The preservation of atmospheric nitrate in snow at Summit, Greenland



[1] There is great interest in using nitrate (NO3) isotopic composition in ice cores to track the history of precursor nitrogen oxides (NOx = NO + NO2) in the atmosphere. NO3, however, can be lost from the snow by surface processes, such as photolysis back to NOx upon exposure to sunlight, making it difficult to interpret records of NO3 as a tracer of atmospheric NOx loading. In a campaign consisting of two field seasons (May–June) at Summit, Greenland, high temporal frequency surface snow samples were collected and analyzed for the oxygen isotopic composition of NO3. The strong, linear relationship observed between the oxygen isotopes of NO3, in both 2010 and 2011, is difficult to explain in the presence of significant postdepositional processing of NO3, unless several unrelated variables change in concert. Therefore, the isotopic signature of NO3 in the snow at Summit is most feasibly explained as preserved atmospheric NO3 deposition.

1 Introduction

[2] Ideally, nitrate (NO3), a major ion present in ice cores, could be used to trace the history of its precursor nitrogen oxides (NOx = NO + NO2) sources and impact on the atmosphere. Natural (lightning, microbial process in soils, and stratospheric N2O oxidation) and anthropogenic (fossil fuels combustion and biomass/biofuel burning) sources both make important contributions to the global atmospheric burden of NOx; these emissions have a direct impact on the oxidation capacity of the atmosphere through NOx interactions with ozone (O3) and HOx (hydroxyl and hydroperoxyl radicals). Understanding of NO3 concentration records has long been complicated by the fact that NO3 can be postdepositionally processed in surface snow [Honrath et al., 1999]. Recent studies of the isotopic composition of NO3 in snow have aimed to trace the preservation of NOx source and chemistry signals and/or assess the postdepositional loss of NO3 [Blunier et al., 2005; Frey et al., 2009; Hastings et al., 2009].

[3] At Summit, Greenland (72.6°N, 38.5°W), observations of NOx fluxes from the snow have been ascribed to the photolysis of snow NO3 [Dibb et al., 2002; Honrath et al., 1999] with NOx concentrations of up to 50 pptv measured in the boundary layer [Yang et al., 2002]. In addition, it has been suggested, based on nitrate concentration ([NO3]), that 5–25% of NO3 is lost from the snow [Burkhart et al., 2004; Dibb et al., 2007]. This large amount of loss does not fit with current modeling simulations [Thomas et al., 2011], which suggest that very little NO3 loss is required to explain the locally observed atmospheric concentrations of NOx in the summertime boundary layer.

[4] The nitrogen and oxygen isotopic composition of NO3 has been used to document the postdepositional processing of snow NO3 in polar environments. In low snow accumulation environments, such as Dome C, Antarctica (10 cm of snow per year [Rothlisberger et al., 2000]), the isotopes serve as a tracer of postdepositional loss [Blunier et al., 2005; Frey et al., 2009; Rothlisberger et al., 2002]. In higher accumulation environments, such as Summit (65 cm of snow per year [Dibb and Fahnestock, 2004]), work thus far has indicated that the majority of NO3 and its isotopic composition is preserved in the snow [Hastings et al., 2004, 2009; Jarvis et al., 2009]. It is possible, however, that recycling of NO3 occurs locally, i.e., postdepositional loss of NO3 from the snow as NOx, followed by local oxidation of NOx to NO3 and redeposition. At Summit, accumulation occurs year-round [Dibb and Fahnestock, 2004], with riming and fog in addition to snow deposition possible during the spring and summer.

[5] Here we use the isotopic composition of NO3 to quantify the significance of postdepositional processing of NO3 in snow. This work presents the first complete oxygen isotopic measurements, δ18O and Δ17O, for NO3 in surface snow at Summit with both (δ = (Rsample/RVSMOW − 1) * 1000‰, where R = 18O/16O for δ18O and R = 17O/16O for δ17O, Δ17O = δ17O − 0.52 * δ18O‰) reported. The oxygen isotopic composition of atmospheric NO3 is representative of the oxidation pathways of NO3 formation. Δ17O is the measure of deviation from the typical “mass-dependent” relationship between 18O and 17O, and anomalously high Δ17O values are found in atmospheric NO3 as a result of the interaction of NOx with O3.

[6] Both modeling and observational studies suggest that photolysis is the primary driver of NO3 loss from the snowpack; while other postdepositional processes such as evaporation and volatilization may contribute some to NO3 loss, they are not believed to be as important in locations with low [NO3] and temperatures [Frey et al., 2009; Rothlisberger et al., 2002]. The various photolysis pathways (Figure 1) should induce different effects on the oxygen isotopes of NO3 due to fractionation and exchange of oxygen atoms. Based on theoretical calculations [Frey et al., 2009], photolysis of NO3 to gas-phase NO2 (Figure 1, pathway a) should increase δ18O-NO3 in the surface snow, while Δ17O-NO3 should remain constant. The resulting NO2 may be cycled as NOx in the atmosphere and, ultimately, converted back to HNO3 (Figure 1, pathway b). If HNO3 is formed and redeposited locally, this would imprint a local oxidant composition on the snow NO3 (Figure 1, pathway c). It is also possible that NO3 is photolyzed to a compound (e.g., NO2) that remains in the condensed phase (Figure 1, pathway d) and is capable of exchanging oxygen atoms with an oxidant also in the condensed phase. In this case, the resultant NO3 will have an isotopic composition that is similar to that of water (snow).

Figure 1.

A simplified schematic of reaction pathways for photolysis of NO3 in snow. Pathway a represents photolytic loss of NO3, with pathway b showing NOx cycling and pathway c showing redeposition. Pathway d shows an example mechanism for NO3 exchanging O atoms with H2O. The NO2 here should be considered representative of any NO3 photolysis product trapped in the condensed phase that can be oxidized back to NO3. H2O represents any oxidant in the condensed phase that shares the isotopic composition of H2O.

2 Methods

[7] Two field seasons were conducted: 17 May to 22 June 2010 and 24 May to 26 June 2011. Throughout both, surface snow samples, comprising the top 1–2 cm of snow from 100 to 400 cm2 areas, were collected every 4–12 h and included all major snowfall events. At each sampling time, three replicate samples were collected adjacent to each other within a 10 m by 5 m section of the clean air sector using a Lexan scraper, and each snow sample was stored frozen in a high-density polyethylene bottle until analysis, with all materials precleaned in 18 MΩ cm water. The samples were then analyzed for a suite of ion concentrations at the University of New Hampshire, including NO3, on a Dionex ion chromatograph. (Dibb et al. [2007] provided details on the sampling and analytical protocols, as well as data screening criteria.)

[8] The complete isotopic composition of NO3 was measured at Brown University for each sample. The δ18O-NO3 and Δ17O-NO3 are quantified using the bacterial denitrifier method [Casciotti et al., 2002; Kaiser et al., 2007], where nitrous oxide (N2O) generated from NO3 by denitrifying bacteria is used to determine δ18O, and Δ17O is determined by quantitatively decomposing the N2O into molecular oxygen (O2) and nitrogen (N2). The supporting information includes further discussion of measurement techniques and data correction. In contrast to other studies, the δ18O-NO3 and Δ17O-NO3 data should be considered independent since different aliquots of the samples are run separately for δ18O-NO3 (quantified from N2O) and Δ17O-NO3 (quantified from O2). The most realistic picture of precision is obtained from a pooled standard deviation of sample replicates, which is 0.7‰ (n = 271) for δ18O and 0.9‰ (n = 271) for Δ17O. (See Table S1 in the supporting information for additional error statistics.)

[9] In addition to the NO3 analysis, a selection of samples from each field season was analyzed for δ18O-H2O and δD-H2O. The isotopes of water were measured on a Picarro isotopic water liquid analyzer (L1102-i) with a precision better than 0.2‰.

3 Results

[10] The 2010 and 2011 field seasons show markedly different ranges and average values for NO3 concentration and oxygen isotopes (Figure 2). Comparable high temporal resolution measurements of surface snow are not available for Δ17O-NO3, but measurements from 1 and 2 m snowpits at Summit ranged from ~23‰ to 30‰ for summertime snow [Kunasek et al., 2008], similar to the mean values observed in the May–June surface snow. An average Δ17O-NO3 of 26‰ calculated for atmospheric NO3 in June from a global 3-D atmospheric chemistry model simulation [Alexander et al., 2009] also compares generally well with the surface snow mean values found in 2010 and 2011. Interestingly, there is a great deal of variability in Δ17O-NO3 measured in surface snow, even within a single day, which is not represented by previous snow pit data. Jarvis et al. [2009] reported δ18O of surface snow with a range shifted slightly higher (45‰ to 108‰ versus 30‰ to 95‰) than what was observed in the 2010 and 2011 seasons.

Figure 2.

(left) [NO3], (middle) Δ17O, and (right) δ18O-NO3 comparison for the 2010 (n = 277) and 2011 (n = 345) seasons, which show remarkably different average values and ranges. The black lines indicate median values, with the boxes encompassing the upper and lower quartiles. The individual points are more than 1.5 times the interquartile distance.

[11] Despite the differences in range and average values (Figure 2), similar relationships between [NO3], δ18O-NO3, and Δ17O-NO3 were found in both seasons. In both the 2010 and 2011 field seasons, no correlation was found in the surface snow between [NO3] and the oxygen isotopic composition of NO3 (Figure S1). On the other hand, in both seasons, a very strong linear relationship was found between δ18O-NO3 and Δ17O-NO3 (Figure 3a). As outlined below, this relationship is difficult to explain in the presence of significant postdepositional processing of NO3 given what is known about isotope effects of photolytic processes.

Figure 3.

The 2010 and 2011 surface snow Δ17O-NO3 versus δ18O-NO3. (a) All data points for 2010 (blue, solid circles) and 2011 (red, open circles) are shown. The best fit line for all data is Δ17O = 0.46* δ18O − 6.9 (R2 = 0.9). (b) All data (black solid circles) relative to possible oxidant mixing lines are shown (see section 4.2). Note the change in scale for δ18O-NO3 between the graphs.

4 Discussion

[12] In interpreting the isotope results, the discussion below focuses on photolysis as a primary driver of the potential postdepositional processing of NO3 in snow at Summit. The possible impacts of photolytic loss of NO3, exchange of oxygen atoms within the snow and photolytic driven exchange of oxygen atoms in the snow, are considered.

4.1 Isotopic Impacts of NO3 Processing

[13] As discussed above (section 1), different photolysis pathways will induce different isotopic effects. For NO3 deposited to the snow that is then photolyzed, a theoretical fractionation factor, ϵ18, that assumes Rayleigh fractionation, can be used to quantify the change in δ18O with the degree of photolytic loss of NO3 as follows:

display math(1)

[14] For conditions at Dome C, Frey et al. [2009] calculated ϵ18 as −34‰ serving to increase the residual δ18O-NO3 in the snow (δ18Ofinal) and recalculated for average Summit radiation conditions, ϵ18 = −32‰. As a mass dependent process, the loss of NO3 from the snow should have no impact on the Δ17O-NO3. If we assume that some portion of the data presented in Figure 3a reflects direct deposition, photolysis would serve to move the snow composition away from the observed Δ17O-NO3 versus δ18O-NO3 relationship along lines of constant Δ17O by differing amounts depending on the amount of loss. There is, however, considerable uncertainty associated with the calculated oxygen enrichment factor for photolysis of NO3: the calculations are for the gas phase only (i.e., there is no consideration of matrix effects), the quantum yield has no wavelength dependence, and it is assumed that photolyzed NO3 is lost directly to the gas phase only as NO2.

[15] The enrichment in snow δ18O-NO3 due to photolytic loss fits neither with the data from Summit nor with other measurements of photolytic loss made in the laboratory or field [Frey et al., 2009; McCabe et al., 2005]. In laboratory photolysis experiments and in situ snow measurements, depletion in 18O has been observed and assumed to be the result of competing factors of enrichment due to photolysis and mixing of the residual NO3 with a source depleted in 18O. The most likely source of this low δ18O, due to its abundance, is water or an isotopically similar oxidant. Indeed, laboratory studies of photolysis of nitrate have shown, when beginning with a single NO3 source, that photolysis of USGS35 NaNO3 results in a single line for δ17O-NO3 versus δ18O-NO3 that has markedly different slopes when in waters of differing isotopic composition [McCabe et al., 2005].

[16] If NO3 were to simply exchange oxygen atoms with water, with no loss at all, the expected result would be for the δ18O-NO3 and Δ17O-NO3 to be pulled toward that of water. (For our samples, δ18O-H2O = −38‰ to −20‰ and Δ17O-H2O = 0‰.) This would serve to decrease both δ18O and Δ17O of the NO3 in the snow, though not in the ratio that fits with our observations. For example, starting at the mean values for 2011 (δ18O-NO3 = 70.1‰ and Δ17O-NO3 = 25.3‰), if 10% of the oxygen atoms were to exchange with water of δ18O = −38‰, the new isotopic composition would be δ18O-NO3 = 59.3‰ and Δ17O-NO3 = 22.8‰, using equation (2).

display math(2)

[17] If the water δ18O = −20‰, δ18O-NO3 = 61.1‰, and Δ17O-NO3 would remain at 22.8‰. With increasing exchange, the data point would move toward the isotopic composition of water, i.e., a line set by the water composition end point (or range) and the starting composition of the NO3 (e.g., Figure 3b, green dashed line). The slope of increasing exchange would vary from 0.07 to 0.33 depending upon the initial NO3 composition and the composition of water but would never be equal to the observed value of 0.46.

[18] The competing processes of enrichment due to photolysis and mixing with a depleted oxygen source could result in apparent fractionation along the relationship we observe between δ18O-NO3 and Δ17O-NO3, if they happen in a specific ratio. With a higher degree of loss, a larger amount of mixing with the depleted source would be required to maintain this relationship. This seems plausible, as the proposed mechanisms for mixing with water involve branching photolysis, with some fraction of the NO3 becoming NOx and another portion of NO3 following a path that remains in the condensed phase. The condensed phase NO2/NO2 can exchange oxygen atoms with the solvent water and then reform NO3 (Figure 1, pathway d). As more NO3 is photolyzed to gas-phase NOx, more NO3 may also be photolyzed to the condensed-phase substance (NO2), thus increasing the oxygen exchange with water. As long as these reactions occur in the necessary ratios, the linear relationship between δ18O and Δ17O of the residual nitrate can be maintained. For instance, using the range of NO3 loss from snow concentration studies [Burkhart et al., 2004; Dibb et al., 2007], if a 10% loss of NO3 were occurring, 7% of the remaining NO3 oxygen atoms must exchange with water in order to maintain the observed relationship between δ18O-NO3 and Δ17O-NO3, (i.e., applying equation (1) and then equation (2)). At a 25% loss, a 16% exchange is needed, assuming that the water has a δ18O of −30‰ and Δ17O of 0‰.

[19] If the water had a constant δ18O, competing photolytic enrichment and exchange with water would be a logical explanation for the relationship observed between δ18O-NO3 and Δ17O-NO3. The water observed over the May–June 2010 and May–June 2011 seasons, however, varies in δ18O from −38‰ to −20‰. With a 25% photolytic loss and water with δ18O of −20‰, a 22% exchange of remaining oxygen atoms is required to maintain the relationship between δ18O-NO3 and Δ17O-NO3, while with water δ18O of −38‰, a 15% exchange is required to maintain the relationship. The differences in exchange required with varying water isotopic composition change with differing degrees of NO3 loss. If the isotopic composition of the water were to vary in concert with the photolysis of NO3, we would expect to find a relationship between δ18O-H2O and δ18O-NO3, but there is none.

[20] If the degree of NO3 photolysis and the δ18O-H2O were to vary synchronously, that would require them to both be controlled by the same factors. If the only control on sublimation of water, and therefore δ18O-H2O increase, was actinic flux, then it would be possible to relate it to the degree of photolysis of NO3. The δ18O-H2O should, however, be primarily controlled by relative humidity, which should have no effect on NO3 photolysis. In addition, concurrent changes in NO3 photolysis and δ18O-H2O would require δ18O-H2O to reset to the same values each evening before NO3 photolysis restarts in the morning. This is improbable, as the water deposition can come from a variety of sources with different δ18O-H2O, e.g., riming, fog deposition, or fresh snowfall. In addition, if sublimation were driving the change in δ18O-H2O, there should be a change in deuterium excess in the snow [Stichler et al., 2001], but all the samples fall along a line with a slope of 8 (δ18O = 8.0 * δD + 6.0, R2 = 0.99). This indicates that all isotopic changes in water are derived at equilibrium; therefore, sublimation cannot be the source of variation in δ18O-H2O. The most likely source of δ18O-H2O variation is deposition of new water.

[21] Additionally, stratospheric O3 concentration, and therefore UV penetration to the surface at Summit, is an important control on the photolysis of NO3. It is notable that despite significant depletion in stratospheric O3 during spring 2011 compared to spring 2010 [Manney et al., 2011], the observed relationship between Δ17O and δ18O of NO3 is the same in both years (Figure 3a).

[22] In summary, the observed relationship between δ18O-NO3 and Δ17O-NO3 cannot be explained by postdepositional processing of NO3 in the snow, considering our current understanding of the isotopic imprints of the processes discussed above. The oxygen isotopic signals observed in NO3 at Summit are more plausibly explained as representing atmospheric NO3 deposition to Summit.

4.2 Atmospheric Production of NO3

[23] Most linear relationships of the type found between δ18O-NO3 and Δ17O-NO3 at Summit are interpreted as the result of mixing of different oxidants that react with NOx to produce atmospheric NO3 [e.g., Michalski et al., 2004]. The linear relationship between δ18O-NO3 and Δ17O-NO3 suggests isotopic mixing between a high end-member with δ18O = 100‰ and Δ17O = 39‰ and a low end-member with δ18O = 18‰ and Δ17O = 0‰ (Figure 3a). The high end-member likely results from stratospheric O3. The lower end-member is more difficult to identify. The atmospheric oxidant with the closest isotopic composition is molecular oxygen (O2) (δ18O-O2 = 23.9‰, Δ17O-O2 =0‰ versus Vienna standard mean ocean water [Barkan and Luz, 2005]). A mixing line between these two oxidants (O2 and stratospheric O3) (Figure 3b, red solid line) is the best fit to the surface snow data, compared with other oxidants of Δ17O = 0‰. For instance, H2O vapor (or OH in isotopic equilibrium) would have δ18O between −30‰ and −10‰ (Figure 3b, green dashed line shows −20‰), which does not fit the data. If OH maintains some of its O3 character from O(1D) [e.g., Kunasek et al., 2008], the mixing line would remain the same, but the lower endpoint would be moved toward O3. Assuming an equilibrium fractionation of 44‰ between OH and H2O [Michalski et al., 2012] results in an OH-O3 mixing line that is an even worse fit for the data (Figure 3b, orange dotted line). Thus, it would appear that oxygen atoms from stratospheric O3 and atmospheric O2 are the main controls on the isotopic composition of NO3 that is ultimately deposited to Summit. Furthermore, the influence of the stratosphere on NO3 (e.g., NO3 formed in the troposphere via reaction of NOx and stratospheric O3) may account for the higher than expected summertime Δ17O-NO3 at Summit based on models [Alexander et al., 2009; Kunasek et al., 2008].

4.3 NOx Production From Snow NO3

[24] The conclusion, based upon the isotopic data, that the NO3 seen in Summit snow is a direct atmospheric signal that reflects little to no postdepositional loss contrasts with the estimates of NO3 loss based upon snow concentration measurements made in the past [Burkhart et al., 2004; Dibb et al., 2007]. Gas-phase observations and recent modeling work, however, suggest that very small fractions of NO3 are involved in the postdepositional processing at Summit [Honrath et al., 1999; Thomas et al., 2011].

[25] Honrath et al. [1999] calculated that only a tiny amount of the NO3 must photolyze in order to give a 1000 pptv NOx concentration in the interstitial snow (firn) air (i.e., 6 × 10−11% of the NO3 in a 5 µmol L−1 snow sample must be converted to NOx). These numbers, however, were never translated into boundary layer concentrations.

[26] In a 1-D model that matches well with observed NOx concentrations, Thomas et al. [2011] showed that only 0.10% of the NO3 in the top 10 cm of snow is required to be lost over a 3 day period in order to explain the NOx concentrations measured in the boundary layer at Summit. Assuming a summer accumulation rate of 5.1 cm mo−1 [Dibb and Fahnestock, 2004], the top 10 cm of snow will be entirely replaced by fresh snow in less than 60 days. In that case, a loss of 2.1% of the NO3 in the snow is required to account for the measured NOx concentrations, backing our interpretation that the postdepositional processing of NO3 is small in magnitude and has little to no effect on the isotopes observed.

[27] The contrast between prior snow concentration measurements and the isotopic measurements, as well as the modeled NO3 loss, is difficult to reconcile. We have demonstrated that large photolytic losses of NO3 are not driving these measurements. Evaporative loss or volatilization of HNO3 will also not account for the discrepancy, as HNO3 concentrations in the atmosphere would have to be 4–10 times larger than the NOx concentrations, which is inconsistent with measurements at Summit [Dibb et al., 1998; Honrath et al., 2002]. The lower amount of NO3 loss predicted from NOx concentrations in air fits with the isotope data, while loss predicted from snow concentration measurements does not. It is possible that the calculations based on snow concentrations are confounded by the spatial heterogeneity of NO3 or by fluctuations in water content (e.g., evaporation). The isotopes of NO3, therefore, present a more sensitive record of NO3 chemistry than concentration alone in the snow at Summit.

5 Conclusions and Implications

[28] The isotopic composition of NO3 in the snow at Summit, Greenland, is largely preserved and is representative of atmospheric NO3 deposition to Summit. NO3 at Summit shows a mix of oxidation processes by stratospheric O3 and an unknown oxidant of low Δ17O and δ18O that is isotopically similar to O2, which influence NOx cycling and the formation of NO3. It remains uncertain what accumulation rate is required to preserve the NO3 signal, but at Summit, photolytically driven postdepositional processing is so small in magnitude that it does not have a significant effect on NO3 concentration or isotopes in the snow. Simultaneous observations of gas-phase species and isotopes of NO3 in air and snow may distinguish whether the signal represents regionally formed or long-range transported NO3. In high accumulation areas, such as Summit, isotopic records of ice-core NO3 can be interpreted as a preserved atmospheric signal and used as a tracer of past NOx and atmospheric oxidation conditions.


[29] This work was supported by the National Science Foundation under grant 0909374 (Arctic Natural Sciences). We thank C. Corr and E. Scheuer for their help with field sampling and ion analysis and Polar Field Services for their support with logistics.

[30] The Editor thanks two anonymous reviewers for their assistance in evaluating this paper.