Interactive comment on “ The sensitivity of the oxygen isotopes of ice core sulfate to changing oxidant concentrations since the preindustrial ”

Line 24-28, p20609: Due to the widely varying model approaches and the nonlinearity of oxidant chemistry, a proxy is needed for model validation of PI oxidant concentrations. In this study, we consider the oxygen isotopes of atmospheric sulfate extracted from ice cores as a potential constraint for oxidant concentrations in a global model. The rational for the study of mass-independent isotopes is supposed to be provided by these 2 sentences? Please, can the authors give us some background information


Preindustrial oxidants
The concentrations of the tropospheric oxidants, including ozone (O 3 ), hydroxyl radical (OH), and hydrogen peroxide (H 2 O 2 ), are described collectively as the "oxidizing capacity of the atmosphere" (Thompson, 1992).Variations in the oxidizing capacity impact the lifetimes of chemically and radiatively important reduced trace gases, such as carbon monoxide (CO), methane (CH 4 ), and halocarbons.However, the extent to which the oxidizing capacity of the troposphere has changed between the preindustrial Introduction

Conclusions References
Tables Figures

Back Close
Full Holocene (PI; ∼1850 CE) and present-day (PD) due to anthropogenic activity remains highly uncertain.
Measurements of H 2 O 2 concentrations (Sigg and Neftel, 1991) and CH 2 O/CH 4 (Staffelbach et al., 1991) in ice cores have been studied as potential proxies of the past oxidizing capacity of the atmosphere.However, both are sensitive to post-depositional processing.There are also reconstructions of O 3 measurements from the nineteenth century made with the Sch önbein method (Marenco et al., 1994;Pavelin et al., 1999) and oxidation of arsenate (Volz and Kley, 1988).These reconstructions suggest very low surface O 3 concentrations, on the order of 5-15 ppbv.
Due to the challenges in interpreting these records, the community relies on models to quantify the past oxidizing capacity of the atmosphere.However, between model studies, the fractional change in global mean oxidant concentrations between PI and PD (PD-PI) scenarios varies greatly (e.g., Wang and Jacob, 1998;Thompson et al., 1993;Grenfell et al., 2001;Lamarque et al., 2005).Models predict PD-PI changes of +30 to +65% in mean O 3 , −33% to +10% in OH, and +40% to +140% in H 2 O 2 .Intermodel variability in oxidants is due to differing meteorological fields, differing choices of CH 4 concentrations, uncertainties in PI biogenic and biomass-burning emissions, and inherent variations in PD models related to NO x and volatile organic compound (VOC) emissions and stratosphere-troposphere exchange of O 3 (Wu et al., 2007).In addition to inter-model variability, Mickley et al. (2001) points out that most PI simulations overestimate O 3 relative to the late-1800s measurements described above.To achieve agreement with these measurements, models require dramatically lower emissions of soil and lightning NO x and higher emissions of biogenic VOCs.
Due to the widely varying model approaches and the nonlinearity of oxidant chemistry, a proxy is needed for model validation of PI oxidant concentrations.In this study, we consider the oxygen isotopes of atmospheric sulfate extracted from ice cores as a potential constraint for oxidant concentrations in a global model.Introduction

Conclusions References
Tables Figures

Back Close
Full 2 Oxygen isotopic composition of sulfate In the gas phase, sulfate (SO 2− 4 ) forms through oxidation of sulfur dioxide gas (SO 2 ) by OH.In water, SO 2 dissolves, speciating into SO 2 •H 2 O+HSO − 3 +SO 2− 3 , the total of which is described as S(IV).Aqueous-phase sulfate forms by the oxidation of S(IV) by O 3 , H 2 O 2 , and by molecular oxygen (O 2 ) catalyzed by metals (mainly Fe and Mn).Also, SO 2− 4 forms through heterogeneous-phase oxidation by O 3 on alkaline sea salt and dust aerosols.Additional sulfate-formation pathways are thought to be minor contributors to the global sulfur budget (Faloona, 2009).
The PD model is validated against all available annual, seasonal, and monthly measurements of ∆   water pH=4.5, 5.0, and 5.5.The best-fit pH values from the base simulations are used for all further sensitivity studies.Alexander et al. (2009) showed that metal-catalyzed sulfate formation is dominated by anthropogenic metals in mid-to high-latitudes during winter and is necessary to model the observed seasonal ∆ 17 O SO 4 cycle in Arctic aerosol.Simulations are run both with and without metal-catalyzed oxidation to test the influence of increasing metal emissions between the PI and PD.Modeled metal emissions are scaled to mineral dust and primary anthropogenic sulfate, following Alexander et al. (2009).Thus, only metal in dust is included in the PI.
Two sensitivity studies explore uncertainties in PI emissions of oxidant precursors and their impact on ∆ 17 O SO 4 .The first uses PD biomass burning in a simulation with an otherwise PI configuration.The second sensitivity study is modeled on the work of Mickley et al. (2001).We halve lightning and biomass-burning NO x and double biogenic VOCs from their PI values to test whether modeled preindustrial oxidants can be consistent with both late-1800s O 3 measurements and ice core ∆ 17 O SO 4 .Thompson et al., 1993;Grenfell et al., 2001), however ice core H 2 O 2 measurements are qualitatively consistent with our results (Frey et al., 2006;Sigg and Neftel, 1991).The sensitivity of ∆ 17 O SO 4 to changing oxidants is best assessed using fractional changes in regional oxidants because ∆  Amazon region due to a decrease in the fraction of sulfate formed in the gas-phase.

Results and discussion
In the PD, we assume cloud water pH values of 4.5 in the Northern Hemisphere and 5.0 in the Southern Hemisphere, while in the PI, both hemispheres are assumed to have a cloud-water pH of 5.0, due to the lack of anthropogenic acid emissioins.In both time periods, the model agrees well with the Northern Hemisphere ∆ 17 O SO 4 measurements, while slightly underestimating (by 0.4‰) the WAIS-Divide observations.Simulations at pH=5.5 yield unrealistically high ∆ 17 O SO 4 values.A PD Northern Hemisphere pH of 5.0 does not produce the observed decrease in the ∆ 17 O SO 4 at Summit, Greenland (Table 1).The assumed pH values are consistent with SO 2− 4 and NO − 3 trends in Greenland and Antarctic ice cores.All simulations described below use these pH assumptions.Sulfate deposited at Alert, Canada is likely emitted from Eurasia and the Arctic, whereas sulfate at Site-A, Greenland is influenced by North America and the northern midlatitudes, as it is better exposed to the free troposphere due to its high elevation (Hirdman et al., 2010).Therefore, we consider the change in oxidants over the entire northern midlatitudes ( 30• -60 • N), where we find modeled PD-PI changes of +51%, −7%, and +72%, respectively, for O 3 OH, and H 2 O 2 .There is little variation in these values compared with several other sub-regions (North America, Greenland, >60 • N) of the Northern Hemisphere.At Site-A, we take the PI period to be prior to 1837 CE, as ∆ 17 O SO 4 increases in the Site A record in the late-1800s due to increased North American biomass burning (Alexander et al., 2004), a condition not considered in these simulations.In spite of the 51% increase in O 3 due to anthropogenic activity, there is a PD-PI decrease of 0.5‰ in the measured ∆ At WAIS-Divide, the PI period is taken to be 1850 CE, as the ∆ 17 O SO 4 measurements prior to 1850 are influenced by the 1810 and Tambora (1815 CE) volcanic eruptions and may not represent tropospheric chemistry (Kunasek et al., 2010).We assume that most sulfate deposited at WAIS-Divide, Antarctica originates from oxidation of DMS emitted from the Southern Ocean (Patris et al., 2000).Across the Antarctic region (>60 on the glacial-interglacial timescale in an Antarctic ice core (Alexander et al., 2002).
Modeling the oxygen isotopic composition of both sulfate and nitrate in future work will Introduction

Conclusions References
Tables Figures

Back Close
Full Screen / Esc Printer-friendly Version Interactive Discussion Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Screen / Esc Printer-friendly Version Interactive Discussion Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper |

17 O
SO 4 from aerosol, precipitation, and firn samples (Fig. 1).The model captures the spatial variability in ∆ 17 O SO 4 measurements (e.g.latitudinal gradient).The best agreement is achieved when cloud water pH values of 4.5 and 5.0 are assumed for the Northern and Southern Hemispheres, respectively.This is consistent with the 20611 Discussion Paper | Discussion Paper | Discussion Paper | Due to the uncertainties in PI conditions, we conduct four sensitivity studies that involve varying cloud water pH, metal-catalyzed oxidation, biomass burning emissions, and biogenic VOC and NO x emissions.The choice of global cloud water pH contributes to uncertainty in ∆ 17 O SO 4 , since S(IV)+O 3 is highly pH-dependent.Because the Northern Hemisphere has seen an increase in the acidity of precipitation since the PI(Mayewski et al., 1986) while the Southern Hemisphere has not(Cragin et al., 1987), simulations are run at a bulk cloud 20612 Discussion Paper | Discussion Paper | Discussion Paper |

Figure
Figure 2a-c shows the PD-PI change in annual mean tropospheric oxidant concentrations.Between the PI and PD simulations, global annual mean (3-year average) tropospheric O 3 increases by 42% (32-45 ppbv), OH decreases by 10% (1.3-1.2×10 6molecules cm −3 ), and H 2 O 2 increases by 58% (0.58-0.92 ppbv).Our basecase simulations represent reasonably mainstream changes in O 3 and OH compared to previous modeling studies.Very few studies report PD-PI H 2 O 2 changes (e.g.,Thompson et al., 1993; Grenfell et al., 2001), however ice core H 2 O 2 measurements are qualitatively consistent with our results(Frey et al., 2006; Sigg and Neftel, 1991).The sensitivity of ∆ 17 O SO 4 to changing oxidants is best assessed using fractional 17 O SO 4 depends on the fraction of SO 2Discussion Paper | Discussion Paper | Discussion Paper | by each pathway.
Screen / Esc Printer-friendly Version Interactive Discussion Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | help to further constrain paleo-oxidants, as all non-oxidant factors that impact sulfate or nitrate formation are mutually exclusive.Discussion Paper | Discussion Paper | Discussion Paper |Table 1. Difference in ∆ 17 O SO 4 between PD simulations and the base PI simulation (pH=5.0).Sensitivity studies are also relative to the base PI simulation, and are at a cloud-water pH of 5.0.Italics indicate agreement with measurements within the analytical uncertainty of the measurements.Simulations without metal-catalyzed oxidation by O 2 are indicated by "no met".Simulation (pH)Site-A and alert WAIS-
. The resulting mean ∆ 17 O of atmospheric sulfate (∆ 17 O SO 4 ) depends on the ∆ 17 O transferred to sulfate by each oxidant and the fraction of sulfate formed through each oxidation pathway.The latter depends on oxidant concentrations, cloud liquid water content, Introduction 17O SO 4 of the PI and PD atmosphere.The model is driven by assimilated Goddard Earth Observing System (GEOS) 1989-1991 meteorology.The PD simulation relies on the standard GEOS-Chem emissions inventories -GEIA fossil fuel, fertilizer, biogenic, and biofuel emissions(Wang et al.,     1998)scaled by national energy and CO 2 emission data to 1989-1991, MEGAN 2.0 biogenics increased acidity of Northern Hemisphere precipitation by anthropogenic emissions of nitric and sulfuric acid precursors.When comparing measured and modeled ∆ 17 O SO 4 on the PI to PD timescale, we focus on the PD-PI change of ∆ 17 O SO 4 , rather than the absolute value, to mitigate 3 , and we assume a value at the upper end of the range (35‰).In polar regions, stratospheric O 3 intrusions may increase ∆ 17 O O 3 .Also,Morin et al. (2007)has postulated the non-zero ∆ 17 O of OH formed from O 3 may not be eliminated by isotopic exchange with water vapor in polar regions because of the low water vapor content.These effects may offset modeled ∆ 17 O SO 4 from its true value, but the influence will be similar in both time periods, so the difference in ∆ 17 O SO 4 is largely unaffected.By assuming no change in meteorology, any change in the modeled ∆ 17 O SO 4 is due to a change in oxidant concentrations, pH, or metal emissions.
Table 1 compares the PD-PI change in modeled ∆ 17 O SO 4 to both Arctic (Site-A and Alert) and Antarctic (WAIS-Divide) measurements.Figure 2d illustrates the PD-PI change in the annual mean ∆ 17 O SO 4 of deposited sulfate across the globe.The largest ∆ 17 O SO 4 decrease occurs in the Eurasian Arctic due to increased 17O SO 4 .The model reproduces this decrease in Introduction O SO 4 .Coincident with increased PD O 3 production from anthropogenic precursors, increased anthropogenic metal emissions, primarily from coal-fired power plants, increase the fraction of Site-A sulfate formed by S(IV)+O 2 (∆17O O 2 =0‰) from 7% (PI) to 24% (PD).This increase reduces the fraction of sulfate formed by O 3 and H 2 O 2 since the PI.If metal-catalyzed oxidation is not included, PD ∆17O SO 4 at Alert is overestimated by 0.6‰.The decrease in pH between the PI and PD also decreases the fraction of SO 2− 4 formed by O 3 in the aqueous phase.Both increases in anthropogenic metals and a decrease in pH are needed to explain the observations.
Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | the ∆ 17 • S), relative modeled PD-PI changes in O 3 , OH, and H 2 O 2 are +22%, −16%, and +52%, respectively.Despite these oxidant changes, only a very slight change (−0.2‰) in the ∆ 17 O SO 4 at WAIS-Divide is modeled, consistent with observations.The greater increase in H 2 O 2 than O 3 suppresses S(IV)+O 3 .Thus, slightly increased oxidation by H 2 O 2 , and decreased oxidation by OH and O 3 results in little net change in ∆ 17 O SO 4 due to offsetting effects of ∆ 17 O OH and ∆ 17 O O 3 .That is, ∆ 17 O SO 4 at WAIS-Divide is sensitive to changing oxidant concentrations, but the oxidants change such that there is little net effect on ∆ 17 O SO 4 .The ∆ 17 O SO 4 is used to assess the two sensitivity studies of PI emissions.Increasing biomass burning emissions to PD levels causes changes relative to the PI base simulation of +9%, −4% and +18% in global O 3 , OH, and H 2 O 2 , respectively.However, ∆ 17 O SO 4 is not impacted at either WAIS-Divide or Site-A, because the oxidant changes associated with biomass burning, particularly OH and O 3 , are restricted to low latitudes.Following Mickley et al. (2001), we reduce NO x emissions and double biogenic VOC emissions to try to reproduce late-1800s O 3 measurements.This changes global O 3 , Introduction OH, and H 2 O 2 by −14%, −42%, and +62%, respectively.Surface O 3 concentrations at the sites of late-1800s O 3 measurements are reduced from the base PI simulation by ∼5 ppbv, but are still ∼5 ppbv higher than both those reported by Mickley et al. (2001) and the measurements.These changes do impact the polar regions, resulting in reductions in the modeled ∆ 17 O SO 4 by >0.2‰ in both regions, relative to the base PI simulation, due to increases in H 2 O 2 .The modeled ∆ 17 O SO 4 disagrees with WAIS-Divide measurements and with the PI mean at Site-A (Table 1), although it does fall within the range of Site-A PI measurements.These results suggest that a low bias in the late-1800s O 3 reconstructions may be responsible for the discrepancy with PI O 3 modeling results.
made insensitive to PD-PI changes in oxidants by the increased importance of oxidation by O 2 catalyzed by anthropogenic metals in the PD and the decreased pH.Finally, modeled PI oxidant concentrations cannot be consistent with both ice core ∆ 17 O SO 4 and late-1800s O 3 measurements, suggesting that these O 3 measurements are biased low.This method for modeling ∆ 17 O SO 4 is now being applied to glacial-interglacial timescales.In contrast to the PD-PI transition, ∆ 17 O SO 4 varies dramatically (by 3.5‰)