Atmospheric and snow nitrate isotope systematics at Summit, Greenland: the reality of the post-depositional effect

20 The effect of post–depositional processing on the preservation of snow nitrate isotopes at Summit, Greenland remains a subject of debate which hinders the interpretations of ice–core nitrate concentrations and isotope records. Here we present the first year–round observations of atmospheric aerosol nitrate and its isotopic compositions at Summit, and compare them with published surface snow and snowpack 25 observations. The atmospheric δN(NO3 ) remained negative throughout the year, ranging from –3.1 ‰ to –47.9 ‰ with a mean of (–14.8 ± 7.3) ‰, and displayed no apparent seasonality that is different from the distinct seasonal δN(NO3 ) variations observed in snowpack. The spring average aerosol δN(NO3 ) was (–17.9 ± 8.3) ‰, significantly depleted compared to snowpack spring average of (4.6 ± 2.1) ‰, with 30 surface snow δN(NO3 ) of (–6.8 ± 0.5) ‰ that is in between. The differences in aerosol, surface snow and snowpack δN(NO3 ) are best explained by the photo-driven post–depositional processing of snow nitrate, with potential contributions from https://doi.org/10.5194/tc-2021-355 Preprint. Discussion started: 10 December 2021 c © Author(s) 2021. CC BY 4.0 License.


Introduction
Ice-core nitrate and its isotopes are potential proxies to constrain atmospheric  Savarino et al., 2016;Wolff et al., 2008). Nitrate is chemically reactive in snow upon exposure to sunlight and thus its deposition to snow is not irreversible (Blunier et al., 2005;Erbland et al., 2013;Frey et al., 2009).
Numerous studies across Antarctica and Greenland have observed decreases in snow nitrate concentrations with depth in the snowpack (Erbland et al., 2013;Frey et al., 2009;55 Mulvaney et al., 1998;Röthlisberger et al., 2000) and/or emissions of NOx and HONO from snowpack (Dibb et al., 1998;Frey et al., 2015;Honrath et al., 2002;Jones et al., 2013;Frey et al., 2009;Geng et al., 2015;Jiang et al., 2021;Shi et al., 2015;Winton et al., 2020). 65 Post-depositional processing of snow nitrate is mainly initiated by photolysis Erbland et al., 2013;Frey et al., 2009;Zatko et al., 2016). The evaporation of nitrate from snow grains may also contribute but this process has not been directly observed/evidenced in the field. Observations and/or modelling of snowpack nitrate concentration and isotope profiles across many different sites (e.g., 70 Summit in Greenland, Dronning Maud Land (DML) and Dome A/Dome C in Antarctica) in general agree that photolysis dominates post-depositional processing (Erbland et al., 2013;Frey et al., 2009;Geng et al., 2015;Jiang et al., 2021;Winton et al., 2020;Shi et al., 2015;Shi et al., 2019). The degree of the photo-driven post-depositional processing is influenced by three main factors including snow accumulation rate, surface actinic 75 flux and light penetration depth in snow (i.e., the photic zone where actinic flux decreases exponentially) (Zatko et al., 2013). Snow and ice-core nitrate isotope records have shown variations in δ 15 N(NO3 -) in response to varying snow accumulation rate as well as light-absorbing impurities (e.g., BC, dust, etc.) that influences light penetration depth in snow. For example, Geng et al. (2014) found correlations between δ 15 N(NO3 -) 80 and snow accumulation rate across the GISP2 ice core record, except in periods with very low snow accumulation rate (<0.08 m ice a -1 ) and high dust concentrations when δ 15 N(NO3 -) was correlated with dust concentration. These correlations reflect the effect of snow accumulation rate and snow light absorbing impurities on the degree of postdepositional processing, respectively. At the West Antarctica ice sheet divide, where 85 snow accumulation rate is high (0.24 m ice a -1 ) at present, a decreasing trend in snow accumulation rate since 2400 yr BP led to an increasing trend in the degree of postdepositional processing as indicated by elevated δ 15 N(NO3 -) (Sofen et al., 2014).
Variations in surface actinic flux (especially the UVB radiation) would also induce changes in the degree of post-depositional processing and leave signals in the preserved 90 nitrate in snow and ice cores. Previous studies (Erbland et al 2013;Frey et al., 2009;McCabe et al. 2007) proposed that δ 15 N(NO3 -) preserved in snow and ice cores may serve as a proxy of total column ozone (TCO) due to its influence on surface UVB https://doi.org/10.5194/tc-2021-355 Preprint. Discussion started: 10 December 2021 c Author(s) 2021. CC BY 4.0 License. radiation, while a recent study suggested the preserved δ 15 N(NO3 -) is more sensitive to snow accumulation rate and light penetration depth, but less to TCO (Winton et al.,95 2020). Nevertheless, in periods with relatively constant snow accumulation rate but distinct surface actinic flux, e.g., the switch of the polar night and polar day over a year, and the Antarctic ozone hole period, changes in the degree of post-depositional processing and thus the associated isotope effects are expected. Using a snow column photochemical model (the TRANSITS model by Erbland et al., (2015)), Jiang et al. 100 (2021) explicitly quantified the effects of post-depositional processing on snow nitrate and its isotopes on seasonal scale at Summit, Greenland. Owing to the seasonal differences in surface actinic flux, the model predicted a seasonal variation in δ 15 N(NO3 -) snowpack similar to the observations. On annual scale, the model predicted a ≈ 4 % net nitrate mass loss, which is within the range estimated by previous studies 105 (Burkhart et al., 2004;Dibb et al., 2007) but is subject to uncertainties in the fraction of the snow-sourced nitrate exported from the region. In contrast, the model predicted minimum changes in Δ 17 O of snow nitrate on both seasonal and annual scale because the photo-driven post-depositional processing affects Δ 17 O mainly from the cage effect (i.e., the intermediate photo-products (NO2and NO2) exchange with water oxygen or 110 react with radicals such as OH in snow grains to regenerate nitrate before being emitted to the atmosphere) which is however minimum at Summit given the high snow accumulation. The study by Jiang et al. (2021) further suggested that seasonal δ 15 N(NO3 -) variations in snowpack at Summit, Greenland is caused by photo-driven post-depositional processing, an alternative to previous interpretations that attributed collected atmospheric and surface snow samples in May and June. The process of photolysis of snow nitrate to NOx, oxidation of snow-sourced NOx to nitrate, followed 125 by re-deposition of snow-sourced nitrate will render the isotopic composition of atmospheric and surface snow nitrate similar to each other. Nitrate at depth but still in the photic zone experiences photolysis, but is isolated from surface deposition, making post-depositional loss more apparent in the isotope observations. Therefore, in order to reflect the full picture of post-depositional processing, considering snow samples 130 covering the entire photic zone (~ 40 cm at Summit) is necessary (Jiang et al., 2021).
To thoughtfully evaluate the effects of post-depositional processing at Summit, Greenland, and to verify the modeling results by Jiang et al. (2021), nitrate isotopes in the atmosphere and in snow covering a full cycle of polar night vs. polar day (i.e., a cycle with distinct actinic flux variations) are necessary. Here, we present the first year-135 round observations of nitrate isotopes in the air at Summit, and compare them with similar observations in surface snow and in snow at depth (i.e., snowpack) to conduct a comprehensive evaluation on the seasonality in nitrate isotopes in both air and in snow, as has already been done in Antarctica (Erbland et al., 2013;Frey et al., 2009;Winton et al., 2020). These observations provide information regarding the evolution of nitrate 140 isotopes from atmospheric nitrate to its final preservation in snowpack, which is critical to assess the post-depositional changes of nitrate isotopes.

aerosol sampling and measurements
glass fiber filters was first extracted by 18 MΩ water via centrifugation using Millipore Centricon™ filter units. The samples were then measured for nitrate concentrations by 155 ion chromatography. Among these samples, 54 out of 97 were determined to be valid by comparing the extracted nitrate concentration with blank. These samples were then individually concentrated on a 0.3 mL resin bed with anionic exchange resin (Bio-Rad™ AG 1-X8, chloride form) and eluted with 5 × 2 mL of NaCl solution (1M). The isotopic compositions of each sample were determined by using the bacterial denitrifier 160 method. Briefly, NO3in each sample was converted to N2O by denitrifying bacteria under anaerobic conditions. N2O was then thermally decomposed into N2 and O2 on a gold tube heated at 800 ℃. The N2 and O2 were then separated by a gas chromatography column and injected into an isotope ratio mass spectrometer (Thermo Finnigan™ MAT  (Fibiger et al., 2013;Fibiger et al., 2015;Geng et al., 2014;Hastings et al., 2004;Jarvis et al., 2009;Kunasek et al., 2008). Details about these data (e.g., sample type, depth, age) and the corresponding references are listed in Table   1. Note that in some early publications only the seasonal averages instead of the original data with finer resolution were available. These data were compiled to produce a dataset 175 including all seasons for nitrate in the air, surface snow and snowpack by averaging samples covering multiple years and/or by different groups to reduce the spatial and temporal heterogeneities. For samples with resolution finer than monthly, we compiled them as their mass-weighted monthly averages (if the mass information for each sample is known), and for samples with coarser than monthly resolution, seasonal averages 180 were used.
For aerosol and surface snow samples, age information was indicated as the time of sampling. Snowpack samples require a conversion from depth to age. The snowpack https://doi.org/10.5194/tc-2021-355 Preprint. Discussion started: 10 December 2021 c Author(s) 2021. CC BY 4.0 License. samples from Hastings et al. (2004) and Kunasek et al. (2008) were dated by seasonal binning according to measured accumulation rate and water isotopes, and their age 185 information was used as is. For samples from Geng et al. (2014), we recalculated the dating by bamboo stake measured snow accumulation data (Burkhart et al., 2004;Dibb et al., 2004;Kuhns et al, 1997) constrained by snow density and further justified by seasonal peaks of Na + and Cl -/Na + ratio. This is similar to the dating method in Hastings et al. (2004) and the only difference is which proxy was used as the seasonal marker.

190
Briefly, we used the bamboo stake measurements of weekly snow accumulation at Summit and the snowpack density profile to estimate the deposition timing of each samples in the 2.1 m snowpack that was collected in July of 2007. We first converted the thickness of each sample (referred to as Dm) to a fresh snow thickness (referred to as Df) by the following equation: Where is the real snow density at each depth from field measurement (Geng et al., 2014), and is the fresh snow density (0.32 g.cm -3 ; Dibb et al., 2004). These fresh snow thicknesses were then stacked to construct an idealized snow depth profile without densification due to compaction and/or metamorphism. This idealized depth 200 profile was then matched the stacked depth by the observed average weekly snow accumulation rate to determine the exact age for each sample. A previous study showed that using the stack measured accumulation rate is capable of reconstructing the vertical profile of snowpack nitrate (Burkhart et al., 2004). This dating method has uncertainties, mostly owing to the large variability of measured accumulation rate among different 205 stakes (Burkhart et al., 2004). To reduce the uncertainties in our dating results, we calculated their monthly average and compared with aerosol and/or surface snow data with a similar or coarser time resolution. The compiled δ 15 N and Δ 17 O data in monthly resolution display seasonal patterns similar to their original seasonal variations observed in snowpack, and the Cl -/Na + ratio of the compiled samples also displays 210 summer high and winter low as has been previously observed snowpack or firn cores (Geng et al., 2014), corroborating the dating method in terms of capturing the Table1. Nitrate isotope data information and references.  resolutions are plotted in Figure 2. These compiled data of atmospheric, surface snow 270 and snowpack averages should represent the status of nitrate before deposition, after deposition, and archival, respectively. To validate our dating results on the snowpack data, we also plotted the resampled monthly snowpack Na + concentration and Cl -/Na + ratio. As shown in Figure 2e, the Na + concentration and Cl -/Na + ratio displayed clear winter and summer peak, respectively, indicating a general reliability of our dating 275 method. We also calculated the accumulated UV-B * daily dose for nitrate deposited in different weeks of a year using Eq (2)

Year-round atmospheric nitrate concentrations and isotopes at Summit,
where A(t) and ze represent the weekly snow accumulation rate and e-folding depth This gives a first order estimation of the total radiation (i.e., the degree of postdepositional processing) that the archived nitrate experienced at Summit.  Table 1.   lower the export fraction of the snow-sourced nitrate compared to summer, which tends to lower the spring atmospheric δ 15 N(NO3 -) as more snow-sourced nitrate with extremely low δ 15 N will accumulate in the local boundary layer. Honrath et al. (2002) found that at Summit, in summer the snow-sourced nitrate (their measured form was starting time of photolysis), and ta is the time for snow nitrate to reach a depth below the snow photic zone (i.e., the archival layer) (3 times the e-folding depth). t is the time 420 variable between t0 and ta. ε and J represent the N isotope fractionation factor and nitrate photolysis rate constant for snow nitrate at surface conditions, respectively. Both ε and J varies seasonally owing to the timely-varied actinic flux, while J also varies with depth which is constrained by the exponential term. ze and A(t) represent the e-folding depth and snow accumulation rate, respectively. Essentially Eq(3) is the same as Eq (2), 425 because they both describe the total actinic flux received by a specific snow layer before archival, but Eq(3) provides a direct way to evaluate the induced isotope effects on δ 15 N.  The δ 15 N(NO3 -) of atmospheric nitrate was depleted by (9.8 ± 5.1) ‰ relative to surface snow nitrate during spring. In summer, the enrichment was (9.1 ± 5.1) ‰, and  (Dibb et al., 2010;Fibiger et al., 2016;Dibb et al., 1998 and this study). This may imply potential contamination during sampling of the gas-phase HNO3, which remains to be explored and confirmed.

Discussion
Nevertheless, the enrichments of δ 15 N(NO3 -) in surface snow compared to 460 atmospheric nitrate and its seasonal difference (larger in the summer half year) also imply the effect of the photo-driven post-depositional processing. Erbland et al. (2013) also observed enriched δ 15 N(NO3 -) in surface snow nitrate compared to atmospheric nitrate at Dome C, Antarctica. At Dome C, the seasonal pattern of the surface snowatmosphere enrichments was similar to that at Summit, being the largest in Austral 465 spring (~ 30 ‰) and the smallest in Austral winter (~ 10 ‰). In addition, the enrichment at Dome C was observed throughout the year, and even in winter there was still a ~ 10 ‰ enrichment. The elevated enrichment of δ 15 N(NO3 -) in surface snow nitrate compared to atmospheric nitrate in spring/summer observed both at Summit and Dome C suggest the role of photolysis as proposed by Erbland et al. (2013). Compared to surface snow, 470 atmospheric nitrate is more influenced by snow-sourced nitrate which is severely depleted in δ 15 N (-60 to -100 ‰, Jiang et al. (2021)). In addition, surface snow nitrate has experienced photolysis which tends to increase its δ 15 N relative to the originally deposited nitrate. Winton et al. (2020) also suggested that at DML low snow accumulation rate and ample solar radiation tends to alter the original deposited nitrate 475 signal through photolysis even for the skin layers (defined as the upper most 0.5 cm snow). At Summit, although the snow accumulation rate is high compared to the East Antarctic plateau, unless frequent snowfall occurs to wash out atmospheric nitrate to refresh the surface snow δ 15 N(NO3 -), dry deposition of atmospheric nitrate is unable to influence the budget of nitrate in surface snow (1-3 cm) and disturb its δ 15 N(NO3 -) even 480 in a period of a few weeks (Jiang et al. 2021). This is because (i) snow is a much larger reservoir of nitrate compared to the atmosphere and (ii) the nitrate dry deposition flux https://doi.org/10.5194/tc-2021-355 Preprint. Discussion started: 10 December 2021 c Author(s) 2021. CC BY 4.0 License.

485
Different from Summit, around +10 ‰ enrichment in surface snow δ 15 N(NO3 -) compared to atmospheric nitrate exists at Dome C during winter in the absence of sunlight. Erbland et al. (2013) attributed this winter enrichment to nitrogen isotope fractionation during nitrate deposition which increases δ 15 N(NO3 -) in the deposited nitrate compared to the atmospheric pool, and suggested this also contributes to the 490 observed surface snow to atmospheric enrichment in spring/summer. However, the Summit data indicated no such enrichment in the winter, and this appears to be in conflict with the suggested deposition fractionation by Erbland et al. (2013). Although detailed physical mechanism leading to the deposition fractionation remains unknown, we speculated that the fractionation might be related to the form of deposition. Given 495 the large difference in snow accumulation rate at Summit (250 kg m -2 a -1 ) and Dome C (25 kg m -2 a -1 ), their main nitrate deposition mechanism might be quite different. At Dome C, nitrate concentration in the skin layer is mainly controlled by kinetic adsorption and co-condensation of atmospheric nitrate (Bock et al., 2016;Frey et al., 2009;Chan et al., 2018). While at Summit, the dominant mechanism for nitrate 500 incorporation into snow grain is the surface uptake during wet scavenging of atmospheric nitrate (Röthlisberger et al., 2002). Since wet deposition can efficiently scavenge atmospheric nitrate, a more complete removal of atmospheric nitrate at Summit compared to Dome C may occur, which would induce little to no isotope fractionation in δ 15 N due to mass balance. However, for surface snow that continues to 505 incorporate atmospheric nitrate via co-condensation or dry deposition (adsorption/desorption) after snowfall events, isotope fractionation could occur and leads to detectable enrichments in surface snow nitrate. The surface snow to atmospheric nitrate enrichments of δ 15 N(NO3 -) at Summit also appears to support the speculated role of fractionation during nitrate deposition. As shown in Figure 2b, the 510 maximum enrichments occurred in spring/summer, which was also the time with the lowest weekly average snow accumulation rate in a year (Burkhart et al., 2004;Jiang https://doi.org/10.5194/tc-2021-355 Preprint. Discussion started: 10 December 2021 c Author(s) 2021. CC BY 4.0 License. et al, 2021) and presumably more nitrate dry deposition occurred which leads to large isotope fractionation effect.
In summary, the systematic differences in δ 15 N(NO3 -) between atmospheric, 515 surface snow and snowpack samples are consistent with the expected effects of the photo-driven post-depositional processing, while the occurrence and mechanism(s) of nitrogen isotope fractionation during deposition and its contribution to the surface snow-atmospheric δ 15 N(NO3 -) enrichment need to be further explored and confirmed. In the following discussion, we focus on the processes occurring at the air-snow interface and in snow and their effects on Δ 17 O(NO3 -). Frey et al. (2009) proposed that nitrate in the uppermost layer of snow should 530 reach equilibrium with atmospheric nitrate to maintain consistent isotope ratios.
However, the large difference between atmospheric and surface snow δ 15 N(NO3 -) at Dome C Antarctica and Summit Greenland suggests no equilibrium. Conversely, an equilibrium in Δ 17 O(NO3 -) appears to exist. Erbland et al. (2013) made year-round observations of atmospheric nitrate and nitrate in the skin layer at Dome C, and found 535 that Δ 17 O(NO3 -) in the skin layer was similar to atmospheric Δ 17 O(NO3 -) except in spring when Δ 17 O(NO3 -) was ~ 5 ‰ higher than the former. This was explained by a reservoir effect by Erbland et al. (2013), as the surface snow is always a much larger reservoir for nitrate relative to the atmosphere, and there might be a delay in skin layer nitrate variations compared to the changes in atmospheric nitrate. low snow accumulation rate such as Dome C (Erbland et al., 2013;Frey et al., 2009), where nitrate stays in the photic zone for 4 to 5 years. In comparison, at Summit, the cage effect is negligible (< 0.3 ‰ upon archival, calculated by Jiang et al. (2021)) owing its fast archival (less than a half year) given the high snow accumulation rate, the archived snow nitrate should carry similar Δ 17 O signal to its deposited value at the 565 surface, which is in turn determined by atmospheric nitrate. Therefore, snowpack Δ 17 O(NO3 -) should be very similar to that of atmospheric nitrate, as is observed ( Figure   2c). However, this doesn't mean that snow nitrate Δ 17 O(NO3 -) can be directly linked to primary nitrate. Locally reformed nitrate under sunlight in the summer half year would possess low Δ 17 O compared to primary nitrate deposited earlier in the season (Kunasek 570 et al., 2009;Jiang et al., 2021) and contributes to the local atmospheric nitrate budget (Jiang et al., 2021). BrO, HO2, RO2, etc), and atmospheric water, as well as fractionations during formation (Michalski et al., 2012). Additionally, snow nitrate photolysis also directly influences δ 18 O with a fractionation factor calculated to be -34 ‰ by Frey et al. (2009), but does not affect Δ 17 O owing to its mass-independent nature. Some of these processes act to 585 enrich δ 18 O (e.g., photolysis) while others act to deplete δ 18 O (e.g., OH oxidation and/or exchange with water).
In general, under sunlight, nitrate formed from NO2 + OH reaction possesses lower 590 δ 18 O than that formed from N2O5 hydrolysis under dark conditions. The latter involves more oxygen atoms transferred from O3 which possesses very high δ 18 O (90-120 ‰, Johnston et al., 1997;Krankowsky et al., 1995). As a result, higher winter δ 18 O(NO3 -) and lower summer δ 18 O(NO3 -) should be expected, as observed for the atmospheric nitrate in this study and many others (Erbland et al., 2013;Savarino et al., 2007;Walters 595 et al., 2019). This is also why we noted that the δ 18 O(NO3 -) data in Jarvis et al. (2009) should be treated with caution as it indicated a summer maximum, which is difficult to understand given current knowledge. In the following discussion, we do not attempt to describe this discrepancy in Jarvis et al. (2009)  to Δ 17 O(NO3 -), there was also an (quasi) equilibrium in δ 18 O(NO3 -) between atmospheric and skin layer snow nitrate observed at Dome C, Antarctica (Erbland et al., 2013). Atmospheric gas-phase and surface snow nitrate δ 18 O(NO3 -) at Summit has been 605 reported by Jarvis et al. (2009) andFibiger et al. (2015) for spring and summer months.
While the Jarvis et al. (2009) study suggested that the surface snow nitrate δ 18 O(NO3 -) was on average 40 ‰ higher than atmospheric gas phase HNO3, the Fibiger et al. (2016) study found that surface snow δ 18 O(NO3 -) was lower than atmospheric nitrate in one year but higher in another. The atmospheric gas-phase nitrate δ 18 O(NO3 -) reported by 610 Jarvis et al. (2009) andFibiger et al. (2016) were also lower than the atmospheric δ 18 O(NO3 -) data reported by this study. The seasonal atmospheric and surface snow δ 18 O(NO3 -) data at Summit also didn't indicate an equilibrium. Overall, the proposed equilibrium between atmospheric and surface snow nitrate δ 18 O(NO3 -) is not supported by current observations.

615
Because of the lack of sufficient surface snow samples, and the inconsistency among the limited observations by Jarvis et al. (2009) andFibiger et al. (2015), we are unable to assess the potential oxygen isotope fractionation effects during nitrate deposition. But we note that this could also alter δ 18 O(NO3 -) in analogy with δ 15 N(NO3 -) and therefore this point needs to be further explored. After deposition, the post-620 depositional processing will impact the snow δ 18 O(NO3 -) in a similar matter as it impacts δ 15 N. The typical photolysis isotope fractionation factor ( 18 εp) for 18 O at Summit was calculated to be -32.8 ‰ using the ZPE shift method following Frey et al. (2009). Using the maximum loss fraction of 21% for spring snow from Jiang et al. (2021) and applying the Rayleigh equation, we calculated a maximum PIE of 7.7 ‰ for These simplified calculations suggest that there might be a difference in atmospheric δ 18 O(NO3 -) and snowpack δ 18 O(NO3 -) at Summit, but the magnitude and direction depend on the relative degrees of photolysis fractionation, the cage effect, and also other processes mentioned above (e.g., the kinetic isotope fractionation during 650 secondary nitrate formation).   where the photolysis of snow nitrate has been unambiguously shown to be dominant (Erbland et al., 2013). This emphasizes again that, when evaluating the degree of postdepositional processing, one should consider samples covering the all depth of the photic zone, not only surface samples.

Conclusion
In this study, we reported the first year-round atmospheric nitrate isotopes measurements for Summit, Greenland. The atmospheric δ 15 N(NO3 -) displayed systematic differences from surface snow and snowpack δ 15 N(NO3 -) values at Summit compiled from the literature. In general, atmospheric, surface snow, and snowpack 700 δ 15 N(NO3 -) diverged when there was sunlight but converged in the absence of sunlight.
The gradual enrichments in δ 15 N(NO3 -) from atmospheric nitrate to surface snow nitrate, and finally to snowpack nitrate can only be explained by the effect of the photo-driven post-depositional processing, and the enrichment after deposition can also be quantitatively explained by the photo-induced effect (PIE). We proposed a simplified 705 method for estimating PIE that can quickly assess the degree of δ 15 N(NO3 -) enrichment from the time of deposition to preservation in snow beneath the snow photic zone. In the end, we note the limitations of the compiled data. These data were collected by different groups at different time, and with different sampling methods as well as 725 different temporal resolutions. Although theoretically, the seasonality of the isotopes should be similar in different years or for samples collected and measured by different groups, and the heterogeneity of the samples was reduced by taking weighted average, there were some aspects and inconsistencies in the data that are difficult to interpret.
Simultaneous collection of atmospheric, surface snow and snowpack samples with 730 similar resolution for at least one complete year in the future should be conducted. This will provide a more consistent and solid dataset to improve or confirm the current understanding of nitrate preservation and isotope variations at Summit, Greenland. This is not only important for nitrate isotope record interpretation at this site, but also for other sites with similar or higher snow accumulation rate such as the WAIS (West 735 Antarctic Ice Sheet) Divide.
Data availability. The atmospheric aerosol nitrate isotope data and the compiled dataset will be provided upon direct request to the corresponding author.