Chapter 10. The Role of Tropospheric

Ozone in Climate Change and 

UV-B Effects
 

10.1. In tro d u c tio n 

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

2 incoming solar radiation and outgoing infrared radiation. Tropospheric O3 makes up only a small
 

3 portion of the total column of O3, but it has important incremental effects on the overall radiation 

4 budget. This chapter assesses the specific role that tropospheric O3 plays in the earth's radiation 

5 budget and how perturbations in tropospheric O3 might affect climate through its role as a
 

6 greenhouse gas (Section 10.2), and health, ecology and welfare through its role in shielding the
 

7 earth’s surface from solar ultraviolet radiation (Section 10.3).
 

10.2. Effects of Tropos pheric Ozone on Climate 

10.2.1. Background 

8 Tropospheric O3 is a major greenhouse gas, and increases in its abundance may contribute to 

9 climate change (IPCC, 2007, 092980). Models calculate that the global burden of tropospheric O3 

10 has doubled since the preindustrial era (Gauss et al., 2006, 630275), while observations indicate that 

11 in some regions O3 may have increased by factors as great as 4 or 5 (Marenco et al., 1994, 047733; 

12 Staehelin et al., 1994, 055369). These increases are tied to the rise in emissions of O3 precursors 

13 from human activity, mainly fossil fuel consumption and agricultural processes. The impact on 

14 climate of the O3 change since preindustrial times has been estimated to be about 25-40% of 

15 anthropogenic CO2 impact and about 75% of anthropogenic CH4 impact (IPCC, 2007, 092980), 

16 ranking it third in importance of the greenhouse gases. In the 21st century as the Earth’s population 

17 continues to grow and energy technology spreads to developing countries, a further rise in the global 

18 burden of tropospheric O3 is possible, with consequences for future climate. 

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

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

21 Nations Environment Programme (UNEP) formed the Intergovernmental Panel on Climate Change 

22 (IPCC) in 1988. The IPCC supports the work of the Conference of Parties (COP) to the United 

23 Nations Framework Convention on Climate Change (UNFCCC). The IPCC periodically brings 

Note: Hyperlinks to the reference citations throughout this document will take you to the NCEA HERO database (Health and 

Environmental Research Online) at http://epa.gov/hero. HERO is a database of scientific literature used by U.S. EPA in the process of 

developing science assessments such as the Integrated Science Assessments (ISAs) and the Integrated Risk Information System (IRIS). 

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1 together climate scientists from member countries of WMO and the United Nations to review 

2 knowledge of the physical climate system, past and future climate change, and evidence of human3 induced climate change. IPCC climate assessment reports are issued every 5 to 7 years. 

4 This chapter draws in part on the fourth IPCC Assessment Report (AR4) (IPCC, 2007, 

5 092980), as well as other peer-reviewed published research. Section 10.2.2 reviews the physics and 

6 chemistry of climate change and radiative forcing, together with evidence of climate change in the 

7 recent past and projections of future climate change. It also offers a brief description of tropospheric 

8 O3 as compared to other greenhouse gases. Section 10.2.3 describes factors that influence the 

9 magnitude of O3 effects of climate. Section 10.2.4 considers the competing effects of O3 precursors 

10 on climate. Sections 10.2.5 and 10.2.6 describe the effects of changing tropospheric O3 on present11 day and future climate, respectively. Finally, Section 10.2.7 presents a summary of the effects of 

12 tropospheric O3 on climate. 

10.2.2. Ph ys ic s a nd Ch e mis try o f Clima te Cha n ge an d Radia tive Fo rcing 

10.2.2.1. Phys ics of Greenhous e Gas es 

13 The Earth’s climate depends upon the flux of energy from the sun and its redistribution in the 

14 earth-atmosphere-ocean system. Radiant energy from the sun enters the atmosphere in a range of 

15 wavelengths, but peaks strongly in the shortwave (visible) part of the spectrum. Most solar energy at 

16 very short wavelengths (e.g., ultraviolet) is absorbed at high altitudes by gases such as stratospheric 

17 O3. About 30% of incoming solar radiation is reflected back to space, mainly by clouds or surfaces 

18 with high albedo (reflectivity), such as snow, ice, and desert sand. In the troposphere, gases and 

19 particles can interact with a fraction of the incoming solar radiation, but for the most part the 

20 troposphere is transparent to shortwave radiation. Thus about 70% of shortwave solar radiation 

21 penetrates to the Earth’s surface and is absorbed. About one-third of the absorbed energy is then re22 emitted in the longwave (infrared) portion of the spectrum. The rest goes into evaporating water or 

23 soil moisture or emerges as sensible heat. 

24 The troposphere is opaque to the outgoing longwave radiation. Polyatomic gases such as CO2, 

25 CH4, and O3 absorb and re-emit the radiation upwelling from the Earth’s surface, reducing the 

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

27 Earth's surface. This phenomenon, known as the “Greenhouse Effect,” was first quantified in the 19th 

28 century (Arrhenius, 1896, 043125), and gives rise to the term “greenhouse gas”. 

10.2.2.2. Clima te Change in the Recent Pas t 

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

30 cores show that concentrations of the long-lived greenhouse gases CO2, CH4, and N2O have been 

31 relatively stable. Unlike these greenhouse gases, O3 is not preserved in ice, and no record of it before 

32 the late 1800s exists. Models, however, suggest that it, too, has remained relatively constant during 

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1 this time period (Thompson, 1992, 043463; Thompson et al., 1993, 029510). The stable mix of 

2 greenhouse gases in the atmosphere has kept the global mean temperature of the Earth close to 15oC. 

3 Without the presence of greenhouse gases in the atmosphere, the Earth’s temperature would be about 

4 30oC cooler, or -15oC. Since the start of the Industrial Revolution, human activity has led to 

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

6 According to the IPCC AR4 (IPCC, 2007, 092980), we now have “very high confidence” that the net 

7 effect of anthropogenic greenhouse gas emissions since 1750 has led to warming, and it is “very 

8 likely” that human activity contributed to the 0.76oC rise in global mean temperature observed over 

9 the last century. The increase of tropospheric O3 may have contributed 0.1-0.3oC warming to the 

global climate during this time period (Hansen et al., 2005, 190596; Mickley et al., 2004, 057416). 

11 Global cooling due to anthropogenic aerosols (IPCC, 2007, 092980) has likely masked the full 

12 warming effect of the anthropogenic greenhouse gases. 

10.2.2.3. Proje c tions of Fu ture Clima te Cha nge 

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

14 (IPCC, 2007, 092980). Even at constant concentrations of greenhouse gases in the atmosphere, 

temperatures are expected to increase by about 0.1oC per decade, due to the slow response of oceans 

16 to the warming applied so far. It is likely that the Earth will experience longer and more frequent 

17 heat waves in the 21st century, together with more frequent droughts and/or heavy precipitation 

18 events in some regions, due to perturbations in the hydrological cycle that result from changing 

19 temperatures (IPCC, 2007, 092980). Sea levels could increase by 0.3-0.8 m by 2300 due to thermal 

expansion of the oceans. The extent of Arctic sea ice is expected to decline, and contraction of the 

21 Greenland ice sheet could further contribute to the sea level rise (IPCC, 2007, 092980). 

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

23 major uncertainty involves future trends in the anthropogenic emissions of greenhouse gases or their 

24 precursors. For the IPCC AR4 climate projections, a set of distinct “storylines” or emission pathways 

was developed (IPCC, 2000, 080704). Each storyline took into account factors such as population 

26 growth, mix of energy technologies, and the sharing of technology between developed and 

27 developing nations, and each resulted in a different scenario for anthropogenic emissions. When 

28 these trends in emissions are applied to models, these scenarios yield a broad range of possible 

29 climate trajectories for the 21st century. 

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

31 representation of climate and, especially, climate feedbacks. A rise in surface temperatures would 

32 perturb a suite of other processes in the earth-atmosphere-ocean system, which may in turn either 

33 amplify the temperature increase (positive feedback) or diminish it (negative feedback). One 

34 important feedback involves the increase of water vapor content of the atmosphere that would 

accompany higher temperatures (Bony et al., 2006, 630272). Water vapor is a potent greenhouse gas; 

36 accounting for the water vapor feedback may increase the climate sensitivity to a doubling of CO2 by 

37 nearly a factor of two (Held and Soden, 2000, 630279). The ice-albedo feedback is also strongly 

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1	 positive; a decline in snow cover and sea ice extent would diminish the Earth’s albedo, allowing 

2	 more solar energy to be deposited to the surface (Holland and Bitz, 2003, 630280; Rind et al., 1995, 

3	 630285). A final example of a climate feedback involves the effects of changing cloud cover in a 

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

surface temperatures (Soden and Held, 2006, 631183).
 

10.2.2.4. Me tric s of Po te n tia l Clima te Cha nge 

6 Two different metrics are frequently used to estimate the potential climate impact of some 

7 perturbation such as a change in greenhouse gas concentration: (1) global warming potential (GWP); 

8 and (2) radiative forcing (RF). 

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

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

11 magnitude of this forcing relative to that of CO2. GWP is most useful for comparing the potential 

12 climate impacts of long-lived gases, such as N2O or CH4. Since tropospheric O3 has a lifetime on the 

13 order of weeks to months, GWP is not seen as a valuable metric for quantifying the importance of O3 

14 on climate (Forster et al., 2007, 092936). 

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

16 at the surface when a perturbation is introduced in the earth-atmosphere-ocean system. In the global 

17 mean, radiative forcing of greenhouse gases at the tropopause (top of the troposphere) is roughly 

18 proportional to the surface temperature response 

19 (Committee on Radiative Forcing Effects on Climate; Climate Research Committee; National 

Research Council et al., 2005, 057409; Hansen et al., 2005, 190596). It thus provides a useful metric 

21 for policymakers for assessing the response of the earth's surface temperature to a given change in 

22 the concentration of greenhouse gas. Positive values of radiative forcing indicate warming in a test 

23 case relative to the control; negative values indicate cooling. The units of radiative forcing are 

24 energy flux per area, or W/m2. 

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

26 model to model. However, radiative forcing does not take into account the climate feedbacks that 

27 could amplify or dampen the actual surface temperature response, depending on region (Section 

28 10.2.2.3). Quantifying the change in surface temperature requires a climate simulation in which all 

29 important feedbacks are accounted for. As these processes are not well understood, the surface 

temperature response to a given radiative forcing is highly uncertain and can vary greatly among 

31 models and even from region to region within the same model. 

10.2.2.5. Tro pos pheric Ozone as a Greenhouse Gas 

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

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

34 organic compounds (VOCs) in the presence of nitrogen oxide radicals (NOX = NO + NO2; see 

Section 3.2). It is also supplied by vertical transport from the stratosphere. The lifetime of O3 in the 

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1 troposphere is typically a few weeks, resulting in an inhomogeneous distribution that varies 

2 seasonally; the distribution of the long-lived greenhouse gases like CO2 and CH4 are much more 

3 uniform. The longwave radiative forcing by O3 is mainly due to absorption in the 9.6 μm window, 

4 where absorption by water vapor is weak. It is therefore less sensitive to local humidity than the 

5 radiative forcing by CO2 or CH4, for which there is much more overlap with the water absorption 

6 bands (Lenoble, 1993, 630969). And unlike other major greenhouse gases, O3 absorbs in the 

7 shortwave as well as the longwave part of the spectrum. 

8 Figure 10-1 shows the main steps involved in the influence of tropospheric O3 on climate. An 

9 increase in the emissions of O3 precursors leads to an increase in the burden of tropospheric O3. The 

10 added O3 then perturbs the radiative balance of the atmosphere, an effect quantified by the radiative 

11 forcing metric. This forcing results in climate change, usually expressed as a change in surface 

12 temperature. Climate change can also perturb tropospheric O3, as will be discussed in Section 

13 10.2.6.3. As shown in Figure 10-2, the IPCC (IPCC, 2007, 092980) reports a radiative forcing of 

14 0.35 W/m2 for the change in tropospheric O3 since the preindustrial era, ranking it third in 

15 importance after the greenhouse gases CO2 (1.66 W/m2) and CH4 (0.48 W/m2). The error bars 

16 encompassing the tropospheric O3 radiative forcing estimate range from 0.25 to 0.65 W/m2, making 

17 it relatively more uncertain than the long-lived greenhouse gases. 

Figure 10-1. Flow chart for the effects of tropospheric ozone on climate. ﷒Emissions of the ozone 

precursors CO, VOCs, CH4, and NOX lead to production of tropospheric ozone. A

change in the burden of tropospheric ozone perturbs the radiative balance of the 

atmosphere, leading to radiative forcing. The earth-atmosphere-ocean system

responds to the forcing with a change in climate. Climate change, in turn, can affect

the abundance of tropospheric ozone through multiple mechanisms. Units shown 

are those typical for each quantity, with the climate response expressed as a change

in surface temperature. 

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Source: Used with permission from Cambridge University Press, IPCC (IPCC, 2007, 092980) 

Figure 10-2. Global average radiative forcing (RF) estimates and ranges in 2005 for 

anthropogenic CO2, CH4, ozone and other important agents and mechanisms, 

together with the typical geographical extent (spatial scale) of the 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 forcing factors not included here are

considered to have a very low LOSU. 

10.2.3. Fa c to rs th at Influe n ce th e Effec t of Tro p os ph e ric Ozone on Climate 

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

2 changes in tropospheric O3. They include: (1) trends in the burden of tropospheric O3; (2) the effect 

3 of surface albedo on O3 forcing; (3) the effect of vertical distribution on O3 forcing; (4) feedback 

4 factors that can alter the climate response to O3 forcing; and (5) the indirect effects of tropospheric 

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

6 surface, but that topic is beyond the scope of this assessment. 

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10.2.3.1. Tre n ds in the Burde n of Tro p os p he ric Ozone 

1 To first order, the effect of tropospheric O3 on climate is proportional to the change in O3 

2 burden. The earth's surface temperatures are most sensitive to O3 perturbations in the mid to upper 

3 troposphere. This section therefore focuses mainly on observed O3 trends in the free troposphere or 

4 in regions far from O3 sources, where a change in O3 concentrations may indicate change throughout 

the troposphere. Data from ozonesondes, mountaintops, and remote surface sites are discussed, as 

6 well as satellite data. 

Observed Trends in Ozone Since the Preindustrial Era 

7 Measurements of O3 at two European mountain sites dating from the late 1800s to early 1900s 

8 show values at about 10 ppb, about one-fifth the values observed today at similar sites (Marenco et 

9 al., 1994, 047733; Pavelin et al., 1999, 087296). The accuracy of these early measurements is 

questionable however, in part because they exhibit O3 concentrations equivalent to or only a couple 

11 of parts per billion greater than those observed at nearby low-altitude sites during the same time 

12 period (Mickley et al., 2001, 080134; Volz and Kley, 1988, 041650). A larger vertical gradient in 

13 tropospheric O3 would be expected because of its stratospheric source and its longer lifetime aloft. In 

14 another study, Staehelin et al. (1994, 055369) revisited observations made in the Swiss mountains 

during the 1950s and found a doubling in O3 concentrations from that era to 1989-1991. 

16 Routine observations of O3 in the troposphere began in the 1970s with the use of balloon17 borne ozonesondes, but even this record is sparse. Trends from ozonesondes have been highly 

18 variable and dependent on region (Logan et al., 1999, 631175). Over most sites in the U.S., 

19 ozonesondes reveal little trend. Over Canada, observations show a decline in O3 between 1980 and 

1990, then a rebound in the following decade (Tarasick et al., 2005, 631184). Ozonesondes over 

21 Europe give a mixed picture, with Hohenpeissenberg in Germany showing declines through the 

22 troposphere in recent decades, while Zugspitze, also in Germany, exhibiting small increases 

23 (Oltmans et al., 2006, 180188). Over Japan, O3 in the lower troposphere increased about 0.2­

24	 0.4 ppb/y during the 1990s (Naja and Akimoto, 2004, 631178). 

Ground-based measurements in remote regions provide a record of background tropospheric 

26 O3 extending as far back as the 1980s or, for ship measurements, the late 1970s. Springtime O3 

27 observations from several mountain sites in the western U.S. show a positive trend of about of 0.5­

28 0.7 ppb/y since the 1980s (Cooper et al., 2010, 380093; Jaffe et al., 2003, 052229). Ship-borne O3 

29 measurements for the time period 1977 to 2002 indicate increases of 0.1-0.7 ppb/y over the tropical 

and South Atlantic, but no significant change over the North Atlantic (Lelieveld et al., 2004, 

31 630578). The lack of trend for the North Atlantic would seem at odds with O3 observations at Mace 

32 Head on the west coast of Ireland, which show a significant positive trend of about 0.5 ppb/y from 

33 1987 to 2003 (Simmonds et al., 2004, 631182). Over Japan, O3 at a remote mountain site has 

34 increased 1 ppb/y from 1998 to 2003 (Tanimoto, 2009, 620751), a rate more than double that 

recorded by ozonesondes in the lower troposphere over Japan during the 1990s (Naja and Akimoto, 

36 2004, 631178). 

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

2 total column amounts of tropospheric O3. In contrast to the surface data from ships, tropospheric O3 

3 columns from the Total Ozone Mapping Spectrometer (TOMS) show no trend over the tropical 

4 Atlantic for the period 1980-1990 (Thompson and Hudson, 1999, 631185). Over the Pacific, a 

longer, 25 year record of TOMS data again reveals no trend over the tropics, but shows increases in 

6 tropospheric column O3 of about 2-3 Dobson Units (DU [1 DU = 2.69 × 1016 molecules of O3/cm2]) 

7 at midlatitudes in both hemispheres (Ziemke et al., 2005, 631193); for comparison, the tropospheric 

8 O3 burden averages about 34 DU. 

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

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

11 difficulty is that the O3 trends depend on several factors, not all of which can be well characterized. 

12 These factors include (1) trends in emissions of O3 precursors, (2) variation in the stratospheric 

13 source of O3, (3) changes in solar radiation resulting from stratospheric O3 depletion, and (4) trends 

14 in tropospheric temperatures (Fusco and Logan, 2003, 051229). The positive trends in the western 

U.S. and over Japan are consistent with the rapid increase in emissions of O3 precursors from 

16 mainland Asia and transport of pollution across the Pacific (Cooper et al., 2010, 380093; Tanimoto, 

17 2009, 620751). The satellite trends over the northern mid-latitudes are consistent with this picture as 

18 well (Ziemke et al., 2005, 631193). Increases in tropospheric O3 in the Southern Hemisphere are also 

19 likely due to increased anthropogenic NOX emissions, especially from biomass burning. The declines 

in O3 over Europe can be at least partly explained by decreases in O3 precursor emissions there 

21 (Jonson et al., 2005, 630282), though recent O3 depletion in the lower stratosphere may also 

22 contribute to the decreases by reducing stratospheric input to the troposphere (Fusco and Logan, 

23 2003, 051229). 

Calculation of Ozone Trends for the Recent Past 

24 Attempts to simulate trends in tropospheric O3 allow us to test current knowledge of O3 

processes and to predict with greater confidence trends in future O3 concentrations. Time-dependent 

26 emission inventories of O3 precursors have also been developed (e.g., Lamarque et al., 2010, 630289 

27 for 1850-2000; Van Aardenne et al., 2001, 055564 for 1890-1990). These inventories allow for the 

28 calculation of changing O3 burden over time. 

29 One recent multi-model study calculated an increase in the O3 burden since preindustrial times 

of 8-14 DU, or about 30-70% (Gauss et al., 2006, 630275). The large spread in modeled estimates 

31 reveals our limited knowledge of processes in the pristine atmosphere. Models typically overestimate 

32 the late nineteenth and early twentieth century observations available in surface air and at mountain 

33 sites by 50-100% (Kiehl et al., 1999, 047917) (Lamarque et al., 2005, 630287; Mickley et al., 2001, 

34 080134; Shindell et al., 2003, 057417). Reconciling the differences between models and 

measurements will require more accurate simulation of the natural sources of O3 (Mickley et al., 

36 2001, 080134) and/or implementation of novel sinks such as bromine radicals, which may reduce 

37 background O3 in the pristine atmosphere by as much as 30% (Parrella et al., In Press, 664506). 

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

2 equatorward shift in the distribution of tropospheric O3 in the Northern Hemisphere due to the 

3 industrialization of societies at low-latitudes (Berntsen et al., 2000, 047916; Lamarque et al., 2005, 

4 630287). By constraining a model with historical (1950s-2000) observations, Shindell et al. (2002, 

5 080130) calculated a large increase of 8.2 DU in tropospheric O3 over polluted continental regions 

6 since 1950. Their result appears consistent with the large change in tropospheric O3 since 

7 preindustrial times implied by the observations from the late 1800s (Marenco et al., 1994, 047733; 

8 Pavelin et al., 1999, 087296). 

10.2.3.2.	 The Effe c t o f Su rfa c e Albe do on Ozone Forcing 

9 The Earth's surface albedo plays a role in O3 forcing. Through most of the troposphere, 

10 absorption of incoming shortwave solar radiation by O3 is small relative to its absorption of outgoing 

11 longwave terrestrial radiation. However, over surfaces characterized by high albedo (e.g., over snow, 

12 ice, or desert sand), incoming radiation is more likely to be reflected than over darker surfaces, and 

13 the probability that O3 will absorb shortwave solar energy is therefore larger. In other words, energy 

14 that would otherwise return to space may instead be deposited in the atmosphere. Several studies 

15 have shown that transport of O3 to the Arctic from mid-latitudes leads to radiative forcing estimates 

16 greater than 1.0 W/m2 in the region, especially in summer (Liao et al., 2004, 057414; Mickley et al., 

17 1999, 047918; Shindell et al., 2006, 631181). Because the Arctic is especially sensitive to radiative 

18 forcing through the ice-albedo feedback, the large contribution in the shortwave to the total radiative 

19 forcing in the region may be important. 

10.2.3.3.	 The Effec t of Vertical Dis tribution on Ozone Forcing 

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

21 increases in surface temperature (Lacis et al., 1990, 037834; Wang et al., 1980, 674821). This is a 

22 result of the colder temperature of the tropopause relative to the rest of the troposphere and 

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

24 power of its temperature1, the colder the added O3 is relative to the earth’s surface, the weaker the 

25 radiation emitted and the greater the “trapping” of longwave radiation in the troposphere. 

10.2.3.4.	 Fe e dba c k Fa c tors tha t Alte r the Climate Res pons e to Cha nges in Ozone

Forcing 

26 Estimates of radiative forcing provide a first-order assessment of the effect of tropospheric O3 

27 on climate. In the real atmosphere, climate feedbacks and transport of heat alter the sensitivity of 

28 Earth’s surface temperature to addition of tropospheric O3. Assessment of the full climate response to 

29 increases in tropospheric O3 requires use of a climate model to simulate these interactions. 

1 As described by the Stefan-Boltzmann law, an ideal blackbody--which the atmosphere approximates--absorbs at all wavelengths and reradiates proportional to the fourth power of its temperature. 

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

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

preindustrial times occurring over the northern mid-latitudes (more on this in Section 10.2.5). If 

climate feedbacks are particularly powerful, they may obscure or even erase the correlation between 

regional radiative forcing and climate response (Boer and Yu, 2003, 630271; Harvey, 2004, 190598). 

For example, several model studies have reported that the horizontal pattern of surface temperature 

response from 2000-2100 trends in short-lived species (including O3) closely matches the pattern 

from the trends in the long-lived greenhouse gases over the same time period (Levy H et al., 2008, 

631174; Shindell et al., 2007, 521350; Shindell et al., 2008, 190393). This correspondence occurs 

even though the patterns of radiative forcing for the short-lived and long-lived species differ 

significantly. In a separate paper, Shindell (2007, 521350) found that Arctic temperatures are 

especially sensitive to the mid-latitude radiative forcing from tropospheric O3. 

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

some consistency with the O3 forcing. For example, Mickley et al. (2004, 057416) examined the 

change in O3 since preindustrial times and found greater warming in the Northern Hemisphere than 

in the Southern Hemisphere (+0.4oC versus +0.2oC), as well as higher surface temperatures 

downwind of Europe and Asia and over the North American interior in summer. For an array of 

short-lived species including O3, Shindell and Faluvegi (2009, 631180) found that radiative forcing 

applied over northern mid-latitudes yield more localized responses due to local cloud, water vapor, 

and albedo feedbacks than radiative forcing applied over the tropics. 

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

distribution of tropospheric O3. The previous section (Section 10.2.3.3) described the greater impact 

of O3 added to the upper troposphere (near the tropopause) on radiative forcing, relative to additions 

in the mid- to lower troposphere. However, warming induced by increased O3 in the upper 

troposphere could stabilize the atmosphere to some extent, limiting the transport of heat to the 

Earth’s surface and mitigating the impact of the added O3 on surface temperature (Christiansen, 

1999, 047920; Joshi et al., 2003, 193752). Hansen et al. (1997, 043104) determined that allowing 

cloud feedbacks in a climate model meant that O3 enhancements in the mid-troposphere had the 

greatest effect on surface temperature. 

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

relative to another. For example, Mickley et al. (2004, 057416) found a greater temperature response 

to CO2 forcing than to an O3 forcing of similar global mean magnitude, due in part to the relatively 

weak ice-albedo feedback for O3. Since CO2 absorbs in the same bands as water vapor, CO2 forcing 

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

radiative forcing amplifies the ice-albedo feedback in the case of CO2, and the greater midtroposphere radiative forcing allows for greater surface temperature response, relative to that for O3. 

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10.2.3.5. Indire c t Effe c ts of Tro pos phe ric Ozone on the Carbon Cycle 

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

2 damaging plant life in ways discussed in Chapter 9, increases in tropospheric O3 may depress the 

3 land-carbon sink of CO2, leading to accumulation of CO2 in the atmosphere and ultimately warming 

4 of the Earth's surface. Sitch et al. (2007, 093294) calculated that this indirect warming effect of O3 

on climate has about the same magnitude as the O3 direct effect. Their results suggest a doubled 

6 sensitivity of surface temperatures to O3 forcing, compared to current model estimates. 

10.2.4.	 Co mp e tin g Effe cts o f Ozone Precurs ors on Climate 

7 Changes in O3 precursors affect not just O3 concentrations, but also other species that have 

8 importance to the radiative balance of the earth's climate system. For example, an increase in CO or 

9 VOCs would lead to a decrease in hydroxyl (OH) concentrations. Since OH is a major sink of the 

greenhouse gas CH4, a decline in OH would lengthen the CH4 lifetime, enhance the CH4 burden, and 

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

12 OH in certain locations, shortening the CH4 lifetime and leading to surface cooling (Fuglestvedt et 

13 al., 1999, 047431). 

14 Analyzing the net radiative forcing per unit emission for a suite of O3 precursors, Shindell and 

Faluvegi (2009, 631180) calculated positive (+0.25 W/m2) radiative forcing from the increase in 

16 anthropogenic emissions of CO and VOCs since preindustrial times, as well as for CH4 (+1 W/m2). 

17 In contrast, they found negative (-0.29 W/m2) radiative forcing from anthropogenic emissions of 

18 NOX due mainly to the link between NOX and CH4. Other studies have found a near cancellation of 

19 the positive O3 forcing and the negative CH4 forcing that arise from an incremental change in 

anthropogenic NOX emissions (Fiore et al., 2002, 051221; Fuglestvedt et al., 1999, 047431; Naik et 

21 al., 2005, 193194). In addition, Wild et al. (2001, 193196) found that an increase in surface NO 

22 emissions would lead to net cooling, while an increase in aircraft NO emissions would lead to net 

23 warming. 

24 These results point out the need for careful assessment of net radiative forcing involving 

multiple pollutants in developing climate change policy (Unger et al., 2008, 631186). Naik et al. 

26 (2005, 193194) has calculated that a carefully combined reduction of CO, VOCs, and NOX emissions 

27 could lead to net cooling. In addition, several studies point to CH4 as an attractive target for 

28 emissions control since CH4 is itself an important precursor of O3 (Fiore et al., 2002, 051221; West 

29 et al., 2007, 622733). Fiore et al. (2002, 051221) found that reducing anthropogenic CH4 emissions 

by 50% would lead to a global radiative cooling of -0.37 W/m2, mostly from CH4. 

10.2.5.	 Calculating Ra diative Fo rcin g a nd Clima te Re s po n s e to Pa s t Tren d s

in Tro po s ph e ric Ozone 

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

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

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

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

change in tropospheric O3 since 1750 (Forster et al., 2007, 092936), ranking it third in importance 

among greenhouse gases after CO2 and CH4. The O3 forcing could, in fact, be as large as 0.7 W/m2

, 

if reconstructions of preindustrial and mid-20th century O3 based on the measurement record are 

valid (Mickley et al., 2001, 080134; Shindell and Faluvegi, 2002, 080130). In any event, Unger et al. 

(2010, 387104) showed that present-day O3 forcing can be attributed to emissions from across many 

economic sectors, including on-road vehicles, household biofuel, power generation, and biomass 

burning. As much as one-third of the radiative forcing from the 1890 to 1990 change in tropospheric 

O3 could be due to increased biomass burning (Ito et al., 2007, 608803). 

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

data. Using data from TOMS, Worden et al. (2008, 631188) estimated a reduction in clear-sky 

outgoing longwave radiation of 0.48 W/m2 by O3 in the upper troposphere over oceans in 2006. This 

radiative forcing includes contributions from both anthropogenic and natural O3. Assuming that the 

burden of O3 has roughly doubled since preindustrial times (Gauss et al., 2006, 630275), the total O3 

forcing estimated with TOMS is consistent with that obtained from models estimating just the 

anthropogenic contribution. 

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

complexity of feedbacks, as mentioned in Sections 10.2.2.3 and 10.2.3.4. In their model study, 

Mickley et al. (2004, 057416) reported a global mean increase of 0.28oC since preindustrial times, 

with values as large as 0.8oC in continental interiors. For the time period since 1870, Hansen et al. 

(2005, 190596) estimated a much smaller increase in global mean surface temperature (0.11oC), but 

they implemented 1880s anthropogenic emissions in their base simulation and also took into account 

trends in both stratospheric and tropospheric O3; the modeled decline of lower stratospheric O3, 

especially over polar regions, cooled surface temperatures in this study, counteracting the warming 

effect of increasing tropospheric O3. 

Figure 10-3 shows the Hansen et al. (2005, 190596) results as reported in Shindell et al. (2006, 

631181). In that figure, summertime O3 has the largest radiative impact over the continental interiors 

of the Northern Hemisphere. In winter, the impact of tropospheric O3 is greatest over the snow and 

ice regions of the Arctic, where the probability of O3 absorption of shortwave radiation is high 

(Section 10.2.3.2). Shindell et al. (2006, 631181) estimated that the change in tropospheric O3 over 

the 20th century could have contributed about 0.3oC to annual mean Arctic warming and as much as 

0.4-0.5oC during winter and spring. Over eastern China, Chang et al. (2009, 630273) calculated a 

surface temperature increase of 0.4oC to the 1970-2000 change in tropospheric O3. It is not clear, 

however, to what degree regional changes in O3 burden influenced this response, as opposed to more 

global changes. 

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Source: Used with permission from American Geophysical Union, Shindell et al. (2006, 631181) 

Figure 10-3. Ensemble average 1900-2000 surface temperature trends (oC per century) in

response to tropospheric ozone changes and 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 area-weighted global averages in the same units as the 

plots. 

10.2.6.	 Calculating the Radiative Forcing and Climate Res pons e to Future 

Tre nd s in Trop o s p h eric 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 will also 

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

4 atmosphere. This section describes the following issues: (1) projected trends in the anthropogenic 

5 emissions of O3 precursors; (2) the effects of these emissions on the tropospheric O3 burden; (3) the 

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1 effects of changing climate on tropospheric O3; and (4) radiative forcing and climate response to 21st 

2 century trends in tropospheric O3. 

10.2.6.1. Emis s io ns o f Anthro po ge nic Ozone Precursors a cross the 21st Century 

3 The IPCC SRES effort devised scenarios for short-lived O3 precursors as well as the well4 mixed greenhouse gases including NOX, CO, and VOCs (IPCC, 2000, 080704). Using the IMAGE 

socioeconomic model, Streets et al. (2004, 190423) provided speciation for NOX and VOCs and 

6 allocated the trends in emissions over 17 regions and 8 economic sectors for the 2000-2050 time 

7 period. The worst-case IPCC scenario, A2, features continued dependence on fossil fuels, rapid 

8 population growth, and little sharing of technology between developed and developing nations. By 

9 2100 in this scenario, global NOX emissions increase by a factor of 3.4 and CO emissions and CH4 

by ~2.7, relative to 2000 (IPCC, 2000, 080704). Most of these increases in emissions occur over 

11 developing countries. For example over Asia, NOX emissions in the A2 scenario increase by more 

12 than a factor of 4 by 2100. The more moderate A1B scenario has global NOX and CO emissions 

13 increasing by 25% and 90%, respectively by 2100, but global CH4 emissions decreasing by 10%. In 

14 the B1 scenario, with its emphasis on clean and efficient technologies, global emissions of NOX, CO, 

and CH4 all decrease by 2100 relative to the present day (-40%, -60%, and -30%, respectively). 

16 Other emissions scenarios have been recently developed to describe trends in the short-term 

17 (up to 2030). The Current Legislation (CLE) scenario provides trends consistent with existing air 

18 quality regulations; the Maximum Feasible Reduction (MFR) scenario seeks to reduce emissions of 

19 O3 precursors to the maximum extent possible. Emission source changes relative to the present day 

for CLE, MFR, and A2 are given in Stevenson et al. (2006, 089222). 

21 For the Fifth Assessment Report (IPCC AR5), a new set of scenarios has been developed: the 

22 Representative Concentration Pathways (RCPs) (Moss et al., 2010, 664501). The RCPs will explore 

23 for the first time approaches to climate change mitigation. The scenarios are designed to achieve 

24 radiative forcing targets of 2.6, 4.5, 6.0 and 8.5 W/m2 by 2100, and have been designated RCP 2.6, 

RCP 4.5, RCP 6.0, and RCP 8.5 (RCP 2.6 is also known as RCP3-PD.) In all scenarios, global 

26 anthropogenic NOX emissions decline 30-50% during the 21st century, though RCP 8.5 shows a peak 

27 during the 2020s at a value ~15% greater than that of 2000. Global anthropogenic VOC and CO 

28 emissions are relatively flat during the 2000-2050 time range, and then decline by 30-50% by the 

29 end of the century. For CH4, global mean emission trends for the four RCP scenarios differ 

significantly across the 21st century, with RCP 8.5 showing a tripling of emissions by 2100, and 

31 RCP 2.6 showing the emissions cut by half in this time range. All these global trends, however, 

32 contain some regional variation. For example, Asian emissions of both NOX and VOCs show 

33 significant increases in the near term (2030s to 2050s). Plots of the RCP trends can be found at 

34 http://iiasa.ac.at/web-apps/tnt/RcpDb/dsd?Action=htmlpage&page=about (RCP, 2009, 677552). 

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10.2.6.2. Impa c t o f 21st Cen tury Tren ds in Emis s io ns o n Trop os p heric Ozone 

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

2 emissions of its precursors. As shown in Table 10-1, a recent multi-model study found increases in 

3 the tropospheric O3 burden of 15% and 6% for the IPCC A2 and CLE scenarios respectively for the 

4 2000-2030 time period, and a decrease for the MFR scenario of 5% (Stevenson et al., 2006, 089222). 

5 These results indicate that the growth in tropospheric O3 between 2000 and 2030 could be reduced or 

6 even reversed, depending on emission controls. For the relatively moderate A1B emissions scenario 

7 over the 2000-2050 time period, Wu et al. (2008, 190039) calculated a change in O3 burden of about 

8 20%. Looking further into the 21st century, Gauss et al. (2003, 094204) reported O3 burden changes 

9 of 30-60% in response to application of the A2p anthropogenic emissions over the 2000-2100 time 

10 period (the A2p scenario was a preliminary version of the A2 scenario). Using the A2 scenario for 

11 the same 100 year time period, Pyle et al. (2007, 630284) projected a 50% increase in the O3 burden, 

12 consistent with Gauss et al. (2003, 094204) and with Liao et al. (2006, 664500) who calculated an O3 

13 change of 60% for the same conditions. Given the large (+40 ppb) monthly mean increases in 

14 surface O3 that the A2 or A2p scenarios would yield over Asia and elsewhere by the end of the 21st 

15 century (Prather et al., 2003, 047879), these 100-yr projections of the O3 burden would lead to 

16 extremely unhealthy air quality. 

17 As noted above, the RCP scenarios of AR5 show long-term declines in the global mean 

18 emissions of O3 precursors, with some regional increases in the near-term. As of this writing, no 

19 model study has reported the response of the tropospheric O3 burden to any of the IPCC AR5 

20 scenarios. 

Table 10-1. 2000-2030 changes in anthropogenic emissions, and CH4 and tropospheric ozone burdens,

and the associated tropospheric ozone forcing for three scenarios; values are ensemble 

means 

Current Legislation Maximum Feasible Reduction Scenario IPCC A2 (CLE) (MFR) 

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

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

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

Percent change in tropospheric O3 burden +15% +6% -5% 

Radiative forcing due to O3 changea

(W/m2) 0.3 0.18 -0.05 

a

Includes radiative forcing due to corresponding CH4 change. 

Source: Adapted from Stevenson et al. (2006, 089222). 

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10.2.6.3.	 Impa c t o f 21st Cen tury Clima te o n Trop os p he ric Ozone 

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

2 tropospheric O3 can be traced to changing emissions. Model studies show that climate change so far 

3 has likely had little impact on the tropospheric O3 (e.g., Grenfell et al., 2001, 664496). In the future, 

4 however, climate change is expected to bring large changes in a suite of variables that could affect 

O3 production, loss, and transport. For example, increased water vapor in a warming atmosphere is 

6 expected to enhance OH concentrations, which in remote, NOX-poor regions will accelerate O3 loss 

7 rates (Johnson et al., 1999, 052390). 

8 In the 2050s A1B climate, Wu et al. (2008, 629684) calculated a 5 ppb decrease in surface O3 

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

biogenic precursor of O3. Model studies have calculated 21st-century increases in isoprene emissions 

11 ranging from 25-50%, depending on climate scenario and time horizon (Wu et al., 2008, 190039, and 

12 references therein). These studies however did not take into account the effects of changing climate 

13 and CO2 burden on vegetation extent, which could have large consequences for biogenic emissions 

14 (Heald et al., 2008, 191617; Sanderson et al., 2003, 630286). In any event, enhanced isoprene 

emissions will increase O3 concentrations in VOC-limited regions, but decrease O3 in NOX-limited 

16 regions (Pyle et al., 2007, 630284; Sanderson et al., 2003, 630286; Wu et al., 2008, 190039). 

17 Convection frequencies and lightning flash rates will also likely change in a changing climate, with 

18 consequences for lightning NOX emissions and O3 concentrations in the upper troposphere (Price and 

19 Rind, 1994, 630283; Sinha and Toumi, 1997, 047932). While Wu et al. (2008, 190039) calculated an 

increase in lightning NOX by 2050 due to enhanced deep convection, Jacobson and Streets (2009, 

21 630281) projected a decrease in lightning NOX due to a declining cloud ice in their future 

22 atmosphere. Finally, changes in transport processes will almost certainly accompany global climate 

23 change. For the 2050 A1B climate, Wu et al. (2008, 629684) showed that flattening of the meridional 

24 temperature gradient in a warming world would lead to slower intercontinental transport of 

tropospheric O3. For the A2 climate in 2100, Zeng and Pyle (2003, 047492) projected an 80% 

26 increase in the flux of stratospheric O3 into the troposphere, relative to the present-day. 

27 Taken together, these climate-driven processes could have significant effects on the burden 

28 and distribution of tropospheric O3. As shown in Wu et al. (2008, 629684), model projections of the 

29 change in O3 burden due solely to future climate change range from -12% to +3%, depending on the 

model, scenario, and time horizon. 

10.2.6.4.	 Ra dia tive Forc in g a n d Clima te Re s p ons e fro m 21st Ce ntury Tre nds in 

Tro pos pheric Ozone 

31 In the near term (2000-2030), Stevenson et al. (2006, 089222) estimated an O3 forcing of near 

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

33 (2008, 613861), following the moderate A1B scenario, calculated a radiative forcing of 0.12 W/m2 

34 from the 2000-2030 change in tropospheric O3, about the same as that derived by Stevenson et al. 

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1 (2006, 089222) for the CLE scenario. Over the longer term (2000 to 2100) for the A1B scenario, 

2 Gauss et al. (2003, 094204) reported large positive radiative forcing (0.40 to 0.78 W/m2) due to the 

3 change in tropospheric O3, as shown in Figure 10-4. Normalized radiative forcing for these model 

4 calculations fell within a relatively narrow range, 0.032 to 0.040 W/m2/DU, indicating that the 

5 largest uncertainty lies in the model-calculated changes in O3 burden. Applying the A2 scenario, 

6 Chen et al. (2007, 630274) estimated a global mean radiative forcing of 0.65 W/m2 from 

7 tropospheric O3 by 2100, consistent with the Gauss et al. (2003, 094204) results. These studies took 

8 into account only the impact of changing emissions on tropospheric O3. In their calculations of the 

9 2000-2100 radiative forcing from O3 in the A2 scenario, Liao et al. (2006, 664500) found that 

10 inclusion of climate effects on tropospheric O3 reduced their radiative forcing estimate by 20%. 

Source: Used with permission from American Geophysical Union, Gauss et al. (2003, 094204) 

Figure 10-4. Global mean radiative forcing estimates calculated by a set of models for the 2000­ 2100 change in tropospheric ozone. ﷒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. 

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

12 future atmosphere to a suite of short-lived species (e.g., Levy H et al., 2008, 631174; Shindell et al., 

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1 2007, 521350; Shindell et al., 2008, 190393). Few studies, however, have calculated the climate 

2 response to changes in tropospheric O3 alone in the future atmosphere. For the A2 atmosphere, Chen 

3 et al. (2007, 630274) estimated a global mean surface temperature increase of +0.34oC by 2100 in 

4 response to the change in O3. The largest temperature increases in this study, as much as 5oC, 

occurred over the populous regions of Asia and the Middle East and downwind of biomass burning 

6 regions in South Africa and South America. 

10.2.7. Su mma ry of the Effe cts o f Tro p os p he ric Ozone on Climate 

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

8 developed world has successfully reduced emissions of O3 precursors in recent decades, many 

9 developing countries have experienced large increases in precursor emissions and these trends are 

expected to continue, at least in the near term. Projections of radiative forcing due to changing O3 

11 over the 21st century show wide variation, due in large part to the uncertainty of future emissions of 

12 source gases. In the near-term (2000-2030), projections of O3 radiative forcing range from near zero 

13 to +0.3 W/m2, depending on the emissions scenario (Stevenson et al., 2006, 089222). Reduction of 

14 tropospheric O3 concentrations could therefore provide an important means to slow climate change 

in addition to the added benefit improving surface air quality. 

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

17 have competing effects on the greenhouse gas CH4, complicating emissions reduction strategies. A 

18 decrease in CO or VOC emissions would enhance OH concentrations, shortening the lifetime of 

19 CH4, while a decrease in NOX emissions could depress OH concentrations in certain regions and 

lengthen the CH4 lifetime. Recent research, however, has shown that a carefully combined reduction 

21 of CO, VOCs, and NOX emissions could lead to net cooling (Naik et al., 2005, 193194). In addition, 

22 abatement of CH4 emissions would provide a straightforward means to address climate change since 

23 CH4 is itself an important precursor of background O3 (Fiore et al., 2002, 051221; West et al., 2006, 

24 196558; West et al., 2007, 622733). 

Important uncertainties remain regarding the impact of O3 on future climate change. To 

26 address these uncertainties, further research is needed to: (1) enhance our knowledge of the natural 

27 atmosphere; (2) interpret observed trends of O3 in the free troposphere and remote regions; (3) 

28 understand the relationship between regional O3 forcing and regional climate change; and (4) 

29 determine the optimal mix of emissions reductions that would act to limit future climate change. 

10.3. UV-B Re la te d Effe c ts a n d Tro p os p h e ric Ozone 

10.3.1. Background 

Ultraviolet (UV) radiation emitted from the Sun contains sufficient energy when it reaches the 

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

32 living organisms and materials. Atmospheric O3 plays a crucial role in reducing exposure to solar 

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1 UV radiation at the Earth's surface. Stratospheric O3 is responsible for the majority of this shielding 

2 effect, as approximately 90% of total atmospheric O3 is located there over mid-latitudes (Crist et al., 

3 1994, 668881; Kar et al., 2010, 670423). Investigation of the supplemental shielding of UV radiation 

4 provided by tropospheric O3 is important for quantifying UV exposure and the incidence of related 

5 human health effects, ecosystem effects, and materials damage. The role of tropospheric O3 in 

6 shielding of UV radiation is discussed in this section. 

10.3.2. Ph ys ic s o f UV Radia tio n a nd Flu x 

7 Solar UV radiation is subdivided into classes based on wavelength: UV-A refers to wavelengths from 

8 400-315 nm; UV-B from 315-280 nm; and UV-C from 280-100 nm. Since the energy possessed by a 

9 photon is inversely proportional to its wavelength, UV-A radiation is the least energetic and UV-C is 

10 the most energetic, with UV-B falling in-between. The wavelength determines how the photons 

11 interact with the complex mixture of gases, clouds and particles present in the atmosphere (see 

12 Figure 10-5). UV-A radiation can be scattered but is not absorbed to any meaningful degree by 

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

14 atmosphere. UV-C is almost entirely blocked by the Earth’s upper atmosphere, where it participates 

15 in photoionization and photodissociation processes. Since UV-A is less energetic and does not 

16 interact with O3 and UV-C is almost entirely blocked by stratospheric O3, UV-B is the most 

17 important band of UV radiation to consider in relation to tropospheric O3 shielding. 

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Figure 10-5. Diagram of the factors that determine human exposure to ultraviolet radiation. 

Solar flux has a temporal dependence, while radiative scattering and absorption have strong 

wavelength, path length, and/or particle concentration dependencies. These combine to create 

nonlinear effects on UV flux at the Earth’s surface. Thus, careful quantification of atmospheric 

absorbers and scatterers, along with a well-resolved description of the physics of these interactions, 

is necessary for predicting the impact of ground-level O3 on UV flux. Chapter 10 of the 2006 O3 

AQCD (U.S. EPA, 2006, 088089) describes in detail several key factors that influence the 

spatiotemporal distribution of ground-level UV radiation flux, including: (1) long-term solar activity 

including sunspot cycle; (2) solar rotation; (3) the position of the Earth in its orbit around the sun; (4) 

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

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

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

medium, particle size (for aerosols and clouds), and the altitude at which these processes are 

occurring. These properties are sensitive to meteorology, which introduces additional elements of 

temporal dependency in ground-level UV radiation flux. As seen in data collected by the Global 

Ozone Monitoring by Occultation of Stars (GOMOS) instrument onboard the European Space 

Agency’s ENVISAT satellite (Figure 10-6), atmospheric O3 density undergoes wide natural variation 

on relatively short timescales, particularly at mid-latitudes (Kyrola et al., 2010, 667819). 

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Source: Used with permission from Copernicus Publications, Kyrola et al., (2010, 667819). 

Figure 10-6. Monthly stratospheric ozone number density (scaled by 1 × 1012 molecules/cm3), in 

3 latitude belts as a function of time (August 2002 - December 2008) and altitude 

(15-40 km) from the Global Ozone Monitoring by Occultation of Stars (GOMOS)

instrument onboard the European Space Agency’s ENVISAT satellite. 

Latitude belts: 30oN-50oN (top), 10oS-10oN (middle), 30oS-50oS (bottom). White space in the

panels means that there are not enough data available. 

1 The lower atmospheric pressure in the stratosphere means fewer gas molecules are present that 

2 can absorb or scatter radiation. Stratospheric clouds and aerosols are also thinner and more dispersed 

3 than those in the troposphere. In the language of the radiative transfer literature, these conditions 

4 make the stratosphere a “single scattering” regime for UV radiation. The troposphere, due to its high 

5 gas and particle concentrations is referred to as a “multiple scattering” regime. In practical terms, 

6 UV radiation traverses the stratosphere with a substantially lower probability of encountering a gas 

7 molecule, cloud, or aerosol particle than it would in the troposphere. The multiple scattering of UV 

8 radiation in the troposphere accounts for the “disproportionate” role that tropospheric O3 is said to 

9 play in absorbing UV radiation versus stratospheric O3 on a molecule per molecule basis (Balis et al., 

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1 2002, 055023; Bruhl and Crutzen, 1989, 012518; Crist et al., 1994, 668881; Zerefos et al., 2002, 

2 055169). 

3 Latitude and altitude are primary variables in defining UV-B flux at the Earth’s surface, 

4 immediately followed in importance by clouds, surface albedo, PM concentration and composition, 

and then by gas phase pollution. Of all these variables, only latitude and altitude can be defined with 

6 small uncertainty in any effort to develop a UV climatology for use in a public health benefits 

7 analysis relevant to the areas not presently attaining the NAAQS for O3. Cloud cover, and its effect 

8 on surface UV flux, continues to be extremely difficult to define and predict. Particulate matter and 

9 gas-phase tropospheric pollutants are subject to similarly high degrees of uncertainty in predicting 

their relative concentration distributions, but recent advancements have been made (e.g., Bais et al., 

11 2005, 669135; Bergstrom et al., 2004, 669158; Goering et al., 2005, 669164). Land cover and, 

12 consequently, surface albedo is highly variable at the geographic scales relevant to NAAQS 

13 attainment. 

14 The 2006 WMO assessment (WMO, 2006, 669178) reported that global average total column 

O3 had declined by 3.5% from pre-1980 concentrations due to the presence of anthropogenic O3­

16 depleting substances in the atmosphere. In the period 2002-2005, no additional declines were found 

17 in the global average due to bans on and reduced emissions of O3-depleting substances. The report 

18 found that O3 depletion has a strong latitude and seasonal dependence: total column O3 declined by 

19 ~3% in the Northern Hemisphere, declined by ~6% in the Southern Hemisphere, and remained 

essentially unchanged over the tropics relative to pre-1980 total column O3 abundances. Polar 

21 stratospheric O3 depletion is more complex and exhibits large interannual variations driven by 

22 changes in meteorology. 

10.3.3. Hu ma n Exp o s u re a nd Su s ce p tibility to Ultraviole t Radia tio n 

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

24 Chapter 10 of the 2006 O3 AQCD (U.S. EPA, 2006, 088089) and are summarized here. These factors 

included outdoor activity, occupation, age, gender, geography, and protective behavior. Outdoor 

26 activity and occupation both influenced the amount of time people spend outdoors during daylight 

27 hours, the predominant factor for exposure to solar UV radiation. Participation in outdoor sports 

28 (e.g., basketball, soccer, golf, swimming, cycling) significantly increased UV radiation exposure 

29 (Moehrle, 2001, 057502; Moehrle et al., 2000, 057503; Thieden et al., 2004, 057558; Thieden et al., 

2004, 057557). Occupations that substantially increased exposure to UV radiation included farming 

31 (Airey et al., 1997, 057458; Schenker et al., 2002, 057522), fishing (Rosenthal et al., 1988, 057513), 

32 landscaping (Rosenthal et al., 1988, 057513), construction (Gies and Wright, 2003, 057477), 

33 physical education (Vishvakarman et al., 2001, 057614), mail delivery (Vishvakarman et al., 2001, 

34 057614), and various other occupations that require workers to spend the majority of their day 

outdoors during peak UV radiation hours. 

36 Age and gender were found to be factors that influence human exposure to UV radiation, 

37 particularly by influencing other factors of exposure such as outdoor activity and risk behavior. 

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1 Studies indicated that females generally spent less time outdoors and, consequently, had lower UV 

2 radiation exposure compared to males (Gies et al., 1998, 057478; Godar et al., 2001, 057165; 

3 Shoveller et al., 1998, 057524). The lowest exposure to UV radiation among Americans in the Godar 

4 et al. (2001, 057165) study was received in females during their child raising years (age 

5 22-40 years); the highest exposure was observed in males aged 41-59 years. A similar Canadian 

6 survey found that younger adult males had the greatest exposures to UV radiation (Shoveller et al., 

7 1998, 057524). 

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

9 radiation. In the U.S. study by Godar et al. (2001, 057165), northerners and southerners were found 

10 to spend an equal amount of time outdoors; however, the higher solar flux at lower latitudes 

11 significantly increased the annual UV radiation dose for southerners. The annual UV radiation doses 

12 in southerners were 25 and 40% higher in females and males, respectively, compared to northerners. 

13 Other studies also have shown that altitude and latitude influence personal exposure to UV radiation 

14 (e.g., Kimlin et al., 1998, 057491; Rigel et al., 1999, 057511). 

15 Protective behaviors such as using sunscreen (e.g., Nole and Johnson, 2004, 057505), wearing 

16 protective clothing (e.g., Rosenthal et al., 1988, 057513), and spending time in shaded areas (e.g., 

17 Moise et al., 1999, 057504) were shown to reduce exposure to UV radiation. In one study, the use of 

18 sunscreen was associated with extended intentional UV radiation exposure (Autier et al., 1999, 

19 057459); however, a follow-up study indicated that sunscreen use increased duration of exposures to 

20 doses of UV radiation that were below the threshold level for erythema (Autier et al., 2000, 057069). 

21 Given these and other factors that potentially influence UV radiation exposure, the 2006 O3 

22 AQCD (U.S. EPA, 2006, 088089) listed the following subpopulations potentially at risk for higher 

23 exposures to UV radiation: 

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

25  Individuals who participate in outdoor sports and activities; 

26  Individuals who work outdoors with inadequate shade (e.g., farmers, construction 

27 workers, etc.); and 

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

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

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

31 (Sliney and Wengraitis, 2006, 651896). 

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10.3.4. Hu ma n Health Effe cts d u e to UV-B Radiation 

Chapter 10 of the 2006 O3 AQCD (U.S. EPA, 2006, 088089) covered in detail the human 

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

cancer, ocular damage, and immune system suppression. These adverse effects, along with protective 

effects of UV radiation through increased production of vitamin D are summarized in this section. 

For additional details, the reader is referred to Chapter 10 of the 2006 O3 AQCD (U.S. EPA, 2006, 

088089) and references therein. 

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

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

(Matsumura and Ananthaswamy, 2004, 057184). Many studies discussed in the 2006 O3 AQCD 

(U.S. EPA, 2006, 088089) found skin type to be a significant risk factor for erythema. Additional 

risk factors include atopic dermatitis (ten Berge et al., 2009, 651897). 

Skin cancer is another prevalent health effect associated with UV radiation. Exposure to UV 

radiation is considered to be a major risk factor for all forms of skin cancer (Diepgen and Mahler, 

2002, 093593; Gloster and Brodland, 1996, 057479). Ultraviolet radiation is especially effective in 

inducing genetic mutations and acts as both a tumor initiator and promoter. Keratinocytes have 

evolved DNA repair mechanisms to correct the damage induced by UV; however, mutations can 

occur, leading to skin cancers that are appearing with increasing frequency (Hildesheim and Fornace, 

2004, 057168). The relationship between skin cancer and chronic exposure to UV radiation is further 

explored in Chapter 10 of the 2006 O3 AQCD (U.S. EPA, 2006, 088089). 

Ocular damage from UV radiation exposure includes effects on the cornea, lens, iris, and 

associated epithelial and conjunctival tissues. The region of the eye effected by exposure to UV 

radiation depends on the wavelength of the incident UV radiation. Depending on wavelength, 

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

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

Experimental studies have suggested that exposure to UV radiation may suppress local and 

systemic immune responses to a variety of antigens (Clydesdale et al., 2001, 057105; Garssen and 

Van Loveren, 2001, 057161; Selgrade et al., 1997, 036165). In rodent models, these effects have 

been shown to worsen the course and outcome of some infectious diseases and cancers (Granstein 

and Matsui, 2004, 057747; Norval et al., 1999, 036089). Results from human clinical studies suggest 

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

induction (Caforio et al., 2000, 080058; Lindelof et al., 2000, 080084; Ullrich, 2005, 095635). There 

is also evidence that UV radiation has indirect involvement in viral oncogenesis through the human 

papillomavirus (Pfister, 2003, 057515), dermatomyositis (Okada et al., 2003, 057769), human 

immunodeficiency virus (Breuer-McHam et al., 2001, 057736) and other forms of 

immunosuppression (Selgrade M-JK; Smith et al., 2001, 057454). 

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

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

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1 (Holick, 2004, 057691). Vitamin D deficiency can cause metabolic bone disease among children and 

2 adults, and also may increase the risk of many common chronic diseases, including type I diabetes 

3 mellitus and rheumatoid arthritis (Holick, 2004, 057691). Substantial in vitro and toxicological 

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

forms of cancer (Freedman et al., 2002, 035530; Garland et al., 1990, 025242; Giovannucci, 2005, 

6 074099; Gorham et al., 1990, 000682; Grant, 2002, 034981; Grant, 2002, 025244; Grant and 

7 Garland, 2004, 075093; Hanchette and Schwartz, 1992, 025257; Hughes et al., 2004, 074101; John 

8 et al., 1999, 057489; John et al., 2005, 670422; Lefkowitz and Garland, 1994, 025253; Smedby et 

9 al., 2005, 669175; Studzinski and Moore, 1995, 057554). In some studies, UV-B related production 

of vitamin D had potential beneficial immunomodulatory effects on multiple sclerosis, insulin11 dependent diabetes mellitus, and rheumatoid arthritis (Ponsonby et al., (2002, 080061); Cantorna, 

12 (2000, 080060)). More details on UV-B protective studies are provided in Chapter 10 of the 2006 O3 

13 AQCD (U.S. EPA, 2006, 088089). 

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

on Non-Ionizing Radiation Protection (ICNIRP) agreed that some low-level exposure to UV 

16 radiation has health benefits (ICNIRP, 2004, 057187). However, the adverse health effects of higher 

17 UV exposures necessitated the development of exposure limits for UV radiation. The ICNIRP 

18 recognized the challenge in establishing exposure limits that would achieve a realistic balance 

19 between beneficial and adverse health effects. As concluded by ICNIRP (2004, 057187), "[t]he 

present understanding of injury mechanisms and long-term effects of exposure to [UV radiation] is 

21 incomplete, and awaits further research". 

10.3.5. Ec o s ys te m a nd Ma te rials Da ma g e Effe cts Due to UV-B Radiation 

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

23 Environmental Effects Assessment Panel (UNEP, 2009, 669084) lists many ecosystem and materials 

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

damage effects from UV-B radiation per se is out of the scope of this assessment. However, a brief 

26 summary of some mid-latitude effects is provided in this section to provide context for UV-B related 

27 issues pertaining to tropospheric O3. The reader is referred to the UNEP report (UNEP, 2009, 

28 669084) and references therein for further details. All of these UV-B related ecosystem and materials 

29 effects can also be influenced by climate change through temperature and other meteorological 

alterations, making quantifiable predictions of UV-B effects difficult. 

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

32 and plant cover, changes in biodiversity, susceptibility to infection, and increases in natural UV 

33 protective responses. In general, however, these effects are small for moderate UV-B increases at 

34 mid-latitudes. A field study on wheat in southern Chile found no substantial changes in crop yield 

with moderate increases in UV-B radiation (Calderini et al., 2008, 668893). Similarly, field studies 

36 on silver birch (Betula pendula) in Finland found no significant effects in photosynthetic function 

37 with increases in UV-B radiation (Aphalo et al., 2009, 668923). Subtle, but important, changes in 

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1 habitat and biodiversity have also been linked to increases in UV-B radiation (Mazza et al., 2010, 

2 668911; Obara et al., 2008, 668913; Wahl, 2008, 668921). Some plants have natural coping 

3 mechanisms for dealing with changes in UV-B radiation (Brown and Jenkins, 2008, 668892; Favory 

4 et al., 2009, 668897; Ioki et al., 2008, 668904; Jenkins, 2009, 668905), but these defenses may have 

5 costs in terms of reduced growth (Clarke and Robinson, 2008, 668928; Phoenix et al., 2000, 668915; 

6 Semerdjieva et al., 2003, 668919; Snell et al., 2009, 668920). 

7 Aquatic ecosystem effects from increased UV-B radiation include sensitivity in growth, 

8 immune response, and behavioral patterns of aquatic organisms. One study looking at 

9 coccolithophores, an abundant phytoplankton group, found a 25% reduction in cellular growth with 

10 UV-B exposure (Gao et al., 2009, 668899). Exposure to relevant levels of UV-B radiation has been 

11 shown to modify immune response, blood chemistry, and behavior in certain species of fish (Holtby 

12 and Bothwell, 2008, 668903; Jokinen et al., 2008, 668906; Markkula et al., 2009, 613291). Adverse 

13 effects on growth and development from UV-B radiation have also been observed for amphibians, 

14 sea urchins, mollusks, corals, and zooplankton (Croteau et al., 2008, 668894; Croteau et al., 2008, 

15 603785; Garcia et al., 2009, 668935; Marquis and Miaud, 2008, 668908; Marquis et al., 2008, 

16 668909; Oromi et al., 2008, 668914; Romansic et al., 2009, 515730). 

17 Biogeochemical cycles, particularly the carbon cycle, can also be influenced by increased UV18 B radiation. A study on high latitude wetlands found UV-induced increases in CO2 uptake through 

19 soil respiration (Haapala et al., 2009, 607260) while studies on arid terrestrial ecosystems found 

20 evidence for UV-induced release of CO2 through photodegradation of above-ground plant litter 

21 (Brandt et al., 2009, 668891; Caldwell et al., 2007, 668927; Henry et al., 2008, 668902; Zepp et al., 

22 2007, 668896). Changes in solar UV radiation may also have effects on carbon cycling and CO2 

23 uptake in the oceans (Brewer and Peltzer, 2009, 669197; Fritz et al., 2008, 668898; Hader et al., 

24 2007, 668901; Meador et al., 2009, 668912; Zepp et al., 2008, 668922) as well as release of 

25 dissolved organic matter from sediment and algae (Mayer et al., 2009, 668910; Riggsbee et al., 2008, 

26 668917). Additional studies showing effects on these and additional biogeochemical cycles including 

27 the water cycle and halocarbon cycle can be found in the UNEP report (UNEP, 2009, 669084) and 

28 references therein. 

29 Materials damage from increased UV-B radiation include UV-induced photodegradation of 

30 wood (Kataoka et al., 2007, 670425) and plastics (Pickett et al., 2008, 668916). These studies and 

31 others summarizing photo-resistant coatings and materials designed to reduce photodegradation of 

32 materials are summarized in the UNEP report (UNEP, 2009, 669084) and references therein. 

10.3.6.	 UV-B Rela te d Effec ts As s ociate d with Cha ng e s in Trop o s p he ric

Ozone Concentrations 

33 There are multiple complexities in attempting to quantify the relationship between changes in 

34 tropospheric O3 concentrations and UV radiation exposure, as described above. Furthermore, 

35 quantifying the relationship between UV radiation and health or welfare effects is complicated by the 

36 uncertainties involved in the selection of an action spectrum and appropriate characterization of dose 

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(e.g., peak or cumulative levels of exposure, timing of exposures, etc.) The lack of published studies 

that critically examine these issues together--that is the incremental health or welfare effects 

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

concentrations--reflects the significant challenges in this field. 

As reported in the 2006 O3 AQCD (U.S. EPA, 2006, 088089), one analysis by Lutter and Wolz 

(1997, 082672) attempted to estimate the effects of a nationwide 10 ppb reduction in seasonal 

average tropospheric O3 on the incidence of nonmelanoma and melanoma skin cancers and cataracts 

in humans. Their estimate, however, depended upon several simplifying assumptions, ranging from 

an assumed generalized 10-ppb reduction in O3 column density, national annual average incidence 

rates for the two types of skin cancer, and simple, linear biological amplification factors. 

Specifically, the decrease of 10 ppbv in seasonally averaged O3 concentrations is likely an 

overestimate since it doesn't account for the influence of background O3 coming from the global 

accumulation or generation of regional chemistry (Adamowicz et al., 2004, 670421). Further, the 

methodologies used in this analysis have ignored area-specific factors that are important in 

estimating the extent to which small, variable changes in ground-level O3 mediate long-term 

exposures to UV-B radiation. 

A handful of studies have addressed the relationship between changes in tropospheric pollutant 

concentrations and UV-B radiation exposure, providing some additional insight. A study by Palancar 

and Toselli (2002, 057207) looked at changes in measured UV-B radiation in relation to ground-level 

air pollutants during several air pollution episodes in Cordoba, Argentina. They found that changes 

in aerosol concentrations explained the majority of UV-B radiation fluctuations, and that changes in 

tropospheric O3 and SO2 had little effect. Repapis et al. (1998, 038015) performed a similar study on 

UV-B exposures during high and low air pollution days in Athens, Greece. They found cloud cover 

and aerosols to be the major factors in observed UV-B exposures reductions. Studies by Acosta and 

Evans (2000, 670420) in Mexico City and Koronakis et al. (2002, 129938) in Athens, Greece both 

found significant reductions in surface-level UV exposures during pollution episodes. Both these 

studies include tropospheric O3 as a potential driver for the reductions, but neither study was able to 

quantify the influence of individual atmospheric components involved in the observed attenuation in 

UV-B radiation. 

In the absence of studies specifically addressing UV-B related health effects from a reduction 

in tropospheric O3, inferences were made in the 2006 O3 AQCD (U.S. EPA, 2006, 088089) on the 

basis of studies focused on stratospheric O3 depletion. Studies included in that review examined the 

potential effect of stratospheric O3 depletion on the risk of erythema (Longstreth et al., 1998, 

001200), skin cancer (De Gruijl, 1995, 057471; Longstreth et al., 1995, 055174; Madronich and 

De Gruijl, 1993, 055183; Slaper et al., 1996, 055128; Urbach, 1997, 086255), nonmelanoma skin 

cancer (Longstreth et al., 1995, 055174; Slaper et al., 1996, 055128), and cataracts (Longstreth et al., 

1995, 055174). Note that several of the concerns expressed above in relation to the Lutter and Wolz 

(1997, 082672) analysis are relevant to these analyses as well. Furthermore, these studies have a 

high degree of uncertainty due to inadequate information on the action spectrum and dose-response 

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relationships. As a result, caution is advised when assessing and interpreting the quantitative results 

of health risks due to stratospheric O3 depletion in the context of tropospheric O3 shielding. 

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

ground-level O3 that would result from attainment of the O3 NAAQS have not been directly 

assessed, they would be expected to be small or nonexistent given the above findings and the fact 

that tropospheric O3 makes up only ~10% of the total atmospheric O3 column at mid-latitudes (Kar 

et al., 2010, 670423). Furthermore, O3 present in the planetary boundary layer makes up only ~10% 

of tropospheric O3 (Thompson et al., 2007, 090796) and the NAAQS has only a fractional influence 

on those ground-level O3 concentrations (i.e., it is not the intent of the NAAQS to entirely eliminate 

ground-level O3). The net result is a very small influence on total column O3 through attainment of 

the O3 standard. In addition, the health benefits of UV-B in the production of vitamin D suggests that 

increased risks of human disease due to a slight excess in UV-B radiation exposure may be offset by 

the benefits of enhanced vitamin D production. However, as with other impacts of UV-B on human 

health, this beneficial effect of UV-B has not been studied in sufficient detail to allow for a credible 

health benefits assessment. Hence, the above mentioned health and welfare effects associated with 

UV-B exposures resulting from changes in ground-level O3 concentrations would likely be small or 

nonexistent based on current information. 

More reasonable estimates of the human health impacts of enhanced UV-B penetration 

following reduced ground-level O3 concentrations require both (a) a solid understanding of the 

multiple factors that define the extent of human exposure to UV-B, and (b) well-defined and 

quantifiable links between human disease and UV-B exposure. Within the uncertain context of 

presently available information on UV-B surface fluxes, a risk assessment of UV-B-related health 

effects would need to factor in human habits (e.g., daily activities, recreation, dress, and skin care) in 

order to adequately estimate UV-B exposure levels. Little is known about the impact of variability in 

these human factors on individual exposure to UV radiation. Furthermore, detailed information does 

not exist regarding the relevant type (e.g., peak or cumulative) and time period (e.g., childhood, 

lifetime, or current) of exposure, wavelength dependency of biological responses, and interindividual variability in UV resistance. In conclusion, the effect of changes in surface-level O3 

concentrations on UV-induced health outcomes cannot yet be critically assessed within reasonable 

uncertainty. The reader is referred to the U.S. EPA 2002 Final Response to Court Remand (2003, 

015702) 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. 

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References
 

A list of all references considered for inclusion in this chapter can be found at 

http://hero.epa.gov/index.cfm?action=search.do&submit=Search&group_id=478 

Acosta, L. R.; Evans, W. F. J. (2000). Design of the Mexico City UV monitoring network: UV-B measurements at ground 

level in the urban environment. J Geophys Res, 105: 5017-5026. http://dx.doi.org/10.1029/1999JD900250 670420 

Adamowicz, V.; Dales, R.; Hale, B. A.; Hrudey, S. E.; Krupnick, A.; Lippman, M.; McConnell, J.; Renzi, P. (2004). Report 

of an expert panel to review the socio-economic models and related components supporting the development of 

Canada-Wide Standards (CWS) for particulate matter (PM) and ozone to the Royal Society of Canada. J Toxicol 

Environ Health B Crit Rev, 7: 147-266. http://dx.doi.org/10.1080/10937400490253238 670421 

Airey, D. K.; Wong, J. C. F.; Fleming, R. A.; Meldrum, L. R. (1997). An estimate of the UV-B exposure for outdoor 

workers during a south-east Queensland summer. Health Phys, 72: 544-549. 057458 

Aphalo, P. J.; Vapaavuori, E. M.; de la Rosa, T. M.; Lehto, T. (2009). Does supplemental UV-B radiation affect gas 

exchange and RuBisCO activity of Betula pendula Roth. seedlings grown in forest soil under greenhouse 

conditions? Plant Ecol Divers, 2: 37-43. http://dx.doi.org/10.1080/17550870902780299 668923 

Arrhenius, S. (1896). On the influence of carbonic acid in the air upon the temperature of the ground. Philos Mag, 41: 237­

276. http://www.rsc.org/images/Arrhenius1896_tcm18-173546.pdf. 043125 

Autier, P.; Dore, J. F.; Negrier, S.; Lienard, D.; Panizzon, R.; Lejeune, F. J.; Guggisberg, D.; Eggermont, A. M. (1999). 

Sunscreen use and duration of sun exposure: a double-blind, randomized trial. J Natl Cancer Inst, 91: 1304-1309. 

057459 

Autier, P.; Dore, J.-F.; Reis, A. C.; Grivegnee, A.; Ollivaud, L.; Truchetet, F.; Chamoun, E.; Rotmensz, N.; Severi, G.; 

Cesarini, J. P.; EORTC Melanoma Co-operative Group (2000). Sunscreen use and intentional exposure to 

ultraviolet A and B radiation: a double blind randomized trial using personal dosimeters. Br J Cancer, 83: 1243­

1248. 057069 

Bais, A. F.; Kazantzidis, A.; Kazadzis, S.; Balis, D. S.; Zerefos, C. S.; Meleti, C. (2005). Deriving an effective aerosol 

single scattering albedo from spectral surface UV irradiance measurements. Atmos Environ, 36: 1093-1102. 

http://dx.doi.org/10.1016/j.atmosenv.2004.09.080 669135 

Balis, D. S.; Zerefos, C. S.; Kourtidis, K.; Bais, A. F.; Hofzumahaus, A.; Kraus, A.; Schmitt, R.; Blumthaler, M.; Gobbi, G. 

P. (2002). Measurements and modeling of photolysis rates during the photochemical activity and ultraviolet 

radiation (PAUR) II campaign. J Geophys Res, 107: 8138. http://dx.doi.org/10.1029/2000JD000136 055023 

Bergstrom, R. W.; Pilewskie, P.; Pommier, J.; Rabbette M.; Russell, P. B.; Schmid, B.; Redemann, J.; Higurashi, A.; 

Nakajima, T.; Quinn, P. K. (2004). Spectral absorption of solar radiation by aerosols during ACE-Asia. J Geophys 

Res, 109: D19S15. http://dx.doi.org/10.1029/2003JD004467 669158 

Berntsen, T. K.; Myhre, G.; Stordal, F.; Isaksen, I. S. A. (2000). Time evolution of tropospheric ozone and its radiative 

forcing. J Geophys Res, 105: 8915-8930. http://dx.doi.org/10.1029/1999JD901139 047916 

Boer, G. J.; Yu, B. (2003). Climate sensitivity and response. Clim Dynam, 20: 415-429. http://dx.doi.org/10.1007/s00382­

002-0283-3 630271 

Bony, S.; Colman, R.; Kattsov, V. M.; Allan, R. P.; Bretherton, C. S.; Dufresne, J. L.; Hall, A.; Hallegatte, S.; Holland, M. 

M.; Ingram, W.; Randall, D. A.; Soden, B. J.; Tselioudis, G.; Webb, M. J. (2006). How well do we understand and 

evaluate climate change feedback processes? J Clim, 19: 3445-3482. 

http://journals.ametsoc.org/doi/abs/10.1175/JCLI3819.1. 630272 

Brandt, L. A.; Bohnet, C.; King, J. Y. (2009). Photochemically induced carbon dioxide production as a mechanism for 

carbon loss from plant litter in arid ecosystems. J Geophys Res, 114: G02004. 

http://dx.doi.org/10.1029/2008jg000772 668891 

Note: Hyperlinks to the reference citations throughout this document will take you to the NCEA HERO database (Health and 

Environmental Research Online) at http://epa.gov/hero. HERO is a database of scientific literature used by U.S. EPA in the process of 

developing science assessments such as the Integrated Science Assessments (ISAs) and the Integrated Risk Information System (IRIS). 

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Breuer-McHam, J.; Simpson, E.; Dougherty, I.; Bonkobara, M.; Ariizumi, K.; Lewis, D. E.; Dawson, D. B.; Duvic, M.; 

Cruz, P. D. Jr (2001). Activation of HIV in human skin by ultraviolet B radiation and its inhibition by NF"kappa"B 

blocking agents. Photochem Photobiol, 74: 805-810. 057736 

Brewer, P. G.; Peltzer, E. T. (2009). Limits to marine life. Science, 324: 347-348. http://dx.doi.org/10.1126/science.1170756 

669197 

Brown, B. A.; Jenkins, G. I. (2008). UV-B signaling pathways with different fluence-rate response profiles are 

distinguished in mature Arabidopsis leaf tissue by requirement for UVR8, HY5, and HYH. Plant Physiol, 146: 576­

588. http://dx.doi.org/10.1104/pp.107.108456 668892 

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