# Oxygen isotope mass balance of atmospheric nitrate at Dome C, East Antarctica, during the OPALE campaign

Joël E Savarino, William Vicars, Matthieu Legrand, Susanne Preunkert, Bruno Jourdain, Markus Michael Frey, Alexandre Kukui, Nicolas Caillon, Jaime Gil Roca

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Joël E Savarino, William Vicars, Matthieu Legrand, Susanne Preunkert, Bruno Jourdain, et al. Oxygen isotope mass balance of atmospheric nitrate at Dome C, East Antarctica, during the OPALE campaign. Atmospheric Chemistry and Physics Discussions, 2015, 15, pp.24041-24083. 10.5194/acpd-15-24041-2015. insu-03577055

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# Oxygen isotope mass balance of atmospheric nitrate at Dome C, East Antarctica, during the OPALE campaign

J. Savarino<sup>1,2</sup>, W. C. Vicars<sup>1,2,a</sup>, M. Legrand<sup>1,2</sup>, S. Preunkert<sup>1,2</sup>, B. Jourdain<sup>1,2</sup>, M. M. Frey<sup>3</sup>, A. Kukui<sup>4,5</sup>, N. Caillon<sup>1,2</sup>, and J. Gil Roca<sup>4,5</sup>

<sup>1</sup>Université Grenoble Alpes, Laboratoire de Glaciologie et Géophysique de l'Environnement (LGGE), 38000 Grenoble, France

<sup>2</sup>CNRS, Laboratoire de Glaciologie et Géophysique de l'Environnement (LGGE), 38000 Grenoble, France

<sup>3</sup>British Antarctic Survey, Natural Environment Research Council, Cambridge, UK

<sup>4</sup>Laboratoire Atmosphère, Milieux et Observations Spatiales (LATMOS), UMR8190, CNRS-Université de Versailles Saint Quentin, Université Pierre et Marie Curie, Paris, France

<sup>5</sup>Laboratoire de Physique et Chimie de l'Environnement et de l'Espace (LPC2E), UMR6115 CNRS-Université d'Orléans, 45071 Orléans CEDEX 2, France

<sup>a</sup>now at: Technical Services Program, Air Pollution Control Division, Colorado Department of Public Health and Environment, Denver, CO, USA

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Abstract

Introduction

Conclusions

References

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Abstract

Variations in the stable oxygen isotope composition of atmospheric nitrate act as novel tools for studying oxidative processes taking place in the troposphere. They provide both qualitative and quantitative constraints on the pathways determining the fate of atmospheric nitrogen oxides (NO + NO₂ = NOₓ). The unique and distinctive ¹⁷O-excess (Δ¹⁷O = δ¹⁷O - 0.52 × δ¹⁸O) of ozone, which is transferred to NOₓ via oxidation, is a particularly useful isotopic fingerprint in studies of NOₓ transformations. Constraining the propagation of ¹⁷O-excess within the NOₓ cycle is critical in polar areas where there exists the possibility of extending atmospheric investigations to the glacial/interglacial time scale using deep ice core records of nitrate. Here we present measurements of the comprehensive isotopic composition of atmospheric nitrate collected at Dome C (East Antarctic plateau) during the austral summer of 2011/12. Nitrate isotope analysis has been here combined for the first time with key precursors involved in nitrate production (NOₓ, O₃, OH, HO₂, RO₂, etc.) and direct observations of the transferrable Δ¹⁷O of surface ozone, which was measured at Dome C throughout 2012 using our recently developed analytical approach. Assuming that nitrate is mainly produced in Antarctica in summer through the OH + NO₂ pathway and using concurrent measurements of OH and NO₂, we calculated a Δ¹⁷O signature for nitrate in the order of (21–22 ± 3)‰. These values are lower than the measured values that ranged between 27 and 31‰. This discrepancy between expected and observed Δ¹⁷O(NO₃⁻) values suggests the existence of an unknown process that contributes significantly to the atmospheric nitrate budget over this east Antarctic region.

1 Introduction

The search for ice core proxies related to past change of oxidative properties of the atmosphere is motivated by the need to model simulations of ozone and OH changes over preindustrial–industrial and glacial–interglacial timescales. Early ice

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core reconstructions of oxidants based on H₂O₂ (Sigg and Neftel, 1991) and HCHO (Staffelbach et al., 1991) measurements were hampered by the occurrence of post-depositional alteration of H₂O₂ and HCHO concentrations in the upper snowpack prior to preservation in the ice (Hutterli et al., 2003). More recently, the ¹⁷O-excess (Δ¹⁷O = δ¹⁷O - 0.52 × δ¹⁸O) of nitrate, a unique isotopic signature inherited from ozone via bimolecular chemical reactions in the atmosphere, has shown promise as a conserved proxy for past oxidant concentrations (McCabe et al., 2005; Alexander et al., 2004). The Δ¹⁷O signal of nitrate reflects the relative importance of NOₓ transformation mechanisms and recent studies suggest that the measurement of this isotopic signal in ice cores and ancient sediments may provide relevant information regarding the role of ozone in the overall oxidative capacity of the paleo-atmosphere (McCabe et al., 2007; Michalski et al., 2004, 2003; Savarino et al., 2007; Thiemens, 2006). Parallel studies of the nitrogen isotope ratios (δ¹⁵N) of nitrate in polar ice and snow suggest that this isotopic tracer may serve as a proxy for past variations in natural sources of atmospheric NOₓ (Hastings et al., 2009, 2005; Jarvis et al., 2008). However, the chemical and physical factors governing the oxygen and nitrogen isotopic composition of atmospheric nitrate and its nitrogen oxide precursors are not fully understood (Alexander et al., 2009; Morin et al., 2009). This is particularly true in the polar troposphere, where the UV-photolysis of trace species present in the snowpack or marine aerosols initiates complex boundary layer oxidation processes involving reactive halogen species (Bloss et al., 2010; Grannas et al., 2007; Wang et al., 2007) and results in distinctive Δ¹⁷O and δ¹⁵N signatures in atmospheric nitrate (Morin et al., 2012, 2007, 2008).

Here we present measurements of δ¹⁵N and δ¹⁷O and δ¹⁸O isotopic composition of atmospheric nitrate collected at Dome C between October 2011 and January 2012. These measurements were conducted within the framework of the OPALE project (Oxidant Production over Antarctic Land and its Export, Preunkert et al., 2012), which has provided an opportunity to combine nitrate isotopic observations with a wealth of meteorological and chemical observations, including measurements of species involved in nitrate production (NOₓ, O₃, OH, HO₂, RO₂, etc.). The primary objective of this study

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was to reconcile observations of $\Delta^{17}$O for atmospheric nitrate at Dome C with quantitative predictions based on nitrate isotope mass balance and atmospheric chemistry parameters, a unique opportunity offered by the OPALE campaign.

# 2 Methods

## 2.1 Site description and scientific context

Dome C is situated 3233 m a.s.l. on the East Antarctic Plateau (75°06′ S, 123°23′ E), approximately 1100 km from the coastal research station Dumont d'Urville and 560 km from the Vostok station. Deep ice cores were extracted at Dome C in the framework of the European Project for Ice Coring in Antarctica (EPICA) covering approximately 800 000 yr (EPICA-community-members, 2004) and Vostok covering the last 420 000 years (Petit et al., 1999). In parallel, studies aiming to understand the meteorological, chemical, and physical factors governing the variability of trace constituents preserved in the ice were initiated. This was done for aerosol in the framework of the French environmental observation service CESOA (Etude du cycle atmosphérique du Soufre en relation avec le climat aux moyennes et hautes latitudes Sud, http://cesoa.ore.fr) dedicated to the study of the sulfur cycle (Jourdain et al., 2008; Preunkert et al., 2008).

Although the Antarctic plateau is extraordinarily dry, cold, and far removed from sources of anthropogenic emissions, first atmospheric measurements of oxidants conducted in 1998–1999 during the ISCAT (Investigation of Sulfur Chemistry in the Antarctic Troposphere) field campaign revealed a high level of photochemical activity. For example, the average summertime OH concentration ($2 \times 10^6$ cm$^{-3}$) over the South Pole was found to be similar to that of the tropical marine boundary layer (MBL) (Mauldin et al., 2001). Unexpectedly high levels of nitric oxide (NO) were also detected, with concentrations one to two orders of magnitude higher than that typically observed in other remote regions (Davis et al., 2001). Model simulations revealed that the large OH

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concentrations observed at South Pole were a result of the elevated NO level, which catalyzes a rapid cycling of HO₂ to OH (Chen et al., 2004, 2001). The high concentrations of NOₓ were also inferred to drive in situ photochemical production of ozone during the ISCAT campaign (Crawford et al., 2001). Surface ozone and NOₓ measurements at Dome C suggest a similar level of enhanced oxidant production during November–January (Frey et al., 2013, 2014; Legrand et al., 2009).

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The high levels of photochemical activity observed at South Pole and Dome C are now understood in terms of NOₓ release from the snowpack (Honrath et al., 2000, 1999; Jones et al., 2001, 2000; Zhou et al., 2001). This process is initiated by the photolysis of nitrate, which can lead to large fluxes of NO₂, NO, and HONO from permanently sunlit snow (Anastasio and Chu, 2009; Grannas et al., 2007; Jacobi and Hilker, 2007; Legrand et al., 2014; Frey et al., 2013). Observed and modeled NOₓ production rates are largely capable of explaining the high levels of photochemical activity observed on the Antarctic plateau during spring (France et al., 2011; Liao and Tan, 2008; Wang et al., 2007) although detailed and speciation of nitrogen oxides chemistry remain largely unknown in this rich NOₓ/poor VOCs environment (Kukui et al., 2014; Frey et al., 2014; Legrand et al., 2014; Davis et al., 2008).

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## 2.2 High-volume sampling

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Bulk atmospheric samples were collected at Dome C on glass fiber filters using a high-volume air sampler (HVAS), which was installed on a platform 1 m above the ground. The HVAS was run by applying an average STP flow rate of 0.7 m³ min⁻¹ that ensures the collection of a sufficient amount of nitrate for isotopic analysis. The atmospheric NO₃⁻ collected on glass fiber filters represents the sum of atmospheric particulate NO₃⁻ (*pNO₃⁻*) and gaseous nitric acid (HNO₃) (Frey et al., 2009; Savarino, unpublished data). 11 HVAS samples were obtained during the OPALE campaign (from October 2011 to January 2012). After each collection period, filters were removed from the HVAS and placed in clean 50 mL centrifuge tubes, which were sealed in plastic bags and stored at -20°C. Upon arrival at our laboratory in Grenoble, atmospheric filter samples were

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extracted in 40 mL of ultra-pure water via centrifugation using Millipore Centricon™ filter units. Nitrate concentration was then determined for each filter extract solution using a colorimetric technique (Frey et al., 2009).

## 2.3 Ozone collection

The nitrite-coated filter technique for ozone isotope analysis has been described in detail in (Vicars et al., 2012; Vicars and Savarino, 2014). The principle of ozone collection underlying this technique is the filter-based chemical trapping of ozone via aqueous phase reaction with nitrite (Adon et al., 2010; Geyh et al., 1997; Koutrakis et al., 1993; Krzyzanowski, 2004):

$$NO_2^- + O_3 \rightarrow NO_3^- + O_2 \quad (R1)$$

By coupling this routine ozone measurement technique with recent analytical developments in the $\Delta^{17}O$ analysis of nanomole quantities of nitrate (Kaiser et al., 2007), the $^{17}O$-excess transferred by one of the two $O_3$ terminal atoms through bimolecular chemical reactions, denoted $\Delta^{17}O(O_3)_{term}$, as well as ozone's bulk $\Delta^{17}O$ value, denoted $\Delta^{17}O(O_3)_{bulk}$ ($= 2/3 \times \Delta^{17}O(O_3)_{term}$ since all $^{17}O$-excess is located at the two $O_3$ terminal atoms (Bhattacharya et al., 2008; Janssen and Tuzson, 2006), can be inferred from the oxygen isotopic composition of the nitrate produced via R1 within the coated filter matrix.

Ambient ozone collections were performed using an active air sampler consisting of 1/4" Teflon™ (PFA) tubing connecting three main sampler components: (i) a standard low-volume vacuum pump (Welch™, Model 2522C-02) equipped with a volume counter and needle valve (or flow meter) for flow rate regulation, (ii) a closed PFA filter holder assembly (Savillex™) containing a pre-coated 47 mm glass fiber sampling substrate (Whatman™, GF/A type), and (iii) an open-faced PFA filter holder assembly containing a 47 mm PTFE membrane filter (Zylon™, 5 µm pore size) for the removal of particulate species upstream of the coated filter. Glass fiber sampling substrates were
coated prior to sample collection with 1 mL of a nitrite-based ozone sampling solution (0.1 M NaNO₂, 0.05 M K₂CO₃, 0.1 M glycerol) (Koutrakis et al., 1993), allowed to dry at 75°C for approximately 10 min, and then stored frozen in the dark in individual plastic PetriSlide™ containers (Millipore™). Samples were collected by loading pre-coated filters into the sampling filter holder, which was then connected to the prefilter (upstream) and needle valve/pump (downstream) and covered in aluminum foil to limit light exposure, a step that is necessary to limit the blank production rate during sample collection (Vicars et al., 2012). Air was then pumped through the sampling system at a target flow rate of 3.0 L min⁻¹.

Sampling was conducted during the OPALE campaign (December 2011–January 2012) from a climate-controlled shelter, and a total of 28 samples were collected. However, due to difficulties in the application of our analytical technique to the unique environmental conditions encountered in Antarctica, the results obtained from these samples were inconclusive (i.e. unrealistic variability) due to the lack of light protection of the filter holder (Vicars et al., 2012, 2013). Sampling and isotopic analysis of ozone was therefore repeated in 2012, and a complete annual record of Δ¹⁷O(O₃) was obtained (n = 60). Procedural filter blanks were also collected at regular intervals and were subjected to all of the same preparation, storage, handling, and analytical procedures as field samples. After sample collection, filter samples and procedural blanks were returned to their containers, which were covered in aluminum foil and stored at –20°C before processing and analysis.

Filter samples were extracted in 18 mL of deionized water (18.2 MΩ, hereafter referred to as “MQ water”). In order to remove the excess (i.e., unreacted) nitrite reagent from the sample extracts, the solutions were treated with 1 mL of a 1 M sulfamic acid solution and then neutralized with a corresponding addition of high-purity sodium hydroxide (Granger and Sigman, 2009; Vicars et al., 2012). Extract solutions were then filtered via centrifugation using Millipore Centricon™ assemblies. The nitrate extracted from the coated filter samples was then subjected to isotopic analysis, as described in the following section.

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The comprehensive isotopic composition of nitrate ($^{15}$N/$^{14}$N, $^{17}$O/$^{16}$O, $^{18}$O/$^{16}$O) was measured on a Finnigan™ MAT253 isotope ratio mass spectrometer (IRMS), equipped with a GasBench II™ and coupled to an in-house built nitrate interface (Morin et al., 2009). Nitrate in both the nitrite-coated filter and aerosol sample extracts was prepared for isotopic analysis by conversion to N₂O via the bacterial denitrifier method (Casciotti et al., 2002; Kaiser et al., 2007; Michalski et al., 2002; Sigman et al., 2001). The detailed analytical procedure has been described elsewhere (see Morin et al., 2009) and is here briefly presented.

Denitrifying bacteria (*Pseudomonas aureofaciens*) were cultured in nitrate-amended soy broth and incubated for 5 days in stoppered glass bottles. Bacterial cultures, after concentration by centrifugation and re-suspension, were dispensed as 2 mL aliquots into 20 mL glass vials, which were then crimped and purged with helium for 3 h. Approximately 100 nmol of sample nitrate was then injected into the purged vials and conversion of the sample nitrate to nitrous oxide (N₂O) via bacterial denitrification was allowed to proceed overnight. The N₂O sample was then cryo-focused in a liquid nitrogen trap and introduced into a gold furnace where it was thermally decomposed at 900 °C into O₂ and N₂. Following separation via gas chromatography, the O₂ and N₂ sample gases were directed into the ionization chamber of the IRMS. All analytical steps were simultaneously performed on nitrate isotopic standards and their equimolar mixtures (International Atomic Energy Agency USGS 32, USGS 34, and USGS 35), which were prepared in an identical background matrix as the samples. Individual analyses were normalized through comparison with these three nitrate reference materials (Coplen, 2011; Werner and Brand, 2001). All isotopic enrichment values for nitrate are reported relative to VSMOW and air N₂ for oxygen and nitrogen, respectively. The overall accuracy of the method is estimated as the standard deviation of the residuals from the linear regression between the measured reference materials and

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their expected values. For the results reported here, the average uncertainty obtained for δ¹⁸O, Δ¹⁷O, and δ¹⁵N were 1.6, 0.5, and 1.0‰, respectively.

## 2.5 Complementary measurements

Concurrent chemical measurements conducted at Dome C during the campaign include HONO (Legrand et al., 2014), HOₓ (=OH + HO₂ + RO₂) radicals (Kukui et al., 2014), O₃ (Legrand et al., in preparation) and NO and NO₂ (Frey et al., 2014). Photolysis rate coefficients and meteorological parameters were also documented.

# 3 Results and discussion

## 3.1 Isotope ratios of ozone and atmospheric nitrate

Atmospheric nitrate concentrations observed at Dome C during the campaign are presented in Fig. 1, the corresponding nitrate Δ¹⁷O and δ¹⁵N values in Fig. 2. Atmospheric nitrate concentrations ranged between 20 and 90 ng m⁻³, with the maximum values occurring in mid December 2011, concurrent with the period of intensive of atmospheric sampling of the OPALE field campaign. These values are in good agreement with those observed during the 2007–2008 and 2009–2010 field studies conducted at Dome C by Frey et al. (2009) and Erbland et al. (2013), respectively.

Δ¹⁷O values for atmospheric nitrate ranged between 27.3 and 32.4‰, δ¹⁵N between -42.8 and 1.7‰. The observed strongly depleted δ¹⁵N(NO₃⁻) values are in good agreement with those previously reported and having unambiguously attributed to the transformation of local snowpack NOₓ emissions via photochemistry in the boundary layer, which led to peaks in atmospheric nitrate concentration during the period from October to December (Erbland et al., 2013). As seen in Fig. 2, variations in Δ¹⁷O and δ¹⁵N were negatively correlated (r value of -0.86) and again show similar amplitude and phase than those reported in previous studies (Erbland et al., 2013; Frey et al., 2009).

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A time series showing the year-round record of $\Delta^{17}\text{O}(\text{O}_3)_{\text{bulk}}$ at Dome C in 2012 is presented in Fig. 3. $\Delta^{17}\text{O}(\text{O}_3)_{\text{bulk}}$ averaged $24.9 \pm 1.9\%$ over 2012, corresponding to $\Delta^{17}\text{O}(\text{O}_3)_{\text{term}}$ values of $37.4 \pm 1.9\%$. As shown in Fig. 4, these $\Delta^{17}\text{O}(\text{O}_3)_{\text{bulk}}$ values are consistent with those observed in Grenoble (France), as well as with measurements conducted along a latitudinal transect from 50° S to 50° N in the Atlantic Ocean (Vicars and Savarino, 2014). Although the $\Delta^{17}\text{O}(\text{O}_3)_{\text{bulk}}$ seasonal cycle reveals some interesting features, like the winter maximum, probably in response of the permanent winter darkness and stratospheric air mass intrusions, a complete description is beyond the scope of the present paper. What should be kept in mind here is the pretty stable $\Delta^{17}\text{O}(\text{O}_3)_{\text{bulk}}$ value close to 26‰ that can be considered as representative for the OPALE campaign hold in November–January.

## 3.2 Nitrate isotope mass balance

The availability of a large database of trace chemical species measurements at Dome C during a portion of the OPALE field campaign (December 2011) offers a unique opportunity to compare observed $\Delta^{17}\text{O}(\text{NO}_3^-)$ values in the atmosphere to estimated ones calculated from concurrent observations. As discussed at length in recent studies (for example, by Morin et al., 2011 and Vicars et al., 2013), the $^{17}\text{O}$-excess transfer functions associated with the various nitrate production pathways (i.e., $\Delta^{17}\text{O}(\text{NO}_3^-)_i$ values) can be estimated as a function of the $\Delta^{17}\text{O}$ of nitrate precursor gases (i.e., $\text{NO}_x$, $\text{O}_3$, OH, etc.) using mass balance calculations that trace the origin of oxygen atoms transferred during the chemical transformation of $\text{NO}_x$ in the atmosphere. All atmospheric nitrate production channels involve either $\text{NO}_2$ or a $\text{NO}_x$ reservoir species derived from $\text{NO}_2$ (e.g., $\text{N}_2\text{O}_5$). The first step in determining the $\Delta^{17}\text{O}$ signature of each pathway is therefore a quantitative assessment of the steady state $\Delta^{17}\text{O}$ value of $\text{NO}_2$, which is typically calculated as a function of the $\Delta^{17}\text{O}$ value of $\text{O}_3$ and the reaction dynamics involved in the conversion of NO to $\text{NO}_2$. As Dome C in summer is

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permanently under sunlight, photochemical inter-conversion of NOₓ is permanent:

$$NO_2 + hv \rightarrow NO + O \quad (R2)$$

$$NO + O_3 \rightarrow NO_2 \quad (R3)$$

$$NO + HO_2/RO_2 \rightarrow NO_2 \quad (R4)$$

At photochemical steady state (i.e. Reactions R2–R4 being faster than NO₂ net sink reactions), an assumption that can be reasonably applied throughout the day at Dome C during summer, we have (Morin et al., 2011):

$$\Delta^{17}O(NO_2) = \alpha \times (1.18 \times \Delta^{17}(O_3)_{\text{bulk}} + 6.6) \quad (1)$$

where the term in bracket represents the laboratory deduced anomaly transfer function of the NO + O₃ reaction (Savarino et al., 2008), Δ¹⁷O(O₃)ₗₖₗₘ the ¹⁷O-excess of the bulk O₃ and α represents the fraction of the atmospheric NO₂ reservoir that has been produced through oxidation by O₃ rather than HO₂/RO₂ at photochemical equilibrium (Alexander et al., 2009; Michalski et al., 2003; Morin et al., 2011; Röckmann et al., 2001):

$$\alpha = \frac{k_{NO+O_3}[NO][O_3]}{k_{NO+O_3}[NO][O_3] + k_{NO+HO_2}[NO][HO_2]^*} \quad (2)$$

with [HO₂]⁺ = [HO₂] + [RO₂].

It is important to note here that Eqs. (1) and (2) although established under the NOₓ steady-state approximation, are independent of NO₂ concentration for which a bias in measurement cannot be ruled out. Indeed, as discussed by Frey et al. (2013, 2014), bias in NO₂ measurements is suspected partly because it remains difficult to explain the observed ratio of NO₂/NO which is systematically higher (up to a factor of 7) than predicted by calculations made by assuming photochemical steady state considering the NO₂ photolysis and reaction of NO with O₃, HO₂/RO₂ and BrO. Equation (2) also assumes that [HO₂]⁺ is predominantly formed by the reaction H + O₂

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and R + O₂ during the OPALE campaign (Kukui et al., 2014), resulting in the formation of [HO₂]⁺ devoid of any significant ¹⁷O-excess (Morin et al., 2011). Using OPALE measurements of NO, O₃, OH and HO₂/RO₂ (Frey et al., 2014; Kukui et al., 2014), along with temperature dependent reaction kinetics data obtained from (Atkinson et al., 2004), we have calculated the diurnally mass averaged trend in α for the month of December 2011 at Dome C. Measurements of Δ¹⁷O(O₃)bulk at Dome C during the OPALE campaign averaged 25±2‰, corresponding to Δ¹⁷O(O₃)term values of 37±2‰ (Fig. 4). Samples collected in December indicate Δ¹⁷O(O₃)bulk values close to 26‰ (Δ¹⁷O(O₃)term = 3/2 × Δ¹⁷O(O₃)bulk = 39–40‰, Fig. 3), and we have therefore adopted a Δ¹⁷O(O₃)term value of 40‰ in the subsequent mass balance calculations, in good agreement with the predicted value from a 1-D atmospheric model (Zahn et al., 2006). The diurnally mass average trend in Δ¹⁷O(NO₂) calculated using a Δ¹⁷O(O₃)bulk value of 26‰ and Eq. (2) is shown in Fig. 5. No trend is observed during the OPALE campaign and on average the predicted value is Δ¹⁷O(NO₂) = 31±2‰ throughout December, corresponding to average α value of 0.83. In other words, at steady state, the concentrations of O₃ and HO₂⁺ measured during OPALE predicts that around 83% of NO₂ is formed via Reaction (R3) (see also Table 1). In the absence of the α dilution effect introduced by the HO₂⁺ reaction, Δ¹⁷O(NO₂) would equal 37‰, a value 8‰ lower than an estimation obtained from modeling only NOₓ-O₃ chemistry at standard temperature and pressure (Michalski et al., 2014). This difference is essentially explained by the use of different Δ¹⁷O(O₃)bulk (32‰ Michalski's simulation, 26‰ for our observations), which possibly correspond to different conditions of the two studies.

By accounting for the origin of the oxygen atom transferred during the conversion of NO₂ to nitrate, the Δ¹⁷O signature of the nitrate produced through different reaction mechanisms can be calculated. For summer conditions at Dome C, it is reasonable to assume that the dominant atmospheric nitrate formation pathway is the gas-phase association of NO₂ and the OH radical (Alexander et al., 2009):

$$NO_2 + OH + M \rightarrow HNO_3 + M \quad (R5)$$

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leading to the following ¹⁷O-excess mass balance (Michalski et al., 2003; Morin et al., 2011):

$$ \Delta^{17}\text{O}(\text{NO}_3^-) = \frac{2}{3}\Delta^{17}\text{O}(\text{NO}_2) + \frac{1}{3}\Delta^{17}\text{O}(\text{OH}) \quad (3) $$

In order to predict the Δ¹⁷O value of the nitrate produced through Reaction (R5) by mass balance, the isotopic composition of tropospheric OH must be known. The OH radical participates in a rapid isotopic exchange with atmospheric water vapor, which represents a very large oxygen reservoir relative to OH, with a Δ¹⁷O that is negligible compared to ozone or nitrate (Luz and Barkan, 2010). This exchange tends to erase the ¹⁷O-excess of OH under humidity and temperature conditions typical of the mid-latitudes (Dubey et al., 1997); therefore, the Δ¹⁷O of OH is normally assumed to be zero in modeling studies applied to these regions. As discussed by Morin et al. (2007), this assumption of Δ¹⁷O(OH) = 0 is not valid under the low humidity conditions encountered in the polar atmosphere. The degree of isotopic equilibration between OH and H₂O can be determined as a function of the relative rates of the isotope exchange reaction and the main OH sink reactions:

$$ \beta = \frac{L}{L + k_{\text{H}_2\text{O}+\text{OH}}[\text{H}_2\text{O}]} \quad (4) $$

where L represents the total chemical loss rate of OH. β is the factor relating the initial Δ¹⁷O transferred to OH upon its formation, denoted Δ¹⁷O(OH)prod., to its steady state Δ¹⁷O value (Morin et al., 2007):

$$ \Delta^{17}\text{O}(\text{OH}) = \beta \times \Delta^{17}\text{O}(\text{OH})_{\text{prod.}} \quad (5) $$

In plain words, Eqs. (4) and (5) predict that when the isotopic exchange reaction dominates over OH chemical losses (i.e., β << 1), the steady state Δ¹⁷O value of OH will be equal to that of water (i.e., Δ¹⁷O ≈ 0‰). Conversely, when water vapor concentrations are low and the rate of chemical loss is large relative to the rate of the isotopic
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exchange, then $\Delta^{17}\text{O(OH)} = \Delta^{17}\text{O(OH)}_{\text{prod.}}$. Kukui et al. (2014), using a Master Chemical Mechanism box model, constrained by the OPALE meteorological conditions and concurrent chemical observations, give the rate of the OH chemical sources and sinks. $\text{NO}_2$ as measured by Frey et al. (2014) represents at most only ca. 10% (equivalent of ca. 1‰) of the total sink of OH, which is predominantly dominated by reactions with $\text{CO}$, $\text{CH}_4$, aldehydes and to a lesser extent by reactions with $\text{O}_3$, $\text{H}_2$, and NO. Thus, the possible overestimation of $\text{NO}_2$ concentration has only a minor effect on $\beta$ calculation and is well embedded within the total uncertainty of such calculation. To assess the value of $\Delta^{17}\text{O(OH)}$, we have computed $\beta$ for the conditions found during the OPALE campaign using the same 0-D box model that is used to evaluate the budget of OH and $\text{RO}_2$ during the OPALE campaign (see Kukui et al., 2014 and Table 1) and used the exchange kinetic rates given in Dubey et al. (1997). The absolute water vapor concentration is deduced from relative humidity and temperature measurements using Bolton, (1980) (i.e. $P_{\text{water}} = 6.112 \times e^{\frac{(17.67 \times (T - 273))}{T - 29.5}}$, with $P_{\text{water}}$ in hPa and $T$ in K). The results of this calculation (Fig. 6), indicate that $\beta$ varies between 0.70±0.10 (1 $\sigma$) and 0.30±0.10 from midnight to noon for conditions prevailing during the OPALE campaign, suggesting that on daily average basis approximately 43% of the $\Delta^{17}\text{O}$ value originally present in OH is preserved from exchange with $\text{H}_2\text{O}$, consistent with estimates for an Arctic site described by Morin et al. (2007).

The value of $\Delta^{17}\text{O(OH)}_{\text{prod.}}$ is more difficult to assess because of the interplay of $\text{HO}_x$ family and the different sources involved in OH formation. In the rich $\text{NO}_x$ atmosphere at Dome C in summer, the $\text{O}^{1}\text{D} + \text{H}_2\text{O}$ reaction forming OH is a minor reaction pathway. When multiple pathways are involved in the production of OH, $\Delta^{17}\text{O(OH)}_{\text{prod.}}$ can be estimated by a simple isotope mass balance equation where $\Delta^{17}\text{O(OH)}_{\text{prod.}} = \sum_i P_i \times \Delta^{17}\text{O}_i$ with $P_i$ the relative production rate of the $i$th reaction pathway with respect to the total production rate and $\Delta^{17}\text{O}_i$ its associated $^{17}\text{O}$-excess (Morin et al., 2011). Observations at Dome C during the OPALE campaign indicate that the photolysis of HONO and the $\text{HO}_2 + \text{NO}$ reaction may represent the most sig-
nificant sources of OH at Dome C during the period of seasonal snowpack emissions (Kukui et al., 2014). However, the measurement of HONO (around pmol mol⁻¹) during OPALE, probably biased by HO₂NO₂ interference (Legrand et al., 2014) is incompatible with the HOₓ (= OH + HO₂/RO₂) radical budget. Best agreement is achieved when HONO at Dome C is assumed to originate from snow emissions with the emission strength evaluated by Legrand et al. (2014). Using a 1-D model, Kukui et al. (2014) show that the concentrations of HONO corresponding to about 20–30% of measured HONO are consistent with those calculated from the budget analysis of OH radicals with the concentrations of NO₂ either calculated from NO measurements assuming PSS or observed by Frey et al. (2014). Therefore, the production of OH by HONO photolysis is adjusted consequently and the 0-D box model (Kukui et al., 2014) is used to calculate all other production rates of OH. Note that even lowering HONO to 20–30% of the measured values, this species remains the major primary source of radicals at Dome C. Applying the isotope ¹⁷O-excess transfer (Morin et al., 2011) and the OHₚrod. isotope mass balance, Δ¹⁷O(OH)ₚrod. on average equals 5 ± 2‰ (1 σ). Because the major process leading to the emission of HONO from the snowpack is the photolysis of nitrate, which possesses a Δ¹⁷O value of approximately 32‰, both in the snow “skin layer” (Erbland et al., 2013) and in the top 10 cm of snow (Frey et al., 2009), we have assumed that Δ¹⁷O(HONO)ₐtm = Δ¹⁷O(NO₃⁻)ₐsnow as both oxygen atoms of HONO can be tracked back to the nitrate. An example of the isotope mass balance calculation is given in Table 1. Figure 7 shows the diurnally-integrated average of the Δ¹⁷O(OH). Δ¹⁷O(OH) varies in a narrow range, between 1 and 3‰. An estimation of the Δ¹⁷O signature for the NO₂ + OH channel, Δ¹⁷O(NO₃⁻)R₂, that accounts for the ¹⁷O-excess carried by the OH radical results in values ranging between 20 and 23‰. Averaging over the same time period as the nitrate atmospheric sampling, diurnally-integrated average Δ¹⁷O(NO₃⁻) values of 21–22‰ ± 3‰ can be estimated for December (Table 2). These values are 6–8‰ lower than the observed atmospheric values for Δ¹⁷O(NO₃⁻) (27–30‰ during OPALE, Fig. 2 and Table 2). The source of discrepancy between ob-

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served and modeled Δ¹⁷O(NO₃⁻) is presently unknown. A critical evaluation may offer nevertheless some clues.

# 4 Discussion

## 4.1 Alternative sources of NO₂

A possible explanation for the underestimation of Δ¹⁷O(NO₃⁻) involves the potential role of reactive halogen chemistry in the troposphere over the Antarctic plateau (Bloss et al., 2010; Morin et al., 2008). Reactive halogen oxides (XO = BrO, ClO, IO, etc.) are produced through the reaction of halogen radicals (X) with ozone, a pathway that plays an important role in the catalytic process responsible for ozone depletion events (ODEs) observed in the Arctic boundary layer since the 1980s (Fan and Jacob, 1992; Simpson et al., 2007):

$$X + O_3 \rightarrow XO + O_2 \quad (R6)$$

In terms of the chemical budget of NOₓ, the impact of XO can occur via two chemical mechanisms (see Sect. 4.2 for the second mechanism involving the formation of halogen nitrate, XONO₂). First, XO can oxidize NO to NO₂, a pathway that competes with the NO + O₃ and NO + HO₂/RO₂ reactions in terms of NO oxidation:

$$XO + NO \rightarrow X + NO_2 \quad (R7)$$

For conditions typical of the Antarctic boundary layer, 1 pmol mol⁻¹ of XO has roughly the same oxidizing power as 4 nmol mol⁻¹ of ozone in terms of NO oxidation (Atkinson et al., 2007). Therefore, when halogen oxides are present at relevant levels, the R7 reaction can result in concentrations of NO₂ that are higher than that predicted from the balance between NO₂ destruction via photolysis and production through the reaction
of NO with O₃ or HO₂/RO₂ (i.e., the extended Leighton mechanism):

$$ \frac{[NO_2]}{[NO]} = \frac{k_{NO+O_3}[O_3] + k_{NO+HO_2}[HO_2] + k_{NO+RO_2}[RO_2] + k_{NO+XO}[XO]}{j_{NO_2}} \quad (6) $$

The interaction of XO in the NOₓ cycle at Dome C would have important implications for the Δ¹⁷O of atmospheric nitrate. The production of halogen oxide radicals proceeds through a direct transfer of a terminal oxygen atom from ozone to the XO product (Zhang et al., 1997). Therefore, it is expected that the Δ¹⁷O of XO is equal to Δ¹⁷O(O₃)ₜerm., which means that the reaction of NO with XO is roughly equivalent to the NO + O₃ reaction in terms of Δ¹⁷O transfer to NO₂ (note that the NO + XO transfer is greater than NO + O₃ as in the later case, part of the central O₃ atom participates in the reaction). The participation of XO species in the oxidation of NO thus leads to a greater Δ¹⁷O transfer to NO₂ by effectively increasing the value of α. However, on the Antarctic plateau, BrO did not exceed 2–3 pmol mol⁻¹ at most during OPALE campaign (Frey et al., 2014). Including BrO chemistry would only increase α by 2% (due to the specific form of α ≡ 1/(1+x)) of which is to low to significantly influence Δ¹⁷O(NO₂) and ultimately Δ¹⁷O(NO₃⁻). In the absence of measurements of other halogens we cannot completely rule out a role of the halogen chemistry there. However, even with α = 1, its maximum but unrealistic value due to the high concentration of HO₂, Δ¹⁷O(NO₃⁻) would reach the range of 23–25‰, in better agreement with the observations but still significantly lower. Similarly, in the event of a non isotopic steady state of NO₂ (Michalski et al., 2014), it is very unlikely that Δ¹⁷O(NO₂) could reach values greater than its primary snow nitrate source (i.e. Δ¹⁷O(NO₂) > Δ¹⁷O(NO₃⁻)ₐnow = 30–35‰ in summer at Dome C, (Erbland et al., 2013; Frey et al., 2009), still leaving the predicted Δ¹⁷O(NO₃⁻) underestimated with respect to atmospheric observations.
## 4.2 Alternative oxidation pathways of NO₂

Considering Reaction (R5) as the main source of HNO₃, an alternative approach is to consider that OH bears a higher Δ¹⁷O than the estimate calculated previously. Assuming a β of one, which seems again unrealistic, will increase Δ¹⁷O(NO₃⁻) by 1‰ at most (Table 2), still insufficient to explained atmospheric observations as NO + HO₂ remain a major source of OH, independently of the assumed β.

Alternatively, if measured HONO concentrations are considered instead of those assumed to constrain by the HOₓ budget (i.e. 4 times lower than measured), average Δ¹⁷O(NO₃⁻) values of 23–24‰ are calculated (Table 2), again systematically lower than the observed range of 27–30‰. However, given the significant uncertainty surrounding the isotopic composition of HONO and its relative contribution to total OH production at Dome C, it is not possible to make a firm conclusion in this regard.

Therefore, neither the common sources of NO₂ nor the daytime formation of HNO₃ seems to be able to explain the high Δ¹⁷O(NO₃⁻) values of atmospheric nitrate observed at Dome C in summer. When this observation is taken together with the high NO₂/NO ratio observed by Frey et al. during two summer seasons at Dome C (Frey et al., 2013, 2014), clearly our current understanding of the NOₓ chemistry on the Antarctic plateau seems to be incomplete.

There are several other processes that possibly account for the disagreement between the measurements and mass balance calculations. Indeed, in addition to its impact on NOₓ cycling through the R7 pathway, an increasing body of evidence points towards reactive halogen chemistry as a major NOₓ sink and source of nitrate via the production and subsequent hydrolysis of XNO₃ species (Sander et al., 1999; Savarino et al., 2013; Vogt et al., 1996):

$$XO + NO_2 + M \rightarrow XNO_3 + M \quad (R8)$$

$$XNO_3 + H_2O_{(liquid)} \rightarrow HNO_{3(aq)} + HOX \quad (R9)$$

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A critical analysis of the CHABLIS data led Bauguitte et al. (2012) to conclude that the Reactions (R8) and (R9) pathway exerted predominant control over the chemical loss rate of NOₓ during the campaign, despite the significant uncertainties involved in the parameterization of the uptake processes (Finlayson-Pitts, 2009). This implies that XNO₃ uptake may also represent a significant source of nitrate at Dome C should halogen oxide radicals (XO) be present at the required concentration. Experimental (Gane et al., 2001) and theoretical (McNamara and Hillier, 2001) studies suggest that the oxygen atom initially associated with XO combines with the N atom of NO₂ to form nitrate, thus transferring the isotopic signature of both XO and NO₂. The specific Δ¹⁷O value induced by XNO₃ hydrolysis can thus be expressed as follows (Morin et al., 2007):

$$ \Delta^{17}\text{O}(\text{NO}_3^-)_{\text{R6}} = \frac{2}{3}\Delta^{17}\text{O}(\text{NO}_2) + \frac{1}{3}\Delta^{17}\text{O}(\text{O}_3)_{\text{term}} \quad (7) $$

efficiently bypassing the OH ¹⁷O-excess budget. Through consideration of the increased Δ¹⁷O transfer associated with Reaction (R8), the observations of Δ¹⁷O(NO₃⁻) during December can be reconciled with the values calculated by mass-balance if approximately 10–20 % of total nitrate production is assumed to occur via XNO₃ hydrolysis. However, no sufficient halogen concentration has been observed on the Antarctic plateau to sustain such chemical pathway but we note that chlorine chemistry has never been probed on the Antarctic plateau.

There is increasing body of evidence that heterogeneous hydrolysis of NO₂ can be a possible source of HONO and HNO₃ in acidic conditions (Finlayson-Pitts, 2009), with the potential to explain the difference between the calculated and measured atmospheric Δ¹⁷O(NO₃⁻) values. This mechanism would represent a source of nitrate with a Δ¹⁷O value roughly equivalent to the nitrate originally present in the surface snow (i.e., 30–35‰), a signature significantly higher than that induced by Reaction (R5). If this production mechanism is active at the air–snow interface at Dome C and results in the slow emission of nitrate to the atmosphere via physical release after its formation,

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it would act to increase the $\Delta^{17}\text{O}$ value of nitrate in the boundary layer relative to the local $\Delta^{17}\text{O}(\text{NO}_3^-)_{\text{R5}}$ oxidation signature. However, considering the propensity of nitric acid to stick on snow (Crowley et al., 2010), the snowpack to act as a sink rather than a source of nitric acid (Dibb, 2004; Erbland et al., 2013) and the fast $\text{NO}_x$ recycling that should take place within the snowpack, it is very unlikely that $\Delta^{17}\text{O}(\text{NO}_3^-)$ could be explained by a direct nitric acid emissions from snow, which has been ultimately shown to be limited (Slusher et al., 2010; Erbland et al., 2013; Berhanu et al., 2014).

A critical analysis of $\Delta^{17}\text{O}(\text{NO}_3^-)$ shows in fact that such high values correspond mainly to the nighttime chemistry of $\text{NO}_x$ (Michalski et al., 2003; Morin et al., 2008). Nighttime chemistry involves species like $\text{N}_2\text{O}_5$ and $\text{NO}_3$ in the process of forming $\text{HNO}_3$ and again efficiently by-pass the OH pathway. It is conceivable that below the photic zone, within the snowpack, $\text{N}_2\text{O}_5$ and $\text{NO}_3$ could be produced when $\text{O}_3$ and $\text{NO}_2$ are transported at depth but there is no reason to think that such dark $\text{NO}_x$ chemistry could in a way or in another survive the photic zone transition and thus influences the overlying atmosphere.

Stratospheric nitrate deposited to the surface snow during winter, which has been observed to possess $\Delta^{17}\text{O}(\text{NO}_3^-)$ values in the range of 35–41 ‰ (Erbland et al., 2013) and possibly more, may act to buffer the $\Delta^{17}\text{O}$ of the atmospheric nitrate reservoir via evaporation late into the spring and summer. However, this seems again unlikely given the rapidity of $\text{NO}_x$ cycling and oxidative loss at Dome C during this time (Frey et al., 2013; Legrand et al., 2009).

While it is presently difficult to determine the precise nature of the process(es) leading to the relatively large $^{17}$O-excess values observed for atmospheric nitrate at Dome C, the correlation observed between the $\delta^{15}\text{N}$ and $\Delta^{17}\text{O}$ values of atmospheric nitrate (see Sect. 3.1) provides at least one direct line of evidence that the high $\Delta^{17}\text{O}(\text{NO}_3^-)$ values observed during spring and early summer could be associated with snowpack emissions of $\text{NO}_x$. Considering only samples collected at Dome C between October and December, both those reported here and those collected in 2009 and described by Erbland et al. (2013), a strong anticorrelation ($r = -0.90$) is observed between the
$\delta^{15}$N and $\Delta^{17}$O values of atmospheric nitrate (Fig. 8). In other words, the atmospheric nitrate sampled in early spring, which is heavily depleted in $^{15}$N due its formation from snowpack NO$_x$ emissions, possessed consistently higher $\Delta^{17}$O values than the nitrate sampled directly after this period of maximum snow photochemistry. This finding suggests that the mechanism producing enhanced $\Delta^{17}$O(NO$_3^-$) values observed during this time is tightly coupled in time and space with the intensity of NO$_x$ emissions from the snowpack, an observation very similar to that of Morin et al. (2012), who detected a similar relationship between $\delta^{15}$N and $\Delta^{17}$O for atmospheric nitrate in the springtime boundary layer over Barrow, Alaska (71°N). The authors of this study attributed the observed correlation to the coupling of snowpack NO$_x$ emissions and reactive halogen chemistry, suggesting that these two processes were interrelated and mutually strengthening. In the case of the OPALE 2011–2012 data, the correlation between $\delta^{15}$N and $\Delta^{17}$O could arise from any of the potential pathways previously discussed. For example, as proposed by Morin et al. (2012), the R8 and R9 pathways may be enhanced under periods of intense snowpack emissions. Alternatively, a correlation could result from an increased contribution to total OH production from the photolysis of HONO, which is co-emitted with NO$_x$ via nitrate photochemistry (Grannas et al., 2007) and may induce a larger $^{17}$O-excess in OH as compared to the conventional O($^1$D) + H$_2$O pathway. Furthermore, the hydrolysis of NO$_2$ in snow, should it contribute significantly to nitrate production at Dome C, is likely amplified during periods when concentrations NO$_2$ are high in the snowpack interstitial air due to nitrate/nitrite photochemistry. Therefore, while the processes responsible for driving the formation of atmospheric nitrate at Dome C during summer cannot be unambiguously identified, the isotopic results presented here clearly indicate that snowpack emissions result in enhanced $\Delta^{17}$O transfer to nitrate. Our understanding of NO$_x$ chemistry above the snow surface at Dome C is therefore incomplete.

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Constraining the propagation of ozone's ¹⁷O-excess signature within the NOₓ cycle is critical in polar areas where the opportunity is offered to extend atmospheric investigations based on Δ¹⁷O measurements to the glacial/interglacial time scale using deep ice core records of nitrate. However, the factors governing the present-day isotopic composition of atmospheric nitrate over the Antarctic plateau remain poorly understood, primarily due to the complex nature of the boundary layer photochemistry initiated during spring by NOₓ emissions from the snowpack.

An isotopic mass balance performed for atmospheric nitrate during December 2011, informed by in situ oxidant concentration measurements conducted within the framework of the OPALE field study, suggests the existence of an unexpected process bypassing the commonly accepted daytime chemistry of NO₂ (i.e. NO₂ + OH) that contributes significantly to the atmospheric nitrate budget over Dome C. The strong negative correlation observed between the δ¹⁵N and Δ¹⁷O values of nitrate between October and December suggests that this unknown process is enhanced during periods of intense emissions from the snowpack. Potential explanations for this observation include: (i) an increased Δ¹⁷O transfer from OH due to its formation from the photolysis of HONO released from the snowpack, (ii) heterogeneous hydrolysis of NO₂ due to the high concentrations of NO₂ in the snowpack interstitial air, and (iii) the co-emission of reactive halogen species that act as an intermediate in the transfer of Δ¹⁷O from ozone to nitrate. Further research is needed to solve the many inconsistencies (e.g. high NO₂/NO ratio, high concentration of NO₂, unresolved HONO atmospheric concentration, interference such as HO₂NO₂, isotope mass balance) observed during the OPALE experiments.

**Acknowledgements.** The research leading to these results has received funding from the European Community's Seventh Framework Programme (FP7/2007–2013) under the grant agreement number 237890. We would like to thank INSU for its financial support for lab experiments through its LEFE program. The Agence nationale de la recherche (ANR) is gratefully acknowledged for its financial support through the OPALE project (contract NT09-451281). The Insti-

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tute Polaire Paul-Emile Victor (IPEV) supported the research and polar logistics through the program SUNITEDC No. 1011. This work has been partially supported by a grant from Labex OSUG@2020 (Investissements d'avenir - ANR10 LABX56. We would also like to thank all the field team members present during the OPALE campaign. Meteorological data were obtained from "IPEV/PNRA: Routine Meteorological Observation at Station Concordia".

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<table><caption>**Table 1.** Rate of production and sink of OH and mass balance calculation of Δ¹⁷O.</caption><thead><tr><th>Conditions for 19 Dec 2011, 3:45 p.m.<br>OH = 3.96 molecules cm<sup>-3</sup></th><th><sup>a</sup>Median rate in 10<sup>5</sup><br>molecules cm<sup>-3</sup> s<sup>-1</sup></th><th>Δ¹⁷O<sub>i</sub> in %</th></tr></thead><tbody><tr><td colspan="3"><strong>Net sources of OH</strong></td></tr><tr><td>P1 HONO + hν → OH + NO</td><td>5.1<sup>b</sup></td><td>32</td></tr><tr><td>P2 H<sub>2</sub>O<sub>2</sub> + hν → 2OH</td><td>1.7</td><td>2</td></tr><tr><td>P3 O<sub>3</sub> + hν + H<sub>2</sub>O → 2OH</td><td>0.6</td><td>20</td></tr><tr><td>P4 CH<sub>3</sub>OOH + hν → HO<sub>2</sub> + OH</td><td>0.3</td><td>0</td></tr><tr><td colspan="3"><strong>Recycling RO<sub>2</sub> → OH</strong></td></tr><tr><td>P5 NO + HO<sub>2</sub> → NO<sub>2</sub> + OH</td><td>7.7</td><td>0</td></tr><tr><td>P6 HO<sub>2</sub> + O<sub>3</sub> → OH + 2O<sub>2</sub></td><td>0.4</td><td>0</td></tr><tr><td colspan="3"><strong>Net sink of OH</strong></td></tr><tr><td>L1 CO + OH → HO<sub>2</sub> + CO<sub>2</sub></td><td>6.3</td><td></td></tr><tr><td>L2 CH<sub>4</sub> + OH → CH<sub>3</sub>O<sub>2</sub> + H<sub>2</sub>O</td><td>2.6</td><td></td></tr><tr><td>L2 HCHO + OH → HO<sub>2</sub> + CO</td><td>0.8</td><td></td></tr><tr><td>L4 CH<sub>3</sub>CHO + OH → CH<sub>3</sub>CO<sub>3</sub></td><td>0.9</td><td></td></tr><tr><td>L5 O<sub>3</sub> + OH → HO<sub>2</sub> + O<sub>2</sub></td><td>0.6</td><td></td></tr><tr><td>L6 H<sub>2</sub> + OH + O<sub>2</sub> → HO<sub>2</sub> + H<sub>2</sub>O</td><td>0.60</td><td></td></tr><tr><td>L7 CH<sub>3</sub>OOH + OH → CH<sub>3</sub>O<sub>2</sub> + H<sub>2</sub>O</td><td>0.5</td><td></td></tr><tr><td>L8 H<sub>2</sub>O<sub>2</sub> + OH → HO<sub>2</sub> + H<sub>2</sub>O</td><td>0.3</td><td></td></tr><tr><td colspan="3"><strong>Net OH losses</strong></td></tr><tr><td>L9 NO<sub>2</sub> + OH → HNO<sub>2</sub></td><td>3.9</td><td></td></tr><tr><td>L10 NO + OH → HONO</td><td>0.6</td><td></td></tr><tr><td>L11 OH + RO<sub>2</sub> → products</td><td>0.5</td><td></td></tr><tr><td>L12 OH + RO<sub>2</sub>NO<sub>2</sub> → products</td><td>0.6</td><td></td></tr><tr><td>L13 OH + HONO → NO<sub>2</sub> + H<sub>2</sub>O</td><td>0.2</td><td></td></tr><tr><td>L14 OH + HNO<sub>3</sub> → H<sub>2</sub>O + NO<sub>3</sub></td><td>0.0</td><td></td></tr><tr><td colspan="3"><strong>Isotope exchange</strong></td></tr><tr><td>E<sub>1</sub> HQ + H<sub>2</sub>O ⇌ HO + H<sub>2</sub>Q</td><td>24.3</td><td></td></tr><tr><td colspan="3"><strong>NO<sub>2</sub> main source</strong></td></tr><tr><td>N1 NO + O<sub>3</sub> → NO<sub>2</sub> + O<sub>2</sub></td><td>27.0</td><td>37</td></tr><tr><td colspan="3"><sup>17</sup>O-excess NO<sub>2</sub></td></tr><tr><td>α = (N1/(N1 + P5))</td><td>0.78</td><td></td></tr><tr><td>Δ<sup>17</sup>O(NO<sub>2</sub>)</td><td></td><td>29</td></tr><tr><td colspan="3"><sup>17</sup>O-excess OH</td></tr><tr><td>Δ<sup>17</sup>O(OH)<sub>prod</sub> = (ΣP<sub>i</sub> × Δ<sup>17</sup>O<sub>i</sub>)/ΣP<sub>i</sub></td><td></td><td>5.8</td></tr><tr><td>β = ΣL<sub>i</sub>/(ΣL<sub>i</sub> + E<sub>1</sub>)</td><td>0.43</td><td></td></tr><tr><td>Δ<sup>17</sup>O(OH)</td><td></td><td>2.5</td></tr></tbody></table>

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<footnote><sup>a</sup>: Production rates obtained from a 0-D box model (see Kukui et al., 2014 for details).</footnote>

<footnote><sup>b</sup>: HONO production rate divided by a factor 4 to balance the HO<sub>2</sub> radical budget (see Kukui et al., 2014; Legrand et al., 2014 for justification).</footnote>

<footnote><sup>c</sup>: HONO is assumed to be formed by the photodissociation of nitrate in snow. Δ¹⁷O(NO₃)ₜnow is therefore assigned to HONO. The rest of the ¹⁷O-excess transfer (i.e. P₂ to P₆ and N₁) follows the rules established in (Morin et al., 2011) and a Δ¹⁷O(O₃)ₜbulk = 26‰.</footnote>
<table><caption>**Table 2.** Comparison of measured and calculated $\Delta^{17}\text{O}(\text{NO}_3^-)$ values.</caption><thead><tr><th rowspan="2">Sampling Period</th><th rowspan="2">Measured</th><th colspan="4">Calculated</th></tr><tr><th>α constrained<br>by observations<br>$\Delta^{17}\text{O}(\text{OH})*$<br>based on $\text{HO}_x$ budget</th><th>α = 1<br>$\Delta^{17}\text{O}(\text{OH})*$<br>based $\text{HO}_x$ budget</th><th>α constrained<br>by observations<br>$\beta = 1$</th><th>α constrained<br>by observations<br>$\Delta^{17}\text{O}(\text{OH})$ based on<br>observed HONO</th></tr></thead><tbody><tr><td>10 Dec–16 Dec</td><td><strong>29.6</strong></td><td>21.9</td><td>25.6</td><td>22.6</td><td>27.0</td></tr><tr><td>16 Dec–23 Dec</td><td><strong>29.0</strong></td><td>21.0</td><td>25.6</td><td>21.7</td><td>26.3</td></tr><tr><td>23 Dec–30 Dec</td><td><strong>27.8</strong></td><td>21.6</td><td>25.4</td><td>22.0</td><td>25.7</td></tr><tr><td>30 Dec–2 Jan</td><td><strong>27.3</strong></td><td>21.5</td><td>25.3</td><td>22.4</td><td>24.9</td></tr></tbody></table>

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<footnote>* HONO production rate divided by a factor 4 to balance the $\text{HO}_x$ radical budget (see Kukui et al. (2014) and Legrand et al. (2014) for justification).</footnote>
<figure><img src="image_131.png" /><figcaption>**Figure 1.** Atmospheric nitrate concentrations observed between October 2011 and January 2012. The samples collected during the intensive measurement period of the OPALE campaign (December 2011) are indicated with open circles.</figcaption></figure>

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<figure><img src="image_136.png" /><figcaption>**Figure 2.** $\Delta^{17}$O (primary y axis) and $\delta^{15}$N (secondary y axis) of atmospheric nitrate collected between October 2011 and January 2012. The samples collected during the intensive measurement period of the OPALE campaign (December 2011) are indicated with open symbols.</figcaption></figure>

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<figure><img src="image_141.png" /><figcaption>**Figure 3.** $\Delta^{17}\text{O(O}_3)_{\text{bulk}}$ values for the 60 ambient air sampling done at Dome C throughout 2012. Vertical error bars refer to the total uncertainty estimated for the technique (±1.7‰).</figcaption></figure>

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<figure><img src="image_150.png" /><figcaption>**Figure 4.** Comparison of $\Delta^{17}\text{O}(\text{O}_3)_{\text{bulk}}$ values obtained at Dome C with those previously reported by Vicars and Savarino (2014) at other sites. The box plot indicates the interquartile range (box) and the median (line), maximum, and minimum values. The mean value is denoted by a circle.</figcaption></figure>
<figure><img src="image_151.png" /><figcaption>**Figure 5.** Quantitative assessment of the daily averaged trend in the $\Delta^{17}\text{O}$ of NO$_2$ at Dome C during December 2011 derived from concurrent measurements of ozone, NO, and HO$_2$/RO$_2$.</figcaption></figure>

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<figure><img src="image_156.png" /><figcaption>**Figure 6.** December 2011 time-series for $\beta$, the fraction of the $^{17}$O-excess originally associated with the OH radical that is preserved against isotopic exchange with water.</figcaption></figure>

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<figure><img src="image_161.png" /><figcaption>**Figure 7.** Same as Fig. 5 but for $\Delta^{17}\text{O}$ of OH.</figcaption></figure>

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<figure><img src="image_171.png" /><figcaption>**Figure 8.** Relationship observed between the $\delta^{15}$N and $\Delta^{17}$O values of atmospheric nitrate present at Dome C between October and December during both the 2009 and 2011 summer campaigns. Error bars indicate the typical analytical uncertainties associated with the measurements.</figcaption></figure>