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Biogeochemistry of Seasonally Snow-Covered Catchments (Proceedings of a Boulder Symposium,
July 1995). IAHSPubI.no. 228,1995.
3
Relationship between atmospheric composition and
snow composition for HC1 and HN0 3
FLORENT DOMINE & EMMANUEL THIBERT
Laboratoire de Glaciologie et Géophysique de l'Environnement, BP 96,
54 rue Molière, F-38402 St Martin d'Hères Cedex, France
Abstract The diffusion coefficients and equilibrium solubilities of HC1
and HN0 3 in ice have been measured in the laboratory as a function of
trace gas partial pressure and temperature. To understand the
mechanisms of incorporation of HC1 and HN0 3 in natural snow, these
laboratory data are compared with Greenland measurements of HC1 and
HNO3 in the atmosphere and in fresh snow. The HC1/H20 ratio in
Greenland snow seems to be determined by kinetics, as it can be
explained by considering the number of sticking collisions on the ice
surface. Equilibration does not happen because of the slow diffusion of
HC1 in ice. A preliminary conclusion from limited data is that HN0 3 in
Greenland snow seems to be in equilibrium with the atmosphere,
probably because the diffusion of HN0 3 in ice is fast. Our laboratory
data are also applied to the understanding of meltwater composition.
INTRODUCTION
Before deposition, snow acquires its composition by interacting with the atmosphere.
Atmospheric trace compounds present in the aerosol phase are incorporated mainly by
scavenging (Barrie, 1991) while trace compounds present in the gas phase can be
incorporated by a variety of mechanisms which are poorly understood, but include cocondensation with water vapor while snow crystals grow from the gas phase and riming
(Barrie, 1991). After deposition, the composition of snow changes (Jaffrezo et al.,
1994), in part because of interactions with the atmosphere, which include dry and occult
deposition, and evaporation of trace constituents and water.
Understanding of the interactions between trace gases and snow is therefore
necessary to predict snow composition as a function of atmospheric composition and vice
versa. This is important to reconstruct paleoclimate from polar ice core data (Legrand,
1995).
The study of the incorporation of atmospheric trace gases in snow is also applicable
to snowmelt chemistry: knowing the location of trace constituents in snow may help
elucidate the mechanism of elution of ions from melting snow (e.g. Brimblecombe et
al, 1987).
To contribute to these studies, we have built an apparatus to measure the diffusion
and solubility of trace gases in ice. The results are used to predict the equilibrium
volume composition of ice as a function of air composition. Results on HC1 and HN0 3
are presented, along with applications to atmospheric processes.
Florent Dominé & Emmanuel Thibert
4
EXPERIMENTAL
The apparatus used to measure the diffusion and solubility of trace gases in ice has been
described in Dominé et al. (1994). Briefly, laboratory-grown cylindrical ice single
crystals 8 cm in diameter were cut to a length of 4 cm, placed in a stainless steel
chamber, and exposed to a flow of the desired trace gas that was diluted in watersaturated nitrogen. After a diffusion time of 1 to 3 weeks, the ice crystals were sliced
in sections 25 /xm thick which were melted and analyzed by ion chromatography.
RESULTS
Typical diffusion profiles of HC1 and HN03 are shown in Fig. 1. Diffusion profiles were
fitted to equation (1) (Dominé et al., 1994):
X(x,t) = Xn l-erf
(1)
2JDt
where X(x,t) is the trace gas mole fraction at a distance x from the surface after a
diffusion time t, X0 is the trace gas equilibrium solubility, D is the diffusion coefficient,
and erf is the error function. Diffusion and solubility data are reported in Table 1. Our
ice crystals contained defects, probably small angle boundaries, which acted as diffusion
short circuits. The variability of the defect density in different crystals accounts for the
variability in D, and the values of Table 1 must be taken as upper limits. The
crystallographic system of ice is hexagonal, and diffusion was studied parallel (//) and
perpendicular ( 1 ) to the c axis. Within experimental error, D// and D± are equal.
Contrary to D values, X values are very little affected by short circuits (Dominé et al.,
1994).
- j — i — 1 _
X
o
§
o-
Distance from the surface, x (10" cm)
Fig. 1 Diffusion profiles of HC1 and HN0 3 in ice. Profile numbers refer to the numbers
in Table 1.
Relationship
between atmospheric
and snow composition
for HCl and HN03
Table 1 Experimental results of the diffusion of gaseous HCl and H N 0 3 in ice.
HCl
-7.5
_Lc
0.267
4.32
1.24
7.8
Trace gas
Temperature (°C)
Direction of diffusion
P gas (10- 3 Pa)
Diffusion time (10 ^s)
Xgas do" 6 )
D(10- 12 cm 2 /s)
HCl
-8
lie
1.07
6.05
2.42
6.4
HCl
-8
1c
1.07
6.05
2.44
7.2
HCl
-8
He
1.67
12.6
2.66
2.8
HCl
-15
He
6.67
12.9
5.32
14
HCl
15
1c
06.67
412.9
15.52
7.5
HCl
-8
Xc
1.67
12.6
2.55
2.6
HCl
-15
He
0.267
7.18
1.67
5.0
(2)a
HCl
-15
Polycrystal
0.56
10.4
1.98
9.4
(D a
Trace gas
Temperature (°C) Direction of diffusion
P gas (10- 3 Pa)
Diffusion time (10 -*s)
Xgas do" 6 )
D(10- 12 cm 2 /s)
a
HCl
-15
//c
4.0
18.4
3.89
2.8
HCl
-25
He
4.0
27.7
6.64
1.22
HCl
-25
Xc
4.0
27.7
7.03
1.14
HNO3
-15
He
10.7
5.82
0.286
110
HNO3
-15
He
5.37
5.83
0.225
76
HCl
-15
Xc
1.08
15.5
2.75
2.3
HCl
-15
He
1.69
7.56
3.64
13
(3)a
HNO3
-15
Xc
5.37
5.83
0.20
84
The diffusion profiles of experiments ( 1 ), (2) and (3) are shown in Fig. 1.
The solubility values of HCl and HN0 3 obtained at -15°C have been plotted in
Fig. 2 as a function of trace gas partial pressure, PHC1 or P HNOj . The relationship
between P and equilibrium trace gas mole fraction in ice, X0, is (Hanson &
Mauersberger, 1988a):
0.01
0.1
Mixing ratio (ppbv)
1
10
X
100
1
J
10"
HCl
•=
10"0 -i
«
M
o>
CJ
es
H 10" -4
HNo,
a
HNO3 in fresh snow
1
10--=j
a
•Si
"o
HCl in fresh snow
HT-t
10"6
10"5
10"4
10°
Trace gas partial pressure (Pa)
10"'
Fig. 2 Solubilities of HCl and HN0 3 in ice, as a function of partial pressure.
Measurements of fresh Greenland snow are also reported. The solid line is equations
(3). The dashed line is to suggest that HN0 3 in Greenland snow may be in equilibrium
with the atmosphere according to equation (6), and is speculative (see text).
6
Florent Dominé & Emmanuel Thibert
X0 = A(T)Pl/n
(2)
where A is a coefficient that depends only on temperature T, and n is the number of
entities associated with HC1 in ice. These may include molecular HC1, H + , CI", and
defects such as Bjerrum L defects. For HC1, a fit of the data obtained at -15 °C yields,
with PHC] in Pascals:
XO.HCI = (3-3
± 0.7)
x 10-5 (PHC,) 1 ^ 2 - 6 7 * 0 1 5 )
(3)
The available data from the HN0 3 measurements do not allow the establishment of
a similar relationship between X0 HN0 and P HN0 •
APPLICATION TO THE ATMOSPHERE
Within this work, we determined solubility values for HC1 and HN0 3 in ice under
thermodynamic equilibrium. In the atmosphere, however, snow and gas phase are not
necessarily in equilibrium. Only few data of simultaneous measurements of P in the
atmosphere and X in the snow, that would be necessary to test whether the incorporation
of trace gases in snow is ruled by equilibrium thermodynamics, exist. Dominé et al.
(1995) and Silvente & Legrand (1995) report such measurements, for respectively HC1
and HN0 3 , that were performed at Summit, Greenland, in the summer of 1991 - PHC1
and P H N 0 were measured at ground level using dénuder tubes. Particulate chloride and
nitrate were also measured, but these compounds represented less than 10 and 20% of
gas phase HC1 and HN0 3 , respectively. XHCl and XHN0 were measured in snow during
or just after snow falls (fresh snow), and in snow that had remained on the ground for
several days (aged snow).
Because postdeposition processes can affect snow composition in a manner that is
not well understood at present (Jaffrezo et al., 1994; Dominé et al., 1995), this
discussion will be limited to the relationship between the compositions of fresh snow and
air. Typical values for fresh snow are shown in Fig. 2. The analysis of other ions such
as Na + in snow (Whitlow et al., 1992) show that the marine contribution to chloride and
nitrate were small and that the measurements correspond to dissolved HC1 and HN0 3 .
The incorporation of HC1 and HN0 3 in snow appear to involve different processes, and
these two gases will be discussed successively.
Incorporation of HC1 in ice
The Greenland PHCI values are 2 orders of magnitude lower than the lowest PHC1 used
in our experiments. Our data at -15°C show that, at the very low XHC] found in our
experiments, the solubility of HCi in ice follows a pseudo Henry's law, as observed in
electrolyte solutions, according to equation (3), which is valid over a range of PHC1
varying by a factor of 25. It is reasonable to assume that the pseudo Henry ' s law remains
valid at lower XUCI, and we have extrapolated our HCI results to the lower Greenland
values, as shown in Fig. 2. For the measured PHC1 background value of 1.8 x 10~6 Pa,
Relationship between atmospheric and snow composition for HCl and HN03
1
(40 ng nr 3 STP, 27 pptv) XHC1 in fresh snow is 5 x 10"9 (10 ng g"1) which is a factor of
46 lower than the equilibrium value of 2.3 x 10"7.
Nonequilibrium processes then took place during the formation of Greenland snow.
These may include riming, dissolution in the quasi-liquid layer (QLL) at the surface of
ice, and co-condensation.
No riming was observed at the surface of the Greenland snow crystals (Silvente &
Legrand, 1995; Dominé et al., 1995). Valdez et al. (1989) studied experimentally the
incorporation of S0 2 in ice growing from the vapor phase. They proposed a model
where S0 2 dissolved in the QLL according to Henry's law. A concentration gradient
develops in the QLL, with higher S0 2 concentration at the ice-QLL interface due to
solute exclusion from the ice lattice, and S0 2 is incorporated in the ice according to a
partition coefficient Kd, which is the ratio of the concentration in the ice over the
concentration in the QLL at the ice-QLL interface. This model, however, predicts that
the trace gas mole fraction in ice is always greater than or equal to the equilibrium
concentration, depending on the ice growth rate. This is contrary to what was observed
for HCl in Greenland, and we conclude that the model of Valdez et al. (1989), which
assumes equilibrium dissolution in the QLL, does not apply in this case.
The processes taking place at the air-QLL interface may be ruled by kinetics rather
than by equilibrium. The composition of snow may indeed be determined by the number
of molecules that hit the surface, i.e. by co-condensation. This has been invoked by Sigg
et al. (1992) to explain the incorporation of H 2 0 2 in ice. They suggested:
^H202
=
(4)
^*H 2 0 2 ^H 2 0
To be more accurate, we prefer to also take into account the sticking coefficients,
agas, of the trace gas on ice, as well as the number of colllisions as predicted by the
kinetic theory of gases. This predicts:
Y
_ ^HCl aHCl
HC1
A
^HJO^HJO^
^fL
M
nc\
(5)
where M is the molecular mass. The mean Greenland temperature during the
measurements was — 15°C, implying Pno = 165.7 Pa. The background PHC1 was 1.8
X 10"6 Pa, but it is probably more justified to consider the PHC1 value of the free
troposphere, where the snow formed. The value was probably higher (Dominé et al.,
1995) and we adopt PHC, = 3.6 x 10"6 Pa.
The sticking coefficient data of water on ice, aH 0 of Haynes et al. (1992) yield
a H 0 ~ 0.5 at 258 K ( —15°C) after extrapolation. Leu (1988) measured aHC1 and
obtained 0.4 at 200 K. Van Doren et al. (1990) measured aHCI on water at 273 K and
obtained 0.18. Assuming that, due to the presence of a QLL, the surface of ice at 258 K
is not very different from a water surface, we adopt aHC1 = 0.2 on ice at 258 K.
Including the above two factors, the growth of ice crystals when PHCI^H O = 2-2
x 10 s should result in XHC1 = 6.2 x 10"9, while XHC1 = 5 x 10"9 was measured. These
values are in reasonable agreement, and co-condensation may then explain the chloride
content of the fresh snows analyzed in Greenland.
8
Florent Dominé & Emmanuel Thibert
Solid phase diffusion will tend to increase XHC1 towards the equilibrium value. The
snow crystals were dendritic and about 40 /xm thick (Dominé et al., 1994). The D values
of Table 1 are upper limits, and we use D = 10"12 cm2 s"1 for HC1 diffusion in snow, to
find that the diffusion distance of HC, (Dt)112, is 1 /xm in 3 h. This will result in an Xncl
increase from 5 to about 8 x 1CT9. This modification does not invalidate the cocondensation hypothesis.
We then conclude that incorporation of HC1 in snow crystals growing from the gas
phase is better explained by kinetics of the ice-air interface, followed by solid state
diffusion during snowfall than by equilibrium thermodynamics or by dissolution in the
QLL.
Incorporation of HN0 3 in ice
In the case of HN0 3 , any conclusion will only be tentative because of the fewer data
obtained. We note, however, that Greenland snow contained more HN0 3 than HCl,
despite the lower solubility we measured and the lower partial pressure measured in
Greenland.
To further stress the difference in incorporation with HCl, we also note that a cocondensation scenario is incompatible with the observations, as it predicts XHN0 = 7
x 10~10, i.e. a factor of about 40 lower than what was measured. This ^ HNO value is
obtained using aHNO = 0.2 on ice at 258 K (Van Doren et al., 1990).
These considerations suggest that HN0 3 in Greenland snow might have been in
equilibrium with the atmosphere. This is clearly very speculative and needs to be
confirmed by more experimental data. This suggestion, however, appears to be
compatible with the behavior we expect of HN03 in ice. Because of its size, we expect
HN0 3 to create more defects in the ice lattice than HCl, and «HNO should be greater
than «Hci- A least square fit of the experimental and Greenland data yields:
*HNO3=7.5X10~7(PHNO3)1/4-32
(6)
The value « HN0 = 4.32 thus obtained is compatible with our expectations,
strenghtening somewhat our equilibrium hypothesis which, again, should still be taken
as a very preliminary statement. In any case, HN0 3 in Greenland snow was closer to
equilibrium than HCl, probably because of the faster diffusion of HN0 3 in ice. Indeed,
with D = 10"10 cm2 s"1 the diffusion distance will be 10 ttm in 3 h, while within the
crystals observed, the farthest distance to the surface is 2 /xm.
Application to meltwater composition
The present results and their interpretation add quantitative aspects to the explanation,
based on the difference in solubilities, given by Brimblecombe et al. (1987, 1988) for
the preferential elution of nitrate and sulfate relative to chloride in melting laboratory
snow.
Relationship between atmospheric and snow composition for HCl and HN03
9
These authors used ice with an initial content of sulfate, nitrate and chloride of 10~4
mole"1 each, (X = 1.96 x 1CT6). For nitrate, this is well above the solubility in ice for
F HNOj values up to 0.0107 Pa at T = -15°C. At -15°C, little more HN0 3 can be
incorporated in ice, because at P HN0 > 0.015 Pa, ice will melt and form a solution
with Z HNOj « 0.08 (Hanson & Mauersberger, 1988b). At T > -15°C, the solubility
will be even lower. The ice of Brimbelcombe et al. (1987) was then supersaturated in
nitrate by a factor of at least 6. Partial melting then happened, and at least 5/6 of the
nitrate was located in the water, formed at the grain boundaries.
Because the solubility of chloride in ice is much higher than that of nitrate, it
probably would not induce melting by itself. In the absence of melting, chloride does not
seem to be incorporated in grain boundaries as suggested by an experiment (see Table 1)
that used polycrystalline ice with crystals about 1 cm in size, and where no enhanced
solubility was noted. In the presence of nitrate-induced meltwater, some chloride
partitioned to the liquid, explaining why the initial meltwater fractions contained
chloride. When the initial melt water had been removed, most of the nitrate had been
lost while the chloride incorporated in the ice lasted untill the final meltwater fractions.
Data on the solubility of sulfate in ice are lacking. Mulvaney et al. (1988), however,
observed that in Antarctic ice sulfate was located at grain boundaries. Because solubility
decreases as temperature increases, sulfate is then also located at grain boundaries at
temperatures higher than in Antarctica, which would explain why it behaved like nitrate.
This reasoning also predicts that at concentrations higher than the chloride solubility
limit, chloride, nitrate and sulfate should behave similarly, as reported by Brimblecombe
etal. (1988).
CONCLUSION
Our experimental data show that HCl and HN0 3 have very different solubilities and
diffusion coefficients in ice. These data have implications for the understanding of fresh
snow and meltwater composition. Our work suggests that the growth rate of snow
crystals in the atmosphere should influence snow composition, and experiments similar
to those of Valdez et al. (1989) using a variety of trace gases would be useful. The
simultaneous measurements of atmospheric and snow composition are also essential to
understand snow composition.
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Florent Dominé & Emmanuel
Thibert
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