3-D water properties and geostrophic circulation on the eastern Bering Sea shelf

Deep-Sea Research II ∎ (∎∎∎∎) ∎∎∎–∎∎∎

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3-D water properties and geostrophic circulation on the eastern Bering Sea shelf Edward D. Cokelet n Pacific Marine Environmental Laboratory, National Oceanic and Atmospheric Administration, 7600 Sand Point Way NE, Seattle, WA 98115, United States

art ic l e i nf o

Keywords: Alaska Bering Sea Temperature Salinity Geostrophy Circulation

a b s t r a c t The National Oceanic and Atmospheric Administration's (NOAA) Alaska Fisheries Science Center bottom trawl survey samples demersal fish at over 350 sites on the eastern Bering Sea continental shelf each summer on a 37
37 km2 grid. Rugged conductivity–temperature-depth sensors (CTDs) were added to the net hauls to obtain gridded data sets of temperature and salinity measurements for 2008–2010. Results reveal the three-dimensional thermohaline structure of the shelf including the Cold Pool and areas of fresher water around St. Matthew Island and in Bristol Bay. Horizontal gradients are often strongest roughly along the 50-m and 100-m isobaths that traditionally separate the inner- and outershelf from the middle-shelf centered along the 70-m isobath. The summer mixed layer depth is less than 30 m over much of the region. It reaches the bottom along the Alaska Peninsula in water depths greater than 70 m, showing that the boundary of the well-mixed, inner shelf is not always at the 50-m isobath. The greatest upper-to-lower layer density difference is found across the shelf north of 59°N. The salinity difference is the main contributor to this density difference over most of the region in 2008 and 2010, but the temperature difference dominates in 2009 due to decreased ice melt and reduced freshening near St. Matthew Island. The geostrophic velocity relative to the bottom shows northwestward flow seaward of the 100-m isobath and northwestward transports integrated across the shelf of 0.10–0.25
106 m3/s. In 2008 and 2010 there was clockwise circulation in a region of less-saline water around St. Matthew Island. In 2009 that fresher lens did not exist, and flow was more concentrated along the 100-m isobath bringing saltier water across the shelf. Published by Elsevier Ltd.

1. Introduction The eastern Bering Sea continental shelf is a large, productive marine ecosystem supporting phytoplankton, zooplankton, fin and shellfish, sea birds, marine mammals and people. It sustains coastal communities and a large fishing fleet that provides more than half of the United States’ wild-caught seafood (Wiese et al., 2012; Wiseman et al., 2009). NOAA's Alaska Fisheries Science Center has conducted standardized bottom trawl surveys each summer on the eastern Bering Sea continental shelf since 1982, trawling at over 350 sites to determine the composition, distribution and abundance of demersal fish, shellfish and epibenthic invertebrates (Lauth, 2011; Stauffer, 2004). Besides providing an important long-term view of this resource needed for fisheries management, the survey has made concomitant temperature measurements to help understand the progression between cold and warm years and the size of the Bering Sea Cold Pool

n

Fax: þ 1 206 526 6485. E-mail address: [email protected]

defined by water in which the summer temperature is less than 2 °C (Wyllie-Echeverria and Wooster, 1998). The repeated, regular nature of the bottom trawl survey presented an opportunity to add other oceanographic measurements at low cost with no increase in ship time to further our understanding of the physical processes that affect this ecosystem. Such measurements can provide the basis for an enhanced, long-term program to study physical changes in the Bering Sea. Funded by the North Pacific Research Board's (NPRB) Bering Sea Integrated Ecosystem Research Program (BSIERP), we added conductivity– temperature–depth (CTD) instruments to the net hauls in 2008– 2010. The purpose of this paper is to present the temperature and salinity measurements, the mass density field, their gradients, the mixed layer depth and the bottom-referenced geostrophic circulation on a three-dimensional grid covering a large portion of the eastern Bering Sea continental shelf as a background for the biological studies. Such gridded measurements are rare in physical oceanography and provide a detailed look at the water properties and circulation in the region in a manner usually only obtained by numerical circulation models. They provide a plan view of the region's oceanography at a resolution not possible with a series of

http://dx.doi.org/10.1016/j.dsr2.2016.08.009 0967-0645/Published by Elsevier Ltd.

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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widely spaced oceanographic sections, such as those from the extensive Bering Sea Study (e.g., Mordy et al., 2012). In Section 2 of this paper we describe the CTDs that we employed on the trawl nets, the survey sampling strategy and independent field checks on CTD accuracy. Section 3 begins with maps of each annual survey's progress on the shelf. Temperature (Section 3.1) and salinity (Section 3.2) measurements define the thermohaline field and its gradients (Section 3.3). Section 3.4 deals with the mixed layer depth and density stratification. The overall heat, salt and mass content and their corresponding volumeaveraged temperature, salinity and density are studied in Section 3.5. We compute the geostrophic velocity fields and associated transport (Section 3.6) and compare them to current-meter measurements (Section 3.7). The paper ends with the conclusions in Section 4. Appendix A describes in detail the CTD preparation and data processing.

2. Methods The CTDs needed to be small and light so as to reduce their effect on how the bottom trawl nets fished, be rugged enough to survive hundreds of trips up and down the trawl ramps on the typical 38- to 45-m-long commercial fishing trawlers employed as contract vessels, have long battery life to minimize the chances of waterproof-seal failures when the instruments were opened at sea for battery changes, and sample rapidly to resolve sharp pycnoclines. Falmouth Scientific Instruments (FSI) NXIC CTDs (later Teledyne RD Instruments (TRDI) Citadel CTD-NVs (Teledyne RD Instruments, 2013)) met those requirements.1 Each CTD was 67 cm long, 8 cm in diameter and weighed 3.8 kg in air with a titanium case. It had a rugged thermistor and non-external inductive conductivity sensor (NXIC). The batteries lasted most of the summer field season, requiring only one battery change per instrument. The instruments were deployed at their maximum sampling rate of 15 Hz. In a standard net haul, the ship steamed at 1.5 m/s while the net was lowered to the bottom at a vertical speed of  0.3 m/s, towed for 30 min, and brought up to the surface at  0.3 m/s (Stauffer, 2004). The CTDs were attached to the headrope 2.5 m above the bottom, and measurements made while the net was on bottom will be referred to as bottom measurements. The CTDs were factory calibrated at the beginning of each field season. See Appendix A for details on CTD accuracy, ruggedization and data processing. Each year two ships (F/V Aldebaran and either F/V Arcturus, Alaska Knight or Vesteraalen) conducted the bottom trawl survey, sailing parallel north-south tracks in tandem, 37 km apart with trawls every 37 km, repeating the same basic pattern each year (Fig. 1). During some years, more finely spaced stations were added around islands, and in 2010 the survey extended north to cover the entire U.S. Bering Sea shelf (Lauth, 2011). For each survey year, CTD casts (321 in 2008, 292 in 2009 and 377 in 2010) were gridded onto an equispaced (1/3° latitude
2/3° longitude) grid between 54.6667°N and 66.0000°N and 178.6667°W and 157.6667°W using a Laplacian algorithm with a spline-smoothing parameter of 2 and interpolating no more than 1 grid point away from any data point (Denbo, 1993; Hankin et al., 1992, 1991). Some CTD casts were lost owing to equipment failure, dead batteries or corrupt data. Unless otherwise mentioned, the results that follow are based upon the gridded data set. The survey sampled in June–August 2008–2010 from the 30-m isobath to the shelf break (  180 m) occupying a volume of over 30,000 km3. The only study with somewhat similar coverage is that of Danielson et al. (2011) who used CTD casts from the U.S. 1

Reference to trade names does not imply endorsement by NOAA.

Bering-Aleutian Salmon International Survey (BASIS) in midAugust to early October 2002–2007. BASIS sampled a smaller, shallower area with coarser spacing from the Alaska Peninsula to 64°N at  56-km spacing and from the 20-m isobath to 172°W at  50-km spacing. The BASIS sampling grid was not covered fully each year, and the number of CTD casts varied between 128 and 166. Danielson et al. (2011) gridded the BASIS measurements between the 20 and 70-m depth contours and from 57°N to 62.5°N with a nominal latitude
longitude x depth spacing of 0.5°
0.8°
1 m, occupying a volume of 9300 km3. In addition to the CTDs, F/V Aldebaran was equipped with an underway seawater sampling system consisting of a Sea-Bird Electronics SBE 38 digital oceanographic thermometer and an SBE 45 MicroTSG thermosalinograph to measure salinity. Water was sampled every 60 s from the ship's sea chest, drawing water from approximately 3 m. Discrete salinity samples were taken daily, analyzed in the laboratory after the field season and used to calibrate the thermosalinograph. These temperature and salinity measurements (Cokelet, 2012a, 2012b, 2012c) provided an independent check on the CTD measurements. Fig. 2 shows time series of the underway system and the CTD measurements for the 2009 bottom trawl survey. The CTD measurements track the nearsurface variations very well with correlation coefficients, r2, for temperature and salinity over the three years exceeding 0.99 and 0.98, respectively. The mean differences between the ship's sea chest and CTD measurements for 2008, 2009 and 2010 were 0.22, 0.13 and 0.15 °C for temperature and  0.058, 0.056 and 0.068 (PSS-78, Fofonoff, 1985) for salinity. The CTDs were not recalibrated against the underway measurements owing to the uncertainties of comparing values from the ship's sea chest located in the engine room to CTDs towed in the ship's turbulent wake. The CTD cast data and supporting metadata with plots are available on the BSIERP Data Archive (Cokelet, 2013a, 2013b, 2013c). The oceanographic mooring current measurements referred to in Section 3.7 were made with bottom-mounted, 300 or 600 kHz Teledyne RD Instruments acoustic Doppler current profilers with a 4-m bin size (Stabeno et al., 2012b) . Near-bottom currents were measured with single-point, Aanderaa RCM-9 acoustic Doppler current meters.

3. Results and discussion The NOAA bottom trawl survey required about two months to cover the eastern Bering Sea continental shelf. It can best be characterized as a scan of the region, not a snapshot. It is important to know when the measurements were made for comparisons to numerical model results and to evaluate the synopticity between adjacent samples used to compute gradients. Fig. 1 shows maps of the survey dates for each of the years 2008–2010. The survey began in early June in the southeast corner of the region and moved toward the northwest. While the sampling was not synoptic over the entire region, adjacent sites were typically sampled no more than 2–3 days apart. 3.1. Temperature The area covered by the bottom trawl survey provides a large-scale picture of ocean properties. Fig. 3 (right column) shows the gridded bottom temperature that is defined to be the deepest temperature within 10 m of the bottom or within 10% of the bottom depth for water depths greater than 100 m. The most striking feature is the tongue of cold water with temperatures below 2 °C, the so-called Cold Pool (Wyllie-Echeverria and Wooster, 1998), centered along the muchstudied 70-m isobath section (Stabeno et al., 2012a, 2010) and created by sea-ice formation and melting during the previous winter.

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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Fig. 1. Maps of each year's bottom trawl dates. The black dots represent the CTD cast sites. The long-term EcoFOCI mooring sites (M2, M4, M5 and M8) and 70-m isobath section are shown in red. Depths are contoured at 30, 50, 100, 200, 500 and 2000 m.

The bottom temperature is determined by an interplay between winter cooling, mixing due to wind and tide, and ice cover and persistence (Sullivan et al., 2014). The water is near freezing (o  1.7 °C) beneath areas of persistent winter ice extent. The three study years are considered to be cold years on the Bering Sea shelf (Stabeno et al., 2012b). The Cold Pool is bounded to seaward by warmer, deeper, outer-shelf water in which winter cooling cannot penetrate to the bottom owing to less sea ice and related cooling, more density stratification, and greater depth. To shoreward, the Cold Pool is bounded by the surface mixed layer that reaches the bottom on the inner shelf and is maintained by tidal and wind mixing. We have few actual measurements at bottom depths shallower than 30 m in 2008 and 2009, but those that exist show warmer water near the coast due to summer heating of fresher water derived from coastal rivers that is confined along the coast by shallow fronts (Danielson et al., 2011). In 2010 the measurements extended into Norton Sound where the warmest water was found in shallow depths late in the season. Fig. 3 shows the temperature at 5 m, representative of the surface mixed layer. It varies between 0.54 and 15.20 °C. From

somewhat north of Nunivak Island southward in water depths less than about 50 m, the near-surface temperature reflects the bottom temperature because the mixed layer depth reaches the bottom (see Section 3.4). At greater bottom depths the surface waters are warmer owing to summer warming and density stratification. The warming from southeast to northwest tracks somewhat with the survey age (Fig. 1) and is influenced by summer solar heating during the cruises’ progress. Danielson et al. (2011) calculated that a 800 km
250 km
45 m volume of the inner and middle shelf gained an average of 3.0
1020 J of heat through surface heat flux in May–August of 2002–2007. That works out to an average flux of 140 W/m2. If that surface heat flux were absorbed by a surface mixed layer of depth H, then that layer's rate of temperature T T increase would be Δ Δt ¼ 3:0=H °C /d with the depth in m, the specific heat of sea water of 4020 J/(kg °C) (Fofonoff and Millard, 1983) and a typical mass density of 1025 kg/m3. The measured near-surface temperature in 2010 south and southwest of St. Matthew Island was about 2 °C warmer than in 2008 and 2009 (Fig. 3), but the measurements were made about 10 days later in 2010

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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Fig. 2. Time series of the underway (lines) and CTD (circles) measurements of (A) temperature and (B) salinity at 3 m on F/V Aldebaran, summer 2009.

(Fig. 1). With a typical mixed layer depth of 15 m (see Section 3.4), the temperature increase formula gives 2 °C, in approximate agreement with the observations. 3.2. Salinity The unique aspect of adding CTDs to the bottom trawl survey is the measurement of salinity over the region. Fig. 4 shows the salinity at 5 m and near the bottom. Salinity is a more conservative quantity than temperature in summer, being affected slowly by circulation and river inflow near the coast. In the most general sense, salinity increases seaward. The lowest near-surface salinity (Fig. 4, left column) was found in shallow water near shore, diluted by coastal runoff. Low salinity water (S o31.5) occurred in Bristol Bay, southeast of Cape Newenham, all three years due to runoff from the Kvichak and other rivers. The lowest salinity (So27.5) was measured in shallow Norton Sound in 2010 owing to summer Yukon River flow. The highest upper layer salinity (32.5 oS o33.0 in 2010) was found at the shelf break at the edge of the saltier Bering Sea basin whose surface waters are in that range (32.5 oS r32.9) in summer (Sayles et al., 1979). The greatest bottom salinity (33.0 o S) was also found at the shelf break (Fig. 4, right column). Salinity varied substantially along isobaths over much of the shelf, which is in contrast to the conclusions of Danielson et al. (2011) who studied the warmer years of 2002–2007. South of  61°N and in water depths less than  50 m, the surface mixed layer reached the bottom

(Section 3.4) due to mixing; therefore the surface salinity matched the bottom salinity. This led to a tongue of saltier water (S431.5, Fig. 4) extending northward in  30–50-m water depths in 2009 and 2010, but not in 2008 when the water column salinity was less (30.5oSo31.5). The northward tongue is consistent with the northward movement of inshore drifters in 2008 and 2009 (Danielson et al., 2011). Seaward of this region the water was stratified, and the surface and bottom salinities were more decoupled. A mid-shelf salinity minimum (So31.5) existed southward of 60°N between basininfluenced saltier water on the outer shelf and vertically mixed saltier water on the inner shelf (Fig. 4 and Section 3.4). This is caused by a progression of winter cooling, cycles of sea-ice advance and retreat, and strong winds that mix cold and less-saline water to the bottom there (Sullivan et al., 2014). Following this, spring ice melt freshens the surface layer and stratifies the water column, making the salinity minimum zone larger at the surface than at the bottom (Fig. 4). In latespring and summer, solar heating during a time of weaker winds warms the surface layer (Sullivan et al., 2014). Freshening from summer river runoff would have a different signature with freshest water near the coast and confined to the surface at mid-shelf where tidal mixing could not penetrate to the bottom. Year-to-year differences in the northern mid-shelf surface salinity can be accounted for by the distribution and persistence of sea ice. For example, the near-surface salinity southwest of St. Matthew Island was higher in 2009 (31.5–33.0) than in 2008 and 2010 (30.0–32.0, Fig. 4). Sea ice covers the middle shelf north of 58°N an average of five months each year, and south of 57°N less than one month (Stabeno et al., 2012a). Fig. 5 shows a time series of sea-ice concentration from daily satellite microwave radiometer measurements (Comiso, 2000, updated 2015) along a 1070-km transect passing through St. Matthew and St. Lawrence Islands to Cape York (65.40°N, 167.53°W) on the Alaskan Coast. Also shown are the weekly averaged wind velocities at St. Matthew Island derived from the six-hourly ERA-Interim global atmospheric reanalysis wind field (Dee et al., 2011). In general, northerly winds cause the ice to advance seaward until it melts, cooling warmer water on the middle-to-outer shelf (Pease, 1980). During a 23-day period in February–March 2009, southerly winds forced the ice edge to retreat nearly 200 km to St. Matthew Island. The measurements show that over the three-year record in Fig. 5, the iceedge boundary (one-percent ice concentration line) was correlated (r ¼0.25, p o0.005) with the daily along-axis wind at St. Matthew Island but lagged it by 11 days. To determine the quantitative effect of melting sea ice, consider a mass of ice, M i , with salinity Si (measured in mass of salt per mass of melt water) and mass density ρi melting in a mass of seawater, M w , with initial salinity Sw and density ρw to form a resulting melted mixture of salinity Sm and density ρm . The conservation of total and salt mass can be expressed as Mi þ Mw ¼ Mm M i Si þ M w Sw ¼ M m Sm :

ð1Þ

Combining and rearranging these two equations gives an expression for the difference between the initial salinity of the seawater and the resulting mixture

ΔS ¼ S w  S m ¼

Mi ðS w  S i Þ Mi þ Mw

ð2Þ

(Moore and Wallace, 1988). Now consider this balance on an areal basis for a field of sea ice floating in water. The sea-ice mass per unit area can be expressed as P i H i ρi where P i is the areal concentration of sea ice, varying between 0 and 1, and H i is the ice thickness. Similarly, the mass of seawater per unit area in a layer of thickness H w , equal to or greater than the depth to which the ice melt has an effect, can be expressed as H w ρw P i H i ρi because P i H i ρi is the mass of seawater displaced by the free-floating sea ice.

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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Fig. 3. Maps of the ocean temperature at 5 m depth (left column) and the sea bottom (right column) for the 2008, 2009 and 2010 bottom trawl surveys. The black dots represent the CTD cast sites. The long-term EcoFOCI mooring sites (M2, M4, M5 and M8) and 70-m isobath section are shown in red. Depths are contoured at 30, 50, 100, 200, 500 and 2000 m.

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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Fig. 4. Maps of the salinity at 5 m depth (left column) and the sea bottom (right column) for the 2008, 2009 and 2010 bottom trawl surveys. The black dots represent the CTD cast sites. The long-term EcoFOCI mooring sites (M2, M4, M5 and M8) and 70-m isobath section are shown in red. Depths are contoured at 30, 50, 100, 200, 500 and 2000 m.

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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Fig. 5. Time series of the sea ice concentration with 25
25 km2 resolution along a transect of heading 25°T passing through St. Matthew and St. Lawrence Islands to Cape York on the Alaskan Coast. Stick plots show the weekly average wind speed and direction (from) at St. Matthew Island. Lower ice concentration polynyas occur downwind of Cape York and St. Lawrence Island. Northerly winds cause the ice to advance, but a 23-day period of southerly winds in February–March 2009 caused the ice to retreat. Table 1 Sea ice spring melt parameters near St. Matthew Island. Year

Date of Pi ice melt onset

2008 31 March 2009 13 April 2010 7 April

Hw (m) ρi (kg/ m3)

ρw (kg/ m3)

Si Sm

16

890

1025

6

31.22 32.60 1.38 0.97

0.82 25 0.89 14

890 890

1025 1024

6 6

32.19 32.60 0.51 0.55 31.24 32.60 1.36 0.93

0.99

Sw

ΔS

Hi (m)

Combined, these expressions give for the salinity decrease resulting from melting sea ice in seawater

ΔS ¼ S w  S m ¼ P i

H i ρi ðSw  Si Þ: H w ρw

ð3Þ

If the salinity decrease has been measured, then solving for the thickness of sea ice that melted gives Hi ¼

H w ρ w ΔS : P i ρi S w  S i

ð4Þ

Consider the region extending southwest from St. Matthew Island (58.50°N–60.33°N, 178.00°W–170.00°W) during the period of ice melt when the ice concentration begins its final decrease from its last substantial maximum to zero in the spring. Based upon the sea-ice concentration measurements (Table 1) spring ice melt began between 31 March and 13 April with initial concentrations ranging between 0.82 in 2009, after the ice had been pushed back and then recovered from the southerly wind event, to 0.99 in 2008. Assume that the freshening is confined to the upper mixed layer whose areal average depth H w , density ρw , and final salinity Sm are provided by the bottom trawl survey measurements. The sea-ice salinity Si and densityρi are available from the average of measurements by Sullivan et al. (2014) during 2006– 2009. Based upon Bering Sea Ecosystem Study measurements along the MN (St.-Matthew-Nunivak-Island) transect at 60°N, just south of St. Matthew Island, the typical mixed-layer salinity in early April was 32.60 (Stabeno et al., 2013a, 2013b, 2013c). These values lead to a mixed-layer salinity decrease of  1.37 and a springtime ice melt of 0.95 m for 2008 and 2010, but a corresponding salinity decrease of only 0.51 and an ice melt of 0.55 m

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for 2009 (Table 1). Therefore, sea-ice melt can decrease mixedlayer salinities in this region, and the higher salinity observed in 2009 can be explained by decreased ice concentration owing to earlier southerly winds. By calculating the spring ice melt at a single point (mooring M5) with P i ¼ 1, Sullivan et al. (2014) found values in a similar range in 2008 and 2010. Ice transport by ocean currents and winds during ice retreat has been neglected, but can also play a role. In 2010 when measurements are available north of 62°N, highsalinity water (S 432.5) was observed at the bottom on three sides of St. Lawrence Island and extended northeastward to Bering Strait (Fig. 4). The warmer (  1 oTo 2 °C) water to the north and east of the island (Fig. 3) is consistent with Anadyr Water that flows eastward from the Gulf of Anadyr and then northward through Bering Strait (Coachman et al., 1975; Danielson et al., 2006; Schumacher et al., 1983; Takenouti and Ohtani, 1974). The nearfreezing water (T o 1°C, Fig. 3) to the southwest of St. Lawrence Island could be Anadyr Water or remnant salty water left over from brine rejection during sea-ice formation the previous winter. Currents are weak on the middle shelf giving a typical residence time of months since winter-spring sea-ice formation (Danielson et al., 2012; Stabeno et al., this issue). The numerical climate model of Cheng et al. (2014) shows a net freshwater flux out of the ocean in that region during ice production which is equivalent to a net increase in salinity due to brine rejection. The salinity of the well-mixed area south of Nunivak Island was different in 2009 and 2010. The water was fresher (S o31.0, Fig. 4) than its surroundings in 2009 and saltier (S 432.5) in 2010. The fresher water in 2009 could be due to spring sea-ice melt or runoff from the Kuskokwim River transported seaward as revealed by drifter measurements (Danielson et al., 2011). The saltier water in 2010 could have originated from winter ice formation and brine rejection in the Nunivak polynya or in shallow Kuskokwim Bay to the east. The differences merit further investigation, but are beyond the scope of this paper. 3.3. Gradients
Fig.
6 shows the magnitudes of the horizontal gradients,
∇H T B
and j∇H SB j, of the bottom temperature, T B , and bottom salinity, SB . These show the entire region, regardless of water depth, and the conditions that bottom-dwelling organisms encounter. The gradients are similar if one examines them at constant depth. Bands of strong temperature gradient (j∇T B j Z0.030°C/km) lie approximately along the 100-m isobath in all three years and along the 50-m isobath in 2009 and 2010. Strong salinity gradient bands

(
∇SB
Z 0:012=km) lie along the 100-m isobath, but are not so clearly demarcated. These bands are shelf-wide affirmation of the middle and inner fronts separating the outer, middle and inner domains of the continental shelf as first observed along isolated cross-shelf sections (Coachman, 1986; Kinder and Schumacher, 1981; Schumacher et al., 1979). In simple terms, the outer domain has mixed upper and lower layers separated by a diffuse density gradient layer, the middle domain has two mixed layers separated by a sharper gradient region, and the inner domain is mixed top to bottom. The mid-shelf, 70-m-isobath CTD section, occupied several times a year by NOAA's Pacific Marine Environmental Laboratory (Stabeno et al., 2012a, 2010), is a region of weakly varying temperature and salinity. 3.4. Mixed layer depth and stratification The upper and lower layers of the water column under summer stratified conditions define separate habitats – the upper with ample light for photosynthesis but stripped of essential nutrients, and the lower layer on the mid- to outer-shelf deficient in light but

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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Fig. 6. Maps of the horizontal temperature gradient magnitude, j∇H T B j (left column), and salinity gradient magnitude, j∇H SB j (right column), at the sea bottom for the 2008, 2009 and 2010 bottom trawl surveys. The black dots represent the CTD cast sites. The long-term EcoFOCI mooring sites (M2, M4, M5 and M8) and 70-m isobath section are shown in red. Depths are contoured at 30, 50, 100, 200, 500 and 2000 m.

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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Fig. 7. Maps of the mixed layer depth (left column), and upper-to-lower-layer density difference, Δσt (right column), for the 2008, 2009 and 2010 bottom trawl surveys. White hatching represents areas where the upper mixed layer reaches the ocean bottom. The black dots represent the CTD cast sites. The long-term EcoFOCI mooring sites (M2, M4, M5 and M8) and 70-m isobath section are shown in red. Depths are contoured at 30, 50, 100, 200, 500 and 2000 m.

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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salinity effects is  α ddzθ=β dS , but this quantity can vary between  1 dz and 1 depending on the signs of the numerator and denominator as the salinity gradient approaches zero, much like the trigonometric function tan(x) as x approaches 790°. This inconvenient metric can be rendered useful by defining its inverse tangent, called the Turner angle (Ruddick, 1983), Tu, by   dθ dS  45 3 ð6Þ T u  tan  1  α =β dz dz

Fig. 8. The Turner angle, Tu, for the 2010 bottom-trawl-survey CTD casts, plotted against the contributions of the upper-to-lower-layer salinity and temperature differences. The red dashed curves are isolines of the Brunt–Vaisala frequency squared, N2, in increments of 3
10  4 s  2. Salinity differences contribute more than temperature differences to vertical density differences for negative Tu, and vice versa for positive Tu.

with ample nutrients. Following Danielson et al. (2011), we define the mixed layer depth to be that depth at which the density anomaly, σt ¼ ρ(S, T, p ¼0)–1000 (where ρ is the fluid mass density in kg/m3 and p is the pressure in dbar (decibar ¼ 104 Pa) above atmospheric), first exceeds its average value over the upper 5 m by 0.1 kg/m3. This usually coincides with the top of the pycnocline. If upon reaching the bottom, the density difference criterion is not met, then the mixed layer depth is set to the bottom depth, and the water column is deemed well mixed. The mixed layer depth (Fig. 7, left column) varied between 3 and 79 m during the three bottom-trawl surveys. It was less than 30 m over most of the region but deepened south of Nunivak Island, paralleling the coast, and near the shelf break at some sites. The mixed layer reached the bottom (Fig. 7, white hatched pattern), often in water depths up to 50 m, the traditional outer boundary of the well-mixed inner shelf (Coachman, 1986; Kachel et al., 2002; Kinder and Schumacher, 1981; Ladd and Stabeno, 2012; Schumacher et al., 1979), but results show that the depth at which the mixed layer reaches the bottom can both extend beyond, and at other sites not reach, the 50-m isobath. It is instructive to compute the density difference between the upper and lower layers and to determine the individual contributions of temperature and salinity. The densities at 5 m and at 30 m below the mixed layer depth, or at the bottom if shallower, are representative of the upper and lower layers. The latter was chosen through experimentation to be well below the pycnocline that can be diffuse on the outer shelf. Fig. 7 (right column) shows the lower-toupper layer density difference. It is much greater north of 58°N and around St. Matthew Island in 2008 and 2010 than in 2009. To determine the individual contributions of potential temperature, θ, and salinity to the potential density, ρ, of an incompressible fluid consider its vertical derivative     dρ S; θ dS dθ ¼ ρ β α ð5Þ dz dz dz where β ¼ ρ1 ∂∂Sρ is the haline contraction coefficient, and α ¼  ρ1 ∂∂ρθ is the thermal expansion coefficient. The ratio of temperature to

that varies between  90° and 90° for stable density gradients and equals 0° when the temperature and salinity gradients contribute equally to the density gradient. The derivatives in Eq. (6) were approximated by finite differences between the upper and lower layers, and alpha and beta were averaged over the layers. The Turner angle has an advantage over the stability-index (Ladd and Stabeno, 2012) and the Brunt–Vaisala-frequency (Danielson et al., 2011) methods to measure the effects of temperature and salinity on density stratification. In both applications, those authors compared their indices in two hypothetical cases, one with assumed uniform temperature in the water column and one with assumed uniform salinity. Tu is a measure of the actual stratification without the assumptions of temperature and salinity uniformity. Fig. 8 shows an example of the Turner angle for each CTD cast from the 2010 bottom trawl survey, plotted on axes representing the salinity and temperature contributions. The Bunt–Vaisala frequency is a function of the axes coordinates as shown by dashed red lines isolines (McDougall et al., 1988). For negative Turner angles, salinity differences outweigh temperature differences, and for positive Turner angles, temperature outweighs salinity. For Turner angles between –45° and 45°, as is usually the case, salinity increases and temperature decreases with depth, both increasing the density difference and water column stability. For Turner angles less than  45°, temperature increases with depth, and for Turner angles greater than 45°, salinity decreases with depth – both reducing water column stability. Beyond  90° and þ90°, the water column becomes statically unstable. Fig. 9 shows maps of the Turner angle for the three bottom trawl surveys. On the whole, salinity differences dominate the density differences between the upper and lower layers (Tu o0) over a little more than half of the area sampled in 2008 (58%) and 2010 (55%); whereas in 2009 temperature differences (Tu 40) dominate over 70% of the area. Some of this variation in areal proportion is affected by the differences in sampling areas between the years, but the main effect comes from the distribution of salinity. The low-surface-salinity region in the vicinity of St. Matthew Island in 2009 is smaller than in 2008 and 2010 due to less sea-ice melt, as discussed in Section 3.2 (Figs. 4 and 5), and the density difference is less (Fig. 7). This leads to smaller vertical salinity gradients in 2009 that reduce the relative salinity contribution to the density gradient allowing the temperature gradient to dominate (Tu 40). On a finer scale there are regional differences in the relative contributions of temperature and salinity to the density gradient. The temperature gradient dominates (Tu 4 0) along the 70-m isobath in summer south of about 59°N (Fig. 9) as found by Ladd and Stabeno (2012) and Stabeno et al. (2012a). North of there, salinity dominates in 2008 due to lower surface salinity caused by ice melt (Fig. 4), temperature dominates in 2010 due to warmer surface temperatures (Fig. 3), and a mix of both occurs in 2009. Our observations reveal a more complex regional interplay between temperature and salinity than has been observed along the 70-m isobath alone. Similar to the 70-m line, the Turner angle is positive over much of the southern middle shelf. It is negative along much of the coast between Nunivak Island and Bristol Bay where coastal river inflows contribute to the stratification. Where the surface mixed layer reaches the bottom, what little density difference that exists can be dominated by either temperature or salinity

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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11

Fig. 9. Maps of the Turner angle, Tu, for the 2008, 2009 and 2010 bottom trawl surveys. Salinity differences contribute more than temperature differences to vertical density differences for negative Tu, and vice versa for positive Tu. Black hatching represents areas where the upper mixed layer reaches the ocean bottom. The black dots represent the CTD cast sites. The long-term EcoFOCI mooring sites (M2, M4, M5 and M8) and 70-m isobath section are shown in red. Depths are contoured at 30, 50, 100, 200, 500 and 2000 m.

differences (Fig. 9, hatched areas). In all three years, salinity dominates the density difference near Unimak Pass (Tu o0) due to inflow from the Gulf of Alaska of the Alaska Coastal Current in which salinity controls density owing to river inflows along the current's path (Aagaard et al., 2006; Royer, 1982; Schumacher and Kinder, 1983; Stabeno et al., this issue, 2002). The middle shelf stratification partitioning carries onto the outer shelf in 2008 and 2009. However in 2010 the outer shelf stratification is salinity dominated in a band of negative Turner angle; whereas the middle shelf is temperature dominated (Tu 40). This is because lower salinity surface water (31.5o So32.0) is farther offshore, seaward of the 100-m isobath north of Pribilof Canyon, in 2010 than in the other years, leading to a larger vertical salinity difference. Measurements made north of 62°N, in 2010 only, show salinity gradient domination (Tu o0) in most regions, especially in shallow Norton Sound where the surface waters are diluted by Yukon River inflow (Fig. 9).

3.5. Heat and salt content Given the regional water-property measurements, one can compute the mass (M), heat (H) and salt (S) contents of a volume of fluid (V with differential element dV ¼ dxdydz) as given by Z   M¼ ρ ! x ; t dV ð7Þ Z H¼

   h   i ! ! ρ ! x ; t C p x ; t T x ; t þ 273:15 dV

ð8Þ

    ! ! S x ;t ρ x ;t S¼ dV ð9Þ 1000 ! where x ¼ hðx;y; zÞ are  the  spatial   coordinates; i  ρ isthe in-situ mass ! ! ! ! density; C p S x ; t ; T x ; t ; p x ; t ¼ C p x ; t is the specific heat at constant pressure p; T is the temperature in degrees Celsius; T þ273:15 is the absolute temperature in Kelvin; and S is the Z

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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Table 2 Layer averages over common regions on the eastern Bering Sea shelf. Region

Year 2008

2009

2010

Volume (km3)

Mixed Layer Below Mixed Layer Cold Pool Total

11,050.1 18,459.9 15,694.3 29,510.0

11,712.0 17,798.0 13,550.6 29,510.0

8987.4 20,522.6 11,311.8 29,510.0

Mass (Gt)

Mixed Layer Below Mixed Layer Cold Pool Total

11,327.3 18,938.8 16,098.5 30,266.2

12,006.2 18,259.8 13,900.5 30,266.0

9212.1 21,052.6 11,602.2 30,264.7

Heat Content (1022 J)

Mixed Layer Below Mixed Layer Cold Pool Total

1.2539 2.0795 1.7636 3.3334

1.3323 2.0053 1.5221 3.3376

1.0236 2.3172 1.2705 3.3408

Salt Content (Gt)

Mixed Layer Below Mixed Layer Cold Pool Total

356.22 608.12 512.75 964.34

379.19 586.24 443.52 965.43

290.20 674.48 368.35 964.68

In-Situ Density (σ, kg/m3) Mixed Layer Below Mixed Layer Cold Pool Bottom Total

25.09 25.94 25.75 25.81 25.62

25.12 25.95 25.82 25.76 25.62

25.00 25.82 25.68 25.88 25.57

Temperature (°C)

Mixed Layer Below Mixed Layer Cold Pool Bottom Total

3.20 1.25 0.46 1.19 1.98

3.97 1.30 0.37 1.41 2.36

4.28 1.90 0.31 1.44 2.63

Mixed Layer Below Mixed Layer Cold Pool Bottom Total

31.45 32.11 31.85 31.84 31.86

31.58 32.11 31.91 31.81 31.90

31.50 32.04 31.75 31.95 31.87

Salinity

salinity (Fofonoff and Millard, 1983; Gill, 1982; Willis et al., 2003). The heat content is computed from the absolute temperature. Some investigators (e.g. Danielson et al., 2011) reference the heat content to 273.15 K (0 °C), but that reference heat   content  varies ! ! with time owing to the time-dependent ρ x ; t C p x ; t terms in Eq. (8). This can have a substantial influence on the variation of the heat-content anomaly about its temporal mean value. Of particular interest is how the volume, mass, heat and salt contents change over time. We consider the grid points common to the 2008–2010 annual surveys and compute the contents of four regions: the total shelf (that part which we measured), the surface mixed layer, below the mixed layer, and the Cold Pool (Table 2). Fig. 10A shows time series of the layer volumes. The mixed-layer volume increased slightly in 2009, but decreased in 2010. The volume below the mixed layer did just the opposite to keep their sum (the total volume) constant. The mass, heat and salt contents are strongly influenced by the layer volumes, such that plots of those quantities (not shown but values given in Table 2) closely resemble Fig. 10A but with appropriate ordinates. In general between the 2008 and 2010 surveys the total eastern Bering Sea shelf warmed by 0.65 °C (Fig. 10B), and its density decreased owing to the warming (Fig. 10D). The salinity increase in 2009, caused by decreased sea-ice melt, counteracted the effects of warming on the density that year. The heat content (referenced to

0 °C) increased by 78
1018 J (or 74
1018 J referenced to 0 K, Table 2) from 2008 through 2010, which reversed a trend of decreasing late-summer heat content of 92
1018 J in 2004–2007 in the upper 100 m of the water column as deduced by Danielson et al. (2011) over a larger area. The depth-averaged summer temperature at mooring M2 decreased from 2008 through 2009 (Stabeno et al., 2012b); where as the summer, shelf-wide surveys gave a temperature increase. Danielson et al. (2011) showed that temperature and salinity trends may not be the same over different parts of the shelf. Based upon long-term M2 measurements, 2001–2005 was a warm period of low ice extent, and 2007–2010 was a cold period of high ice extent. Stabeno et al. (2012b) ascribe warming and cooling on the southern shelf to low and high sea-ice extent, respectively. Danielson et al. (2011) argue that the variability of along-isobath summer heat flow accounts for much of the variability in late-summer heat content. All these effects are important and merit further study, but are beyond the scope of this paper. Between 2008 and 2010, the Cold Pool's volume decreased by 28% (Fig. 10A) and temperature cooled by 0.15 °C, while the other layers warmed (Fig. 10B). A shrinking Cold Pool provides a larger habitat for fish such as walleye pollock (Gadus chalcogrammus) that avoid very cold water (De Robertis and Cokelet, 2012). In some years the average water density at the sea bottom was less than that below the mixed layer and the Cold Pool (Fig. 10D), but that does not mean that the water column was statically unstable. The Cold Pool and the area below the mixed layer did not extend onto the inner shelf because the mixed layer reached to the bottom there (hatched area in Fig. 7). Therefore that low-density, inner-shelf water did not contribute to the Cold-Pool and belowthe-mixed-layer densities, but it did contribute to the average bottom density, decreasing it. 3.6. Geostrophic velocity and circulation Horizontal density gradients are balanced by geostrophic ocean currents under the assumption of steady flow over periods greater than the inertial period of  14 h at 60°N (Gill, 1982; Neumann and Pierson, 1966). Neighboring grid points in the bottom-trawl surveys were typically sampled no more than 2–3 days apart (Fig. 1); therefore velocities need only be assumed time-invariant over a few days for the geostrophic approximation to be valid. However if the geostrophic velocity vectors are interpreted together over the entire survey region, then there is an implicit assumption of steadiness over the survey's duration. Geostrophic velocities were computed relative to the sea bottom (the greatest common pressure between the north-south and east-west grid point pairs). For the bottom to be a level of no motion, isobars tilt upward and isopycnals tilt downward to the right of the flow leaving denser water on the flow's left. Assuming the bottom to be a level of no motion is the most restrictive assumption that makes the computed geostrophic currents only one element of the complete flow field that includes shorter period wind-driven and tidal currents. Fig. 11 shows the geostrophic velocity vectors at 5 m, referred to the bottom, overlaying the density anomaly, σt, field. Seaward of the 100-m isobath, there is a band of relatively strong, organized northwestward flow (speed43 cm/s) north of 58°N all three years. Organized flow along the 100-m isobath has been observed in satellite-tracked drifter trajectories and by geostrophic velocity calculations along the MN hydrographic transect (Stabeno et al., this issue). The 100-m isobath bends away from the shelf break north of the Pribilof Islands, leading to cross-shelf flow. In 2009 the flow is confined nearer the 100-m isobath because saltier, denser water (Fig. 4) has intruded across the shelf. South of the Pribilof Islands the flow turns offshore near Pribilof Canyon as has been observed in current meter measurements (Stabeno et al., this issue). In 2010, the

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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Fig. 10. Time series of the volume, mean temperature, salinity and in-situ density anomaly (σ ¼ρ  1000 kg/m3) for the mixed layer, below the mixed layer, the total shelf, the Cold Pool and the ocean bottom. Volume (solid lines) and area (dashed line) averages are over gridded CTD sites common to the 2008–2010 bottom trawl surveys.

band of stronger flow extends from Pribilof Canyon to the Russia–USA convention line. In 2008 and 2010 there is clockwise circulation around St. Matthew Island owing to the low-density water there caused by low salinity in 2008, due to ice melt, and by high temperature and low salinity in 2010 (Figs. 3, 4 and 8). Farther south along the middle shelf, geostrophic flow is weaker near mooring M2 (speedo1.5 cm/s) and M4 (speedo2.5 cm/s). Flow is weak in the areas measured landward of the 50-m isobath and southeast of Nunivak Island, except for a coastal flow south of Cape Newenham set up by the pressure gradient due to low-salinity, lowdensity water from Kvichak River flow into Bristol Bay. The bottom trawl survey only sampled north of St. Lawrence Island in 2010, but measurements show strong, bottom-referenced geostrophic flow (speed43 cm/s) of salty Anadyr Water from the west flowing along the Russia–USA convention line toward Bering Strait (Coachman et al., 1975; Danielson et al., 2014; Sambrotto et al., 1984). Reed and Stabeno (1996) computed a composite mean geostrophic circulation relative to a pressure of 50 dbar on the Bering Sea shelf south of 58°N and west of 161°W by combining 1299 summer (June–September) CTD casts taken between 1975 and 1989. We repeated the calculation using the 2008–2010 bottom trawl survey measurements (not shown). The dynamic topography was similar in magnitude in all three years with geopotential heights varying between 0.110 and 0.160 dynamic meters consistent with past computations (Reed and Stabeno, 1996). However the circulation patterns differed. Reed and Stabeno found northward flow from Unimak Pass to near the Pribilof Islands that then looped south and east across the shelf to near the site of presentday mooring M2 before turning north again along the 50-m isobath. West of this loop was northward flow near the shelf break. We found the latter flow in all three years, but only in 2010 did we

find flow somewhat similar to the former. In the other two years the flow was slower and more chaotic, especially in 2008. Overall, our results show that the bottom-referenced geostrophic flow is generally northwestward and similar to flow patterns derived from satellite-tracked drifters (Stabeno et al., this issue). To compute the northwestward or poleward volume transport, we define four northeast–southwest sections, S1–S4, denoted by white lines in Fig. 11. Ideally these sections should be synoptic and cross the continental shelf from the 200-m isobath to the shoreline so that the total along-shelf transport could be computed. However the available observations do not span the shelf, and the sites where the geostrophic velocities can be computed vary from survey to survey owing to some missing CTD casts. Therefore we chose sections that intersect geostrophic velocity sites (and their interpolates) common to all years. This allows for year-to-year transport comparisons at equivalent sections but may not provide a good measure of total along-shelf transport comparisons in a given year. White arrows in Fig. 11 represent the net poleward volume transport through each section that varied between 0.08 and 0.23 Sverdrup (Sv ¼106 m3/s). For comparison, the May–September transport entering the Bering Sea through Unimak Pass has been estimated to be 0.25 Sv from combined geostrophic calculations referenced to current meter observations (Stabeno et al., this issue). The flow leaving the Bering Sea through Bering Strait near 66°N has been measured at 0.8 Sv (Roach et al., 1995) with June and July values of 1.3 and 1.1 Sv, respectively (Woodgate et al., 2005), but much of this flow (Stabeno et al., this issue) comes from the Anadyr Current on the western side of Bering Strait that draws water from the Bering Slope Current, seaward of our study area. Woodgate and Aagaard (2005) estimate that the fresher Alaska Coastal Current part of the flow on the eastern side of Bering Strait is 0.08 70.02 Sv.

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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Fig. 11. Maps of the density anomaly, σt, at 5 m with geostrophic velocity vectors referred to the bottom for the 2008, 2009 and 2010 bottom trawl surveys. White arrows represent the net transport across sections S1–S4, and black dots mark the CTD cast sites. The long-term EcoFOCI mooring sites (M2, M4, M5 and M8) and 70-m isobath section are shown in red. Depths are contoured at 30, 50, 100, 200, 500 and 2000 m.

Fig. 12 shows the volume transport time series. The curves with positive values represent the net poleward (polewardþ equatorward) transport, and those with negative values represent the equatorward components, which are much smaller. Sections S1 and S2 come closest to spanning the shelf, being nearly closed at their landward ends owing to shallow depths there, but they have the greatest transport differences and variations. In 2008 and 2009 when S1 exceeded S2 in transport, there was a net seaward outflow between the two sections’ seaward ends to conserve mass. In 2010 the situation reversed with landward inflow between the sections. The net transport through a section is primarily governed by the flow at its deep, seaward end. Eddies in the Bering Slope Current seaward of the shelf break can transport water on or off the shelf (Ladd, 2014; Mizobata et al., 2002, 2008; Okkonen, 2001; Okkonen et al., 2004). Currents in such eddies reach depths of 1000 m or more. The interaction between eddies and the Bering Slope Current sets up complicated flow patterns that are beyond the scope of this paper (Ladd, 2014); however sea surface height anomaly (SSHA) measurements

from satellite altimeters (Envisat, Jason-1 and Jason-2) can provide some insight (Colorado Center for Astrodynamics Research, 2015; Leben et al., 2002). In July 2008 there was a weak (5-cm-maximumSSHA) anticyclonic eddy impinging on the shelf south of the Pribilof Islands that would induce seaward flow in Pribilof Canyon (not shown, Colorado Center for Astrodynamics Research, 2015; Leben et al., 2002). In July 2009 there were two eddies that would induce seaward flow at the shelf break between S1 and S2: a 12-cmminimum-SSHA cyclonic eddy centered southeast of S1 at 168°W and a 10-cm-maximum-SSHA anticyclonic eddy centered off the seaward end of S2 at 174°W (Colorado Center for Astrodynamics Research, 2015). In July 2010 the eddy arrangement favored on-shelf flow with a counter-rotating eddy pair (an 8-cm-minimum-SSHA cyclonic eddy centered at 172°W and a 12-cm-maximum-SSHA anticyclonic eddy centered at 170°W) whose between-eddy induced flow was aimed directly onto the shelf at 171°W (see Ladd, 2014, Fig. 8B). These on-off shelf flows are not represented well in our geostrophic analysis because the CTD casts do not go beyond

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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the shelf break. Also the assumption of vanishing geostrophic flow at the bottom may not be valid where the depth changes abruptly, and a deep eddy impinges upon the shelf. However the gains and losses

Fig. 12. Time series of the volume transport through sections S1–S4 as shown in Fig. 11.

15

across the shelf break are reflected in the along-shelf transports in the cross-shelf sections. Fig. 13 shows cross-sections of the temperature, salinity, density anomaly and velocity, vN, normal to sections S1–S4. Vertical dashed lines represent the positions of the 100- and 50-m depth contours that traditionally mark the middle and inner fronts, respectively. The Cold Pool (outlined in black) is broader at middepth than along the bottom in the majority of sections and extends seaward of section S4 in all years. The coldest zone (T o 1°C) is larger along sections S3 and S4 that are north of 58°N and most influenced by sea ice. Along sections S1 and S2 there is a well-mixed region in the vicinity of, but not always at, the 50-m depth contour, but some vertical stratification can exist landward of there as implied by Fig. 7 and discussed in Section 3.4. The middle front is identified by tilted isopycnals over the full-depth range of each section at approximately the position of the 100-m depth contour, but its indicial depth contour can vary by 20 m or more. Landward of there, the isopycnals are flatter in the upper 20–30 m of the water column, characteristic of the middle shelf domain. The geostrophic velocity field depends on the horizontal pressure gradient that is reflected in the steepness and closeness of the isopycnals. Steep, closely spaced isopycnals lead to stronger flow. In 2008 (Fig. 13A) and 2009 (Fig. 13B) section S1's normal velocities reach their maxima (4 4 cm/s) seaward of the 100-m depth contour with the region of strong northwestward flow

Fig. 13. Sections S1–S4 of water temperature T, density anomaly σt with salinity contours, and normal velocity vN for the bottom trawl surveys of (A) 2008, (B) 2009 and (C) 2010. Thick, black curves highlight the Cold Pool (T r 2 °C) in the T sections. Mixed layer depths are shown by black, dashed curves. Vertical dashed lines represent the positions of the 100- and 50-m depth contours.

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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Fig. 14. The measured ocean current speed and direction with 95% confidence limits (red curves) at PMEL moorings compared to those quantities computed from geostrophy 1 (black curves). The measurements were averaged over the timescale ð f R0 Þ C58 days centered on the time of the nearest bottom trawl survey CTD cast to a mooring. (A) Mooring M4, 27 May–24 July 2008, (B) Mooring M2, 11 May–8 July 2009, (C) Mooring M5, 4 June–1 August 2009, and (D) Mooring M5, 1 July–28 August 2010.

probably extending past the end of the section, but in 2010 (Fig. 13C) the northwestward velocity is weaker and shallower, leading to a decrease in the transport. A comparison of T, S and σt on section S1 shows that the flattened isopycnals in 2010 are due

to a reduced horizontal salinity gradient and an over-50-m deep lens of isothermal water. The enhanced flow through S2 in 2010 reflects the steeper isopycnals that year owing to a larger salinity gradient (compare Fig. 13A, B and C between 50 and 100 km).

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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Maximum flow through section S3 is just seaward of the 100-m depth contour. Flow through section S4 is largest in 2009 owing to a stronger horizontal salinity gradient caused by decreased sea-ice melt and an intrusion of saltier basin water (Fig. 11 and Section 3.2). 3.7. Geostrophic velocities and mooring measurements The bottom-referenced geostrophic velocities can be compared with ocean current measurements. PMEL has maintained a series of moorings – M2, M4, M5 and M8 – along the 70-m isobath for over 10 years (Stabeno et al., this issue, 2012b). We interpolated the gridded geostrophic velocities to the mooring sites at the times corresponding to the nearest CTD casts and compared them to the measured currents whenever possible. It is not appropriate to compare the instantaneous measured current to the geostrophic flow; the comparison must be performed over a dynamically relevant time scale. If an oceanic flow field is subject to new forcing by wind or a changing mass distribution, it will respond quickly on the geostrophic adjustment time scale (on the order of the inertial period, 2π =f where f is the Coriolis parameter) but shed transient disturbances via propagating inertia-gravity waves over a longer time scale given by ð f R0 Þ  1 where R0 ¼ U=ð f LÞ is the Rossby number, U is a typical velocity scale, and L is a typical length scale (Blumen, 1972; Reznik et al., 2001; Zeitlin et al., 2003). This longer time scale is the appropriate one for the velocity comparisons. We take U¼ 2 cm/s as a typical mid-depth velocity scale from geostrophic velocity profiles and L¼100 km as a typical length scale of horizontal density variations to give R0 ¼ 160
10  3 and ð f R0 Þ  1 C 58 days. Different techniques are required to compute the uncertainties in the current-meter time series than in the geostrophic velocity spatial fields. The hourly current-meter time series were low-pass filtered with a least-squares-linear filter with 35-hour-cutoff period to remove autocorrelations due to the tides. The uncertainty or error in the mean low-pass-filtered current meter speed,  1=2 q ¼ u2 þ v2 where the overbar represents an average over the time scale ðf R0 Þ  1 , is characterized by its standard error given by the square root of the ratio of its variance to the number of independent samples in the average, N  ¼ N=ð2τÞ, where N is the number of samples and τ is the integral time scale (Leith, 1973; Taylor, 1921). Computing the uncertainty in the mean current direction, ϕ ¼ tan  1 ðu=vÞ, requires a special approach owing to the 360° periodicity and application of a wrapped-normal distribution. Fisher's (1987, 1993) methods were used to compute the circular dispersion and circular standard error with the effective sample size, N  , accounted for. The geostrophic velocity spatial uncertainty was computed from horizontal gradients and by application of the method of propagation of errors (Ku, 1969; Meyer, 1975). The uncertainty, Δu, in the eastward velocity, u, is given by  2  2 ∂u ∂u ðΔuÞ2 ¼ ðΔxÞ2 þ ðΔyÞ2 ð10Þ ∂x ∂y where Δx and Δy are the east-west and north-south spacings of the interpolated data grid. Similar results hold for Δv. By the propagation of errors, the uncertainty in the current speed, q ¼ ðu2 þ v2 Þ1=2 , is given by ðΔqÞ2 ¼

u2 ðΔuÞ2 þ v2 ðΔvÞ2 ; q2

ð11Þ

and the uncertainty in the current direction is given by ðΔϕÞ2 ¼

v2 ðΔuÞ2 þ u2 ðΔvÞ2 q4

ð12Þ

17

Fig. 14 shows vertical profiles of the current speed and direction on the four occasions in the summers of 2008–2010 when the CTD casts and the current measurements coincided. Error bars represent 95% confidence intervals defined as 71.96 times the standard errors as described above. The appropriately averaged mooring and geostrophic current speeds (left column) varied between 0 and 3 cm/s, and they agree within overlapping 95% confidence limits for moorings M4 in 2008 (Fig. 14A), M2 in 2009 (Fig. 14B) and M5 in 2010 (Fig. 14D). The current directions (Fig. 14, right column) were more variable, but they also agree within overlapping error bars in those three cases. The speed and direction do not agree at M5 in 2009 (Fig. 14C). In fact, measured M5 currents show considerable vertical or baroclinic structure in 2009 and 2010 (Fig. 14C and D) with directional shifts of  120° over the depth range. The geostrophic velocities reflect that in 2010 (Fig. 14D), but not in 2009 (Fig. 14C). In 2009 at M5, the observed velocity did not vanish at the sea bottom; measurements show an appreciable mean velocity of 2 cm/s there (Fig. 14C). Mooring M5 is near to St. Matthew Island (Fig. 11), and it is likely that water piled up against the island leading to a depth-independent barotropic component that was not compensated by the baroclinic component at the bottom. Otherwise the measured and geostrophic currents are in reasonable agreement. In general, mean currents along the 70-m isobath are weak. It would be advantageous to obtain current measurements seaward of the 100-m isobath where the flow is stronger (Fig. 11) and compare those to geostrophy.

4. Conclusions This BSIERP project to augment the Bering Sea bottom trawl survey with ruggedized CTD measurements was very successful. We obtained 990 CTD casts between 2008 and 2010 with no increase in ship-time resources. Independent temperature and salinity measurements from an underway-oceanographic measurement system provided field checks on the CTD's accuracy near the sea surface, a zone of large horizontal variability. The nature of the bottom trawl survey, on its regular three-dimensional grid that is occupied in the same way each year, provides a unique data set that is rarely encountered in oceanography. The grid lends itself to straightforward analysis and interpretation, and will be useful for comparison with numerical models (Hermann et al., this issue), which are often expressed on a grid. Results reveal the threedimensional thermohaline structure of the continental shelf between the 30-m and 200-m isobaths from the Alaska Peninsula to 62°N in 2008–2009 and to 65°N in 2010. There are a few older surveys of the eastern Bering Sea that cover a series of hydrographic sections. One common thread that emerges is the existence of the cold bottom-water tongue – the Cold Pool bounded by the 2°C-isotherm – produced by winter sea ice on the middle shelf. Using USCGT Redwing hydrographic casts from the summers of 1939–40, Dodimead et al. (1963) found the 2 °C-isotherm reached as far south as about the 50-m isobath along the Alaska Peninsula (to 57.8°N, 160°W in June 1939 and 56.0°N, 162°W in July 1940). From the summer 1956-67 cruises of R/V Oshoro Maru, Takenouti and Ohtani (1974) showed that the 2 °C-isotherm's southernmost extent varied between 58.1°N in 1958, following a strong El Niño, and 55.3°N in 1959, with a typical extent around 56.5°N. Reviewing June 1963-73 Oshoro Maru and 1976-82 PROBES observations, Coachman (1986) showed the Cold Pool extended to between 57.7°N and 55.3°N, but could not be detected south of 59°N in 1969. In June 1976 following the previous cold La Niña winter, it reached the Alaska Peninsula at 55.0°N. During our 2008-10 field years the Cold Pool extended to 55.5-55.8°N bounded by the Alaska Peninsula, about as far south

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

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E.D. Cokelet / Deep-Sea Research II ∎ (∎∎∎∎) ∎∎∎–∎∎∎

as any previous observations and consistent with a cold period. All of these authors mentioned the mixed-layer depth and the wellmixed inner shelf, but none had the breadth of measurements necessary to map these features shelf-wide. The new measurements give a fuller view of the salinity and density structure on the shelf than those along the oft-measured 70-m transect alone, or at other widely spaced hydrographic sections. They enable the computation of horizontal gradients that are often strongest roughly along the 50-m and 100-m isobaths that traditionally demarcate the inner and middle fronts. The sharp temperature gradient between the Cold Pool and the surrounding waters can provide a boundary to fish habitat (De Robertis and Cokelet, 2012). The CTD measurements have been used to help establish the effects of ocean habitat on the distribution forage fish across the Bering Sea shelf, in particular the effects of surface-tobottom temperature differences (Hollowed et al., 2012). The geostrophic velocity vectors show that cross-shelf advective pathways exist between the outer (depth 4100 m) and middle shelf (50o depth o100 m). A regional view of the summer mixedlayer depth shows it to be less than 30 m over much of the region. It reaches deeper to the bottom along the coast southeast of Nunivak Island often in water depths up to 50 m, and results show that the surface-to-bottom-mixed region can extend beyond, and at other sites not reach, the 50-m isobath. The greatest upper-tolower layer density difference was found across the shelf north of 59°N, but its horizontal extent was much less in 2009 than in 2008 and 2010. Using the Turner angle, we have shown that the salinity difference is the main contributor to the vertical density difference over most of the shelf especially near Unimak Pass, at the shelfbreak, around St. Matthew Island and near the coast, but the temperature difference dominates in 2009 due to decreased ice melt and reduced freshening near St. Matthew Island. The regional measurements provide time series of the shelf-wide volume, mass, heat content, salt content, temperature, salinity and mass density that give integrated metrics to track physical changes in the eastern Bering Sea. Horizontal density gradients give the geostrophic velocity relative to the bottom, providing the first comprehensive, shelfwide circulation maps of the geostrophic circulation in the eastern Bering Sea. Velocity vectors show northwestward flow seaward of the 100-m isobath. Northwestward transports integrated across the shelf range from 0.10 to 0.25 Sv. In 2008 and 2010 there was clockwise circulation in a region of less-dense water around St. Matthew Island. In 2009 that less-dense lens did not exist, and flow was concentrated along the 100-m isobath bringing saltier water across the shelf. The bottom-referenced geostrophic velocity vectors agreed with current-meter measurements in three out of four comparisons. The 2008–2010 measurements reported upon here were made during relatively cold years on the eastern Bering Sea continental shelf (Stabeno et al., 2012b). The measurement time series should continue. Shelf-wide comparisons between cold and future warm years will be of particular interest.

system on Aldebaran. The members of NOAA's Alaska Fisheries Science Center Bering Sea bottom trawl survey, led by Robert Lauth, graciously attached the CTDs to their nets, downloaded data daily, looked after the equipment at sea and took daily salinity samples. Dennis Holzer and PMEL's Machine/Fabrication Shop built the CTD polypropylene cases, and AFSC's net shed constructed the net bags. Scott McKeever and William Floering analyzed the salinity samples. Tiffany Vance provided the positions of the Alaskan rivers shown on the maps. The response to comments by an anonymous reviewer and discussions with Phyllis Stabeno strengthened the paper. The Ferret program for analysis and graphics (http://ferret.pmel.noaa.gov/Ferret/) was used in this paper (Hankin et al., 1992, 1991). This is BEST-BSIERP Bering Sea Project publication No. 178, NPRB publication No. 580, contribution No. 4130 from the Pacific Marine Environmental Laboratory and contribution EcoFOCI-N817 from NOAA's North Pacific Climate Regimes and Ecosystem Productivity research program.

Appendix A Although the Citadel CTDs were designed to be rugged, they needed extra protection during net deployment and recovery across the fishing vessel trawl ramps. We constructed protective cases from 160-mm-diameter Asahi/America Proline Pro-150 polypropylene pipe with a 14.6 mm wall thickness (Fig. A1). Each case was placed in a nylon mesh bag and attached to the trawl net's headrope (Fig. A2) with the conductance cell facing into the flow as the net fished. This configuration represented a compromise between CTD protection and thermal and salinity lags owing to added thermal mass and reduced flow through the conductivity cell. The manufacturers specified temperature, conductivity and pressure accuracies of 70.005 °C, 70.009 mS/cm and 0.05% of full scale, respectively. These convert to salinity and pressure accuracies of 70.011 (PSS 78, Fofonoff (1985)) at 5 °C and 70.25 dbar for the 500m pressure transducers that the CTDs carried. We estimate that the uncertainty in a dynamic field environment is probably twice that or about 70.01 °C for temperature and 70.02 for salinity. As the CTD is lowered or towed, water flows through the inductive cell and past the thermistor; therefore no pump is required, thus enhancing battery life. However the residence time of a water parcel within the

Acknowledgments This research was funded in part by the North Pacific Research Board under the Bering Sea Integrated Ecosystem Research Program (Project B62) which together with the US National Science Foundation's Bering Ecosystem Study (BEST) formed the Bering Sea Project. I thank the captains and crews of F/Vs Aldebaran, Arcturus, Alaska Knight and Vesteraalen and the fishing vessel owner, Trident Seafoods, Seattle. Antonio Jenkins and David Strausz designed and installed the underway seawater sampling

Fig. A1. NXIC CTD with protective polypropylene case components and net bag.

Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i

E.D. Cokelet / Deep-Sea Research II ∎ (∎∎∎∎) ∎∎∎–∎∎∎

Fig. A2. Photograph of an NXIC CTD mounted inside a 160-mm diameter polypropylene pipe and attached to the trawl net's headrope.

inductive cell and the time lag between the conductance and temperature measurements vary with flow speed. To determine varying salinity accurately, the changing temperature at the instant and location of the conductance measurement must be accounted for because temperature effects dominate conductance measurements, and salinity plays a secondary role. The data processing scheme comprised several steps. First, pressure, temperature, conductivity, salinity and battery voltage data were downloaded daily at sea with software (CTDPro) provided by the manufacturer. Second, outlying values were eliminated during laboratory processing as follows. The CTDs were set up to begin recording when they entered salt water. Data values with conductivities below 5 mS/cm, owing to initial startup transients, were removed at this processing stage. During initial net deployment and final retrieval, the CTDs remained on the surface for a few minutes behind the moving ship. These measurements in the ship's wake did not represent the undisturbed water column; therefore we removed values for pressures less than 1 dbar. Owing to data handling errors within the CTDs’ factory software, especially in the early years of this study, some files had repeated data blocks making it appear that the sampling time did not progress or had jumped backwards. These sequences were detected and removed. The remaining values were linearly interpolated onto an equispaced time axis with an increment of 1/15th second – the instrument's sampling interval. The third step in our processing scheme involved correcting the measured pressures, temperatures and conductivities for sensor lag times. The sensor response times were as follows: pressure 0.025 s, temperature 0.100 s and conductivity 0.05 s at a flow rate of 1 m/s (Teledyne RD Instruments, 2009). We used the method of Fofonoff et al. (1974) to correct for these lags. In general, the true value of a variable, f T ðtÞ, is approximated by its measured value, f M , by

df M ð13Þ f T ¼ f M þ τf dt where τf is the response time for variable f ; and the bracket

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represents a smoothing operation applied to reduce noise in the time derivative. The CTD's pressure measurements were noisier than the other variables; therefore we smoothed them first with a 15-point (1s duration) Hanning window (Press et al., 1986) before computing the three-point centered time derivative and the “true” pressure in (13). For temperature we computed the three-point centered time derivative of the measured temperature and then smoothed it with a 5point Hanning window before applying (13). For conductivity, we assumed that the more massive conductivity cell's temperature lagged that of the “true” temperature measured with a downstream thermistor. We experimented with lagging the “true” temperature and its 5-point-Hanning-smoothed derivative by various times, used that and the measured conductivity to compute the salinity and checked for salinity spikes. By trial and error, we found that a lag time between the “true” temperature and the conductivity temperature of 0.3 s gave the best results in general, but it did not always entirely eliminate salinity spikes. The fourth step in CTD data processing involved matching the CTD time series with net haul times and positions. The pressure time series was smoothed with a Hanning window whose width in time (51 points) was chosen to be equivalent to 1 dbar at the sampling frequency of 15 Hz and a typical vertical net speed of 0.3 dbar/s. We designated that portion of each net haul a down- or upcast when the down- or upward velocity exceeded 0.08 dbar/s, respectively. Net haul logs gave the time and geographic position when each haul reached and left the sea bottom, and the down- and upcasts were assigned those times and positions. The down- and upcast time series were linearly interpolated onto pressure axes with one-dbar increments, missing values above 3 dbar were assigned their nearest-neighbor values from below. Usually the downcast was chosen to represent each CTD station, but each was compared graphically with its upcast and the latter was chosen if it had smaller spurious density inversions owing to salinity spiking.

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Please cite this article as: Cokelet, E.D., 3-D water properties and geostrophic circulation on the eastern Bering Sea shelf. Deep-Sea Res. II (2016), http://dx.doi.org/10.1016/j.dsr2.2016.08.009i