SOLA, 2006, Vol. 2, 013 016, doi:10.2151/sola.2006 004
13
Atmospheric Aerosol Deposition on Snow Surfaces
and Its Effect on Albedo
Teruo Aoki1, Hiroki Motoyoshi2, Yuji Kodama3, Teppei J. Yasunari3,
Konosuke Sugiura4, and Hiroshi Kobayashi5
1
Meteorological Research Institute, Tsukuba, Japan
2
The Graduate University for Advanced Studies, Tokyo, Japan
3
Hokkaido University, Sapporo, Japan
4
Institute of Observational Research for Global Change, JAMSTEC, Yokohama, Japan
5
University of Yamanashi, Kofu, Japan
This paper investigated how atmospheric aerosol
deposition into snowpack affects snow albedo reduction
using radiation budget observation, atmospheric aerosol
monitoring, and snow pit work during the winter of
2003/2004 in Sapporo, Japan. The mass concentration
of snow impurities was less than 10 parts per million by
weight (ppmw) in the core accumulation season and
exceeded 100 ppmw in the melting season due to a
heavy dust event on March 11 to 12, 2004. The relationship between the visible albedo and snow impurities
suggested the considerable effect of snow impurities in
reducing the visible albedo even during the accumulation season in Sapporo. After the dust event, high
impurity concentrations of approximately 600 ppmw in
the top snow layer were maintained, but the visible
albedo still decreased because the increasing snow grain
size continued to reduce the visible albedo. The size distribution of snow impurities measured from a snow
sample with a Coulter counter during the dust event
was compared with the size distributions of snow impurities calculated as wet and dry depositions from a laser
optical particle counter for atmospheric aerosols. The
result confirmed that the contribution from wet depositions was important except for giant particles with a
radius larger than 2.5 µm.
1. Introduction
The important factors affecting albedo of snow with
optically sufficient depth are snow grain size, snow impurities (water-insoluble solid particles), solar zenith
angle, and cloud cover (Warren 1983). The effects of
snow impurities, snow grain size, and snow aging on
broadband albedos were identified from long-term radiation budget measurements on the snow surface and
snow pit work (Aoki et al. 2003). The atmospheric
effects such as cloud cover and effects of geometric conditions on albedo were investigated by Aoki et al. (1999)
using the radiative transfer model for the atmospheresnow system. The main cause of visible albedo reduction is snow impurities originating from atmospheric
aerosols (Warren and Wiscombe 1980; Aoki et al. 2000).
It was recently reported that absorptive anthropogenic
aerosols containing black carbon could reduce snow
albedo and thus a possible reason for warming in the
Arctic (Hansen and Nazarenko 2004). Possible snow
albedo reduction by black carbon contamination was
revealed by radiation measurements on the snow
surface at Barrow, Alaska, (Aoki et al. 1998) and in
Japanese urban areas (Motoyoshi et al. 2005). Natural inCorresponding author: Teruo Aoki, Meteorological Research
Institute, 1-1 Nagamine, Tsukuba, Ibaraki 305-0052, Japan. Email: [email protected]. ©2006, the Meteorological Society
of Japan.
soluble aerosols such as mineral dust are also main
snow impurities during the melting season due to the
effect of Asian dust.
To accurately simulate future climate conditions of
the Earth, especially for the Arctic, a physically based
snow albedo model that calculates the effect of snow
impurities on albedo in the land surface process is necessary. For this purpose, the deposition of atmospheric
aerosols into snow should be investigated in field experiments. We conducted continuous radiation budget
measurement and atmospheric aerosol monitoring, as
well as frequent snow pit work, during the winter of
2003/2004 in Sapporo, Japan. The albedo reduction by
aerosol deposition is discussed based on these data. We
also focus on the wet and dry depositions of dust particles that were transported from the Chinese desert into
the snow during the heavy dust event of March 11 to
12, 2004.
2. Observation site and instruments
Field measurements were carried out from December
2003 to March 2004 at the meteorological observation
field (43°04´56´´N, 141°20´30´´E) of the Institute of Low
Temperature Science, Hokkaido University in Sapporo,
Japan, which is located in an urban area of Sapporo
City. In the radiation budget observation, upward and
downward radiant flux densities were measured with
pyranometers (CM-21, Kipp & Zonen) in the shortwave
(=0.305 2.8 µm) and the near-infrared regions (=0.715
2.8 µm). Flux densities in the visible region (=0.305
0.715 µm) were calculated as the differences between the
flux densities for the former two spectral regions. Each
radiation component was sampled every ten seconds,
and one-minute-averaged values were stored in a data
logger. For analyses of albedos, 30-minute averaged
values were used in this study.
The number density of the atmospheric aerosols was
measured with a laser optical particle counter (OPC: TD100, SIGMATEC Inc.). The instrument has six channels,
counting aerosol particles with radii of > 0.15, 0.25, 0.5,
1, 1.5, and 2.5 µm. Sample air was taken at a flow rate of
0.001 m3/min, and the number of aerosol particles accumulated in each bin during 1 minute were recorded. In
this study, 30-minute averaged values with a correction
for attenuation by tube were analyzed. Snow pit work
was performed twice a week to determine snow type,
temperature, density, and snow grain size in each snow
layer, together with snow sampling for measuring snow
impurities. The snow impurities were filtered using
Nuclepore filters, and the concentrations of snow impurities were estimated by weighing the Nuclepore filters
with a balance. The size distribution of snow impurities
was measured with a Coulter counter (CC: Multisizer 3,
Beckman Coulter Inc.) for radii ranging from 0.7 to 21
µm.
14
Aoki et al., Aerosol on Snow Surfaces and Albedo
Fig. 1. Visible (VIS), shortwave (VIS+NIR), and near infrared
(NIR) albedos averaged from 1131 to 1200 LT and snow depth
at 1200LT during the winter of 2003/2004 in Sapporo, Japan.
Fig. 2. Mass concentrations of snow impurities measured from
snow samples and snow grain size at the snow surface
measured from snow pit work. Red circles indicate the snow
sampling layer 0 to 2 cm, and blue circles denote 0 to 10 cm.
Definition of snow grain size is the same as r2 in Aoki et al.
(2003).
3. Results and discussion
3.1 Relationship between broadband albedo and snow impurities throughout the winter
The snow period in the winter of 2003/2004 in
Sapporo was from the beginning of December to the end
of March, for which the broadband albedos averaged
from 1131 to 1200 LT close to local solar noon for the
visible, near-infrared, and shortwave wavelength
regions, and snow depths are shown in Fig. 1. During
this period, the solar zenith angle varies from 66.6 to
40.8° at local solar noon. Albedos remained stable from
the end of December to the end of February (accumulation season) and then decreased in March (melting
season). During the accumulation season, the short-term
variation of the near-infrared albedo exceeded that of
the visible albedo. These variations are due mainly to
the effects of snow physical parameters and cloud cover.
The primary snow physical parameters affecting the
albedo are snow impurities (mainly for the visible
albedo) and snow grain size (for the near-infrared
albedos) (Wiscombe and Warren 1980; Warren and
Wiscombe 1980). The effect of cloud cover on snow
albedo was theoretically investigated by Aoki et al.
(1999). In that investigation, the shortwave albedo was
increased by 0.06 by cloud cover for a snow grain radius
of 50 µm at a solar zenith angle of 50°. Since the measured shortwave albedo fluctuation frequently exceeds
the cloud effect (0.06) during the accumulation season,
the short-term albedo variations should contain the
effects of snow physical parameters. In particular, the
near-infrared albedo fluctuation should be affected by
the variation of snow grain size, as will be shown in Fig.
2.
Fig. 3. Relationship between mass concentration of snow impurities and visible albedos averaged from 1131 to 1200 LT at
each day. Meanings of color for data points are the same as in
Fig. 2.
Figure 2 presents the mass concentration of snow
impurities for different snow layers and the snow grain
size measured from snow pit work during the winter.
The measured snow grain size is defined in the same
way as r2 in Aoki et al. (2003), which is one-half the
branch width of dendrites or one-half the dimension of
the narrower portion of broken crystals. This dimension
corresponds with the optically equivalent snow grain
radius (Aoki et al. 1998, 2000). The values of snow impurities for the top layer (0 to 2 cm) were generally higher
than those for the layer of 0 to 10 cm because the snow
impurities are supplied by the deposition of atmospheric
aerosols as suggested by Aoki et al. (2000). During the
core accumulation season (January and February), both
concentrations were less than 10 ppmw. They began to
increase from early March and suddenly exceeded 100
ppmw. This abrupt rise in concentration was due to the
heavy dust event on March 11 to 12. After the dust
event, the high values of mass concentration of around
600 ppmw in the top layer remained, but the visible
albedo shown in Fig. 1 still decreased. This could have
been caused by the effects of snow depth, liquid water
in the snow, and snow grain size. The liquid-equivalent
depth of snow for which snowpack becomes optically
semi-infinite is 20 cm for a snow grain radius of 1000
µm (Wiscombe and Warren 1980). Since the measured
liquid-equivalent depth was about 18 cm on March 26,
the effect of snow depth seems to appear in the last
several days of March. Increasing liquid water would
increase the optically equivalent snow grain size
because water is optically very similar to ice in the
visible region (Wiscombe and Warren 1980). Visible
albedo reduction basically depends on snow impurities,
but the reduction rate per unit mass concentration
depends on snow grain size (Warren and Wiscombe
1980). The measured snow grain size increased after the
dust event (Fig. 2), and the visible albedo decreased
rapidly (Fig. 1). The visible albedo reduction after the
dust event for the period with enough snow depth
would thus be due to the effect of increased snow grain
size, where liquid water would also enhance this effect.
The relationships between the visible albedo and
snow impurities for both sampling layers are illustrated
in Fig. 3, which clearly demonstrates the snow impurity
dependence of the visible albedo. The data for albedos of
less than 0.7 were observed during and after the dust
event, and high albedos were recorded in the accumulation season, as depicted in Fig. 1. The highest visible
albedo is about 0.92, which is lower than that at Kitami
(∼0.98), Japan (Aoki et al. 2003). Since the measured
concentrations of impurities for the highest albedos at
Sapporo (3 to 4 ppmw) are roughly in the same range as
those at Kitami (Fig. 9 in Aoki et al. 2003), the result in
Fig. 3 suggests the considerable effect of snow
SOLA, 2006, Vol. 2, 013 016, doi:10.2151/sola.2006 004
Fig. 4. Aerosol number densities for four channels that counted
particles with radii larger than 0.15, 0.25, 1.0, and 2.5 µm from
March 9 to 13, 2004. A heavy dust event was observed on
March 11 to 12. A to C indicate the periods used to estimate the
dry and wet depositions of aerosols. Colored periods with
weather symbols mean snowfall (blue), rainfall (green), and
sleet (light green).
impurities in reducing the visible albedo even during
the accumulation season in Sapporo because the cloud
effect on the visible albedo is less than that for the
shortwave albedo mentioned at Fig. 1.
3.2 Heavy dust event on March 11 to 12, 2004
A dust event was observed on March 11 to 12, 2004.
Figure 4 presents the aerosol number density measured
with an OPC for the four selected channels on five days
including the dust event. Dust variation is clearly
evident for particles larger than 2.5 µm. When rainfall
was observed (at around 0500 LT on March 11) just
before the dust event, all aerosol counts temporarily decreased. After that, dry deposition of dust began, and
the number density of particles larger than 2.5 µm
exceeded 106 particles/m3 at around local noon. After
the peak in the aerosol density, snowfalls were identified from a rain gauge record and radiation record
(1530LT on March 11 to 0230LT on March 12 and 0700
to 0800LT on March 12). During snowfall periods, the
number density of particles larger than 2.5 µm decreased.
Snow pit work with snow sampling conducted from
1010 to 1120LT on March 12 revealed that the new
snowfall during the dust event was 3 cm in depth, and
this snow layer contained a large amount of dust particles. This means the dust had fallen by wet deposition
together with the snowfall. Snow in the top layer was
sampled for this new dust-containing snow layer. The
mass concentration of snow impurities in the snow
sample was 583.3 ppmw. Broadband albedo variations
corresponding to the dust event are shown in Fig. 5. The
increased near-infrared albedo with the snowfall on
March 11 to 12 is unnatural. We checked the data on
March 11 to 12 and found that the measured downward
flux decreased from 1530LT on March 11 compared
with the variation of upward flux. The most probable
cause is that the glass dome of the near-infrared
pyranometer was covered with snow. Measured albedos
after local solar noon on March 12 should be correct.
The visible albedo was approximately 0.7 on March 10
before the dust event and decreased to 0.5 in the early
afternoon of March 12 at the end of dust event.
3.3 Deposited dust particle size distribution
The size distribution of snow impurity particles
(mainly dust) was measured for the snow sample during
the dust event with a CC in the radius range from 0.7 to
21 µm. This snow sample was collected from the snow
layer that had fallen during the dust event, so it should
15
Fig. 5. Half-hourly mean albedos in the visible (VIS), shortwave
(VIS+NIR), and near-infrared (NIR) spectrum from March 9 to
13, 2004.
contain the dry and wet depositions of dust particles
from the beginning of the first snowfall (1530LT on
March 11) to the snow sampling time (1120LT on March
12). During the dust event, atmospheric aerosols were
measured with an OPC, as shown in Fig. 4. Using these
data, we estimated how the size distributions of dry and
wet depositions of dust particles contributed to the collected snow sample. Although a CC counts only waterinsoluble solid particles and an OPC measures all
particles including water-soluble ones, the amount of
water-soluble aerosols contained in giant particles with
radius > 1.0 µm would be relatively less during the dust
event. We focused on the size distribution of giant particles. When the aerosol size distribution is calculated
from the OPC data with six channels, the distributions
for the five smaller size bins can be calculated as the differences between adjacent channels. However, the
largest OPC channel is for radius > 2.5 µm. We assumed
the upper boundary of this channel to be 5.0 µm. Even if
the dust size distribution followed the coagulation mode
of OPAC mineral dust (log-normal with mode radius
rmod=1.90 µm and width of the distribution =2.15, Hess
et al. 1998), the number of dust particles contained in
the radius range of 2.5 to 5.0 µm is 97.5 % of the total
dust particles larger than 2.5 µm.
The dry deposition was calculated by accumulating
the aerosol flux from the beginning of snowfall during
the dust event to snow sampling time (period-A in Fig.
4), during which the aerosol flux was calculated by the
aerosol number density measured with an OPC and
terminal velocity as a function of particle size. We considered the wet deposition to be supplied from the total
aerosol particles in the air column below the snowfall
layer in the atmosphere. The aerosol number density
measured with an OPC during snowfall therefore does
not contribute to the wet deposition of aerosols. We thus
assumed the aerosol number density contributing to
wet deposition to be the difference between averaged
aerosol densities measured with an OPC during the nonsnowfall period (period-B in Fig. 4) and the snowfall
period (period-C in Fig. 4). For the former period, we
employed the time from the beginning of the dust event,
not from the beginning of the snowfall, because the
aerosol density before snowfall is important for wet
deposition. The atmospheric aerosol layer from which
particles were to be deposited was assumed to be from
the surface to heights of 1km, 2km, and 3km.
Figure 6 compares the calculated number size distributions of deposited aerosols with the measured one.
Calculated dry depositions, except for OPC size bin-5
and bin-6, are lower than any wet depositions due to the
low terminal velocity of small particles. Wet deposition
is naturally sensitive to the air thickness of the aerosol
deposition layer. In size bins 4 to 6, total aerosol density
16
Aoki et al., Aerosol on Snow Surfaces and Albedo
Fig. 6. Number size distribution of snow impurities (plus
symbols) in the snow sample collected on March 12, 2004,
during the dust event, measured with a CC. The dry and wet
depositions of aerosols were estimated from OPC data for each
size bin. Colored lines represent the depositions of dry (blue),
wet (green), and “dry + wet” (red). For the wet deposition, the
levels containing aerosols to be deposited were assumed to be
from the surface to heights of 1 km, 2 km, and 3 km; for “dry +
wet” deposition, only the result for the surface to 3km is shown.
Units are the number of particles in a snow column of 1.0 m2
area with 3 cm depth (snowfall depth) for the snow sample and
the number of particles that fall onto a snow surface of 1.0 m2
for aerosol depositions.
Fig. 7. Cloud-top height during the dust event estimated from
brightness temperatures of IR channel on board the GOES-9
satellite and corresponding GPV data provided from the JMA
at the grid point nearest to Sapporo.
results suggested that the concentration of snow impurities could be treated as a predictor in a physically
based snow albedo model in the land-surface process by
calculating the aerosol transport process in general circulation models.
Acknowledgments
determined by “Dry + Wet (3 km)” agrees well with that
for the snow sample, where the contribution from wet
deposition is important for bin-4 and bin-5, and the contribution of dry deposition exceeds that of wet deposition (3 km) for the largest bin (size bin-6). In size bin-3,
however, the calculated aerosol density is lower than
the measured one. A CC counts only water-insoluble
solid particles, whereas an OPC measures all types of
particles. This result is not consistent with the disagreement in bin-3. One possible reason is that aggregated
dust particles in the atmosphere broke into smaller particles suspended in melted snow water when measured
with a CC.
We examined the cloud-top height to validate the appropriateness of the value of 3 km for air thickness for
wet deposition. Figure 7 depicts the cloud-top height estimated from the satellite data and grid point values
(GPV) of numerical weather prediction data calculated
by the Japan Meteorological Agency (JMA) at the grid
point nearest to Sapporo during the dust event. The
observed cloud-top height was approximately 3 km
during the snowfall periods, consistent with our
assumed air thickness.
4. Conclusion
We found that the effect of the atmospheric aerosol
deposition into the snowpack on snow albedo reduction
is important for snow albedo variation in Sapporo,
where absorptive atmospheric aerosols are frequently
deposited into the snow. For the melting season in particular, the visible albedo was efficiently reduced by the
reduction in the visible albedo with increasing snow
grain size even if snow impurity concentration did not
change. Liquid water contained in snow would also
enhance this effect by increasing the optically equivalent snow grain size.
The size distribution of snow impurities in the snow
sample collected during the dust event was estimated
by a simple model for the dry and wet depositions of atmospheric aerosols. Comparing the measured size distribution with the calculated one revealed that the
contribution from wet depositions was important,
except for giant particles with radii > 2.5 µm. These
We would like to thank Keiko Konya, Syosaku
Kanamori, Gaku Yamazaki, Nobuyuki Imanishi, Shinya
Shishido, Nobuyuki Tsubonuma, and Reiko Kurosawa
at the Institute of Low Temperature Science of
Hokkaido University for snow pit work throughout the
winter. This work was conducted as part of the ADEOSII/GLI Cal/Val experiment supported by the Japan
Aerospace Exploration Agency, and the Experimental
Research Fund for Global Environment Conservation
supported by the Ministry of the Environment of Japan.
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Manuscript received 31 October 2005, accepted 17 December 2005
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