Barrow Canyon volume, heat, and freshwater fluxes revealed by long-term mooring observations between 2000 and 2008

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1 JOURNAL OF GEOPHYSICAL RESEARCH: OCEANS, VOL. 118, , doi: /jgrc.20290, 2013 Barrow Canyon volume, heat, and freshwater fluxes revealed by long-term mooring observations between 2000 and 2008 Motoyo Itoh, 1 Shigeto Nishino, 1 Yusuke Kawaguchi, 1 and Takashi Kikuchi 1 Received 5 December 2012; revised 21 March 2013; accepted 25 June 2013; published 10 September [1] Barrow Canyon, in the northeast Chukchi Sea, is a major conduit for Pacific Water to enter the interior Arctic basins. Assemblies of annual (September 2000 to August 2008) temperature, salinity, and velocity data acquired from a mooring array in the mouth of Barrow Canyon and high-resolution hydrographic and velocity transects along the mooring array in 2002, 2010, and 2011 have enabled a direct computation of volume, heat, and freshwater fluxes. Annual mean volume transport through Barrow Canyon was 0.45 Sv, which consisted of 0.44 Sv of Pacific Water and 0.01 Sv of Atlantic Water. Annual mean Pacific Water transport through Barrow Canyon represents 55% of the long-term mean Pacific Water inflow through the Bering Strait. During summer, more of the Pacific inflow was advected to an eastern path as the Alaskan Coastal Current that flows along the Alaskan coast to Barrow Canyon. The freshwater flux through Barrow Canyon was 904 km 3 /yr, which is equivalent to 5% of the freshwater content of the Canada Basin. The annual averaged heat flux displayed substantial interannual variability, ranging from 0.93 to 3.02 TW, which could melt 88, ,000 km 2 of 1 m thick ice. A relationship exists between the measured Barrow Canyon transport and local winds such that under southerly winds the northward flow of water through the canyon increases. Such wind conditions are induced by the weaker sea level pressure contrast between the Arctic and North Pacific, caused by a decrease in the pressure over the Arctic and an increase over the North Pacific. Citation: Itoh, M., S. Nishino, Y. Kawaguchi, and T. Kikuchi (2013), Barrow Canyon volume, heat, and freshwater fluxes revealed by long-term mooring observations between 2000 and 2008, J. Geophys. Res. Oceans, 118, , doi: /jgrc Introduction [2] Interest in Pacific Water flowing from the Bering Strait into the Arctic basins has increased markedly in recent years. There is evidence of increasing heat entering the Arctic Ocean through the Bering Strait [Woodgate et al., 2006, 2010]. The heat significantly contributes to both sea-ice melt in summer and a decrease in sea-ice formation during winter because this water typically subsides just below the surface mixed layer in the Canada Basin [Steele et al., 2004; Shimada et al., 2006]. During the extreme Arctic sea-ice retreat in 2007, the heat flux through the Bering Strait was sufficient to cause one third of the seasonal Arctic sea-ice loss [Woodgate et al., 2010]. Additionally, Pacific Water contributes up to 75% of the freshwater input to the Canada Basin [Yamamoto-Kawai et al., 2008]. Decadal variation in this freshwater storage in the Canada Basin is believed to affect global thermohaline 1 Research Institute for Global Change, Japan Agency for Marine-Earth Science and Technology, Yokosuka, Japan. Corresponding author: M. Itoh, Research Institute for Global Change, Japan Agency for Marine-Earth Science and Technology, Yokosuka, Kanagawa , Japan. (motoyo@jamstec.go.jp) American Geophysical Union. All Rights Reserved /13/ /jgrc circulation by freshwater transport to the deep convection area in the northern North Atlantic [Komuro and Hasumi, 2007]. Pacific Water, being rich in nutrients, also influences productivity and the biological pump in the Canada Basin [Nishino et al., 2011]. [3] Pacific Water from the Bering Strait crosses the wide and shallow Chukchi shelf and then flows into the Arctic basins (Figure 1). Although seasonal and interannual variations in water volume and freshwater and heat fluxes through the Bering Strait have been well documented since 1990 [Roach et al., 1996; Woodgate et al., 2005a], north of the strait the exact pathways by which Pacific Water reaches the Arctic basins are still uncertain. Over the Chukchi shelf, the flow from the Bering Strait is believed to follow three different branches [Weingartner et al., 2005]. One route is the Alaskan Coastal Current in the eastern Chukchi Sea [Paquette and Brouke, 1974]. Another branch extends to the west through the Hope Valley into Herald Canyon [Weingartner et al., 1998], and a third branch flows through the Central Channel between the Herald and Hanna shoals [Weingartner et al., 2005]. In summer and early fall when the volume and heat fluxes through the Bering Strait increase to their maximum values, an important part of the flow is the Alaskan Coastal Current, which is forced by the pressure difference between the Pacific and Arctic oceans and modified by wind and buoyancy forcing. The Alaskan Coastal Current flows through the eastern path along the 4363

2 Figure 1. The Chukchi Sea, showing major topographic features, obtained from the International Bathymetric Chart of the Arctic Ocean. The white arrows show schematic pathways of Pacific Water inflow. The rectangular region indicates the location of the enlarged view of the Barrow Canyon region presented in the lower right panel. Solid circles denote the sites of the mooring array. Solid lines indicate the hydrographic transects undertaken by the R/V Mirai and USCGC Healy. The location of the NCEP wind data used in section 6 is indicated by a plus symbol. Alaskan coast to Barrow Canyon [Paquette and Brouke, 1974]. [4] Mooring observations in Barrow Canyon have revealed persistent northward flow that is strong in summer and weak in winter [Aagaard and Roach, 1990; Weingartner et al., 1998, 2005]. In summer the Alaskan Coastal Current carries warm, fresh Alaskan Coastal Water (ACW) into the canyon as a surface-intensified jet [Paquette and Brouke, 1974; Munchow and Carmack, 1997]. In fall and winter, cooling and ice formation over the Chukchi Sea lead to the formation of cold and relatively saline winter waters with salinities of (hereafter referred to as Pacific Winter Water (PWW)) [Weingartner et al., 1998]. During the strong northerly wind conditions in late fall and winter, the northward flow of Pacific Water is retarded, and warm deep water of Atlantic origin (Atlantic Water (AW)) in the Arctic basins is transported occasionally into the canyon [Aagaard and Roach, 1990]. Water-mass modification and currents along Barrow Canyon have been well documented from mooring observations [e.g., Weingartner et al., 1998, 2005]. However, the transport through Barrow Canyon into the Arctic basins has remained uncertain because the mooring coverage was, in general, too sparse for adequate spatial resolution of the flow field. [5] Munchow and Carmack [1997], Weingartner et al. [2005], Pickart et al. [2005], Okkonen et al. [2009], and Shroyer and Plueddemmann [2012] reported ship-based synoptic observations of spatial hydrographic fields in Barrow Canyon that capture the entire width of the northeastward flow through the canyon. The synoptic estimates of Barrow Canyon transport vary widely and exceeded 1.0 Sv (1 Sv ¼ 10 6 m 3 s 1 ) in September 1993 [Munchow and Carmack, 1997]. However, these ship-based observations were limited to the summer and early fall because of the severe weather and sea-ice conditions in the Arctic Ocean in other seasons. Despite their importance to the Arctic Ocean climate, in situ observations of Barrow Canyon and estimates of the transport are limited, particularly with regard to understanding the temporal variation. [6] Our study focuses on the quantitative estimate of volume transport through Barrow Canyon by combining mooring observations with hydrographic surveys, in addition to hydrographic and ocean current conditions. We conducted year-round mooring measurements at one station from September 2000 to September 2001 and at three stations from October 2001 to August 2008 in the mouth of Barrow Canyon (Figure 1). High-resolution hydrographic and velocity data along the mooring array were also obtained in September 2002, September 2010, and July and October Section 2 shows the observations and data in detail. These hydrographic data could improve the estimate of property fluxes obtained from the mooring observations alone by interpolating the discrete mooring observations with the hydrographic data. In section 3, we investigate the synoptic hydrography and current conditions across Barrow Canyon in summer and early fall. Section 4 describes the seasonal variation of the water masses and flow field in Barrow Canyon based on mooring data. The volume transports through Barrow Canyon are quantitatively estimated and discussed in section 5. Section 6 examines the relationship between Barrow Canyon transport and wind. This relationship allows us to extend our estimation of Barrow Canyon transport from 1985 to We conclude with a summary in section

3 2. Data [7] The locations of our moorings are shown in Figure 1. Year-round oceanographic moorings have been deployed at one location in the center of Barrow Canyon (Stn. BCC) from September 2000 to September 2001 and at three locations in the eastern shelf (Stn. BCE), the center (Stn. BCC), and the western shelf (Stn. BCW) of the canyon from October 2001 to August Briefly, the array spanned a total distance of approximately 20 km, with horizontal spacing of 10 km (Figure 2). The configurations of the array are shown in Table 1. The configurations from September 2000 to September 2006 have been also reported previously [Itoh et al., 2012]. [8] To estimate transport through Barrow Canyon, temperature, salinity, and velocity data were linearly interpolated to every 1 m between the observational levels. The shallowest data were extended from the shallowest observational level to the surface, and the deepest data were extended from the deepest observational level to the bottom. With regard to sea-surface temperature (SST), daily data recorded by the Advanced Microwave Scanning Radiometer of the Earth Observing System (AMSR-E) were available during ice-free months (August-October) from 2002 onward, although the horizontal resolution was not good enough to resolve narrow ACW temperature differences. We therefore adopted warmer values from AMSR-E SST or measured the temperature using the shallowest instrument at the mooring site during ice-free months. [9] Overall, the equipment yielded very good data as shown in Figure 2 and by Itoh et al. [2012]. However, some data shortfalls occurred at Stn. BCC. The temperature and salinity instrument moored at a depth of 50 m failed from September 2002 to September We used averaged data for Stns. BCE and BCW to represent temperature and salinity at 50 m depth at Stn. BCC over that period. The upwardlooking acoustic Doppler current profiler (ADCP) moored at 150 m depth at Stn. BCC also failed from September 2002 to September Since the velocity at Stn. BCC largely corresponded to that at Stn. BCE rather than at Stn. BCW, we used the velocity data from Stn. BCE to represent Stn. BCC in the upper 100 m over that period. In addition, the ADCP that measured the velocity profile between the surface and 250 m at Stn. BCC also failed from June 2004 to September Therefore, there are approximately 16 months of missing data for depths deeper than 60 m at Stn. BCC. [10] Hydrographic surveys using a conductivitytemperature-depth (CTD) sensor or expendable CTD (XCTD) sensor were undertaken along the mooring array in Barrow Canyon in September 2002 and September 2010 by the R/V Mirai andinjulyandoctober2011bytheuscgchealy (Figure 1). These surveys collected high-resolution temperature, salinity, and velocity measurements and were labeled from I to X (Figure 3). Temperature and salinity profiles were taken at nine stations across Barrow Canyon (during surveys I, III, VI, VIII, IX, and X), as shown in Figure 4. We used velocity data obtained from shipboard 75 and 150 khz ADCPs on R/V Mirai and USCGC Healy, respectively. The shipboard ADCP data were corrected for bias and misalignment errors following the procedure of Joyce [1989]. We neglected tidal currents in the analysis of shipboard ADCP data because barotropic tidal currents are very weak in Barrow Canyon [Mountain et al., 1976]. [11] To clarify the relationship between flux in Barrow Canyon and local winds, we used six hourly wind data at 10 m height and atmospheric sea level pressure (SLP) from National Centers for Environmental Prediction/National Center for Atmospheric Research (NCEP/NCAR) reanalysis data to examine long-term variability. Note that the closest grid point for NCEP reanalysis data used in this study was located roughly 100 km to the southwest of the mooring site shown in Figure 1. Figure 2. Configuration of the typical Barrow Canyon mooring array in relation to the bottom topography. The mooring names are indicated at the top of Figure 2. The east side of the canyon is on the right of each panel. Observational levels of (a) temperature and salinity and (b) velocity are indicated by triangles. Vertical lines at Stn. BCC show the depth range of velocity data observed by the ADCP. Percentages indicate the temporal coverage of data in relation to the full record length of September 2000 to August 2008 for Stn. BCC and October 2001 to August 2008 for Stns. BCE and BCW. Details of the record lengths and observational depths are also given in Figure 4 of Itoh et al. [2012]. 4365

4 Table 1. Mooring Configurations a Station Period (mo/d/yr) Instrument Nominal Depth (m) Stn. BCE 30 September 2001 to 9 September 2002 ADCP and CM 67/87 (8 80) CT 48/67/89 9 September 2002 to 5 October 2003 ADCP 80 (3 75) CT and T 52/80/86 5 October 2003 to 1 October 2005 CM 55/75 CT 35/55/92 2 October 2005 to 30 September 2006 CM 40/84 CT 34/40/86 1 October 2006 to 7 October 2007 CM 46/76 CT and T 42/46/76/92 8 October 2007 to 9 September 2008 CM 47/75 CT 42/47/75/92 Stn. BCC 9 September 2000 to 1 October 2001 ADCP and CM 150 (11 139)/195 b /235 b CT and T 40/80/150/158/198/235/ September 2001 to 9 September 2002 ADCP and CM 72/157 (16 152) b /162 ( ) CT and T 52/72/90/132/157/162/190/256 9 September 2002 to 5 October 2003 ADCP 245 ( ) CT 89/158 b /252 5 October 2003 to 1 October 2005 ADCP and CM 57/240 (29 229) b CT 37/78/120 2 October 2005 to 1 October 2006 ADCP 265 (8 240) CT 35/42/72/112/170/265 2 October 2006 to 7 October 2007 ADCP 144 (17 133)/153 ( ) CT 41/48/78/118/144/149/153/178/249/271 7 October 2007 to 31 August 2008 ADCP 142 (16 132)/151 ( ) CT 35/42/72/112/170/265 Stn. BCW 30 September 2001 to 8 September 2002 ADCP 147 (13 141) CT and T 40/60/80/147/150 b /160 9 September 2002 to 5 October 2003 ADCP 146 (16 140) CT 49/82/146/152 5 October 2003 to 2 October 2005 CM 49/89/140 CT 29/49/89/140/158 2 October 2005 to 1 October 2006 CM 46/76/150 CT 39/76/115/151 1 October 2006 to 7 October 2007 CM 52/81/146 CT 45/52/81/121/146/158 8 October 2007 to 9 September 2008 CM 50/80/143 CT 43/50/80/120/143/155 a The acronyms ADCP, CM, CT, and T denote the acoustic Doppler current profiler, current meter, conductivity-temperature sensor, and temperature sensor, respectively. Most of the mooring data were recorded at hourly intervals. Nominal depth of CM and T is underlined. Values in brackets show the depth range of the ADCP. b Sensors that stopped during the observation period. 3. Synoptic Hydrographic and Flow Fields in 2002, 2010, and 2011 [12] In September 2002, September 2010, and July and October 2011, high-resolution temperature, salinity, and velocity surveys along our mooring array in Barrow Canyon were undertaken from the R/V Mirai and USCGC Healy. Here we describe in detail the properties of the current, temperature, and salinity in Barrow Canyon in summer and early fall Water Properties and Throughflow in Barrow Canyon [13] Figure 3 shows the vertically averaged current fields that were detected by the shipboard ADCP. Northeastward flows were generally concentrated over the eastern slope of the canyon for all surveys, with the exception of survey X in late October The flow field on the western side of the canyon was generally weaker than that observed on the eastern side and was also more variable. From daylong repeat surveys on 29 and 30 September 2010 (surveys IV, V, VI, and VII), it was determined that flows along the canyon quickly responded to changes in wind. [14] Figure 4 shows the temperature, salinity, ADCP velocity, and geostrophic velocity. The water in Barrow Canyon is composed of three water masses: warm and fresh ACW (salinity, S 32.0), cold and moderately salty PWW (close to freezing and S 33.1), and warm and salty AW (S 34.8). Warm and fresh ACW was evident in surveys I, III, VI, and VIII. At the shallower depths of less than 100 m, ACW is trapped in the eastern side of the canyon, with a horizontal width of km (Figure 4a). The axis of the strong northeastward-flowing Alaskan Coastal Current lies above the shelf slope and is displaced farther offshore than the core of warm ACW (Figures 4a and 4c). The surface intensification of the Alaskan Coastal Current is partly 4366

5 Figure 3. Vertically averaged velocity vectors obtained by a ship-mounted ADCP for 10 repeat transects of surveys (a) I on 8 September 2002, (b) II and III on 12 September 2010, (c) IV and V on 29 September 2010, (d) VI on 29 September 2010 and VII on 30 September 2010, (e) VIII on 22 July 2011, (f) IX on 21 October 2011, and (g) X on 23 October The gray lines denote 25, 50, 100, 150, 200, 250, and 500 m isobaths. Vectors follow the scale in Figure 3a. Included are the daily-averaged surface wind vectors at Pt. Barrow. The compass circles indicate wind velocities of 4 m/s. explained by a baroclinic geostrophic current caused by a sharp density (salinity) front associated with the existence of warm and fresh ACW over the shelf (Figures 4b 4d). During surveys IX and X in late October 2011, fresh, cold water at 0 C appeared in Barrow Canyon, suggesting that the ACW had cooled by mid-october. Additionally, the coastal current was barotropic because of the horizontal density surface. A strong southward wind continued from 20 to 26 October The current system is in transition toward upwelling beginning on transect IX (21 October 2011) because the isopycnals appear to tilt upward toward the coast. Then the coastal current reversed and cores of southwestward flow were evident in both sides of Barrow Canyon (Figures 3g and 4c) during survey X (23 October 2011). [15] During surveys I, III, and VI in September and survey VIII in July, lateral shears of the along-canyon flow as strong as f, where f is the local Coriolis parameter ( s 1 ), were measured on the shoreward and seaward sides of the core of the Alaskan Coastal Current (Figure 4c). This suggests that the cross-canyon Rossby number, which is the ratio of the relative vorticity derived from the cross-canyon gradient of the along-canyon velocity to the local Coriolis parameter, is occasionally very large and that it is therefore likely nonnegligible ageostrophic effects exist. The values are larger than those observed in Barrow Canyon in September 1993 [Munchow and Carmack, 1997] and in July August 2002 [Pickart et al., 2005]. In contrast, during surveys IX and X in late October, the lateral shear of the along-canyon flow was less than 0.5f. 4367

6 Figure 4. Vertical sections of (a) temperature ( C), (b) salinity, (c) velocity (m/s) observed by a shipmounted ADCP, and (d) geostrophic velocity (m/s) referenced to 300 m for surveys I, III, VI, VIII, IX, and X. The east side of the canyon is on the right of each panel. Northeastward down-canyon flow is positive. Tick marks at the top of Figure 4 correspond to CTD or XCTD stations. The mooring sites are also indicated at the top of Figure 4. (a) White lines indicate ACW/PWW and PWW/AW water-mass boundaries, which correspond to contours of salinity (S ¼ 32.5 and 34.0) because hypersaline PWW did not appear for all surveys. Included are the daily-averaged surface wind vectors at Pt. Barrow. The compass circles indicate wind velocities of 4 m/s. (b) White lines indicate contours of potential density ( ¼ 23, 24, 25, 26, 27). Volume transport values calculated from (c) ADCP velocity and (d) geostrophic velocity integrated from the easternmost station to Stn. BCW are indicated at the lower left of the figures. (c) Solid white lines indicate contours of the cross-canyon Rossby number, Ro ¼ 0.6, 1.0, and 1.4. Dashed white lines indicate contours of the cross-canyon Rossby number, Ro ¼ 1.4, 1.0, and 0.6. (d) To calculate the geostrophic velocity at the point shallower than the reference depth, the value of the geopotential anomaly at the bottom layer is extrapolated by extending that of the next western (next eastern) station horizontally from the central of the canyon toward the eastern (western) side, following the procedure of Itoh and Ohshima [2000]. [16] Cold PWW typically occupies depths of m in the canyon, except for the eastern shelf where thick ACW extends down to m depth (Figures 4a and 4b). Warm and saline AW is dominant at m depth in the central canyon. During survey III, a relatively strong northeastward flow was observed in the Atlantic layer as well as in the upper layer on the eastern side of the canyon. In contrast, the AW flow was reversed, being on the western side of the canyon during surveys I and VI. Because Barrow Canyon is about 3 times wider than the ratio of inertial deformation, such baroclinic flows can become established in the canyon, as suggested by Munchow and Carmack [1997]. 4368

7 3.2. Transport Estimation From Synoptic Hydrographic Surveys in Barrow Canyon [17] Here we examine the properties of fluxes through Barrow Canyon, using the hydrographic and shipboard ADCP survey data collected in 2002, 2010, and The volume flux peaked approximately 8 16 km from the easternmost station, mainly because of the strong Alaskan Coastal Current, and decreased to almost zero at a distance of approximately km (Figures 5a and 5b). Therefore, we defined the total transport through Barrow Canyon as volume transport integrated from the easternmost station to Stn. BCW. By combining hydrographic and shipboard ADCP data, we also estimated freshwater and heat fluxes through Barrow Canyon (Figures 5c and 5d). Note that the reference salinity used to calculate the freshwater flux was 34.8, following Aagaard and Carmack [1989]. Since we were interested in the heat available for sea-ice melting, we Figure 5. (a) Horizontal profiles of volume transport (m 2 / s) per unit horizontal length for surveys I X. (b) Horizontal profiles of depth-averaged flows along the canyon (m/s) in 0 to 50 m intervals over Barrow Canyon for surveys I X. Northeastward down-canyon flow is positive. The solid black line indicates the velocity averaged for surveys I IX, when the Alaskan Coastal Current was dominant. Horizontal profiles of (c) freshwater flux (m 2 /s) and (d) heat flux per unit horizontal length (MJ/ms) for surveys I, III, VI, VIII, IX, and X. (e) Bottom topography of this line. The mooring locations are indicated at the top of Figure 5. used the freezing temperature depending on its salinity as a reference to calculate the heat content and flux. [18] We here define three water masses, ACW, PWW, and AW, in order to examine the composition of the property fluxes for surveys I X. A water mass with salinity of S < 32.5 is defined as ACW. The boundary of PWW and AW is about S 34.0 (e.g., Figure 4). However, enhanced brine rejection within the coastal polynya occasionally increases the salinity of PWW to form the hypersaline PWW (S > 34.0), and this hypersaline PWW enters the Arctic basins through Barrow Canyon [Weingartner et al., 1998]. Thus, we here define warm water (T > freezing temperature þ1 C) with S > 34.0 as AW. Water masses with 32.5 < S < 34.0 and cold water (T < freezing temperature þ1 C) with S > 34.0 are defined as PWW. We note that this definition of PWW in Barrow Canyon is not appropriate for early winter because less saline PWW with S < 32.5 enters the Chukchi Sea through the Bering Strait [Woodgate et al., 2005a, 2005b] and appears in Barrow Canyon in December and January [Woodgate et al., 2005b; Weingartner et al., 2005]. [19] Volume, heat, and freshwater fluxes through Barrow Canyon, which were estimated from the hydrographic and shipboard ADCP surveys, are summarized in Figure 5 and Table 2. The volume flux was highly variable (from 0.20 to 2.22 Sv) over all surveys. When very warm ACW was dominant (surveys I and III), the estimated volume flux through Barrow Canyon increased by 30 50% due to a large baroclinic geostrophic current, as shown in Figures 4c and 4d. Daylong repeat surveys on 29 and 30 September 2010 (surveys IV, V, VI, and VII) also highlighted the high variability of the volume flux, which changed by a factor of 2 within a day, possibly owing to changes in the wind. Furthermore, when a strong northerly wind continued from 20 to 26 October 2011, volume transport changed from 0.99 (survey IX) to 0.20 Sv (survey X) within 2 days. [20] Heat and freshwater fluxes also varied among the six surveys (Figures 5c and 5d and Table 2). Horizontal distributions of the volume and freshwater fluxes had similar shapes, suggesting that the variation in the freshwater flux was predominantly caused by variations in the distribution of velocity rather than salinity. When warm ACW and the strong Alaskan Coastal Current were dominant over the eastern shelf and slope of Barrow Canyon (surveys I, III, and VIII), large heat and freshwater fluxes were estimated. [21] Table 2 shows volume, heat, and freshwater fluxes for three water masses: ACW, PWW, and AW. Of these water masses, the largest volume flux was found for ACW in the surveys conducted in July and September (surveys I, III, VI, and VIII), as listed in Table 2. The contribution of PWW to the total volume flux was also significant. The AW volume flux was the smallest (sometimes negative) in all of the surveys. ACW also dominated both freshwater and heat fluxes, particularly the larger heat fluxes (e.g., surveys I, III, and VIII). This allowed an approximate estimation of the heat flux through Barrow Canyon to be made within the upper 100 m layer, as discussed in section 5.4. [22] Munchow and Carmack [1997] reported that the Barrow Canyon heat flux referenced to 0 C was 5 TW, which was estimated from a synoptic survey on 9 September Heat fluxes, which were adjusted to the reference value used by Munchow and Carmack [1997], during 4369

8 Table 2. Volume Transport (Sv) Through Barrow Canyon, Calculated From High-Resolution Shipboard ADCP Data Obtained for 10 Repeat Transects of Surveys I X a Volume Transport (Sv) Freshwater Flux (msv) Heat Flux (TW) Survey Total ACW PWW AW Total ACW PWW AW Total ACW PWW AW I II 2.00 III IV 1.34 V 1.11 VI VII 0.50 VIII IX X a Freshwater flux (msv) and heat flux (TW) calculated from high-resolution ship-based CTD and ADCP data were obtained during surveys I, III, VI, VIII, IX, and X. Fluxes of Alaskan Coastal Water (ACW, S < 32.5), Pacific Winter Water (PWW, 32.5 < S < 34.0 and cold water (T < freezing temperature þ1 C) with S > 34.0), and Atlantic Water (AW, T > freezing temperature þ1 C with S > 34.0) are also listed. The fluxes are integrated from the easternmost station to Stn. BCW. surveys I and III in early September 2002 and 2010 were 2 4 times larger than the fluxes observed in early September 1993, mainly because of the warming of ACW. It is likely that Barrow Canyon heat fluxes in early September, when the ACW was warmer in our surveys than in 1993, has increased. In surveys VI, IX, and X, all in late September and October, the heat fluxes through Barrow Canyon were small (7.9, 5.5, and 1.2 TW) because the ACW had cooled (Table 2 and Figure 4a). The heat flux during survey VI in late September 2010 was similar in magnitude to that observed in late September 1993 [Munchow and Carmack, 1997]. 4. Seasonal Variability of Current, Temperature, and Salinity Revealed by Mooring Observations [23] Here we use mooring data to examine the seasonal variability of current, temperature, and salinity in Barrow Canyon. The monthly mean depth-averaged velocity vectors at all mooring sites are shown in Figure 6. As explained previously, the major flow in Barrow Canyon was essentially parallel to the isobaths. Northeastward flow was particularly evident at Stns. BCE and BCC on the eastern side of the canyon from April to September (23 cm/ s), with a maximum of 43 cm/s in July. In contrast, the current speed was lower (4 cm/s) and more variable at Stn. BCW on the western side of the canyon. Northeastward flow was smaller in fall and winter than in summer at all locations. [24] Figures 7a 7c show the typical seasonal variability on the vertical structure of temperature, salinity, and alongcanyon velocity at Stn. BCC between October 2006 and September Cold PWW was observed from the surface to m from late November to mid-june (Figures 7a and 7b). During early winter from November to January, PWW near the surface had low salinity, between S ¼ 31.0 and 32.0, because the salinity of water entering Bering Strait was relatively low in this period [Woodgate et al., 2005a]. The salinity of PWW occasionally increased to S ¼ in early spring due to salinization within coastal polynyas [Weingartner et al., 1998]. Related to northeastward flow of PWW into the Canada Basin, Figure 7c shows that the maximum flow depth sank to the intermediate layer at a depth range of m from November to mid-june. This occurred because PWW is denser than local surface water and thus subsides below the surface water. The maximum velocity across Barrow Canyon was observed at Stn. BCC where a core of PWW was present during this period (Figure 6). Figure 7a shows that warm subsurface water started to intrude in the middle of June and the warmest ACW was observed in August and September. Figure 7c shows that the northeastward current was strong with a maximum near the surface in summer because of the flow of warm ACW. From spring to summer, the velocity core migrated eastward from Stns. BCC to BCE (Figure 6). 5. Estimation of Transport From Mooring Data in Barrow Canyon [25] In this section we examine the seasonal and interannual variabilities in volume, freshwater, and heat fluxes using the time series of moored temperature, salinity, and velocity data at Stns. BCE, BCC, and BCW for the period Interpolation/Extrapolation of Mooring Data Onto the Barrow Canyon Section [26] First, to estimate accurate Barrow Canyon fluxes, we considered how to interpolate/extrapolate the velocity data measured at Stns. BCE, BCC, and BCW onto the section across Barrow Canyon. We note that the mooring section does not capture all of the transport as indicated in the ADCP sections (Figure 5) because either side of the mooring section should be extrapolated. Velocity distribution data from the easternmost hydrographic station to Stn. BCW are needed to estimate the fluxes through Barrow Canyon because the strong along-canyon current decreased to almost zero at Stn. BCW (Figures 5a), as mentioned in section 3.2. Therefore, we set a 28 km section along the mooring array, which consisted of 13 points at km horizontal intervals, as shown in Figure 8. The following three methods were applied to interpolate/extrapolate velocity distributions. In method I, we used the data for Stns. 4370

9 Figure 6. Monthly mean vertically averaged velocity vectors and associated standard error ellipses using mooring data for at Stns. BCE, BCC, and BCW in (a) January, (b) February, (c) March, (d) April, (e) May, (f) June, (g) July, (h) August, (i) September, (j) October, (k) November, (l) and December. For Stn. BCC data from July 2004 and September 2005 were not used because of missing data for the deeper depths. The gray lines denote 25, 50, 100, 150, 200, 250, and 500 m isobaths. Vectors follow the scale in Figure 6a. BCE, BCC, and BCW as the velocity values for the eastern, central, and western sections of the canyon with widths of 11, 11, and 6 km, respectively (Figure 8a). In method II, velocity data were linearly interpolated between the mooring stations. Furthermore, a velocity profile measured at Stn. BCE was extended to the easternmost point (Figure 8b). Method III is similar to method II except for the velocity extrapolation in the eastern side of the canyon. A reduction factor, which indicates the velocity distribution slowing to 20% of the speed within 10 km of the eastern side of Stn. BCE (solid black line in Figure 5b), was applied for the points from Stn. BCE to the easternmost point (Figure 8c). Using the velocity distribution calculated by methods I III, we could estimate the volume flux across Barrow Canyon. [27] To evaluate these methods, we examined the differences in the volume transports by comparing the highresolution ship-based data obtained in 2002, 2010, and 2011 with the same data subsampled to match the mooring horizontal sensor locations. Volume transports were calculated from the high-resolution shipboard ADCP data (Table 2) and were compared with the values estimated from the subsampled velocity data using interpolation methods I III (Table 3). Volume flux estimations were sensitive to the different interpolation methods adopted as shown in Table 3. In general, methods I and II overestimated the volume flux. This was likely to be a result of the misleading velocity distribution in the inshore region east of Stn. BCE. Estimations using method III were the most reliable of the examined methods. The average error in method III was þ0.067 Sv, which was much less than those of methods I and II. We therefore applied method III for the calculation of volume transport in this study. [28] Estimations of freshwater and heat fluxes are sensitive to the interpolation methods used for temperature and salinity determination in addition to the velocity distribution. Temperature and salinity data were linearly interpolated between mooring stations, and values from Stn. BCE were extended toward the eastern side. To assess the error in the fluxes resulting from this method, we examined the differences in the freshwater and heat fluxes by comparing the high-resolution ship-based data with the same data subsampled to match the mooring vertical and horizontal sensor locations (Figure 2). The subsampled data provided pseudomooring data for the comparison. Fluxes were calculated from the high-resolution hydrographic and 4371

10 Figure 7. Depth-time plots of (a) temperature ( C), (b) salinity, and (c) along-canyon velocity (m/s) at mooring Stn. BCC from October 2006 to September Northeastward down-canyon flow is positive. The black arrows indicate the nominal depth of the temperature and salinity sensors. (a) Sea-surface temperature was obtained from AMSR-E data. Periods with sea-ice cover (>80%) are shaded at the top. (a) The white lines indicate contours of temperature (T ¼ freezing temperature þ1 C) delimiting the cold water. (b) White lines indicate contours of salinity (S ¼ 32.5 and 34.0). (c) The white lines in Figure 7c indicate the 10 day running average for the maximum flow depth. Figure 8. Schematic views of (a) method I, (b) method II, and (c) method III used to interpolate/extrapolate the velocity distribution across Barrow Canyon using the mooring data at Stns. BCE, BCC, and BCW. Locations of moorings are indicated at the top. The grid location used in section 5.1 is also indicated at the top of Figure 8. The horizontal intervals of the grid are km. Velocity data measured at Stns. BCE, BCC, and BCW were interpolated/ extrapolated at every 1 m depth. shipboard ADCP data (Table 2) and were compared with the values estimated from the pseudomooring data after linearly interpolating the temperature and salinity as described above and using interpolation method III for velocity distribution (Table 3). It should be noted that vertical resolution of temperature and salinity was poor at Stn. BCC from September 2002 to September 2003 [see Itoh et al., 2012, Figure 4], and the errors of the estimated fluxes for this reduced resolution are larger than when the resolution was increased. We determined errors for both normal hydrographic data coverage (e.g., October 2006 to September 2007) and for the poorest hydrographic data coverage (e.g., September 2002 to September 2003). The average errors for freshwater and heat fluxes for the six surveys from normal (poorest) data resolution were 14% (19%) and 3% (16%) of the total fluxes, respectively (Table 3). [29] We also noted that the fluxes though Barrow Canyon, calculated from data for Stns. BCE, BCC, and BCW, correlated well with fluxes calculated at Stn. BCC alone [Itoh et al., 2012]. Therefore, for September 2000 to September 2001, when only Stn. BCC data are available, the properties of fluxes through Barrow Canyon can be determined by applying a coefficient determined from a linear regression when data from all stations are available Volume Transport [30] Figure 9a shows a time series of volume transport through Barrow Canyon, which was estimated using Method III from mooring data from September 2000 to 4372

11 Table 3. Differences in the Volume Transport (Sv) Between Values Calculated From High-Resolution Shipboard ADCP Data in Table 2 and Values Estimated From the Subsampled Velocity Data Using Interpolation Methods I III a Volume Transport Difference (Sv) Freshwater Flux Difference (msv) Heat Flux Difference (TW) Survey Method I Method II Method III Normal Case Worst Case Normal Case Worst Case I þ0.243 [31%] þ0.114 [15] þ0.042 [5] þ10 [10%] 47 [45%] þ1.40 [7%] þ8.93 [45%] II þ0.614 [31%] þ0.518 [25] þ0.326 [16] III þ0.505 [23%] þ0.390 [18] þ0.179 [8] þ11 [7%] 10 [7%] þ1.19 [3%] þ6.34 [18%] IV þ0.073 [5%] þ0.032 [2] [6] V þ0.233 [21%] þ0.176 [16] þ0.068 [6] VI þ0.446 [86%] þ0.423 [81] þ0.293 [56] 8 [14%] 18 [31%] 0.88 [11%] 3.18 [40%] VII þ0.028 [6%] þ0.034 [7] [15] VIII þ0.252 [16%] þ0.169 [11] þ0.026 [2] þ18 [16%] þ5 [5%] 0.25 [1%] 4.17 [20%] IX þ0.100 [10%] þ0.071 [7] [1] þ2 [3%] þ6 [11%] 0.57 [10%] þ1.41 [3%] X [44%] [65] [52] þ5 [47%] þ14 [127%] þ0.76 [61%] þ1.05 [85%] Average þ0.240 [22%] þ0.179 [16%] þ0.067 [6%] þ6 [14%] 8 [19%] þ0.27 [3%] þ1.73 [16%] RMS [20%] [18%] [13%] 8 [19%] 20 [46%] 0.88 [8%] 4.7 [44%] a Differences in freshwater flux (msv) and heat flux (TW) between values calculated from high-resolution ship-based CTD and ADCP data (Table 2) and values estimated from the pseudomooring data after linearly interpolating the temperature and salinity and using interpolation method III for velocity distribution. Positive values indicate that the estimated values are larger than the observed values. RMS indicates the root mean square. Bracketed values for surveys I X represent the difference values in Table 3 divided by the corresponding transports or fluxes from Table 2. Bracketed values for average and RMS are average and RMS values in Table 3 divided by the average values corresponding transports or fluxes for all surveys in Table 2. August Volume transport is highly variable with short-term changes of Sv. This corresponds to our discussions of the results from the synoptic surveys in section 3. The seasonal variation in the transport is illustrated in Figure 10a. Monthly mean volume transport was typically highest in summer (1.06 Sv in July) and decreased in autumn and winter (0.09 Sv in November). Annual mean volume transport through Barrow Canyon was 0.45 Sv, which consisted of 0.44 Sv of Pacific Water and 0.01 Sv of Atlantic Water. Annual mean Pacific Water transport through Barrow Canyon (0.44 Sv) is equivalent to 55% of the long-term average through the Bering Strait (0.8 Sv) [Woodgate et al., 2005a]. Pacific Water transports through the Barrow Canyon were 0.67 Sv for the May October period and 0.19 Sv for the November April period. These are equivalent to 68% and 29% of the Bering Strait transport for similar periods according to the climatology given by Woodgate et al. [2005a]. Note the important point here is that Woodgate et al. [2005a] give climatological values, which do not include the same years as in our paper. This point needs to be emphasized because the comparisons are not for the same time periods. Comparison of the volume transport through the Bering Strait and Barrow Canyon suggests that the Bering inflow is dominantly advected by the eastern path as the Alaskan Coastal Current and emanates from Barrow Canyon into the Canada Basin during summer, in contrast with a flow toward Herald Canyon or through the Central Channel that is more important in late autumn and winter. The integrated Pacific Water inflow through Barrow Canyon for a year was km 3. Because the total volume of Pacific Water in the Canada Basin was calculated to be approximately km 3, from 30 to 250 m within 71 N 80 N and 130 W 160 W [e.g., Itoh et al., 2012, Figures 2 and 3], Barrow Canyon transport is likely to supply 9% of the total volume of Pacific Water in the Canada Basin annually Freshwater Flux [31] In terms of salinity, the summer/autumn CTD sections indicate that 15 to 20 m thick surface mixed layers are fresher than layers below as shown in Figure 4b. The mooring array cannot monitor surface layers. Therefore, to evaluate the contributions of surface freshening, an additional estimated freshwater flux is needed. Here we assumed that the salinity of the 20 m thick surface mixed layer was 0.7, 1.2, and 2.0 fresher than values observed at 30 to 40 m depth at Stns. BCE, BCC, and BCW, respectively, from June to October, when surface stratification is expected to be distinct because of the presence of fresh meltwater from sea ice and ACW. As a result, an additional msv (1 msv ¼ 10 3 m 3 s 1 ) of freshwater flux should be included, although this is not a substantial contribution to the summer/autumn freshwater flux (Figure 9b). [32] Figures 9b and 10b show that the annual mean freshwater flux during was 29 msv (904 km 3 / yr). This flux tended to be highest in summer (80 msv in July) and lower in winter (6 msv in November) as shown in Figure 10b. Freshwater fluxes through Barrow Canyon were strongly linked to the volume fluxes, with a high correlation (r ¼ 0.97), and were typically about 7% of the volume flux. The annual mean freshwater flux was equivalent to 37% of the freshwater inflow through the Bering Strait [Woodgate et al., 2005a] and could supply 5% of the freshwater content of the Canada Basin [Proshutinsky et al., 2009]. In addition to Pacific Water inflow, river discharges and sea-ice meltwater are also significant sources of freshwater in the Arctic Ocean [Aagaard and Carmack, 1989]. The freshwater flux through Barrow Canyon was nearly one third of the total river discharge into the Arctic Ocean (3200 km 3 /yr) [Serreze et al., 2006] and twice the discharge from North American rivers directly into the Canada Basin (418 km 3 /yr) [Lammers et al., 2001]. The annual mean Barrow Canyon freshwater flux was equivalent to the 4373

12 Figure 9. Seven day (gray line) and 30 day (red line) smoothed time series of (a) volume transport (Sv), (b) freshwater flux (msv), and (c) heat flux (TW) through Barrow Canyon calculated from mooring data at Stns. BCE, BCC, and BCW. The blue line indicates 30 day smoothed fluxes calculated from mooring data at Stn. BCC. (b) The revised freshwater flux for the surface layer during summer (from June to October) is indicated by the green line. (c) Heat flux in the upper 100 m is indicated by the green line. Annual averaged values for each year calculated by data at one mooring site (Stn. BCC) and at three mooring sites (Stns. BCE, BCC, and BCW) are indicated in blue and red characters, respectively, at the top of Figure 9. freshwater flux contained by 966,000 km 2 of 1 m thick ice and 41% of the change in summer ice extent (2,380,000 km 2 ) between 2011 and the average for Heat Flux [33] The heat flux of the whole water column was almost the same as that of upper 100 m water column as shown in Figure 9c because ACW, which is the dominant heat source, is typically present above 100 m as shown in Figures 4a and 7a. Therefore, we used the heat flux for the upper 100 m to represent the heat flux through Barrow Canyon in this study. The heat flux through Barrow Canyon was greatest in summer and the weekly mean heat flux reached 30 TW owing to the presence of the warm ACW in Figures 9c and 10c. It typically falls to zero in mid-october and remains at zero during the winter because the water temperature was generally close to the freezing point, with the exception of deep warm AW (Figure 7a). The heat flux rises above zero in July because of the warm ACW arriving at Barrow Canyon. The annual averaged heat flux through Barrow Canyon into the Arctic basins displayed substantial interannual variability, ranging from 0.93 to 3.02 TW during the period of , which is sufficient to melt 1 m thick ice over an area of 88, ,000 km 2. Woodgate et al. [2010] recorded a heat flux through the Bering Strait of 6 19 TW ( J/yr) from 2001 to Comparison of the heat fluxes through the Bering Strait and Barrow Canyon indicates that the ACW would lose 3 10 TW ( J/yr) within the Chukchi Sea, if we assume that 68% of Bering Strait throughflow takes an eastern path and exits the Chukchi Sea via Barrow Canyon during summer. [34] We next examine in detail the interannual variation in heat fluxes through Barrow Canyon. Figure 11 shows the heat flux, heat content, and volume transport of the upper 100 m in Barrow Canyon from June to October in The interannual variation in heat flux is governed by 4374

13 Figure 10. Monthly averaged (a) volume transport (Sv), (b) freshwater flux (msv), and (c) heat flux (TW) through Barrow Canyon calculated from mooring data. The solid black lines indicate averaged values. Annual averaged values for the duration are indicated in the upper right corner. variations in the volume transport and heat content of water through Barrow Canyon. Summer heat content monotonically increased from 2001 to 2007, with the exception of Within the period, the Barrow Canyon heat content was at a maximum in 2007 and was 1.5 times larger than the averaged heat content from 2001 to Within the general trend of summertime SST increases over the past few decades, the positive temperature anomalies in the northern Chukchi Sea were extraordinarily high in 2007 [Steele et al., 2008], which is consistent with our results. Perovich et al. [2007] established that decreasing sea-ice cover in the region was related to increased solar energy reaching the upper ocean in the Chukchi Sea over the past decade. In addition, Woodgate et al. [2010] showed that the Bering Strait heat flux and heat content have increased recently, reaching a record maximum in 2007 for the period Recent warming of the Chukchi Sea, including Barrow Canyon, can be attributed in part to recent increases in heat flux and heat content measured across the Bering Strait. However, the Barrow Canyon heat flux in 2007 was as high as those in both 2002 and This was largely a result of the smaller volume transport in 2007 as shown in Figure 11. [35] By contrast, the heat flux for the period studied was at its lowest in 2006, when the heat content was smallest, although volume transport was not small in this year (Figure 11). This is consistent with measurements of summertime SST in the northern Chukchi Sea, which reached minimum values in 2006 during the period [Steele et al., 2008]. However, the Bering Strait heat flux and heat content in 2006 were not small, as suggested by Woodgate et al. [2010]. Figure 11 shows that the summer sea-ice extent in the Chukchi Sea was largest in It can be assumed that more of the oceanic heat of ACW was consumed during sea-ice melting on the way from the Bering Strait to Barrow Canyon. Additionally, the presence of sea ice would prevent the absorption of incoming solar Figure 11. Time series of the mean summer (June October) conditions in Barrow Canyon. The blue curve is volume transport (Sv), the green curve is heat flux (TW), and the red curve is heat content in Barrow Canyon (MJ). The magenta curve is the averaged sea-ice area (km 2 ) in the eastern Chukchi Sea (155 W 170 W, 66 N 72 N) from June to September. 4375

14 radiation, and thus, less heating of the surface water would occur in the Chukchi Sea. These observations suggest that year-to-year variations in sea-ice extent in the Chukchi Sea as well as variations in heat advection through the Bering Strait into the Chukchi Sea are critical to the process of heat transport through Barrow Canyon and into the interior Arctic basins. 6. Relationship Between Barrow Canyon Transport and Atmospheric Circulation [36] The mean flow through Barrow Canyon is primarily forced by the sea-surface pressure gradient between the Pacific and Arctic oceans, with variations mainly caused by changes in local wind [Weingartner et al., 2005; Woodgate et al., 2005b; Kawaguchi et al., 2011]. Here we examine the relationship between wind and the volume transport through Barrow Canyon. Six hourly wind measurements at a height of 10 m and atmospheric sea-level pressure were obtained from NCEP/NCAR reanalysis data. Runningaveraged data over 30 days of both Barrow Canyon transport and the wind measurements were used in the analysis. We computed the coherence between the wind and volume transport for different wind-component projections and found that the maximum coherence across the array occurs for along the coast winds (directed 10 from true north, T). A comparison of Barrow Canyon transport to the wind speed along the coast is shown in Figure 12a. The data are significantly correlated (r ¼ 0.69), which is consistent with the results of Weingartner et al. [2005] and Woodgate et al. [2005b]. This high correlation coefficient indicates that it is possible to estimate the Barrow Canyon flux from the wind speed along the coast. [37] Figure 12a also shows the linear regression line from a least-squares fit between Barrow Canyon transport and the wind speed along the coast. Barrow Canyon volume transport (V) was approximated by a linear function of the wind speed along the coast (w) as follows: V ¼ P þ m w ð1þ Figure 12. (a) Scatterplots of the monthly averaged volume transport (Sv) and the NCEP 10 m wind component (m/s) along the coast (10 T). The solid line is the leastsquares regression line, and the dashed lines show the 95% confidence intervals. The correlation coefficient is indicated in the upper left. (b) Seasonal variations of the correlation coefficient (black), slope m (red), and intercept P (blue) for the linear fit between the volume transport and NCEP 10 m wind component along the coast (10 T). The error bars of the slope and intercept P are the 95% confidence limits. with an empirical constant (P) and slope (m), which is related to the coupling of the flux and wind. Note that P represents the transport without wind forcing, i.e., driven by large-scale pressure gradients. Although there was a high correlation between the flux and wind, we found most of the volume flux data during summer (winter) had a positive (negative) bias with respect to the linear regression line as indicated in Figure 12a, suggesting seasonality in this relationship. [38] The seasonal variation in the relationship between Barrow Canyon transport and the wind speed along the coast is shown in Figure 12b. Non-wind-related transport through Barrow Canyon (P) and slope (m) were calculated for each month. The results show that non-wind-related transport (P) reaches its maximum in summer and then weakens in winter. Non-wind-related transport is typically larger than the actual mean flow estimated from mooring data, indicating that synoptic wind generally slows the northeastward flow through Barrow Canyon in all seasons (Figures 10a and 12b). Strong northeasterly winds, which oppose the pressure-head-driven flow, generally dominate in winter. Therefore, both the stronger northeasterly wind and weaker large-scale pressure gradients result in the low transport that was observed in winter. Conversely, higher sea-surface pressure gradients and a weaker northeasterly wind result in the larger transport that was observed in summer. The slope (m) varies between 0.12 and 0.24 during the annual cycle shown in Figure 12b. However, this difference is within the error bar. [39] Figure 13a presents estimates of the volume transport based on the wind correlation together with direct measurements during The agreement between the volume transport calculated from the wind data and that estimated from direct measurement for four complete years of data (2002, 2003, 2006, and 2007) is within 0.01 Sv. The mean value of Barrow Canyon volume transport estimated by using NCEP wind data was 0.49 Sv during Figure 13b shows the freshwater flux estimated based on the correlation between the transport and freshwater flux (as discussed in section 5.3). The agreement of the 4376

15 Figure 13. (a) A time series of 30 day smoothed volume transport (Sv) for , extended through a correlation of NCEP wind data and volume transport in Figure 12b, as indicated by gray lines, with annual averages (circles) and their 95% confidence limits (error bars). Direct measurements from the mooring array at Stn. BCC for and Stns. BCE, BCC, and BCW for are indicated by red and blue lines. (b) Same as (a) except for the freshwater flux (msv). A time series for was extended through a correlation of volume transport and freshwater flux. (c) NCEP 10 m wind speed along the coast (10 T) (m/s). Positive (negative) values indicate southerly (northerly) wind. The blue lines are data averaged from January to March. The green lines are data averaged from April to June. The yellow lines are data averaged from July to September. The red lines are data averaged from October to December. Open (solid) circles indicate 5 years of high (low) wind for each season. freshwater flux calculated from the wind data and that estimated from the four complete years of direct measurements is better than 1.5 msv. The mean value of Barrow Canyon freshwater flux estimated using NCEP wind data was 32 msv during [40] To clarify the relationship between the Barrow Canyon flux and wind, we next examine the SLP pattern over the Arctic Ocean and North Pacific. Atmospheric circulation over the Chukchi Sea is dominated by the interplay between the Aleutian Low and Beaufort High, which display large seasonal variations in position and intensity, as shown in Figures 14a 14d. During fall and winter, the Aleutian Low is large and at its most intense, resulting in stronger northerly winds around Barrow Canyon. During spring and summer, the Aleutian Low is located farther north, and on average is smaller and weaker, or even absent. The Beaufort High is also weaker in summer. In contrast, the North Pacific High is located farther north and is larger over the North Pacific. Correspondingly, the northerly winds around Barrow Canyon are weaker in spring and summer. [41] The interannual variability in the average wind speed near Barrow Canyon for winter (January-March), spring (April-June), summer (July-September), and fall (October-December) is shown in Figure 13c. For each season, we selected 5 years of high and low southerly wind speeds, resulting in high and low Barrow Canyon outflow, and then conducted a composite analysis in terms of SLP and wind patterns (Figures 14e 14h). A decrease in pressure over the Arctic and an increase in pressure over the North Pacific and Alaskan continent were found for all seasons. Such atmospheric patterns induce the weaker SLP contrast between the Arctic and North Pacific, which would damp the northerly wind and consequently result in higher Barrow Canyon outflow from the Chukchi Sea into the Canada Basin. During fall and winter, these patterns are 4377

16 Figure 14. Mean sea level pressure (SLP) and 10 m height wind data for (a) January March, (b) April June, (c) July August, and (d) October December from 1985 to Differences in SLP and wind velocity between high and low Barrow Canyon transport for (e) January March, (f) April June, (g) July August, and (h) October December. more intense because of large interannual variation in the intensity of the Aleutian Low, which is consistent with the result of Kawaguchi et al. [2011]. [42] Figure 13c implies that the Barrow Canyon flux was small in summer during owing to the relatively strong northerly wind. Annual averaged heat flux through Barrow Canyon was maintained only in summer (Figures 9c and 10c), as described in section 5. In addition, variations in the heat flux through Barrow Canyon are driven by changes both in volume transport and in heat content (Figure 11). Therefore, the strong northerly winds in recent summers may have restrained the heat transport through Barrow Canyon into the Canada Basin, although the summer heat content in the Chukchi Sea, including Barrow Canyon, was high in [e.g., Steele et al., 2008; Proshutinsky et al., 2010, 2011]. 7. Summary [43] We established moorings to measure temperature, salinity, and water velocity from September 2000 to August 2008 in the mouth of Barrow Canyon, which is a major conduit for Pacific Water to spread into the interior Arctic basins (Figures 1 and 2). Synoptic high-resolution ship-based surveys along the mooring array were also conducted in September 2002, September 2010, and July and October 2011 (Figures 3 and 4). A northeastward downcanyon flow was evident in the central (Stn. BCC) and eastern shelf (Stn. BCE) regions of the canyon from spring to early fall (Figures 3 and 6). Northeastward flow in the intermediate layer at m was generally stronger than that in the near surface from April to June owing to the flow of PWW (Figure 7). By contrast, from July to September, surface intensification of the northeastwardflowing Alaskan Coastal Current, which carries warm and fresh ACW, was evident. The surface intensification of this current arises largely from a baroclinic geostrophic current, formed by a sharp density slope caused by the presence of warm and fresh ACW on the shelf (Figure 4). [44] Assemblies of annual mooring data and highresolution data collected by ship in summer and early fall along the mooring array enabled quantitative estimations of volume, freshwater, and heat fluxes in Barrow Canyon (Figures 9 and 10). The annual mean volume transport through Barrow Canyon was 0.45 Sv, which consisted of 0.44 Sv of Pacific Water and 0.01 Sv of Atlantic Water. Annual mean Pacific Water transport through Barrow Canyon represents 55% of the long-term mean Pacific Water inflow through the Bering Strait. Volume transport is typically large in summer and small in winter. More of the Pacific inflow is advected by the eastern path as the Alaskan Coastal Current along the Alaskan coast to Barrow Canyon during summer. The freshwater flux through Barrow Canyon was 904 km 3 /yr, which is equivalent to 5% of the freshwater content within the Canada Basin. Seasonal cycles in freshwater fluxes were predominantly caused by variations in volume transport rather than in salinity. The annual averaged heat flux displayed substantial interannual variability and ranged from 0.93 to 3.02 TW, which is enough to melt 1 m thick ice over an area of 88, ,000 km 2. [45] A relationship exists between the measured Barrow Canyon transport and the local winds along the canyon such that under southerly winds, the northward flow of water through the canyon increases (Figure 12). Non-windrelated transport through Barrow Canyon, which is caused by the sea-surface pressure gradient alone, displayed substantial seasonal variations, with a maximum (1.0 Sv) in summer and a minimum (0.4 Sv) in winter. Therefore, both a higher sea-surface pressure gradient and weaker northerly winds generated the larger transport observed in summer. By the correlation of wind and volume transport, we estimated a time series of volume and freshwater fluxes for (Figures 13a and 13b). Composite analysis indicated that a decreasing SLP over the Arctic and increasing SLP over the North Pacific would induce a weaker SLP contrast between the Arctic and North Pacific, which would damp the northerly winds and consequently result in 4378