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10 THE ROLE OF TROPOSPHERIC OZONE IN CLIMATE 

CHANGE AND UV-B EFFECTS

10.1 Introduction

1 Atmospheric O3 plays an important role in the Earth’s energy budget by interacting with 

2 incoming solar radiation and outgoing infrared radiation. Over mid-latitudes, 

3 approximately 90% of the total atmospheric O3 column is located in the stratosphere (Kar 

4 et al., 2010; Crist et al., 1994). Therefore, tropospheric O3 makes up a relatively small 

5 portion (~10%) of the total column of O3 over mid-latitudes, but it does play an important 

6 role in the overall radiation budget. The next section (Section 10.2) briefly describes the 

7 physics of the earth’s radiation budget, providing background material for the subsequent 

8 two sections assessing how perturbations in tropospheric O3 might affect (1) climate 

9 through its role as a greenhouse gas (Section 10.3), and (2) health, ecology and welfare 

10 through its role in shielding the earth’s surface from solar ultraviolet radiation 

11 (Section 10.4). The concluding section in this chapter (Section 10.5) includes a summary 

12 of effects assessed in this chapter along with their associated causal determinations.

10.2 Physics of the Earth’s Radiation Budget

13 Radiant energy from the sun enters the atmosphere in a range of wavelengths, but peaks 

14 strongly in the visible (400-750 nm) part of the spectrum. Longer wavelength infrared 

15 (750 nm-1 mm) and shorter wavelength ultraviolet (100-400 nm) radiation are also 

16 present in the solar electromagnetic spectrum. Since the energy possessed by a photon is 

17 inversely proportional to its wavelength, infrared (IR) radiation carries the least energy 

18 per photon, and ultraviolet (UV) radiation carries the most energy per photon. UV 

19 radiation is further subdivided into classes based on wavelength: UV-A refers to 

20 wavelengths from 400-315 nm; UV-B from 315-280 nm; and UV-C from 280-100 nm. 

21 By the same argument above describing the relationship between photon wavelength and 

22 energy, UV-A radiation is the least energetic and UV-C is the most energetic band in the 

23 UV spectrum.

24 The wavelength of radiation also determines how the photons interact with the complex 

25 mixture of gases, clouds, and particles present in the atmosphere (see Figure 10-1). UV-A 

26 radiation can be scattered but is not absorbed to any meaningful degree by atmospheric 

27 gases including O3. UV-B radiation is absorbed and scattered in part within the 

28 atmosphere. UV-C is almost entirely blocked by the Earth’s upper atmosphere, where it 
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1 participates in photoionization and photodissociation processes including absorption by 

2 stratospheric O3.

3 Since UV-A radiation is less energetic and does not interact with O3 in the troposphere or 

4 the stratosphere and UV-C radiation is almost entirely blocked by stratospheric O3, UV-B 

5 radiation is the most important band to consider in relation to tropospheric O3 shielding. 

6 Furthermore, tropospheric O3 plays a “disproportionate” role in absorbing UV-B 

7 radiation compared with stratospheric O3 on a molecule per molecule basis (Balis et al., 

8 2002; Zerefos et al., 2002; Crist et al., 1994; Bruhl and Crutzen, 1989). This effect results 

9 from the higher atmospheric pressure present in the troposphere, resulting in higher 

10 concentrations of gas molecules present that can absorb or scatter radiation. For this 

11 reason, the troposphere is referred to as a “multiple scattering” regime for UV absorption, 

12 compared to the “single scattering” regime in the stratosphere. Thus, careful 

13 quantification of atmospheric absorbers and scatterers, along with a well-resolved 

14 description of the physics of these interactions, is necessary for predicting the effects of 

15 tropospheric O3 on UV-B flux at the surface.

16 Solar flux at all wavelengths has a temporal dependence, while radiative scattering and 

17 absorption have strong wavelength, path length, and gas/particle concentration 

18 dependencies. These combine to create nonlinear effects on UV flux at the Earth’s 

19 surface. Chapter 10 of the 2006 O3 AQCD (U.S. EPA, 2006b) describes in detail several 

20 key factors that influence the spatiotemporal distribution of ground-level UV radiation 

21 flux, including: (1) long-term solar activity including sunspot cycle; (2) solar rotation; (3) 

22 the position of the Earth in its orbit around the sun; (4) atmospheric absorption and 

23 scattering of UV radiation by gas molecules and aerosol particles; (5) absorption and 

24 scattering by stratospheric and tropospheric clouds; and (6) surface albedo. The 

25 efficiencies of absorption and scattering are highly dependent on the concentration of the 

26 scattering medium, particle size (for aerosols and clouds), and the altitude at which these 

27 processes are occurring. These properties are sensitive to meteorology, which introduces 

28 additional elements of spatial and temporal dependency in ground-level UV radiation 

29 flux.
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Source: 2006 O3 AQCD (U.S. EPA, 2006b).

Figure 10-1 Diagram of the factors that determine human exposure to 

ultraviolet radiation.

1 About 30% of incoming solar radiation is directly reflected back to space, mainly by 

2 clouds or surfaces with high albedo (reflectivity), such as snow, ice, and desert sand. 

3 Radiation that does penetrate to the Earth’s surface and is absorbed can be re-emitted in 

4 the longwave (infrared) portion of the spectrum; the rest goes into evaporating water or 

5 soil moisture or emerges as sensible heat. The troposphere is opaque to the outgoing 

6 longwave radiation. Polyatomic gases such as water vapor, CO2, CH4, and O3 absorb and 

7 re-emit the radiation upwelling from the Earth’s surface, reducing the efficiency with 

8 which that energy returns to space. In effect, these gases act as a blanket warming the 

9 Earth’s surface. This phenomenon, known as the “Greenhouse Effect,” was first 

10 quantified in the 19th century (Arrhenius, 1896), and gives rise to the term “greenhouse 

11 gas.” The most important greenhouse gas is water vapor.
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10.3 Effects of Tropospheric Ozone on Climate

10.3.1 Background

1 As a result of its interaction with incoming solar radiation and outgoing longwave 

2 radiation, tropospheric O3 plays a major role in determining climate, and increases in its 

3 abundance may contribute to climate change (IPCC, 2007c). Models estimate that the 

4 global average concentration of O3 in the troposphere has increased 30-70% since the 

5 preindustrial era (Gauss et al., 2006), while observations indicate that in some regions 

6 tropospheric O3 may have increased by factors as great as 4 or 5 (Marenco et al., 1994; 

7 Staehelin et al., 1994). These increases are tied to the rise in emissions of O3 precursors 

8 from human activity, mainly fossil fuel consumption and agricultural processes.

9 The effect on climate of the tropospheric O3 change since preindustrial times has been 

10 estimated to be about 25-40% of the anthropogenic CO2 effect and about 75% of the 

11 anthropogenic CH4 effect (IPCC, 2007c), ranking it third in importance behind these two 

12 major greenhouse gases. In the 21st century, as the Earth’s population continues to grow 

13 and energy technology spreads to developing countries, a further rise in the global 

14 concentration of tropospheric O3 is likely, with associated consequences for human health 

15 and ecosystems relating to climate change.

16 To examine the science of a changing climate and to provide balanced and rigorous 

17 information to policy makers, the World Meteorological Organization (WMO) and the 

18 United Nations Environment Programme (UNEP) formed the Intergovernmental Panel on 

19 Climate Change (IPCC) in 1988. The IPCC supports the work of the Conference of

20 Parties (COP) to the United Nations Framework Convention on Climate Change 

21 (UNFCCC). The IPCC periodically brings together climate scientists from member 

22 countries of WMO and the United Nations to review knowledge of the physical climate 

23 system, past and future climate change, and evidence of human-induced climate change. 

24 IPCC climate assessment reports are issued every five to seven years.

25 This section draws in part on the fourth IPCC Assessment Report (AR4) (IPCC, 2007c), 

26 as well as other peer-reviewed published research. Section 10.3.2 reviews evidence of 

27 climate change in the recent past and projections of future climate change. It also offers a 

28 brief comparison of tropospheric O3 relative to other greenhouse gases. Section 10.3.3

29 describes factors that influence the magnitude of tropospheric O3 effects on climate. 

30 Section 10.3.4 considers the competing effects of O3 precursors on climate. Finally, 

31 Section 10.3.5 and Section 10.3.6 describe the effects of changing tropospheric O3

32 concentrations on past and future climate. Downstream effects resulting from climate 

33 change, such as ecosystem responses, are outside the scope of this assessment, which 
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1 focuses rather on the effects of changes in tropospheric O3 concentrations on radiative 

2 forcing and climate.

10.3.2 Climate Change Evidence and the Influence of Tropospheric Ozone

10.3.2.1 Climate Change in the Recent Past

3 From the end of the Last Ice Age 12,000 years ago until the mid-1800s, observations 

4 from ice cores show that concentrations of the long-lived greenhouse gases CO2, CH4, 

5 and N2O have been relatively stable. Unlike these greenhouse gases, O3 is not preserved 

6 in ice, and no record of it before the late 1800s exists. Models, however, suggest that it, 

7 too, has remained relatively constant during this time period (Thompson et al., 1993; 

8 Thompson, 1992). The stable mix of these greenhouse gases in the atmosphere, together 

9 with water vapor, has kept the global mean temperature of the Earth close to 15oC. 

10 Without the presence of greenhouse gases in the atmosphere, the Earth’s global mean 

11 temperature would be about 30oC cooler, or -15oC.

12 Since the start of the Industrial Revolution, human activity has led to observable

13 increases of greenhouse gases in the atmosphere, mainly through fossil fuel combustion. 

14 According to the IPCC AR4 (IPCC, 2007c), there is now “very high confidence” that the 

15 net effect of anthropogenic greenhouse gas emissions since 1750 has led to warming, and 

16 it is “very likely” that human activity contributed to the 0.76oC rise in global mean 

17 temperature observed over the last century. The increase of tropospheric O3 may have 

18 contributed 0.1-0.3oC warming to the global climate during this time period (Hansen et 

19 al., 2005; Mickley et al., 2004). Global cooling due to anthropogenic aerosols (IPCC, 

20 2007c) has likely masked the full warming effect of the anthropogenic greenhouse gases 

21 on a global scale.

10.3.2.2 Projections of Future Climate Change

22 The IPCC AR4 projects a warming of ~0.2oC per decade for the remainder of the 21st 

23 century (IPCC, 2007c). Even at constant concentrations of greenhouse gases in the 

24 atmosphere, temperatures are expected to increase by about 0.1oC per decade, due to the 

25 slow response of oceans to the warming applied so far. It is likely that the Earth will 

26 experience longer and more frequent heat waves in the 21st century, together with more 

27 frequent droughts and/or heavy precipitation events in some regions, due to perturbations 

28 in the hydrological cycle that result from changing temperatures. Sea levels could 
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1 increase by 0.3-0.8 meters by 2300 due to thermal expansion of the oceans. The extent of 

2 Arctic sea ice is expected to decline, and contraction of the Greenland ice sheet could 

3 further contribute to the sea level rise (IPCC, 2007c).

4 Projections of future climate change are all associated with some degree of uncertainty. A 

5 major uncertainty involves future trends in the anthropogenic emissions of greenhouse 

6 gases or their precursors. For the IPCC AR4 climate projections, a set of distinct 

7 “storylines” or emission pathways was developed (IPCC, 2000). Each storyline took into 

8 account factors such as population growth, mix of energy technologies, and the sharing of 

9 technology between developed and developing nations, and each resulted in a different 

10 scenario for anthropogenic emissions. When these trends in emissions are applied to 

11 models, these scenarios yield a broad range of possible climate trajectories for the 21st 

12 century.

13 A second factor bringing large uncertainty to model projections of future climate is the 

14 representation of climate and, especially, climate feedbacks. A rise in surface 

15 temperatures would perturb a suite of other processes in the earth-atmosphere-ocean 

16 system, which may in turn either amplify the temperature increase (positive feedback) or 

17 diminish it (negative feedback). One important feedback involves the increase of water 

18 vapor content of the atmosphere that would accompany higher temperatures (Bony et al., 

19 2006). Water vapor is a potent greenhouse gas; accounting for the water vapor feedback 

20 may increase the climate sensitivity to a doubling of CO2 by nearly a factor of two (Held 

21 and Soden, 2000). The ice-albedo feedback is also strongly positive; a decline in snow 

22 cover and sea ice extent would diminish the Earth’s albedo, allowing more solar energy 

23 to be retained at the surface (Holland and Bitz, 2003; Rind et al., 1995). A final example 

24 of a climate feedback involves the effects of changing cloud cover in a warming 

25 atmosphere. Models disagree on the magnitude and even the sign of this feedback on 

26 surface temperatures (Soden and Held, 2006).

10.3.2.3 Metrics of Potential Climate Change

27 Two metrics frequently used to estimate the potential climate effect of some perturbation 

28 such as a change in greenhouse gas concentration are: (1) radiative forcing; and (2) global 

29 warming potential (GWP). These metrics differ in a fundamental way as described below.

30 Radiative forcing is a change in the radiative balance at a particular level of the 

31 atmosphere or at the surface when a perturbation is introduced in the earth-atmosphere32 ocean system. In the global mean, radiative forcing of greenhouse gases at the tropopause 

33 (top of the troposphere) is roughly proportional to the surface temperature response 

34 (Hansen et al., 2005; NRC, 2005). It thus provides a useful metric for policymakers for 
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1 assessing the response of the earth’s surface temperature to a given change in the 

2 concentration of a greenhouse gas. Positive values of radiative forcing indicate warming 

3 in a test case relative to the control; negative values indicate cooling. The units of 

radiative forcing are energy flux per area, or W/m2 4 .

5 Radiative forcing requires just a few model years to calculate, and it shows consistency 

6 from model to model. However, radiative forcing does not take into account the climate 

7 feedbacks that could amplify or dampen the actual surface temperature response, 

8 depending on region. Quantifying the change in surface temperature requires a climate 

9 simulation in which all important feedbacks are accounted for. As some of these 

10 processes are not well understood, the surface temperature response to a given radiative 

11 forcing can be highly uncertain and can vary greatly among models and even from region 

12 to region within the same model.

13 GWP indicates the integrated radiative forcing over a specified period (usually 100 years) 

14 from a unit mass pulse emission of a greenhouse gas or its precursor, and is reported as 

15 the magnitude of this radiative forcing relative to that of CO2. GWP is most useful for 

16 comparing the potential climate effects of long-lived gases, such as N2O or CH4. Since 

17 tropospheric O3 has a lifetime on the order of weeks to months, GWP is not seen as a 

18 valuable metric for quantifying the importance of O3 on climate (Forster et al., 2007). 

19 Thus, this assessment focuses on radiative forcing as the metric of climate influence 

20 resulting from changes in tropospheric O3.

10.3.2.4 Tropospheric Ozone as a Greenhouse Gas

21 Tropospheric O3 differs in important ways from other greenhouse gases. It is not emitted 

22 directly, but is produced through photochemical oxidation of CO, CH4, and nonmethane 

23 volatile organic compounds (VOCs) in the presence of nitrogen oxide radicals 

24 (NOX = NO + NO2; see Chapter 3, Section 3.2 for further details on the chemistry of O3

25 formation). It is also supplied by vertical transport from the stratosphere. The lifetime of 

26 O3 in the troposphere is typically a few weeks, resulting in an inhomogeneous 

27 distribution that varies seasonally; the distribution of the long-lived greenhouse gases like 

28 CO2 and CH4 are much more uniform. The longwave radiative forcing by O3 is mainly 

29 due to absorption in the 9.6 μm window, where absorption by water vapor is weak. It is 

30 therefore less sensitive to local humidity than the radiative forcing by CO2 or CH4, for 

31 which there is much more overlap with the water absorption bands (Lenoble, 1993). And 

32 unlike other major greenhouse gases, O3 absorbs in the shortwave as well as the 

33 longwave part of the spectrum.
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1 Figure 10-2 shows the main steps involved in the influence of tropospheric O3 on climate. 

2 Emissions of O3 precursors including CO, VOCs, CH4, and NOX lead to production of 

3 tropospheric O3. A change in the abundance of tropospheric O3 perturbs the radiative 

4 balance of the atmosphere, an effect quantified by the radiative forcing metric. The earth5 atmosphere-ocean system responds to the radiative forcing with a climate response, 

6 typically expressed as a change in surface temperature. Finally, the climate response 

7 causes downstream climate-related health and ecosystem effects, such as redistribution of 

8 diseases or ecosystem characteristics due to temperature changes. Feedbacks from both 

9 the climate response and downstream effects can, in turn, affect the abundance of 

10 tropospheric O3 and O3 precursors through multiple mechanisms. Direct feedbacks are 

11 discussed further in Section 10.3.3.4; the downstream climate effects and their feedbacks 

12 are extremely complex and outside the scope of this assessment.
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Note: Figure includes the relationship between precursor emissions, changes in tropospheric ozone abundance, radiative forcing, 

climate response, and climate effects. Units shown are those typical for each quantity illustrated. Feedbacks from both the climate 

response and climate effects can, in turn, affect the abundance of tropospheric ozone and ozone precursors through multiple 

feedback mechanisms. Climate effects and their feedbacks are deemphasized in the figure since these downstream effects are 

extremely complex and outside the scope of this assessment.

Figure 10-2 Schematic illustrating the effects of tropospheric ozone on climate.

The IPCC (2007c) reported a radiative forcing of 0.35 W/m2 1 for the change in 

2 tropospheric O3 since the preindustrial era, ranking it third in importance after the 

greenhouse gases CO2 (1.66 W/m2) and CH4 (0.48 W/m2 3 ). Figure 10-3 shows the global 

4 average radiative forcing estimates and uncertainty ranges in 2005 for anthropogenic 

5 CO2, CH4, O3 and other important agents and mechanisms. The error bars encompassing 

the tropospheric O3 radiative forcing estimate in the figure range from 0.25 to 0.65 W/m2 6 , 

7 making it relatively more uncertain than the long-lived greenhouse gases.
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Note: Figure shows the typical geographical extent (spatial scale) of the radiative forcing and the assessed level of scientific 

understanding (LOSU). The net anthropogenic radiative forcing and its range are also shown. These require summing asymmetric 

uncertainty estimates from the component terms, and cannot be obtained by simple addition. Additional radiative forcing factors not 

included here are considered to have a very low LOSU.

Source: Reprinted with permission of Cambridge University Press (IPCC, 2007c).

Figure 10-3 Global average radiative forcing (RF) estimates and uncertainty 

ranges in 2005 for anthropogenic CO2, CH4, ozone and other 

important agents and mechanisms.

10.3.3 Factors that Influence the Effect of Tropospheric Ozone on Climate

1 This section describes the main factors that influence the magnitude of the climate 

2 response to changes in tropospheric O3. They include: (1) trends in the concentration of 

3 tropospheric O3; (2) the effect of surface albedo on O3 radiative forcing; (3) the effect of 

4 vertical distribution on O3 radiative forcing; (4) feedback factors that can alter the climate 

5 response to O3 radiative forcing; and (5) the indirect effects of tropospheric O3 on the 

6 carbon cycle. Trends in stratospheric O3 may also affect temperatures at the Earth’s 

7 surface, but aside from issues relating to stratospheric-tropospheric exchange discussed in 

8 Chapter 3, Section 3.4.1.1, stratospheric O3 assessment is beyond the scope of this 

9 document.
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10.3.3.1 Trends in the Concentration of Tropospheric Ozone

1 To first order, the effect of tropospheric O3 on global surface temperature is proportional 

2 to the change in tropospheric O3 concentration. The earth’s surface temperatures are most 

3 sensitive to O3 perturbations in the mid to upper troposphere. This section therefore 

4 focuses mainly on observed O3 trends in the free troposphere or in regions far from O3

5 sources, where a change in O3 concentrations may indicate change throughout the 

6 troposphere. Data from ozonesondes, mountaintops, and remote surface sites are 

7 discussed, as well as satellite data.

Observed Trends in Ozone since the Preindustrial Era

8 Measurements of O3 at two European mountain sites dating from the late 1800s to early 

9 1900s show values at about 10 ppb, about one-fifth the values observed today at similar 

10 sites (Pavelin et al., 1999; Marenco et al., 1994). The accuracy of these early 

11 measurements is questionable however, in part because they exhibit O3 concentrations 

12 equivalent to or only a couple of parts per billion greater than those observed at nearby 

13 low-altitude sites during the same time period (Mickley et al., 2001; Volz and Kley, 

14 1988). A larger vertical gradient in tropospheric O3 would be expected because of its 

15 stratospheric source and its longer lifetime aloft. In another study, Staehelin et al. (1994)

16 revisited observations made in the Swiss mountains during the 1950s and found a 

17 doubling in O3 concentrations from that era to 1989-1991.

18 Routine observations of O3 in the troposphere began in the 1970s with the use of balloon19 borne ozonesondes, but even this record is sparse. Trends from ozonesondes have been 

20 highly variable and dependent on region (Logan et al., 1999). Over most sites in the U.S., 

21 ozonesondes reveal little trend. Over Canada, observations show a decline in O3 between 

22 1980 and 1990, then a rebound in the following decade (Tarasick et al., 2005). 

23 Ozonesondes over Europe give a mixed picture. European ozonesondes showed increases 

24 in the 1970s and 1980s, with smaller increases or even declines since then (Oltmans et 

25 al., 2006; Logan et al., 1999). Over Japan, O3 in the lower troposphere increased about 

26 0.2-0.4 ppb/year during the 1990s (Naja and Akimoto, 2004).

27 Ground-based measurements in remote regions provide a record of tropospheric O3, but 

28 like ozonesonde data are sparse before the 1970s. Springtime O3 observations from 

29 several mountain sites in the western U.S. show a positive trend of about of 0.5-

30 0.7 ppb/year since the 1980s (Cooper et al., 2010; Jaffe et al., 2003). Ship-borne O3

31 measurements for the time period 1977 to 2002 indicate increases of 0.1-0.7 ppb/year 

32 over much of the Atlantic south of 40oN, but no appreciable change north of 40oN 

33 (Lelieveld et al., 2004). The lack of trend for the North Atlantic would seem at odds with 
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1 O3 observations at Mace Head (53oN) on the west coast of Ireland, which show a 

2 significant positive trend of about 0.5 ppb/year from 1987 to 2003 (Simmonds et al., 

3 2004). Over Japan, O3 at a remote mountain site has increased 1 ppb/year from 1998 to 

4 2003 (Tanimoto, 2009), a rate more than double that recorded by ozonesondes in the 

5 lower troposphere over Japan during the 1990s (Naja and Akimoto, 2004). At Zugspitze, 

6 a mountain site in Germany, O3 increased by 12% per decade during the 1970s and 

7 1980s, consistent with European ozonesondes (Oltmans et al., 2006). Since then, O3

8 continues to increase at Zugspitze, but more slowly. What little data exist for the 

9 Southern Hemisphere point to measurable increases in tropospheric O3 in recent decades, 

10 as much as ~15% at Cape Grim in the 1989-2004 time period (Oltmans et al., 2006).

11 The satellite record is now approaching a length that can be useful for diagnosing trends 

12 in the total tropospheric O3 column (details on the use of satellites to measure 

13 tropospheric O3 are covered in Chapter 3, Section 3.5.5.5). In contrast to the surface data 

14 from ships, tropospheric O3 columns from the Total Ozone Mapping Spectrometer 

15 (TOMS) show no trend over the tropical Atlantic for the period 1980-1990 (Thompson 

16 and Hudson, 1999). Over the Pacific, a longer, 25 year record of TOMS data again 

17 reveals no trend over the tropics, but shows increases in tropospheric column O3 of about 

2-3 Dobson Units (DU)1 18 at mid-latitudes in both hemispheres (Ziemke et al., 2005).

19 Interpreting these recent trends in tropospheric O3 is challenging. The first difficulty is 

20 reconciling apparently contradictory trends in the observations, e.g., over tropical oceans. 

21 A second difficulty is that the O3 trends depend on several factors, not all of which can be 

22 well characterized. These factors include (1) trends in emissions of O3 precursors, 

23 (2) variation in the stratospheric source of O3, (3) changes in solar radiation resulting 

24 from stratospheric O3 depletion, and (4) trends in tropospheric temperatures (Fusco and 

25 Logan, 2003). Recent positive trends in the western U.S. and over Japan are consistent 

26 with the rapid increase in emissions of O3 precursors from mainland Asia and transport of 

27 pollution across the Pacific (Cooper et al., 2010; Tanimoto, 2009). The satellite trends 

28 over the northern mid-latitudes are consistent with this picture as well (Ziemke et al., 

29 2005). Increases in tropospheric O3 in the Southern Hemisphere are also likely due to 

30 increased anthropogenic NOX emissions, especially from biomass burning (Fishman et 

31 al., 1991). Recent declines in summertime O3 over Europe can be partly explained by 

32 decreases in O3 precursor emissions there (Jonson et al., 2005), while springtime 

33 increases at some European sites are likely linked to changes in stratospheric dynamics 

34 (Ordonez et al., 2007). Over Canada, Fusco and Logan (2003) found that O3 depletion in 

 1 The Dobson Unit is a typical unit of measure for the total O3 in a vertical column above the Earth’s surface. One DU is equivalent 

to the amount of O3 that would exist in a 1 μm (10–5 m) thick layer of pure O3 at standard temperature (0oC) and pressure (1 atm), 

and corresponds to a column of O3 containing 2.69 × 1020 molecules/m2

. A typical value for the amount of ozone in a column of the 

Earth’s atmosphere, although highly variable, is 300 DU and approximately 10% (30 DU) of that exists in the troposphere at mid 

latitudes.
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1 the lowermost stratosphere may have reduced the stratospheric flux of O3 into the 

2 troposphere by as much as 30% from the early 1970s to the mid 1990s, consistent with 

3 the trends in ozonesondes there.

Calculation of Ozone Trends for the Recent Past

4 Simulations of trends in tropospheric O3 provide a means for testing current knowledge 

5 of O3 processes and predicting with greater confidence trends in future O3 concentrations. 

6 Time-dependent emission inventories of O3 precursors have also been developed for 

7 1850-2000 (Lamarque et al., 2010) and for 1890-1990 (VanAardenne et al., 2001). These 

8 inventories allow for the calculation of changing O3 concentration over time.

9 One recent multi-model study calculated an increase in the O3 concentration since 

10 preindustrial times of 8-14 DU, or about 30-70% (Gauss et al., 2006). The large spread in 

11 modeled estimates reveals the limitations in knowledge of processes in the pristine 

12 atmosphere. Models typically overestimate the late nineteenth and early twentieth century 

13 observations available in surface air and at mountain sites by 50-100% (Lamarque et al., 

14 2005; Shindell et al., 2003; Mickley et al., 2001; Kiehl et al., 1999). Reconciling the 

15 differences between models and measurements will require more accurate simulation of 

16 the natural sources of O3 (Mickley et al., 2001) and/or implementation of novel sinks 

17 such as bromine radicals, which may reduce background O3 in the pristine atmosphere by 

18 as much as 30% (Yang et al., 2005c).

19 For the more recent past (since 1970), application of time-dependent emissions reveals an 

20 equatorward shift in the distribution of tropospheric O3 in the Northern Hemisphere due 

21 to the industrialization of societies at low-latitudes (Lamarque et al., 2005; Berntsen et 

22 al., 2000). By constraining a model with historical (1950s-2000) observations, Shindell 

23 and Faluvegi (2002) calculated a large increase of 8.2 DU in tropospheric O3 over 

24 polluted continental regions since 1950. This trend is not captured in standard chemistry 

25 models, but is consistent with the change in tropospheric O3 since preindustrial times 

26 implied by the observations from the late 1800s (Pavelin et al., 1999; Marenco et al., 

27 1994).

10.3.3.2 The Effect of Surface Albedo on Ozone Radiative Forcing

28 The Earth’s surface albedo plays a role in O3 radiative forcing. Through most of the 

29 troposphere, absorption of incoming shortwave solar radiation by O3 is small relative to 

30 its absorption of outgoing longwave terrestrial radiation. However, over surfaces 

31 characterized by high albedo (e.g., over snow, ice, or desert sand), incoming radiation is 
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1 more likely to be reflected than over darker surfaces, and the probability that O3 will 

2 absorb shortwave solar radiation is therefore larger. In other words, energy that would 

3 otherwise return to space may instead be retained in the atmosphere. Several studies have 

4 shown that transport of O3 to the Arctic from mid-latitudes leads to radiative forcing 

estimates greater than 1.0 W/m2 5 in the region, especially in summer (Shindell et al., 2006; 

6 Liao et al., 2004b; Mickley et al., 1999). Both the high surface albedo of the Arctic and 

7 the large solar zenith angles there (which increase the path length of incoming sunlight) 

8 lead to strong shortwave forcing in the region. Because the Arctic is especially sensitive 

9 to radiative forcing through the ice-albedo feedback, the large contribution in the 

10 shortwave solar spectrum to the total radiative forcing in the region may be important.

10.3.3.3 The Effect of Vertical Distribution on Ozone Radiative 

Forcing

11 In the absence of feedbacks, O3 increments near the tropopause produce the largest 

12 increases in surface temperature (Lacis et al., 1990; Wang et al., 1980). This is a result of 

13 the colder temperature of the tropopause relative to the rest of the troposphere and 

14 stratosphere. Since radiation emitted by the atmosphere is approximately proportional to 

the fourth power of its temperature1 15 , the colder the added O3 is relative to the earth’s 

16 surface, the weaker the radiation emitted and the greater the “trapping” of longwave 

17 radiation in the troposphere.

10.3.3.4 Feedback Factors that Alter the Climate Response to 

Changes in Ozone Radiative Forcing

18 Estimates of radiative forcing provide a first-order assessment of the effect of 

19 tropospheric O3 on climate. In the atmosphere, climate feedbacks and transport of heat 

20 alter the sensitivity of Earth’s surface temperature to addition of tropospheric O3. 

21 Assessment of the full climate response to increases in tropospheric O3 requires use of a 

22 climate model to simulate these interactions.

23 Due to its short lifetime, O3 is heterogeneously distributed through the troposphere. Sharp 

24 horizontal gradients exist in the radiative forcing of O3, with the greatest radiative forcing 

25 since preindustrial times occurring over the northern mid-latitudes (more on this in 

26 Section 10.3.5 and Section 10.3.6). If climate feedbacks are particularly powerful, they 

27 may obscure or even erase the correlation between regional radiative forcing and climate 

 1 As described by the Stefan-Boltzmann law, an ideal blackbody--which the atmosphere approximates--absorbs at all wavelengths

and re-radiates proportional to the fourth power of its temperature.
Draft – Do Not Cite or Quote 10-15 June 2012

1 response (Harvey, 2004; Boer and Yu, 2003). The transport of heat through the 

2 atmosphere, though not technically a feedback, may also weaken the correlation between 

3 forcing and climate response. Several model studies have reported that the horizontal 

4 pattern of surface temperature response from 2000-2100 trends in predicted short-lived 

5 species (including O3) closely matches the pattern from the trends in the long-lived 

6 greenhouse gases over the same time period (Levy et al., 2008; Shindell et al., 2008; 

7 Shindell et al., 2007). This correspondence occurs even though the patterns of radiative 

8 forcing for the short-lived and long-lived species differ substantially. In a separate paper, 

9 Shindell et al. (2007) found that Arctic temperatures are especially sensitive to the mid10 latitude radiative forcing from tropospheric O3.

11 Other studies have found that the signature of warming due to tropospheric O3 does show 

12 some consistency with the O3 radiative forcing. For example, Mickley et al. (2004)

13 examined the change in O3 since preindustrial times and found greater warming in the 

14 Northern Hemisphere than in the Southern Hemisphere (+0.4oC versus +0.2oC), as well 

15 as higher surface temperatures downwind of Europe and Asia and over the North 

16 American interior in summer. For an array of short-lived species including O3, Shindell 

17 and Faluvegi (2009) found that radiative forcing applied over northern mid-latitudes yield 

18 more localized responses due to local cloud, water vapor, and albedo feedbacks than 

19 radiative forcing applied over the tropics.

20 Climate feedbacks can also alter the sensitivity of surface temperature to the vertical 

21 distribution of tropospheric O3. The previous section (Section 10.3.3.3) described the 

22 greater effect of O3 added to the upper troposphere (near the tropopause) on radiative 

23 forcing, relative to additions in the mid- to lower troposphere. However, warming 

24 induced by increased O3 in the upper troposphere could stabilize the atmosphere to some 

25 extent, limiting the transport of heat to the Earth’s surface and mitigating the effect of the 

26 added O3 on surface temperature (Joshi et al., 2003; Christiansen, 1999). Hansen et al. 

27 (1997) determined that allowing cloud feedbacks in a climate model meant that O3

28 enhancements in the mid-troposphere had the greatest effect on surface temperature.

29 Finally, climate feedbacks can amplify or diminish the climate response of one 

30 greenhouse gas relative to another. For example, Mickley et al. (2004) found a greater 

31 temperature response to CO2 radiative forcing than to an O3 radiative forcing of similar 

32 global mean magnitude, due in part to the relatively weak ice-albedo feedback for O3. 

33 Since CO2 absorbs in the same bands as water vapor, CO2 radiative forcing saturates in 

34 the middle troposphere and is also shifted toward the drier poles. A poleward shift in 

35 radiative forcing amplifies the ice-albedo feedback in the case of CO2, and the greater 

36 mid-troposphere radiative forcing allows for greater surface temperature response, 

37 relative to that for O3.
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10.3.3.5 Indirect Effects of Tropospheric Ozone on the Carbon Cycle

1 A proposed indirect effect of tropospheric O3 on climate involves the carbon cycle. By 

2 directly damaging plant life in ways discussed in Chapter 9, increases in tropospheric O3

3 may depress the land-carbon sink of CO2, leading to accumulation of CO2 in the 

4 atmosphere and ultimately warming of the Earth’s surface. Sitch et al. (2007) calculated 

5 that this indirect warming effect of O3 on climate has about the same magnitude as the O3

6 direct effect. Their results suggest a doubled sensitivity of surface temperatures to O3

7 radiative forcing, compared to current model estimates.

10.3.4 Competing Effects of Ozone Precursors on Climate

8 Changes in O3 precursors can affect the radiative balance of the atmosphere through 

9 multiple (and sometimes competing) mechanisms. For example, the O3 precursor CH4 is 

10 itself a powerful greenhouse gas. O3 and its other precursors also exert a strong control on 

11 the oxidizing capacity of the troposphere, and so can affect the lifetime of gases such as 

12 CH4 (Derwent et al., 2001). For example, an increase in CO or VOCs would lead to a 

13 decrease in hydroxyl (OH) concentrations. Since OH is a major sink for CH4, a decline in 

14 OH would lengthen the CH4 lifetime, enhance the CH4 concentration, and amplify 

15 surface warming. A rise in NOX emissions, on the other hand, could lead to an increase in 

16 OH in certain locations, shortening the CH4 lifetime and causing surface cooling 

17 (Fuglestvedt et al., 1999). O3 can itself generate OH through (1) photolysis leading to 

18 excited oxygen atoms followed by reaction with water vapor and (2) reaction with HO2.

19 Figure 10-4 shows the radiative forcing associated with a suite of anthropogenic 

20 emissions, including O3 precursors (IPCC, 2007b). The emission-based radiative forcing 

for CH4, which includes the CH4 effect on O3 production, is +0.9 W/m2 21 , or nearly double 

22 that of the CH4 abundance-based radiative forcing shown in Figure 10-3. Figure 10-4 also 

shows a warming from anthropogenic CO and VOC emissions of +0.27 W/m2 23 and a net 

cooling of -0.21 W/m2 24 for NOX emissions. The net cooling for NOX occurs mainly due to 

25 the links between NOX and CH4. Consistent with these results, Shindell and Faluvegi 

(2009) calculated positive (+0.25 W/m2 26 ) radiative forcing from the increase in 

27 anthropogenic emissions of CO and VOCs since preindustrial times, as well as for CH4

(+1 W/m2). In contrast, Shindell and Faluvegi (2009) found negative (-0.29 W/m2 28 ) 

29 radiative forcing from anthropogenic emissions of NOX. Other studies have found a near 

30 cancellation of the positive O3 radiative forcing and the negative CH4 radiative forcing 

31 that arise from an incremental increase in anthropogenic NOX emissions (Naik et al., 

32 2005; Fiore et al., 2002; Fuglestvedt et al., 1999). The net effect of aircraft NOX on 

33 climate is especially complex (Isaksen et al., 2001; Wild et al., 2001). Stevenson (2004)
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1 calculated that aircraft NOX leads to short-term net warming via O3 production in the cool 

2 upper troposphere, but long-term net cooling because of CH4 loss.

3 OH production from O3 precursors can also affect regional sulfate air quality and climate 

4 forcing by increasing gas-phase oxidation rates of SO2. Using the A1B scenario in the 

5 IPCC AR4, Unger (2006) reported that by 2030, enhanced OH from the A1B O3

6 precursors may increase surface sulfate aerosol concentrations by up to 20% over India 

7 and China, relative to the present-day, with a corresponding increase in radiative cooling 

8 over these regions. In this way, O3 precursors may impose an indirect cooling via sulfate 

9 (Unger, 2006).

10 Taken together, these results point out the need for careful assessment of net radiative 

11 forcing involving multiple pollutants in developing climate change policy (Unger et al., 

12 2008). Many studies point to CH4 as a particularly attractive target for emissions control 

13 since CH4 is itself an important precursor of O3 (West et al., 2007; Fiore et al., 2002). 

14 Fiore et al. (2002) found that reducing anthropogenic CH4 emissions by 50% would lead 

to a global negative (-0.37 W/m2 15 ) radiative forcing, mostly from CH4. In later research, 

16 Fiore et al. (2008) reported that CH4 reductions would most strongly affect tropospheric 

17 O3 column amounts in regions of strong downwelling from the upper troposphere 

18 (e.g., around 30oN) and in regions of NOX-saturated conditions.

19 The magnitude of the radiative forcing from the change in tropospheric O3 since the 

20 preindustrial era is uncertain. This uncertainty derives in part from the scarcity of early 

21 measurements and in part from limited knowledge regarding processes in the natural 

atmosphere. As noted previously, the IPCC AR4 reports a radiative forcing of 0.35 W/m2 22

23 from the change in tropospheric O3 since 1750 (Forster et al., 2007), ranking it third in 

24 importance behind the greenhouse gases CO2 and CH4. The O3 radiative forcing could, in 

fact, be as large as 0.7 W/m2 25 , if reconstructions of preindustrial and mid-20th century O3

26 based on the measurement record are valid (Shindell and Faluvegi, 2002; Mickley et al., 

27 2001). In any event, Unger et al. (2010) showed that present-day O3 radiative forcing can 

28 be attributed to emissions from many economic sectors, including on-road vehicles, 

29 household biofuel, power generation, and biomass burning. As much as one-third of the

30 radiative forcing from the 1890 to 1990 change in tropospheric O3 could be due to 

31 increased biomass burning (Ito et al., 2007a).

32 These calculated radiative forcing estimates can be compared to those obtained from 

33 satellite data. Using data from TOMS, Worden et al. (2008) estimated a reduction in 

clear-sky outgoing longwave radiation of 0.48 W/m2 34 by O3 in the upper troposphere over 

35 oceans in 2006. This radiative forcing includes contributions from both anthropogenic 

36 and natural O3. Assuming that the concentration of O3 has roughly doubled since 

37 preindustrial times (Gauss et al., 2006), the total O3 radiative forcing estimated with 
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1 TOMS is consistent with that obtained from models estimating just the anthropogenic 

2 contribution.

Note: Values represent radiative forcing in 2005 due to emissions and changes since 1750. (S) and (T) next to gas species 

represent stratospheric and tropospheric changes, respectively.Source: Reprinted with permission of Cambridge University Press 

(IPCC, 2007b).

Figure 10-4 Components of radiative forcing for emissions of principal gases, 

aerosols, aerosol precursors, and other changes.
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10.3.5 Calculating Radiative Forcing and Climate Response to Past Trends 

in Tropospheric Ozone

1 Calculation of the climate response to the O3 radiative forcing is challenging due to 

2 complexity of feedbacks, as mentioned in Section 10.3.2.2 and Section 10.3.3.4. In their 

3 modeling study, Mickley et al. (2004) reported a global mean increase of 0.28oC since 

4 preindustrial times, with values as large as 0.8oC in continental interiors. For the time 

5 period since 1870, Hansen et al. (2005) estimated a much smaller increase in global mean 

6 surface temperature (0.11oC), but they implemented 1880s anthropogenic emissions in 

7 their base simulation and also took into account trends in both stratospheric and 

8 tropospheric O3,. The modeled decline of lower stratospheric O3, especially over polar 

9 regions, cooled surface temperatures in this study, counteracting the warming effect of 

10 increasing tropospheric O3.

11 Figure 10-5 shows the Hansen et al. (2005) results as reported in Shindell et al. (2006). In 

12 that figure, summertime O3 has the largest radiative effect over the continental interiors 

13 of the Northern Hemisphere. Shindell et al. (2006) estimated that the change in 

14 tropospheric O3 over the 20th century could have contributed about 0.3oC to annual mean 

15 Arctic warming and as much as 0.4-0.5oC during winter and spring. Over eastern China, 

16 Chang et al. (2009) calculated a surface temperature increase of 0.4oC to the 1970-2000 

17 change in tropospheric O3. It is not clear, however, to what degree regional changes in O3

18 concentration influenced this response, as opposed to more global changes.
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Note: Figure includes the input radiative forcing (W/m2), as computed by the NASA GISS chemistry-climate model. Values are 

surface temperature trends for the annual average (top left), June–August (bottom left), and December-February (bottom right) and 

annual average tropopause instantaneous radiative forcing from 1880 to 1990 (top right). Temperature trends greater than about 

0.1oC are significant over the oceans, while values greater than 0.3oC are typically significant over land, except for northern middle 

and high latitudes during winter where values in excess of about 0.5oC are significant. Values in the top right corner give areaweighted global averages in the same units as the plots.

Source: Reprinted with permission of American Geophysical Union (Shindell et al., 2006).

Figure 10-5 Ensemble average 1900-2000 radiative forcing and surface 

temperature trends (oC per century) in response to tropospheric 

ozone changes.

10.3.6 Calculating Radiative Forcing and Climate Response to Future 

Trends in Tropospheric Ozone

1 Future trends in tropospheric O3 concentrations depend in large part on what pathways in 

2 energy technology the world's societies will follow in coming decades. The trends in O3

3 will also depend on the changes in a suite of climate-sensitive factors, such as the water 

4 vapor content of the atmosphere. This section describes the following issues: 

5 (1) projected trends in the anthropogenic emissions of O3 precursors; (2) the effects of 

6 these emissions on the tropospheric O3 concentrations; (3) the effects of changing climate 
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on tropospheric O3; and (4) radiative forcing and climate response to 21st 1 century trends 

2 in tropospheric O3.

10.3.6.1 Emissions of Anthropogenic Ozone Precursors Across the 

21st Century

3 The IPCC SRES effort devised scenarios for short-lived O3 precursors as well as the 

4 well-mixed greenhouse gases including NOX, CO, and VOCs (IPCC, 2000). Using the 

5 IMAGE socioeconomic model, Streets et al. (2004) provided speciation for NOX and 

6 VOCs and allocated the trends in emissions over 17 regions and 8 economic sectors for 

7 the 2000-2050 time period. The worst-case IPCC scenario, A2, features continued 

8 dependence on fossil fuels, rapid population growth, and little sharing of technology 

9 between developed and developing nations. By 2100 in this scenario, global NOX , CO 

10 and CH4 emissions increase by a factor of 3.5, 2.6, and 2.9, respectively, relative to 2000 

11 (IPCC, 2000). Most of these increases in emissions occur over developing countries. For 

12 example over Asia, NOX emissions in the A2 scenario increase by more than a factor of 

13 four by 2100. The more moderate A1B scenario has global NOX and CO emissions 

14 increasing by 25% and 90%, respectively by 2100, but global CH4 emissions decreasing 

15 by 10%. In the B1 scenario, with its emphasis on clean and efficient technologies, global 

16 emissions of NOX, CO, and CH4 all decrease by 2100 relative to the present day (-40%, 

17 -60%, and -30%, respectively).

18 Other emissions scenarios have been recently developed to describe trends in the short19 term (up to 2030). The Current Legislation (CLE) scenario provides trends consistent 

20 with existing air quality regulations; the Maximum Feasible Reduction (MFR) scenario 

21 seeks to reduce emissions of O3 precursors to the maximum extent possible. Emission 

22 source changes relative to the present day for CLE, MFR, and A2 are given in Stevenson 

23 et al. (2006).

24 For the Fifth Assessment Report (IPCC AR5), a new set of climate futures has been 

25 developed: the Representative Concentration Pathways (RCPs) (Moss et al., 2010). The 

26 RCPs will explore for the first time approaches to climate change mitigation. The RCPs 

are designed to achieve radiative forcing targets of 2.6, 4.5, 6.0 and 8.5 W/m2 27 by 2100, 

28 and have been designated RCP 2.6, RCP 4.5, RCP 6.0, and RCP 8.5, respectively (RCP 

29 2.6 is also known as RCP3-PD.) The trends in O3 precursors for the RCP scenarios were 

30 determined by climate policies implicit in each scenario and by plausible assumptions 

31 regarding future air quality regulations. These scenarios were chosen to map the wide 

32 range of climate outcomes presented in the literature and represent only four of many 

33 possible scenarios that would lead to the specific radiative forcing targets; a wide range 
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1 of socioeconomic conditions could be consistent with each forcing pathway (Moss et al., 

2 2010). Therefore, they should not be interpreted as forecasts of future conditions, but 

3 rather as plausible climate and socio-economic futures.

4 Plots and comparisons of the RCP trends are available on the RCP website (RCP, 2009). 

In all RCPs, global anthropogenic NOX emissions decline 30-50% during the 21st 5 century, 

6 though RCP 8.5 shows a peak during the 2020s at a value ~15% greater than that of 

7 2000. Global anthropogenic VOC and CO emissions are relatively flat during the 2000-

8 2050 time range, and then decline by 30-50% by the end of the century. For CH4, global 

mean emission trends for the four RCP projections differ substantially across the 21st 9

10 century, with RCP 8.5 showing a tripling of emissions by 2100, and RCP 2.6 showing the 

11 emissions cut by half in this time range. RCP 4.5 and 6.0 show a peak in CH4 emissions 

12 in the middle of the century before dropping by the end of the century to just below 2000 

13 emission levels. All these global trends, however, contain some regional variation. For 

14 example, Asian emissions of both NOX and VOCs show large increases in the near term 

15 (2030s to 2050s).

10.3.6.2 Impact of 21st Century Trends in Emissions on 

Tropospheric Ozone

16 Due to its short lifetime, tropospheric O3 will respond readily to changes in 

17 anthropogenic emissions of its precursors. As shown in Table 10-1, a recent multi-model 

18 study found increases in the tropospheric O3 concentration of 15% and 6% for the IPCC 

19 A2 and CLE scenarios respectively for the 2000-2030 time period, and a decrease for the 

20 MFR scenario of 5% (Stevenson et al., 2006). These results indicate that the growth in 

21 tropospheric O3 between 2000 and 2030 could be reduced or even reversed, depending on 

22 emission controls. For the relatively moderate A1B emissions scenario over the 2000-

23 2050 time period, Wu et al. (2008a) calculated a change in O3 concentration of about 

24 20%.

25 As noted above, the RCP scenarios show large variations in their future projections of 

26 global mean CH4 emissions, but mainly declines in the emissions of the other O3

27 precursors across the 21st century. In one of the first efforts to assess the effect of these 

28 emission trends on global O3 abundances, Lamarque et al. (2011) found that the large 

29 CH4 increase in the RCP 8.5 scenario would drive a 15% enhancement of the 

30 tropospheric O3 burden by 2100, relative to the present-day, leading to a global mean 

radiative forcing of +0.2 W/m2 31 . By contrast, the global O3 burden would decrease in the 

other three RCPs, with declines in forcing ranging from -0.07 to -0.2 W/m2 32 .
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Table 10-1 2000-2030 changes in anthropogenic emissions, and CH4 and 

tropospheric ozone concentrations, and the associated 

tropospheric ozone forcing for three scenarios.

Scenario IPCC A2a Current Legislation (CLE)a Maximum Feasible Reduction 

(MFR)a

Percent change in NOX emissions +96% +18% -53%

Percent change in CO emissions +62% -16% -53%

Percent change in CH4 concentration +23% +19% 0%

Percent change in tropospheric O3

concentration

+15% +6% -5%

Radiative forcing due to O3 changeb 

(W/m2)

0.3 0.18 -0.05

a

Values are ensemble means.

b

Includes radiative forcing due to corresponding CH4 change. 

Source: Adapted from Stevenson et al. (2006).

10.3.6.3 Impact of 21st Century Climate on Tropospheric Ozone

1 For the time period from the 1800s to the present-day, most of the increase in the 

2 concentration of tropospheric O3 can be traced to changing emissions. Model studies 

3 show that climate change so far has likely had little effect on the tropospheric O3 (e.g., 

4 Grenfell et al., 2001). In the future, however, climate change is expected to bring large 

5 changes in a suite of variables that could affect O3 production, loss, and transport. For 

6 example, increased water vapor in a warming atmosphere is expected to enhance OH 

7 concentrations, which in remote, NOX-poor regions will accelerate O3 loss rates (Johnson 

8 et al., 1999).

9 In the 2050s A1B climate, Wu et al. (2008b) calculated a 5 ppb decrease in surface O3

10 over oceans. A rise in temperatures will also likely promote emissions of isoprene, an 

important biogenic precursor of O3. Model studies have calculated 21st 11 -century increases 

12 in isoprene emissions ranging from 25-50%, depending on climate scenario and time 

13 horizon (Wu et al., 2008a and references therein). These studies however did not take 

14 into account the effects of changing climate and CO2 concentration on vegetation extent, 

15 which could have large consequences for biogenic emissions (Heald et al., 2008; 

16 Sanderson et al., 2003). In any event, enhanced isoprene emissions will increase O3

17 concentrations in VOC-limited regions, but decrease O3 in NOX-limited regions (Wu et

18 al., 2008a; Pyle et al., 2007; Sanderson et al., 2003). Convection frequencies and 

19 lightning flash rates will also likely change in a changing climate, with consequences for 

20 lightning NOX emissions and O3 concentrations in the upper troposphere (Sinha and 

21 Toumi, 1997; Price and Rind, 1994). While Wu et al. (2008a) calculated an increase in 

22 lightning NOX by 2050 due to enhanced deep convection, Jacobson and Streets (2009)
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1 projected a decrease in lightning NOX due to a declining cloud ice in their future 

2 atmosphere. Finally, changes in transport processes will almost certainly accompany 

3 global climate change. For the 2050 A1B climate, Wu et al. (2008b) showed that 

4 flattening of the meridional temperature gradient in a warming world would lead to 

5 slower intercontinental transport of tropospheric O3. For the A2 climate in 2100, Zeng 

6 and Pyle (2003) projected an 80% increase in the flux of stratospheric O3 into the 

7 troposphere, relative to the present-day.

8 Taken together, these climate-driven processes could have appreciable effects on the 

9 concentration and distribution of tropospheric O3. As shown in Wu et al. (2008b), model 

10 projections of the change in O3 concentration due solely to future climate change range 

11 from -12% to +3%, depending on the model, scenario, and time horizon.

10.3.6.4 Radiative Forcing and Climate Response from 21st Century 

Trends in Tropospheric Ozone

12 In the near term (2000-2030), Stevenson et al. (2006)estimated an O3 forcing of near zero 

for MFR, 0.18 W/m2 for CLE, and +0.3 W/m2 13 for the A2 scenario (Table 10-1). Menon et 

14 al. (2008), following the moderate A1B scenario, calculated a radiative forcing of 

0.12 W/m2 15 from the 2000-2030 change in tropospheric O3, about the same as that derived 

16 by Stevenson et al. (2006)for the CLE scenario. Over the longer term (2000 to 2100) for 

17 the A1B scenario, Gauss et al. (2003)reported large positive radiative forcing (0.40 to 

0.78 W/m2 18 ) due to the change in tropospheric O3, as shown in Figure 10-6. Normalized 

19 radiative forcing for these model calculations fell within a relatively narrow range, 0.032 

to 0.040 W/m2 20 DU, indicating that the largest uncertainty lies in the model-calculated 

21 changes in O3 concentration. Applying the A2 scenario, Chen et al. (2007b) estimated a 

global mean radiative forcing of 0.65 W/m2 22 from tropospheric O3 by 2100, consistent 

23 with the Gauss et al. (2003) results. These studies took into account only the effect of 

24 changing emissions on tropospheric O3. In their calculations of the 2000-2100 radiative 

25 forcing from O3 in the A2 scenario, Liao et al. (2006) found that inclusion of climate 

26 effects on tropospheric O3 reduced their radiative forcing estimate by 20%.
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Note: Shown are the components of radiative forcing in W/m2. SW = shortwave component; LW = longwave component; Net = total 

forcing; t = tropospheric ozone changes only; and t + s = both tropospheric and stratospheric changes.

Source: Reprinted from Gauss et al. (2003), American Geophysical Union.

Figure 10-6 Global mean radiative forcing estimates calculated by a set of 

models for the 2000-2100 change in tropospheric ozone.

1 Several studies have included tropospheric O3 in their investigations of the response in 

2 the future atmosphere to a suite of short-lived species (e.g., Levy et al., 2008; Shindell et 

3 al., 2008; Shindell et al., 2007). Few studies, however, have calculated the climate 

4 response to changes in tropospheric O3 alone in the future atmosphere. For the A2 

5 atmosphere, Chen et al. (2007b) estimated a global mean surface temperature increase of 

6 +0.34oC by 2100 in response to the change in O3. The largest temperature increases in 

7 this study, as much as 5oC, occurred over the populous regions of Asia and the Middle 

8 East and downwind of biomass burning regions in South Africa and South America.
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10.4 UV-B Related Effects and Tropospheric Ozone

10.4.1 Background

1 UV radiation emitted from the Sun contains sufficient energy when it reaches the Earth to 

2 break (photolyze) chemical bonds in molecules, thereby leading to damaging effects on 

3 living organisms and materials. Atmospheric O3 plays a crucial role in reducing exposure 

4 to solar UV radiation at the Earth’s surface. Stratospheric O3 is responsible for the 

5 majority of this shielding effect, as approximately 90% of total atmospheric O3 is located 

6 there over mid-latitudes (Kar et al., 2010; Crist et al., 1994). Investigation of the 

7 supplemental shielding of UV-B radiation provided by tropospheric O3 is necessary for 

8 quantifying UV-B exposure and the incidence of related human health effects, ecosystem 

9 effects, and materials damage. The role of tropospheric O3 in shielding of UV-B radiation 

10 is discussed in this section.

10.4.2 Human Exposure and Susceptibility to Ultraviolet Radiation

11 The factors that potentially influence UV radiation exposure were discussed in detail in 

12 Chapter 10 of the 2006 O3 AQCD (U.S. EPA, 2006b) and are summarized here. These 

13 factors included outdoor activity, occupation, age, gender, geography, and protective 

14 behavior. Outdoor activity and occupation both influenced the amount of time people 

15 spend outdoors during daylight hours, the predominant factor for exposure to solar UV 

16 radiation. Age and gender were found to be factors that influence human exposure to UV 

17 radiation, particularly by influencing other factors of exposure such as outdoor activity 

18 and risk behavior. Studies indicated that females generally spent less time outdoors and, 

19 consequently, had lower UV radiation exposure on average compared to males. 

20 Geography influences the degree of solar UV flux to the surface, and hence exposure to 

21 UV radiation. Higher solar flux at lower latitudes increased the annual UV radiation dose 

22 for people living in southern states relative to northern states. Altitude was also found to 

23 influence personal exposure to UV radiation. Protective behaviors such as using 

24 sunscreen, wearing protective clothing, and spending time in shaded areas were shown to 

25 reduce exposure to UV radiation. Given these and other factors that potentially influence 

26 UV radiation exposure, the 2006 O3 AQCD (U.S. EPA, 2006b) listed the following 

27 subpopulations potentially at risk for higher exposures to UV radiation:

28  Individuals who engage in high-risk behavior (e.g., sunbathing);

29  Individuals who participate in outdoor sports and activities;
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1  Individuals who work outdoors with inadequate shade (e.g., farmers, 

2 construction workers, etc.); and

3  Individuals living in geographic areas with higher solar flux including lower 

4 latitudes (e.g., Honolulu, HI) and higher altitudes (e.g., Denver, CO).

5 The risks associated with all these factors are, of course, highly dependent on season and 

6 region (Sliney and Wengraitis, 2006).

10.4.3 Human Health Effects due to UV-B Radiation

7 Chapter 10 of the 2006 O3 AQCD (U.S. EPA, 2006b) covered in detail the human health 

8 effects associated with solar UV-B radiation exposure. These effects include erythema, 

9 skin cancer, ocular damage, and immune system suppression. These adverse effects, 

10 along with protective effects of UV radiation through increased production of vitamin D 

11 are summarized in this section. For additional details, the reader is referred to Chapter 10 

12 of the 2006 O3 AQCD (U.S. EPA, 2006b) and references therein.

13 The most conspicuous and well-recognized acute response to UV radiation is erythema, 

14 or the reddening of the skin. Erythema is likely caused by direct damage to DNA by UV 

15 radiation. Many studies discussed in Chapter 10 of the 2006 O3 AQCD (U.S. EPA, 

16 2006b) found skin type to be a significant risk factor for erythema. Skin cancer is another 

17 prevalent health effect associated with UV radiation. Exposure to UV radiation is 

18 considered to be a major risk factor for all forms of skin cancer. Ocular damage from UV 

19 radiation exposure includes effects on the cornea, lens, iris, and associated epithelial and 

20 conjunctival tissues. The region of the eye affected by exposure to UV radiation depends 

21 on the wavelength of the incident UV radiation. Depending on wavelength, common 

22 health effects associated with UV radiation include photokeratitis (snow blindness; short 

23 wavelengths) and cataracts (opacity of the lens; long wavelengths).

24 Experimental studies reviewed in Chapter 10 of the 2006 O3 AQCD (U.S. EPA, 2006b)

25 have shown that exposure to UV radiation may suppress local and systemic immune 

26 responses to a variety of antigens. Results from human clinical studies suggest that 

27 immune suppression induced by UV radiation may be a risk factor contributing to skin 

28 cancer induction. There is also evidence that UV radiation has indirect involvement in 

29 viral oncogenesis through the human papillomavirus, dermatomyositis, human 

30 immunodeficiency virus and other forms of immunosuppression.

31 A potential health benefit of increased UV-B exposure relates to the production of 

32 vitamin D in humans. Most humans depend on sun exposure to satisfy their requirements 

33 for vitamin D. Vitamin D deficiency can cause metabolic bone disease among children 
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1 and adults, and also may increase the risk of many common chronic diseases, including 

2 type I diabetes mellitus and rheumatoid arthritis. Substantial in vitro and toxicological 

3 evidence also support a role for vitamin D activity against the incidence or progression of 

4 various forms of cancer. In some studies, UV-B related production of vitamin D had 

5 potential beneficial immunomodulatory effects on multiple sclerosis, insulin-dependent 

6 diabetes mellitus, and rheumatoid arthritis. More details on UV-B protective studies are 

7 provided in Chapter 10 of the 2006 O3 AQCD (U.S. EPA, 2006b).

8 In establishing guidelines on limits of exposure to UV radiation, the International 

9 commission on Non-ionizing Radiation Protection (ICNIRP) agreed that some low-level 

10 exposure to UV radiation has health benefits (ICNIRP, 2004). However, the adverse 

11 health effects of higher UV exposures necessitated the development of exposure limits 

12 for UV radiation. The ICNIRP recognized the challenge in establishing exposure limits 

13 that would achieve a realistic balance between beneficial and adverse health effects. As 

14 concluded by ICNIRP (2004), “[t]he present understanding of injury mechanisms and 

15 long-term effects of exposure to [UV radiation] is incomplete, and awaits further 

16 research.”

10.4.4 Ecosystem and Materials Damage Effects Due to UV-B Radiation

17 A 2009 progress report on the environmental effects of O3 depletion from the UNEP, 

18 Environmental Effects Assessment Panel (UNEP, 2009) lists many ecosystem and 

19 materials damage effects from UV-B radiation. An in-depth assessment of the global 

20 ecosystem and materials damage effects from UV-B radiation per se is out of the scope of 

21 this assessment. However, a brief summary of some mid-latitude effects is provided in 

22 this section to provide context for UV-B related issues pertaining to tropospheric O3. The 

23 reader is referred to the UNEP report (UNEP, 2009) and references therein for further 

24 details. All of these UV-B related ecosystem and materials effects can also be influenced 

25 by climate change through temperature and other meteorological alterations, making 

26 quantifiable predictions of UV-B effects difficult.

27 Terrestrial ecosystem effects from increased UV-B radiation include reduced plant 

28 productivity and plant cover, changes in biodiversity, susceptibility to infection, and 

29 increases in natural UV protective responses. In general, however, these effects are small 

30 for moderate UV-B increases at mid-latitudes. A field study on wheat in southern Chile 

31 found no substantial changes in crop yield with moderate increases in UV-B radiation 

32 (Calderini et al., 2008). Similarly, field studies on silver birch (Betula pendula) in 

33 Finland found no measurable effects in photosynthetic function with increases in UV-B 

34 radiation (Aphalo et al., 2009). Subtle, but important, changes in habitat and biodiversity 
Draft – Do Not Cite or Quote 10-29 June 2012

1 have also been linked to increases in UV-B radiation (Mazza et al., 2010; Obara et al., 

2 2008; Wahl, 2008). Some plants have natural coping mechanisms for dealing with 

3 changes in UV-B radiation (Favory et al., 2009; Jenkins, 2009; Brown and Jenkins, 2008; 

4 Ioki et al., 2008), but these defenses may have costs in terms of reduced growth (Snell et 

5 al., 2009; Clarke and Robinson, 2008; Semerdjieva et al., 2003; Phoenix et al., 2000).

6 Aquatic ecosystem effects from increased UV-B radiation include sensitivity in 

7 growth, immune response, and behavioral patterns of aquatic organisms. One study 

8 looking at coccolithophores, an abundant phytoplankton group, found a 25% reduction in 

9 cellular growth with UV-B exposure (Gao et al., 2009a). Exposure to relevant levels of 

10 UV-B radiation has been shown to modify immune response, blood chemistry, and 

11 behavior in certain species of fish (Markkula et al., 2009; Holtby and Bothwell, 2008; 

12 Jokinen et al., 2008). Adverse effects on growth and development from UV-B radiation 

13 have also been observed for amphibians, sea urchins, mollusks, corals, and zooplankton 

14 (Garcia et al., 2009; Romansic et al., 2009; Croteau et al., 2008b; Croteau et al., 2008a; 

15 Marquis et al., 2008; Marquis and Miaud, 2008; Oromi et al., 2008). Increases in the flux 

16 of UV-B radiation may also result in an increase in the catalysis of trace metals including 

17 mercury, particularly in clear oligotrophic lakes with low levels of dissolved organic 

18 carbon to stop the penetration of UV-B radiation (Schindler et al., 1996). This could then 

19 alter the mobility of trace metals including the potential for increased mercury 

20 volatilization and transport within and among ecosystems.

21 Biogeochemical cycles, particularly the carbon cycle, can also be influenced by 

22 increased UV-B radiation. A study on high latitude wetlands found UV-induced increases 

23 in CO2 uptake through soil respiration (Haapala et al., 2009) while studies on arid 

24 terrestrial ecosystems found evidence for UV-induced release of CO2 through 

25 photodegradation of above-ground plant litter (Brandt et al., 2009; Henry et al., 2008; 

26 Caldwell et al., 2007; Zepp et al., 2007). Changes in solar UV radiation may also have 

27 effects on carbon cycling and CO2 uptake in the oceans (Brewer and Peltzer, 2009; 

28 Meador et al., 2009; Fritz et al., 2008; Zepp et al., 2008; Hader et al., 2007) as well as 

29 release of dissolved organic matter from sediment and algae (Mayer et al., 2009; 

30 Riggsbee et al., 2008). Additional studies showing effects on these and additional 

31 biogeochemical cycles including the water cycle and halocarbon cycle can be found in 

32 the UNEP report (UNEP, 2009) and references therein.

33 Materials damage from increased UV-B radiation include UV-induced 

34 photodegradation of wood (Kataoka et al., 2007) and plastics (Pickett et al., 2008). These 

35 studies and others summarizing photo-resistant coatings and materials designed to reduce 

36 photodegradation of materials are summarized in the UNEP report (UNEP, 2009) and 

37 references therein.
Draft – Do Not Cite or Quote 10-30 June 2012

10.4.5 UV-B Related Effects Associated with Changes in Tropospheric 

Ozone Concentrations

1 There are multiple complexities in attempting to quantify the relationship between 

2 changes in tropospheric O3 concentrations and UV-B exposure. The 2006 O3 AQCD 

3 (U.S. EPA, 2006b) described a handful of studies addressing this relationship, but none 

4 reported quantifiable effects of tropospheric O3 concentration fluctuations on UV-B 

5 exposure at the surface. Further quantifying the relationship between UV-B exposure and 

6 health or welfare effects is complicated by the uncertainties involved in the selection of 

7 an action spectrum and appropriate characterization of dose (e.g., peak or cumulative 

8 levels of exposure, timing of exposures, etc.) The lack of published studies that critically 

9 examined these issues together--that is the incremental health or welfare effects 

10 attributable specifically to UV-B changes resulting from changes in tropospheric O3

11 concentrations--lead to the prior conclusion that the effect of changes in surface-level O3

12 concentrations on UV-induced health outcomes could not be critically assessed within 

reasonable uncertainty (U.S. EPA, 2006b).

1 13

14 A recent study by Madronich et al. (2011) used CMAQ to estimate UV radiation 

15 response to changes in tropospheric O3 under different control scenarios projected out to 

16 2020. This study focused on southeastern U.S. and accounted for spatial and temporal 

17 variation in tropospheric O3 reductions, an important consideration since most controls 

18 are focused on reducing O3 in populated urban areas. The contrasting control strategies 

19 considered in this study included a historical scenario designed to meet an 84 ppb 8-h 

20 daily max standard and a reduced scenario designed to bring areas predicted to exceed a 

21 similarly designed 70 ppb standard into attainment. A biologically effective irradiance 

22 was estimated by multiplying the modeled UV irradiance by a sensitivity function (action 

23 spectrum) for the induction of nonmelanoma skin cancer in mice corrected for human 

24 skin transmission, then integrating over UV wavelengths. The average relative change in 

25 skin cancer-weighted surface UV radiation between the two scenarios was 0.11 ± 0.03% 

26 over June, July and August. Weighting by population, this estimate increased to 

27 0.19 ± 0.06%. Madronich et al. (2011) report that their estimated UV radiation increment 

28 is an order of magnitude less than that reported in an earlier study by Lutter and Wolz 

29 (1997) with the main reason for the discrepancy coming from the unrealistic uniform 

30 10 ppb reduction in O3 assumed in the former study. Madronich et al. (2011) did not 

31 attempt to link their predicted increase in UV radiation to a predicted increase in skin 

32 cancer incidence, however, due to several remaining and substantial uncertainties.

 1 The reader is referred to the U.S. EPA 2003 Final Response to Court Remand (U.S. EPA, 2003) for detailed discussions of the 

data and scientific issues associated with the determination of public health benefits resulting from the attenuation of UV-B by 

surface-level O3.
Draft – Do Not Cite or Quote 10-31 June 2012

1 Quantitatively estimating human health and welfare effects directly attributed to changes 

2 in UV-B penetration resulting from changes in ground-level O3 concentrations will 

3 require both (a) a solid understanding of the multiple factors that define the extent of 

4 exposure to UV-B, and (b) well-defined and quantifiable links between UV-B exposure 

5 and human disease and welfare effects. Detailed information does not exist regarding the 

6 relevant type (e.g., peak or cumulative) and time period (e.g., developmental, lifetime, or 

7 current) of exposure, wavelength dependency of biological responses, and 

8 inter-individual variability in UV resistance.

9 Although the UV-B related health effects attributed to marginal reductions in 

10 tropospheric or ground-level O3 have not been directly assessed to date, they would be 

11 expected to be small based on current information indicating a negligibly small effect of 

12 potential future changes in tropospheric O3 concentrations on ground-level UV-B 

13 radiation. In conclusion, the effect of changes in surface-level O3 concentrations on 

14 UV-induced health and welfare outcomes cannot yet be critically assessed within 

15 reasonable uncertainty.

10.5 Summary and Causal Determinations

10.5.1 Summary of the Effects of Tropospheric Ozone on Climate

16 Tropospheric O3 is a major greenhouse gas, third in importance after CO2 and CH4. While 

17 the developed world has successfully reduced emissions of O3 precursors in recent 

18 decades, many developing countries have experienced large increases in precursor 

19 emissions and these trends are expected to continue, at least in the near term. Projections 

20 of radiative forcing due to changing O3 over the 21st century show wide variation, due in 

21 large part to the uncertainty of future emissions of source gases. In the near-term (2000-

2030), projections of O3 radiative forcing range from near zero to +0.3 W/m2 22 , depending 

23 on the emissions scenario (Stevenson et al., 2006). Reduction of tropospheric O3

24 concentrations could therefore provide an important means to slow climate change in 

25 addition to the added benefit of improving surface air quality.

26 It is clear that increases in tropospheric O3 lead to warming. However the precursors of 

27 O3 also have competing effects on the greenhouse gas CH4, complicating emissions 

28 reduction strategies. A decrease in CO or VOC emissions would enhance OH 

29 concentrations, shortening the lifetime of CH4, while a decrease in NOX emissions could 

30 depress OH concentrations in certain regions and lengthen the CH4 lifetime.
Draft – Do Not Cite or Quote 10-32 June 2012

1 Abatement of CH4 emissions would likely provide the most straightforward means to 

2 address climate change since CH4 is itself an important precursor of background O3

3 (West et al., 2007; West et al., 2006; Fiore et al., 2002). A reduction of CH4 emissions 

4 would also improve air quality on its own right. A set of global abatement measures 

5 identified by West and Fiore (2005) could reduce CH4 emissions by 10% at a cost 

6 savings, decrease background O3 by about 1 ppb in the Northern Hemisphere summer, 

and lead to a global net cooling of 0.12 W/m2 7 . West et al. (2007) explored further the 

8 benefits of CH4 abatement, finding that a 20% reduction in global CH4 emissions would 

9 lead to greater cooling per unit reduction in surface O3, compared to 20% reductions in 

10 VOCs or CO.

11 Important uncertainties remain regarding the effect of tropospheric O3 on future climate 

12 change. To address these uncertainties, further research is needed to: (1) improve

13 knowledge of the natural atmosphere; (2) interpret observed trends of O3 in the free 

14 troposphere and remote regions; (3) improve understanding of the CH4 budget, especially 

15 emissions from wetlands and agricultural sources, (4) understand the relationship 

16 between regional O3 radiative forcing and regional climate change; and (5) determine the 

17 optimal mix of emissions reductions that would act to limit future climate change.

18 The effect of the tropospheric O3 change since preindustrial times on climate has been 

19 estimated to be about 25-40% of anthropogenic CO2 effect and about 75% of 

20 anthropogenic CH4 effect according to the IPCC. There are large uncertainties in the 

21 radiative forcing estimate attributed to tropospheric O3, making the effect of tropospheric 

22 O3 on climate more uncertain than the effect of the long-lived greenhouse gases. Overall, 

23 the evidence supports a causal relationship between changes in tropospheric O3

24 concentrations and radiative forcing.

25 Radiative forcing does not take into account the climate feedbacks that could amplify or 

26 dampen the actual surface temperature response. Quantifying the change in surface 

27 temperature requires a complex climate simulation in which all important feedbacks and 

28 interactions are accounted for. As these processes are not well understood or easily 

29 modeled, the surface temperature response to a given radiative forcing is highly uncertain 

30 and can vary greatly among models and from region to region within the same model. In 

31 light of these uncertainties, the evidence indicates that there is likely to be a causal 

32 relationship between changes in tropospheric O3 concentrations and effects on 

33 climate.
Draft – Do Not Cite or Quote 10-33 June 2012

10.5.2 Summary of UV-B Related Effects on Human Health, Ecosystems, 

and Materials Relating to Changes in Tropospheric Ozone 

Concentrations

1 UV radiation emitted from the Sun contains sufficient energy when it reaches the Earth to 

2 break (photolyze) chemical bonds in molecules, thereby leading to damaging effects on 

3 living organisms and materials. Atmospheric O3 plays a crucial role in reducing exposure 

4 to solar UV radiation at the Earth’s surface. Ozone in the stratosphere is responsible for 

5 the majority of this shielding effect, as approximately 90% of total atmospheric O3 is 

6 located there over mid-latitudes. Ozone in the troposphere provides supplemental 

7 shielding of radiation in the wavelength band from 280-315 nm, referred to as UV-B 

8 radiation. UV-B radiation has important effects on human health and ecosystems, and is 

9 associated with materials damage.

10 There is a lack of published studies that critically examine the incremental health or 

11 welfare effects (adverse or beneficial) attributable specifically to changes in UV-B 

12 exposure resulting from perturbations in tropospheric O3 concentrations. While the 

13 effects are expected to be small, they cannot yet be critically assessed within reasonable 

14 uncertainty. Overall, the evidence is inadequate to determine if a causal relationship 

15 exists between changes in tropospheric O3 concentrations and effects on health 

16 and welfare related to UV-B shielding.

10.5.3 Summary of Ozone Causal Determinations

17 The evidence reviewed in this chapter describes the recent findings regarding the climate 

18 and UV-B related effects of changes in tropospheric O3 concentrations. Table 10-2

19 provides an overview of the causal determinations for each of the categories evaluated.

Table 10-2 Summary of ozone causal determinations for climate and UV-B 

effects.

Effects Causal Determination

Radiative Forcing Causal relationship

Climate Change Likely to be a causal relationship

Health and Welfare Effects Related to UV-B 

Shielding

Inadequate to determine if a causal relationship 

exists
Draft – Do Not Cite or Quote 10-34 June 2012

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