Current Pollution Reports

, Volume 1, Issue 2, pp 57–69 | Cite as

Paleo-Perspectives on Potential Future Changes in the Oxidative Capacity of the Atmosphere Due to Climate Change and Anthropogenic Emissions

  • Becky Alexander
  • Loretta J. Mickley
Air Pollution (S Wu, Section Editor)
Part of the following topical collections:
  1. Topical Collection on Air Pollution


The oxidizing capacity of the atmosphere, defined as the global mean tropospheric abundance of the hydroxyl radical (OH·), strongly influences air pollution by controlling the lifetimes of gaseous pollutants and the production of particulate matter. Predicting future changes in OH· due to anthropogenic emissions and climate change is of interest to air quality managers, but it is difficult because of multiple competing effects. Models of atmospheric chemistry suggest that these competing effects buffer significant change in OH· in the past and in the near future. However, proxy-based observations for past changes in OH· and other oxidants over the preindustrial-industrial and glacial-interglacial time scales suggest much larger changes than models estimate. Model sensitivity studies show that variability in past and future OH· is highly sensitive to relative emissions of reactive nitrogen and carbon, water vapor, lightning, and stratospheric ozone, implying that one or more of these variables is highly sensitive to climate.


Hydroxyl radical Ozone Atmospheric chemistry 


The atmosphere plays a key role in biogeochemical cycles on Earth by oxidizing reduced trace gases emitted from natural and anthropogenic sources. The oxidizing capacity of the atmosphere is nominally defined as the global mean tropospheric abundance of the hydroxyl radical (OH·). Other oxidants, such as the nitrate radical (NO3·), ozone (O3), hydrogen peroxide (H2O2), the sum of peroxy radicals (HO2· + RO2·), and reactive halogens (e.g., Cl·, BrO·) also play a key role in the oxidizing capacity of the atmosphere, and the abundance of each oxidant is influenced by the other oxidants through highly non-linear oxidant cycling reactions (Fig. 1). For example, O3 acts as a precursor to OH·, and the oxidation of CH4 by OH· provides a source of RO2·. The oxidizing capacity of the atmosphere affects air quality by determining the lifetime and formation of gas-phase pollutants such as carbon monoxide (CO) and O3 and by contributing to the formation of particulate matter via gas-to-particle conversion. Thus, information about how the oxidizing capacity of the atmosphere responds to human activity and natural phenomena is of key interest for air pollution concerns.
Fig. 1

Schematic of the coupling between HO x (= OH· + HO2· + RO2·), NO x (= NO + NO2), and reactive halogens (X = Br, Cl) and impact on OH·, O3, and H2O2 production in the troposphere. Parameters in blue (where NMHC = non-methane hydrocarbons) lead to OH· destruction, while parameters in red lead to OH· production

An understanding of potential future changes in the oxidizing capacity of the atmosphere is also relevant for air quality managers, as variability in the oxidizing capacity of the atmosphere may influence the ability of any given region to remain in compliance with local air quality standards. The abundance of OH· may change due to changes in anthropogenic emissions of pollutants or changes in climate. Anthropogenic emissions of NOx ( = NO + NO2), originating primarily from fossil fuel combustion [1], drive production of both tropospheric O3 and OH·. Anthropogenic emissions of reduced carbon species such as CO (from fossil-fuel production) and methane (CH4, from agriculture and other activities) serve as a source of O3 but a sink of OH·. As a result of increasing anthropogenic emissions of NO x and reduced carbon gases, O3 concentrations have increased over the twentieth century [2, 3], while global OH· concentrations are thought to have remained relatively stable, at least for the last 30 years [4]. The relative stability of OH· is due to the buffering effect of simultaneous increases in anthropogenic OH· sources (NO x ) and sinks (reduced carbon gases) [5].

The response of OH to the current and future warming climate resulting from increases in anthropogenic greenhouse gases is unclear due to competing effects. OH· may increase in a future, warmer world due to increases in water vapor abundance and in convection and thus lightning NO x [6, 7]. On the other hand, OH· may decrease due to increases in temperature-dependent emissions of biogenic hydrocarbons and rate constants for OH·-destruction reactions [8]. Anthropogenic land use change could also perturb biogenic emissions, with consequences for OH [9]. OH· would also be sensitive to changes in stratospheric O3 due to its influence on ultraviolet (UV)-induced photochemical reactions in the troposphere [10].

Projections of the response of the tropospheric oxidative capacity to future change have typically relied on three-dimensional chemistry-climate modeling frameworks (e.g., Naik et al. [11]). Future scenarios of greenhouse gases and anthropogenic land use change are applied to these models, together with emissions scenarios of short-lived gases and aerosols of importance to OH· and other oxidants. Many future scenarios have been developed with support of the Intergovernmental Panel on Climate Change (IPCC) [12] and have permitted cross-model comparison, thus deepening our awareness of uncertainties in model assumptions and processes. To date, these cross-model comparisons have revealed large discrepancies in projections of future oxidative capacity (e.g., Voulgarakis et al. [13]).

One solution for dealing with uncertainties in future model predictions is to look to the past as an analogue. Future changes in oxidant abundances will likely be very different from what occurred in the past due to different or even opposite forcings. For example, past climates are characterized by minimal anthropogenic impacts and, for the last ~100,000 years, cooler temperatures. However, knowledge of past variability in tropospheric oxidants can provide information on the sensitivity of tropospheric oxidant abundance over a variety of time scales, and thus inform potential future changes. Here, we review current knowledge of the oxidation capacity of the atmosphere from the paleo-perspective, focusing on both observational constraints and modeling results. This review is followed by a discussion of the implications of what we have learned from the paleo-perspective for understanding the key uncertainties and for informing estimates of possible future changes in the oxidizing capacity of the atmosphere. We end by discussing the most promising future directions for improving our understanding of potential changes in the oxidative capacity of the atmosphere.

A Paleo-Perspective on Changes in the Oxidation Capacity of the Atmosphere

Paleo Observations

Polar ice-core measurements of long-lived greenhouse gases such as carbon dioxide (CO2) and CH4 and aerosols such as sulfate and nitrate provide evidence of the extent and degree of human influence on the chemical composition of the atmosphere. Anthropogenic activities, such as the burning of fossil fuels, have increased the concentration of CO2 and CH4 by 43 and 150 %, respectively [14, 15, 16]. Concentrations of CO2 and CH4 during the Last Glacial Maximum (LGM) were 30 % and 40–50 % of their preindustrial values, respectively [17, 18, 19]. Aerosol concentrations from Greenland ice cores for the 1970–1990s show increases in nitrate of ~100 % and in sulfate of 320 % since the preindustrial era, demonstrating the impacts of fossil fuel burning [20, 21]. Greenland ice-core nitrate and sulfate concentrations have leveled off or declined since 1990 due to emissions reductions in the USA [22]. In contrast, continued increases in sulfate concentrations over the past 50-years as observed in a Tibetan ice core [23] are consistent with increasing trends in Asian emissions of SO2 [22]. The lack of similar increases in Antarctic ice cores [24] shows that anthropogenic influence on aerosol abundance is by far most important in the Northern Hemisphere. Oxidants such as OH· and O3, however, are not directly preserved in ice cores, so much less is known about how human activity has influenced their abundances. Here, we briefly review potential proxies for tropospheric oxidants. Table 1 provides a summary of the observational constraints on changes in oxidant abundances in the present day and LGM relative to the preindustrial Holocene, based on proxies that are thought to be well-preserved during the time scale of interest as discussed below.
Table 1

Percent changes in oxidant abundances in the present day (PD) and last glacial maximum (LGM) relative to the preindustrial Holocene (PIH), inferred from observations and calculated by models


Change in present day (PD – PIH)/PIH

Change at LGM (LGM – PIH)/PIH





Surface O3

+200 to +300 % (Schönbein method, ~19 sites worldwide) [2, 25, 26, 27, 28]

+11 to +48 % (global) [11, 29, 30, 31, 32, 33]


−17 to −31 % (global) [30, 34, 35]


+60 % (Greenland ice cores) [36, 37, 38]

+50 % (Antarctica ice cores) [39, 40]

+31 to +58 % (global) [29, 30]


−36 ± 21 % (global) [30]


−20 % (inferred from Greenland ice-core HCHO/CH4) [40]

−16 to +7 % (global) [11, 29, 30, 31, 33, 41]

> +50 % (inferred from Greenland and Antarctic ice-core CH4) [42]

−29 to +28 % (global) [30, 34, 35, 43]

+68 to +120 % (46–66°S) [44]


−60 to −90 % (inferred from Antarctic ice-core nitrate Δ17O) [24]



+ (inferred from Antarctic ice-core sulfate Δ17O) [24]

+35 to +42 % (46–66°S) [44]

− (inferred from Antarctic ice-core sulfate Δ17O) [45]




+35 to +42 % (46–66°S) [44]


Except for surface ozone, observations are from the mid- to high latitudes. Modeled values represent tropospheric annual mean changes, except as noted, and are from the recent (last decade) literature only


In the late nineteenth century, O3 abundance was measured around the world using either the Schönbein method (e.g., [25]) or a wet-chemical method involving the oxidation of arsenite in neutral aqueous solution [26]. The Schönbein method employs paper strips impregnated with potassium iodide and starch and is sensitive to artefacts from humidity and other oxidants [26, 27]. The arsenite method is also subject to both positive artefacts from other oxidants such as H2O2 and negative artefacts from reducing gases such as SO2 [26]. Volz and Kley [26] reproduced and recalibrated the most reliable of the nineteenth-century observations, those from Paris from 1876–1910 [2]. Their work suggests very low preindustrial surface O3 concentrations on the order of 10 ppbv, implying a doubling of background surface O3 abundance in rural areas over this 100-year period [26]. Although the accuracy of these nineteenth-century observations has been debated [28, 46], they suggest at least a factor of 2–3 increase in surface O3 in the late twentieth century, compared to its preindustrial abundance [2].

Hydrogen Peroxide

Hydrogen peroxide (H2O2) is the only oxidant that can be directly measured in polar ice cores and was first measured in both Greenland and Antarctic ice by Neftel et al. [47]. Hydrogen peroxide forms through the self-reaction of HO2·, which in turn is formed through the oxidation of reduced trace gases such as CO and CH4 by OH·. Once formed, H2O2 has an atmospheric lifetime of a few days and is lost from the atmosphere via photolysis, chemical reactions in the gas- (e.g., H2O2 + OH·) and aqueous-phase (e.g., HSO3  + H2O2), and wet and dry deposition to the Earth’s surface. H2O2 is reversibly deposited to snow, meaning it can evaporate back into the overlying atmosphere [48, 49] or undergo ultraviolet (UV)-induced destruction (photolysis) [50, 51] and production [52, 53]. Dust may also catalyze H2O2 decomposition in the ice at depth [54]. This so-called post-depositional processing complicates interpretation of ice-core H2O2 concentrations. Frey et al. [55] suggest that variability in ice-core H2O2 concentrations may represent atmospheric variability only under conditions of relatively stable temperatures and UV radiation at locations with low temperatures (<−31 °C) and high snow accumulation rates (>30 cm year−1).

H2O2 measurements from Greenland and Antarctic ice cores show 50–60 % increase over the past 100–200 years, with most of that increase occurring since the early 1970s [36, 37, 38, 39, 40]. Möller [56] interprets most of the variability in the Greenland record to a changing sink from aqueous-phase reaction of H2O2 with dissolved SO2 (HSO3 ) from coal combustion, with a slight decrease in H2O2 from 1850–1970 as SO2 anthropogenic emissions rose [22]. Implementation of the Clean Air Act of 1970 and its amendments led to a decline in SO2 emissions and a subsequent increase in H2O2. In contrast, the increase in Antarctic ice-core H2O2 is thought to be driven by an increase in local, tropospheric photolysis rates resulting from decreases in overhead (stratospheric) O3-column abundance, with a smaller but significant impact from surface O3 abundance [39].

The Isotopes of CH4

CH4 is well-preserved in gas bubbles in ice cores, providing a rich history of its atmospheric abundance in the past. The main natural source of CH4 is from wetlands, with smaller sources from ruminants, termites, thermokarst lakes, and biomass burning [57]. Major anthropogenic sources are from fossil-fuel extraction, agriculture (primarily rice), ruminant livestock, and waste treatment [57]. The main (~85 %) sink for CH4 in the troposphere is oxidation by OH· [57]. Minor sinks include oxidation in the stratosphere and in soils [57, 58]. The carbon and hydrogen isotopic composition of CH4 is determined by its relative source strengths, as different sources have different isotopic signatures, and the strength of its sinks, which all act to enrich the isotopic composition of remaining atmospheric CH4. Oxidation of CH4 by OH· leads to large enrichments in the hydrogen isotopes (δD) on the order of 200 ‰ [59, 60] and smaller enrichments in the carbon isotopes (δ13C) on the order of 4–5 ‰ [60, 61] relative to what was emitted. Although tropospheric Cl· is thought to be only a minor sink (3–4 %) for CH4 [57, 62, 63], it induces an isotopic fractionation more than an order of magnitude larger than does reaction with OH· [64, 65]. The small soil sink (5 % [57]) of CH4 induces an isotopic fractionation about four times larger than the OH· sink [66, 67].

As with ice-core CH4 (e.g., Chappellaz et al. [68]), the observed variability in CH4 isotopes from Craig et al. [69], and many subsequent analyses were interpreted by assuming a constant sink of CH4 or by making a priori assumptions about changes in the sink strength [14, 70, 71, 72, 73, 74, 75, 76, 77, 78]. Fischer et al. [79], however, determined that the most likely explanation for the observed increase in δ13C(CH4) and δD(CH4) in the LGM compared to the preindustrial Holocene was a significantly reduced (>50 %) atmospheric lifetime of CH4 in the LGM. Though a unique interpretation of ice-core methane isotope records is not possible [73], simultaneous measurements of 13C/12C and D/H of CH4 may be used to examine possible scenarios and place bounds on past [72, 80] and present [81, 82] variability in the relative importance of the OH· sink.


Formaldehyde (HCHO) is an oxidation product of the CH4 + OH· reaction, and in remote, unforested regions, the main source of HCHO is from oxidation of CH4 by OH·. HCHO is lost from the atmosphere via photolysis and oxidation by OH·. Due to interactions between CH4, HCHO, and OH·, the HCHO/CH4 ratio in ice cores could potentially serve as a proxy for past OH abundance, with higher values indicating higher OH· [42]. Like H2O2, HCHO is reversibly deposited to ice sheets via evaporation [83] and photochemical loss [84], hindering its interpretation in terms of past atmospheric abundance. Assuming that the transfer of HCHO between the snow/ice and the overlying atmosphere was constant over time, observations of HCHO/CH4 from Greenland snow and firn suggest that OH· was 30 % higher in the preindustrial, qualitatively consistent with HCHO observations from Law Dome, Antarctica [40]. HCHO observations from a South Greenland and Antarctic (Byrd) ice core imply, in turn, that OH· was a factor of 2–4 lower in the LGM relative to the preindustrial [42]. In contrast, ice cores from central Greenland (Summit) [38] and Antarctica (Law Dome) [40] show higher HCHO concentrations in the last glacial period, with HCHO variability at Summit paralleling that of calcium (Ca2+). However, the large change in snow accumulation rates in Antarctica and Greenland on the glacial-interglacial time scale likely precludes constant preservation of HCHO.

Nitrate Concentration and Its Nitrogen Isotopes

Nitrate (NO3 ) is a major sink for atmospheric NO x and its concentration in ice cores can provide insight into past variability in NO x concentrations. Since NO x plays an important role in the formation and cycling of tropospheric oxidants (Fig. 1), information about its past variability can shed light on the oxidation capacity of paleo atmospheres. The atmospheric lifetime of NO x against oxidation to nitrate is on the order of 1 day [85], and the atmospheric lifetime of nitrate is on the order of 3–4 days, making polar ice-core nitrate concentrations an indicator of mid- to high-latitude atmospheric variability. The (83 ± 34) % increase in Greenland ice-core nitrate concentrations [20] is smaller than the ~150 % increase in fossil-fuel NO x emissions between 1940–1970 from North America [86], the main aerosol source region to Summit, Greenland. The difference is likely due to some depositional loss of nitrate to the surface during transport and possibly also in part due to post-depositional loss from the Greenland snowpack (see below). The lack of a similar increase in nitrate concentrations in a West Antarctic ice core [24] suggests little change in NO x abundance in the mid- to high-southern latitudes due to anthropogenic activities. Antarctic ice-core records show higher nitrate concentrations in the glacial period relative to the preindustrial Holocene [87, 88], while Greenland ice-core records snow little glacial-interglacial variability [89, 90].

The interpretation of ice-core nitrate concentration records during times of major climate change has been hampered by the influence of post-depositional processing of snow nitrate. Snow nitrate is lost via evaporation [91] and photolysis [92] and the degree of post-depositional loss is determined mainly by the snow accumulation rate and the concentration of UV-absorbing impurities in the snow [93, 94]. Higher snow accumulation rates lead to less post-depositional loss by more quickly burying snow nitrate below the photic zone. During periods when the snow accumulation rate and the concentration of UV-absorbing snow impurities have not changed significantly, such as over the past several hundreds [95] to several thousands [24] of years, variability in ice-core nitrate concentrations likely represents atmospheric variability of nitrate and NO x . However, during periods of major changes in snow accumulation rate, such as over glacial-interglacial cycles, variability in ice-core nitrate concentrations may not reflect its past atmospheric abundance.

The photolysis of snowpack nitrate induces a large fractionation on its nitrogen isotopes (δ15 N(NO3 )) on the order of 50 ‰, emitting isotopically light NO x into the overlying boundary layer and leaving the nitrate remaining in the snow enriched in δ15N by up to 350 ‰ [94, 96, 97]. As the enrichment in ice-core δ15N(NO3 ) is proportional to the fraction of nitrate lost from the snow due to photolysis, simultaneous measurement of ice-core nitrate concentrations and δ15N(NO3 ) may allow for a semi-quantitative reconstruction of atmospheric nitrate concentrations. Such a reconstruction may be feasible in regions such as across much of the East Antarctic plateau where snow accumulation rates are low and the variability in δ15N(NO3 ) is driven mainly by photolytic loss [93, 94].

Oxygen Isotopes of Sulfate and Nitrate

The mass-independent fractionation evident in atmospheric O3 affects the oxygen isotopic composition (Δ17O = δ17O – 0.52 × δ18O) of nitrate and sulfate in the atmosphere and provides a tool to assess their formation pathways. O3-dominated reactions lead to high Δ17O values in the nitrate and sulfate products, while reactions dominated by HO x (OH·, HO2· + RO2·, or H2O2) lead to low Δ17O. The value of Δ17O in nitrate and sulfate from ice cores can thus provide information on the relative abundance of O3/HO x in paleo atmospheres.

Nitrate formation occurs via oxidation of NO x by HO x and O3. The Δ17O of NO x is determined by the relative abundance of O3/(HO2· + RO2·) during NO x cycling reactions, while the Δ17O of nitrate (Δ17O(NO3 )) is determined by the Δ17O of NO2 and the oxidant (OH or O3) involved in the oxidation of NO2 to HNO3 [98]. Since Δ17O(NO3 ) is most sensitive to the Δ17O of NO2, the variability of Δ17O(NO3 ) can to first order be interpreted as variability in the relative O3/(HO2· + RO2·) abundance.

Photolytic-driven loss of snow nitrate itself does not directly impact Δ17O(NO3 ) since it involves mass-dependent isotopic behavior. However, some of the snow-sourced NO x is re-oxidized and re-deposited to the snowpack, and this recycling tends to erase the oxidant signature in Δ17O(NO3 ) derived from the primary NO x -source regions, replacing it with a local oxidant signature. In regions of low snow accumulation rates and high photolytic-driven recycling such as the East Antarctic plateau, ice-core Δ17O(NO3 ) measurements [94] likely indicate a local oxidant signature. In West Antarctica and Greenland, where snow accumulation rates are higher, the oxidant signature from the primary NOx-source regions is more likely preserved [24, 99].

Ice-core Δ17O(NO3 ) measurements from a Greenland ice core indicate a peak in Δ17O(NO3 ) beginning in the mid-1800s and ending in the early 1900s [100]. This peak coincided with indicators of biomass burning records from a Greenland ice core [101], suggesting that biomass burning in North America increased the relative abundance of O3 over (HO2· + RO2·) in that region. In contrast, a 2400-year ice-core record from West Antarctica shows a gradual decrease in Δ17O(NO3 ) over most of the record, with a more rapid decrease between the mid-nineteenth century and the present day [24]. The long-term gradual decrease in Δ17O(NO3 ) is likely due to an increase in the amount of photolytic-driven recycling, consistent with a decrease in snow accumulation rate over this same time period. The rapid recent decrease in Δ17O(NO3 ) over the last 150 years may be due to the estimated 60–90 % decrease in the O3/(HO2· + RO2·) ratio in the extratropical Southern Hemisphere, when anthropogenic activities increased the sum of (HO2· + RO2·) more rapidly than O3.

Sulfate formation occurs via oxidation of SO2 in the gas-phase through oxidation by OH· and in cloud droplets by H2O2 and O3. Due to its atmospheric lifetime of 5–6 days [102], ice-core Δ17O(SO4 2−) measurements can provide information about the mid- to high-latitude atmosphere. Unlike nitrate, sulfate does not undergo post-depositional processing and both its concentration and isotopic composition are well-preserved throughout the ice-core record. Variability in the relative abundance of the oxidants OH·, H2O2, and O3 influences Δ17O(SO4 2−) [29] because each oxidant possesses a unique Δ17O value that is transferred to sulfate during its formation [103]. However, clouds also affect Δ17O(SO4 2−) by influencing the relative importance of gas- versus aqueous-phase production. In addition, as the oxidation rate of dissolved SO2 by O3 is highly pH dependent, cloud pH determines the relative importance of O3 versus H2O2 oxidation.

In their analysis of a West Antarctic ice core, Sofen et al. [24] detected a step increase in Δ17O(SO4 2−) of 1.1 ‰ in the early nineteenth century, a time when no significant change in climate, ice-core pH, and sulfate and H2O2 concentrations occurred. The cause of the increase is unclear [24].

Alexander et al. [100] analyzed oxygen isotopes in a Greenland ice core and found a peak in both Δ17O(SO4 2−) and Δ17O(NO3 ) beginning in the mid-1800s and ending in the early 1900s. This peak coincided with indicators of biomass burning [101], suggesting that enhanced O3 from agricultural clearing in North America increased the importance of O3 oxidation of SO2 and NO x . In addition, Δ17O(SO4 2−) values in 1980 were slightly higher than values before 1800, implying an increased importance for O3 oxidation in the industrial era compared to preindustrial times.

Discrete measurements of Δ17O(SO4 2−) from an East Antarctic ice core (Vostok) over a full glacial-interglacial cycle (~130 kyr before present) show significantly lower values during the glacial period compared to the interglacial periods before and after [45]. The observed variability in Δ17O(SO4 2−) was interpreted to be driven by a 40 % increase in the amount of sulfate formed by OH· oxidation relative to O3 oxidation. However, the role of cloud fraction and liquid water content in driving the observed variability in Δ17O(SO4 2−) on this time scale could not be discerned with available observations.

Oxygen Isotopes of Carbon Monoxide

Carbon monoxide (CO) originates from incomplete combustion of fossil fuels and biomass and from the oxidation of reduced carbon gases such as CH4 and biogenic hydrocarbons [104]. Its tropospheric lifetime of ~2 months is determined by OH· oxidation [105]. CO serves as the largest (70 % [106]) sink of OH· and is an O3 precursor, rendering its abundance an important control on the oxidation capacity of the atmosphere. Similar to the long-lived greenhouse gases CO2 and CH4, CO is preserved in air bubbles in ice cores. In situ production of CO in Greenland ice cores has been observed [107] and is thought to originate from photo-oxidation of organic material [107, 108, 109], complicating the reconstruction of past CO in the Northern Hemisphere. Alternative explanations for in situ CO production involve the influence of cosmic rays [110] or microbial activity [107]. In situ CO production in Antarctic snow and ice is thought to be minimal in part due to much lower concentrations of organic material [111]. Antarctic firn and ice-core observations of CO abundance since 250 BC show century-long trends that may reflect changes in biomass-burning sources [e.g., 60, 105]. Isotopic measurements of δ18O(CO) and δ13C(CO) are consistent with this interpretation [111, 112], as these isotopic signatures are determined in part by CO sources. Δ17O(CO) is mainly influenced by its degree of reaction with OH· [113, 114]. Δ17O(CO) has not yet been measured in ice cores, but its dependence on OH· abundance and its preservation in Antarctic ice renders it a potential proxy for the oxidation capacity of paleo atmospheres [115].


Table 1 summarizes model estimates of the changes in oxidant abundances in the present day and LGM relative to the preindustrial Holocene published within the last decade [11, 29, 30, 31, 32, 33, 34, 35, 41, 43, 44, 116]. A complete list of all model results to date, except for the recent results of Achakulwisut et al. [44] and John et al. [116], are in Murray et al. [30]. Differences between models are due mainly to differences in assumptions about past emissions and climate and are also sensitive to assumptions about HO x recycling during isoprene oxidation [44, 117].

Table 1 reveals that there is no model consensus on the sign of the change in OH· in the LGM or the present day, relative to the preindustrial. Most models calculate decreases or no change in present day OH· relative to the preindustrial due to increases in OH· sinks (CO and CH4) that are partly or completely compensated by increases in NO x emissions. All models calculate an increase in present day O3 and H2O2 [11, 29, 30, 31, 32, 33, 41]. Modeled increases in O3 are much smaller than the Schönbein measurements [2, 25, 26, 27, 28] suggest because the models cannot reproduce the low preindustrial O3 abundance. Parrella et al. [32] show that inclusion of bromine chemistry in model calculations lowers preindustrial O3 to be more in line with the Schönbein observations, though the directional impact of reactive bromine on tropospheric O3 abundance depends on assumptions in the chemical mechanism [118]. Modeled decreases in present day OH· [11, 29, 31, 33, 116] are smaller than the Greenland HCHO/CH4 observations imply [40]. Modeled increases in present-day H2O2 [29, 30] are similar to those in the Greenland and Antarctic observations [36, 37, 38, 39, 40]. Modeled changes in the O3/OH· ratio over the Southern Ocean [30, 44] are qualitatively consistent with ice-core observations of sulfate Δ17O [24]. Finally, modeled changes in the O3/HO2· ratio over South America [44] are in the right direction but much smaller than nitrate Δ17O observations suggest [24].

Modeled LGM OH· abundances tend to be higher than in the preindustrial Holocene [34, 35, 43], though the modeled increases are much smaller than the ice-core CH4 isotopes suggest [79]. Modeled LGM increases in OH· and decreases in O3 over the Southern Ocean [30, 44] are qualitatively consistent with ice-core sulfate Δ17O observations [45].

Implications for the Future

Drivers of Change in the Oxidative Capacity of the Future Atmosphere

The drivers of the oxidative capacity of the troposphere will likely change in future decades. Warmer surface temperatures will enhance evaporation, thereby increasing water vapor and OH· production [119]. Lightning frequency will also change, but the sign and magnitude of that change are uncertain, since convection depends not just on moisture content but also the vertical temperature gradient [120]. Greenhouse gas cooling in the lower stratosphere could accelerate the Brewer-Dobson circulation, reducing the overhead O3 column in the tropics but thickening the column over the extra-tropics [121]. Future changes in ozone-depleting substances will also impact the stratospheric O3 abundance [122]. Such changes in overhead O3 would influence tropospheric photolysis rates as well as the stratospheric flux of O3 into mid-latitudes. Changes in aerosol abundance resulting from changing emissions of aerosol precursors or lifetimes will also impact oxidants through heterogeneous reactions of N2O5 [123] and HO2 [124] on their surface.

Changes in the biosphere will also perturb the oxidative capacity. Widespread deforestation could affect oxidants by disrupting the hydrological cycle and reducing biogenic emissions [9, 125]. Warmer temperatures, on the other hand, will enhance biogenic emissions, though the effect of increasing CO2 on plant physiology may modulate that increase [126]. Wetland and peatland emissions of methane are likely to rise as microbial activity accelerates in a warming climate, and such trends may already be occurring [127]. Depending on the ecosystem and local meteorology, wildfire frequency could increase in future decades, enhancing emissions of CO, hydrocarbons, and NO x [128].

Finally, trends in the anthropogenic emissions of O3 and aerosol precursors could have large effects on the oxidative capacity. Future scenarios of precursor emissions have been devised to probe the response of atmospheric composition to trends in human activity. For example, the Representative Concentration Pathway (RCP) scenarios were designed to achieve the radiative forcing targets of 2.6, 4.5, 6.0, and 8.5 W m−2 by 2100 and have been designated RCP2.6, RCP4.5, RCP6.0, and RCP8.5, respectively [12]. In all RCP scenarios, global anthropogenic NO x emissions decline 30–50 % over the course of the twenty-first century. Trends in global anthropogenic hydrocarbon and CO emissions are flat during the 2000–2050 time range, and then also show declines by 30–50 % by 2100. While the global methane burden for three of the RCP scenarios decreases by 2100, RCP8.5 projects a doubling of the methane over the century. Finally, in all scenarios, concentrations of chlorofluorocarbons and other O3-depleting substances in the stratosphere decrease substantially over the twenty-first century.

Tests of Modeled Oxidative Capacity over the 1980–2010 Timeframe

Multi-model time slice simulations conducted in support of the Atmospheric Chemistry Climate Model Intercomparison Project (ACCMIP) for 1980 and 2000 were analyzed to test our understanding of the impact of climate change and anthropogenic emissions on the oxidative capacity of the atmosphere [11]. Results from this multi-model study can be compared to the OH· global abundance inferred from observations of methyl chloroform, a trace gas whose main sink is tropospheric OH·. The models suggest that global OH· has increased 3.5 ± 2.2 % since 1980, mainly due to increasing NO x emissions. Observations, however, imply a possible decrease in global mean OH· in the 1980–2000 timeframe [129, 130], with little trend thereafter [4]. Comparison of the modeled interhemispheric ratios (N/S) of tropospheric mean OH· to that inferred from observations reveals another discrepancy, with models calculating an average N/S ratio of 1.28 and observation-derived ratios ranging between 0.85–1.0 [11, 131]. These spatial and temporal discrepancies between modeled and inferred OH· may arise from model underestimates of CO concentrations in the Northern Hemisphere, incomplete understanding of the photochemistry of biogenic species, or neglect of the heterogeneous uptake of HO2· on aerosols [11, 44].

Projections of the Oxidative Capacity in the Twenty-First Century

Model studies have been extended beyond 2000 in an effort to characterize the effects of changing climate and emissions on the tropospheric oxidative capacity. Such studies indicate that climate change alone could increase global mean OH· by 4–13 % by 2050 mainly because of higher water vapor abundance and increased lightning NO x frequency (Wu et al. [119], and references therein). Most of these studies, however, did not consider the effect of increased methane emissions from wetlands and peatlands, which could be appreciable [132] but is highly uncertain [133, 134].

When changes in anthropogenic emissions of key precursors are considered, the impact on OH· depends on the emission scenario. In three of the RCPs, the effects of declining anthropogenic NO x and hydrocarbons tend to cancel each other, allowing the influence of climate change to dominate the OH· trend. In RCP8.5, however, the large rise in anthropogenic methane swamps the climate effect and leads to strong decreases. For example, John et al. [116] determined that greater water vapor abundance by 2100 enhances global OH· by 7–9 % relative to the present day for RCP2.6 and RCP4.5, with little change in RCP6.0. In contrast, global OH· declines by nearly 15 % in the CH4-rich RCP8.5 scenario over the same time period. The multi-model study of Voulgarakis et al. [13], however, reveals large uncertainties in such projections, with intermodel differences as large as the differences between scenarios. While the three scenarios with the lowest target radiative forcing all project modest increases in OH· by 2100, the uncertainties swamp the median result [13]. RCP2.6, for example, yields a change in global OH· of +3.7 ± 9.0 % by 2100 relative to the present day. These intermodel differences are smaller in the RCP8.5 scenario on account of the strong influence of the doubled methane burden, leading to a decrease in global OH· of −11.3 ± 7.7 % by 2100. The decrease in OH· in RCP8.5 lengthens the methane lifetime, resulting in a positive feedback for this key greenhouse gas [13, 135].

Besides OH·, the other oxidants are also expected to respond to global change. Results from multi-model ensembles suggest that the tropospheric O3 burden may decrease by 5–16 % by 2100 in RCP2.6, RCP4.5, and RCP6.0, as anthropogenic emissions of NO x and hydrocarbons decline [33, 136]. In contrast, models predict that tropospheric O3 could increase by 15–20 % in RCP8.5 due to the doubling of methane [33, 136, 137]. Consideration of the effects of increasing stratospheric O3 also enhances tropospheric O3, mainly in the Northern Hemisphere [121]. Trends in H2O2 in the future atmosphere are expected to follow trends in OH·, though warmer temperatures could accelerate H2O2 production via HO2· + HO2· → H2O2 [138]. NO3· concentrations will likely decline in the future atmosphere, following projections in anthropogenic NO x emissions [139]. Taken together, these projections of the oxidative capacity of future atmosphere underscore the uncertainty not just in current scientific knowledge but also in the socioeconomic paths that societies will follow.


Although all of the observational constraints for past atmospheres are uncertain to varying degrees, they suggest a common theme that models underestimate the sensitivity of oxidant (OH·, O3) abundances to anthropogenic and climate impacts (Table 1). In general, models tend to underestimate changes implied by a variety of paleo-observational constraints on both short (past ~200 years) and longer (past ~100,000 years) time scales. There is high confidence that anthropogenic activities have increased the abundance of tropospheric O3; however, models cannot reproduce the magnitude of the change implied by the observations over the past 100 years. While ice-core observations of HCHO/CH4 imply a decrease in OH· over the last 200 years, models cannot yet agree on either the sign or magnitude of the change in OH·. A recent global model [44] underestimates the decrease in the O3/(HO2· + RO2·) ratio inferred by ice-core observations of Δ17O(NO3 ) in South America over the past 150 years [24]. Few observational constraints exist during large climate changes such as on the glacial-interglacial time scale. However, as with recent time scales, global models cannot agree on the sign or magnitude of the change in global mean OH from the LGM to the preindustrial. Although highly uncertain, paleo-observations [45, 79] suggest 50 % higher global OH· abundance in the LGM relative to the preindustrial, an enhancement larger than any reported by models. This result contrasts with future modeling studies that predict higher OH· due to the effects of a warming climate alone (Wu et al. [8], and references therein).

Model sensitivity studies can point to the most important parameters that influence oxidant abundances on each time scale. Naik et al. [11] found that global OH· abundance is most sensitive to the relative emissions of NO x and CH4 over the 1980–2000 time frame. This result has important implications for future OH·, as estimates of CH4 emissions vary considerably between the different RCP scenarios. Murray et al. [30] found that three parameters can explain most of the modeled variability in global-mean OH· over the glacial-interglacial time scale: (1) stratospheric O3 abundance via its influence on the photolysis frequency of tropospheric O3, (2) water vapor, and (3) lightning emission of NO x , largely consistent with conclusions of Holmes et al. [135] who examined future sensitivities. Achakulwisut et al. [44] demonstrated that the response of global-mean OH· to climate change is particularly sensitive to uncertainties in isoprene oxidation chemistry. In light of these model studies, the large changes in oxidants in past atmospheres inferred from proxy data suggest that stratospheric O3, water vapor, and/or lightning NO x emissions may be much more sensitive to climate change than is currently estimated.

Can we gain information from the paleo-records about climate-induced variability of the most influential forcings for the oxidation capacity of the atmosphere? Polar stratospheric O3 abundance will influence ice-core δ15N(NO3 ) in regions of low snow accumulation rate such as the East Antarctic Plateau where variability in δ15N(NO3 ) is governed by UV-photolytic loss [93, 94]. Information on stratospheric O3 over polar regions can, in turn, inform our understanding of past variability in global stratospheric O3. Such a proxy could be easily calibrated by measuring δ15N(NO3 ) before and during the modern-day stratospheric O3 depletion. For lightning NO x , measurements of δ15N(NO3 ) could again prove useful, but from regions of high snow accumulation rates such as Greenland and West Antarctica. In these regions, ice-core δ15N(NO3 ) could track changes in the emission strength of lightning NO x , which is thought to have a relatively high δ15N signature [140], compared to that of other natural NO x sources such as soil microbes [141] and biomass burning [95]. However, care must be taken when interpreting such data sets since the signature of δ15 N from NO x sources may be overwritten by subsequent atmospheric or post-depositional processing [95]. Variability in water vapor abundance is governed by sea-surface temperature, making knowledge of this parameter essential for evaluating past global-mean OH· abundance. Finally, present-day field campaigns combined with modeling studies are required to advance our understanding of HO x cycling during oxidation of isoprene, especially under low-NO x conditions.

Of all proxies for past oxidant abundances, Δ17O of nitrate and CO appear the most promising proxies over longer time scales due to their unique sensitivity to oxidant abundances combined with their high degree of preservation in polar ice. Only two records of ice-core Δ17O(NO3 ) have been published to date [24, 100], and these suggest significant anthropogenic influence on the O3/(HO2· + RO2·) ratio in the mid- to high-latitudes [24]. Measurements of ice-core Δ17O(NO3 ) on the glacial-interglacial time scale are in progress. Δ17O(CO) has not yet been measured in ice cores, but a modeling study suggests that Δ17O(CO) may indeed be a useful indicator of past OH· abundance [115]. As all paleo-climate and paleo-chemistry proxies are inherently uncertain, information from multiple proxies will be required for robust constraints on past variability in the oxidizing capacity of the atmosphere. Such constraints can be used to inform potential future changes in the oxidizing capacity of the atmosphere and the implications for air pollution.



B. Alexander acknowledges the support from NSF-AGS 1103163, NSF-PLR 1106317, NSF-PLR 1244817, and NSF-PLR 1446904. L. J. Mickley acknowledges the support from NSF-AGS 1102880.

Conflict of Interest

Becky Alexander and Loretta J. Mickley declare that they have no conflict of interest.


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Copyright information

© Springer International Publishing AG 2015

Authors and Affiliations

  1. 1.Department of Atmospheric SciencesUniversity of WashingtonSeattleUSA
  2. 2.School of Engineering and Applied SciencesHarvard UniversityCambridgeUSA

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