Isotope ratios of ozone and atmospheric nitrate
Atmospheric nitrate concentrations observed at Dome C during the campaign are
presented in Fig. 1, and 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 ‰, and those for 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 to those reported in previous studies (Erbland et al.,
2013; Frey et al., 2009).
Δ17O (primary y axis) and δ15N
(secondary y axis) of atmospheric nitrate collected between November 2011
and January 2012. The samples collected during the intensive measurement
period (December 2011–January 2012) are indicated with open symbols.
A time series showing the year-round record of Δ17O(O3)bulk at Dome C in 2012 is presented in Fig. 3.
Δ17O(O3)bulk averaged 24.9 ± 1.9 ‰,
derived from
Δ17O(O3)term values of 37.4 ± 1.9 ‰. As
shown in Fig. 4, these Δ17O(O3)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 Δ17O(O3)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 quite stable Δ17O(O3)bulk value close
to 26 ‰ that can be considered as representative of the OPALE
campaign held in November–January.
Δ17O(O3)bulk values for the 60 ambient air
samplings done at Dome C throughout 2012. Vertical error bars refer to the
total uncertainty estimated for the technique (±1.7 ‰).
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 Δ17O(NO3-) values
in the atmosphere to 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
17O-excess transfer functions associated with the various nitrate
production pathways (i.e., Δ17O(NO3-)i values) can be
estimated as a function of the Δ17O of nitrate precursor gases
(i.e., NOx, O3, OH, etc.) using mass balance calculations that
trace the origin of oxygen atoms transferred during the chemical
transformation of NOx in the atmosphere. All atmospheric nitrate
production channels involve either NO2 or a NOx reservoir species
derived from NO2 (e.g., N2O5). The first step in determining
the Δ17O signature of each pathway is therefore a quantitative
assessment of the steady-state Δ17O value of NO2, which is
typically calculated as a function of the Δ17O value of O3 and
the reaction dynamics involved in the conversion of NO to NO2. As Dome C
in summer is permanently under sunlight, photochemical inter-conversion of
NOx continues:
Comparison of Δ17O(O3)bulk values obtained
at Dome C with those previously reported by Vicars and Savarino (2014) at
other sites. Box plots indicate the interquartile range (box) and the
median (line), maximum, and minimum values. The mean value is denoted by a
circle.
NO2+hv→NO+O,NO+O3→NO2,NO+HO2/RO2→NO2.
At photochemical steady state (i.e., 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)
Δ17ONO2=α×1.18×Δ17(O3)bulk+6.6,
where the term in bracket represents the laboratory-deduced anomaly transfer
function of the NO + O3 reaction (Savarino et al., 2008),
Δ17O(O3)bulk 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):
α=kNO+O3NOO3kNO+O3NOO3+kNO+HO2NOHO2∗,
with [HO2]* = [HO2] + [RO2].
Example of mass balance calculation of Δ17O for
19 December 2011 at 15:45 local time (UTC + 8 h).
Conditions for 19 December 2011, 15:45
aMedian rate in
Δ17Oic
OH = 3.96 × 106 molecules cm-3
105 molecules cm-3 s-1
in ‰
Net sources of OH
P1
HONO + hv → OH + NO
5.1b
32
P2
H2O2 + hv → 2 OH
1.7
2
P3
O3 + hv + H2O → 2 OH
0.6
20
P4
CH3OOH + hv → HO2 + OH
0.3
0
Recycling RO2 → OH
P5
NO + HO2 → NO2 + OH
7.7
0
P6
HO2+O3 → OH + 2O2
0.4
0
Net sink of OH
L1
CO + OH → HO2 + CO2
6.3
L2
CH4 + OH → CH3O2 + H2O
2.6
L2
HCHO + OH → HO2 + CO
0.8
L4
CH3CHO + OH → CH3CO3
0.9
L5
O3 + OH → HO2 + O2
0.6
L6
H2 + OH + O2 → HO2 + H2O
0.60
L7
CH3OOH + OH → CH3O2 + H2O
0.5
L8
H2O2 + OH → HO2 + H2O
0.3
Net OH losses
L9
NO2 + OH → HNO3
3.9
L10
NO + OH → HONO
0.6
L11
OH + RO2 → products
0.5
L12
OH + RO2NO2 → products
0.6
L13
OH + HONO → NO2 + H2O
0.2
L14
OH + HNO3 → H2O + NO3
0.0
Isotope exchange
E1
HQ + H2O ⟺ HO + H2Q
24.3
NO2 main source
N1
NO + O3 → NO2 + O2
27.0
37
17O-excess NO2
α = (N1/N1 + P5)
0.78
Δ17O(NO2)
29
17O-excess OH
Δ17O(OH)prod = (∑Pi⋅Δ17Oi)/∑Pi
5.8
β=∑Li/(∑Li+E1)
0.43
Δ17O(OH)
2.5
a Production rates obtained from a 0-D box model (see
Kukui et al., 2014, for details). b HONO production rate divided by
a factor of 4 to balance the HOx radical budget (see Kukui et al., 2014,
and Legrand et al., 2014, for justification). c HONO is assumed to
be formed by the photodissociation of nitrate in snow.
Δ17O(NO3-)snow 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 Δ17O(O3)bulk=26 ‰.
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, 2015), 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 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., 2015; 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 = 3/2Δ17O(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 of 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 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 ‰ in Michalski's
simulation, 26 ‰ for our observations), which possibly corresponds
to different conditions of the two studies.
Quantitative assessment of the daily averaged trend in the
Δ17O of NO2 at Dome C during December 2011–January 2012
derived from concurrent measurements of ozone, NO, and HO2 / RO2.
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):
NO2+OH+M→HNO3+M,
leading to the following 17O-excess mass balance (Michalski et al.,
2003; Morin et al., 2011):
Δ17ONO3-=23Δ17ONO2+13Δ17OOH.
In order to predict the Δ17O value of the nitrate produced through
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
midlatitudes (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:
β=LL+kH2O+OH[H2O],
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):
Δ17O(OH)=β×Δ17O(OH)prod.
In plain words, Eqs. (4)–(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 exchange, Δ17O(OH) = Δ17O(OH)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. NO2 as measured by Frey et
al. (2015) represents at most only ca. 10 % (equivalent of ca.
1 ‰) of the total sink of OH, which is predominantly dominated by
reactions with CO, CH4, aldehydes and to a lesser extent by reactions
with O3, H2, and NO. Thus, the possible overestimation of NO2
concentration has only a minor effect on β calculation and is well
embedded within the total uncertainty of such calculation. To assess the
value of Δ17O(OH), we have computed β 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 RO2 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.,
Pwater=6.112×e(17.67×(T-273)T-29.5, with
Pwater in hPa and T in K). The results of this calculation
(Fig. 6) indicate that β varies between 0.70 ± 0.10 (1σ)
and 0.30 ± 0.10 from midnight to noon for conditions prevailing
during the OPALE campaign, suggesting that, on a daily average basis,
approximately 43 % of the Δ17O value originally present in OH
is preserved from exchange with H2O, consistent with estimates for an
Arctic site described by Morin et al. (2007).
December 2011 time series for β, the fraction of the
17O excess originally associated with the OH radical that is preserved
against isotopic exchange with water.
Comparison of measured and calculated Δ17O(NO3-) values.
Sampling period
Measured
Calculated
α constrained
α=1
α constrained
α constrained
by observations
by observations
by observations
Δ17O(OH)* based
Δ17O(OH)*
β=1
Δ17O(OH) based
on HOx budget
based HOx budget
on observed HONO
10–16 Dec
29.6
21.9
25.6
22.6
27.0
16–23 Dec
29.0
21.0
25.6
21.7
26.3
23–30 Dec
27.8
21.6
25.4
22.0
25.7
30 Dec–2 Jan
27.3
21.5
25.3
22.4
24.9
* HONO production rate divided by a factor of 4 to balance the HOx radical budget (see Kukui et al., 2014, and Legrand et al., 2014, for justification).
The value of Δ17O(OH)prod is more difficult to assess
because of the interplay between HO and HO2, and the different sources
involved in OH formation. In the NOx-rich atmosphere at Dome C in
summer, the O(1D) + H2O reaction forming OH is a minor reaction
pathway. When multiple pathways are involved in the production of OH,
Δ17O(OH)prod can be estimated by a simple isotope mass
balance equation where
Δ17O(OH)prod = ∑iPi×Δ17Oi, with Pi the relative production rate of the ith
reaction pathway with respect to the total production rate and
Δ17Oi its associated 17O excess (Morin et al., 2011).
Observations at Dome C during the OPALE campaign indicate that the photolysis
of HONO and the HO2 + NO reaction may represent the most significant
sources of OH at Dome C during the period of seasonal snowpack emissions
(Kukui et al., 2014). However, the measurement of HONO (around few
10 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. (2015). Therefore, the production of OH by HONO photolysis is
consequently adjusted and the 0-D box model (Kukui et al., 2014) is used to
calculate all other production rates of OH. Note that, even when 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 ‰
(1σ). 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-)snow 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 observed and
modeled Δ17O(NO3-) during OPALE is presently unknown, but
we note that such underestimation of the modeled Δ17O(NO3-)
versus the observed Δ17O(NO3-) was also pointed out in 3-D
modeling of the nitrate 17Oexcess (Alexander et al., 2009). A
critical evaluation may nevertheless offer some clues.
Same as Fig. 5 but for Δ17O of OH.