Beaufort shelf break eddies and shelf basin exchange of Pacific summer water in the western Arctic Ocean detected by satellite and modeling analyses

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 116,, doi:10.1029/2010jc006259, 2011 Beaufort shelf break eddies and shelf basin exchange of Pacific summer water in the western Arctic Ocean detected by satellite and modeling analyses Eiji Watanabe 1,2 Received 17 March 2010; revised 15 May 2011; accepted 1 June 2011; published 26 August 2011. [1] Mesoscale eddies and shelf basin exchange of Pacific summer water in the western Arctic Ocean are examined using satellite data sets and an eddy resolving coupled sea ice ocean model. Several surface eddy like features along the Beaufort shelf break are detected by the Moderate Resolution Imaging Spectroradiometer (MODIS) sea surface temperature and the Global Imager (GLI) radiance image. The QuikSCAT sea wind vectors indicate that summertime shelf wide wind is an important factor for interannual variations in the eddy properties. A realistic numerical experiment reveals that the origin of shelf break warm eddies and timing of their generation can be classified into three types. Type I eddies are frequently produced in the vicinity of the Barrow Canyon throughout the summer season by a combination of outflow of Pacific summer water with a low potential vorticity from the Barrow Canyon and lateral velocity shear of the canyon jet. Surface wind variability modulates the timing of generation. Type II and Type III eddies originate from an eastward current along the Beaufort shelf break, although origin depth and background hydrographic structure differ between them. Type II eddies are spawned from the shelf slope at middepth during early summer. Type III eddies are generated from the surface isopycnal front over the shelf break during late summer. The shift from Type II eddies to Type III eddies is caused by surface buoyancy input in the upstream region. A sensitivity experiment using different atmospheric forcing data suggest that the mechanism controlling Pacific water transport from the Chukchi shelf to the Canada Basin differs depending on surface wind fields in the shelf region. The mesoscale eddies primarily induce the shelf to basin transport under weak or westerly wind conditions during summer, while wind driven Ekman transport is a major driver in the easterly wind regime. Year 2003 (2007) corresponds to the former (latter) case. Citation: Watanabe, E. (2011), Beaufort shelf break eddies and shelf basin exchange of Pacific summer water in the western Arctic Ocean detected by satellite and modeling analyses, J. Geophys. Res., 116,, doi:10.1029/2010jc006259. 1. Introduction [2] The Pacific water that passes through the Bering Strait is a predominant source of heat, fresh water, and nutrients in the Arctic Ocean. Its transport has a potential to affect both the sea ice variation and marine ecosystem. The Bering Strait throughflow is primarily driven by a meridional gradient of sea surface height between the Bering and Chukchi seas, although the volume transport seasonally and interannually varies according to changes in the wind stress field [Aagaard et al., 2006]. A significant part of the Pacific water travels northward along several major features of bottom topography over the Chukchi shelf and is then transported into the Canada Basin [Weingartner et al., 2005]. Shimada et al. [2006] 1 International Arctic Research Center, University of Alaska Fairbanks, Fairbanks, Alaska, USA. 2 Now at Japan Agency for Marine Earth Science and Technology, Yokohama, Japan. Copyright 2011 by the American Geophysical Union. 0148 0227/11/2010JC006259 propose that ocean heat transport from the Alaskan northern coast toward the Northwind Ridge has been promoted by the intensified westward ocean surface stress under fragile sea ice for several recent years. This scenario implies the existence of a positive feedback mechanism between sea ice retreat and ocean heat transport, which reduces net sea ice growth. The nutrient rich Pacific water undergoes physical and biogeochemical modifications through densification due to winter surface cooling and sea ice brine rejection, seasonal primary production, and interactions with organic rich sediments in the Chukchi Sea. The buoyant Pacific summer water flows along the Alaskan northwestern coast as a surface intensified current from late summer to early autumn. The winterdeformed Pacific water flows eastward as a bottom intensified shelf edge current along the Chukchi and Beaufort shelf breaks from late winter to early summer. These currents are intrinsically constrained along the isobaths by geostrophic dynamics. Possible mechanisms controlling cross shelf transport of the shelf water are suggested to be frontal instabilities, wind driven Ekman transport, and dense water 1of16

plumes [Pickart, 2004; Spall, 2007; Chao and Shaw, 2003]. [3] Many studies have focused on the contribution of mesoscale eddy activities over the Chukchi and Beaufort shelf breaks to the transport of Pacific water into the basin interior. The surface and halocline layers above 300 m depth in the Canada Basin are full of a number of anticyclonic eddies [Manley and Hunkins, 1985]. The surface intensified warm core eddies observed in the vicinity of the Barrow Canyon have a signal of the Alaskan Coastal Current. Hydrographic composites over the Beaufort shelf break in each season constructed by Pickart [2004] imply that the surface intensified jet forms silicate rich warm core eddies detected in the offshore side of the shelf break. Pickart [2004] suggests that these eddies are produced by mixed baroclinic and barotropic instability from late summer to early autumn. The detailed physical and chemical properties of a cold core eddy observed on the Chukchi Sea continental slope are surveyed by Mathis et al. [2007] and Kadko et al. [2008]. Both studies mention that the bottom intensified shelf edge current forms the cold eddies, which play a significant role in the transport of carbon, oxygen, and nutrients associated with the Pacific winter water into the upper halocline of the Canada Basin. Spall et al. [2008] combine high resolution mooring data and an idealized numerical experiment in the southern Beaufort Sea. Their analysis of energetics points out that anticyclonic cold core eddies are generated by baroclinic instability of the shelf break jet. Thus the mesoscale warm and cold eddies and their roles in the transport of shelf water into the basin interior are occasionally reported. However, the exact place and period of the generation of Beaufort shelf break eddies and eddy induced transport are still uncertain. [4] In this research, the process of Pacific water transport from the Chukchi shelf to the Canada Basin is investigated using an eddy resolving coupled sea ice ocean model, whose horizontal resolution is about 2.5 km, and realistic experimental design. The author s previous study indicates that the Pacific water transport across the Beaufort shelf break differs depending on summertime sea ice extent [Watanabe and Hasumi, 2009]. When the sea ice margin retreats toward the basin interior in late summer, a narrow jet through the Barrow Canyon (hereafter referred to as the Barrow Canyon jet) is intensified, and correspondingly enhanced eddy activities promote the shelf to basin transport. When sea ice remains in the shelf throughout the year, the sea ice ocean stress plays a great role in braking the jet and in consequent restraint of the eddy activities. The dependence of the jet strength on surface buoyancy flux, such as absorption of shortwave radiation and sea ice melting, is negligible compared with that on surface momentum flux. Watanabe and Hasumi [2009] propose that the shelf to basin transports of heat and fresh water associated with the Pacific summer water are promoted following the northward shift of summer sea ice margin for several recent years. [5] The present study attempts to clarify (1) mechanisms controlling the origin of Beaufort shelf break eddies and timing of their generation and (2) the relationship of Pacific water transport with atmospheric conditions in the western Arctic Ocean. Satellite data and numerical simulation results are analyzed, and the interannual variability is then discussed. This paper is organized as follows: The method of satellite data and modeling analyses is described in section 2. Eddyrelated features and surface wind properties detected by satellite data sets are presented in section 3. The origin of Beaufort shelf break warm eddies and timing of their generation are examined using a simulation result under realistic atmospheric forcing in section 4. The relationship of the Barrow Canyon jet with surface wind variability and the interannual variability of shelf to basin Pacific water transport are discussed in section 5. The findings obtained in the present work are summarized in section 6. 2. Method 2.1. Satellite Data Analysis [6] Spatial and temporal features of warm water spread in the western Arctic Ocean and the Beaufort shelf break eddies are investigated using following satellite data sets: the Advanced Microwave Scanning Radiometer for Earth Observing System (AMSR E), the Moderate Resolution Imaging Spectroradiometer (MODIS), the Global Imager (GLI), and the QuikSCAT Scatterometer from 2003 to 2008. [7] Sea ice concentration (SIC) is obtained from the AMSR E data set. The AMSR E is an eight channel dual polarization passive microwave radiometer. The daily SIC data from June 2002 to present are available at the International Arctic Research Center of the University of Alaska Fairbanks (IARC) Japan Aerospace Exploration Agency (JAXA) Information System (IJIS) data archive. To overlay the SIC on the MODIS sea surface temperature (SST) images, the original gridded data in polarstereo projection are interpolated to a geographical grid of 0.125 in zonal direction and 0.04 in meridional direction. The grid interval of constructed data is from 5.9 km at latitude 65 N to 2.4 km at 80 N. [8] The MODIS is a visible and infrared sensor covering 36 spectral bands. In the present study, the 8 day composites of level 3 binned 11 mm brightness temperature (BT11) are checked to detect summertime warm water spread in the southern Beaufort Sea. It is assumed that the BT11 precisely represents SST. The scenes with a spatial resolution of 4 km in the entire Chukchi Sea and the southern Beaufort Sea from June 2003 to October 2008 are obtained from the National Aeronautics and Space Administration (NASA) ocean color Web site (http://oceancolor.gsfc.nasa.gov) and are digitally processed using the NASA s Sea Viewing Wide Field of View Sensor (SeaWiFS) Data Analysis System (SeaDAS) software. The BT11 images are interpolated to the same geographical grid as the AMSR E SIC. [9] To address the detailed structure of mesoscale eddies, the MODIS level 2 Local Area Coverage (LAC) of the BT11 in the southern Beaufort Sea is used. The level 2 product has a spatial resolution of 1 km. The snapshot scene presented in this paper was obtained in September 2003, when distinct eddy like features were visible in the MODIS level 3 BT11 (see section 3.1). The data process method is the same as the level 3 product. In addition, sea ice floes likely driven by oceanic mesoscale eddies are focused on using a highresolution optical radiance image of the GLI. The GLI has 23 bands in the visible and near infrared region (VNIR), 6 bands in the short wave length infrared region (SWIR), and 7 bands in the middle and thermal infrared region. Both the VNIR and SWIR bands have a spatial resolution of 250 m at the nadir. The GLI level 1B images from April to October 2003 are obtained from the IJIS archive. The RGB browse is 2of16

Figure 1. Model bathymetry (m). The locations shown include Chukchi Sea (C.S.), Herald Canyon (H.C.), Chukchi Plateau (C.P.), Northwind Ridge (N.R.), and Mackenzie River mouth (M.R.). constructed from 825, 545, and 460 nm visible bands to visualize sea ice floes. The BT11 in thermal infrared bands is used to display the SST pattern. [10] The impact of atmospheric circulation fields on the shelf wide ocean circulation is also a target in this paper. The QuikSCAT sea wind vectors are acquired from the Physical Oceanography Distributed Active Archive Center at NASA s Jet Propulsion Laboratory (http://podaac.jpl.nasa.gov/). The daily level 3 gridded data set over the entire globe from 2003 to 2008 is processed using the Hierarchical Data Format (HDF) library version 4.2. The spatial resolution is 0.25 0.25, which is equivalent to 10 km in zonal direction and 28 km in meridional direction at 70 N. All the ascending daily data with a rain flag of 0 in 160 W 180 W and 66 N 72 N are used. The target region contains most of the Chukchi Sea. The undefined grid points due to sea ice cover are excluded from the analysis. 2.2. Modeling Analysis [11] The model and experimental design are basically the same as Watanabe and Hasumi [2009] except for model domain and atmospheric forcing data including river runoff. The detail is presented below. The coupled sea ice ocean model for the numerical experiments is Center for Climate System Research Ocean Component Model (COCO) version 3.4 developed at the University of Tokyo [Hasumi, 2006]. The sea ice part 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 rheology of Hunke and Dukowicz [1997]. The ocean part is a free surface ocean general circulation model formulated on the spherical coordinate system. The model incorporates 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 advection, respectively. The turbulence closure scheme of Noh and Kim [1999] is applied for calculating vertical viscosity and diffusion coefficients. The background vertical diffusion coefficient varies with depth from 0.1 10 4 m 2 s 1 at the top level to 3.0 10 4 m 2 s 1 at the bottom level. To represent mesoscale eddy activities well, the horizontal eddy viscosity is parameterized by the Smagorinsky type biharmonic formulation whose constant coefficient is set to 3 [Griffies and Hallberg, 2000], and an enstrophy preservation scheme of Ishizaki and Motoi [1999] is adopted. The background horizontal viscosity and diffusion coefficients are 5.0 10 2 m 2 s 1 and 1.0 m 2 s 1, respectively. [12] The model domain contains the entire Chukchi Sea and the southern area of the Canada Basin (Figure 1). The bathymetry is constructed from the International Bathymetric Chart of the Arctic Ocean [Jakobsson et al., 2000]. The model s spherical coordinate system is rotated so that the singular points are on the equator. The horizontal resolution is about 2.5 km, and there are 25 vertical levels. The vertical grid width varies from 2 m at the top level to 1000 m at the bottom level (Table 1). [13] The atmospheric forcing components are obtained from the National Centers for Environmental Prediction/ National Center for Atmospheric Research (NCEP/NCAR) reanalysis data set [Kalnay et al., 1996]. Two experiments are performed in the present work. The reanalysis daily data from March 2003 (2007) to February 2004 (2008) is given in the 2003 (2007) case to investigate the ocean response under different wind regimes. The satellite analyses show different shelf wide wind field and SST distribution including eddylike features in these years. The wind stress is calculated from the sea level pressure (SLP) following a formulation adopted in the Arctic Ocean Model Intercomparison Project (AOMIP) (http://www.whoi.edu/page.do?pid=30565). The river water discharge is prescribed as surface freshwater flux at the Mackenzie River mouth (Figure 1). The monthly gauged and ungauged volume flux based on the Global River Data Center is listed on the AOMIP Web site (http://www.whoi.edu/page. do?pid=30587). In the marginal region of the model domain except for the Bering Strait, a sponge boundary condition is Table 1. Depth Levels of the Model Level Depth (m) 1 2 2 5 3 10 4 15 5 20 6 30 7 40 8 50 9 60 10 80 11 100 12 125 13 150 14 175 15 200 16 250 17 300 18 350 19 400 20 500 21 700 22 1000 23 2000 24 3000 25 4000 3of16

Figure 2. (a) MODIS 8 day mean level 3 sea surface temperature on 6 13 September 2003 ( C). Scene ID is A20032492003256. AMSR E sea ice area is overlaid by sky blue shade. (b) MODIS level 2 sea surface temperature on 9 September 2003 ( C). Scene ID is A2003252130000. The region corresponds to a yellow trapezoid in Figure 2a. applied for the ocean part, and an open boundary condition is applied for the sea ice part; no ocean current normal to the lateral boundary is applied, and the normal gradient of sea ice velocity is set to zero. The horizontal viscosity coefficient is enlarged, and the temperature and salinity at all depths are restored to the monthly mean of the Polar Science Center Hydrographic Climatology (PHC) data [Steele et al., 2001] in the sponge region. According to these procedures, artificial ocean reflective waves, sea ice ridge, and polynya owing to the model boundary are depressed. [14] The Pacific water inflow with seasonal cycle is provided at the Bering Strait based on the hydrographic observation of Woodgate et al. [2005a]. Specific values are prescribed to northward velocity and salinity at the strait so that the annual mean inflow and salinity of the Pacific water are 0.8 Sv (1 Sv 10 6 m 3 s 1 ) and 31 psu, respectively. The inflow reaches a maximum in June and a minimum in December, and its seasonal amplitude is 0.4 Sv. Salinity reaches a maximum in March and a minimum in September, and its seasonal amplitude is 1 psu. Temperature is kept at the freezing point from January to June, and it reaches a maximum of 5 C in September. The zonal gradient of temperature and salinity is taken into account so that warmer and fresher water passes along the Alaskan side. The total freshwater transport through the Bering Strait referenced to a salinity of 34.8 psu is 2755 km 3 yr 1, which is almost equal to the estimation of Woodgate and Aagaard [2005]. [15] To visualize the Pacific water pathway, a virtual passive tracer is provided at the Bering Strait so that the Pacific water concentration is kept at 100% at the strait throughout the integration period in each experiment. Advection and diffusion of the tracer are calculated by the same formulation as ocean temperature and salinity. As shown in section 4.3, the Pacific water transport across the Chukchi and Beaufort shelf breaks reaches a maximum from late summer to early autumn and becomes a minimum in midwinter. Hence the model is integrated for 1 year from March to the next February in all the experiments. March and February are regarded as the beginning and ending of the annual transport, respectively. The model is initiated from temperature and salinity fields in the PHC March data and from no sea ice and ocean circulation. The initial distribution of sea ice thickness is constructed by multiplying an arbitrary factor of 1 m to the annual mean sea ice concentration over the period from 1972 to 2004, which is obtained from the National Ice Center Arctic sea ice chart [National Ice Center, 2006]. All the experiments have no spin up stage to minimize the disturbance of hydrographic structure in the Canada Basin due to the solid lateral boundary of model domain. 3. Satellite Data Analysis 3.1. Beaufort Shelf Break Eddies [16] The 8 day mean MODIS SST and AMSR E sea ice area from 6 to 13 September 2003, are displayed in Figure 2a. The sea ice margin retreats offshore far from the Chukchi and Beaufort shelf region from late summer to early autumn. The warm water area where the SST is above 0 C spreads over the entire Chukchi Sea and north of the Alaskan northern coast. SST in the sea ice marginal region is kept almost at the freezing point. There are a few eddy like warm core patches cores along the Beaufort shelf break, which likely contribute to shelf basin heat exchange. The weekly mean field may underestimate the spatial scale of the warm core patches because water mass could move on a daily timescale. The MODIS level 2 snapshot SST image reveals the detailed warm core structure (Figure 2b). The field shows that these eddy like features characterized by a 0 C isothermal line have an anticyclonic structure, and their diameter is about 70 km. The horizontal scale is significantly larger than an internal Rossby radius of deformation in the polar region. It should be noted that the eddy properties derived from satellite images represent only the surface features, so the spatial scale of each eddy could be underestimated. [17] In the area north of the Barrow Canyon, which corresponds to an upstream region of summertime eastward current along the Beaufort shelf break, an eddy pair like feature is detected by the high resolution GLI measurement. The snapshot GLI radiance image obtained on 8 July 2003, captures a few curved sea ice edges and oval shaped fragile ice pieces. This scene implies that the Alaskan Coastal Current pushes sea ice floes northeastward, and an eddy pair like pattern appears north of Point Barrow (Figure 3). The GLI SST displays that the pair is composed of a warm part on 4of16

than mesoscale eddies are assumed to account for a large part of the shelf to basin transport in 2007. Figure 3. GLI level 1B radiance on 8 July 2003. RGB image is constructed using 865, 545, and 460 nm channels. GLI sea surface temperature is overlaid in open water areas ( C). Scene ID is A2GL20307085609OD1. the inshore side and a cold part on the offshore side. This structure indicates that the instability of the Barrow Canyon jet is a formation mechanism of the Beaufort shelf break eddies. The MODIS and GLI images propose that a part of mesoscale eddies detected in the southern Beaufort Sea might originate in the vicinity of the Barrow Canyon mouth. [18] The eddy like warm water features north of Point Barrow are captured by the MODIS SST in summer 2005 and 2007 (not shown). In 2008, a multiyear ice ridge is hooked by the sea coast near Point Barrow (K. Shimada, personal communication, 2009) and seems to prevent eddy generation. It is difficult to track eddy like signals because of many missing values due to cloud cover in 2004 and 2006. When the field of view is extended to the shelf wide region, the isothermal line is found to be parallel to the Chukchi and Beaufort shelf breaks during summer in most years. However, the warm water area widely spreads northwestward over the Chukchi Plateau and the isothermal line cuts across the shelf basin boundary in September 2007. Some factors other 3.2. Surface Wind Condition [19] It is expected that the strength of the Barrow Canyon jet depends on local and shelf wide wind fields. The response of the Pacific water circulation to wind stress variation over the Chukchi shelf is assessed by Winsor and Chapman [2004]. Their barotropic model demonstrates that strong northeasterly to easterly winds are able to completely reverse the eastward currents along the Alaskan coast and the Beaufort shelf break and to promote a northward flow of Pacific water across the shelf edge of the Chukchi Sea. Although the barotropic model cannot explicitly reproduce the shelf basin exchange, it is suggested that such a strong wind can cause the transport of surface shelf water toward the Canada Basin interior changing the eddying regime. To assess the potential role of shelf wide wind fields in the Pacific water transport, summertime SLP derived from the NCEP/NCAR reanalysis data set is compared between 2003 and 2007 (Figures 4a 4c). In the climatological field, which corresponds to a long term mean from 1948 to 2008, the Beaufort High covers a broad area of the Canada Basin (Figure 4a). This SLP pattern induces moderate easterly wind in the Chukchi Sea. The Beaufort High records up to 1020 hpa in 2007, while the SLP gradient from the Chukchi Sea to the Canada Basin is the opposite in 2003 (Figures 4b and 4c). The geostrophic wind associated with the SLP fields flows in an eastward (westward) direction in 2003 (2007) over the shelf region. The daily QuikSCAT sea wind vectors in the entire Chukchi Sea is analyzed to visualize the shelf wide wind properties from June to September 2003 and 2007, respectively. The characteristics of wind direction and speed are represented by the probability density function (PDF) in the combined spatial and temporal dimension (Figure 5). Total size of the PDF is equal to the sum of numbers of valid grid points in a target region (see section 2.1) and days in each month. The PDF in each range of wind direction is further categorized by the range of wind speed. The PDF histogram shows that northwesterly surface wind is frequently intensified in July and September 2003. The maximum of daily mean wind velocity exceeds 10 m s 1. In contrast, easterly wind is remarkably dominant throughout the summer, especially in September Figure 4. NCEP/NCAR sea level pressure averaged from June to August(a) 1948 2008 mean, (b) 2003, and (c) 2007 (hpa). 5of16

Figure 5. Surface wind properties in the Chukchi Sea detected by QuikSCAT level 3 daily data set from June to September 2003 and 2007. Target region is 160 W 180 W and 66 N 72 N. Horizontal axis is wind direction ( ). Northerly wind direction is set to 0. Positive (negative) values correspond to easterly (westerly) wind. Vertical axis denotes the probability density function (PDF) of wind direction in the spatial and temporal total (%). Blue, pink, and red bars show the PDF of wind speed ranging 0 5, 5 10, 10+ m s 1, respectively. 2007. The easterly wind associated with high SLP in the central Beaufort Sea has a potential to directly promote a northward ocean current across the shelf edge. [20] Based on the analysis so far, one scenario in the shelfto basin transport can be proposed. When westerly wind becomes intensified in the Chukchi Sea during summer, the Alaskan Coastal Current and consequent generation of mesoscale warm eddies might be enhanced. Under easterly wind conditions, the wind driven Ekman transport may become dominant for the Pacific water pathway from the Chukchi Sea to the Canada Basin interior. 4. Modeling Analysis 4.1. Ocean Hydrographic and Circulation Fields 4.1.1. Overall Pattern [21] The simulation result in the 2003 case is compared with observational data. The sea ice margin is located far from the Alaskan northern coast in September (Figure 6a). It is suggested that the summertime sea ice cover is one of the crucial indices for fluctuation of the Barrow Canyon jet [Watanabe and Hasumi, 2009]. In the 2003 case, the ice free period, when the SIC is below 0.2 in the Barrow Canyon, is from 29 June to 25 October. The total number of open water days is 119. The daily AMSR E SIC shows the annual open water days vary between 97 and 151 from 2003 to 2008. Even when the SIC criterion for definition of the ice free period is changed to 0.1, the tendency is essentially similar. The relationship between sea ice condition and the Barrow Canyon jet is presented in section 4.1.3. [22] The SST field in the 2003 case shows that the warm water area spreads over the entire Chukchi Sea and north of the Alaskan northern coast in September (Figure 6a). There are several warm core eddies along the Beaufort shelf break. Their horizontal scale is up to 80 km. The spatial SST pattern is close to that detected by the MODIS satellite in 2003, as shown in section 3.1. The sea surface height (SSH) maximum characterized by a low salinity pool is located east of the Northwind Ridge (Figure 6b). The minimum sea surface salinity of 29.8 psu is located slightly east compared with the PHC September data. Correspondingly, the anticyclonic Beaufort Gyre covers a major area of the Canada Basin. The Figure 6. (a) Sea surface temperature ( ) and sea ice area (sky blue shading, dimensionless) and (b) sea surface height (shading, cm) and sea surface salinity (black contours, psu) on 22 September in the 2003 case. White contours in Figure 6a show bottom topography (m). 6of16

Figure 7. (a) Barotropic ocean current in the Chukchi Sea averaged in JJA (m2 s 1). (b) Seasonal cycle of mass transport across Barrow Canyon (A B) and Chukchi eastern (C D) and western (D E) sections shown as yellow lines in Figure 7a (Sv). The 10 day running mean is applied to each transport. (c) Vertical profiles of horizontal velocity (shading, cm s 1) and potential temperature (contours, C) averaged in August. All (Figures 7a 7c) are those in the 2003 case. gyre velocity at 30 m depth is about 2 5 cm s 1. Since the baroclinic Rossby waves hardly propagate in the polar region due to small b, the barotropic response is important in the present experiment. Another SSH maximum artificially arises along the northern boundary of the model domain. The high SSH may partly affect the Beaufort Gyre, but it has a minor impact on the Chukchi Sea circulation. The elevated SSH in the Chukchi Sea eastern shelf reflects the buoyant Alaskan Coastal Current. The detail of circulation in the shelf is mentioned in subsequent sections. 4.1.2. Chukchi Sea [23] The northward Bering Strait throughflow is diverted into three branches following major features of bottom topography over the broad, shallow Chukchi Sea (Figure 7a). One branch travels northwest toward the Herald Canyon. The other two branches trace northeast through the Central Channel and along the Alaskan coastline, respectively. This overall circulation pattern in the Chukchi Sea is consistent with previous observations [Weingartner et al., 2005; Woodgate et al., 2005b] and regional model experiments [Winsor and Chapman, 2004; Spall, 2007]. In addition, there exist a lot of small scale pathways controlled by complex topography and surface wind variations. The western passage in the Chukchi Sea splits into two branches near the head of the Herald Canyon. The eastern branch is trapped along the northern flank of the Herald Shoal and the corresponding eastward current eventually joins the Alaskan Coastal Current south of the Hanna Shoal. This topographic current is continuously maintained, as Spall [2007] indicates the existence of zonal flows. The western branch is directed downward along the Herald Canyon. A significant part of the western component occasionally flows eastward over the shelf and merges with the eastern branch. It is assumed that this stream is dominated by surface wind variations. The remaining component reaches the Herald Canyon mouth and then turns eastward along the northern edge of the Chukchi Sea shelf. The Central Channel throughflow also has bifurcation toward the Alaskan coast. Part of the flow is incorporated into the easternmost branch and creates a strong narrow jet in the Barrow Canyon during late summer. The other part, arriving at the Chukchi Sea shelf break, flows westward along the Beaufort Gyre. [24] Figure 7b shows the seasonal cycle of the simulated volume transport across the Barrow Canyon and the Chukchi eastern and western sections. A 10 day running mean is applied to the transport to remove conspicuous daily variability. The Barrow Canyon section is defined by a line between a location northwest of the canyon (156.6 W, 71.9 N) and Point Barrow (156.3 W, 71.3 N). The Chukchi eastern section corresponds to a path between Cape Lisburne (166.2 W, 69.2 N) and the Herald Shoal (171.7 W, 70.6 N). The Chukchi western section corresponds to a channel between the Herald Shoal and Wrangel Island (179.2 W, 71.4 N). In March, the northwestward transport across the Chukchi western section is dominant ( 0.5 Sv) compared with the eastern passage between Cape Lisburne and the Herald Shoal ( 0.2 Sv). The flow through the Barrow Canyon is still weak ( 0.2 Sv). From April to July, the major currents 7 of 16

Figure 8. (a) Time series of strength of Barrow Canyon jet (m s 1 ) in the 2003 and 2007 cases. (b) Wind stress anomalies regressed to jet strength (vector) and correlation coefficients between eastward wind stress and jet strength (orange contour, see section 5.1 for details). The unit vector of wind stress is 0.02 Pa. (c) Time series of eastward wind stress at location shown by a red cross in Figure 8a where maximum correlation appears (Pa). Both Figures 8b and 8c are those in the 2003 case. shift eastward and the transport across the Chukchi eastern section becomes exceeding the western one. The eastern transport gradually increases until July and decreases afterward. The Barrow Canyon throughflow reaches a maximum ( 1.5 Sv) at the end of July. The seasonal cycle of these eastern transports primarily reflects the prescribed Bering Strait throughflow (see section 2.2). After November, all the components have remarkable daily weekly fluctuations due to surface wind variability, while most of the Pacific water passes through the Chukchi western section. In winter, the Barrow Canyon throughflow frequently shows negative values, which indicate an up canyon flow from the Canada Basin onto the Chukchi shelf. The annual mean transport through the Herald Canyon, which is close to the transport across the Chukchi western section, is within the previous estimates of Spall [2007] (0.4 Sv), Weingartner et al. [2005] (0.28 Sv), and Woodgate et al. [2005b] (0.28 Sv). The seasonal amplitude of the volume transport across the Chukchi western section is relatively small, possibly explained by the atmospheric circulation field. In spring, strong easterly wind promotes the flow to the western route in the Chukchi Sea, although the Bering Strait inflow is small. In summer, when the Bering Strait inflow increases, weakening of easterly wind reduces the ratio of the flow toward the Herald Canyon compared with the eastern transport. The large amplitude of fluctuation during the winter season is also caused by surface wind variability. 4.1.3. Barrow Canyon [25] The hydrographic structure in the Barrow Canyon is similar to the previous experiments [Watanabe and Hasumi, 2009] and in situ observation [Pickart et al., 2005]. The model result broadly captures a steep isopycnal front between the warm Pacific summer water in the upper layer and the underlying cold Chukchi shelf water. The Barrow Canyon jet has daily variability with a maximum velocity of 2.1 m s 1 during summer (Figure 8a). The ice free period mentioned in section 4.1.1 indicates that an abrupt increase of the jet strength at the end of June is induced by sea ice retreat from the canyon surface, as with Watanabe and Hasumi [2009]. The monthly mean temperature in the Barrow Canyon section increases from 1.7 C in May to 2.3 C in September. The salinity decreases from 31.9 psu in June to 31.5 psu in September. Hence the Barrow Canyon throughflow obtains buoyancy during this period. In winter, the Atlantic origin water whose potential density is more than 27s occasionally intrudes from the Canada Basin into a deep layer in the canyon accompanied with the upstream flow mentioned in section 4.1.2. This type of Atlantic water intrusion was previously detected by hydrographic observations of Münchow and Carmack [1997] and Okkonen et al. [2009]. 4.1.4. Herald Canyon [26] Pickart et al. [2010] investigate flow of the Pacific summer and winter water masses through the Herald Canyon using in situ data along four cross canyon transects in summer 2004. The survey detected hydraulic activities including an eastward switch of the winter water path, a corresponding sudden increase in layer thickness, and production of mixed summer and winter water masses. The COCO model reproduces major features of their findings. For example, at the head of the Herald Canyon, lateral shear of barotropic currents composed of a poleward summer water jet over the eastern flank of the canyon and a dense winter water flow on the western side of the canyon arises below seasonal pycnocline at 10 20 m depth (Figure 7c). However, the Herald Canyon jet has a velocity between 10 and 30 cm s 1 in the 2003 case, which is clearly smaller than the geostrophic velocity estimated by Pickart et al. [2010]. The model demonstrates that the flow on the eastern side of the canyon is 8of16

Figure 9. Daily ocean surface velocity (vector) and relative vorticity (shading, s 1) for the (a) 21 July, (b) 6 August, (c) 20 September, and (d) 7 October in the 2003 case. The unit vector is 50 cm s 1. diverted into the central Chukchi Sea along the northwestern flank of the Herald Shoal. The swift jet on the eastern side of the canyon, advecting the summer water in the southern part of the canyon, has been considerably cooled before passing through the northern part. The Long Strait is a potential gate for water exchange between the East Siberian Sea and the Chukchi Sea. The flow through the Long Strait seems to have a minor contribution to the Herald Canyon throughflow, especially in spring and summer, since the volume transport through the Long Strait is estimated to be negligible ( 0.01 Sv) [Woodgate et al., 2005b]. 4.2. Beaufort Shelf Break Eddies 4.2.1. Development Process [27] The simulated daily ocean velocity fields in the surface layer capture a lot of anticyclonic eddies whose diameter is approximately 70 km along the Beaufort shelf break (Figures 9a 9d). This horizontal scale is consistent with the warm core eddies detected by the MODIS SST (Figure 2). North of the Barrow Canyon, a pair of anticyclonic and cyclonic eddies is also produced. The strength of each eddy can be represented by relative vorticity, which reaches 1.2 f. f is the Coriolis coefficient at the latitude where the eddies stay ( 1.4 10 4 s 1). The maximum rotational velocity is about 0.9 m s 1. [28] To investigate the origin of the shelf break eddies and timing of their generation, the development process of the eddies is tracked. 4.2.1.1. Stage I (21 July): Generation [29] An energetic eddy pair composed of a warm core anticyclone and a cold core cyclone (ED1) is spawned from the tip of the intensified Barrow Canyon jet (Figure 9a). The GLI image catches such an eddy pair, as referred in section 3.1. The maximum jet velocity exceeds 1 m s 1 (Figure 8a). A small warm core anticyclone (ED2) is located east of the eddy pair. 4.2.1.2. Stage II (6 August): Development [30] ED1 and ED2 migrate eastward, developing their kinetic energy and spatial scale up to 60 km (Figure 9b). The cyclonic component of ED1 significantly shrinks. The anticyclonic component of ED1 has a maximum vorticity of 0.7f and rotational velocity reaching 0.8 m s 1. The eastward shelf break current becomes highly intensified with meandering during this period. The additional anticyclone (ED3) appears along the meander. The distance between their generation places is about 120 km. The isobaths do not curve so sharply at the location of these eddies. 4.2.1.3. Stage III (20 September): Maximum Vorticity and New Generation [31] A lot of anticyclones including newly formed eddies (ED4 9) are clustered offshore of the shelf break (Figure 9c). The relative vorticity of ED1 reaches a maximum of 1.2 f. 4.2.1.4. Stage IV (7 October): Maximum Size [32] ED1 absorbs another eddy pair (ED4) originating in the Barrow Canyon jet. Its horizontal scale is increased to more than 80 km (Figure 9d). 4.2.1.5. Stage V (November): Decay [33] Each eddy gradually shrinks and loses their vorticity due to sea ice drag and lateral friction. Thus the lifetime of these eddies can be estimated to be several months. [34] The time series of the relative vorticity fields along the Chukchi and Beaufort shelf breaks illustrates spatial and temporal variations of the mesoscale eddy behaviors (Figure 10a). North of the Barrow Canyon mouth, eddy pairs composed of an anticyclone and a cyclone are frequently yielded. East of the canyon, timing of eddy generation is placed at the end of July and the middle of September. In each event, a few anticyclones are simultaneously spawned from the eastward shelf break current. Here, the eddy pair originating in the Barrow Canyon mouth is named Type I eddy. The eddy groups generating along the shelf break in early and late summer are named Type II and Type III eddies, respectively. 9 of 16

Figure 10. Spatial and temporal variations in relative vorticity at 30 m depth along the Chukchi and Beaufort shelf breaks in the (a) 2003 and (b) 2007 cases (s 1). The horizontal axis corresponds to east west location. The right end is 140 W and the left end is 160 W. Vertical axis shows the time series from July to October. 4.2.2. Type I Eddy [35] The mechanisms controlling generation and development of Type I eddy are focused on. A few previous studies propose that the outflow of a low potential vorticity (PV) water from a strait toward an offshore high PV region could induce nonlinear evolution of an eastward coastal current due to upstream propagating instability waves along the PV front and produce a large anticyclonic eddy near the strait (Figure 11a). The original theory of Kubokawa [1991] concludes that the anticyclonic eddy could be created when the layer thickness of an inshore low PV water is sufficiently larger than the thickness of an offshore high PV water. Shimada et al. [2001] refer to this theory as a possible formation mechanism of the Beaufort shelf break warm eddies. They assume that the Barrow Canyon plays a role in the strait, and the Pacific summer water corresponds to the low PV water. Here this scenario is addressed using the result in the 2003 case. The inshore layer thickness of the Pacific summer water is represented by the depths of potential density of 24.5s and 26s on the inshore side of the canyon (Figure 11b). The layer thickness rapidly thickens according to the northward transport of the buoyant Pacific water at the beginning of July (Figure 11c). The change in isopycnal depth on the offshore side is relatively small during the summer season. During this period, the Barrow Canyon jet exists along the PV front (Figure 11b). The difference in layer thickness between inshore and offshore sides of the canyon axis supports the possibility of eddy formation along the PV front from July to September, as explained by the Kubokawa s [1991] theory. Thus the transport of the low PV Pacific summer water through the Barrow Canyon might be regarded as a formation mechanism of Type I eddies. [36] The Barrow Canyon jet has an ageostrophic component, while the theory of Kubokawa [1991] assumes that the outflow satisfies a quasi geostrophic balance. The lateral shear of the Barrow Canyon jet reaches 1.2 f along both sides of the jet axis at the end of July (Figure 11d) when one of the Figure 11. (a) Schematic image of Kubokawa s [1991] theory. A low PV water outflow from a strait or a canyon toward an offshore high PV region creates an anticyclonic eddy when the PV difference is large. (b) Potential density (shading, kg m 3) and horizontal ocean velocity (black contours, cm s 1) in the Barrow Canyon section on 21 July. (c) Seasonal cycle of layer thickness between 24.5s and 26s depths of potential density averaged on the inshore and offshore sides of the canyon (three grids from the canyon axis, respectively) (m). (d) Ocean velocity and its lateral shear at 30 m depth on 31 July (s 1). Unit vector is 50 cm s 1. Figures 11b 11d are those in the 2003 case. 10 of 16

Figure 12. Vertical structure of potential vorticity (shading, 10 9 m 1 s 1 ) and potential density (contours, kg m 3 ) on (a) 12 July and (b) 10 September in the 2003 case. Type I eddy pairs arises (Stage II). The corresponding Rossby number occasionally becomes so large that the jet could be unstable by its lateral velocity shear. The time series of maximum velocity through the Barrow Canyon shows that the jet is frequently intensified throughout the summer season (Figure 8a). The jet strength has its maximum above 1 m s 1 on 4, 18, and 28 July; 5 and 25 August; and 11 September, whenever the signal of Type I eddy emerges. In addition to the increase of the Bering Strait throughflow, local daily westerly wind assists the jet intensification. The rapid disappearance of cyclonic parts of the eddy pairs is simply explained by their density anomaly. The cold core cyclone is denser than the surrounding area due to local upwelling in the eddy interior at the time of its generation. However, the density gradient around the cyclone becomes reduced according to its advection toward the offshore cold and saline region. Although D Asaro [1988] hypothesizes that frictional torques on the inshore wall of the Barrow Canyon lead to the generation of anticyclonic eddies, the axisymmetric profile of jet velocity in the 2003 case indicates the minor contribution of bottom friction to the eddy generation. Thus it is suggested that the generation of Type I eddies is induced by a combination of the outflow of the low PV Pacific summer water from the Barrow Canyon and the barotropic instability of the canyon jet due to its lateral velocity shear. 4.2.3. Type II and Type III Eddies [37] Figures 12a and 12b show the vertical structure of potential vorticity and potential density at the time when Type II and Type III eddies are generated. Both eddies are characterized by low PV. There exist remarkable differences in origin depth and background hydrographic structure between the Type II and Type III eddies. The Type II eddy is spawned from the shelf break wall at the depth of 50 to 150 m. The Type III eddy is generated from the isopycnal front between shelf and basin in the surface layer. The other eddies have similar structures in each type. The bottom trapped cold core eastward flow exists along the Beaufort shelf break from May to early July. The jet intensity is so weak during the period that no instabilities could arise. In the 2003 case, sea ice retreats from the northwestern coast of Alaska by the end of June when the intensification of northeasterly wind and thermal forcing promotes sea ice melting along the coastal region. The sea ice reduction results in freshening of the Pacific water passing through the Barrow Canyon. In the latter half of July, the warming of the canyon outflow is caused by the enhanced atmospheric radiative fluxes into the ocean surface in the upstream region. The canyon outflow including the Pacific water, which has a potential density of 24s 27s, intrudes along the shelf break wall at the depth of 60 120 m and creates the bottom trapped jet at the generation place of Type II eddies. The central jet axis is located at 100 m depth. The jet strength is about 15 cm s 1. Spall et al. [2008] mention that Pacific summer water is possibly dense enough to flow eastward as a bottom trapped boundary current. After the generation of Type II eddies, the warmer and fresher water originating in the surface layer of the Barrow Canyon is involved into the eddies that still stay along the shelf break. The eddies gradually obtain buoyancy through the transformation process and are detached from the shelf break toward the offshore surface layer. The relationship between the buoyancy and the detachment is still unclear. In late summer, the more lightened Pacific water lifts a warm core jet to the surface layer along the Beaufort shelf break. The Type III eddies are produced by this surface intensified jet. The jet strength is more than 30 cm s 1, which is obviously larger than the bottom trapped jet detected in early summer. The jet width is about 20 km. Vigorous reinforcement of the Barrow Canyon outflow and consequent shelf break current support the energetic eddy productions (Figure 8a). The existence of such surface and bottom intensified jets along the shelf break is indicated by previous mooring observations, and it is suggested that the intense narrow jet is capable of producing mesoscale eddies by baroclinic instability [Pickart, 2004; Spall et al., 2008]. These warm core eddies survive until the end of November, when sea ice covers the entire Beaufort Sea. 4.3. Pacific Water Transport [38] The distribution of a virtual tracer associated with the Pacific water provided at the Bering Strait demonstrates that the Pacific water is basically transported by northward current over the Chukchi shelf (Figure 13a). The Pacific water content is calculated by the integration of concentration of the Pacific water tracer from the ocean surface to a sea bottom in each grid. The Alaskan Coastal Current primarily transports the Pacific water northeastward during summer, and a significant part of the tracer seems to be then transported into the Canada Basin interior by mesoscale eddy activities over the Beaufort shelf break. The Pacific water tracer field in October shows a large distinct eddy with a horizontal scale of 80 km 11 of 16

Figure 13. Pacific water content on 10 October in the (a) 2003 and (b) 2007 cases (m). The definition is described in section 4.3. Shelf basin boundary is shown by a pink line. Barrow Canyon and 152 W sections are shown by yellow lines. These sections are referred in Figure 14. in the southern Beaufort Sea (Figure 13a). A few mesoscale eddy like features also appear along the shelf break. Some of the Pacific water traces into the Chukchi Sea western branch and flows into the Herald Canyon and the Central Channel. These portions eventually turn to flow eastward over the Chukchi Sea northern shelf. The geostrophic current along isobaths prevents the Pacific water from being separated offshore. The simulated Pacific water distribution in the Chukchi Sea is mostly consistent with the reconstructed fields of Panteleev et al. [2010], which demonstrate that the fraction of virtual particles passing between the Herald Shoal and Cape Lisburne surpasses the Herald Canyon branch. [39] The spatial and temporal variations of the shelf to basin transport of the Pacific water almost coincide with the eddy fields from late summer to early autumn. Figure 14 shows the seasonal cycle of volume transport weighted by the Pacific water concentration throughout the integration period. In this calculation, the shelf basin boundary is defined by isobaths of 2000 m east of the southeastern corner of the Northwind Ridge and of 500 m west of the corner shown by a pink line in Figure 13. The net transport has several peaks at the time when the eddy activities are enhanced. The seasonal mean of the shelf to basin Pacific water transport from July to October is 0.17 Sv, while the Barrow Canyon throughflow is 0.37 Sv (Table 2). A part of the Pacific water passing through the Barrow Canyon sometimes directly intrudes onto the eastern shelf during the summer season, as discussed by Okkonen et al. [2009]. However, the transport on the shelf across the 152 W section is notably less than the shelf to basin transport (Table 2). This volume is relatively small compared to the mooring estimate by Nikolopoulos et al. [2009] where the annual mean is 0.13 ± 0.08 Sv. This discrepancy is probably caused by definition of the Pacific water. Nikolopoulos et al. [2009] determine the Pacific/ Atlantic water boundary by the high PV layer at about 180 m depth and regard both summer and winter water masses as the Pacific water. In contrast, the simulated Pacific water tracer representing the summer water is mainly distributed above the upper halocline in the 2003 case. Note that both estimates indicate that the Pacific water scarcely reaches the Canadian northern shelf. Therefore, it is found that approximately half of the Pacific water flowing through the Barrow Canyon during summer is transported into the Canada Basin; the rest remains in the Alaskan Beaufort shelf or is transported westward toward the Chukchi Sea shelf break. Although Spall [2007] assumes a possibility of offshore transport in the upstream region of the Barrow Canyon, it appears that only a little fraction of the Pacific water tracer splits westward before it passes through the canyon in the 2003 case. [40] In winter, robust easterly wind associated with the strengthened SLP over the central Beaufort Sea promotes westward advection of most of the Pacific water. Some of the Pacific water tracers go northward across the shelf break and are counted as shelf to basin transport. This transport accounts for a noticeable peak of 0.42 Sv in December (Table 2). The Herald Canyon north of the Chukchi western section becomes the primary passage of the Pacific water tracer rather than the Barrow Canyon (Table 2). The eddy like Pacific water pool in the southern Beaufort Sea moves and diffuses toward the Northwind Ridge. Spall [2007] calculates that the annual northward transport from the Chukchi Sea to the Canada Basin interior is 0.22 Sv in his normal experiment and 0.03 Sv in the no wind case. The difference would correspond to the offshoreward Ekman transport. However, the estimated volume might be underestimated due to lack of the eddy induced transport in his model, which does not explicitly reproduce the Beaufort shelf break eddies. Figure 14. Seasonal cycles of Pacific water transport across (a) shelf basin boundary and (b) Barrow Canyon and 152 W sections in the 2003 case. 12 of 16