Ocean Modelling 71 (2013) 43–53
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Ocean Modelling
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Linkages among halocline variability, shelf-basin interaction, and wind
regimes in the Beaufort Sea demonstrated in pan-Arctic Ocean modeling
framework
Eiji Watanabe ⇑
Japan Agency for Marine-Earth Science and Technology, 3173-25, Showa-machi, Kanazawa-ku, Yokohama, Kanagawa 236-0001, Japan
a r t i c l e
i n f o
Article history:
Available online 5 January 2013
Keywords:
Halocline layer
Shelf-break process
Anti-cyclonic/cyclonic wind patterns
Pacific-origin summer and winter water
Beaufort Sea
a b s t r a c t
To address the mechanisms controlling halocline variability in the Beaufort Sea, the relationship between
halocline shoaling/deepening and surface wind fields on seasonal to decadal timescales was investigated
in a numerical experiment. Results from a pan-Arctic coupled sea ice-ocean model demonstrate reasonable performances for interannual and decadal variations in summer sea ice extent in the entire Arctic
and in freshwater content in the Canada Basin. Shelf-basin interaction associated with Pacific summer
and winter transport depends on basin-scale wind patterns and can have a significant influence on halocline variability in the southern Beaufort Sea. The eastward transport of fresh Pacific summer water
along the northern Alaskan coast and Ekman downwelling north of the shelf break are commonly
enhanced by cyclonic wind in the Canada Basin. On the other hand, basin-wide anti-cyclonic wind
induces Ekman upwelling and blocks the eastward current in the Beaufort shelf-break region. Halocline
shoaling/deepening due to shelf-water transport and surface Ekman forcing consequently occur in the
same direction. North of the Barrow Canyon mouth, the springtime down-canyon transport of Pacific
winter water, which forms by sea ice production in the Alaskan coastal polynya, thickens the halocline
layer. The model result indicates that the penetration of Pacific winter water prevents the local upwelling
of underlying basin water to the surface layer, especially in basin-scale anti-cyclonic wind periods.
Ó 2012 Elsevier Ltd. All rights reserved.
1. Introduction
Halocline variability is important for both physical and biogeochemical aspects in the western Arctic Ocean. Carmack et al. (2008)
described the details of the stratified structure of the Canada Basin,
which consists of a seasonal mixed layer in the surface 40 m, halocline layers from 40 to 200 m, and the Atlantic layer below
200 m. Sea ice meltwater and meteorological water from precipitation and river water discharge are mostly included in the surface
mixed layer. The mixed and halocline layers in the western Arctic
are both substantially occupied by Pacific-origin summer and winter water masses.
The warm Pacific summer water intrudes into the upper portion
of halocline in the southern Canada Basin, and the corresponding
ocean heat transport has the potential to eventually reduce net
thermal sea ice production over its pathway (Shimada et al.,
2006). The shelf-to-basin transport of cold, nutrient-rich Pacific
winter water into the lower part of halocline layer is suggested
to suppress vertical heat transfer from the underlying Atlantic
layer and to be a source of primary phytoplankton production in
the Canada Basin (Kadko et al., 2008). Nishino et al. (2011) indi⇑ Tel.: +81 45 778 5675.
E-mail address: [email protected]
1463-5003/$ - see front matter Ó 2012 Elsevier Ltd. All rights reserved.
http://dx.doi.org/10.1016/j.ocemod.2012.12.010
cated that freshwater accumulation in the surface Ekman layer
limited nutrient supply from the underlying halocline layers and
hence inhibited phytoplankton growth within the Beaufort Gyre,
which is an anti-cyclonic circulation in the Canada Basin. On the
other hand, the shallower halocline with higher nitrate concentration in the surface euphotic zone maintains the biological productivity outside the Beaufort Gyre. These previous studies proposed
that halocline depth was a key index of various types of sea iceocean dynamics and marine ecosystem in the western Arctic
Ocean.
It is pointed out that halocline variability with the vertical shifting of isopycnal layers in the Canada Basin is significantly controlled by wind stress fields on seasonal, interannual, and
decadal timescales (Proshutinsky et al., 2009). Anti-cyclonic wind
circulation produces Ekman convergence in the upper ocean, the
deepening of halocline layers, and a consequent increase in the
freshwater content of water column, whereas cyclonic wind circulation induces freshwater release accompanied by the upwelling of
halocline layers (Proshutinsky et al., 2002). The upwelling flow can
transport nutrients and plankton from the halocline layers to the
surface euphotic zone (Fiechter and Moore, 2009; Mundy et al.,
2009). However, the extent to which wind forcing directly and
indirectly accounts for halocline variability in the Arctic region remains uncertain. The magnitude and phase of variations would be
44
E. Watanabe / Ocean Modelling 71 (2013) 43–53
not always synchronized with wind forcing (e.g., Itoh et al., 2012).
Ocean circulation and fine-scale mixing processes have the potential to modify the hydrographic structure that responds to surface
Ekman forcing. A variety of halocline water masses originate from
the Chukchi and Beaufort shelf regions (Aagaard et al., 1981; Shimada et al., 2001; Steele et al., 2004). It is expected that the
shelf-basin interaction of the eastward current comprising the Pacific summer water and the penetration of Pacific winter water
have important influences on upper and lower halocline variability
north of the Beaufort shelf break, and that such shelf-break processes are also related to surface wind fields (Pickart, 2004; Shimada et al., 2005).
Yang (2006, 2009) examined the variability in wind-driven Ekman transport in the Beaufort Sea using an idealized analysis. Yang
(2006) concluded that robust anti-cyclonic wind and sea ice motion induced strong Ekman-driven coastal upwelling during autumn and winter in the southern Beaufort Sea. Yang (2009)
extended this analysis of the seasonal cycle to interannual and decadal timescales and showed significant variability in the Ekman
transport. In his method, concentration and motion data for sea
ice derived from satellite passive microwave sensors and ice-station drift buoys were processed for the estimation of Ekman transport. These data sources continue to have limitations regarding
spatial resolution and accuracy during the ice-melting season. In
addition, the absence of geostrophic ocean circulation, variable Ekman layer depth, and bottom bathymetry can generate crucial
biases in the estimated variability. Alternatively, numerical modeling can be a strong tool for comprehensive analyses.
The present study focused on halocline variability associated
with basin-scale wind fields in the Beaufort Sea using a pan-Arctic
coupled sea ice-ocean general circulation model. In particular, the
linkages of the direct impact of local Ekman forcing and the indirect wind effects via the Pacific summer and winter water transport on halocline structure were investigated. This paper is
organized as follows. Section 2 outlines the model configuration
and experimental design. Section 3 describes the model performance for major basin-scale sea ice and freshwater variations
and for western Arctic shelf processes. Shelf-break processes related to the Pacific summer and winter water transport are discussed in Section 4. Summary and discussion are presented in
Section 5.
2. Model and experimental design
The model used for the analyses is the Center for Climate System Research Ocean Component Model (COCO) version 3.4 developed at the University of Tokyo (Hasumi, 2006). The COCO model
has been widely utilized for regional and global climate studies
(Proshutinsky et al., 2011; Komuro and Hasumi, 2007; Oka et al.,
2009). The sea ice part of the model has both thermodynamic
and dynamic components. The zero-layer formulation of Semtner
(1976) is adopted for the thermodynamic part. In the dynamic part,
equations for momentum, mass, and concentration are taken from
Mellor and Kantha (1989). The internal sea ice stress is calculated
based on the elastic-viscous-plastic rheology of Hunke and Dukowicz (1997) with the ice strength parameter of 2:0 102 N m1 .
The ocean part is a free-surface ocean general circulation model
(OGCM) that incorporated the uniformly third-order polynomial
interpolation algorithm (UTOPIA) (Leonard et al., 1994) and the
quadratic upstream interpolation for convective kinematics with
estimated streaming terms (QUICKEST) (Leonard, 1979) for horizontal and vertical tracer advection, respectively. The subgrid-scale
eddy-induced transport of tracers is parameterized by the combination of the isopycnal diffusion (Cox, 1987) and isopycnal layer
thickness diffusion scheme of Gent and McWilliams (1990) (GM).
The isopycnal and GM diffusion coefficients are both fixed to the
constant value of 1:0 103 m2 s1 . In the surface mixed layer, the
turbulence closure scheme of Noh and Kim (1999) is applied for
diagnosing vertical viscosity and diffusion coefficients. The background horizontal viscosity and diffusion coefficients are
5:0 103 m2 s1 and 1:0 10 m2 s1 , respectively. The background
vertical viscosity coefficient is 1:0 104 m2 s1 . The background
vertical diffusion coefficient varied with depth from
0:1 105 m2 s1 at the top level to 3:0 105 m2 s1 at the bottom level.
The momentum exchange term at the ice-ocean interface is formulated in the quadratic form of relative velocity. The drag coefficient is set to 5:0 103 following McPhee (1980), and the rotation
angle is taken to be 25°. It is expected that wind-driven Ekman
pumping/suction should be an important part of mechanisms controlling the halocline depth. The Ekman contribution is generally
represented by the wind stress curl. In most areas of the Arctic
Ocean, sea ice-ocean stress instead of wind stress functions as a direct forcing at the ocean surface. The Ekman divergence/convergence was calculated as follows:
1
qo f
5 ss ;
where ss is the ocean surface stress represented by the sum of simulated sea ice-ocean stress in an ice-covered area and wind stress in
an open water area in each grid, qo is ocean density, and f is the
Coriolis parameter at each latitude.
The model domain contains the entire Arctic Ocean, the Greenland–Iceland–Norwegian (GIN) seas, and the North Atlantic Ocean
(Fig. 1). The bathymetry is constructed from the merged product of
the International Bathymetric Chart of the Arctic Ocean (IBCAO)
and the Earth Topography Five-Minute Gridded Elevation Dataset
(ETOPO5), which is available at the Arctic Ocean Model Intercomparison Project (AOMIP) website [http://www.whoi.edu/page.do?pid = 30630]. The horizontal resolution is one-fourth degree (about
25 km) in the rotated spherical coordinate system, and there are 28
vertical levels (Table 1). The atmospheric forcing components are
constructed from the National Centers for Environmental Prediction/National Center for Atmospheric Research (NCEP/NCAR)
reanalysis daily dataset from 1979 to 2008 (Kalnay et al., 1996).
The 13 river water discharges listed in the AOMIP website
[http://www.whoi.edu/page.do?pid = 30587] are prescribed as
surface freshwater flux at each river mouth. In the marginal region
of the model domain, except for the Bering Strait, a sponge boundary condition is applied: the horizontal diffusion coefficient is enlarged by an order of magnitude, and the temperature and
salinity at all depths are restored to the monthly mean data of
the Polar Science Center Hydrographic Climatology (PHC) 3.0
(Steele et al., 2001). The Pacific water inflow with its seasonal cycle
is provided at the Bering Strait based on the hydrographic observations of Woodgate et al. (2005c), following Watanabe and Hasumi
(2009). The specific values are prescribed to northward velocity
and salinity so that the total freshwater transport through the Ber3
ing Strait referenced to a salinity of 34.8 psu is 2755 km yr1 ,
which is almost equal to the estimation of Woodgate and Aagaard
(2005). The water temperature is kept at the freezing point from
January to June. The interannual variation in Pacific water properties at the strait is not given because there are no sufficient yearround data for decadal periods.
First, the model is spun up initialized from the PHC temperature
and salinity fields, no ocean circulation, and no sea ice for 10 years
using the atmospheric forcing in 1979. The restoring of temperature and salinity for the whole depth to the PHC monthly mean
is applied for the first five years of the spin-up stage and is turned
off, except for the marginal region of the model domain, thereafter.
E. Watanabe / Ocean Modelling 71 (2013) 43–53
45
Fig. 1. Bathymetry of the COCO model [m]. The locations include Ellesmere Island (E. I.) and Victoria Island (V. I.). Yellow lines A and B are referred to in Figs. 7 and 11,
respectively. A red cross denotes the location north of the Barrow Canyon mouth referred to in Fig. 10.
Table 1
Depth levels of the COCO model [m].
2
100
700
5
125
1000
10
150
1500
15
175
2000
20
200
2500
30
250
3000
40
300
3500
50
350
4000
60
400
80
500
The total sea ice volume in the entire model domain reaches an
equilibrium state after the 10-year spin up. The decadal experiment for 30 years is then performed using the daily atmospheric
forcing from 1979 to 2008.
The modeled total sea ice extent and current properties at the
Barrow Canyon are directly compared with satellite and mooring
measurements. The monthly sea ice extent data constructed at
the National Snow and Ice Data Center are downloaded from the
University of Colorado website (ftp://sidads.colorado.edu/DATASETS/NOAA/G02135/). The September mean values from 1979 to
2008 are picked up in Section 3.1.1. The modeled sea ice extent
is also defined as total area where sea ice concentration in each
grid is above 0.15. The Barrow Canyon mooring data deployed by
the University of Alaska Fairbanks (UAF) are obtained at the Earth
Observing Laboratory website (http://data.eol.ucar.edu/codiac/dss/
id=62.316). The data period of two moorings covers 657 days from
August 4, 2002 to May 23, 2003 and from September 15, 2003 to
September 13, 2004. The daily-mean near-bottom current direction is categorized to 12 bins every 30 °T. The total days in each
bin are addressed in Section 3.2.1.
3. Model performance for sea ice and ocean properties
3.1. Sea ice and freshwater variability in the Arctic basin
3.1.1. Sea ice state
The sea ice properties simulated by the COCO model capture the
major observed features, including the rapid sea ice decrease that
occurred in the 1990s and 2000s (Kwok and Rothrock, 2009). The
interannual variations in the September mean sea ice extent are
close to a passive microwave product of the Special Sensor Microwave/Imager (SSM/I) (Cavalieri et al., 1996) (Fig. 2a). The correlation coefficient is 0.93, although the COCO result slightly
underestimates the sea ice extent after the late 1990s. The spatial
distribution of sea ice thickness from the COCO model is consistent
with the submarine-based estimates (Belchansky et al., 2008;
Fig. 2. (a) Time series of September mean total sea ice extent from 1979 to 2008 in
(solid line) the COCO model and (dashed line) the SSM/I product [106 km2]. (b)
Spatial distribution of modeled sea ice thickness averaged from 1979 to 2008 [cm].
The white contours show the sea ice margin.
Rothrock et al., 2008), in which sea ice is the thickest along the
northern Canadian Arctic Archipelago and gradually thins toward
the Siberian shelves and Alaskan coast (Fig. 2b). On the other hand,
the model somewhat overestimates sea ice thickness in the East
Siberian Sea. A similar bias in the peripheral seas commonly appears in the AOMIP models, probably owing to the lack of fast ice
formation (Johnson et al., 2012). In addition, the unrealistic ridging
in the Siberian shelves may be partly attributed to the limited abil-
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E. Watanabe / Ocean Modelling 71 (2013) 43–53
Fig. 3. Modeled (shade) sea surface height [cm] and (purple contours) freshwater content above the 34.8 psu isohaline depth [m] averaged for the (a) 1980s, (b) 1990s, and (c)
2000s, respectively. White and black-dashed contours show bottom topography and the Beaufort Gyre Region defined in Section 3.1.2, respectively.
ity of the traditional elastic-viscous-plastic rheology of Hunke and
Dukowicz (1997) to represent the local features of sea ice drift in
the pan-Arctic Ocean model.
3.1.2. Freshwater distribution
The sea surface height (SSH) shows a dipole pattern between
the Arctic basin and the North Atlantic throughout the integration
period from 1979 to 2008 (Fig. 3a–c). The maximum height is located in the Canada Basin, where the oceanic Beaufort Gyre circulates anti-cyclonically. The SSH is widely low in the Barents, GIN,
and Labrador seas. Another maximum along the Alaskan coast in
the eastern Chukchi Sea captures the signal of the northeastward
current of Pacific-origin water. The coastal currents east of Greenland and west of the Scandinavian Peninsula are also accompanied
by higher SSH than the offshore pelagic regions. These spatial SSH
patterns reflect the underlying freshwater content (FWC). The FWC
is calculated from the following formula:
Z
0
H
Sref S
dz;
Sref
where Sref is a reference salinity set at 34.8 psu, and H is the 34.8
psu isohaline depth in each model grid. The Beaufort Gyre Region
is defined by the closed area inside 70.5°N to 80.5°N and 130°W
to 170°W and where a water depth is greater than 300 m, following
Proshutinsky et al. (2009). The COCO model demonstrates a remarkable freshening that widely occurred in the 2000s, as was suggested
by previous studies (Proshutinsky et al., 2009; Rabe et al., 2011)
(Fig. 3). The decadal mean FWC averaged in the Beaufort Gyre Region is 14.7 m in the 1980s, 15.2 m in the 1990s, and 15.9 m in
the 2000s. These amounts are slightly smaller than the observational estimates of 16.6 to 17.9 m reported by Proshutinsky et al.
(2009). A possible factor explaining this difference is the fact that
the modeled FWC maximum is located somewhat on the western
side of the Canada Basin, and the defined Beaufort Gyre Region does
not cover the entire freshwater pool. In addition, there are large
uncertainties in Arctic riverine water inflow, and the prescribed
freshwater flux of Mackenzie River discharge may cause higher
salinity bias in the southern Beaufort Sea. The isohaline depth of
33 psu, which is a key salinity of the halocline in the western Arctic
(Shimada et al., 2005), is located at 110–150 m depth under seasonal stratification in the Beaufort Gyre Region and has a deepening
trend in the 2000s in accordance with recent mooring measurements collected by the Beaufort Gyre Observational System (BGOS)
(Proshutinsky et al., 2009) (not shown). Thus, the abrupt deepening
of halocline layers is considered to largely account for the noticeable FWC increase in the 2000s, as indicated by Rabe et al. (2011).
3.2. Shelf processes in the western Arctic
The model’s ability to represent major ocean currents tracing
local bathymetry, the winter coastal polynya, and eddy-induced
transport over the western Arctic shelf region will now be
addressed.
3.2.1. Shelf circulation in the Chukchi Sea
In the COCO model, which had a grid spacing of about 25 km,
coastal and bottom topography are moderately smoothed. It is
hence difficult to resolve the intense jet streams with widths of
10–15 km along the Barrow Canyon and the Beaufort shelf break
that were explicitly demonstrated by the eddy-resolving regional
version of the COCO model (Watanabe and Hasumi, 2009; Watanabe, 2011). On the other hand, the actual transport routes of Pacific-origin summer/winter water are not confined to narrow coastal
currents and shelf-break jets. Usually, the shelf-break jets are
rather produced along the isopycnal front between shelf and basin
water masses (Pickart, 2004). The bottom-trapped bifurcation of
northward pathway over the Chukchi shelf is reproduced in even
the present medium-resolution version of the COCO model
(Fig. 4a and b). The eastern branch traces the isobaths in the Central Channel and Alaskan coastal region, but the western branch is
directed toward the Herald Canyon. A significant part of each passage eventually merges with the westward circulation along the
anti-cyclonic Beaufort Gyre.
The simulated current direction at the Barrow Canyon is compared with in situ mooring velocity data (Fig. 4c). The Barrow Canyon section defined here is composed of four model grid points
between Point Barrow (157°W, 71°N) and the western side of the
canyon (158°W, 71°N). The unit of current direction used here is
a true bearing (°T). 0 °T, 90 °T, 180 °T, and 270 °T correspond to
north, east, south, and west, respectively. The near-bottom current
direction of the UAF mooring data dominantly ranges between
30 °T and 90 °T. On the other hands, the modeled current velocity
averaged below 30-m depth of the Barrow Canyon section tends
to have the direction of 60–120 °T for the same period (2002 to
2004). Although this eastward bias might be attributed to the
smoothing of canyon shape, the presented model experiment
reproduces the dominance of northeastward flow through the
Barrow Canyon.
The volume transport across the Barrow Canyon section is calculated as by Watanabe (2011) (Fig. 4d). The Barrow Canyon section is composed of four model grid points between Point Barrow
(157°W, 71°N) and the western side of the canyon (158°W,
71°N). The 30-year mean transport of 0.17 ± 0.12 Sv is comparable
E. Watanabe / Ocean Modelling 71 (2013) 43–53
47
Fig. 4. Horizontal ocean velocity averaged in the top 30 m over the Chukchi Sea. The 30-year means in (a) March and (b) September are plotted. The unit vector of velocity is
10 cm s1. The white contours correspond to water depths between 30 and 60 m and have an interval of 3 m. (c) Total days in each category of current direction [°T] at the
Barrow Canyon calculated from (solid bar) the COCO model and (dashed bar) the UAF mooring. (d) Seasonal cycle of monthly mean volume transport across the Barrow
Canyon section [Sv]. See the solid lines in (a-b) and Section 3.2.1 for the exact location.
to the estimate of Woodgate et al. (2005b). The seasonal cycle of
monthly mean transport is then obtained by simply averaging
the yearly values from 1979 to 2008. During mid-winter and early
spring, major currents pass over the western side of the Chukchi
Sea. The volume transport across the Barrow Canyon has negative
values that correspond to the southwestward transport diverging
from the anti-cyclonic Beaufort Gyre. From late spring to early
summer, the principal pathway shifts toward the eastern Chukchi
Sea. The down-canyon transport reaches 0.74 Sv in June. The
strong dependence of the Pacific water pathway on surface wind
direction in the Chukchi Sea was shown by an idealized barotropic
ocean model (Winsor and Chapman, 2004) and by sensitivity
experiments in the eddy-resolving framework (Watanabe, 2011).
Similarly, the seasonal transitions of volume transport can be explained by shelf-wide wind fields, which may possibly be related
to basin-scale atmospheric regimes on monthly to decadal timescales (see Section 4). The predominant easterly wind forces the
bottom-trapped current to the western passage in the Chukchi
Sea during winter. The summertime weakening of easterly wind
or rather westerly wind allows the dynamical preference of eastern
route, where the SSH increases toward the Alaskan coast (Fig. 3).
3.2.2. Sea ice formation in the Alaskan coastal polynya
The coastal polynya of the northeastern Chukchi shelf from
Cape Lisburne to Point Barrow is a major source region for cold
lower-halocline water (Cavalieri and Martin, 1994; Dethleff,
2010). Dense water formation caused by sea ice production occurs
annually from late autumn to early spring (Tamura and Ohshima,
2011). It has been observed that the transformed Pacific-origin
water had an intraseasonally varying salinity that ranges between
32 and 33.5 psu (Itoh et al., 2012), and it sometimes reached a
hypersaline mode higher than 34 psu (Weingartner et al., 1998).
During spring, a portion of the shelf water transported northward
over the continental slope produces a bottom-trapped eastward
current along the northern edge of the Chukchi shelf (Pickart
et al., 2010). The offshoreward intrusion of cold dense water into
the halocline layer in the Canada Basin is accomplished by ventila-
tion along the isopycnal surface and the baroclinic eddies spawned
from the shelf-break jet (Woodgate et al., 2005a; Spall et al., 2008).
It is thought that a series of processes acting on the Pacific winter
water also affects the hydrographic structure of the southern Beaufort Sea (Ivanov and Watanabe, 2012).
It should be noted that the tranditional pan-Arctic Ocean models have reproduced the fundamental features of polynya-type ice
formation, although the lack of exact shape of wind-driven coastal
polynya is unavoidable. The COCO model also captures the emergence of an open water area enclosed by winter sea ice cover along
the northwestern coast of Alaska (Fig. 5a and b). The modeled winter sea ice production ranges within reasonable values as compared with satellite estimates, as already analyzed by Ivanov and
Watanabe (2012). For example, when the coastal shelf shallower
than 30-m depth between Cape Lisburne (166°W, 68°N) and Point
Barrow (157°W, 71°N) is defined as a target region, the seasonal total of sea ice production reaches 132 km3 from September 2000 to
April 2001 in the COCO model. This amount records the maximum
in the 30-year integration and is consistent with previous assessments in the Alaskan coastal polynya derived from satellite-based
algorithm tools (Martin et al., 2004; Tamura and Ohshima, 2011).
Long-term average sea ice freezing in the Alaskan coastal region
continues from October to April (Fig. 5c). Although the ice production peak of 20 km3 mon-1 in mid-winter is lower than that of Tamura and Ohshima (2011), this discrepancy may be caused by the
lack of ocean heat flux in their algorithm. The wind-driven up-canyon flow of warm Atlantic-origin water is a possible ocean heat
source south of the Barrow Canyon. The thermodynamic/dynamic
relationship of sea ice production was also addressed by Ivanov
and Watanabe (2012). Their analyses using the pan-Arctic COCO
model indicated that the mechanical divergence/convergence of
sea ice volume in the coastal region accounted for a significant part
of the interannual variations in thermal sea ice production during
the freezing period. Fig. 5a shows that low sea ice concentration is
accompanied by offshoreward sea ice drift. Correspondingly, the
area-mean sea ice concentration fluctuates around 0.95, even in
mid-winter (Fig. 5b).
48
E. Watanabe / Ocean Modelling 71 (2013) 43–53
Fig. 5. Sea ice properties averaged over the 30-year integration period. (a) Monthly sea ice formation in January [m]. Sea ice concentrations of 0.95, 0.96, 0.97, and 0.98 are
overlaid by black contours. Sea ice velocity fields are shown by vectors for every two model grids. The unit vector of velocity is 10 cm s1. (b and c) Seasonal variations in (b)
3
sea ice concentration and (c) net thermal sea ice production (growth minus melting) [km mon1 ] in the Alaskan coastal region, depicted by yellow dots in (a). Note that the
scale of vertical axis in (b) is enlarged between the values 0.8 and 1.
Fig. 6. Freshwater diffusive flux referenced to a salinity of 34.8 psu from the
Chukchi and Beaufort shelf to the Canada Basin [km3]. See Section 3.2.3 and Fig. 1
for the exact definition of shelf-basin boundary. The summertime total during July,
August, and September is plotted at each level.
3.2.3. Shelf-basin exchange in the southern Beaufort Sea
A number of previous studies have indicated that the Beaufort
shelf-break eddies played an important role in the transport of
warm and fresh shelf water into the upper halocline layer of the
Canada Basin (Pickart, 2004; Watanabe and Hasumi, 2009; Watanabe, 2011). Typically, a model grid spacing of a few kilometers is
necessary to represent realistic eddy features having a spatial scale
of O(10 km) in the polar region (Maslowski et al., 2008). Because
decades-long integration by an eddy-resolving basin-scale OGCM
requires considerably enormous task, even using forefront highperformance computer resources, the effects of subgrid-scale baroclinic eddies on tracer transport are sometimes parameterized by
the isopycnal layer thickness diffusion of GM in a coarser-resolution model than the internal Rossby radius of deformation. The
GM scheme accounts for the aspect of transfer from available potential energy to eddy kinetic energy accompanied by baroclinic
instability and has been widely adopted in global OGCMs (Hunke
et al., 2008). From sensitivity experiments using the pan-Arctic
COCO model, Watanabe and Hasumi (2008) reported that a GM diffusion coefficient of Oð103 m2 s1 Þ was effective for the shelf-to-basin transport of relatively fresh Pacific water across the Chukchi
and Beaufort shelf breaks with a reduction of simulated salinity
bias in the Arctic basin interior (e.g., a positive salinity drift from
the initial PHC value in the Canada Basin under climatological
atmospheric forcing).
Here, the lateral freshwater flux resulting from the isopycnal
and GM diffusion terms in the COCO model is defined as the freshwater diffusive flux. A reference salinity is set at 34.8 psu, as for the
FWC in Section 3.1.2. Fig. 6 shows the vertical profile of summertime flux across the 1000-m isobath between 140°W and 160°W
in the southern Beaufort Sea. The diffusive transport of shelf-origin
freshwater had a peak of 26 km3 in the subsurface layer from 50 to
100 m depth and gradually decreases with depth. Thus the GM
scheme indirectly represents the shelf-basin exchange induced
by baroclinic eddies with a vertical scale of 200–300 m (Watanabe
and Hasumi, 2009). Correspondingly, the shelf-water transport
would contribute to stratification in the halocline layer of the
southern Canada Basin, as is discussed in Section 4.1.
4. Halocline variability related to Beaufort shelf-break
processes
The effects of shelf-basin exchanges of Pacific summer and winter water masses on halocline structure in the southern Beaufort
Sea will now be explored.
4.1. Pacific summer water
The eastward current along the Beaufort shelf break is dominated by the Pacific-origin summer water with salinity assumed
to be lower than 32 psu following Itoh et al. (2012). The poleward
E. Watanabe / Ocean Modelling 71 (2013) 43–53
49
Fig. 7. Ocean volume transport through the Barrow Canyon section shown by line A in Fig. 1 and a solid line in Fig. 4 [Sv]. The seasonal averages during (a) April, May, and
June (AMJ) and (b) July, August, and September (JAS) in each year are categorized by two salinity ranges (solid) below 32 psu and (dashed) above 32 psu.
transport of Pacific water carrying heat, fresh water, and nutrients
is influential for hydrographic and biogeochemical structures in
the western Arctic Ocean (Weingartner et al., 2005; Mathis et al.,
2007). The volume transport through the Barrow Canyon discussed
in Section 3.2.1 is regarded as an index of current properties in the
Beaufort shelf-break region. Fig. 7 shows the interannual variation
in the seasonal averages of simulated total volume transport passing across the Barrow Canyon section. In the 2000s, the transport
averaged from July to September reaches a maximum of 1.01 Sv
in 2003 and a minimum of 0.48 Sv in 2007 (Fig. 7b).
To visualize the effects of shelf-water transport along the northern Alaskan coast on hydrographic structures in the southern
Beaufort Sea, the sea ice, ocean, and atmospheric fields in 2003
and 2007 are compared, along with findings by Watanabe
(2011). During the summer of 2003, low sea level pressure (SLP)
is located in the northwestern Canada Basin (Fig. 8a). The westerly
wind associated with the meridional SLP gradient drives an eastward sea ice drift and underlying ocean current along the northern
Alaskan coast. Sea ice cover over the shelf region disappears by the
end of July in the COCO model as detected by the AMSR-E Arctic
Sea-Ice Monitor, and thus surface wind directly promotes the eastward ocean current in August and September. The vertical salinity
profile along the 145°W section shows depressed halocline layers
near the shelf-break region (Fig. 9a). Two possible processes cause
the downward shift of isohaline layers. One is wind-driven Ekman
downwelling (i.e., the direct effect). In fact, net Ekman divergence
is negative in the southernmost Beaufort Sea (Fig. 8a). Ekman
downwelling above 20 m yr1 lasting throughout July, August,
and September is effective at pushing down nearshore water in
the entire column in 2003. The other process is the lateral advection of fresh shelf water (Fig. 7b). The Pacific summer water in
the Chukchi shelf is fresher than the offshore halocline water north
of Alaska. During this period, a vivid eastward current with a velocity exceeding 10 cm s1 appears over the shelf break (Fig. 9a). It
creates an isopycnal front along the shelf-basin boundary, as captured by the high-resolution mooring array of the Western Arctic
Shelf-Basin Interactions (SBI) project (Nikolopoulos et al., 2009).
The SBI observation suggested that the surface-intensified shelfbreak jet generated buoyant eddies via baroclinic instability during
late summer (Pickart, 2004). In the COCO model, the effect of subgrid-scale eddy behavior on tracer transport is parameterized by
Fig. 8. Net Ekman divergence in the basin area with water depth greater than
100 m [m yr1]. Red (blue) shade denotes Ekman divergence (convergence)
calculated from ocean surface stress in the COCO model. The yellow contours
show NCEP/NCAR sea level pressure [hPa]. The fields averaged over July, August,
and September of (a) 2003 and (b) 2007 are shown.
the isopycnal layer thickness diffusion of GM (Section 2). The GM
diffusion scheme approximately represents the equivalent process
even not all as described in Section 3.2.3. The consequent lateral
50
E. Watanabe / Ocean Modelling 71 (2013) 43–53
region (Fig. 9b). The 32 and 33 psu isohaline depths near the shelf
break are both shallower by approximately 30 m in 2007 than
2003. The disappearance of eastward current along the northern
Alaskan coast indicates that the Pacific summer water does not
reach east of Point Barrow and that lateral freshwater transport
is restricted. Actually, the northward freshwater diffusive flux is almost zero during the summer season of 2007 (Fig. 9b).
Westerly wind accompanied by low SLP in the Canada Basin
promotes the eastward current in the Beaufort shelf-break region
and the corresponding lateral freshwater transport during the
summers of 1983, 1986, 1988, 1992, and 1994 as well as 2003 in
this experiment (Fig. 7b). In these years, Ekman convergence outside the low SLP also contributes locally to halocline deepening
along the shelf break. On the other hand, an inactive or even westward current and Ekman upwelling in the southern basin occurs in
1990, 1995, 1997, 1999, 2005, and 2006 in addition to 2007. In
these years, the Beaufort High prevails even during summer, and
divergent surface winds emerge along its outer side. An exception
occurs in 1996, when Ekman convergence is dominant despite the
existence of westward current along the northern Alaskan coast.
This exception is due to the position of a high SLP center over
the Chukchi Sea shelf break during August and September. Thus
the shoaling/deepening of halocline depth induced by the eastward
current occurs preferably in phase with that due to surface Ekman
forcing along the Beaufort shelf break. Both processes depend on
the SLP pattern in the western Arctic.
Fig. 9. (Right) Vertical salinity profiles along the 145°W section [psu]. Eastward
ocean velocity is overlaid by black contours, and negative values indicate westward
current [cm s1]. The fields averaged over July, August, and September of (a) 2003
and (b) 2007 are shown. (Left) Same property as Fig. 6 but for (a) 2003 and (b) 2007.
freshwater transport results in the downward shift of offshore halocline layers and indirectly affects the hydrography of underlying
layers via Ekman downwelling flow (Fig. 9a).
During the summer of 2007, extremely high SLP covers the entire Canada Basin (Fig. 8b). The anti-cyclonic wind fields intensify
the westward ocean current along the Beaufort Gyre in the southern basin and blocks the eastward current originating from the
Chukchi shelf (Figs. 7b and 9b). This wind pattern induces Ekman
upwelling north of the shelf-basin boundary (Fig. 8b), as shown
by the winter climatology (Yang, 2006). The hydrographic field
along the 145°W section captures the bowl-shaped stratified structure along with the shaoling of halocline depth toward the coastal
4.2. Pacific winter water
The relationship between Pacific winter water and halocline
variability in the southern Beaufort Sea will now be addressed
from the viewpoint of wind patterns using the COCO model result.
In the winter mean atmospheric pattern, a dominant easterly wind
accompanied by prevailing high SLP in the Beaufort Sea drives
dense water formation due to thermal sea ice production in the
Alaskan coastal polynya region (Section 3.2.2). The transformed
shelf water, named the Pacific winter, water is then transported
to the Barrow Canyon. When the halocline layer associated with
the Pacific winter water is defined by a salinity range of 32 to
33.5 psu, as also used by Itoh et al. (2012), the layer thickness north
of the Barrow Canyon mouth (155°W, 72°N) shows decadal variation related to the anti-cyclonic/cyclonic patterns of basin-scale
wind fields (Fig. 10a). A decadal minimum of 50 m occurs in the
1990s. The 32 psu isohaline depth is mostly deeper than 50 m from
Fig. 10. Interannual variations in monthly mean (a) isohaline depth [m] and (b) net Ekman divergence [m mon1] north of the Barrow Canyon mouth. See Section 4.2 and
Fig. 1 for the exact location. The blue and red lines in (a) show 32 and 33.5 psu, respectively.
E. Watanabe / Ocean Modelling 71 (2013) 43–53
51
Fig. 11. Vertical profiles of (shade) potential vorticity [109 m1 s1], (black contours) potential temperature [°C], and (yellow contours) salinity [psu] across the Beaufort
shelf break in July of (a) 1995 and (b) 2001. See Section 4.2 and line B in Fig. 1 for the exact meridional location. The location north of the Barrow Canyon mouth referred to in
Fig. 10 is indicated by black arrows at the top of figures.
1991 to 1997. During this period, the westerly wind transports surface fresh shelf water with a salinity of less than 32 psu to the Barrow Canyon, and the halocline layers north of the canyon mouth
are consequently depressed (Figs. 7 and 10). The shoaling/deepening of 32 psu isohaline is significantly in phase with that of 33.5
psu isohaline. From 1998 to 2008, the 32 psu isohaline depth is
occasionally outcropped to the ocean surface as a result of winddriven upwelling (Fig. 10a and b). Ekman divergence north of the
Barrow Canyon mouth certainly exceeds 5 m mon1 in 1999,
2000, 2001, 2004, 2005, 2007, and 2008 of the last decade of model
integration. Thus it is expected that during this period the easterly
wind promotes the shelf-break upwelling in addition to winter sea
ice formation in the Alaskan coastal polynya.
The mooring array deployed along the 152°W section in the SBI
project captured the signal of cold dense Chukchi shelf water after
April (Spall et al., 2008). Early summer hydrographic properties
hence illuminate the consequence of winter-water intrusion.
Shelf-break profiles under the different wind patterns are compared by featuring the simulated fields in July 1995 and 2001
(Fig. 11). It has been reported that sea ice production in the Alaskan
coastal region was extremely small (large) in the 1994/1995 (2000/
2001) winter (Martin et al., 2004; Itoh et al., 2012). The decadal
experiment using the pan-Arctic COCO model also reproduced
these peaks (Ivanov and Watanabe, 2012). For comparison, the
horizontal axis is set to the line from Point Barrow (157°W,
71°N) to the Canada Basin interior (150°W, 75°N), as shown in
Fig. 1. Fig. 11 shows the contrasting features of halocline structure.
In 1995, strong stratification appears from the ocean surface to halocline depths, except in a thin cold layer at a depth of 60 m near
the Barrow Canyon mouth. On the other hand, in 2001 the intrusion of Pacific winter water characterized by a cold core and low
potential vorticity (PV) increases the halocline layer thickness with
the depression of 33.5 psu isohaline depth. Correspondingly, the
isohaline layer thickness between 32 and 33.5 psu reaches 100 m
(Fig. 10a). The low-PV signal of 1 109 m1 s1 with a water temperature minimum below 1 °C was frequently observed by recent
summer cruises along the 150 °W section (Itoh et al., 2012).
Because these westerly (easterly) wind periods correspond to
the cyclonic (anti-cyclonic) regime of basin-scale wind and ocean
circulation proposed by Proshutinsky et al. (2009), the linkages
among wind patterns, shelf-water transport, and halocline structure in the vicinity of Barrow Canyon mouth can be proposed as
follows. During the cyclonic regime, including the year 1995, both
inactive sea ice formation in the coastal polynya and local Ekman
convergence along the outer side of low SLP allow subsurface strat-
ification overlaid by shelf-origin fresher water in the downstream
region of Pacific winter water. In the anti-cyclonic regime, as in
2001, intense offshoreward wind enhances the sea ice production
and consequent down-canyon transport of transformed Pacific
winter water. The penetration of Pacific winter water with low
PV then prevents the wind-driven upwelling of underlying basin
water to the surface layer north of the Barrow Canyon mouth during the late spring and early summer seasons.
5. Summary and discussion
To address the mechanisms controlling halocline variability in
the western Arctic Ocean, the role of surface wind fields in halocline shoaling/deepening is assessed using the pan-Arctic Ocean
model. The model reproduces significant observed changes in summer sea ice extent over the entire Arctic Ocean and freshwater content in the Beaufort Gyre Region on interannual and decadal
timescales, which include remarkable sea ice decline and freshening in the 2000s. The model results support a strong coupling between the interannual variations in freshwater content and in
the shoaling/deepening of halocline layers in the Canada Basin.
Shelf-basin interaction has a significant influence on the stratified structure of the southern Beaufort Sea (Fig. 12). Eastward
shelf-water transport and local Ekman convergence north of the
Beaufort shelf break are commonly promoted by westerly wind
along the outer side of low SLP in the basin interior. The eastward
current provides the fresh Pacific summer water inflowing from
the Bering Strait for the shelf-break region. The lateral freshwater
transport enhances the subsurface stratification with the deepening of halocline layers in the southern Beaufort Sea. To the contrary, easterly wind accompanied by the prevailing Beaufort High
induces Ekman divergence and blocks the eastward current in
the shelf-break region. High-salinity water in the lower halocline
layer then upwells. Therefore, in the southern Beaufort Sea, the
shoaling/deepening processes of halocline layers induced by
shelf-water transport and directly due to surface Ekman forcing occur in the same direction, commonly depending on basin-scale
anti-cyclonic/cyclonic wind patterns.
The relationship between Pacific winter water and halocline
variability in the southern Beaufort Sea is also addressed. Halocline
thickness north of the Barrow Canyon mouth shows decadal variations in the present model. During the cyclonic period of basinscale wind fields, the inflow of surface fresh shelf water into the
upper portion of halocline layer induced by westerly wind plays
52
E. Watanabe / Ocean Modelling 71 (2013) 43–53
Fig. 12. Schematic image of linkages among basin-scale anti-cyclonic/cyclonic wind patterns, Ekman upwelling (U)/downwelling (D), and the shelf-to-basin transport of
Pacific summer and winter water masses.
a role in the depression of halocline layers. Although the easterly
wind over the Chukchi and Beaufort shelf breaks associated with
the anti-cyclonic wind pattern works for local upwelling, the
northward transport of cold, dense Pacific winter water through
the Barrow Canyon increases the halocline layer thickness north
of the canyon mouth. The down-canyon transport of Pacific winter
water depends in part on dense water formation resulting from sea
ice production in the Alaskan coastal polynya, which is activated
by the easterly wind. These results propose that the entering of Pacific winter water into the basin interior is capable of preventing
the Ekman upwelling of underlying saline basin water. The relative
importance of local upwelling and lateral advection of nutrientrich water for primary phytoplankton production would be the
point at issue in biological studies (Watanabe et al., 2012).
In the present experiment, the water properties of Pacific water
inflow at the Bering Strait (i.e., a lateral model boundary) are fixed
to idealized seasonal cycles, because an insufficient amount of
year-round data covering multi-decadal periods are available. An
increasing trend in heat inflow from the Bering Sea for the late
2000s was recently reported (Woodgate et al., 2010; Mizobata
et al., 2010). The interannual variability of the Bering Strait
throughflow might cause the synchronized fluctuation of shelfwater transport. In the pan-Arctic COCO model, the GM diffusion
scheme is adopted to indirectly reproduce the effects of subgridscale baroclinic eddies on tracer transport. Generally, it is difficult
to measure freshwater transfer by mesoscale eddies directly from
sparse available observations. Long-term data are necessary to
specify statistically significant eddy-induced tracer fluxes because
baroclinic instability is recognized to be a chaotic process.
Although the isopycnal layer thickness diffusivity used for climate
modeling has sometimes been inferred from simulation results of
eddy-resolving models (Rix and Willebrand, 1996; Nakamura and
Chao, 2000), the unrealistic values of diffusion coefficient were
occasionally diagnosed by these approaches. Therefore, although
the present work can be regarded as a practical step toward relating the wind-driven halocline variability with shelf-water transport, further accumulation of observational data and the
improvement of diagnosis method for lateral eddy flux would be
useful for the analyses of shelf-basin exchange and related hydrography in the Arctic Ocean modeling.
Acknowledgment
This research is funded by Grants-in-Aid for Scientific Research
(S) of Japan Society for the Promotion of Science, JFY2010–2014 No.
22221003, ’’Catastrophic reduction of sea-ice in the Arctic Ocean:
its impact on the marine ecosystems in the polar region’’. The
author appreciates Dr. Andrey Proshutinsky at Woods Hole Oceanographic Institution, Dr. Mike Steele at the University of Washington, and another members of the Arctic Ocean Model
Intercomparison Project (AOMIP) for providing fruitful opportunities in polar-region modeling studies. The courteous comments
and suggestions of anonymous reviewers markedly benefited the
presented product.
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