Introduction
As the major sink of atmospheric nitrogen oxides (NOx= NO and
NO2), nitrate (NO3-) is one of the major chemical species
measured in polar snow and ice. The measurements of NO3- in ice
cores may offer potential for understanding the complex atmospheric nitrogen
cycle as well as oxidative capacity of the atmosphere through time
(Legrand and Mayewski, 1997; Alexander et al., 2004; Hastings et al.,
2009; Geng et al., 2017). However, the sources, transport pathways and
preservation of NO3- in Antarctic snowpack are still not well
understood, hampering the interpretation of ice core NO3- records.
The accumulation of NO3- in snow is associated with various
environmental factors and continental, tropospheric and stratospheric
sources could influence NO3- concentrations (Legrand and
Kirchner, 1990; McCabe et al., 2007; Wolff et al., 2008; Lee et al., 2014).
In surface snow, NO3- levels are thought to be linked with snow
accumulation rate, and higher values are usually present in areas with low
accumulation, e.g., East Antarctic plateaus (Qin et al., 1992; Erbland et
al., 2013; Traversi et al., 2017). Unlike sea-salt-related ions (e.g.,
chloride (Cl-), sodium (Na+) and occasionally sulfate
(SO42-)), NO3- does not usually show an elevated level
in coastal Antarctic snow (Mulvaney and Wolff, 1994; Bertler et al.,
2005; Frey et al., 2009), suggesting a negligible contribution from sea-salt
aerosols. However, the marine emissions of alkyl NO3-,
particularly methyl and ethyl NO3-, produced in surface oceans by
microbiological and/or photochemical processes, are thought to be a possible
contribution to Antarctic NO3- (Jones et
al., 1999; Liss et al., 2004). At Halley station in coastal Antarctica,
significant concentrations of organic nitrates (peroxyacetyl nitrate (PAN)
and alkyl NO3-) were observed in the lower atmosphere
(Jones et al., 2011). Organic nitrates dominated the NOy (sum of
reactive nitrogen oxide compounds) budget during the winter and were on par
with inorganic nitrate compounds during the summer. Although not a direct
source of snowpack nitrate, organic nitrates could act as source of NOx
to coastal Antarctica that would ultimately contribute to NO3-
within the snowpack (Jones et al., 2011).
While industrial and/or agricultural emissions have contributed to
increasing NO3- levels in Greenland snow and ice over recent
decades to hundreds of years, the anthropogenic contribution to Antarctic
NO3- is less clear (Mayewski and Legrand, 1990; Hastings et
al., 2009; Felix and Elliott, 2013; Geng et al., 2014). Lightning and
NOx produced in the lower stratosphere have long been thought to play a
major role (Legrand et al., 1989; Legrand and Kirchner,
1990). Recently, adjoint model simulations proposed that tropospheric
transport of NO3- from mid-to-low-latitude NOx sources is an
important source to the Antarctic year round, though less so in austral
spring and summer (Lee et al., 2014). Treatment of
NO3- in snow in the same global chemical transport model suggests
that the recycling of NO3- and/or transport of NOx due to
photolysis of NO3- in the surface snow layer is important in
determining summertime concentrations (Zatko et al.,
2016). The stratospheric inputs of NO3- are thought to result from
N2O oxidation to NO and then formation of NO3- that is deposited
via polar stratospheric cloud sedimentation (Legrand et al., 1989;
Legrand and Kirchner, 1990). The late winter–early spring secondary maximum
of NO3- observed in the atmosphere at coastal and inland locations
has been attributed to the stratospheric source based on the NO3-
stable isotopic composition (Legrand et al., 1989; Savarino et al., 2007;
Frey et al., 2009). At some sites, the snow and ice core NO3-
concentrations were found to be linked with regional atmospheric circulation
(e.g., sea level pressure gradient; Goodwin et al., 2003; Russell et al.,
2006). In general, atmospheric circulation appears to affect snow
NO3- concentrations indirectly through an influence
on the air mass transport and/or snow accumulation rate (Russell
et al., 2004, 2006). In addition, while some studies
suggested that snow/ice NO3- is possibly linked with
extraterrestrial fluxes of energetic particles and solar irradiation, with
solar flares corresponding to NO3- spikes (Zeller et al., 1986;
Smart et al., 2014), other observations and recent modeling studies have
established that there is not a clear connection between solar variability
and NO3- concentrations (Legrand et al., 1989; Legrand and
Kirchner, 1990; Wolff et al., 2008, 2012, 2016; Duderstadt et al.,
2014, 2016). However, the potential
link between the long-term (e.g., centennial to millennial timescales)
variability of NO3- and solar cycles may be present at some
locations (Traversi et al., 2012). In summary, factors influencing
NO3- levels in snow and ice are complicated, and the significance of
the relationship between NO3- and controlling factors varies
temporally and spatially.
Gas-phase and snow concentration studies and recent isotopic investigations
and modeling have shown that NO3-, particularly in snow on the
Antarctic plateau, is a combination of deposition of HNO3 and
post-depositional loss or recycling of NO3-
(e.g., Röthlisberger et al., 2002; Davis et al., 2004; Dibb et al.,
2004; Erbland et al., 2013, 2015; Shi et al., 2015; Bock et
al., 2016; Zatko et al., 2016). Based upon a suite of isotopic studies in
the field and laboratory, it has been demonstrated that under cold, sunlit
conditions ultraviolet photolysis dominates NO3- post-depositional
processing, whereas HNO3 volatilization may become more important at
warmer temperatures > -20 ∘C (Röthlisberger et al.,
2002; Frey et al., 2009; Erbland et al., 2013; Berhanu et al., 2015). In
snowpack, the solar radiation decreases exponentially, with attenuation
described in terms of an e-folding depth (ze) where the actinic flux is
reduced to 37 % (i.e., 1/e) of the surface value. Thus, about 95 % of
snowpack photochemistry is expected to occur above the depth of three times
ze (Warren et al., 2006). Field measurements at Dome C on the
East Antarctic plateau suggest a ze of 10 to 20 cm (France et
al., 2011), and the depth is dependent upon the concentration of impurities
contained in the snow (Zatko et al., 2013). In the inland regions
with low snow accumulation rates, particularly on the East Antarctic
plateaus, photolysis has been shown to lead to significant post-depositional
loss of NO3-, demonstrated by significant enrichment in 15N
of snow NO3- (i.e., high δ15N) (Frey et al., 2009;
Erbland et al., 2013, 2015; Berhanu et al., 2015; Shi et
al., 2015), as well as a decrease in δ18O and Δ17O
due to reformation of NO3- in the condensed phase (Erbland et al.,
2013; Shi et al., 2015, and references therein). The transport and recycling
of NOx sourced from photolysis of snow NO3- in the summertime
has been invoked to model the distribution of snowpack NO3- across
the Antarctic plateau (Zatko et al., 2016). However, snow physical
characteristics play a crucial role in NO3- deposition and
preservation. For instance, summertime concentrations in the surface skin
layer of snow (the uppermost ∼ 4 mm) can be explained as the
result of co-condensation of HNO3 and water vapor, with little to no
photolytic loss in this microlayer (Bock et al., 2016). The combination of
concentration and isotopic studies, along with physical aspects of the snow,
could lead to the reconstruction and interpretation of atmospheric
NO3- over time (e.g., Erbland et al., 2015; Bock et al., 2016),
if there were detailed understanding of the NO3- deposition and
preservation in different environments in Antarctica.
The effects of volatilization of NO3- are uncertain, given that
one field experiment suggests that this process is an active player in
NO3- loss (17 % (-30 ∘C) to 67 % (-10 ∘C) of
NO3- lost after 2 weeks' physical release experiments;
Erbland et al., 2013), while other laboratory and field studies show that
volatilization plays a negligible role in NO3- loss (Berhanu et
al., 2014, 2015). Further investigations are needed to
quantify the effects of volatilization for a better understanding of
NO3- preservation in snow and ice. Based on ze, NO3- at
deeper depths in Antarctic snow (e.g., > 100 cm), well beyond the
snow photic zone, may be taken as the archived fraction. Thus,
NO3- in deeper snow provides an opportunity to
investigate the archived fraction and potential influencing factors (e.g.,
snow accumulation rate). Given that an extensive array of ice core
measurements is unavailable in most of Antarctica, the deeper snow pits (with
depth > 100 cm) may offer a useful way to investigate the
archived NO3-.
In the atmosphere in Antarctica, particularly during spring and summer,
NO3- is found to be mainly in the form of gas-phase HNO3,
with NO3- concentration several times higher in gas phase than in
the particulate phase (Piel et al., 2006; Legrand et al., 2017b; Traversi
et al., 2017). During post-depositional processes, the uptake of gaseous
HNO3 is thought to be important in NO3- concentrations in
surface snow layers (Udisti et al., 2004; Traversi et al., 2014, 2017). Due to the high concentration in summer, HNO3 appears to
play an important role in acidifying sea-salt particles, possibly accounting
for the presence of NO3- in the particulate phase in summer
(Jourdain and Legrand, 2002; Legrand et al., 2017b; Traversi et al.,
2017). It is noted that the significant increase of NO3- during
the cold periods (e.g., Last Glacial Maximum) could be associated with its
attachment to dust aerosol instead of formation of gas-phase HNO3
(Legrand et al., 1999; Wolff et al., 2010).
To date, investigations on spatial and temporal patterns of snow
NO3- have been performed on several traverses in Antarctica (e.g.,
1990 International Trans-Antarctica Expedition and DDU to Dome C; Qin et
al., 1992; Bertler et al., 2005; Frey et al., 2009; Erbland et al., 2013;
Pasteris et al., 2014), but these provide an uneven distribution of snow
NO3- concentrations and leave large regions un-sampled (e.g.,
Lambert Glacier basin and Dome A plateau). Over the past few decades, while
several glaciological observations have been carried out on the Chinese
inland Antarctic traverse route from Zhongshan to Dome A, East Antarctica
(Hou et al., 2007; Ding et al., 2010, 2011; Ma et al., 2010;
Li et al., 2013; Shi et al., 2015), the data on snow chemistry are still
rare, particularly detailed information on NO3-. From 2009 to
2013, we therefore conducted surface snow and snow pit sampling campaigns
along the traverse route, with the main objectives to (1) describe
NO3- distribution in surface snow and snow pits, (2) characterize
the relationship between archived NO3- and snow accumulation rate,
and (3) examine the potential effects of coexisting ions on NO3-
preservation. The results of this study may help to better understand
NO3- deposition and preservation in the snowpack, which is
critical to the interpretation of ice core NO3- records.
Snow pit information on the traverse from coastal Zhongshan Station
to Dome A, East Antarctica.
Snow pit
Latitude,
Longitude,
Elevation,
Distance to
Annual snow
Depth,
Sampling
Sampling
no.
∘
∘
m
coast, km
accumulation,
cm
resolution,
year
kg m-2 a-11
cm
SP1
-70.52
76.83
1613
132
193.2
150
5.0
2010/2011
SP2
-71.13
77.31
2037
200
172.0
150
3.0
2012/2013
SP3
-71.81
77.89
2295
283
99.4
200
5.0
2012/2013
SP4
-72.73
77.45
2489
387
98.3
200
5.0
2012/2013
SP5
-73.40
77.00
2545
452
90.7
200
5.0
2012/2013
SP6
-73.86
76.98
2627
514
24.6
300
2.5
2012/2013
SP7
-74.50
77.03
2696
585
29.2
100
2.0
2012/2013
SP8
-74.65
77.01
2734
602
80.2
180
2.0
2010/2011
SP9
-76.29
77.03
2843
787
54.8
200
2.0
2012/2013
SP10
-76.54
77.02
2815
810
100.7
240
3.0
2010/2011
SP11
-77.13
76.98
2928
879
81.2
200
2.5
2012/2013
SP12
-77.26
76.96
2962
893
83.4
265
5.0
2009/2010
SP13
-77.91
77.13
3154
968
33.3
200
2.0
2012/2013
SP14
-78.34
77.00
3368
1015
87.6
216
3.0
2010/2011
SP15
-78.35
77.00
3366
1017
70.0
162
2.0
2009/2010
SP16
-79.02
76.98
3738
1092
25.4
200
2.5
2012/2013
SP17
-79.65
77.21
3969
1162
46.2
130
2.0
2010/2011
SP18
-80.40
77.15
4093
1250
24.2
300
2.0
2010/2011
SP19
-80.41
77.11
4092
1254
23.7
300
1.0
2009/2010
SP20
-80.42
77.12
4093
1256
23.5
300
2.5
2012/2013
Core 12
-70.83
77.08
1850
168
127.0
–
–
1996/1997
Core 23
-76.53
77.03
2814
813
101.0
–
–
1998/1999
1 Annual snow accumulation rate is obtained from the field bamboo stick
measurements (2009–2013), updated from the report (Ding et al.,
2011). Note that snow accumulation rate at the two ice core sites are
derived from ice core measurements.
2 Core 1, ice core data of previous report (Li et al., 1999; Xiao et
al., 2004).
3 Core 2, ice core data of previous report (Li et al., 2009).
Methodology
Study area (Zhongshan to Dome A traverse)
The Zhongshan to Dome A CHINARE (Chinese National Antarctic Research
Expedition) inland traverse is an important leg of the ITASE (International
Trans-Antarctic Scientific Expedition). The traverse is in the Indian Ocean
sector of East Antarctica, passing through the Lambert Glacier, the largest
glacier in Antarctica. In January 1997 the first Chinese Antarctic inland
expedition reached an area ∼ 300 km from the coast; in January
1998 the traverse was extended to 464 km, and in December 1998 to the Dome
A area ∼ 1100 km from the coast. In the austral 2004/2005
summer, for the first time, the traverse extended to the ice sheet summit,
Dome A, a total distance of ∼ 1260 km. In January 2009, the
Chinese inland research base, Kunlun station (80∘25′1.7′′ S and
77∘6′58.0′′ E; 4087 m above mean sea level), was established
at Dome A, mainly aimed at deep ice core drilling and astronomical
observations. Now, Kunlun base is a summer station, and the CHINARE team
typically conducts an annual inland traverse from the coastal Zhongshan
station to Dome A.
In January 2010, the Dome A deep ice core project was started, and the
construction of basic infrastructure (including drill trench and scientific
workroom) took four summer seasons. The deep ice core drilling began in January
2013, and in total 801 m ice core was recovered by the 2016/2017 season. The
investigation of NO3- deposition and preservation in the snowpack
will be of help to the interpretation of Dome A deep ice core NO3-
records.
Sample collection
During the 2010/2011 CHINARE, surface snow samples (the topmost
∼ 3 cm) were collected at an interval of ∼ 10 km
along the traverse route from Zhongshan to Dome A, using 3.0 cm diameter
high-density polyethylene (HDPE) bottles (volume of 100 mL). The bottles
were pre-cleaned with Milli-Q ultrapure water (18.2 MΩ), until
electrical conductivity of the water stored in bottles (> 24 h)
decreased to < 0.5 µS cm-1. Then, the bottles were dried
under a class 100 super-clean hood at 20 ∘C. Immediately after the
drying procedure, the bottles were sealed in clean PE bags that were not
opened until the field sampling started. At each sampling site (typically
> 500 m away from the traverse route), the bottles were pushed
into surface snow layers in the windward direction. In total, 120 surface
snow samples were collected. In addition, at each sampling site, the upper
snow density (∼ 10 cm) was measured using a density scoop with
a volume of 1000 cm3. As the field blanks, pre-cleaned bottles filled
with Milli-Q water were taken to the field and treated to the same
conditions as field samples (n=3).
On the Dome A plateau, the snow is soft and non-cohesive, and morphology of
the surface snow is different from other areas on the traverse, with a
crystal ice layer extensively developed, in particular on the
sastrugi (Fig. S1 in the Supplement). The depth of the needle-like
crystal ice layer (referred to as “crystal ice” in the following context)
is generally < 1.0 cm. In order to investigate air–snow transfer of
NO3- in this uppermost ∼ 1 cm layer, the crystal ice
was collected using a clean HDPE scoop and then poured into clean, wide-mouth HDPE bottles. Approximately 30 g of crystal ice was collected for each
sample. In total, six crystal ice samples were collected on the traverse near
Dome A plateau.
In addition to surface snow, snow pit samples were collected during CHINARE
inland traverse campaigns in 2009/2010, 2010/2011 and 2012/2013. The
snow pits were excavated manually, and the snow wall in the windward
direction was scraped clean and flat with a clean HDPE scraper. Then the
bottles were pushed horizontally into the snow wall. Snow pit samples were
collected from the base towards the top layer along a vertical line. During
the sampling process, all personnel wore PE gloves and face masks to minimize
potential contamination. Note that the snow pits are generally
> 1 km from the traverse route to avoid possible contamination from the
expedition activities. All information about individual snow pits,
including location, distance from the coast, elevation, snow pit depth,
sampling resolution, collection date and annual snow accumulation rate, is
summarized in Table 1. All together, 20 snow pits (SP1 to SP20 in Fig. 2,
with SP20 corresponding to the location of Kunlun station at Dome A) as 1741
snow samples were collected.
To support understanding of the air–snow transfer of NO3- on the
traverse, atmospheric NO3- was collected on glass fiber filters
(Whatman G653) using a high-volume air sampler (HVAS), with a flow rate of
∼ 1.0 m3 min-1 for 12–15 h, during the inland
traverse campaign in 2015/2016. The NO3- collected on glass fiber
filters is expected to equal the sum of particulate NO3- and
gaseous HNO3, based upon previous investigations in East Antarctica
(Savarino et al., 2007; Frey et al., 2009; Erbland et al., 2013). In
total, 34 atmospheric samples were collected on the traverse. In addition,
two field blanks were collected from filters installed in the HVAS without
pumping and treated as samples thereafter. Detailed information about the
atmospheric sampling is presented in Table S1 in the Supplement.
After sample collection, all filters and snow samples were sealed in clean
PE bags and preserved in clean thermal insulated boxes. All of the samples
were transported to the laboratory under freezing conditions
(< -20 ∘C).
Sample analysis
In the laboratory, three quarters of individual filters were cut into pieces
using pre-cleaned scissors that were rinsed between samples, placed in
∼ 100 mL of Milli-Q water, ultrasonicated for 40 min and
leached for 24 h under shaking. The sample solutions were then filtered
through 0.22 µm ANPEL PTFE filters for concentration analysis. Snow
samples were melted in the closed sampling bottles on a super-clean bench
(class 100) before chemical measurements. Analyses of Na+,
NH4+, K+, Mg2+, Ca2+, Cl-, NO3- and SO42- were performed using a Dionex ICS-3000 ion
chromatography system. The column used for cation analysis (Na+,
NH4+, K+, Mg2+ and Ca2+) was a Dionex column CS12
(2×250 mm), with a guard column CG12 (2×50 mm), while
the anions (Cl-, NO3- and SO42-) were analyzed
using a Dionex column AS11 (2×250 mm) with a guard column AG11
(2×50 mm). The eluent for cations was 18.0 mM methanesulfonic acid, and the gradient elution method was employed for anion analysis, with
eluent of potassium hydroxide. More details on this method are
described in a previous report (Shi et al., 2012). During sample
analysis, duplicated samples and field blanks were synchronously analyzed.
The pooled standard deviation (σp; σp=∑i=1kni-1si2/∑i=1kni-1, where ni and si2 are the size and
variance of the ith samples, respectively, and k is the total number of sample
sets) of all replicate samples run at least twice in two different sample
sets is 0.019 (Cl-), 0.023 (NO3-), 0.037 (SO42-),
0.022 (Na+), 0.039 (NH4+), 0.006 (K+), 0.006 (Mg2+)
and 0.006 (Ca2+) µeq L-1, respectively (n=65 pairs of
samples). Ion concentrations in field blanks (n=3) are generally lower
than the detection limit (3 standard deviations of water blank in the
laboratory).
For Antarctic snow samples, the concentrations of H+ are usually not
measured directly, but deduced from the ion-balance disequilibrium in the
snow. Here, H+ concentration is calculated through ion
balance.
[H+]=[Cl-]+[NO3-]+[SO42-]-[Na+]-[NH4+]-[K+]-[Mg2+]-[Ca2+],
where ion concentrations are in µeq L-1. In addition, the non-sea-salt fractions of SO42- (nssSO42-) and Cl-
(nssCl-) can be calculated from the following expressions, by assuming
Na+ exclusively from sea salt (in µeq L-1).
[nssSO42-]=[SO42-]-0.12×[Na+][nssCl-]=[Cl-]-1.17×[Na+]
It is noted that SO42- fractionation (the precipitation of
mirabilite, Na2SO4 ⋅ 10H2O) may introduce a bias in
nssSO42-, particularly during the winter half year
(Wagenbach et al., 1998a).
Results
NO3- concentration in surface snow
Concentrations of NO3- in surface snow are shown in Fig. 1,
ranging from 0.6 to 5.1 µeq L-1, with a mean of 2.4 µeq L-1.
One standard deviation (1σ) of NO3- concentration in surface snow is 1.1 µeq L-1, with coefficient
of variation (Cv, 1σ over mean) of 0.5, indicating a moderate spatial
variability. About 450 km from the coast, NO3- shows a
slightly increasing trend towards the interior, with low variability, while
NO3- concentrations are higher in the inland region, with a large
fluctuation. It is notable that in the area ∼ 800 km from the
coast, where snow accumulation is relatively high, NO3- concentrations decrease to < 2.0 µeq L-1, comparable to
the values on the coast. Near the Dome A plateau (> 1000 km from
coast), there is a tendency for higher NO3- concentrations
(> 5.0 µeq L-1). Similarly, atmospheric NO3-
concentrations increase from the coast towards the plateau, ranging from 6
to 118 ng m-3 (mean of 38 ng m-3) (Fig. 1).
Concentrations of NO3- in snow (surface snow, crystal
ice and snow pits; on the primary y axis) and atmosphere (on the secondary
y axis), with error bars representing 1 standard deviation of
NO3- (1σ) for individual snow pits. Also shown is the
annual snow accumulation rate on the traverse (red solid line; based on Ding
et al., 2011). Note that NO3- concentration in one crystal ice
sample (red dot) is higher than the maximum value of the primary y axis
(NO3- concentration = 16.7 µeq L-1 in the
parentheses).
Concentrations of NO3- in surface snow across
Antarctica. Note that the values of crystal ice around Dome A were not
included. The data of DDU to Dome C are from Frey et al. (2009). The other
surface snow NO3- concentrations are from compiled data
(Bertler et al., 2005, and references therein). Also illustrated are
the locations of snow pits on the traverse route from Zhongshan to Dome A in
this study (SP1 to SP20, solid short blue line; Table 1).
The full profiles of NO3- concentrations for snow pits
collected on the traverse from the coast to Dome A, East Antarctica (SP1 is
closest the coast; SP20 the furthest inland; see Fig. 2). The details on
sampling of the snow pits refer to Table 1. The numbers in parentheses in
each panel denote the annual snow accumulation rates (kg m-2 a-1).
Note that the scales of x axes for the snow pits SP1–SP9 and SP10–SP20
are different.
The percentage that surface snow NO3- contributes to total ions
(i.e., total ionic strength, sum of Na+, NH4+, K+,
Mg2+, Ca2+, Cl-, NO3- , SO42- and
H+, in µeq L-1) varies from 6.7 to 37.6 % (mean of 27.0 %; Fig. S2 in the Supplement), with low values near the coast and
high percentages on the plateau. A strong relationship was found between
NO3- and the total ionic strength in surface snow (R2=0.55,
p<0.01).
In the crystal ice, the means (ranges) of Cl-, NO3-,
SO42-, Na+, NH4+, K+, Mg2+, Ca2+ and
H+ concentrations are 0.98 (0.62–1.27), 10.40 (8.35–16.06), 1.29
(0.87–2.13), 0.27 (0.21–0.33), 0.24 (0.03–0.56), 0.05 (0.03–0.08), 0.18 (0.15–0.22), 0.18 (0.05–0.57) and 11.75 (9.56–18.12)
µeq L-1, respectively. H+ and NO3- are the most
abundant species, accounting for 46.4 and 41.0 % of the total ions,
followed by SO42- (5.1 %) and Cl- (3.9 %). The other
five
cations, Na+, NH4+, K+, Mg2+ and Ca2+, only
represent 3.6 % of the total ion budget. A significant linear
relationship was found between NO3- and the total ionic strength
(R2=0.99, p<0.01), possibly suggesting that NO3- is
the species controlling ion abundance by influencing acidity of the crystal
ice (i.e., H+ levels). In comparison with surface snow, concentrations
of H+ and NO3- are significantly higher in crystal ice
(independent samples t test, p<0.01), while concentrations of
Cl-, SO42-, Na+, NH4+, K+, Mg2+ and
Ca2+ are comparable in the two types of snow samples (Fig. S2 in
the Supplement). To date, the information on the chemistry of ice
crystal is rather limited but data from the so-called skin layer at Dome C
(top ∼ 4 mm snow), where NO3- concentrations are
in the range of 9–22 µeq L-1 in summertime (Erbland et
al., 2013), are generally comparable to our observations.
NO3- concentrations in surface snow have been widely measured
across Antarctica (Fig. 2), and the values vary from 0.2 to 12.9 µeq L-1, with a mean of 2.1 µeq L-1 (n= 594, 1σ=1.7 µeq L-1) and a median of 1.4 µeq L-1. Most of the data
(87 %) fall in the range of 0.5–4.0 µeq L-1, and only 7 %
of the values are above 5.0 µeq L-1, mainly distributed on the
East Antarctic plateaus. Spatially, NO3- concentrations show an
increasing trend with distance inland, and the values are higher in East
than in West Antarctica. Overall, this spatial pattern is opposite to that
of the annual snow accumulation rate (Arthern et al., 2006), i.e.,
low (high) snow accumulation corresponds to high (low) NO3-
concentrations. It is difficult to compare with NO3-
concentrations derived from the “upper snow layer” in different studies
because each study sampled a different depth (Fig. 2), e.g., 2–10 cm for
DDU-Dome C traverse (Frey et al., 2009; Erbland et al., 2013), 25 cm for
the 1989–1990 International Trans-Antarctica Expedition (Qin et al.,
1992) and 3 cm for this study. The different sampling depths can result in
large differences in NO3- concentration, especially on the East
Antarctic plateaus (e.g., the values of the topmost 1 cm of snow, the
crystal ice in this study, can be up to > 15 µeq L-1;
Fig. 1). Because of this, any comparison of NO3- concentrations in
surface snow collected in different campaigns should be made with caution.
Snow pit NO3- concentrations
Mean NO3- concentrations for snow pits are shown in Fig. 1. From
the coast to ∼ 450 km inland, snow pit NO3- means are
comparable to those of surface snow, whereas NO3- means are lower
in inland snow pits than in surface snow with the exception of sites
∼ 800 km from the coast. In general, the differences between
snow pit NO3- means and the corresponding surface snow values are
small at sites with high snow accumulation (e.g., close to coast), while the
differences are large in low snow accumulation areas (e.g., near Dome A).
The profiles of NO3- for all snow pits are shown in Fig. 3.
NO3- concentrations vary remarkably with depth in pits SP1–SP5,
which are located near the coast. Although SP2 and SP5 show high
NO3- concentrations in the topmost sample, the data from deeper
depths can be compared with the surface values. In addition, NO3-
means for the entire snow pits are close to the means of the topmost layer
covering a full annual cycle of accumulation (i.e., the most recent year of
snow accumulation) at SP1–SP5 (Fig. 4). Given the high snow accumulation
(Fig. 1), NO3- variability in coastal snow pits is likely
suggestive of a seasonal signature (Wagenbach et al., 1998b; Grannas et
al., 2007; Shi et al., 2015). Among the coastal snow pits, water isotope
ratios (δ18O of H2O) of samples at SP2 were also
determined, thus allowing for investigating NO3- seasonal
variability (Fig. S3 in the Supplement). In general, the δ18O(H2O) peaks correspond to high NO3- concentrations
(i.e., NO3- peaks present in summer). This seasonal pattern is in
agreement with previous observations of NO3- in snow/ice and
atmosphere in coastal Antarctica (Mulvaney and Wolff, 1993; Mulvaney et
al., 1998; Wagenbach et al., 1998b; Savarino et al., 2007).
In contrast, most of the inland snow pits show high NO3-
concentrations in the top layer and then fall sharply from
> 2.0 µeq L-1 in top snow to < 0.2 µeq L-1 in the
first meter of depth (Fig. 3). NO3- means for the entire
snow pits are typically lower than those of the most recent year snow layer
(Fig. 4). Similar NO3- profiles for snow pits have been reported
elsewhere in Antarctica, as a result of post-depositional processing of
NO3- (Röthlisberger et al., 2000; McCabe et al., 2007;
Erbland et al., 2013; Shi et al., 2015).
Comparison of the NO3- profile patterns reveals significant
spatial heterogeneity, even for neighboring sites. For instance, sites SP11
and SP12, 14 km apart, feature similar snow accumulation rate (Table 1). If
it is assumed that snow accumulation is relatively constant during the past
several years at SP11 (sampled in 2012/2013), snow in the depth of
∼ 54 cm corresponds to the deposition in 2009/2010 (snow
density = 0.45 g cm-3, from field measurements). NO3-
concentrations are much higher in the top snow of SP12 (sampled in
2009/2010) than in the depth of ∼ 54 cm in SP11 (Fig. 3). This
variation in NO3- profiles on a local scale has been reported,
possibly related to local morphologies associated with sastrugi formation
and wind drift (Frey et al., 2009; Traversi et al., 2009).
It is interesting that higher NO3- concentrations were not found
in the uppermost layer at sites SP7 and SP8 (∼ 600 km from
coast; Fig. 3), where large sastrugi with hard smooth surfaces had
extensively developed (from field observations; Fig. S4 in supporting
information). Snow accumulation rate in this area fluctuates remarkably, and
the values of some sites are rather small or close to zero due to the strong
wind scouring (Fig. 1) (Ding et al., 2011; Das et al., 2013). In this
case, the snow pit NO3- profiles appear to be largely influenced by
wind scour on snow, possibly resulting in missing years and/or intra-annual
mixing.
Discussion
Accumulation influence on NO3-
The preservation of NO3- is thought to be closely associated with
snow accumulation, where most of the deposited NO3- is preserved
at sites with higher snow accumulation (Wagenbach et al., 1994; Hastings
et al., 2004; Fibiger et al., 2013), whereas NO3- may be altered
significantly at sites with low snow accumulation, largely due to photolysis
(Blunier et al., 2005; Grannas et al., 2007; Frey et al., 2009; Erbland
et al., 2013, 2015). In the following discussion, we divide
the traverse into two zones, i.e., the coastal zone (< ∼ 450 km from the coast, including SP1–SP5 and Core 1;
Table 1) and the inland region (∼ 450 km to Dome A, including pits
SP6-SP20 and Core 2; Table 1), following NO3- distribution
patterns in surface snow and snow pits (Sect. 3.1 and 3.2) as well as the
spatial pattern of snow accumulation rate (Fig. 1).
As for snow pits, NO3- levels in top and deeper layers are
comparable near the coast, while NO3- differs considerably between
the upper and deeper snow at inland sites (Figs. 3 and 4). Photochemical
processing is responsible for NO3- distribution in inland snow pits
(Erbland et al., 2013; Berhanu et al., 2015). Considering that the
actinic flux is always negligible below the depth of 1 m, the bottom layers
of the snow pits (i.e., > 100 cm; Table 1) are well below the
photochemically active zone (France et al., 2011; Zatko et al., 2013). In
this case, NO3- in the bottom snow pit, i.e., below the photic
zone, can be taken as the archived fraction without further modification, as
also suggested by previous observations (Frey et al., 2009; Erbland et
al., 2013, 2015). Here, we define NO3- in the
bottom layer covering a full annual cycle of deposition as an approximation
of the annual mean of archived NO3- (i.e., beyond photochemical
processing; denoted as “Carchived” in the following context; Fig. 4),
thus allowing for calculating the archived annual NO3- flux (i.e.,
the product of Carchived and annual snow accumulation rate). Although
there is uncertainty in the calculation of archived NO3- flux due
to interannual variability in NO3- inputs and snow accumulation,
this assumption provides a useful way to investigate the relationship
between preservation of NO3- and physical factors considering that
an extensive array of ice core measurements are unavailable in most of
Antarctica. It is noted that Carchived is generally close to (lower
than) the NO3- means for entire snow pits in coastal (inland)
Antarctica (Fig. 4).
Mean concentrations of NO3- for the entire snow pit depth
(in square), the uppermost layer covering one-year snow accumulation (in
diamond) and the bottom layer covering a full annual cycle of deposition
(archived NO3- concentration, Carchived, in triangle).
NO3- in coastal snowpack
The simplest plausible model to relate flux and concentration of
NO3- in snow to its atmospheric concentration (Legrand,
1987; Alley et al., 1995) can be expressed as
Ftotal=K1Catm+K2CatmA,Ftotal=Cfirn×A,
where Ftotal is snow NO3- flux (µeq m-2 a-1);
Catm is atmospheric concentration of NO3- (µeq m-3); A is annual snow accumulation rate (kg m-2 a-1);
Cfirn is measured firn NO3- concentration (µeq L-1,
here Cfirn=Carchived); K1 is the dry deposition velocity
(cm s-1); and K2 is the scavenging ratio for precipitation
(m3 kg-1), which allows conversion of atmospheric concentration to snow
concentration of NO3- in this study. From Eqs. (4) and (5), firn
NO3- concentration can be expressed as
Cfirn=K1Catm×1/A+K2Catm.
If K1 and K2 are constants, a linear relationship between measured
NO3- concentration (Cfirn) and snow accumulation (A) can be
interpreted using Eq. (6), which assumes regional spatial homogeneity of fresh
snow NO3- levels and dry deposition flux. The slope
(K1Catm) of the linear model represents an approximation of dry
deposition flux of NO3- (i.e., an apparent dry deposition flux),
while the intercept (K2Catm) stands for NO3- concentration
in fresh snowfall. If dry deposition (K1Catm) is much larger than
wet deposition (K2CatmA), the concentration of NO3- in snow
will be proportional to its concentration in the atmosphere. In the
condition of a constant atmospheric concentration, larger snow accumulation
will increase the flux of NO3- but decrease its concentration in
snow. While this linear model is a gross oversimplification of the complex
nature of air–snow exchange of NO3-, it provides a simple approach
to compare the processes occurring on the coast versus those inland. In
addition, this model can provide useful parameter values in modeling
NO3- deposition and preservation on large scales, considering that
observations remain sparse across Antarctica (e.g., Zatko et al., 2016).
The relationship between Carchived of NO3- and snow
accumulation rate is shown in Fig. 5. The linear fit of Carchived vs.
inverse snow accumulation (R2=0.88, p<0.01; Fig. 5a) supports
the assumptions of spatial homogeneity. The intercept and slope of the
linear fit suggest a NO3- concentration in fresh snow and an
apparent NO3- dry deposition flux of 0.7±0.07 µeq L-1 and 45.7±7.8 µeq m-2 a-1, respectively. The
apparent dry deposition flux is opposite to the observation in Dronning Maud
Land (DML) region, where a negative dry deposition flux suggested net losses
of NO3- (Pasteris et al., 2014).
The relationships amongst snow accumulation rate, the archived
concentration (Carchived), and flux of NO3- in coastal (a, b and c)
and inland (d, e and f) Antarctica.
In panel (d), the linear fit in black line (y=-44.5x+2.1) includes
the full data set, while the linear equation in red (y=-27.7x+1.5)
was obtained by excluding two cases (open circles) with snow accumulation
rate larger than 100 kg m-2 a-1 (see the main text). The flux
values are the product of Carchived of NO3- and snow
accumulation rate, namely the archived flux. Least-squares regressions are
noted with solid lines and are significant at p<0.01. Error bars
represent 1 standard deviation (1σ).
Figure 5b shows the archived fluxes of NO3- on the coast, with
values from 104 (at the lowest accumulation site) to 169 µeq m-2 a-1 (at the highest accumulation site). Taking the calculated
NO3- dry deposition flux of 45.7 µeq m-2 a-1, dry
deposition accounts for 27–44 % (mean of 36 %) of total
NO3- inputs, with higher (lower) percentages at lower (higher)
snow accumulation sites. This result is in line with the observations in
Taylor Valley (coastal West Antarctica), where the snowfall was found to be
the primary driver for NO3- inputs (Witherow et al., 2006).
This observation also generally agrees with, but is greater than, that in the
modeling study of Zatko et al. (2016), which predicts a
ratio of dry deposition to total deposition of NO3- in Antarctica
as < 20 % close to the coast, increasing towards the plateaus.
In Fig. 5a and b, the strong linear relationships between NO3-
and snow accumulation support that K1 and K2 are relatively constant
on the coast (Eqs. 4 and 6). The average atmospheric concentration of
NO3- in the coastal ∼ 450 km region is 15.6 ng m-3 in summer (Table S1 in the Supplement). Taking
Catm=15.6 ng m-3, K1 is estimated to be 0.6 cm s-1,
close to a typical estimate for HNO3 deposition velocity to a snow/ice
surface (0.5 cm s-1; Seinfeld and Pandis, 1997). This predicted
K1 value is lower than that calculated for the dry deposition of
HNO3 at South Pole (0.8 cm s-1; Huey et al., 2004). It is noted
that the true K1 value could be larger than the prediction (0.6 cm s-1)
due to the higher values of Catm in summer (i.e., 15.6 ng m-3 for the calculation of K1) than in other seasons (Mulvaney
et al., 1998; Wagenbach et al., 1998b; Savarino et al., 2007). The
scavenging ratio for precipitation (K2) is estimated to be about
0.2×104 m3 kg-1, i.e., 2 m3 g-1.
If it is assumed that NO3- concentration in snow is related to its
concentration in the atmosphere, the scavenging ratio for NO3-
(W) can be calculated on a mass basis from the following expression
(Kasper-Giebl et al., 1999):
W=ρatm×(Cf-snow /Catm),
where ρatm is air density (g m-3), and Cf-snow and
Catm are NO3- concentrations in fresh snow (ng g-1) and
atmosphere (ng m-3), respectively. If taking ρatm≈1000 g m-3 (on average, ground surface temperature t≈255 K,
ground pressure P≈0.08 MPa, in the coastal region),Cf-snow=43 ng g-1 (see discussion above and Sect. 4.2 below) and
Catm=15.6 ng m-3, W is calculated to be ∼ 2700,
generally comparable to previous reports (Barrie, 1985; Kasper-Giebl et
al., 1999; Shrestha et al., 2002). It is noted that the calculation here may
be subject to uncertainty due to the complex transfer of atmospheric
NO3- into the snow. However, the scavenging ratio provides
valuable insights into the relation between NO3- concentrations in
the atmosphere and snow, which might be useful in modeling NO3-
deposition on a large scale.
Figure 5c shows the distribution of flux is negatively correlated with
Carchived of NO3-, which is not surprising since
Carchived is positively related to inverse accumulation (Fig. 5a). Based
on the observed strong linear relationship between NO3- flux and
snow accumulation (Fig. 5b), the archived NO3- flux is more
accumulation-dependent compared to Carchived. This is compatible with
the observations in Greenland (Burkhart et al., 2009), where
accumulation is generally above 100 kg m-2 a-1, similar to the
coastal values in this study.
The relationships between NO3- concentration and inverse
snow accumulation rate in surface snow in coastal (a) and inland
(b) Antarctica. Least-squares regressions are noted with solid line
and are significant at p<0.01.
In terms of surface snow on the coast, NO3- may be disturbed by
the katabatic winds and wind convergence located near the Amery Ice Shelf
(that is, the snow-sourced NOx and NO3- from the Antarctic
plateau possibly contributes to coastal snow NO3-) (Parish and
Bromwich, 2007; Ma et al., 2010; Zatko et al., 2016). In addition, the
sampled ∼ 3 cm surface layer roughly corresponds to the net
accumulation in the past 0.5–1.5 months, assuming an even distribution of
snow accumulation in the course of a single year. This difference in
exposure time of the surface snow at different sampling sites could
possibly affect the concentration of NO3-, although the
post-depositional alteration of NO3- was thought to be minor on
the coast (Wolff et al., 2008; Erbland et al., 2013; Shi et al., 2015).
Taken together, NO3- in coastal surface snow might represent some
post-depositional alteration. Even so, a negative correlation between
NO3- concentration and snow accumulation rate was found at the
coast (R2=0.42, p<0.01; Fig. 6a), suggesting that overall the
majority of the NO3- appears to be preserved and is determined by
snow accumulation.
NO3- in inland snowpack
In comparison with the coast, the correlation between Carchived and
inverse snow accumulation is relatively weak in inland regions (Fig. 5d),
suggesting more variable conditions in ambient concentrations and dry
deposition flux of NO3-. In addition, the relationship of
Carchived vs. inverse accumulation inland is opposite to that on the coast.
Based on current understanding of the post-depositional processing of
NO3-, the negative correlation between Carchived and inverse
snow accumulation (Fig. 5d) suggests losses of NO3-. The slope of
the linear relationship indicates an apparent NO3- dry deposition
flux of -44.5±13.0 µeq m-2 a-1, much larger than that
of DML (-22.0±2.8 µeq m-2 a-1), where the snow
accumulation is generally lower than 100 kg m-2 a-1
(Pasteris et al., 2014). At Kohnen Station (an inland site in East
Antarctica), with snow accumulation of 71 kg m-2 a-1, the emission
flux of NO3- is estimated to be -22.9±13.7 µeq m-2 a-1 (Weller and Wagenbach, 2007), which is also smaller in
comparison with this observation. Weller et al. (2004)
proposed that the loss rate of NO3- does not depend on snow
accumulation rate and the losses become insignificant at accumulation rates
above 100 kg m-2 a-1. Among the inland sites, SP10 and Core2
(∼ 800 km from the coast), characterized by high snow accumulation
rate (> 100 kg m-2 a-1; Table 1 and Fig. 1), exhibit
even higher values of Carchived and archived fluxes of NO3-
than those of the coastal sites. It is noted that these two cases influence
the linear regression significantly (Fig. 5d). If the two sites are
excluded, we can get a linear regression with a slope of -27.7±9.2 µeq m-2 a-1, which is comparable to previous reports in DML
(Pasteris et al., 2014).
The depths of inland snow pits cover several to tens of years snow
accumulation, thus allowing for directly investigating NO3-
emission rate. The difference between NO3- concentrations in the
snow layer accumulated during the most recent year (Fig. 4) and in the snow
accumulated during the year before the most recent year can represent the
loss rate of NO3-. If it is assumed that snow accumulation rate is
relatively constant during recent decades at specific-sites, on average,
36.7±21.3 % of NO3- (in µeq L-1) was lost
during 1 year, with the two sites (SP10 and Core2) with snow accumulation
> 100 kg m-2 a-1 excluded from the calculation. The
percentages are generally higher at the sites with lower snow accumulation
rate. Together with snow accumulation rate, the emission flux of
NO3- is calculated to be -28.1±23.0 µeq m-2 a-1,
close to the linear model prediction (-27.7±9.2 µeq m-2 a-1). Significant losses can account for NO3-
profiles at inland sites, i.e., NO3- concentration decreasing with
increasing depths. Previous observations and modeling works suggested that
photolysis dominates the losses (Frey et al., 2009; Erbland et al., 2013;
Shi et al., 2015). During photolysis of NO3-, some of the
photoproducts (NOx) are emitted into the gas phase (Davis et al.,
2004; France et al., 2011), and these products could undergo reoxidation by
the local oxidants (e.g., hydroxyl radical (OH), NO2 + OH + M → HNO3 + M), forming gas-phase HNO3. In inland Antarctica, the
dominant NO3- species in the atmosphere is gaseous HNO3
during summertime, while particulate NO3- is more important in
winter (Legrand et al., 2017b; Traversi et al., 2017). The high levels of
gas-phase HNO3 in summer support the importance of the re-emission from
snow through the photolysis of NO3- in affecting the atmospheric
NOx/ NO3- budget (Erbland et al., 2013). On the one
hand, the gaseous HNO3 can be efficiently co-condensed with water
vapor onto the extensively developed crystal ice layers on Antarctic
plateaus (e.g., Fig. S1 in the Supplement), leading to an enrichment
of NO3- in surface snow (Bock et al., 2016). On
the other hand, a large concentration of HNO3 would enhance its
reaction with sea salt, leading to elevated particulate NO3-
concentrations (Legrand et al., 2017b). The significant
correlation between NO3- and H+ in inland Antarctic surface
snow (R2=0.65, p<0.01) seems to support the importance of
atmospheric gas-phase HNO3 in affecting surface snow NO3-
concentrations, in particular NO3- levels in the crystal ice
samples (Fig. 1).
Several modeling works have been performed to understand NO3-
recycling processes across Antarctica (e.g., Erbland et al., 2015; Zatko et
al., 2016; Bock et al., 2016), but each employs different
assumptions and large uncertainty remains in quantifying NO3-
recycling and preservation. It is thought that emission and transport
strength are the main factors controlling the recycling of NO3-,
while the former is associated with initial NO3- concentrations,
UV and snow optical properties, and the latter is linked with air mass
movement (Wolff et al., 2008; Frey et al., 2009). As a result,
snow accumulation alone is likely insufficient to account for NO3-
variability in surface snow (i.e., no significant correlation between
NO3- concentration and snow accumulation; Fig. 6b).
Relationships between NO3- and co-existing major ions in
surface snow in coastal (top row, a, b and c) and inland (bottom row,
d, e and f) Antarctica. Least-squares regressions are noted with solid
line and are significant at p<0.01. The four samples with high Na+
concentrations are denoted by blue open circles (b), the same as those in
Fig. 8 (the blue open circles). Note that the four samples were excluded in
the plot of NO3- vs. nssCl- (c).
The archived NO3- fluxes vary considerably among inland sites,
from ∼ 3 to 333 µeq m-2 a-1, with high values
generally corresponding to high snow accumulation (Fig. 5e). However, the
nearly 1:1 relationship between Carchived and NO3- flux (Fig. 5f), suggests that accumulation rate is not the main driver of the archived
NO3- concentration. In inland Antarctica, the archived
NO3- fraction is largely influenced by the length of time that
NO3- was exposed to UV radiation (Berhanu
et al., 2015), which decreases exponentially in the snowpack. The e-folding
depth, ze value, is thought to be influenced by a variety of factors,
such as co-existent impurities (e.g., black carbon), bulk density and grain
size (Zatko et al., 2013). In addition, the snow albedo is also
dependent on snow physical properties (Carmagnola et al., 2013). Taken
together, this suggests that the inland plateau is below a “threshold” of
accumulation rate such that the archived NO3- flux cannot be
explained by snow accumulation rate.
Effects of coexisting ions on NO3-
Atmospheric NO3- in Antarctica is thought to be mainly associated
with mid-latitude sources, re-formed NO3- driven by snow-sourced
photolysis products and/or stratospheric inputs (Savarino et al., 2007;
Lee et al., 2014; Traversi et al., 2017, and references therein). Although
organic nitrates can play an important role in the atmospheric NOy
budget, multi-seasonal measurements of surface snow NO3- correlate
strongly with inorganic NOy species (especially HNO3) rather than
organic (Jones et al., 2011). Here, we investigate whether
NO3- in snow is closely associated with coexisting ions (e.g.,
Cl-, SO42-, Na+, K+, Mg2+ and Ca2+)
since these ions have different main sources; e.g., Cl- and Na+
are predominantly influenced by sea salt, and SO42- is likely
dominated by marine inputs (e.g., sea salt and bioactivity source)
(Bertler et al., 2005). In the snow, Cl-, Na+ and
SO42- are the most abundant ions in addition to NO3-.
In surface snow, the non-sea-salt fraction of SO42- accounted for
75–99 % of its total budget, with a mean of 95 %. The percentages were
relatively higher in inland regions than at coastal sites. On the coast, a
positive relationship was found between nssSO42- and
NO3- (R2=0.32, p<0.01; Fig. 7a). Previous
observations suggest that NO3- and nssSO42- peaks in the
atmosphere and snow are usually present in summer (Jourdain and Legrand,
2002; Wolff et al., 2008; Sigl et al., 2016; Legrand et al., 2017a, b). However, the similar seasonal pattern of the two species is
associated with distinct sources, i.e., SO42- is mainly derived
from marine biogenic emissions while NO3- is influenced by
photolysis and tropospheric transport (Savarino et al., 2007; Lee et al.,
2014; Zatko et al., 2016). In the atmosphere, SO42- is typically
found on the submicron particles, while most of the NO3- is
gaseous HNO3 and the particulate NO3- is mainly on
intermediate size particles (Jourdain and Legrand, 2002; Rankin and
Wolff, 2003; Legrand et al., 2017a, b). Thus, the
correlation between NO3- and SO42- is unlikely explained
by the sources or their occurrence state in the atmosphere (i.e., gaseous
and particulate phases). Laluraj et al. (2010) proposed that the
correlation between nssSO42- and NO3- in ice (R2=0.31, p<0.01) could be associated with fine nssSO42-
aerosols, which provide nucleation centers forming multi-ion complexes with
HNO3 in the atmosphere. This assertion, however, should be examined
further, considering that the complex chemistry of SO42- and
NO3- in the atmosphere is far from understood
(e.g., Wolff, 1995; Brown et al., 2006). Thus far, the mechanism of
nssSO42- influencing NO3- in the snowpack, however, is
still debated, and it cannot be ruled out that nssSO42- further
affects mobilization of NO3- during and/or after crystallization
(Legrand and Kirchner, 1990; Wolff, 1995; Röthlisberger et al.,
2000). It is noted that no relationship was found between nssSO42-
and NO3- in inland snow (Fig. 7d), possibly due to the strong
alteration of NO3- during post-depositional processes, as
discussed in Sect. 4.1.2.
Concentrations of NO3- and Na+ in surface snow
samples on the coast. Four samples with high Na+ concentrations are
denoted by open circles, corresponding to those in Fig. 7b. Note that
Na+ concentrations in two samples, 2.5 and 2.8 µeq L-1 in
parentheses, are above the maximum value of the secondary y axis (Na+
concentration). The sample in the dashed ellipse, with Na+
concentration of 2.8 µeq L-1, is the fresh snowfall.
In comparison with nssSO42- aerosols, the sea-salt aerosols
(Na+) are coarser and can be removed preferentially from the atmosphere
due to a larger dry deposition velocity. High atmospheric sea-salt aerosol
concentrations are expected to promote the conversion of gaseous HNO3
to the particulate phase, considering that most of the NO3- in the
atmosphere is in the gas phase (HNO3). In this case, particulate
NO3- can be efficiently lost via aerosol mechanisms. In addition,
the saline ice favors the direct uptake of gaseous HNO3 to the ice
surface. Changes in partitioning between gas phase (HNO3) and
particulate phase will affect NO3- levels due to the different wet
and dry deposition rates of the two species (Aw and Kleeman, 2003).
Thus, sea-salt aerosols play an important role in the scavenging of gaseous
HNO3 from the atmosphere (Hara et al., 2005), and elevated
NO3- concentrations are usually accompanied by Na+ spikes in
the snowpack (e.g., at Halley station; Wolff et al., 2008). Surprisingly, no
significant correlation was found between Na+ and NO3- in
coastal snow (Fig. 7b). The concentration profiles of NO3- and
Na+ in coastal surface snow are shown in Fig. 8, and NO3-
roughly corresponds to Na+ in some areas, e.g., 50–150 and
300–450 km distance inland, although in general they are not very coherent. It is
noted that amongst the four snow samples with Na+>1.5 µeq L-1 (open circles in Fig. 8), only one sample co-exhibits a
NO3- spike. This is different from observations at Halley station,
where Na+ peaks usually led to elevated NO3- levels in
surface snow in summer (Wolff et al., 2008). Of the four largest
Na+ spikes, one is a fresh snowfall sample (dashed ellipse in Fig. 8),
and this sample shows the highest Na+ concentration (2.8 µeq L-1) and low NO3- (0.75 µeq L-1). It is noted that
NO3- concentration in this fresh snowfall is close to the model
predictions (0.7±0.07 µeq L-1; Sect. 4.1.1), validating
that the simple linear deposition model (i.e., the Eq. 6) can depict well
the deposition and preservation of NO3- in coastal snowpack. At
inland sites, no correlation was found between NO3- and Na+
(Fig. 7e), likely explained by the alteration of NO3-
concentration by post-depositional processing.
In surface snow, nssCl- represents 0–64 % (mean of 40 %) of the
total Cl-. On the coast, it is of interest that nssCl- in the four
samples with the highest Na+ concentrations (open circles in Figs. 7b
and 8) are close to 0, and positive nssCl- values were found for the
other samples. The fractionation of Na+ can occur due to mirabilite
precipitation in sea-ice formation at < -8 ∘C (Marion et
al., 1999), possibly leading to the positive nssCl-. However, even if
all of SO42- in seawater is removed via mirabilite precipitation,
only 12 % of sea-salt Na+ is lost (Rankin et al., 2002).
Considering the smallest sea-ice extent in summertime in East Antarctica
(Holland et al., 2014), the high Cl- / Na+ ratio (mean of
2.1, well above 1.17 of seawater, in µeq L-1) in surface snow is
unlikely from sea-salt fractionation associated with mirabilite
precipitation in sea-ice formation. In this case, nssCl- could be
mainly related to the deposition of volatile HCl, which is from the reaction
of H2SO4 and/or HNO3 with NaCl (Röthlisberger
et al., 2003). Thus, nssCl- in snowpack can roughly represent the
atmospherically deposited HCl. In summertime, most of the dechlorination
(i.e., production of HCl) is likely associated with HNO3 due to its
high atmospheric concentrations (Jourdain and Legrand, 2002; Legrand et
al., 2017b). Accordingly, the observed relationship between NO3-
and nssCl- (Fig. 7c) appears to suggest that HCl production can be
enhanced by elevated HNO3 levels in the atmosphere.
With regard to the crystal ice, no significant correlation was found between
NO3- and the coexisting ions (e.g., Cl-, Na+ and
SO42-), suggesting that these ions are generally less influential
on NO3- in this uppermost thin layer compared to the strong
air–snow transfer process of NO3- (Bock et al., 2016).
It is noted that NO3- accounts for most of the calculated H+
concentrations (81–97 %, mean of 89 %), and a strong linear
relationship was found between them (R2=0.96, p<0.01),
suggesting that NO3- is mainly deposited as acid, HNO3,
rather than in particulate form as salts (e.g., NaNO3 and
Ca(NO3)2). This deduction is in line with the atmospheric
observations at Dome C, where NO3- was found to be mainly in
gaseous phase (HNO3) in summer (Legrand et al.,
2017b). On average, the deposition of HNO3 contributes > 91 % of NO3- in the crystal ice (the lower limit, 91 %,
calculated by assuming all of the alkaline species (Na+,
NH4+, K+, Mg2+ and Ca2+) are neutralized by
HNO3 in the atmosphere), suggesting a dominant role of HNO3
deposition in snow NO3- concentrations. The elevated high
atmospheric NO3- concentrations observed at Dome A (> 100 ng m-3;
77.12∘ E and 80.42∘ S; Table S1 in the Supplement) possibly indicate oxidation of gaseous NOx to HNO3,
providing further evidence that NO3- recycling driven by
photolysis plays an important role in its abundance in snowpack on East
Antarctic plateaus.
Conclusions
Samples of surface snow, snow pits and the uppermost layer of crystal ice,
collected on the traverse from the coast to Dome A, East Antarctica, were
used to investigate the deposition and preservation of NO3- in
snow. In general, a spatial trend of NO3- in surface snow was
found on the traverse, with high (low) concentrations on the plateau
(coast). Similarly, NO3- concentrations in the atmosphere are
higher on the plateau than at coastal sites, with a range of 6 to 118 ng m-3.
Extremely high NO3- levels (e.g., > 10 µeq L-1) were observed in the uppermost crystal ice layer, possibly
associated with re-deposition of recycled NO3- from snow-sourced
NOx. As for the snow pits, NO3- exhibits high levels in the
top layer and low concentrations at deeper depths in the inland region,
while no clear trend was found on the coast.
On the coast, the archived NO3- flux in snow is positively
correlated with snow accumulation rate, but negatively with NO3-
concentration. A linear model can well depict the relationship between
archived NO3- and snow accumulation, supporting that atmospheric
levels and dry deposition fluxes of NO3- are spatially homogeneous
on the coast and that dry deposition plays a minor role in snow
NO3- inputs. The dry deposition velocity and scavenging ratio for
NO3- are estimated to be 0.6 cm s-1 and 2700, respectively. In inland Antarctica, the archived NO3- fluxes,
varying significantly among sites, are largely dependent on NO3-
concentration. A weak correlation between snow accumulation and archived
NO3- suggests variable ambient concentrations and dry deposition
flux of NO3-, and the relationship is opposite to that on the
coast. This supports the idea that post-depositional processing dominates
NO3- concentration and distribution in inland Antarctica
(e.g., Erbland et al., 2013, 2015; Shi et al., 2015;
Zatko et al., 2016).
The major ions, Cl-, SO42- and Na+, originate from
different sources than NO3- but could potentially affect the
scavenging and preservation of NO3-. In coastal surface snow, a
positive correlation between nssSO42- and NO3- suggests
the potential influence of fine aerosols on NO3- formation and/or
scavenging, while the coarse sea-salt aerosol (e.g., Na+) is likely
less influential. In contrast to the coast, NO3- in inland surface
snow is dominated by post-depositional processes, and the effects of
coexisting ions on NO3- appear to be rather minor. In inland
surface snow, the strong relationship between NO3- and H+
suggests a dominant role of gaseous HNO3 deposition in determining
NO3- concentrations.