JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 110, D22304, doi:10.1029/2004JD005580, 2005
Variability of active chlorine in the lowermost Arctic stratosphere
Brett F. Thornton,1 Darin W. Toohey,1 Linnea M. Avallone,1,2 A. Gannet Hallar,1,2,3
Hartwig Harder,4,5 Monica Martinez,4,5 James B. Simpas,4 William H. Brune,4
Makoto Koike,6 Yutaka Kondo,7 Nobuyuki Takegawa,7
Bruce E. Anderson,8 and Melody A. Avery8
Received 5 November 2004; revised 9 June 2005; accepted 8 July 2005; published 16 November 2005.
[1] We examine the variability of ClO in the Arctic upper troposphere and lowermost
stratosphere (UTLS) during the winter of 1999–2000. Data are binned relative to NOy, a
species that is a proxy for photochemical age and a photochemical source of NOx.
Enhancements in the [ClO]/[Cly] ratio relative to values expected from gas-phase
chemistry alone were observed throughout the region and were largest in the coldest
sampled regions, where T < 208 K. At low NOy values, where particles containing NOy
and water were often detected, twilight ClO abundances in the afternoon were nearly a
factor of 3 larger than those in the morning. At higher NOy values, where much lower
particle surface areas were measured, ClO abundances in morning twilight were somewhat
larger than those in the afternoon. These observations are consistent with a daytime
mechanism of rapid heterogeneous activation of inorganic chlorine in particle-rich, lowNOy regions, with slower deactivation in relatively particle-poor, higher-NOy regions
of the lowermost stratosphere. While the data clearly show widespread chlorine activation,
knowledge of the precise value of the [ClO]/[Cly] ratio is limited because of the lack of
available data on inorganic chlorine species, notably HCl, believed to be the dominant
reservoir of inorganic chlorine at these altitudes.
Citation: Thornton, B. F., et al. (2005), Variability of active chlorine in the lowermost Arctic stratosphere, J. Geophys. Res., 110,
D22304, doi:10.1029/2004JD005580.
1. Introduction
[2] Understanding ozone trends in the lowermost stratosphere, the tropopause region and the first few kilometers
above, is vital for accurate predictions of future global
climate change, as ozone in the lowermost stratosphere
plays a crucial role in the oxidative capacity and radiative
balance of the atmosphere [World Meteorological Organization (WMO), 2003]. Ozone trends on the order of a
Dobson unit (DU) decade1, representing upward of 25%
of the overall trends in column ozone detected at midlatitudes [WMO, 2003], have been detected in the midlatitude
lowermost stratosphere (10 – 15 km). The sources of such
decreases could be enhancements of reactive chlorine on
1
Program in Atmospheric and Oceanic Sciences, University of Colorado, Boulder, Colorado, USA.
2
Also at Laboratory for Atmospheric and Space Physics, University of
Colorado, Boulder, Colorado, USA.
3
Now at NASA Ames Research Center, Moffett Field, California, USA.
4
Department of Meteorology, Pennsylvania State University, University
Park, Pennsylvania, USA.
5
Now at Luftchemie, Max-Planck-Institut für Chemie, Mainz, Germany.
6
Department of Earth and Planetary Science, Graduate School of
Science, University of Tokyo, Tokyo, Japan.
7
Research Center for Advanced Science and Technology, University of
Tokyo, Tokyo, Japan.
8
NASA Langley Research Center, Hampton, Virginia, USA.
Copyright 2005 by the American Geophysical Union.
0148-0227/05/2004JD005580$09.00
particle surfaces and increasing trends in chlorine source
gases [Solomon et al. 1997], and/or changes in transport
processes [Fusco and Salby, 1999].
[3] Ozone destruction in the lowermost polar stratosphere
is of concern because of the impact at lower latitudes.
Owing to the long chemical lifetime of ozone in this region,
transport and mixing can spread the influence of highlatitude ozone losses over much of the Northern Hemisphere. In fact, dilution effects as proposed for the spread of
ozone-depleted air from the Arctic polar region [Knudsen et
al., 1998] will likely occur faster in the near-tropopause
region because intrahemispheric transport processes there
are more vigorous than at higher altitudes [Knudsen and
Grooß, 2000]. If the decreases in ozone observed near the
tropopause are due to chemistry, a proposed mechanism
begins with enhancements to reactive chlorine (defined here
as ClOx ClO + OClO at low abundances of Cly because
Cl2O2 and Cl are negligible under these conditions), an
indicator of ozone destruction by breakdown products of
chlorine source gases such as CFCs and methyl chloride.
These enhancements to ClOx have been proposed to occur
in the lowermost stratosphere on ice particles (e.g., cirrus)
or water-rich background aerosols containing sulfate and
nitrate [Borrmann et al., 1996, 1997; Solomon et al., 1997;
Meilinger et al., 2001]. BrO near the midlatitude tropopause
is generally thought to be small, near 1– 2 ppt [Harder et al.,
1998] but even a small increase in this value leads to
considerably larger ozone losses via synergistic reactions
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of BrO and ClOx [e.g., Salawitch et al., 2005]. Consistent
with this mechanism, Keim et al. [1996] found significant
ClO enhancements and NOx reductions in a narrow layer
influenced by Mount Pinatubo aerosols in the midlatitude
lower stratosphere. This is analogous to the ‘‘ozone hole’’
chemistry occurring higher in the stratosphere on polar
stratospheric clouds (PSCs).
[4] The evidence for chlorine activation in the lower
stratosphere during volcanically quiescent years has been
mixed. On the basis of measurements of water vapor and
temperature from the NASA ER-2 aircraft during the
Airborne Arctic Stratospheric Expedition (AASE), Murphy
et al. [1990] noted that ice clouds with the potential to
activate chlorine should occur above the tropopause at high
latitudes. Analyzing data from that same mission, King et
al. [1991] found ClO abundances near 20 km to be quite
low outside of the perturbed polar vortex. On the basis of a
balloon profile that extended down to the tropopause,
Avallone et al. [1993a] noted that ClO at 16 km was
significantly higher than could be explained by a model
that included reactions on background sulfate aerosols
[Brasseur et al., 1990; Rodriguez et al., 1991]. Borrmann
et al. [1997] later examined ClO and particle data from the
ER-2 during AASE-II and found apparent ClO enhancements that were correlated with enhanced particle surface
areas near the tropopause at midlatitudes. The prevalence of
these particles has been debated: using a more extensive
ER-2 data set, Smith et al. [2001] argued that the midlatitude lowermost stratosphere is a region that is subsaturated
with respect to ice, with average ClO abundances of only a
few ppt (parts per trillion by number, also pmol mol1) or
less there. They argued that chlorine activation was unlikely
in this region because of the lack of sufficient particle
surface area. At higher northern latitudes in winter, however,
Hallar et al. [2004] frequently observed ice clouds above the
tropopause, consistent with the observations of Murphy et al.
[1990]. Under these conditions, Thornton et al. [2003]
measured considerable enhancements in ClO. A key difference between the Thornton et al. [2003] observations and
those of Smith et al. [2001] was that the sampled regions of
the Arctic winter tropopause were saturated with respect to
water and exhibited large regions with condensed water and
nitric acid above the tropopause, presumably where tropospheric and stratospheric air masses were mixing.
[5] On the basis of such observations, it is reasonable to
conclude that chlorine activation in the lowermost stratosphere is likely to be an episodic process, as illustrated in
the recent modeling study of Bregman et al. [2002].
Consistent with this view are the analyses of Gierens et
al. [1999], who found ice supersaturation in about 2% of
their observations from commercial aircraft in the lowermost stratosphere during Measurement of Ozone on Airbus
In-service Aircraft (MOZAIC), and those of Goldfarb et al.
[2001] who showed significant occurrences of cirrus at and
above the thermal tropopause over the Observatoire de
Haute Provence (43.9N) between 1997 and 1999. However, water-rich particles alone may not be efficient surfaces
for halogen activation. Simpson et al. [2003] used airborne
measurements of C2Cl4 and ethane in the tropical and
midlatitude upper troposphere over the western Pacific
Ocean in spring 2001 (a notably warmer environment than
the tropopause region) to find an upper limit of 1.5 ppt ClO,
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and suggested that Cl activation on cirrus is not a major
process in the studied region. Recent laboratory work on
coadsorption of HCl and HNO3 on ice showed that HNO3
acidified the ice surface and preferentially displaced HCl,
suggesting that chlorine activation on cirrus particles may be
substantially slowed when HNO3 is in excess of HCl [Hynes
et al., 2002]. Fluckiger et al. [2000] studied HCl diffusion in
ice under UTLS conditions and showed that production of
Cl2 or HOCl depends on the size of the ice particle.
[6] In the lowermost stratosphere, abundances of inorganic chlorine decrease sharply with decreasing altitude,
such that it is difficult to use measurements of ClO alone to
examine chlorine activation near the tropopause. ClO abundances of even a few tens of ppt could represent large
enhancements in available inorganic chlorine (symbolized
here as Cly). In fact, high levels of Cl presumed to come
from in situ activation of HCl and ClNO3 were inferred
above the Arctic winter tropopause by Lelieveld et al.
[1999] using aircraft measurements of ethane and carbon
monoxide as indicators of chlorine atom abundances.
Thornton et al. [2003] used the ratio [ClO]/[Cly] to show
that the fraction of inorganic chlorine in reactive forms was
enhanced near the Arctic winter tropopause poleward of
55N. To develop a better understanding of the distribution
of reactive chlorine and the nature of halogen activation in
the lowermost stratosphere and upper troposphere, here we
further examine the behavior of ClO and the [ClO]/[Cly]
ratio under various conditions at high latitudes during
winter, focusing primarily on the dependences with photochemical age, temperature, particulate loading, and time of
day.
2. Methods
2.1. Measurements
[7] ClO was measured in situ on the NASA DC-8 during
SAGE III – Ozone Loss and Validation Experiment/Third
European Stratospheric Experiment on Ozone 2000
(SOLVE/THESEO 2000), as described by Thornton et al.
[2003]. This chemical conversion/resonance fluorescence
instrument as configured on the DC-8 detected ClOx, or the
sum of ClO + OClO. As shown in auxiliary material1, OClO
can be a dominant form of ClOx in darkness; however,
OClO is photolyzed at low solar zenith angles (SZA < 93),
such that ClOx ClO in sunlight. Further discussion of the
ClO and OClO reaction kinetics inside the instrument flow
system is provided in the auxiliary material.1
[8] As described previously [Thornton et al., 2003], the
measurements of ClO from the DC-8 were carried out in
tandem with measurements of HOx by the ATHOS instrument (Airborne Tropospheric Hydrogen Oxides Sensor), a
technique that required significantly lower pressures than
are commonly used for ClO detection, thereby reducing the
signals from ClO by a factor of 10. The instrument is
relatively insensitive to Cl2O2 because thermal decomposition of this species within the ATHOS flow system and the
reaction of Cl2O2 with NO are very slow (less than 1 s1)
under the conditions employed to measure ClO.
1
Supporting material is available via Web browser or via Anonymous
FTP from ftp://ftp.agu.org/apend/jd/2004JD005580.
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[9] In this paper, where we are mainly interested in
variations of ClO, we report 1s precision based on photon
counting statistics. Overall accuracy is estimated at ±25%
(1s) because of uncertainties in laboratory calibrations as
described elsewhere [e.g., Brune et al., 1988; Toohey et al.,
1993]. This is somewhat larger than for the same instrument
as flown on balloons and the WB-57F aircraft because of
the very low pressures employed by ATHOS [e.g., Thornton
et al., 2003]. For the analysis of trends in the data presented
here, we are mainly concerned with precision, which is a
function of the number of individual measurements that are
averaged.
[10] The in situ NOy measurements have been summarized by Koike et al. [2002] and Kondo et al. [2003]. NOy
(both particle and gas phase) is converted to NO by reaction
with CO on a heated gold catalyst, and NO is detected by
chemiluminescence following reaction with O3 [Kondo et
al., 1997]. Gas phase NOy, measured with a backward
facing inlet, and enhanced particulate total NOy (henceforth
ENOy), measured with a forward facing inlet, were measured at 1 Hz. In this paper, we refer to the ENOy data that
have been corrected for inlet enhancements as TNOy.
Although NOx was also measured in situ during SOLVE/
THESEO 2000, typical values were small, of order 1% of
TNOy, which is near the detection limit. Indeed, NOx must
be low to avoid rapidly sequestering ClO into ClNO3.
[11] TOTCAP (Tropospheric Ozone and Tracers from
Commercial Aircraft Platforms) is a suite of four instruments designed to fit in a half rack of the DC-8. Data from
two of the TOTCAP instruments are employed in this study.
For one estimate of Cly, CFC-11, CFC-12, and Halon-1211
were measured with a two-column gas chromatograph (GC)
approximately every 3.5 min with a precision of 1% and
an accuracy of 2%. As described by Hallar et al. [2004],
total water measurements were obtained with the closedpath tunable diode laser open-path hygrometer (CLH) [May,
1998]. With a subisokinetic inlet system, similar to that
for the ENOy measurements, inertial enhancements are
observed when particles are larger than 5 – 10 mm in
diameter. This greatly increases sensitivity to low concentrations of water commonly seen in the relatively dry
lowermost stratosphere. Inlet enhancement factors for both
the NOy and TOTCAP water instruments were calculated
using simplified ram heating and inlet speed functions [e.g.,
Kondo et al., 2003].
[12] Ozone was measured by observing near-infrared
chemiluminescence from excited-state nitrogen dioxide
formed by the reaction of pure reagent nitric oxide with
ozone in sampled air. This well-established technique is
described by Gregory et al. [1987], and has been adapted
for use on the DC-8 aircraft [e.g., see Avery et al., 2001].
The measurements are performed by combining pure
reagent nitric oxide (NO) with incoming sample air in a
small volume reaction chamber, and measuring the resultant
chemiluminescence. The reaction chamber is maintained at
constant temperature and pressure (25 torr) by buffering
ambient pressure changes with a larger-volume prechamber
maintained at 100 torr. Sampled air enters the aircraft
through a forward facing, Teflon-lined, J-shaped probe that
has been demonstrated to be insensitive to aircraft attitude.
Approximately 2 STD L min1 of air is pulled into the
instrument prechamber from ram-induced flow through the
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probe, and sample flow into the reaction chamber is maintained at 500 STD cm3 min1. The instrument is calibrated
by referencing to the NIST standard ozone photometer.
Measurements are accurate to 5% with a precision of 2%.
[13] Particle surface area was determined from measurements of particle size distributions using the Forward
Scattering Spectrometer Probe (FSSP-300) instrument built
by Particle Measurement Systems, Boulder, Colorado
[Baumgardner et al., 1992]. This instrument uses Mie
scattering of laser light off particles to compute the size
and occurrence of particles in 31 bins between 0.4 and
20 mm in diameter. The calculation assumes particles of a
density of 0.8 g cm3, which underestimates their diameters
somewhat. Therefore the surface areas shown here represent
an underestimate of the true particle surface area.
2.2. Binning and Averaging of ClO
[ 14 ] ClO abundances at DC-8 altitudes throughout
SOLVE/THESEO 2000 were much lower than values
normally observed at higher altitudes. At such low values,
which are near the detection limit of the instrument for short
integration times, photon-counting noise is the dominant
contributor to the variability in individual measurements,
such that it is necessary to average many measurements to
obtain adequate precision. Extensive tests in the laboratory
and probability distribution functions of in-flight raw detector noise confirm the Poisson statistics assumed in our
propagation of errors. Thus the precision is improved by the
square root of the number of individual 40-s data points that
make up each average.
[15] A confounding factor in averaging a large number of
individual observations to achieve adequate precision is
selecting an appropriate coordinate system. In this paper,
we examine trends of ClO with temperature and altitude or
‘‘age’’ of stratospheric air, for which a tracer (or proxy) such
as ozone or NOy is used. Because the measurements of
those tracers are reported every second, those observations
are first averaged for 40 s to match the ClO measurement
cycle, then the measurements of ClO and other variables of
interest are binned with respect to the averaged tracer
values. The bin size of the tracer is increased until the
proper number of individual 40-s measurements is achieved
to reach the desired precision.
[16] During SOLVE/THESEO 2000, the DC-8 measurements were concentrated between 60N and 85N latitude
during winter. In this analysis, we focus on the results from
flights that departed from and returned to Kiruna, Sweden.
Many of these flights were to the north and west of Kiruna.
Therefore the vast majority of observations were obtained in
twilight and in darkness. Although Thornton et al. [2003]
noted that significant abundances of reactive chlorine (most
likely OClO) were observed in darkness, such that the diel
change in ClOx was only about 50%, in this paper we do not
include data in regions that are in complete darkness in
order to avoid obvious ambiguities in interpretation of those
measurements. However, to achieve reasonable precision,
we are forced to include measurements when the sun is very
close to the horizon as viewed from flight altitudes. At low
abundances of ClO and NOx, such as prevailed in the neartropopause region for the duration of SOLVE/THESEO
2000, OClO can be a dominant nighttime reservoir of ClOx,
and the diurnal variation of ClOx = ClO + OClO is
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significantly smaller than that expected at higher altitudes
and lower latitudes [e.g., Brune et al., 1990]. In auxiliary
material we present an example of a model run from
Thornton et al. [2003] that illustrates this relatively low
sensitivity of ClOx values to changing solar zenith angles at
low-NOx conditions.
[17] To characterize the diurnal behavior of reactive
chlorine, we use local solar time to segregate the observations between SZAs of 88 and 93. The averages of data in
the SZA range of 88 to 93 obtained before local solar
noon (i.e., mornings) we refer to as ‘‘sunrise’’; data in that
range after local solar noon (i.e., afternoons) as ‘‘sunsets.’’ It
is important to note that at Arctic latitudes in winter, such
sunrises and sunsets can take many hours, given the limited
distance the sun rises above the horizon during the day. For
each of the averages reported here, we have also computed
the variation in ClO with a high-SZA cutoff that ranges
from 91 to 95. We found no systematic change in the
average ClO abundances with changes of this cutoff over
this range, other than a decrease in variation due to the
concomitant increase in precision as additional data are
included in the average.
2.3. Cly Determination
[18] Knowledge of Cly, or total inorganic chlorine (Cly =
Cl + 2 Cl2 + ClO + OClO + 2 Cl2O2 + HOCl + BrCl +
ClNO3 + HCl, etc.), is useful in the context of the present
work because the ratio of [ClO]/[Cly] is a measure of the
fraction of inorganic chlorine that is partitioned into reactive
(e.g., ozone destroying) forms. In the lowermost stratosphere, models that assume gas-phase processes alone
typically generate about 1% active chlorine in sunlight (that
is, 1% of the inorganic chlorine is ClO) [King et al., 1991].
As shown by Thornton et al. [2003], mixing ratios of ClO
observed at DC-8 altitudes were of order 15 –20 ppt for
most of SOLVE/THESEO 2000. This produced a maximum
[ClO]/[Cly] ratio of 5%, a factor of 5 larger than expected
from gas-phase processes, assuming that Cly abundances
were between 200 and 500 ppt [e.g., Avallone et al., 1993b].
[19] The largest uncertainty in [ClO]/[Cly] is in the [Cly]
term, and is due mainly to the lack of knowledge of
abundances of inorganic chlorine (primarily HCl and
ClNO3) entering the stratosphere. Thornton et al. [2003]
assumed a value of 100 ppt for HCl at the tropopause,
which represented about 50% of the Cly value in the lowest
ozone bin, which was composed primarily of tropospheric
air, and is generally in agreement with early filter measurements of HCl [Farmer et al., 1976; Lazrus et al., 1977]. In
that analysis, which followed the approach of Engel et al.
[1997], the intercept of the Cly versus stratospheric tracer (in
that case, ozone) relationship was about 200 ppt. For most
analyses of ozone chemistry at higher altitudes, where
abundances of Cly are larger than 1000 ppt, uncertainties
in this intercept are not too significant. However, for
measurements near the tropopause, Cly and the uncertainty
in Cly are dominated by this intercept, which represents the
source of tropospheric inorganic chlorine to the Cly budget.
Recent high precision in situ measurements near the subtropical tropopause [Marcy et al., 2004], closer to the source
of air to the stratosphere, found <25 ppt HCl there,
suggesting that previous formulations of Cly significantly
overestimated the contribution of inorganic sources to the
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total stratospheric chlorine burden. Because this relationship
is most valid for lower midlatitudes, and may not be
representative for the Arctic winter/spring conditions
encountered by the DC-8 during SOLVE/THESEO 2000,
we have assumed that tropopause ozone is 50 ppb larger in
the DC-8 region than in the region sampled by Marcy et al.
[2004]. Thus we added 50 ppb to our observed ozone values
before using the Marcy et al. [2004] relationship to determine HCl values for the Arctic region.
[20] In the present paper we again employ the method
used by Engel et al. [1997], in that TOTCAP GC data were
used to determine the slope of the Cly versus a long-lived
tracer curve due to known organic sources. However, we
now consider the intercept of that curve due to the inorganic
sources to be relatively uncertain. Thus we employ an
intercept that varies over a range of values that encompass
those reported by Marcy et al. [2004] and Thornton et al.
[2003]. Specifically, measurements of a select group of
halocarbons are used to scale up to a total organic chlorine
value (e.g., CCly) based on observed relationships between
that group and total CCly from more extensive, but less
frequent, flask sample measurements. Then, the organic
contribution to Cly is determined by subtraction of these
values from a value for CCly at the tropopause based on the
NOAA CMDL flask-sampling network [Montzka et al.,
1999], and allowing for a 1-year lag time for the gases to
reach the tropical tropopause region. The total chlorine
burden is simply the sum of this organic contribution and
the mixing ratio of inorganic chlorine that enters the
stratosphere. This latter value is estimated from measurements of HCl when the organic contribution to Cly is close
to zero (that is, when HCl Cly). For purposes of this
paper, we assume that ozone can be used as a proxy to
determine the region where Cly is composed mainly of
inorganic chlorine.
[21] In addition to the two estimates of Cly derived from
CCly and the inorganic chlorine entering the stratosphere,
we include a third that is based on an approximate empirical
relationship for the midlatitudes [Avallone et al., 1993a;
Thornton, 2004] that Cly 0.1% of the ozone value. We
expect that actual Cly values are likely to fall within the
ranges of these three estimates, but we stress here that until
new measurements (e.g., HCl and ClNO3) are obtained at
these altitudes and seasons, uncertainties in the inorganic
input at the tropopause will dominate the uncertainties in
Cly for the conditions studied here.
3. Results and Discussion
[22] Thornton et al. [2003] found that average ClO
abundances in the lowermost stratosphere over the Arctic
were relatively uniform. About 10– 20 ppt were observed in
daylight over the range of ozone abundances most frequently
encountered by the DC-8. Although the amount of
reactive chlorine was relatively constant over this range of
photochemical ages (e.g., higher ozone indicates greater
photochemical age), the fraction of available chlorine that
was in reactive forms varied by a more than a factor of 2. In
Figure 1a we show the variability in the [ClO]/[Cly] ratio
versus temperature using ozone (shown in Figure 1b) as a
proxy for Cly for three assumed Cly cases: Cly as given by
Thornton et al. [2003]; Cly based on HCl at the tropopause
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as measured by Marcy et al. [2004] and modified for polar
latitudes; and Cly based on the empirical ozone-Cly relationship for the lowermost Arctic stratosphere. The importance of proper choice of Cly is clearly evident. Using
values from Thornton et al. [2003] that assume the highest
tropospheric inorganic chlorine contribution to Cly, there is
no clear trend in [ClO]/[Cly] with temperature. As the
intercept of the Cly versus ozone relationship is reduced,
there is a distinct increase in [ClO]/[Cly] at temperatures
below 205 K, becoming as large as 25% at the lowest
temperatures for the smallest assumed value for the Cly
intercept. This result is due entirely to the nature of the air
masses sampled by the DC-8 during SOLVE/THESEO
2000: The region nearest the tropopause, where Cly values
are smallest, was also where the lowest temperatures
prevailed. An increase in [ClO]/[Cly] with decreasing temperature may be more realistic than the flat behavior
exhibited by the ‘‘base’’ case (that of Thornton et al.
[2003]), which likely overestimates Cly as it is based on
earlier, primarily satellite-derived, assumptions (made before the Marcy et al. [2004] HCl measurements) which
predicted a larger tropospheric inorganic chlorine contribution to Cly. (The data in Figure 1 include measurements
below 93 SZA. Included in the auxiliary material is a
table of mean solar zenith angles for each point plotted in
Figure 1a.)
[23] Figure 1b shows surface area versus temperature. At
205 K and lower temperatures there is a very large increase
in particulate surface areas measured by the FSSP. It is also
in this region where particles contained a significant fraction
of NOy (presumably nitric acid) [Kondo et al., 2003] and
particles containing ice were often observed [Hallar et al.,
2004] (Figures 1c and 1d). It is unlikely that the air masses
with temperatures below 205 K originated higher in the
vortex, where PSC chemistry could have produced the
enhancements of reactive chlorine. This is especially true
for the observations in December 1999, a time when
temperatures were only low enough for PSC formation very
Figure 1. Plots versus temperature of (a) [ClO]/[Cly],
(b) ozone and Forward Scattering Spectrometer Probe
(FSSP) surface areas, and (c) TOTCAP enhanced water
and (d) NOy, both uncorrected for inlet enhancements for all
SOLVE/THESEO 2000 flights that both originated and
terminated in Kiruna, Sweden. Included are data from solar
zenith angle (SZA) < 93 for which both NOy and ClO
measurements are available. In Figure 1a the three [ClO]/
[Cly] curves use Cly based on (1) the Cly-ozone relationship
given by Thornton et al. [2003], a curve with a Cly-ozone
intercept of 200 ppt Cly at zero ozone; (2) the observations
from Avallone et al. [1993a] and Thornton [2004] that Cly 0.1% ozone for this region of the atmosphere, which results
in, of course, a Cly-ozone intercept of 0 ppt Cly at zero
ozone; and (3) the observations of Marcy et al. [2004] that
showed <25 ppt HCl (Cly) near the subtropical tropopause, a relationship with a Cly-ozone intercept of 25 ppt
Cly at zero ozone, Cly being zero when ozone is 50 ppb.
This intercept was scaled slightly upward from the value of
Marcy et al. to account for the expected subtropical-polar
differences in HCl (see text). Thus it is possible that even
this third estimate overestimates Cly, which would imply an
even higher [ClO]/[Cly] ratio.
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high (>24 km) in the vortex and yet abundances of ClOx
were similar to those found in January and later. Nevertheless, for the analysis in this paper we shift from a coordinate
system based on ozone, a species that is destroyed by
photochemical reactions during this period (as given by
Thornton et al. [2003]), to one based on NOy.
[24] Koike et al. [2002] report a very compact relationship
between NOy and the long-lived tracer N2O below 14 km
during the 1999 – 2000 winter. N2O is a good surrogate for
Cly [e.g., Engel et al., 1997, and references therein] whereas
NOy, as the source of NOx, could strongly influence reactive
chlorine. We have chosen to bin the observations with
respect to NOy for the analyses presented here, where our
focus is mainly on variations in ClO and [ClO]/[Cly].
Because redistribution of NOy by sedimentation of HNO3containing particles, as observed at much higher altitudes
[e.g., Popp et al., 2001] could weaken the correlation
between NOy and Cly, we maintain ozone (which correlated
strongly with N2O) as the surrogate for Cly. However, Koike
et al. [2002] have shown that significant redistribution of
NOy occurs only at relatively high NOy values (above
1 ppb). Thus we conclude that NOy is a good observable
for segregating air masses of different histories and origins
in the lowermost stratosphere during the SOLVE/THESEO
2000 observational period.
[25] The lack of values of NOy much greater than 4 ppb
at DC-8 altitudes indicates that unmixed descent of air from
significantly higher altitudes in the polar vortex did not
influence the observations at DC-8 altitudes. Hence it is
unlikely that the chemical effects reported here resulted
from processes occurring higher in the vortex. This is
consistent with analyses that have shown descent in the
winter polar stratosphere to be greatest in the upper, rather
than lower, stratosphere [Schoeberl and Hartmann, 1991],
especially at the vortex edge [Manney et al., 1999]. Furthermore, if denitrified air from higher altitudes (Q 500 K)
were to be transported down to DC-8 altitudes (Q 350 K),
one would expect to see a breakdown in the compact
relationship between NOy and N2O, which, as pointed out
above, was not the case.
[26] One of the interesting features that was evident in the
figures of Thornton et al. [2003] was the apparent excess of
ClOx in darkness relative to values in daylight at low ozone,
whereas at higher ozone values, ClOx in darkness was about
half that in daylight, the latter being what one would expect
from a detailed photochemical model (see auxiliary material
and Thornton et al. [2003]). To examine this in further
detail, in Figure 2a we have plotted ClOx averages at sunrise
and sunset versus TNOy. Clearly visible in this plot is a
change in the behavior of ClOx with increasing TNOy. ClOx
at sunrise increases with TNOy, as might be expected
because of the increase in available chlorine (Cly); however
Figure 2. Plots versus total TNOy of (a) ClO, (b) FSSP
surface areas, (c) TOTCAP enhanced water, and (d) [ClO]/
[Cly]. Cly is calculated as in Figure 1 for case 1 as given by
Thornton et al. [2003] and case 3 using the Marcy et al.
[2004] HCl-ozone relationship scaled for subtropical-polar
differences. Data are binned separately with respect to
sunrise and sunset as described in text. Included are data
from 88 < SZA < 93 for which both NOy and ClO
measurements are available.
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the opposite is true at sunset, where ClOx decreases with
increasing TNOy. Because Cly increases with increasing
TNOy, the [ClO]/[Cly] ratio at sunset is largest at the lowest
values of TNOy. We have examined the histograms of SZAs
for the individual data points plotted in Figure 2 in search of
a possible bias in the sampling, for example, a greater
fraction of measurements at high SZA for those points that
have the smallest ClOx. However, there are no obvious
trends in the histograms that can explain these variations.
Included in the auxiliary material is a table of mean solar
zenith angles for each point plotted in Figure 2a. For the
ensemble of points obtained below 900 ppt TNOy, the mean
SZA for sunrise data was 90.5 and that for sunset data was
89.9, whereas at TNOy above 900 ppt the mean SZA for
sunrise data was 90.2 and that for sunset data was 90.7.
On the basis of the model results shown in the auxiliary
material, this difference of 0.5– 0.6 could not account for
the factor of 2 – 3 differences in ClOx that are clearly
apparent in the low- and high-TNOy regions of Figure 2a.
We conclude that there were other important differences
between the air masses with low and high TNOy and Sun
angles were not the determining factor for the variability of
ClOx reported here.
[27] In Figures 2b and 2c, respectively, we show binned
averages of FSSP surface area and total water (uncorrected
for inlet enhancements to highlight the effects of particles)
with TNOy. What clearly distinguishes the TNOy < 900 ppt
region from that at higher values are the large amounts of
H2O- and HNO3-containing particles. By contrast, the TNOy
> 900 ppt regions exhibited low particle surface areas
(consistent with background aerosol) for the entire mission.
Not only might such high particulate loadings at low TNOy
be sufficient to activate chlorine through heterogeneous
reactions of HCl, HOCl, and ClNO3, higher abundances
of ClO in the evening relative to those in the morning
suggest that this activation may be occurring on timescales
of hours, as we discuss below. Alternatively, heterogeneous
reactions overnight could be producing less photolabile
precursors of ClOx; however, this contradicts a large number
of studies that indicate heterogeneous reactions produce
forms of chlorine that are preferentially more photolabile
than their precursors [Sander et al., 2003]. In the latter case,
we would expect to see higher sunrise abundances of ClOx
in regions laden with particles. Consequently, we conclude
that the significant activation on particles is occurring
during daylight hours. The dramatic increase in [ClO]/
[Cly] from sunrise to sunset in the lowest TNOy bins as
seen in Figure 2d support this. We examine this daylight
activation in more detail below.
[28] At higher TNOy, where total particle surface areas are
much smaller, there is less ClOx in the evening than in the
morning. First, in the absence of rapid heterogeneous
chemistry NOx may be released from photolabile NOy
reservoirs (e.g., NO3, N2O5 and HO2NO2) during daylight.
In addition, slow heterogeneous reactions occurring while
the air parcels are in darkness could be activating a small
amount of HCl, HOCl, and ClNO3 into a more photolabile
form, such as Cl2, that produces a ‘‘surge’’ of ClOx at
sunrise that is subsequently reduced later in the day when
NOx is slowly released from its reservoirs. This behavior
has been seen before at much higher altitudes in regions rich
in ClNO3 and at similar temperatures [Pierson et al., 1999].
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[29] Thornton et al. [2003] have noted that at these lowSun conditions, there may not necessarily be a direct
correlation between particle surface areas and ClOx abundances as simulated at lower latitudes by some models
[Borrmann et al., 1996; Solomon et al., 1997]. This is
because the timescales for chlorine activation by heterogeneous reactions and deactivation by release of NOx from
more stable forms are decoupled. Consequently, at these
high latitudes the presence alone of enhanced chlorine does
not necessarily imply recent activation. However, signatures
of activation and deactivation may be detectable in ensembles of observations, which is essentially what the averages
reported here represent. In this case, we might expect to see
a general increase in reactive chlorine abundances in wet,
particle-rich regions if the air entering those regions has
previously been much warmer and ‘‘drier’’ (i.e., temperatures are dropping from above the frost point). The
opposite would be true in warm and dry regions, where
reactive chlorine should decrease as NOx is released from
longer-lived forms of NOy, and ClNO3 is formed [King et
al., 1991; Toohey et al., 1993].
[30] On particles, HCl is oxidized by ClNO3 or HOCl to
produce Cl2, a highly photolabile species that serves as a
source of ClO in the presence of ozone and sunlight. NOx is
also converted to HNO3, such that HO2 becomes enhanced,
thereby increasing the rate of HOCl production from HO2 +
ClO [Hanisco et al., 2002]. Production of HOCl further
amplifies the rate of heterogeneous conversion of HCl to
active chlorine. As shown below, this process serves to
catalytically produce ClOx in daylight in the presence of
sufficient particle surface area densities (schematically
shown in Figure 3). HO2 is maintained primarily by
production via the reaction of O(1D) with water.
ðR1Þ
HCl þ HOCl ! Cl2 þ H2 O
ðhet:Þ
ðR2Þ
Cl2 þ hv ! 2Cl
ðR3Þ
2ðCl þ O3 ! ClO þ O2 Þ
ðR4Þ
HO2 þ ClO ! HOCl þ O2
ðR5Þ
net : HCl þ HO2 þ 2O3 þ hv ! ClO þ H2 O þ 3O2
[31] In the absence of rapid heterogeneous chemistry,
photochemical production of NO x from NO 3 , N 2 O 5 ,
HO2NO2, and HNO3 converts active chlorine into ClNO3.
The rate of production of NOx should, in some measure,
depend on the total abundance of NOy, which comprises
primarily HNO3 at these latitudes and seasons [Ballenthin et
al., 2003]. Thus the behavior of ClO in the presence of
particles also ought to differ from that in the absence of
particles. The behavior shown in Figure 2a is consistent
with this expectation; average ClO abundances at sunset are
larger than those at sunrise in particle-rich regions of the
atmosphere, whereas the opposite is the case in regions that
are particle-poor.
[32] That the ClO behavior can be described by the
mechanisms outlined above does not prove that heterogeneous reactions are responsible for the increase in ClO
observed at sunset relative to sunrise, nor that production
of NOx in the absence of particles is responsible for the
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Figure 3. Schematic diagram of chlorine activation on particulate surfaces as described in text.
decrease from sunrise to sunset that is observed at higher
TNOy. A better case could be made by contrasting air
masses with differing particulate, but nearly identical photochemical, histories. Unfortunately, the majority of measurements at low TNOy (<600 ppt) contained large numbers
of particles, such that we do not have a sufficient sample of
‘‘low particle’’ cases for statistical significance. Specifically,
where TNOy was less than 600 ppt, particulate enhancements were detected 40% of the time by the NOy instrument
and 68% of the time by the particulate water instrument.
Conversely, for TNOy between 1000 ppt and 2000 ppt there
were particle enhancements only 2% of the time (based on
measurements from both instruments). Consequently, in this
study we assume the behaviors of particle-free and particlerich air masses are reflected by the entirety of high-TNOy
and low-TNOy air masses, respectively.
[33] In Figure 4, we plot the fraction of Cly (assuming a
composition of HCl, ClNO3, and HOCl) that could be
Figure 4. Calculation of the fraction of HCl that would be converted into active forms of chlorine by
heterogeneous reactions over 5 hours for surface areas as measured by FSSP. The data set for Figure 4 is
the same as that shown in Figure 1b, except without averaging in order to highlight the range of values
expected based on the observed variability of surface area rather than an average value. Above 1000 ppt
of TNOy, activation is minimal, even with a maximum assumed gamma value (e.g., low temperature),
whereas a significant fraction of points at TNOy values below 800 ppt exhibit extensive activation, even
for gamma values at the low range (i.e., high temperature) of possible values.
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converted into active forms by heterogeneous processes
such as those outlined schematically in Figure 4. For these
calculations, we use a Langmuir-type expression
ðR6Þ
k ¼ ðg n SÞ=4;
where g is reaction probability, n is molecular velocity, and
S is surface area (from FSSP)
ðR7Þ
½HClt =½HCl0 ¼ et=ð1=kÞð1=t * 3600Þ ;
where t is time (hr) (3600 s hr1) with n = 200 m s1. We
assume that the reaction occurs over a 5-hour period, typical
for a late winter’s day at these high latitudes. Each point in
Figure 4 represents a calculation for an individual 40-s
average of FSSP surface areas. We have assumed surface
reaction probabilities of 0.2 and 0.01, representing the
maximum and minimum values that we might expect for the
reaction of HCl with ClNO3 on surfaces that are most likely
to prevail in this region. The higher value is that
recommended for reaction of HCl with ClNO3 on nitric
acid ice over the temperature range 185 – 210 K [Sander et
al., 2003], whereas the lower value of 0.01 is that estimated
for background aerosol at 200 hPa and T = 205– 210 K
[WMO, 1999]. (Meilinger et al. [2005] have reported a
slightly higher g of 0.3 for reactions on ice particles.
Kärcher and Solomon [1999] noted that the lowering the
temperature from 210 K to 205 K increased the g for HCl
with ClNO3 on liquid aerosols in subvisible tropopause
cirrus from 0.014 to 0.25. As seen in Figure 1a, this
temperature range brackets the upper edge of the high
[ClO]/[Cly] region at T < 208 K.) The reaction probability
for HCl + HOCl is comparable to that for HCl + ClNO3, if
not larger. For example, several studies have examined the
HOCl + HCl reaction on water ice surfaces [Hanson and
Ravishankara, 1992; Abbatt and Molina, 1992; Chu et al.,
1993; Chu and Chu, 1999]. Sander et al. [2003] recommend
an average g = 0.2 with an uncertainty factor of 2 because
of surface porosity uncertainties and an uncertain temperature dependence.
[34] These calculations indicate that even for the smallest
reaction probabilities, at low TNOy values the particulate
surface areas were sufficient to convert a significant fraction
of inorganic chlorine into reactive forms within a few hours.
Under these conditions, the extent of activation of chlorine
may ultimately be limited by the availability of one reactant.
At high TNOy, surface areas were insufficient to activate
more than a few percent of available chlorine even assuming the largest reaction probabilities (that are probably
unrealistic). Although not shown here, the average FSSP
surface areas measurements were 45 mm2 cm3, (with
a few 40-s average values approaching 500 mm2 cm3)
when TNOy < 600 ppt, but surface areas were only
0.1 mm2 cm3 for TNOy > 1000 ppt. It is also important
to consider that under conditions where condensed water
was detected by the TOTCAP instrument (e.g., low TNOy),
the FSSP measurements of surface area are lower limits
[Hallar et al., 2004]. Thus this simple calculation supports
the notion that the different diurnal behaviors of ClOx at
high and low TNOy are due to differences in the surface
areas of particles more so than to differences in the
reactivities of those particles.
D22304
[35] The calculations shown in Figure 4 strongly support
the notion that the [ClO]/[Cly] ratio can be dramatically
increased within a single day in late winter at high latitudes
at observed particulate abundances near the tropopause,
regardless of the nature of the particles. At higher TNOy
(e.g., under lower particulate loadings), the calculations
suggest that approximately days to approximately weeks
are required for similar levels of activation. Under springtime conditions that favor photochemical release of NOx,
surface areas at higher TNOy may not be sufficient to
activate chlorine. At these low altitudes, where reactive
chlorine abundances are a few tens of ppt, release of NOx by
photolysis of 1 – 2 ppb of HNO3 occurs at a rate of
approximately ppt per hour, which is much faster than
activation when surface areas are small (<0.1 mm2 cm3),
as generally observed in the high-TNOy regions.
[36] The DC-8 flight paths during SOLVE/THESEO
2000 tended to favor regions of potential cirrus formation,
i.e., regions of uplift and particle formation. It is interesting
to note that many of the largest localized ClO enhancements
were seen in regions likely to be influenced by the Icelandic
Low, a large semipermanent low-pressure system between
Greenland and Scandinavia. The Icelandic Low represents
the northern half of the North Atlantic Oscillation (NAO), a
large-scale distribution of atmospheric mass over the North
Atlantic. It has been well documented that the ‘‘positive
phase’’ of the NAO is consistent with increased westerly
flow of moist air across the North Atlantic and increased
cyclonic activity in the Icelandic Low region. During the
SOLVE/THESEO 2000 winter, the NAO was in a strongly
positive phase, stronger in fact than all but ten winters since
the 1860s. It is conceivable that the Icelandic Low, by
favoring cyclogenesis and uplift, increased the available
water, and hence the particulate surface area density, in the
tropopause region heavily sampled by the DC-8 during
SOLVE/THESEO 2000 [Thornton, 2004]. Increased ClO
associated with frontal and cyclonic activity was seen in the
model results of Bregman et al. [2002]. Such stratospheretroposphere exchange in a North Atlantic cyclone has been
described by Cooper et al. [2002]. We hope to provide a
more detailed analysis of this possible connection between
enhanced near tropopause ClO and regions of intense
cyclonic activity in a forthcoming paper. We speculate that
areas of enhancements to the [ClO]/[Cly] ratio may be found
well outside the polar regions sampled during SOLVE/
THESEO 2000 in regions of similar air mass characteristics.
4. Conclusions
[37] Measurements of ClO, NOy (gas phase and particulate), particulate surface area, and particulate water near the
tropopause at high northern latitudes during SOLVE/THESEO 2000 suggest that enhancements of the active chlorine
in this region of the atmosphere can result from exposure of
air to water- and nitric-acid-containing particles at cold
temperatures (<208 K). Throughout the SOLVE/THESEO
2000 campaign, regions of high total (gas + particulate)
NOy were characterized by low particle surface area densities, whereas high particle surface area densities prevailed in
regions of low total NOy. The daytime behavior of ClOx is
consistent with a mechanism where heterogeneous activation of chlorine dominates in regions of the atmosphere that
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are particle-rich, and deactivation by photolytically released
NOx prevails in regions that are particle-poor. In the drier
regions, somewhat larger values for ClOx in the morning
than in the afternoon may indicate that slow heterogeneous
conversion is occurring during long periods of polar darkness and that active forms of chlorine so produced are then
converted back into reservoirs during the day.
[38] The [ClO]/[Cly] ratio was largest in the coldest and
wettest air masses corresponding to the Arctic tropopause (as
defined by ozone mixing ratios). However, in this region, the
uncertainty in Cly due to the poorly quantified inorganic
source contribution makes it difficult to determine the extent
of chlorine activation. Using a revised estimate for Cly that is
based on recent high-precision measurements of HCl nearer
to where air enters the stratosphere (the tropical tropopause)
reported by Marcy et al. [2004], results in a fairly substantial
fraction of available inorganic chlorine in reactive form,
20 – 25%. Such production of active chlorine can be
explained on the basis of known heterogeneous chemical
reactions on particles at the levels of surface area measured
simultaneously. These results appear to help reconcile differences in observations of ClO reported by Thornton et al.
[2003] and Smith et al. [2001]. The key difference between
the two sets of observations is the presence or lack of
particles in air masses that can be considered lower stratospheric on the basis of ozone abundances. In the winter polar
region, high humidities and particle densities were observed
often in the first few km above the tropopause, whereas in
the latter study, the observed tropopause region was characterized by low relative humidities.
[39] The largest values for [ClO]/[Cly] reported here are
smaller than the 0.5 predicted by Borrmann et al. [1996]
and Solomon et al. [1997], unless the lowest intercept for
Cly is used, in which case they are comparable. However,
the ClO mixing ratios observed are in good agreement with
the enhancements calculated by Bregman et al. [2002] that
are scattered about the region from 50N to 65N. Improved
measurements of ClO and estimates of Cly will be necessary
to precisely determine the extent of chlorine activation in
the UTLS. Given that relatively high active chlorine abundances were detected near the tropopause throughout
SOLVE/THESEO 2000, further studies of the altitude,
latitude and seasonal dependences of Cly, HCl, ClO, HOx,
ozone, particles, and water will help to elucidate the
mechanisms responsible for halogen activation and improve
our understanding of ozone chemistry in the UTLS.
[40] Acknowledgments. This work was funded by the NASA Upper
Atmosphere Research Program grant number NAG2-1307 and by the
National Science Foundation grant number ATM-9732909. We would like
to thank the NASA Dryden Flight Research Center DC-8 pilots, flight and
ground crews, and mission support personnel too numerous to mention by
name. NAO Index Data were provided by J.W. Hurrell of the Climate
Analysis Section, NCAR, Boulder, USA (1995); index data are available
online at http://www.cgd.ucar.edu/cas/jhurrell/indices.html. D.W.T.
acknowledges support from the University of Colorado for a temporary
leave to carry out this work.
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