless their emissions are very large as in the case of aerosols. Therefore, in general, short-lived greenhouse gases are not significant forcers of climate. Nonetheless, climate metrics for short-lived compounds are calculated, and they tend to be uncertain. Their lifetimes would depend on where and when they are emitted. Their short lifetimes also lead to nonuniform distribution of their abundance in the atmosphere.4 Molecules with a sufficiently long lifetime (roughly >2 years) are usually sufficiently well-mixed in the troposphere (subsequent to their continued emissions for a few years) such that a single atmospheric lifetime could be calculated and is a reasonable approximation for climate metrics. If the molecule were removed rapidly only in the stratosphere, as for example for chlorofluorocarbons (CFCs), the major controlling factor would be the time taken for circulating all of the tropospheric air through the stratosphere. Similarly, if a molecule were removed only in the mesosphere, its atmospheric lifetime will be controlled by the time it takes to pass the entire tropospheric content through the mesosphere. This time is very long, on the orders of centuries, and the lifetime of such a molecule will be very long even if its instantaneous lifetime in the mesosphere is very short. To evaluate atmospheric chemical lifetimes, it is essential to quantify all the processes that contribute to the removal of the species from the atmosphere. Common reactions that lead to removal of gases of climate importance from the atmosphere include those with ozone, OH, Cl, and NO3 radicals; solar photolysis (especially in the UV); and loss through processes involving interactions between trace gases, aerosol particles, and cloud droplets termed heterogeneous and multiphase processes.19 Other loss processes include dry deposition, photolysis by Lyman-α radiation in the upper stratosphere and mesosphere, uptake by plants, etc.20 To quantify the gas phase loss processes, the rate coefficients for the reactions of the molecule with the individual free radicals noted above are needed as a function of temperature, pressure, and occasionally composition. These parameters are almost always obtained via laboratory studies. Evaluations and a compilation of kinetics data for use in atmospheric modeling and studies are available.21 The atmospheric concentrations of the free radicals are also needed for the calculation of the

2.3. Aerosols

2.3.1. Aerosol Distribution and Properties. Distribution and abundance of aerosols in the troposphere are determined by emissions, chemical transformations, transport, and deposition. Aerosols are short-lived in the troposphere, with a tropospheric residence time of usually a few days to a few weeks. (They are longer lived in the stratosphere because they are essentially removed via transport from the stratosphere to the troposphere. They could also be altered via coagulation or condensation.) Therefore, their abundances vary greatly from E

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Figure 6. Complex ways in which aerosols influence radiation in the atmosphere. Pathways by which aerosols influence climate include direct interactions with radiation and indirect interactions via changes in the properties and extent of clouds. In addition, there are many feedbacks in the system due to these changes, and they are noted on the right side of the diagram. (Reproduced with permission from ref 4. Copyright 2013 Cambridge University Press.)

or absorbing and scattering particles can be in the same parcel of air (externally mixed). One of the key properties that determine the effectiveness of a particle in cooling versus heating is the single scattering albedo, which is the fraction of the light that is scattered relative to the extinction (scattering and absorption) caused by the aerosols. Because of these reasons, there are large regional and global differences in the impact of aerosols on radiation and clouds. The uncertainty in the direct radiative forcing by aerosol particles (called clear sky forcing) is due to uncertainties in their optical properties as well as their atmospheric concentrations and distributions.28 Moise et. al.26c address the optical properties of aerosols in this issue. Human activities that emit particles to the atmosphere include combustion of fossil fuels (especially coal and oil), land use change, deforestation, and biomass burning. Therefore, knowledge of their contributions is another factor that leads to uncertainties in anthropogenic forcing and climate metrics due to aerosols. 2.3.2. Aerosol-Cloud Interactions and Influence on Climate. In addition to the aerosol direct radiative effects particles can also affect climate indirectly by changing the properties of clouds (Figure 6).29 Aerosols that contain soluble components can act as cloud condensation nuclei and thus enhance cloud droplets’ formation. An increase in the number of aerosol particles can also increase the number of cloud droplets and consequently decrease the average size of cloud droplets for a given amount of water vapor in the atmosphere. This enhancement creates more cloud droplets that are smaller in size and thus make brighter clouds. Consequently, clouds with a larger number of smaller droplets reflect away more incoming sunlight, and result in an overall surface cooling. As a result the cloud can live longer in the atmosphere and extend to larger sizes.30 Particles that absorb sunlight may evaporate clouds and thus lead to warming.4 It is estimated that the particles currently offset between 20% and 35% of the warming caused by the increasing concentrations of greenhouse gases.4 The largest uncertainty associated with the radiative forcing from aerosols is due to uncertainties in the indirect forcing from the interaction of aerosol particles with clouds.

region to region and from season to season. Consequently, they mostly influence radiation regionally and locally. Exceptions include large dust storms or smoke from large biomass fires, which can have significant effects on continental scales as well as large volcanic eruptions that inject aerosol precursors to the stratosphere that influence on a global scale. Another complexity with aerosols is that they come in different sizes, composition, optical properties, and abilities to take up water. Further, their characteristics change as they “age”, i.e., undergo chemical and physical changes in the atmosphere. All these attributes enhance the complexity of the role of aerosols in the atmosphere.26 The size of atmospheric aerosol ranges from a few nanometers to tens of micrometers.20 Primary aerosols are particles that are directly emitted to the atmosphere while secondary aerosols form in the atmosphere.20 Primary aerosols include black carbon, biological particles, sea salt, and dust. Secondary inorganic aerosols are produced from reactions involving sulfur dioxide, ammonia, and nitric oxide emissions and include sulfate, nitrate, and ammonium containing particles. Secondary organic aerosols (SOAs) are formed via gas phase oxidation of nonmethane hydrocarbons (and their products) with the hydroxyl radical (OH), ozone, nitrate radical (NO3), or photolysis.27 They may also be formed via aqueous phase processes in cloud and fog droplets. Many hydrocarbons in the atmosphere are of biogenic origin. However, anthropogenic pollutants influence their conversion to SOAs and thus can be considered anthropogenic in origin. Furthermore, often aerosols have both inorganic and organic content. There are large uncertainties in our understanding of the processes involved in the formation, growth, and optical properties of SOAs.26c The size and composition of aerosol particles can be modified by chemical reactions in the particles, condensation on to or evaporation of gaseous species from particles, and coagulation of existing particles (see Seinfeld and Pandis20 and Kolb et al.25). Aerosols can simultaneously have both cooling and warming influences on climate. The cooling and warming components of aerosols can be mixed within each particle (internally mixed), F

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Figure 7. Chart of the information flow for calculating the climate metrics for aerosol emissions. Aerosols influence the forcing in two key ways: (1) directly via absorbing and scattering incoming radiation, which is termed clear sky forcing; and (2) indirectly via changing clouds and their properties that influence incoming solar radiation and also the outgoing infrared radiation. The swelling of the particles due to increases in relative humidity and how the particles are mixed adds further complexity. The relative humidity dependence is derived from both laboratory and field studies. The mixing issue is often based on assumptions and/or field observations. The major topic discussed here is limited to the discussion of the optical properties of the aerosol particles and a very brief discussion of the calculation of the radiative forcing. In general, the GWP and GTP of emissions through their formation of aerosols are not usually calculated, and therefore, those arrows are shown in gray.

is possibly the best known of the aerosol influences. The major absorbing aerosol in the current atmosphere is black carbon produced from incomplete combustion that absorbs incoming radiation and thus leads to a warming influence.8a This is also the most uncertain of the aerosol direct influences. The latest IPCC assessment lists the current estimates of the various aerosols to climate forcing.4 On the basis of the above description of how aerosols influence climate, one could envision a flowchart for the calculation of climate metrics of aerosols. We briefly describe the parameters that determine the influence of aerosols and how they are obtained, Figure 7. 2.3.3. Absorption and Scattering by Particles. Atmospheric aerosols scatter both visible (solar) and infrared radiation. While all aerosols scatter radiation, most aerosol components are weak absorbers of visible and UV radiation. The scattering components include common inorganic species such as sodium chloride (NaCl), ammonium sulfate ((NH4)2SO4), ammonium bisulfate (NH4HSO4), and sulfuric acid (H2SO4). The most important aerosol component that absorbs incoming solar radiation is black carbon containing particles such as soot and mineral dust particles that contain absorbing species such as hematite (Fe2O3). In addition, some organic aerosol particle components of aerosols also absorb in the visible part of the spectrum, and less in the UV region. The strongest near-UV absorbing organics include humic like substance (HULIS), nitrated aromatics, polycyclic aromatic hydrocarbons (PAHs), and some complex aldehydes and ketones. To model the interaction of aerosols with radiation, a few basic parameters are needed, and they are described below. 2.3.3.1. Refractive Index. The refractive index is the key parameter that determines how the condensed matter interacts with light. The refractive index dictates the extent of scattering

The estimation of and effects from aerosols, particularly the difference between absorbing (e.g., black carbon8a) and scattering (e.g., sulfates, nitrates) aerosols, has varied over the last two decades in the different assessments. This is partly because of new information on sources of the particles, their concentrations, and resulting direct and indirect effects, and partly because the aerosol issue has been reframed many times as the understanding of particle processes improves. Black carbon is especially complicated; on its own, black carbon is strongly warming, but it also has additional indirect effects on the albedo of snow and ice by surface deposition that further amplifies its effects. However, black carbon is almost always emitted in combination with organic carbonaceous aerosols and gases that form secondary organic aerosols. Therefore, a climate metric for direct emissions of aerosols, or their precursors, has to account for both warming and cooling effects. This issue is discussed in the recent assessment4 and synthesis paper.8a Natural emissions can also alter climate. For example, highly explosive volcanic eruptions that directly inject material (sulfur and crustal material) into the stratosphere are known to cool the Earth system. The cooling persists for a few years after the emission until the sulfate and other particles are removed from the stratosphere. This is the best-known influence of aerosols on climate, with cooling observed subsequent to large volcanic eruptions, and some having led to disastrous consequences.31 Other examples, as noted earlier, include large scale dust storms and large scale biomass burning. The earliest major aerosol that was considered in climate studies was sulfate from the oxidation of SO2 emitted by combustion processes, especially coal burning.32 Since sulfate (either sulfuric acid or ammonium sulfate or other neutralized forms of sulfate) tends to scatter radiation with minimal absorption, it was considered to have a cooling influence. This G

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scattered. As the absorption by a particle increases, its scattering intensity decreases since the absorption hinders scattering. Once the scattering and absorption cross sections are known, it is possible to calculate the so-called single scattering albedo (SSA) which is the ratio between the scattering and the total extinction (scattering and absorption) by a particle.

and absorption of light as a function of the wavelength of the incident light.33 The complex refractive index of a particle is usually described as m = n + ik

(1)

where n and k are the real and imaginary parts of the refractive index and i is the imaginary unit. The real part of the refractive index determines the extent of scattering while the imaginary part determines the extent of absorption. Both m and k are wavelength-dependent. Measurement of refractive indices has a long history in science. Laboratory measurements of refractive indices for surrogates have been carried out; see for example Beyer et al.34 and Moise et al.26c The extent of scattering and absorption depends on the size of the particle. Aerosols are present in various sizes, and there is a distribution of sizes.20 Therefore, the contribution to scattering and absorption by a collection of particles (aerosols, cloud droplets, haze, etc.) is calculated by summing the scattering and absorption cross sections of the particles as a function of the aerosol size and the wavelength. 2.3.3.2. Mie Scattering. Rayleigh scattering is the major contributor to the attenuation of light by particles that are substantially smaller than the wavelength of light. Scattering by larger particles, such as cloud droplets, can be calculated using their geometrical sizes. Scattering is enhanced, however, when the particle size and sunlight wavelength are approximately the same, which is very common for atmospheric aerosols. This enhancement is usually calculated using Mie scattering theory.35 In 1908, Gustav Mie solved Maxwell’s equations for the scattering of an electromagnetic plane wave (light) by homogeneous spherical particles.36 This is also known as the Lorenz−Mie solution. Particles are considered to have a single refractive index (m) in this formulation. The solution provides the direction and magnitude of the scattered light due to the interactions of the particle with the electric field of the light. Due to the size of the particle and interference patterns, the light is scattered in preferred directions relative to the incident direction. Mie calculations provide this information. The effective absorption and scattering cross section, b, by a particle of radius r is given by

b = πr 2 × Q

SSA = σs/(σs + σa)

Here σs is the scattering cross section, and σa is the absorption cross section. The single-scattering albedo is a unitless quantity. An SSA of unity implies that all of the particle’s extinction is attributed to scattering; the lower the single-scattering albedo, the higher the absorption. It is noted that the SSA varies as a function of size, and this quantity is often derived by simultaneously measuring extinction and scattering.39 2.3.3.3. Measurement of Optical Properties of Aerosol Particles. In general, until recently optical properties of aerosol particles were measured in the laboratory. One recent way to do so is to use cavity ring down spectroscopy at discrete wavelengths. In this case, the extinction efficiency is measured for each particle size at a given wavelength, and then, the refractive index is calculated by fitting the measured extinction to Mie theory by varying the refractive index in the fitting procedure.37,40 More recently, broadband cavity enhanced spectrometers have been used to derive the refractive index of aerosols as a function of wavelength.41 Recent developments have allowed for estimation of optical properties from atmospheric measurements, most commonly by Aeronet stations.42 Other articles in this issue specifically address the issue of optical properties of aerosols.26c In most cases, atmospheric particles are composed of several components mixed together homogeneously (“internally mixed”), or have some type of core−shell configurations. Various mixing rules and approximations have been developed for these cases. When the particles do not obey the basic assumptions of Mie scattering theory, more sophisticated methods, such as a T-matrix38 and the discrete dipole approximation, are used to account for these deviations.43 2.3.3.4. Aerosol Growth with Relative Humidity. Aerosol particles swell when relative humidity increases. The change in particle size alters the scattering properties of the aerosols. The major controlling factor in the growth of aerosols is the amount of available water. It is known that particles generally increase in size proportionately to relative humidity (not the partial pressure of water vapor). From a physical chemistry perspective, this can be viewed simply as a dependence on the activity of water since relative humidity is essentially the ratio of the water vapor partial pressure in the atmosphere to the vapor pressure over water at that temperature. Of course, composition of the particles can alter the activity coefficient, and thus, this relationship is not strictly correct. This change in aerosol size with relative humidity is an extremely important factor in determining the influence of aerosols on climate forcing.32a,44 When the humidity reaches large values (>90%), some particles become precursors to the formation of a cloud droplet or ice particles. The eventual conversion of an aerosol particle to a cloud particle occurs for some particles but not all. Such particles are called cloud condensation nuclei. The conversion of an aerosol particle to form an ice particle is very rare. Such aerosol particles that can lead to ice particles are called ice nuclei (IN). It is estimated that only one or two in a thousand aerosol particles are IN. Because of these reasons, the growth of

(2)

where Q is the scattering (or absorption) efficiency at that wavelength and πr2 is the particle’s geometrical cross section. The scattering efficiency is highest when particle size is comparable to the wavelength of the light. The absorption and scattering efficiencies are functions of the wavelength and the composition-dependent complex index of refraction m and the size parameter, x. The size parameter is simply the ratio of the particle’s geometrical cross section and the incident wavelength, and it is given (assuming a spherical particle) by x=

2πr 2 λ

(4)

(3)

More complex calculations, for inhomogeneous particles containing a core surrounded by a shell of material and nonspherical particles, have formulated special solutions for nonhomogenous particles, such as for core−shell geometries.38 The single-particle scattering efficiency is the ratio of the effective scattering cross section of a particle to its physical cross section. The scattering efficiency can exceed unity (and actually approach a value of 2 at the geometrical optics limit), as radiation diffracting around a particle can be diverted and H

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requires a sufficiently accurate capability for evaluating the effects of human-related and natural factors that affect the climate. Such capabilities include complex state-of-the-art climate models that have sufficiently accurate representations of the components of the model, i.e., the atmosphere, its physics and its chemical composition, the oceans, biosphere, cryosphere, etc. Furthermore, they have to capture the interactions between these components. These models need to encapsulate our best understanding of the underlying physical, chemical and biological processes; hence, they are large and require large computing resources. Furthermore, it requires substantial expertise to analyze, diagnose, and synthesize results from the large amount of output that is generated. The number of different cases (e.g., emission scenarios) considered is limited by computational resource as well. Alternatively, simpler models or metrics (that build on the results of the complex models) can be used to obtain information about various options and emission scenarios. For example, in the case of the ozone-layer depletion, it is quite common to use 2-dimensional models to examine options and scenarios for action by society.9a Clearly, trade-offs between relevance on one hand and uncertainty in the choice of impacts to be assessed on the other often exist. The potential uses for climate metrics include the following:49 (1) providing rapid evaluations of multiple potential approaches that are proposed to minimize the impact of human activities on the climate system; (2) evaluating the relative contributions of two or more emissions from different human activities to climate change; (3) evaluating the climate effects of competing technologies or energy usage as well as the contributions of different emissions even in one given economic sector such as aviation; (4) quantifying the relative contributions of countries; (5) establishing a basis for comparing changes in climate effects in different countries or regions; (6) evaluating proposed policies that encourage some beneficial activities and discourage others that are not beneficial; (7) informing and helping industries and countries to determine their own best approaches and practices to meet specific commitments or to reduce climatic impacts. The following are some key required features of a metric: (1) It should not only be scientifically sound, but also be simple to use and easy to understand and communicate. (2) It has to be applicable to the scientific questions or policy issues of interest to the user. (3) It must be a tool for communicating impact information among scientists, people from industry and government, and the public sector policymakers. (4) It should be transparent enough to convey the intended information by itself, i.e., the users (policymakers and managers) should be able to use the metric without further input from the scientific community that derived the metric. (5) It must be simple. (6) The users must have confidence in the scientific integrity of the metric and trust it. (7) Most importantly, it must be used! Therefore, it should be subject to a minimum of uncertainties or have the effects of scientific uncertainties reduced (or at least represented) as much as possible. Given these features, the main concern with developing new metrics is the need to weigh their applicability against ease of understanding its results. When formulating a metric it is important to choose the key parameters that represent the climate change in that metric. The major climate change impacts include changes in precipitation or severe weather or sea level rise. Clearly, there is a need for regional metrics such as regional precipitation changes. However, current climate models have significant

the particle with relative humidity is a very important parameter. Aerosol growth with relative humidity is often measured in the laboratory using proxies for particles present in the atmosphere (see for example Baynard et al.,45 Garland et al.,46 and Gysel et al.47). The improved instrumentation has enabled such measurements in the atmosphere with atmospheric aerosols.48 Indeed, measurement of aerosol growth factors has become routine. However, fundamental approaches to predicting the aerosol growth factors are at their infancy, and improvements in this capability will enable a priori estimation of aerosol growth.

3. METRICS FOR CHANGES IN CLIMATE One advantage of the currently prevalent climate metrics (radiative forcing, GWP, and GTP) is that they are straightforward to communicate. The user can apply these metrics (akin to economic metrics such as the Dow−Jones index, GDP, etc.) easily even though the details of their formulation and calculation are not completely understood by them. GWP and GTP express the integrated impact of the emission of a given gas relative to emission of the same mass of CO2, the most important anthropogenic greenhouse gas. The relative nature of these two metrics (or indices) leads to the cancellation of some of the uncertainties associated with the calculations and with our understanding of the climate system, while highlighting the relative benefit of controlling emissions of specific gases. It is important, however, to recognize that these metrics do not capture the full complexity of the chemistry and physics of the atmosphere and the climate system (e.g., where and when the temperature changes).49 Their simplicity also has another downside since the metrics conceal some important information. For example, they do not show the extent to which they depend on the background atmosphere assumed in their derivation, or how influences by other large perturbations such as a volcanic eruption will affect temperature change. Therefore, caution is warranted when interpreting and using the derived values. Nonetheless, GWPs have found widespread use in international agreements such as the Kyoto Protocol and the Montreal Protocol as well as in several national regulatory actions. These metrics do incorporate some key relevant physical principles and workings of the Earth system. They also include some value judgments. This point is particularly important since the metrics are designed for nonscientists, such as policymakers, and, thus, have incorporated some key inputs selected by the users. The time horizons (see below) used in the metrics’ calculation are the most significant input from the users. Some general questions must first be answered before using a metric. Such questions include the following: What is the metric used for? Is it used for technology or policy considerations? To which scenarios and forcings can the metric be applied? How effective is the metric for the specific use? Is the metric amenable for updating to incorporate advances in scientific understanding when they occur? When either developing a metric or choosing between existing metrics one must balance between the ease of use and the applicability of the metric to the range of climate altering scenarios used in developing the metrics. Further, it is important to consider the limits of scientific understanding embedded in deriving the metric. Development of meaningful metrics for quantifying climate change and instituting mitigation and adaptation policies I

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Figure 8. Measured extinction cross section of polystyrene particles as a function of their size due to Mie scattering. Measured scattering (left y-axis) by polystyrene particles is shown as a function of their geometrical diameter. The geometrical cross section of the particles (left y-axis) is shown as a dotted curve. The measured scattering cross sections were fitted to Mie theory with only the refractive index of the particle as a variable and shown as the solid curve connecting the measured cross sections. The obtained refractive index was 1.455, in close agreement with the known value. The ratio of the measured Mie cross sections to the geometrical cross sections, expressed here as QExt, is shown (right y-axis). (Reproduced from ref 37, Copyright 2004, with permission from Elsevier.)

uncertainties in calculating these impacts. It is also possible to consider economic impacts, such as change in damages and abatement costs, as the parameter of interest. However, such metrics are difficult to calculate due to the significant uncertainties in our knowledge of the economic consequences of climate change, and there is no general consensus on the best approach for such estimations. 50 Therefore, the most commonly considered parameters are related to the forcing of the climate system and the most easily quantifiable global impact, i.e., change in radiative forcing and change in global surface temperature. They have wide acceptance in the science community. Temporal changes are also important. For example, it is possible to consider the change in the absolute value of a climate parameter over a given time period, the integrated change in the parameter over a given time period, and/or the effects of pulsed or sustained emissions. Choosing one versus the other can affect decisions, e.g., whether it is better to reduce emissions of long-lived gases or short-lived forcing agents. Global values tend to be more robust than regional values. Further, climate change is a global issue with global impacts irrespective of the emissions’ origin. Therefore, the metrics that have found widespread use are all generally global metrics. We note, however, that regional metrics are available (see for example Shindell and Faluvegi51), and with improved scientific understanding they are likely to emerge and gain use. Other considerations for the choice of metrics exist. They include factors such as the most appropriate structure for the metric. The structure may have to consider if it is applicable to the targeted range of temperature changes. The choice of the metric will likely depend on how the intended climate policy is to be designed. Another factor is the quantification of input values (due to underlying uncertainties) and the need for value judgments in the choice of parameters within these metrics (e.g., the evaluation of long-term impacts versus short-term impacts). Such value judgments go beyond natural sciences.

In the discussion below, we describe three of the most commonly used metrics for climate change, namely, radiative forcing, GWPs, and GTPs. Other metrics, such as integrated GTPs (iGTPs; Peters et al.52), have been suggested but are not currently in common usage. 3.1. Radiative Forcing

Radiative forcing (RF) is a metric that is commonly used to measure the change in Earth’s radiation balance due to a change in concentration of a greenhouse gas (or other agents). Radiative forcing estimates the influence of a perturbation on the Earth’s radiation change due to anthropogenic and natural activities and processes. They include natural forcing, such as changes in the sun’s output, and forcing due to anthropogenic activities. These changes in RF lead to other effects on climate, adding or offsetting the different influences. Such effects are considered to be feedbacks on the climate system. The distinction between forcing and feedback is sometimes difficult to parse out, since feedbacks come about via the changes in agents that cause forcing. The term “radiative forcing” has been used since the 1980s as a metric of climate change. It has become a central tool for assessing climate change in the national and international arena. The 2001 IPCC report53 described radiative forcing as “a useful concept, providing a convenient first order measure of the relative climatic importance of different agents”. The simplicity of this metric is that it does not need time-consuming and computationally expensive climate model simulations. However, as discussed below, there are significant limitations to the traditional definition of radiative forcing for spatially inhomogeneous perturbations to the climate system, and alternative definitions have been suggested and developed.4 Simply put, radiative forcing represents a “thought experiment”; it is the instantaneous change in the energy escaping to space due to a change in the atmospheric concentration of radiatively active species (such as greenhouse gases or J

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Figure 9. Conceptualization of radiative forcing. The panel on the left shows the incoming radiation that heats the surface. It emits infrared radiation that transmits through the atmosphere, with a fraction that is absorbed by the background atmosphere, and then a certain amount escapes to space. At equilibrium, the incoming energy equals that escaping to space. The center panel shows that the increase in the greenhouse gas abundance reduces the outgoing infrared radiation, as shown by the shorter arrow outside the atmosphere. Of course, the atmosphere adjusts to this perturbation on various time scales, and different parts of the atmosphere do so on different time scales. If nothing else changes, the lower atmosphere and surface reach a new equilibrium (and a warmer temperature as shown by the brown color of the earth’s surface), with the outgoing energy being the same as the incoming energy as shown in the right panel. The wavelength distribution of the outgoing energy will be slightly different from the one in the left panel.

aerosols). This is schematically illustrated in Figure 8 for the case of a greenhouse gas. 3.1.1. Radiative Forcing by Greenhouse Gases. The above description is what is called the top of the atmosphere radiative forcing, Figure 9. This works quite well for the greenhouse gases since they, to a first approximation, interact with only the outgoing radiation. [Ozone in the stratosphere contributes significantly to the absorption of incoming radiation and thus heating the stratosphere; this influence is not included here. Absorption of incoming radiation by ozone is minimal in the troposphere because of the much smaller abundances of ozone there (90% is in the stratosphere). Ozone is also a greenhouse gas, and changes in its tropospheric and stratospheric distributions also affect climate, but not as much as the large changes occurring in concentrations of the well-mixed greenhouse gases like CO2.] In general, the important greenhouse gases are reasonably well-mixed in the atmosphere, and therefore the global top of the atmosphere forcing is a good measure of their influence. Shorter-lived greenhouse gases are, however, unevenly mixed in the troposphere, which makes it difficult to accurately determine their radiative forcing. However, the very short lifetime leads to minimal accumulation of such gases in the atmosphere, and their influence on climate is usually small. Even in these cases, it is not uncommon to calculate climate metrics that are useful in terms of giving a crude measure of the impact such a gas on climate. 3.1.2. Radiative Forcing by Aerosols. Unlike the greenhouse gases, aerosols directly interact with the incoming radiation by scattering and/or absorbing radiation. This introduces some complexity in how we treat aerosols. Figure 10 shows how the aerosols can change the amount of radiation reaching the surface and the fraction that is scattered back to space. Absorption leads to direct heating of the atmosphere. Scattering leads to a reduction in the heating of the surface. In the absence of clouds, these interactions of the aerosols can be calculated and are called “clear sky radiative forcing”. The ways in which aerosols alter clouds and thus radiation are complex, as described earlier. Aerosols are short-lived in the troposphere. Their lifetime depends on where they are in the atmosphere and could vary from a few hours to a few weeks. They are removed by precipitation, direct surface interactions, and possibly by processes such as evaporation. Because of their short lifetimes, aerosols are inhomogeneously distributed in the atmosphere. Yet, unlike the case of the short-lived greenhouse gases,

Figure 10. Schematic of how aerosols influence climate. Incoming radiation in the UV and visible region is either scattered back out of the atmosphere or absorbed (shown by the black arrow and filled black circles) in the atmosphere. Both of these influences reduce the radiation reaching the surface, and thus lead to a cooling effect, as shown by the reduced size of the emission from the surface. However, absorption of the incoming light by absorbing aerosols directly heats the atmosphere and thus warms the surface as well as the atmosphere. The net effect of the presence of the aerosols critically depends on the scattering and absorbing properties of the aerosol.

aerosols have a major influence on climate because their abundance is very large. Indeed the local forcing by aerosols can be tens of W m−2. These large local and regional values are such that they contribute, on the global scale, to values that are a significant fraction of the forcing by CO2. However, for the most part the overall influence of the aerosols is to cool the atmosphere. The various factors that lead to the large uncertainties in the aerosol indirect forcings have been discussed in section 2.3.2. We noted earlier (not discussed further here) the influence of aerosols on clouds and thus indirectly on climate. Currently, aerosols (including their indirect contributions) are calculated to offset the radiative forcing by anthropogenic greenhouse gases by roughly 25− 35%. 3.1.3. Radiative Forcing and Its Connection to Climate Change. The concept of top of the atmosphere radiative forcing provides an estimate of the potential effect on climate from greenhouse gases and other radiatively active species. It can be inferred from radiation flux measurements for the most prominent forcers. The regional radiative forcing by large abundances of aerosols can indeed be calculated from suitably designed measurement. The radiative forcing depends on the infrared absorption properties and the concentration of a specific gas in the atmosphere. The radiative forcing due to a specific gas can be K

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troposphere is strongly coupled to the surface through convection. Therefore, climate models have typically found that the land surface, ocean mixed layer, and troposphere respond together to a radiative forcing from perturbations. This leads to a relatively uniform increase in globally averaged temperature with the radiative forcing. As a result, the steady state form of the heat change equation is

calculated provided that its infrared absorption cross sections are known (see section on Physical Chemistry Considerations above). The calculations are done by radiative transfer models that consider the absorption and scattering in the process of transmitting light through the atmosphere as well as the reflectance of clouds and the surface. The concept of radiative forcing implicitly assumes that the Earth−atmosphere system is approximately in a radiative convective equilibrium (the convective adjustment is the upward transport of heat due to convection that restores the temperature distribution with altitude, also called the lapse rate, while conserving energy).54 The atmospheric heating rate can be derived by assuming a radiative−convective equilibrium: dH ΔT =F− dt λ

ΔT = λF

Using eq 7 it is possible to estimate the change in surface temperature for a given radiative forcing. λ is called the climate sensitivity factor. It describes the temperature change expected for a doubling atmospheric concentration of CO2 from its preindustrial levels. It ranges from 1.5 to 4.5 °C change in surface temperature for a 3.8 Wm−2 increase in the radiative forcing. This range in temperature change gives an estimate of the uncertainty in the climate sensitivity factor. A number of studies have examined the climate sensitivity factor, λ, but they have not succeeded in reducing the range of uncertainty in its value.4,53,58−61 The uncertainties associated with how aerosols and cloud processes are treated in the model appear to be the major cause for the range of values that are obtained in climate models. (See for example Sherwood et al.62) In addition, the varying strength of feedbacks, in particular stratospheric water vapor feedback and the sea icealbedo feedback, have also been suggested as the largest factors in the variability of λ.63 Does the climate sensitivity factor depend on which gas is responsible for the forcing? To a first approximation, this did not appear to be the case. For example, Ramanathan et al.,64 using a one-dimensional radiative convective model, found that λ was almost invariant with the type of the forcing agent. Since then, other climate models have also shown that the equilibrium global mean temperature at the surface varies linearly with global mean change in radiative forcing at the top of the atmosphere. Even though different climate models obtain different values of λ (ranging from 1.5 to 4.5 °C, as mentioned earlier), each model gives an approximately constant value for λ for changes in solar flux as well as atmospheric concentrations of long-lived gases like CO2, CH4, and N2O. Subsequent to studies by Ramanathan et al.,64 a number of investigators (see for example, Wang et al.,65 Hansen et al.,13a,66,67 Jain et al.,67b Naik et al.,67a Forster et al.,13a Gauss et al.,66b Gohar et al.,66c Huang and Ramaswamy,66d Meehl et al.,58a and Tett et al.66f) examined the direct influence of different trace greenhouse gases on climate as well as the influence of reactive gases indirectly through changes in abundances of greenhouse gases. Wang et al.68 noted that global climate models need to consider the many trace gases, which are found at concentrations much smaller than CO2. Wang et al.68 also showed that the influence of CO2 is different from that of other trace gases, because of differences in absorption wavelengths and atmospheric lifetimes. Starting in the 1990s there have been various modifications to the definition of radiative forcing from its original implementation for the global climate system. Subsequent to the finding that different forcers contribute differently to surface temperature change, the concept of radiative forcing has taken these findings into account. In addition, its use has been extended to determine regional and seasonal mean radiative forcing to evaluate the effects of shorter-lived greenhouse gases (such as hydrochlorofluorocarbons, HCFCs) and the shortlived aerosols (Wang et al., and Haywood and Ramaswamy69).

(5)

where H=

∫z

∞

ρCpT dz b

(7)

(6)

is the heat content of the atmosphere, F is the forcing on the system, ΔT is the change in temperature in the system due to the forcing, λ is the climate sensitivity parameter that accounts for the effects of climate feedbacks, ρ is the density of the atmosphere, Cp is the specific heat of air at constant pressure, and zb is the depth that the thermal energy penetrates into the atmosphere. Within the troposphere, the vertical mixing of sensible heat (where input of energy leads to an increase in temperature) and latent heat (where input of energy is used to induce a phase transition and thus there is no change in temperature) by largescale motions and convection is quite rapid in comparison to the time scales associated with responses to radiative changes (e.g., see Ramanathan et al.55). As a result, dynamical processes largely govern the vertical distribution of the tropospheric temperature change, while the mass-weighted tropospheric temperature change is governed by the radiative forcing of the column. However, one does have to account for changes in stratospheric temperature and long-wave emission in order to compute the flux changes at the tropopause and the corresponding effects on climate in the troposphere. In the case of the stratosphere, time scales associated with radiative adjustments are comparable to, or faster than, those associated with dynamical processes. Consequently, the magnitude of the stratospheric climate change is influenced strongly by the vertical distribution of the radiative heating rate perturbation within the stratosphere. Manabe and Strickler56 and Manabe and Wetherald57 analyzed radiative forcing effects on temperature from greenhouse gases in the Earth’s atmosphere, including examining feedbacks on clouds. They showed that changes in solar irradiance, albedo, and atmospheric distribution of certain radiatively active gases and aerosols could affect Earth’s climate. Since then, climate models show that the change in the equilibrium global mean surface air temperature varies roughly linearly with the global-mean radiative forcing at the tropopause and with changes in solar flux at the top of the atmosphere. The tropopause, which is the coldest part of the lower atmosphere, and the boundary for rapid mixing in the lower atmosphere, is roughly where radiation from the Earth escapes to space almost unimpeded. While the stratospheric ozone layer does attenuate the radiation, it responds rapidly to a forcing change at the tropopause, and this adjustment is accounted for. The L

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Wang et al.70 also found that one could use “effective CO2” in place of the specific climate forcers such as methane and nitrous oxide in climate models (as often used still) as a proxy to determine global average surface temperature (as long as the forcing was dominated by well-mixed gases). However, it was shown that it is not adequate to assess future climate changes on a regional scale. Wang et al.70 also emphasized the need for trace gases to be included in regional calculations. Cox et al.71 highlighted the cooling effects of regional anthropogenic aerosols and showed that they were “offsetting a substantial fraction of the global mean response to forcing due to greenhouse gases”. Cox et al.71 also showed that the hemispheric temperature response was considerably less than expected with the inclusion of aerosols, and that the regional forcing also demonstrated substantial differences between forcing and temperature response. It is apparent that the spatial and seasonal distributions of aerosol forcing need to be represented when examining regional climate responses unlike the case for global and annual mean values.71,72 The definition of radiative forcing, as used initially by the 1990 IPCC assessment,58d was the change in net irradiance (in Wm−2) at the tropopause when a greenhouse gas abundance changes after stratospheric temperatures are allowed to readjust to radiative equilibrium. In this case, the surface and tropospheric temperatures are held fixed at the values before the change in greenhouse gas abundance. Radiative forcing from different forcing agents was compared by assuming that the climate sensitivity factor was constant. In other words, a given value of radiative forcing produces the same change in globally averaged surface temperature. As noted above, this approach works well for well-mixed greenhouse gases, solar irradiance, surface albedo, and homogeneously distributed nonabsorbing aerosols as discussed in the 2001 IPCC report.73 However, global mean surface temperature may not respond linearly with RF at the tropopause for agents whose forcing is highly variable from the surface going up, for example, one that absorbs strongly near the surface but only weakly at the top of the troposphere. This relationship also breaks down if the forcing agent is not homogeneously distributed in the troposphere. For example, Hansen et al.66a found that the climate response is sensitive to the altitude and latitude of the forcing. A good example is the aircraft-induced changes in ozone and the effects of contrails that are more prevalent in the upper troposphere and, hence, have very different (even negative) climate sensitivities (see IPCC 1999 special report74). Indirect effects due to unevenly distributed aerosols also may have different climate sensitivities. The uneven distributions in stratospheric composition changes are not as important because the stratosphere’s temperature responds rapidly to forcing. A new concept, effective radiative forcing (ERF), has been developed to compare climate change forcings from various gases and particles (Forster et al.75 and Myhre et al.76). Here we summarize the description of the concept from the IPCC’s fifth Assessment Report.4 Realizing that not all forcings necessarily have the same efficiency or “efficacy” in causing climate to change, the use of ERF accounts better for the effects of efficacy as noted in the IPCC’s fifth Assessment Report.4 Unlike the case of radiative forcing where all surface and tropospheric conditions are assumed to be constant, ERF allows all physical variables, except the sea surface temperature and sea ice extent, to respond to the perturbations. Thus, ERF accounts for the rapid adjustments in the troposphere for changes that do respond rapidly to a change in forcing, e.g.,

change in clouds. Thus, it makes ERF a better indicator of the eventual temperature response, especially short-lived spatially nonuniform forcing agents such as aerosols. The ERF concept includes much of the relative efficacies of rapidly changing, inhomogeneously mixed, quick-responding forcing agents. It has been found that ERF leads to more uniform climate sensitivity across various forcing agents than the traditional RF concept (Myhre et al.76). However, the uncertainty range for ERF estimates tends to be larger than the range for RF estimates (Myhre et al.76), because the rapid adjustments differ across climate models. As to be expected, for well-mixed gases, there is no significant difference between RF and ERF. Many other metrics are based on the concept of radiative forcing (RF). RF has been commonly used to compare different forcing agents (e.g., emissions of gases and particles) affecting climate in many assessments of climate change.4,53,58b−e Radiative efficiency is a commonly used concept.1 Simply, it is a measure of the radiative forcing for a given change in mixing ratio of a greenhouse gas. This concept can be used to compare the relative effects on climate forcing from emissions of different gases. Once calculated for one molecule, radiative efficiencies of other similar molecules can be estimated. 3.2. Global Warming Potentials (GWPs)

Wuebbles, Rodhe, and Derwent58d developed the global warming potential (GWP) metric for the first IPCC assessment; the basic concept was based on the approach used in developing the ozone depletion potential concept by Wuebbles and several previously suggested concepts for GWP-like metrics proposed by Lashof and Ahuja,77 Rodhe,78 Wuebbles,79 and others. The concept developed for IPCC has been extensively utilized, discussed, and criticized ever since. Yet, it is widely used as an emissions metric for climate considerations and is still the general standard for metrics and their use in climate assessments.4,53,58b−e Despite various criticisms of the GWP metric (e.g., Wuebbles,80 Wuebbles et al.,81 Smith and Wigley,82 Fuglestvedt et al.,83 and Godal and Fuglestvedt84), it remains popular and will likely continue to be used in policy considerations. The United Nations Framework Convention on Climate Change (UNFCCC) has adopted GWPs in the Kyoto Protocol, and other government organizations such as the U.S. Environmental Protection Agency have used it in other policy considerations. For a variety of different greenhouse gases, GWPs provide an approach for quantifying the relative potential integrated forcing on climate. In IPCC and other climate assessments, GWPs are defined in terms of the time-integrated RF from the instantaneous release of a unit mass of a gas relative to RF of the same mass of the reference gas, generally taken as carbon dioxide. CO2 is used because it is the gas of most concern to anthropogenic forcing on climate. GWPs for different gases can be compared to evaluate their potential capability for affecting climate over a given time scale. Thus, per unit mass of emission, GWPs are an index for estimating the relative effects on climate for one greenhouse gas compared to another for a specified time period. As a result, GWPs are a better measure of the relative effects on climate than comparing RFs for different gases because GWPs differentiate between gases that would reside in the atmosphere for vastly different amounts of time, from a few days to many centuries, Figure 11. The GWP metric is thus based on the basic science underlying the greenhouse gas effects; however, it does not represent climate or biospheric M

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uncertainties in the derived RF per unit emission and in the assumed lifetime of the greenhouse gas. Efficacies that account for a correction to the forcing effect on climate can be also incorporated as a multiplier on the RF or by using ERF; the latter approach is generally thought to be better for emissions of short-lived gases that are not well-mixed in the atmosphere and have less well-defined forcings on climate. The ozone depletion potentials (ODPs) metric used in policies to protect stratospheric ozone, such as the Montreal Protocol, can be calculated to steady state. However, the AGWP for CO2 cannot be calculated to steady state. Because of its complex removal processes within the carbon cycle, atmospheric carbon dioxide is removed extremely slowly, with its removal from the atmosphere best represented as the sum of a series of exponential removal terms. As a result, GWPs are not calculated to steady state and are derived for selected integration times. However, the choice of integration time is arbitrary and is only loosely based on physical reasons attributed to the working of the Earth’s climate system, and, as a result, can be more a choice made in response to policy considerations. For most policy considerations, including the Kyoto Protocol, a 100-year integration time-horizon has been adopted. This choice allows all greenhouse gas emissions to be mapped into “CO2-equivalent emissions”. As a result, the 100-year time horizon is often reported for analyses of GWPs (1990 IPCC report,58d Wuebbles,80 and Myhre et al.76). This choice is entirely a value judgment that depends on the relative weight assigned to the climate effects of interest over different time scales. GWPs also depend on the background atmosphere assumed for the perturbation (which historically is taken to correspond to the current atmosphere), and on the way the analyses consider feedbacks and indirect effects (Myhre et al.76). Through their design, GWPs provide a relative measure of the total energy added to the climate system by the greenhouse gas under consideration relative to that added by CO2. Alternatively, GWP can be thought of as being approximately the ratio of the equilibrium temperature response from a sustained emission of the gas (or to the integrated temperature response for a pulse emission10,22a,52,76,86) relative to the similar expression for CO2 (this assumes efficacies are equivalent for CO2 and the gas being compared). However, GWPs are not equivalent to the temporal evolution of the temperature response or the response of other climate variables. Additional metrics have been proposed but have not yet been used for policy decisions. Of these, the most discussed alternative metric is global temperature change potential, also referred to as global temperature potential (GTP), which is discussed further in section 3.3. 3.2.1. CO2 Equivalent. One of the ways that GWPs are used is the concept of carbon dioxide equivalent, or CO2 equivalent, which is a metric used to compare the emissions from various greenhouse gases based upon their 100-year GWP. CO2 equivalents are commonly expressed as million metric tons of carbon dioxide equivalents (MMTCO2Eq). Multiplying the tons of the gas by the associated 100-year GWP derives the CO2 equivalent for a gas.

Figure 11. Decay in the radiative forcing from pulsed injections of three greenhouse gases as a function of time after injection. Red and green curves are for two greenhouse gases with atmospheric lifetimes of 15 and 1.5 years. The blue curve is the decay in radiative forcing following the injection of CO2. The lifetime of CO2 is complex with multiple components ranging in values from 5 to 200 years, and the figure used for the currently accepted value is based on carbon cycling modeling calculations discussed in 2013 IPCC assessment.4 The total energy input from the injection to the Earth system from this injection can be calculated for any time horizon from the area under the curve up to that time. The ratios of the areas under the curves for the greenhouse gas of interest and CO2 give the GWP of the greenhouse gas of interest. Reproduced with permission from ref 85. Copyright WMO/UNEP.

feedbacks, and it does not consider other impacts on the environment that could result from the emission of the gas. In practice, GWPs are determined by the following equation: H

GWP(H )i =

∫0 RF(i t )ci dt H

∫0 RFCO2(t )cCO2 dt

=

AGWPi AGWPCO2

(8)

In this equation, i is the gas (or other forcing agent) of interest, H is the time horizon for integration, RF is radiative forcing for i or for CO2, and c is the remaining mass of i or CO2 over time after the initial pulse emission. The figure shows a number of points. First, if the time horizon chosen is 20 years, then the areas under the curves up to those times give the GWP. This value would be much larger than what one would calculate if the integration time horizon is substantially longer. This is because the climatic effect of compound i is almost negligible after about 30 years while that of CO2 continues for a very long time. Second, the choice of CO2 as the reference, decided clearly for policy purposes as the most important anthropogenic greenhouse gas, is problematic because of its rather complex removal processes affecting the removal of CO2 from the atmosphere. Third, the RF that is shown is the direct effect of the GHG on the outgoing radiation. It does not include changes that the GHG emission causes on the composition of the atmosphere. AGWP (also used as a separate metric in the discussion below) is defined as the absolute global warming potential for a particular gas i or CO2. Also, climate sensitivity, the steady state effect on climate for a given forcing, is assumed to be the same for both i and CO2, and cancels out of the GWP equation (the traditional GWP definition assumes the same sensitivity factor, but this assumption can be relaxed to account for different climate sensitivities of different forcings if this were needed). The uncertainties in the GWP concept primarily depend on

MMTCO2 Eq = (million metric tons emission of a gas) × (GWP of the gas) N

(9) DOI: 10.1021/acs.chemrev.5b00010 Chem. Rev. XXXX, XXX, XXX−XXX

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In this way, emissions from different gases can be put onto a common scale (e.g., Shine87). The separate contributions from different emitted gases can then be summed to give the total CO2 equivalent warming and, hence, the total effect on Earth’s climate. Ideally, the determined CO2 equivalent emissions should lead to the same climate effects regardless of the mix of gases being emitted. However, the different gases can have different physical properties, and the choice of using the 100-year integrated GWP does not provide guaranteed equivalence (e.g., see 2013 IPCC report,4 Lauder et al.,88 O’Neill,89 Smith and Wigley,82b and Fuglestvedt et al.90). Therefore, care should be exercised in using these values for different purposes.

AGTPSx (t ) =

−

(13)

τ ≠ αx where α is the time constant for removal of the gas x, A is the RF for a 1 kg change in concentration of gas x, C is the heat capacity of the mixed-layer ocean, and τ is the time constant (λC) for the climate system. As in the derivation of AGWPs, the AGTPs for CO2 are more complicated because of the complexity of its removal processes. The resulting GTP as a function of time for a gas (or other forcing agent), x, is the ratio of AGTP for x divided by AGTP for CO2:

As an alternative to the GWP climate metric, Shine et al.10 proposed the global temperature change potentials (GTPs) metric. Other integrated temperature metrics (e.g., Rotmans and de Elzen91) have not gained wide acceptance. GTPs (Shine et al.,10,92) derive the relative temperature increase per unit mass of emission of a greenhouse gas relative to that for an equivalent mass of emitted CO2 for a chosen integrated time horizon. Following Shine et al.,10 either a pulse (labeled GTPp) or a sustained emission can be used for calculating GTPs. The GTP concept is based on the assumption that the global mean surface temperature can be defined by dΔT (t ) ΔT (t ) = ΔF(t ) − dt λ

⎫ ⎡ ⎛ t ⎞ ⎛ t ⎞ ⎤⎪ 1 ⎜ − ⎟ ⎥⎬ ⎢ exp exp − − ⎜ ⎟ ⎝ τ ⎠⎥⎦⎪ τ −1 − α −1 ⎢⎣ ⎝ αx ⎠ ⎭

for

3.3. Global Temperature Change Potentials (GTPs)

C

⎪ ⎡ αxAx ⎧ ⎛ t ⎞⎤ ⎨ τ ⎢1 − exp⎜ − ⎟⎥ ⎪ ⎣ ⎝ τ ⎠⎦ C ⎩

GTPSx (t ) =

AGTPSx (t ) AGTPCO2 S (t )

(14)

Like GWPs, GTP uses radiative forcing, but it calculates the response of the surface temperature to the radiative forcing for the emitted greenhouse gas. Similar to GWP, it is also a relative change for a gas compared to CO2. The GTP accounts for the thermal inertia and response of the climate system to provide a relative measure of the temperature responses for a specific time horizon, Figure 12. GTP is an end point measure by

(10)

or: ΔT (t ) =

1 C

∫0

t

⎛ t′ − t ⎞ ⎟ dt ′ ΔF(t ′)exp⎜ ⎝ λC ⎠

(11)

In this equation, the exponential term is an impulse response function to a forcing at some initial time t′, and where t is the time horizon chosen for evaluation. ΔT is the derived change in temperature as a function of time, ΔF is the change in RF, C is the heat capacity of the mixed-layer ocean, and λ is the (assumed) steady state climate sensitivity. In this equation, it is assumed that ocean and land respond together and at the same rate; they are together represented by a single heat capacity. This assumption allows the climate system to have a single time constant (thus greatly simplifying the calculation), rather than a slow time constant (ocean) and a fast time constant (land). The derived GTPs can be subject to assumptions and uncertainties associated both with the climate sensitivity and with the ocean heat uptake. For a known time-independent increase (or decrease) in the concentration (S) of a greenhouse gas, the concentration change over time is ⎡ ⎛ t ⎞⎤ ΔX(t ) = αΔS⎢1 − exp⎜ − ⎟⎥ ⎝ α ⎠⎦ ⎣

Figure 12. Temperature response following pulsed injections of three greenhouse gases. Lifetimes are the same as those in Figure 11. The GTP derivation is based on the temperature response at a selected year after pulse emission of the gas (20 or 100 year are shown (vertical lines)). Reproduced with permission from ref 85. Copyright WMO/ UNEP.

deriving the temperature change for a selected year. Like GWPs, the choice of integration time horizon has a strong effect on the derived GTPs. The GTP for various gases can be used to relatively weigh the effects of different emissions to obtain “CO2 equivalents”, akin to what was discussed for GWPs. As seen in Figures 11 and 12, the concepts for constructing GWPs and GTPs are fundamentally different. GTPs can account for physically based processes, like the climate sensitivity and the ocean−atmosphere exchange of heat, that cannot be done with GWPs. GTPs also can account for the slow response of the (deep) ocean, and the resulting effects on

(12)

For the forcing (F) given by AΔX(t), AGTPs (absolute GTP for a sustained emission change) at a particular integration time for a forcing x is derived by O

DOI: 10.1021/acs.chemrev.5b00010 Chem. Rev. XXXX, XXX, XXX−XXX

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Table 1. Atmospheric Lifetimes/Adjustment Times, Plus Values for 20 and 100 Years, and GTP Values for 20, 50, and 100 Years of Some Relevant Greenhouse Gasesa industrial designation or chemical name

chemical formula

carbon dioxide methane fossil methaned nitrous oxide

CO2 CH4 CH4 N2O

CFC-11 CFC-12 CFC-113 CFC-114 CFC-115 HCFC-22 HCFC-141b HCFC-142b HFC-23 HFC-32 HFC-125 HFC-134a HFC-143a HFC-152a HFC-227ea HFC-245fa methyl chloroform carbon tetrachloride methyl chloride methyl bromide halon-1202 halon-1211 halon-1301 halon-2402

CCl3F CCl2F2 CCl2FCClF2 CClF2CClF2 CClF2CF3 CHClF2 CH3CCl2F CH3CClF2 CHF3 CH2F2 CHF2CF3 CH2FCF3 CH3CF3 CH3CHF2 CF3CHFCF3 CHF2CH2CF3 CH3CCl3 CCl4 CH3Cl CH3Br CBr2F2 CBrClF2 CBrF3 CBrF2CBrF2

sulfur hexafluoride PFC-14 PFC-116

SF6 CF4 C2F6

lifetime (years)

GWP 20-yr

b 1 12.4c 84 12.4c 85 121c 264 Various Halocarbons 52 7090 102 10 800 93 6560 189 7710 540 5780 12 5310 9.4 2590 18 5140 228 10 800 5.4 2530 31 6280 14 3810 51 7050 1.6 545 36 5250 7.9 2980 4.8 555 26.0 3480 0.9 40 0.8 9 2.5 719 16 4590 72 7930 28 3920 Fully Fluorinated Species 3200.0 17 500 50 000.0 4880 10 000.0 8210

GWP 100-yr

GTP 20-yr

GTP 50-yr

GTP 100-yr

1 28 30 265

1 67 68 277

1 14 15 282

1 4 6 234

5160 10 300 6080 8580 7310 1780 800 2070 12 500 704 3450 1360 5080 148 3140 882 153 1730 11 2 196 1750 6670 2030

7160 11 300 6830 8180 6210 4230 1900 4530 11 500 1440 6040 3170 7110 191 5140 2040 298 3280 13 3 285 3950 8160 3730

5480 11 000 6510 9010 7500 847 285 1490 13 000 154 3350 771 5390 26 3180 259 32 1570 2