OCS Study

BOEM 2017-065

US Department of the Interior

Bureau of Ocean Energy Management

[Alaska OCS Region

Characterization of the Circulation on 

the Continental Shelf Areas of the 

Northeastern Chukchi and Western 

Beaufort Seas

OCS Study

BOEM 2017-065

US Department of the Interior

Bureau of Ocean Energy Management

[Alaska OCS Region

Characterization of the Circulation on 

the Continental Shelf Areas of the 

Northeastern Chukchi and Western 

Beaufort Seas

December 2017

Authors:

Thomas J. Weingartner1, PI; Robert Pickart2, Co-PI; and Peter Winsor1, Co-PI

Co-Authors:

W. Bryce Corlett2, Elizabeth L. Dobbins1, Ying-Chih Fang1, Cayman Irvine1, Brita Irving1, Min Li2, 

Kofan Lu1, Rachel Potter1, Hank Stascewich1, and Chase Stoudt1

Prepared under BOEM Agreement M12AC00008 by:

1

University of Alaska Fairbanks

College of Fisheries and Ocean Sciences

Fairbanks, AK 99775-7220

2

Woods Hole Oceanographic Institution

Department of Physical Oceanography

Woods Hole, MA 02543

DISCLAIMER

Study collaboration and funding were provided by the US Department of the Interior, Bureau of Ocean 

Energy Management (BOEM), Environmental Studies Program, Washington, DC, under Agreement 

Number M12AC00008. This report has been technically reviewed by BOEM, and it has been approved 

for publication. The views and conclusions contained in this document are those of the authors and should 

not be interpreted as representing the opinions or policies of the US Government, nor does mention of 

trade names or commercial products constitute endorsement or recommendation for use.

REPORT AVAILABILITY

To download a PDF file of this report, go to the US Department of the Interior, Bureau of Ocean Energy 

Management Data and Information Systems webpage (http://www.boem.gov/Environmental-StudiesEnvData/), click on the link for the Environmental Studies Program Information System (ESPIS), and 

search on 2017-065. The report is also available at the National Technical Reports Library at 

https://ntrl.ntis.gov/NTRL/.

CITATION

Weingartner TJ, Pickart R, Winsor P, et al. 2017. Characterization of the Circulation on the Continental 

Shelf Areas of the Northeast Chukchi and Western Beaufort Seas. Fairbanks, AK: US Department of 

the Interior, Bureau of Ocean Energy Management. OCS Study BOEM 2017-065. 221 p.

ABOUT THE COVER

Visible MODIS Terra image from July 13, 2016, acquired from the University of Alaska Fairbanks 

Geographic Information Network of Alaska.
ACKNOWLEDGMENTS

In addition to those mentioned in partnerships (Section A3.0), there are many people who played a role in 

the success of this endeavor. Additional partner funding was provided by Shell, ConocoPhillips Alaska, 

the Alaska Ocean Observing System, and the State of Alaska Coastal Impact Assistance Program. Special 

thanks goes to Molly McCammon (AOOS), John Cologgi, Caryn Rea, and Chris Yetsko (CPAI), A. 

Michael Macrander and Robert Raye (SEPI), Steinar Elroy (Statoil), and Sheyna Wisdom (Olgoonik 

Fairweather) for partner program management efforts. 

We thank Warren Horowitz (BOEM) for his keen interest and advice throughout this program and for his 

assistance in procuring several partners that participated in the project

We also thank the officers and crews of the R/V Norseman II, R/V Westward Wind, and USCGC Healy

for their skill and efficiency in conducting work safely at sea. John Kemp, Frank Bahr, David Leech, and 

Peter Shipton conducted the WHOI and UAF mooring operations. Leandra de Sousa, Billy Adams, 

Willard Neakok, and Robert Suydam deployed drifters under the sponsorship of the North Slope 

Borough–Shell Baseline Studies Program. James Kelly, Kim Martini, and Steve Sweet provided HFR 

field assistance, and logistical support was appreciated from UIC Science. We are grateful to the 

Ukpeagvik Iñupiat, Olgoonik, and Cully Corporations for the use of their lands and the villages of 

Barrow, Wainwright, and Point Lay, Alaska, for their cooperation and support.

Frank Bahr and Dan Torres processed all of the WHOI data, while Carolina Nobre, Peigen Lin, and 

Yiquan Qi assisted in the scientific analysis. Todd Mudge (ASL) oversaw the industry mooring 

operations and the processing of the industry mooring data with assistance from Keath Borg and Dawn 

Sadowy (ASL).

We thank Steve Okkonen and Carin Ashjian for sharing Barrow Canyon mooring data, as well as the sea 

ice remote sensing group at the University of Bremen, the National Snow and Ice Data Center, and the US

National Ice Center for providing access to sea ice data sets and NOAA/OAR/ESRL PSD for NARR 

wind data access. Additional data utilized herein was collected under previous BOEM efforts involving 

Chukchi deployments of HFR and moorings (OCS Study 2012-079), CTD casts and moorings in the 

Hanna Shoal region (OCS Study 2016-047), and drifters in the Chukchi and Beaufort (OCS Study 2015-

022). Industry, Shell/NSB baseline studies, and CSESP also provided program data for synthesis.
i

Contents

List of Figures ...............................................................................................................................................v

List of Tables ............................................................................................................................................ xvii

List of Abbreviations and Acronyms ........................................................................................................ xviii

Abstract.......................................................................................................................................................xx

A. Introduction, Objectives, and Partnerships ........................................................................................A1

A.1 Introduction ..................................................................................................................................A1

A.2 Objectives ....................................................................................................................................A3

A.3 Partnerships .................................................................................................................................A4

B. The Chukchi Slope Current................................................................................................................B1

Abstract..................................................................................................................................................B1

B.1 Introduction ..................................................................................................................................B1

B.2 Data..............................................................................................................................................B3

B.2.1 Shipboard Hydrographic and Velocity Data ..........................................................................B3

B.2.2 Surface Winds.......................................................................................................................B4

B.3 Methods .......................................................................................................................................B5

B.3.1 Gridding and Averaging ........................................................................................................B5

B.3.2 Defining the Currents ............................................................................................................B6

B.3.3 Calculating Fluxes.................................................................................................................B6

B.4 Mean Conditions ..........................................................................................................................B7

B.4.1 Structure................................................................................................................................B7

B.4.2 Transport.............................................................................................................................B11

B.5 Variability....................................................................................................................................B12

B.5.1 Seasonality..........................................................................................................................B12

B.5.2 Interannual Variability..........................................................................................................B13

B.6 Nature of the Chukchi Slope Current .........................................................................................B14

B.6.1 Wind Influence ....................................................................................................................B14

B.6.2 Path and Downstream Evolution.........................................................................................B15

B.7 Pacific Water Mass Budget of the Chukchi Shelf.......................................................................B17

B.8 Summary and Discussion ..........................................................................................................B19

C. Surface Current Patterns in the Northeastern Chukchi Sea and Their Response to Wind Forcing...C1

Abstract..................................................................................................................................................C1

C.1 Introduction ..................................................................................................................................C1
ii

C.2 Data and Methods........................................................................................................................C3

C.2.1 Surface Currents...................................................................................................................C3

C.2.2 Winds ....................................................................................................................................C4

C.2.3 SOM Analysis .......................................................................................................................C4

C.2.4 EOF Analysis ........................................................................................................................C5

C.2.5 Hydrography .........................................................................................................................C5

C.2.6 Subsurface Currents .............................................................................................................C5

C.3 Results .........................................................................................................................................C5

C.3.1 SOM-derived Patterns ..........................................................................................................C5

C.3.2 The Role of Winds: SOM Perspective ..................................................................................C8

C.3.3 The Role of Winds: EOF Perspective .................................................................................C11

C.3.4 Mean Circulation Patterns...................................................................................................C14

C.3.5 Spatial Correlation Structure...............................................................................................C16

C.4 Discussion..................................................................................................................................C16

C.5 Summary....................................................................................................................................C20

D. Transport and Thermohaline Variability in Barrow Canyon on the Northeastern Chukchi Sea Shelf D1

Abstract..................................................................................................................................................D1

D.1 Introduction ..................................................................................................................................D1

D.2 Methods .......................................................................................................................................D3

D.2.1 Moorings ...............................................................................................................................D3

D.2.2 High-Frequency Radars (HFR).............................................................................................D3

D.2.3 Winds ....................................................................................................................................D4

D.3 Results .........................................................................................................................................D4

D.3.1 Kinematic Description ...........................................................................................................D4

D.3.2 Transport: Observations .......................................................................................................D7

D.3.3 Transport: Hindcast and Climatology..................................................................................D11

D.3.4 Thermohaline Variability .....................................................................................................D15

D.4 Discussion..................................................................................................................................D18

D.4.1 Transport Comparisons ......................................................................................................D18

D.4.2 Consequences of Transport Seasonality ............................................................................D21

D.4.3 Hydrographic Implications...................................................................................................D22

D.4.4 Shelf Response to Wind-Forcing ........................................................................................D22

D.4.5 Transport Intermittency .......................................................................................................D23

D.5 Summary....................................................................................................................................D24

E. A Quasi-Lagrangian Perspective on the Surface Circulation of the Chukchi Sea..............................E1

E.1 Introduction ..................................................................................................................................E1
iii

E.2 Methods .......................................................................................................................................E2

E.2.1 Drifter Deployments ..............................................................................................................E2

E.2.2 Momentum Balance Terms ...................................................................................................E3

E.2.3 Wind and Current Variability .................................................................................................E5

E.3 Results .........................................................................................................................................E6

E.3.1 Hydrographic Background.....................................................................................................E6

E.3.2 Surface Circulation..............................................................................................................E10

E.3.3 Dispersion ...........................................................................................................................E24

E.3.4 Barrow Canyon Mesoscale Analyses..................................................................................E26

E.4 Discussion and Summary ..........................................................................................................E30

F. The Summer Hydrographic Structure of the Hanna Shoal Region on the Northeastern Chukchi Sea 

Shelf: 2011 – 2013..................................................................................................................................... F1

Abstract.................................................................................................................................................. F1

F.1 Introduction................................................................................................................................... F1

F.2 Data and Methods ........................................................................................................................ F3

F.2.1 Hydrography.......................................................................................................................... F3

F.2.2 Meteorology; Sea Ice Concentration, Drift, and Thickness; and Satellite-Tracked Drifters... F3

F.3 Results ......................................................................................................................................... F4

F.3.1 Sea Ice Conditions ................................................................................................................ F4

F.3.2 Hydrography: Broad-Scale Perspective ................................................................................ F5

F.3.3 Hydrography: Hanna Shoal................................................................................................. F11

F.3.4 The North Side of Hanna Shoal .......................................................................................... F14

F.4 Discussion .................................................................................................................................. F17

F.5 Conclusion.................................................................................................................................. F20

G. Circulation of the Chukchi Sea Shelfbreak and Slope from Moored Timeseries ...............................G1

Abstract..................................................................................................................................................G1

G.1 Introduction ..................................................................................................................................G1

G.2 Data and Methods .......................................................................................................................G3

G.2.1 Mooring Data ........................................................................................................................G4

G.2.2 Wind Data.............................................................................................................................G6

G.2.3 Atmospheric Reanalysis Fields.............................................................................................G6

G.2.4 Ice Concentration Data.........................................................................................................G6

G.3 Mean and Seasonality of the Circulation and Hydrography of the Shelfbreak and Slope ...........G7

G.4 Volume Transport of the Slope Current and Shelfbreak Jet ......................................................G12

G.5 Extreme States of the Slope Current and Shelfbreak Jet ..........................................................G12

G.5.1 Strong Slope Current and Reversed Shelfbreak Jet...........................................................G13
iv

G.5.2 Strong Shelfbreak Jet and Weak Slope Current.................................................................G16

G.6 Upwelling ...................................................................................................................................G17

G.7 Propagation of Water Mass Signals from Barrow Canyon to the Chukchi Slope ......................G20

H. Physical Oceanography of the Northeastern Chukchi Sea: A Synthesis...........................................H1

H.1 Introduction ..................................................................................................................................H1

H.2 Sea Ice.........................................................................................................................................H1

H.3 Winter Sea Ice Statistical Properties............................................................................................H5

H.4 Hydrography ..............................................................................................................................H10

H.5 Circulation Connections .............................................................................................................H21

H.5.1 Chukchi Sea Shelf ..............................................................................................................H21

H.5.2 Chukchi Sea Shelf and Alaskan Beaufort Sea Shelfbreak .................................................H33

H.6 Summary....................................................................................................................................H40

I. Recommendations for Future Study..................................................................................................... I1

J. References ..........................................................................................................................................J1

Appendix A: Surface Current Patterns in the Northeastern Chukchi Sea and Their Response to Wind 

Forcing.....................................................................................................................................Appendix A -1

Appendix B: Tidal Ellipse Properties for the Chukchi Sea Shelf..............................................Appendix B -1
v

List of Figures

Figure A1. Mean depth-integrated streamlines (white lines) in the Chukchi Sea, after Spall (2007; left). 

The bathymetry is colored and major bathymetric features are labeled. Bathymetric map of the 

northeastern Chukchi Sea, with place names included (right)...................................................................A2

Figure B1. Revised schematic of the circulation of the Chukchi Sea and western Beaufort Sea from 

Brugler et al. (2014), including an extended Chukchi shelfbreak jet and the newly described Chukchi 

slope current based on the results of this study. .......................................................................................B2

Figure B2. Locations of the 46 shipboard sections used in the study, color coded by cruise (see the 

legend and Table B1). The symbols denote the CTD stations comprising each section. The 90 m isobath 

is highlighted bold. .....................................................................................................................................B5

Figure B3. Temperature-Salinity diagram of all of the hydrographic data offshore of the 90 m isobath and 

above 300 m depth. Each point denotes a single measurement, and the color represents the percentage 

of data within each grid cell of 0.1°C temperature by 0.1 salinity. The red lines indicate the different water 

mass boundaries. ACW = Alaskan coastal water; BSSW = Bering summer water; WW = newly-ventilated 

Pacific winter water; RWW = remnant Pacific winter water; MW = meltwater; AW = Atlantic water..........B8

Figure B4. Mean vertical sections constructed from the 46 shipboard transects. (a) Potential temperature, 

and (b) Absolute geostrophic velocity overlain by potential density (contours, kg m-3). The black dashed 

lines denote the 5 cm s-1 velocity anomaly contour outlining the shelfbreak jet and slope current (see text). 

Positive velocities are eastward (the viewer is looking westward). The number of sections contributing to 

the mean at each cross-slope location is plotted in the top panel. ............................................................B9

Figure B5. Relative occurrence of the different water masses in the mean hydrographic section. The black 

dashed lines denote the outline the shelfbreak jet and slope current as in Figure B4.............................B10

Figure B6. Transport of the Chukchi slope current broken down by water mass (see Figure B3 for the 

water masses). Standard errors are indicated by the lines. (a) Volume transport; (b) Heat transport; (c) 

Freshwater transport................................................................................................................................B11

Figure B7. Mean monthly occurrence of water masses on the Chukchi slope offshore of the 90 m isobath 

and shallower than 300 m. Standard errors are indicated. (See Figure B3 for the water masses.) ........B12

Figure B8. Mean monthly Pacific water transport of the Chukchi slope current with standard errors; N = 11 

for July, 10 for August, 5 for September, and 8 for October. (a) Volume transport; (b) Heat transport; (c) 

Freshwater transport................................................................................................................................B13

Figure B9. Interannual change in occurrence of water masses on the Chukchi slope from the early regime 

to the recent regime, offshore of the 90 m isobath and shallower than 300 m. Standard errors are 

indicated. (See Figure B3 for the water masses.)....................................................................................B14

Figure B10. Interannual change in Pacific water transport of the Chukchi slope current from the early 

regime to the recent regime, with standard errors; N = 18 for the early regime, and 16 for the recent 

regime. (a) Volume transport; (b) Heat transport; (c) Freshwater transport. ...........................................B14

Figure B11. Mean along-shelfbreak 10 m wind speed from July – August each year, averaged over the 

Chukchi slope region (155 – 162°W and 72 – 74°N). The standard errors are indicated. Negative wind 

speeds are easterly. ................................................................................................................................B15

Figure B12. Composite vertical sections of absolute geostrophic velocity (color) and potential density 

(contours, kg m-3) during (a) calm wind conditions; and (b) times when the easterly winds exceeded 4 m 

s

-1 prior to and during the occupation of the transect...............................................................................B16
vi

Figure B13. (a) Map showing the position and strength of the Chukchi slope current for each of the 

transects that measured the current. The width of the arrows denotes the lateral scale of the current, and 

the length of the arrows denotes the mean speed of the current averaged over its width and depth. The 

shipboard transects are indicated by the dashed lines. (b) Same as (a) except for the Chukchi shelfbreak 

jet. ............................................................................................................................................................B17

Figure B14. Mass budget of the Chukchi shelf inflows and outflows as discussed in the text, including the 

residual. Positive transports are inflows and negative transports are outflows. Standard errors are 

indicated. .................................................................................................................................................B19

Figure C1. (a) Map of the Chukchi Sea with place names. Blue dots indicate 2012 HFR network grid 

points with more than 60% temporal coverage. Red squares show locations of HFR field sites. The red 

outlined area represents the NARR wind domain used in the SOM and EOF analyses, with the red 

triangle the location representative of the shelf wind time series. Legs 1 and 2 are hydrographic transects 

conducted in September 2013. Black dot north of Wainwright denotes the BC2 mooring. Gray arrows 

show schematic pathways of Pacific waters. Bathymetric contours are drawn from 10 – 100 m at 10 m 

intervals. Place names include HV = Hope Valley, HC = Herald Canyon, HeS = Herald Shoal, CC = 

Central Channel, and HaS = Hanna Shoal. (b) Grid points within the red rectangle are used to determine 

upcanyon and downcanyon flow conditions. .............................................................................................C2

Figure C2. Representative circulation regimes categorized from twelve SOM-derived patterns of surface 

currents (blue vectors) for 2012: (a) northeastward-flowing regime and (b) reversal regime. The frequency 

of regime occurrence is included in each panel, and the 80 m isobath is thicker to define Barrow Canyon. 

The inserted scatter plot denotes vertically-averaged velocities from mooring BC2 (black dot) when the 

flow regime occurred. Polar histograms on the right denote wind velocities associated with the flow 

regime (red triangle). The direction follows oceanographic convention and speed is shaded. The 

percentage indicates frequencies of winds blowing toward that direction. ................................................C6

Figure C2. (continued) (c) northwesterly wind regime and (d) divergent mode. Note that the scales of 

current vectors in (c) and (d) are different than those in (a) and (b). .........................................................C7

Figure C3. Monthly occurrences of SOM-derived circulation regimes for 2010 – 2014 with different 

hatching denoting different years. (a)-(c) Northeastward-flowing regime (NE). (d)-(f) Reversal regime (R). 

(g)-(i) Divergent mode (D). (j)-(l) Northwesterly wind regime (NW). A 15-day data gap in August 2013 is 

the reason for few regime estimates in that month....................................................................................C9

Figure C4. (a) Upper panel: time series of wind vectors in August 2012. The vector direction follows 

oceanographic convention. Lower panel: SOM-derived circulation regimes (black dots) in August 2012 

and normalized data returns (gray line). The abbreviations are: R, reversal regime; D, divergent mode; G, 

pattern G; NW, northwesterly wind regime; NE, northeastward-flowing regime. (b) As in (a), but for 

September 2012. (c) As in (a), but for October 2012...............................................................................C10

Figure C5. (a) EOF Mode 1 of 2012 reconstructed HFR data (blue vectors) and its explained variance. 

For clarity, only subsampled vectors are shown. (b) As in (a), but for Mode 2. (c) EOF Mode 1 of 2012 

NARR winds (black vectors), also subsampled. (d) As in (c), but for Mode 2. ........................................C12

Figure C6. (a) Time series of wind speed with wind speeds ≥6 m s-1 highlighted in gray. (b) Time series of 

wind direction with gray shading corresponding to winds blowing from the northeast quadrant. (c) Time 

series of PC1 of surface currents (black) and PC1 of winds (gray) with gray shaded areas highlighting 

positive PC1 values. PC1 values approaching one indicate currents or winds approximating the Mode 1 

structure. (d) Time series of representative circulation regimes derived from SOM (black dots): R, reversal 

(highlighted in gray); D, divergent mode; G, pattern G; NW, northwesterly wind; NE, northeastwardflowing. The gray shaded area highlights the reversal regime. Vertical red lines in each panel denote 

periods of selected reversal regimes (T1, T2, T3, and T4). Red arrow and bar denote a northeasterly wind 

event described in the text.......................................................................................................................C13
vii

Figure C7. Mean (a) downcanyon and (b) upcanyon surface currents (blue vectors) for 2010 – 2014. 

Bathymetric contours (gray lines) are drawn within 200 m at 40 m intervals, with depths less than 40 m 

drawn at 10 m intervals. The black vector denotes the mean winds, and the polar histogram along 72oN 

denotes wind velocities. Areas circumscribed by dashed lines denote regions (labeled 1 – 4) of distinct 

flow behaviors (see text). For clarity, the vectors were subsampled. ......................................................C15

Figure C8. Spatial distribution of complex correlations of surface currents with respect to a reference grid 

point (large black dot) for (a, c) downcanyon and (b, d) upcanyon flow. Black contour denotes the 0.8 

correlation level. Correlation below the e-folding scale is shaded in cool colors. The 95% significance 

level is estimated from the effective number of degrees of freedom using the integral time scale..........C17

Figure C9. Vertical sections from September 2013 of potential temperature and salinity overlain with 

isopycnals (white contours) for (a, c) Leg 1 and (b, d) Leg 2. Black arrow denotes the location of 71.5°N. 

Note that the shading scales in (a) and (b) and the horizontal scales for Legs 1 and 2 are different......C18

Figure C10. HFR observations (blue vectors) during 12 – 19 September 2011 at indicated UTC time. The 

black vector denotes the wind at the cardinal hour of the snapshot, with the magnitude and direction 

shown in the legend, and the gray shaded vectors are winds for the preceding 12 and 24 hours. 

Bathymetric contours (gray lines) are drawn within 80 m at 10 m intervals. For clarity, the current vectors 

were subsampled.....................................................................................................................................C19

Figure D1. Bathymetric map of the northeastern Chukchi Sea showing mooring locations. The 2010 –

2012 arrays of 6 moorings are shown with black dots and consist of moorings BC1 (closest to the coast) 

through BC6. Additional moorings, summarized in the discussion, are shown using colored symbols. 

HFRs were located in Point Lay, Wainwright, and Barrow. The striped regions encapsulate that portion of 

the HFR mask containing 50% or more good data....................................................................................D2

Figure D2. A) Vertically-averaged velocity vectors and variance ellipses for the moorings at the head of 

Barrow Canyon. Note the velocity scale change between the mean vector and the variance ellipse. B) 

Time series of the along-canyon currents at 20 m depth at mooring BC2. Both figures are based on the 

August 2010 – August 2011 mooring records. ..........................................................................................D5

Figure D3. Vertical sections of the mean annual: a) along-canyon, ur and b) cross-canyon, vr velocities. 

Mean sections of c) ur and d) vr for the downcanyon case, and mean sections of e) ur and f) vr for the 

upcanyon case. Vertical sections of the first EOF eigenvector for g) ur and h) vr. All figures are based on 

the August 2010 – August 2011 period. ....................................................................................................D6

Figure D4. Mean daily along-canyon winds (top) and transport (bottom). The dashed red line on the wind 

and transport time series denotes the record length means of -1.7 m s-1 and 0.19 Sv, respectively. .......D8

Figure D5. Mean monthly along-canyon winds (top) and transport (bottom) are denoted by black lines and 

dots. Crosses and gray lines delineate the 95% confidence limits. Dashed red lines on show the recordlength means in transport (0.19 Sv) and along-canyon wind velocity (1.7 m s-1

). ...................................D10

Figure D6. Panels A and B are the means (filled circles) and standard deviations (error bars) of the 

monthly transport and along-canyon wind velocity, respectively, for 2010 – 2015. Panel C is the mean 

monthly transport (black line and circles) and 95% confidence limits (error bars) based on the January 

1979 – July 2015 GLM hindcast. The gray lines and circles indicate the maximum and minimum values in 

each month. Panel D displays the monthly means (circles) and standard deviations (squares) of the 

pseudostress for the January 1979 – July 2015 period at the three locations used in the GLM. ............D10

Figure D7. Probability density functions of the mean daily transport in Barrow Canyon for winter (October 

– April; top) and summer (May – September; bottom). Each panel includes the value of the mean, sample 

variance (s

2

), and the skewness (Skw) and kurtosis (Kur) coefficients, computed per King and Julstrom 

(1982). The vertical dashed line delineates the upcanyon from the downcanyon class intervals. ..........D11
viii

Figure D8. Mean monthly maps of sea level pressure (contours), wind speed (color), and direction 

(arrows) over the northern Bering and Chukchi seas. The 2011 maps are for: A) June, B) July, C) August, 

and D) September and the 2012 – 2013 maps are for: E) December, F) January, G) February, and H) 

March.......................................................................................................................................................D12

Figure D9. Monthly transport anomalies (gray) for the 1979 – 2015 period based on the GLM hindcast. 

The thick red curve is the 13-month running mean of the anomalies. .....................................................D14

Figure D10. Sea level pressure (contours), surface wind vectors (arrows), and wind speeds (colors) for 

Marches (left) and Octobers (right) in which there was anomalously large upcanyon (top row; contour 

interval of 5 mb) and downcanyon transports (bottom row; contour interval of 2 mb). The maps are 

composites based on the month and years shown to the right of each map, which includes the mean 

monthly transport for those years. ...........................................................................................................D14

Figure D11. The seasonal cumulative probabilities for a) upcanyon and b) downcanyon events based on 

the hindcast transport record. The numbers associated with the seasons in the legend on each graph 

refer to the number of events over the 37-year record. ...........................................................................D15

Figure D12. Time series of salinity (top) and potential temperature, θ, (bottom) at moorings BC1 (black), 

BC2 (blue), and BC5 (red) for the August 2010 – August 2012 period. ..................................................D16

Figure D13. Time series of temperature (bottom) and salinity (middle) from mooring BC2 and the alongcanyon transport (smoothed with a 7-day running mean; top). The horizontal line on the salinity plot 

delineates the 33.1 value. The light grey shading encompasses the periods of upcanyon transports that 

contained lower halocline and Atlantic Water signatures (e.g., S>33.1 and θ>-1.6°C). The light yellow 

shading captures periods of dense water (S>33.1 and θ = freezing point)..............................................D17

Figure D14. Vertically-averaged mean velocity vectors and variance ellipses for currents and winds in A) 

2010 – 2011, B) 2011 – 2012, and C) 2012 – 2013. Red symbols and text are for winds, black are for

currents outside of the blue box, and blue are for currents within the blue boxes. Gray (black) vectors and 

ellipses in each figure denote record lengths of from 6 to 10 (11 to 12) months duration. ......................D19

Figure D15. Mean September – October 2013 surface current map of the western Beaufort Sea shelf and 

the northern end of Barrow Canyon. The mean winds at Barrow were from the east-northeast (as 

indicated by the arrow) at 3 m s-1. The red squares at the mouth of Barrow Canyon are the locations of 

the moorings discussed by Itoh et al. (2013). ..........................................................................................D21

Figure E1. Bathymetric map of the Chukchi Sea showing the drifter deployment locations (triangles) in 

2012 (magenta), 2013 (green), and 2014 (blue). The 2014 and 2013 deployments are used in evaluating 

the momentum balances and the 2014 deployments were used to update relative dispersion estimates 

presented by Weingartner et al. (2015). The villages of Point Hope (PH), Point Lay (PL), Wainwright (W), 

and Barrow (B) are indicated. The black horizontal line offshore of Point Lay shows the August 2013 

trajectory of a glider and the red horizontal line shows the September 2013 location of the Acrobat CTD 

section northeast of Hanna Shoal..............................................................................................................E3

Figure E2. Delineation of the five different classification groups used to examine the fates of drifters 

moving northeastward through Barrow Canyon. A drifter must pass northeastward across the “Start Line” 

in order to be considered in the analysis. The yellow star indicates the location of the point of interest

(POI) on the 100 m isobath........................................................................................................................E5

Figure E3. Map showing the shelfbreak and slope “boxes” used to evaluate mean and eddy kinetic 

energies. These regions are the Chukchi (blue), Barrow Canyon (green), and Beaufort (yellow). ...........E6
ix

Figure E4. Top panel: map showing location (indicated by red line) of the west-to-east transect of glider 

191 from 10 to 14 August 2013. The bottom three panels are the vertical sections of temperature (top), 

salinity (middle), and density (expressed as sigma-t; bottom) along the transect. The westernmost end of 

the transect is on the left of the figure........................................................................................................E7

Figure E5. Vertical sections of temperature, salinity, and σt from an Acrobat transect along the 

northeastern side of Hanna Shoal on 11 September 2013. The western part of the section is on the lefthand side of the figure. ..............................................................................................................................E8

Figure E6. Vertical sections of temperature, salinity, and σt from an Acrobat transect across the top of 

Hanna Shoal on 10 September 2013. The northern part of the section is on the left-hand side of the 

figure..........................................................................................................................................................E9

Figure E7. Vertical sections of temperature, salinity, and σt from an Acrobat transect on the northwest 

side of Hanna Shoal on 10 September 2013. The southeast end of the section is on the left-hand side of 

the figure....................................................................................................................................................E9

Figure E8. Trajectory of the Point Lay 2013 cluster centroid (color-coded by SST) and the 13 individual 

drifters within this cluster (top panel). Additional plots are the time series of the centroid (2nd from top) and 

wind (3rd from top) velocities and SST (bottom). The colored symbols along each trajectory are at 5-day 

intervals and correspond to the date axis on the SST time series plot. Events A, B, and CC1 – 4 are 

discussed in the text. ...............................................................................................................................E11

Figure E9. Time series of the zonal (top) and meridional (bottom) velocity components associated with 

the Stokes (blue), Ekman (red), and geostrophic (yellow) contributions to the surface velocity. The letters 

and vertical lines correspond to the periods in Figure E8 and discussed in the text. ..............................E12

Figure E10. As in Figure E8 but for the Point Lay 2014 cluster. Segments D, E, and F are discussed in 

the text. ....................................................................................................................................................E13

Figure E11. Time series of the zonal (top) and meridional (bottom) velocity components associated with 

the Stokes (blue), Ekman (red), and geostrophic (yellow) contributions to the surface velocity for the Point 

Lay 2014 deployment. The letters and vertical lines correspond to the periods in Figure E10 and are 

discussed in the text. ...............................................................................................................................E14

Figure E12. As in Figure E8 but for the Point Hope 2014 northerly cluster. Segments G, H, and I are 

discussed in the text. ...............................................................................................................................E15

Figure E13. As in Figure E8 but for the Point Hope 2014 southerly cluster. Segments J, K, L, and M are 

discussed in the text. ...............................................................................................................................E16

Figure E14. Time series of the four terms comprising the deformation tensor (div = divergence, RV = 

relative vorticity, STR DEF = stretching deformation, and SHR DEF = shearing deformation. The Point 

Hope North cluster terms are in blue and those for the Point Hope South cluster are red. Units are 10-5 s

-1

.

.................................................................................................................................................................E18

Figure E15. Time series of the terms comprising the absolute vorticity tendencies for the Point Hope 

North (blue) and Point Hope South (red) clusters. From top to bottom these are the rate of change of 

absolute vorticity, stretching vorticity, and the sum of the first two terms. Units are10-9 s

-2

.....................E18

Figure E16. Time series of the zonal (top) and meridional (bottom) velocity components associated with 

the Stokes (blue), Ekman (red), and geostrophic (yellow) contributions to the surface velocity for the Point 

Hope North deployment. The letters and vertical lines correspond to the periods in Figure E12 and are 

discussed in the text. ...............................................................................................................................E19
x

Figure E17. Time series of the zonal (top) and meridional (bottom) velocity components associated with 

the Stokes (blue), Ekman (red), and geostrophic (yellow) contributions to the surface velocity for the Point 

Hope South deployment. The letters and vertical lines correspond to the periods in Figure E13 and are 

discussed in the text. ...............................................................................................................................E20

Figure E18. As in Figure E8 but for the NE Hanna Shoal cluster. Segments N and O are discussed in the 

text. ..........................................................................................................................................................E21

Figure E19. Time series of the zonal (top) and meridional (bottom) velocity components associated with 

the Stokes (blue), Ekman (red), and geostrophic (yellow) contributions to the surface velocity for the NE 

Hanna Shoal deployment. The letters and vertical lines correspond to the periods in Figure E18 and are 

discussed in the text. ...............................................................................................................................E22

Figure E20. As in Figure E8 but for the NW Hanna Shoal cluster. Segments P, Q, and R are discussed in 

the text. ....................................................................................................................................................E23

Figure E21. Time series of the zonal (top) and meridional (bottom) velocity components associated with 

the Stokes (blue), Ekman (red), and geostrophic (yellow) contributions to the surface velocity for the NW 

Hanna Shoal 2013 deployment. The letters and vertical lines correspond to the periods in Figure E20 and 

are discussed in the text. .........................................................................................................................E24

Figure E22. Time series of relative zonal (Dx

2

; top left) and meridional (Dy

2

; top right) dispersion, the 

cross-correlation term (Dxy) bottom left), and the relative dispersion (D2; bottom right) for the Point Lay 

2014 cluster. As indicated in the upper left, each panel includes the mean (μ) value and the mean value ±

one standard deviation (s). ......................................................................................................................E25

Figure E23. Time series of relative zonal (Dx

2

; top left) and meridional (Dy

2

; top right) dispersion, the 

cross-correlation term (Dxy; bottom left), and the relative dispersion (D2; bottom right) for the Point Hope 

northern cluster. Each panel includes the mean (μ) value and the mean value ± one standard deviation 

(s) per Figure E22....................................................................................................................................E25

Figure E24. Time series of relative zonal (Dx

2

; top left) and meridional (Dy

2

; top right) dispersion, the 

cross-correlation term (Dxy; bottom left), and the relative dispersion (D2; bottom right) for the Point Hope 

southern cluster. Each panel includes the mean (μ) value and the mean value ± one standard deviation 

(s) per Figure E22....................................................................................................................................E26

Figure E25. Time series of relative zonal (Dx

2

; top left) and meridional (Dy

2

; top right) dispersion, the 

cross-correlation term (Dxy; bottom left), and the relative dispersion (D2; bottom right) for the entire Point 

Hope cluster. Each panel includes the mean (μ) value and the mean value ± one standard deviation (s)

per Figure E22. ........................................................................................................................................E27

Figure E26. Trajectories and summary statistics for all 1 m drogued drifters according to their turning 

classification.............................................................................................................................................E29

Figure E27. Log-log plots of the relative meridional (Dy

2

; top), zonal (Dx

2

; middle), and horizontal (D2; 

bottom) dispersion based on the 2012, 2013, and 2014 clusters. The black dashed lines are the 95% 

confidence interval on the mean (blue curve). On the bottom plot, the Lagrangian time scale (TL) is ≤2 

days and denoted by the red, dotted line and the red solid line denotes a dispersion growth rate 

proportional to t

2

.......................................................................................................................................E31

Figure F1. A) Bathymetric map of the Chukchi Sea showing principal flow pathways northward from 

Bering Strait. The pathway denoted by the red arrow denotes the Alaskan Coastal Current, which carries 

the freshest and warmest fraction of the Bering Strait outflow and the blue arrows denote the pathways 

for the cooler and saltier fractions. B) Map of Hanna Shoal and the surrounding shelf showing locations of 

2012 (blue) and 2013 (red) CTD stations. C) Location of various current meter and ice profiling sonar 

(IPS) moorings (symbols) and Acrobat towed-CTD transects (blue lines). ............................................... F2
xi

Figure F2. Sea ice concentration maps on 15 June (top) 15 July (middle) and mean daily wind vectors 

(bottom). Panels a, b and c are for 2011, d, e, and f are for 2012, and g, h, and i are for 2013. The 

features labeled in map E are: Herald Valley (HV), Herald Shoal (HeS), Central Channel (CC), Hanna 

Shoal (HaS), and Barrow Canyon (BC). Wind vector plots include mean monthly wind velocity and wind 

speed in parentheses. ............................................................................................................................... F4

Figure F3. Maps of the 15% ice concentration isopleth at biweekly intervals for the August – September 

period (top row) and mean daily wind vectors (bottom row). Panels a and b are for 2011, c and d are for 

2012, and e and f are for 2013. The mean monthly wind velocity and wind speed (in parentheses) are 

listed below the x-axis for each year.......................................................................................................... F6

Figure F4. Plan views of temperature (left) and salinity (right) in August and September of 2012 based on 

averages of the upper and bottom 10 m of the water column. The village of Wainwright is designated with 

a “W”. Filled circles show CTD station locations........................................................................................ F7

Figure F5. Plan views of temperature (left) and salinity (right) from August and September 2011 based on 

averages of the upper and bottom 10 m of the water column. The village of Wainwright is designated with 

a “W”. Filled circles show CTD station locations. The black line in panel c shows the location of the 

vertical sections from 2011 and 2102 shown in Figure F6. ....................................................................... F8

Figure F6. Vertical sections of potential temperature (top row) and salinity (bottom row) collected from 

early to late of September 2011 (a, b) and 2012 (c, d), respectively. The location of both sections is 

shown in Figure F5c. Hanna Shoal is denoted as “HS”. Inverted triangles indicate station locations. ...... F9

Figure F7. Vertical sections of potential temperature (a, b), salinity (c, d), and fluorescence expressed as 

μg l-1 of chlorophyll a (e, f) in September 2012 along Leg M (left) and Leg Q (right). The white contours 

indicate σθ isolines. The arrows at the top of each figure indicate latitude 71.5°N. The locations of these 

transects are shown in Figure F1c. The front along Leg M is located at 71.2°N as indicated. ................ F10

Figure F8. Histograms of density variance at 10 m depth as a function of horizontal length scale based on 

2013 Acrobat CTD tows collected north of the front (gray stippled bars) and south (black bars) of the 

front. The horizontal bars denote the mean density variance in each band for the gray bars. ................ F11

Figure F9. Vertical sections of potential temperature (θ; top) and salinity (S; bottom) along transects 1-12 

(a, b), 2-12 (c, d), and 3-12 (e, f) on the northwestern, northern, and northeastern sides of Hanna Shoal, 

respectively. Figure F1 shows transect locations. Hanna Shoal is on the left of each panel. ................. F12

Figure F10. As in Figure F9 but for transects 1-13 (a, b), 2-13 (c, d), and 3-13 (e, f) on the northwestern, 

northern, and northeastern sides of Hanna Shoal, respectively. Figure F1 shows transect locations. 

Hanna Shoal is on the left of each panel. ................................................................................................ F12

Figure F11. As in Figure F9 but for transects 4-12 (a, b), 5-12 (c, d), and 4-13 (e, f). Panels a, b, e, and f 

are on the southern side of Hanna Shoal, and panels c and d are on the southeastern side. Figure F1 

shows transect locations. Hanna Shoal is on the left of each panel. Note the temperature scale differs 

from that of Figures F9 and F10. ............................................................................................................. F13

Figure F12. The trajectories of 6 satellite-tracked drifters (1 m drogue depth) deployed offshore of Icy 

Cape on 12 August 2012. One of the trajectories is color-coded according to measured SST. The various 

symbols along each trajectory are at 7-day intervals with the dates given on the wind vector time series 

(bottom plot). The 20, 40, 50, 100, and 1,000 m isobaths are labeled. Green triangles signify the 

deployment position and red stars the last good transmission from each drifter..................................... F15
xii

Figure F13. Scatterplots of potential temperature versus salinity over the upper 30 m of the water column 

for sections 1, 2, and 3 in 2012 (left) and 2013 (right). Only stations along the first 65 km of the transect 

are included. Note the scales change between 2012 and 2013. The black slanted line denotes the 

freezing point curve. The larger symbols outlined in yellow depict the mean values along each section in 

the upper 15 m......................................................................................................................................... F16

Figure F14. Baroclinic, geostrophic velocity field for transects a) 3-12, b) 2-13, and c) 3-13. The 

corresponding temperature and salinity sections are shown in Figures F9 e, f; F10 b, d; and F10 e, f, 

respectively. Positive velocities are north-northwestward and negative values are south-southeastward.

................................................................................................................................................................. F16

Figure F15. Each row shows the mean January – March ice and wind velocities (left panels) and regional 

wind vectors and sea level pressure (right side) over the northern Bering and Chukchi seas for 2011 (top 

row), 2012 (middle row), and 2013 (bottom row). Black arrows and ellipses denote the mean ice velocity 

and the corresponding velocity variance ellipses, while the red arrows and ellipses are for the winds... F19

Figure G1. Schematic circulation in the Chukchi Sea (from Corlett and Pickart 2017), showing the three 

pathways by which Pacific water flows poleward through the Chukchi Sea..............................................G2

Figure G2. (a) Bathymetric map of the northeastern Chukchi Sea showing mooring locations used in the 

study. The six moorings comprising the shelfbreak/slope array are shown by the yellow stars. The three 

additional moorings east of Hanna Shoal and the mooring at the head of Barrow Canyon are shown by 

the red and blue stars, respectively. The red line and black coordinate frame indicate the rotated 

coordinate system. The bathymetry is from IBCAO v3. (b) Large-scale map showing the entire Chukchi 

Sea. The region in (a) is indicated by the dashed box. The magenta and cyan boxes delineate the domain 

over which the ice concentration is calculated for the shelfbreak/slope array and for the coastal polynya 

region south of Barrow Canyon, respectively. The mooring sites are shown by the black dots. (c) 

Configuration of shelfbreak/slope moorings in the vertical plane. The origin of the x-axis is Hanna Shoal.

...................................................................................................................................................................G4

Figure G3. Year-long, depth-mean velocity vectors (blue) at the mooring sites and mean 10 m wind vector 

(black) at the Barrow, AK meteorological station. The standard error ellipses are shown (see the scales at 

the lower left). The red line indicates the along-stream direction (see Figure G2). ...................................G7

Figure G4. (a) Year-long mean along-stream velocity section (positive is southeastward). The thick black 

line is the zero velocity contour. The dashed black line shows the boundary of shelfbreak and slope. The 

grey shading denotes regions of no data coverage. The mooring sites are shown at the top. (b) Potential 

temperature-Salinity diagram for all of the hydrographic data. The color represents the percentage of data 

within a 0.1°C by 0.1°C grid. The black lines delimit the different water masses: MW = meltwater; ACW = 

Alaskan Coastal Water; BSSW = Bering Sea Summer Water; RWW = Remnant Winter Water; WW = 

newly-ventilated Winter Water; AW = Atlantic Water.................................................................................G8

Figure G5. Vertical sections of the seasonally averaged along-stream velocity. The presentation is the 

same as in Figure G4a. .............................................................................................................................G9

Figure G6. Depth-time plot of potential temperature and salinity at CS4 (a, c) and CS5 (b, d), overlain by 

potential density (contours)......................................................................................................................G10

Figure G7. Wind rose showing wind speed and direction at Barrow for the duration of the deployment.G11

Figure G8. (a) Daily-mean wind velocity at the Barrow meteorological station (blue vectors). The gray 

shading and red vectors denote periods of northeasterly wind, and the corresponding mean wind velocity, 

preceding the three major occurrences of reduced ice cover at the mooring array site and south of Barrow 

Canyon. (b) Ice concentration timeseries at the array site (cyan curve) and at the location of the polynya 

south of Barrow Canyon (magenta curve). The gray shading indicates the three periods of reduced ice 

concentration. ..........................................................................................................................................G11
xiii

Figure G9. Volume transport timeseries of (a) the Chukchi slope current (blue curve), and (b) the Chukchi 

shelfbreak jet (yellow curve is the full transport and green dashed curve is the positive transport). (c) 

Monthly-averaged slope current transport with standard errors. (d) Monthly-averaged transport and 

standard errors for the full shelfbreak jet (yellow curve) and for times when the jet transport is positive 

(green dashed curve)...............................................................................................................................G13

Figure G10. Composite velocity sections during (a, b) calm wind conditions and (c, d) times when the 

easterly winds exceeded 4 m s-1 for the months of July – October. (a) and (c) are reconstructed sections 

of absolute geostrophic velocities (color) from Corlett and Pickart (2017). (b) and (d) are velocity sections 

(color) from mooring data. The thin black lines are velocity contours......................................................G14

Figure G11. Composite average fields for the strong slope current and reversed shelfbreak jet events with 

southwesterly wind. (a) Vertical section of alongstream velocity. The number of events and their total 

duration is marked on the lower left. (b) Vertical section of alongstream velocity anomaly (composite 

minus the year-long mean). (c) Sea level pressure (color) and 10 m wind vectors from NARR, along with 

the measured wind from the Barrow meteorological station (purple vector). The location of the 

shelfbreak/slope mooring array is indicated by the purple star. (d) Wind stress curl (color). ..................G15

Figure G12. Same as Figure G11, except the strong slope current and reversed shelfbreak jet events with 

northeasterly wind....................................................................................................................................G16

Figure G13. Same as Figure G11, except the strong shelfbreak jet and weak slope current events with 

northeasterly wind....................................................................................................................................G17

Figure G14. Composite average fields prior to the strongest upwelling event (event 9, May 2014). (a) Sea 

level pressure (color) and 10 m wind vectors from NARR, along with the measured wind from the Barrow 

meteorological station (blue vector). The location of the shelfbreak/slope mooring array is indicated by the 

purple star. (b) Along-stream velocity section. (c) Density timeseries at mooring CS2 and FM1 from 3 

days before the event to 3 days after the event, where the bold indicates the time period before the 

upwelling. The dashed lines are the 20-day low-passed curves. (d) Wind stress curl (color). (e) Crossstream velocity section (positive is northeastward). (f) Depth-averaged (0 – 250 m) velocity vectors at the 

mooring sites (blue arrows) and the mean velocity vector of all moorings (large arrow).........................G19

Figure G15. Same as Figure G14, except for the time period during the upwelling. ...............................G19

Figure G16. Same as Figure G14, except for the time period after the upwelling...................................G20

Figure G17. Timeseries of water mass occurrence at the Chukchi shelfbreak/slope moorings (CS2, CS4) 

and the mooring at the head of Barrow Canyon (BC2). See Figures G2 and G3 for mooring locations. 

Note that there is no depth scale for moorings CS2 and BC2 since these sites have a single sensor near 

the bottom. The gray shading indicates times when the flow is in the opposite direction of the predominant 

current (southwestward, southeastward, and northwestward in the three panels, respectively). The black 

triangles denote the time periods when the bulk of Winter Water was present. ......................................G21

Figure G18. Potential temperature of the Winter Water at (a) mooring BC2 at head of Barrow Canyon 

(near the bottom) and CS4 on the Chukchi slope. The light gray shading means that there is no WW 

present. The dark gray shading indicates when the flow is in the opposite direction of the primary current 

(southwestward and southeastward in (a) and (b), respectively). The black triangles denote the time 

periods when the bulk of WW appeared..................................................................................................G22

Figure H1. Map showing the locations of moorings used in this section of the report as a function of year. 

Moorings Cj (Crackerjack), Bu (Burger), Site1 & 2, CPAI 01 & 02, Statoil 3 & 4, and HS 01-06 were all 

deployed and processed under industry support. HSNE and HSNW moorings were supported by the 

BOEM-COMIDA program. WBC and EBC were supported by NSF. Barrow Canyon moorings (BC 1 – 6) 

were supported by BOEM Study 2012-079, and BC2 from 2012 – 2016 was supported by this project. .H2
xiv

Figure H2. The spatial structure of the most significant EOFs of ice concentration for May (top row), June 

(middle row), July (bottom left), and August (bottom right). Where sketched, the white dotted lines outline 

the region of maximum spatial variability in EOF1.....................................................................................H3

Figure H3. The evolution of sea ice concentrations as seen on 30 October (top row), 15 November 

(middle row), and 30 November (bottom row) for 2011 (left column), 2012 (middle column), and 2013 

(right column).............................................................................................................................................H6

Figure H4. Ice displacement as a function of time for the Burger deployments from 2010 – 2011 (top), 

2011 – 2012 (middle), and 2012 – 2013 IPS and ADCP deployments. The stars indicate the starting 

position and correspond to the first 50 km segment in which any ice was recorded. Red dots indicate the 

mid-point of each 50 km segment and every other segment is labeled by date........................................H7

Figure H5. Mean ice thickness and statistical fits as a function of time (50 km segment) for the Burger 

deployments from 2010 – 2011 (top), 2011 – 2012 (middle), and 2012 – 2013 (bottom). Means encircled 

by a red ellipse suggest periods of new, thinner ice formation, and those with a blue ellipse suggest 

periods when thicker ice was advected over the site.................................................................................H8

Figure H6. Probability density functions (pdfs) for selected dates from the Burger 2010 – 2011 (left), 

Burger 2011 – 2012 (middle), and Burger 2012 – 2013 (right) IPS deployments. The y-axes represent the 

percentage of time a particular ice thickness interval (x-axis) occurred over the 50 km long sampling 

interval. Numbers in parentheses in legend indicate the percentage of all observations less than or equal 

to the maximum value of the ice thickness on the x-axis. Note that the x-axis maximum is 2 m for the top 

row and 4 m for the bottom two rows.......................................................................................................H10

Figure H7. Time series of the primary (blue) and secondary (red) mode ice thickness modes along with 

logarithmic statistical through time. The modes are determined from each 50 km segment for the Burger 

deployments from 2010 – 2011 (top), 2011 – 2012 (middle), and 2012 – 2013 (bottom). ......................H11

Figure H8. Potential temperature-salinity (θ/S) diagram illustrating various water masses found on the 

northeastern Chukchi Sea shelf in summer and early fall. ......................................................................H12

Figure H9. Summary figures for Acrobat Leg Y. The left hand side consists of vertical sections of 

temperature (top), salinity (middle), and chlorophyll (bottom). The black contours on each section are 

isolines of potential density plotted at a contour interval of 2. The right hand sides of each figure contain a 

transect location map (top right) and a θ/S plot based on the data from that section. The view is toward 

the southwest with land to the left of each panel. ....................................................................................H13

Figure H10. As in Figure H6 but for Acrobat Leg E. ................................................................................H13

Figure H11. As in Figure H6 but for Acrobat Leg BB...............................................................................H14

Figure H12. As in Figure H6 but for Acrobat Leg Q.................................................................................H14

Figure H13. As in Figure H6 but for Acrobat Leg S. ................................................................................H15

Figure H14. As in Figure H6 but for Acrobat Leg U. ................................................................................H15

Figure H15. Mean surface current vectors for the week of 5 September 2012, as estimated from HFRs.

.................................................................................................................................................................H16

Figure H16. Color-coded (by SST) trajectory of a satellite-tracked drifter released at the location of the 

green triangle on 13 August 2012. The drifter’s last reported position was on 15 November 2012 at the 

location of the red star. ............................................................................................................................H17
xv

Figure H17. Time series of vertically averaged current vectors (top), temperature (middle) and salinity 

(bottom) at 47 m depth from mooring HSNW50-12. The timing of the temperature maximum (Tmax) and 

the descent to the freezing point (Tf) are noted. ......................................................................................H18

Figure H18. Time series of vertically averaged current vectors (top), temperature (middle) and salinity 

(bottom) at 23 m (red) and 47 m (blue) depth from mooring HSNE50-12. The timing of the temperature 

maximum (Tmax) and the descent to the freezing point (Tf) are noted......................................................H18

Figure H19. Time series of bottom temperature and salinity from the COMIDA 2013 – 2014 mooring 

array.........................................................................................................................................................H20

Figure H20. Mean vectors and variance ellipses for the 2010 – 2011 moorings. Full-length records are in 

black, partial-length records are in gray, and NARR wind records are in red..........................................H21

Figure H21. Mean vectors and variance ellipses for the 2011 – 2012 moorings. Full-length records are in 

black, partial-length records are in gray, and NARR wind records are in red..........................................H22

Figure H22. Mean vectors and variance ellipses for the 2012 – 2013 moorings. Full-length records are in 

black, partial-length records are in gray, and NARR wind records are in red..........................................H23

Figure H23. Mean vectors and variance ellipses for the 2013 – 2014 moorings. Full-length records are in 

black, partial-length records are in gray, and NARR wind records are in red..........................................H24

Figure H24. Mean monthly vectors and current ellipses (black) in the topmost ADCP bin (~7 m depth) for 

the months of May, June, July, and August 2011. Comparable quantities are shown for the winds in red. 

Mooring WBC is highlighted by the orange box in each figure. Gray ellipses and vectors indicate less than 

a full month’s worth of data was available. ..............................................................................................H25

Figure H25. The first three EOF modes of the principal axis velocity component for the 2011 – 2012 

mooring array...........................................................................................................................................H26

Figure H26. The first three EOF modes of the principal axis velocity component for the 2012 – 2013 

mooring array...........................................................................................................................................H27

Figure H27. The first three EOF modes of the principal axis velocity component for the 2013 – 2014 

mooring array...........................................................................................................................................H27

Figure H28. The first three EOF modes of the principal axis velocity component for the 2011 – 2012 

mooring array, which includes Bering Strait. ...........................................................................................H29

Figure H29. The first three EOF modes of the principal axis velocity component for the 2012 – 2013 

mooring array, which includes Bering Strait. ...........................................................................................H29

Figure H30. The first three EOF modes of the principal axis velocity component for the 2013 – 2014 

mooring array, which includes Bering Strait. ...........................................................................................H30

Figure H31. Cross-spectral results between vertically averaged currents in Bering Strait and the Central 

Channel for the entire 2008 – 2014 record (left), the individual May – September periods (middle), and the 

October – April periods (right). Top panel is coherence-squared and bottom panel is the phase...........H30

Figure H32. Cross-spectral results between vertically averaged currents in Bering Strait and the Burger 

for the entire 2010 – 2013 record (left), the individual May – September periods (middle), and the October 

– April periods (right). Top panel is coherence-squared and bottom panel is the phase.........................H31

Figure H33. Cross-spectral results between vertically averaged currents in Bering Strait and Barrow 

Canyon for the entire 2011 – 2013 record (left), the individual May – September periods (middle), and the 

October – April periods (right). Top panel is coherence-squared and bottom panel is the phase...........H32
xvi

Figure H34. Climatological annual cycle based on mean monthly values of Barrow Canyon transport (top 

left) and vertically-averaged velocities projected along the principal axis at moorings over the central 

shelf: Burger (Bu), CPAI01 (previously called Site 1), and in the Central Channel at mooring Crackerjack 

(Cj). Dotted horizontal lines on the velocity figures denote the annual mean..........................................H32

Figure H35. Mean monthly transports at BC2 (blue) and BS3 (red) in Sverdrups...................................H33

Figure H36. Coherence squared and phase spectra of mean daily transports at Barrow Canyon and along 

the Beaufort shelfbreak for the entire record (left), May – September (middle) and October – April (right).

.................................................................................................................................................................H34

Figure H37. Linear regressions of BS3 against BC2 mean monthly transports for all months (upper left), 

January – April (upper right), May – September (lower left), and October –December (lower right).......H34

Figure H38. Linear regressions of BS3 against BC2 daily mean transport for all months (upper left), 

January – April (upper right), May – September (lower left), and October – December (lower right)......H35

Figure H39. Monthly distribution of flow regimes.....................................................................................H37

Figure H40. Annual cycle of the four flow regimes characterizing transports in Barrow Canyon and the 

Beaufort shelfbreak..................................................................................................................................H37

Figure H41. Frequency of occurrence of regimes as a function of the direction toward which the wind at 

Barrow (PABR) is blowing........................................................................................................................H38

Figure H42. Mean surface currents estimated from the Point Barrow and Cape Simpson HFR systesm 

under different wind conditions in 2013. The number of hours under which a given wind condition 

prevailed is indicated in the upper right hand corner of each map. .........................................................H39

Figure H43. Coherence squared and phase spectra of mean daily transports at Bering Strait and along 

the Beaufort shelfbreak for the entire record (left), May – September (middle), and October – April (right).

.................................................................................................................................................................H40

Figure H44. Tidal ellipses for the principal semi-diurnal (upper four panels) and diurnal tides on the 

Chukchi Sea shelf....................................................................................................................................H41

Figure H45. Schematic of the mean circulation of the northeastern Chukchi Sea. The red line denotes the 

Alaskan Coastal Water (ACW) front bordering the offshore side of the Alaskan Coastal Current. The blue 

line indicates the meltwater/Bering Sea Summer Water (MW/BSSW) front surrounding Hanna Shoal. 

Black arrows indicate regions where the surface and sub-surface currents are similar in direction and 

magnitude. Green arrows indicate near-bottom currents, and purple arrows indicate near-surface 

currents....................................................................................................................................................H42

Figure A-1. (a)-(h), SOM-derived patterns No. 1–8 of surface currents (blue vectors) and paired winds 

(black vectors) for 2012 data. Pattern number and its relative frequency of occurrence are shown in the 

lower right corner. The 80 m isobath is thicker to define Barrow Canyon. Note that the scales of current 

vectors in Patterns No. 3–8 are different. ...............................................................................Appendix A - 2

Figure A-1. (continued) (i)-(l), SOM-derived patterns No. 9–12 of surface currents (blue vectors) and 

paired winds (black vectors) for 2012 data. ............................................................................Appendix A - 3

Figure A-2. (a) Upper panel: time series of wind vectors in August 2012. The vector direction follows 

oceanographic convention. Lower panel: SOM-derived patterns (black dots) in August 2012 and 

normalized data returns (gray line). Pattern numbers correspond to Figure A-1. Pattern 0 denotes times 

when data return is zero. (b) As in (a), but for September 2012. (c) As in (a), but for October 2012.

................................................................................................................................................Appendix A - 4
xvii

List of Tables

Table B1. Hydrographic and velocity data sources, delineated between cruise (by row) and section (by 

semicolons). VMADCP = vessel-mounted ADCP; LADCP = lowered ADCP. ...........................................B4

Table C1. Summary of the EOF correlation analysis from the reconstructed fields. The correlations are all 

significant at the 95% significance level using the effective number of degrees of freedom derived from 

the integral time scale..............................................................................................................................C11

Table D1. Mooring positions and depths for the Barrow Canyon (BC) array and locations of NARR wind 

gridpoints. ..................................................................................................................................................D4

Table D2. Row 1 is the linear regression model used to compute Barrow Canyon transports (T) from 

mooring BC2. Rows 2 and 3 are the winter and summer generalized linear models (GLM) used to 

hindcast Barrow Canyon transport based on winds listed in Table D1 over the Bering Sea shelf (C40), 

Bering Strait (A3) and Barrow Canyon (BC2). Terms in parentheses are the 95% confidence limits on the 

model coefficients using an integral time scale of 4 days. The wind components (V, U) and lags at each 

site are defined in the footnote. .................................................................................................................D8

Table E1. Mean and Eddy Kinetic Energy comparisons for 1 m drogued Microstar drifters. ..................E28

Table E2. Mean and Eddy Kinetic Energy comparisons for 10 m drogued SVP drifters. ........................E28

Table G1. Mooring information ..................................................................................................................G5

Table G2. Statistics for the two types of extreme events considered in the text: (i) strong slope current 

and reversed shelfbreak jet; (ii) strong shelfbreak jet and weak slope current. The first column indicates 

the quadrant from which the wind was blowing. The percentage in parentheses means the fraction of the 

event length relative to the total length in the last row. The underlined percentages represent the fraction 

of total length relative to the entire year-long duration of the record. ......................................................G15

Table G3. Statistics of all upwelling events. The rows show the event number, length (in days) of each 

event, value of the upwelling index (UI, unit: kg m-3 h), mean wind direction, and sign of the mean wind 

stress curl condition (+ and – mean positive and negative curl, respectively) in the vicinity of mooring 

array (cyan box in Figure G2b) for the day before each event. ...............................................................G18

Table H1. Number of days observed and the frequency of occurrence of different flow regimes in Barrow 

Canyon (BC2) and the Beaufort shelfbreak (BS3)...................................................................................H36

Table A-1. A summary of the original SOM patterns and how they are grouped as circulation regimes in 

the main manuscript. ..............................................................................................................Appendix A - 1

Table B-1. M2 Tidal Ellipse Parameters .................................................................................Appendix B - 1

Table B-2. S2 Tidal Ellipse Parameters..................................................................................Appendix B - 2

Table B-3. N2 Tidal Ellipse Parameters..................................................................................Appendix B - 3

Table B-4. K2 Tidal Ellipse Parameters..................................................................................Appendix B - 4

Table B-5. O1 Tidal Ellipse Parameters .................................................................................Appendix B - 5

Table B-6. P1 Tidal Ellipse Parameters..................................................................................Appendix B - 6

Table B-7. K1 Tidal Ellipse Parameters..................................................................................Appendix B - 7

Table B-8. Q1 Tidal Ellipse Parameters .................................................................................Appendix B - 8
xviii

List of Abbreviations and Acronyms

ACC Alaskan Coastal Current

ACW Alaska Coastal Water

ADCP Acoustic Doppler Current Profiler

AON Arctic Observing Network

AOOS Alaska Ocean Observing System

AMSR Advanced Microwave Scanning Radiometer

AVHRR Advanced Very High Resolution Radiometer

AW Atlantic Water

BC Barrow Canyon

BOEM Bureau of Ocean Energy Management

BSEE Bureau of Safety and Environmental Enforcement

BSSW Bering Sea Summer Water

BSW Bering Sea Water

CCW Counterclockwise

CIAP Coastal Impact Assistance Program

CMP Coastal Moored Profiler

CODE Coastal Ocean Dynamics Experiment

COMIDA Chukchi Sea Offshore Monitoring in Drilling Area

CSESP Chukchi Sea Environmental Studies Program

CSW Chukchi Summer Water

CTD Conductivity-Temperature-Depth

CW Clockwise

CWP Coastal Winched Profiler

DBO Distributed Biological Observatory

DINEOF Data Interpolating Empirical Orthogonal Functions

EBC Eastern Barrow Canyon

EHSS East Hanna Shoal Shelf

EKE Eddy Kinetic Energy

EOF Empirical Orthogonal Function

ETOPO Earth Topography

GLM Generalized Linear Model

HFR High-Frequency Radar

IOOS Integrated Ocean Observing System

IPS Ice Profiling Sonar

ISCAT Inductive Sacrificial MicroCat

MKE Mean Kinetic Energy

MW Meltwater

NARR North American Regional Reanalysis

NASA National Aeronautics and Space Administration

NCEP National Centers for Environmental Prediction

NIC National Ice Center

NOAA National Oceanic and Atmospheric Administration
xix

NODC National Oceanographic Data Center

NSB North Slope Borough

NSF National Science Foundation

OI Optimal Interpolation

PABR Barrow Wiley Post-Will Rogers Memorial Airport

PC Principal Component

PI Principal Investigator

POI Point of Interest

PV Potential Vorticity

R/V Research Vessel

ROR Ratio of Overlapping Radials

RPM Remote Power Module

RWW Remnant Winter Water

SLP Sea Level Pressure

SOM Self-Organizing Map

SMMR Scanning Multichannel Microwave Radiometer

SSM/I Special Sensor Microwave/Imager

SSMIS Special Sensor Microwave Imager/Sounder

SST Sea Surface Temperature

SVP Surface Velocity Program

T/C/P Temperature/Conductivity/Pressure

T/S Temperature/Salinity

UAF University of Alaska Fairbanks

USCGC United States Coast Guard Cutter

VMADCP Vessel-Mounted Acoustic Doppler Current Profiler

WBC Western Barrow Canyon

WHOI Woods Hole Oceanographic Institution

WW Winter Water
xx

Abstract

This program utilized a wide variety of observations obtained from moored oceanographic instruments, 

satellite-tracked drifters, hydrography (both recent, from towed-CTD and autonomous underwater 

vehicles, and historic, from ship measurements), high-frequency radars (HFR), meteorological buoys, and 

satellite imagery to develop an improved understanding of the physical oceanography of the northeastern 

Chukchi shelf, as well as exchanges between the Chukchi/Beaufort shelves and the adjacent basin. Our 

measurements spanned the shelf region from Point Hope northward to the Chukchi and Beaufort 

continental slopes and from the Alaskan coast to the US-Russian Convention Line. Our primary focus 

area was the shelf and shelfbreak region between the Central Channel and the western Beaufort Sea and 

north to the shelfbreak. 

Throughout the year and over most of the shelf south of 71.5°N, including within the Central Channel and 

along the northwest side of Hanna Shoal (as far north as the 50 m isobath), the flow is weakly sheared 

vertically, such that the surface and sub-surface flows coincide with respect to flow direction, if not 

magnitude. Exceptions to this may involve large shears between drifting ice and the sub-surface flow 

and/or large shears during the melt season. During melt, the near-surface stratification is very strong 

(similar to the stratification found on the shelf east of Hanna Shoal), such that the sub-surface and surface 

circulations are uncoupled from one another. This is the situation encountered on the East Hanna Shoal 

Shelf (EHSS) where vertical shears are large year-round. Other regions of strong vertical shear include 

the Beaufort and Chukchi shelfbreaks and slopes and portions of Barrow Canyon in summer and fall. 

We have examined the strength, persistence, and variability of the model-predicted (Winsor and Chapman 

2004; Spall 2007) clockwise circulation around Hanna Shoal, including southwestward flow over the 

shelf between the Barrow Canyon and Hanna Shoal. Substantial elements of the models are largely 

correct, but there are important differences between the models and observations. We found that the 

component of the clockwise circulation on the northwest side of Hanna Shoal extends through the water 

column as predicted by the models and includes cold Bering Sea Water (BSW) and Winter Water (WW) 

along the bottom. Both of these water masses are transported from the Central Channel and Herald 

Canyon. However, on the EHSS, only the bottom waters, consisting of BSW and WW, engage in the 

clockwise circulation. The upper layer flows westward and/or northwestward over the EHSS. This flow is

a consequence of density gradients that extend northeastward and eastward from Hanna Shoal, which

drive a counterclockwise baroclinic, geostrophic circulation in the upper layer that transports meltwater

(MW) and/or modified Alaskan Coastal Water. The latter is transported onto the EHSS via Ekman 

transport from Barrow Canyon and/or by onshelf transport from the Chukchi shelfbreak. The different 

flow regimes on either side of Hanna Shoal have two effects. First, these result in a zonal exchange of 

water masses that maintains the strong stratification year-round on the EHSS. Second, we infer that the 

shelf north of Hanna Shoal is a region of zonal flow convergence wherein waters must be transported 

toward the shelfbreak. This transport, which certainly involves the surface layers and possibly also bottom 

waters, should affect the magnitude and structure of the flow over the Chukchi shelfbreak.

Models also suggest that the clockwise flow around Hanna Shoal extends south of Hanna Shoal in the 

form of a southwestward drift that eventually turns eastward across the central shelf and continues toward 

the head of Barrow Canyon. The observations indicate that this is not the case. Indeed, we found that 

there was no net transport between the EHSS and the central shelf. Instead there is a southeastward flow 

along the south side of Hanna Shoal that transports water (usually WW) toward the head of Barrow 

Canyon. This flow essentially represents a dynamic boundary that blocks communication between the 

EHSS and the central shelf. As a consequence of this boundary and the weak mean flow on the northern 

side of the Shoal, the residence time of bottom waters on the shelf east of the Shoal is ~9 months, much 

longer than elsewhere in the Chukchi Sea. 
xxi

As noted, the inferred zonal convergence north of the Shoal provides a possible mechanism for 

shelf/slope exchange in which MW, WW, and BSW are transported offshelf in this area. There are 

however other mechanisms that foster exchange. Our retrospective analysis of hydrographic data from the 

Chukchi shelfbreak indicates that this consists of a surface flow that carries Chukchi Sea water masses 

(the Pacific-derived summer waters, WW, and MW) westward and an eastward sub-surface flow that is 

confined to the shelfbreak and which transports WW eastward. The waters flowing westward are largely 

derived from the Barrow Canyon outflow. The WW that comprises the eastward subsurface flow are 

mainly derived from northward flows through the Central Channel and Herald Canyon, although some 

probably form in winter around the periphery of Hanna Shoal. The potential vorticity structure of the 

shelfbreak circulation satisfies the conditions for baroclinic instability, which should lead to eddy-driven 

exchanges between the shelfbreak and basin as is known to occur frequently along the Beaufort Sea 

shelfbreak.

Our results suggest that most of the Chukchi shelf water masses that enter the head of Barrow Canyon are 

drawn from the shelf south of Hanna Shoal. Of the 0.2 Sv average transport we estimate to be flowing 

into the head of the canyon, at least ~0.1 Sv flows eastward across the central shelf from the Central 

Channel. The remainder presumably flows northeastward along the coast. Although we are unable to 

estimate the transports involved, we infer that additional waters enter the canyon along its length 

(downcanyon from the head) due to spillage of bottom waters from the EHSS or from westward-flowing 

waters from the Beaufort Sea. There also appears to be, on average, an upcanyon flow along the west side 

of the canyon. This flow probably transports waters from the subsurface shelfbreak current flowing 

eastward along the Chukchi slope. Most likely this inflow proceeds only part way up the canyon because 

it continuously loses mass to the downcanyon transport along the east side of the canyon. WW comprise 

the bulk of the water masses spilling into the canyon and/or transported upcanyon from the shelfbreak. 

The circulation and transports in the Central Channel, across the central shelf, and in Barrow Canyon vary 

annually and are (mostly) in-phase with the annual transport cycle in Bering Strait. For the 2010 – 2015 

period, we found that the downcanyon transport is large (0.4 – 0.6 Sv) from May through August, weak 

(≤0.2 Sv) from September through November and from February through April, and reversed (e.g., 

upcanyon at ~0.1 Sv) in December and January. The annual cycle over the central shelf is in-phase with 

that in Barrow Canyon and includes a reversed (westward) flow in December and January. In contrast, the 

flow in the Central Channel is northward in all months. These results suggest that in early winter the 

circulation over the northeastern Chukchi Sea shelf undergoes a major reorganization: waters transported 

upcanyon may extend onto the central shelf, and nearshore waters are transported westward into the 

Central Channel and even the western Chukchi Sea. Once in the Central Channel, these nearshore waters 

would be transported northward. This reorganization provides a means by which WW, formed in the 

latent heat polynyas along the northwest coast of Alaska in December and January, can be carried far 

offshore and eventually to Hanna Shoal.

This seasonal reorganization of the flow was likely reflected during synoptic coastal upwelling events 

such as the one captured by satellite-tracked drifters August 2014. During that event, the cross-shelf 

transport of surface coastal waters was a response to various competing mechanisms including the 

geostrophic transport associated with the mean Pacific-Arctic pressure gradient and, in response to the 

winds, surface Ekman drift, shelf-wide pressure gradient adjustments, and Stokes’ drift, which although 

small, is not negligible. The drifter observations showed that surface nearshore waters can be transported 

from the coast to either the Central Channel and/or the head of Herald Canyon in less than a month. 

Similarly, self-organized mapping analytics applied to the HFR observations clearly indicate that waters 

transported up Barrow Canyon can easily be carried westward onto the central shelf and/or southward 

along the Alaskan coast. In fall, the coastal upwelling events and Barrow Canyon reversals likely lead to 

cooling of nearshore waters and thus enhance ice formation. 
xxii

The shelf circulation also includes a complex variety of mesoscale circulations. One example (observed 

about 10% of the time) is the “divergent” mode, which occurs during transitions from northeasterly to 

southeasterly winds. In such events, a pair of counter-rotating vortices form in which the eastward 

onshore flow from the central shelf splits upon approaching the coast; some water enters a cyclonic vortex 

to the north of Wainwright and some enters an anticyclonic vortex offshore of Point Lay. The water 

entering the anticyclonic vortex recirculates over the central shelf. Two other examples are the inferred 

shedding or generation of eddies as the Alaskan Coastal Current rounds Point Hope and the collision of 

the westward-flowing Beaufort shelf waters with the along-canyon flow in Barrow Canyon. There are 

also numerous MW fronts that occur on this shelf, which we suggest are of two types (at least). The first 

type consists of what appears to be numerous small sub-mesoscale (≤6 km) fronts that are shallow (10 –

15 m thick) and found over Hanna Shoal and the EHSS. There is another, more prominent, type of 

mesoscale front, which in August and September extends from the western end of Hanna Shoal eastward 

to the head of Barrow Canyon along the south side of the Shoal and then northward along the western half 

of Barrow Canyon. These fronts separate moderately-stratified Bering Sea Summer Waters to the south 

(and east in Barrow Canyon) from the strongly-stratified, two-layered MW/WW structure to the north. 

The fronts are ~15 km wide, baroclinically unstable, and generate both anticyclonic and cyclonic eddies 

at the surface and anticyclonic eddies primarily within the pycnocline north of the front.

These mesoscale and sub-mesoscale motions play an important role in the lateral and vertical mixing of 

water masses and in the dispersion of dissolved and suspended materials. To a large degree, the mesoscale 

phenomena that we have identified on the shelf are linked to the distribution of MW. Over the 

northeastern shelf, the summer MW pool is tied to ice conditions over Hanna Shoal, which can vary 

considerably from year-to-year. We suggest that the rate of ice retreat over the Shoal in summer depends 

less on summer winds and the shelf circulation than it does on winter conditions that affect the grounding 

of thick ice on the Shoal.
A1

A. Introduction, Objectives, and Partnerships

A.1 Introduction

This report provides a summary and synthesis of data collected over the shelf and shelfbreak of the 

northeastern Chukchi and western Beaufort seas from 2012 through 2014. The study was an expansion of 

an earlier effort, “Application of High Frequency Radar to Potential Hydrocarbon Development Areas in 

the Northeast Chukchi Sea” that was jointly supported by BOEM, ConocoPhillips, and Shell between 

2009 and 2011. As described below, the present study was also made possible through a variety of 

partnerships. The field components of both this and the earlier study included gliders (autonomous 

underwater vehicles), high frequency radars (HFRs), shipboard measurements of the hydrographic and 

velocity fields, satellite imagery, satellite-tracked drifters, and year-round moorings.

The Chukchi and Beaufort seas are the northernmost shelves bordering Alaska. Although properly a part 

of the western Arctic Ocean, the Chukchi shelf is atmospherically and oceanographically linked to the 

Pacific Ocean. These connections influence the wind and wave regimes, ice distribution, and the water 

masses and circulation characteristics of both seas. The atmospheric connection is primarily via two 

centers of action: the Aleutian Low and Beaufort High. Both of these features vary in time, position, 

strength, and interact with polar air masses that affect regional meteorological conditions. The 

oceanographic link is via the mean northward flow through Bering Strait, which is sustained by the largescale pressure gradient between the Pacific and Atlantic Oceans (Coachman et al. 1975; Aagaard et al. 

2006).

The northward transport of mass, heat, nutrients, carbon, and organisms through the Strait bequeath this 

region with biophysical characteristics unique among arctic shelves. For example, in comparison to most 

other arctic shelves, the spring retreat of ice occurs earlier and fall ice formation is delayed in the Chukchi 

Sea because of the northward heat flux through the strait. Woodgate et al. (2006). Shimada et al. (2006), 

and Steele et al. (2010) contend that this heat flux may be an important source of interannual variability in 

the ice cover over the western Beaufort Sea. Similarly, the enormous biological productivity of the region 

(Walsh et al. 1989; Grebmeier and McRoy 1989; Springer and McRoy 1993; Arrigo and van Dijken 

2015), including its ability to support large and diverse marine mammal populations (Citta et al. 2015), is 

attributed to the carbon and nutrient loads carried through Bering Strait.

The shallow (~50 m) Chukchi shelf extends ~800 km northward from Bering Strait to the shelfbreak 

along the 200 m isobath. Figure A1a (from Spall 2007) is a model-based depiction of the mean circulation 

over the shelf and suggests that the northward flow through Bering Strait proceeds along three principal 

pathways (Figure A1b) with each pathway associated with a distinct bathymetric feature: Herald Canyon, 

the Central Channel, and Barrow Canyon. The troughs are separated from one another by shoals. Herald 

Shoal separates Herald Canyon from the Central Channel and Hanna Shoal lies between Barrow Canyon 

and the Central Channel.

The westernmost branch enters Herald Canyon (Coachman et al. 1975; Walsh et al. 1989; Hansell et al. 

1993; Pickart et al. 2010) and while some of it appears to spread eastward across the northern shelf, a 

significant portion of the water exiting Herald Canyon forms an eastward-flowing shelfbreak jet along the 

edge of the Chukchi Sea (e.g., Mathis et al. 2007; Pickart et al. 2010). Existing observations, although 

limited in time and space, suggest that the flow passing through the Central Channel is steered eastward 

towards Hanna Shoal (Johnson 1989; Münchow and Carmack 1997; Gong and Pickart 2015). Upon 

reaching the Shoal a portion of the water appears to flow anti-cyclonically around the Shoal, consistent 
A2

Figure A1. Mean depth-integrated streamlines (white lines) in the Chukchi Sea, after Spall (2007; left). 

The bathymetry is colored and major bathymetric features are labeled. Bathymetric map of the 

northeastern Chukchi Sea, with place names included (right).

with the numerical models of Winsor and Chapman (2004) and Spall (2007). Recent evidence suggests, 

however, that part of the water passes to the south of the Shoal (Pickart et al. 2016). Ultimately all of the 

water is believed to enter the head of Barrow Canyon. Upon exiting the canyon – in the absence of winds 

– some of the water turns eastward again and forms a shelfbreak jet along the edge of the Beaufort Sea 

(Nikolopoulos et al. 2009; von Appen and Pickart 2012). New evidence from the present study indicates, 

however, that a significant portion of the water exiting Barrow Canyon turns westward and forms a jet 

over the Chukchi continental slope (see below). In the presence of easterly winds, models suggest that 

much of the water entering the Pacific sector of the Arctic Ocean ultimately ends up exiting the Chukchi 

shelf via Herald Canyon. 

According to the models and observations a portion of the water that enters the Central Channel flows 

eastward toward the Alaskan coast across the central shelf, i.e., before reaching the latitude of Hanna 

Shoal (Paquette and Bourke 1974; Weingartner et al. 2005). Limited hydrography led Weingartner et al. 

(2013a) to suggest that, in summer, some of the water flowing northward in the Central Channel spreads 

eastward along its length as far north as the southwest side of Hanna Shoal. As these waters flow 

eastward they encounter dense (cold, salty) bottom waters along the south and southeast sides of Hanna 

Shoal. These dense waters were formed the previous winter either locally or elsewhere on the Chukchi 

shelf or in the northern Bering Sea. Model results (Figure A1) suggest that the source of this dense water 

is from the east side of Hanna Shoal, but Pickart et al. (2016) suggest that some of this dense water is 

transported eastward along the south side of Hanna Shoal. 

At this point, several questions arise with respect to a number of the inferences drawn from the models 

and the limited observations pertaining to the shelf circulation in the vicinity of Hanna Shoal. For 

example, to what extent is the eastward flow partitioned between the shelf and shelfbreak? What are the 

downstream (e.g., eastward) fates of these flows; does all of the water enter Barrow Canyon, and, if so, 

where and how is this accomplished? How do the magnitudes and the horizontal and vertical structure of 

these flows vary in time? And finally, how do they vary with respect to winds and circulation elsewhere 

on the shelf?

The third branch of Bering Sea water flows northeastward along the Alaskan coast towards Barrow 

Canyon at the junction of the Chukchi and Beaufort shelves (Mountain et al. 1976; Paquette and Bourke 
A3

1981). In summer, this flow includes the northward extension of the Alaskan Coastal Current (ACC) that 

originates south of Bering Strait (Aagaard et al. 1985; Aagaard 1988; Münchow et al. 2000). At the head 

of Barrow Canyon, the ACC is joined by waters flowing eastward from the central shelf and, as suggested 

by the model streamlines, with water flowing southward from the east side of Hanna Shoal. The merged 

flow then continues downcanyon as a narrow, but strong, coastal jet (Aagaard and Roach 1990; Pickart et 

al. 2005). Hence in summer and fall, the canyon outflow contains a horizontally- and vertically-structured 

complex of water masses (Pickart et al. 2005; Shroyer and Plueddemann 2012) that includes warm, dilute 

Alaskan Coastal Water (carried by the ACC); cold, dilute waters due to melting sea ice; dense winter 

waters; relatively warm and salty Bering Sea Water, and mixtures of each. Although the importance of 

Barrow Canyon as a potential drain for a broad portion of the Chukchi Sea shelf has long been 

recognized, there were no comprehensive efforts to measure the transport that enters the head of the 

canyon. 

The flow field described above is substantially influenced by winds, which on average blow from the 

east-northeast. Consequently the mean flow tends to oppose the mean winds. Earlier work (Weingartner 

et al. 2005) suggested that flow variations tended to be coherent with wind variations on synoptic time 

scales. Those observations, however, were based on a limited number of moorings, with most 

measurements made near the seabed. Thus, there was concern that the measurements may not reflect the 

near-surface currents. We expect there may be a thin, wind-forced layer across which the currents are 

sheared. Presumably, the thickness of this layer (and the dynamics that govern the motion in this layer) 

varies spatially over the shelf and throughout the year with seasonal changes in stratification, ice cover, 

wave climate, and wind velocity. This last point is crucial in considering the ultimate development of oil 

spill trajectory models for the Chukchi shelf. Models that incorporate only wind-forcing to compute 

surface drift could produce serious discrepancies, where the surface currents arise in response to forcing 

by both the winds and the large-scale (and opposing) pressure field.

A.2 Objectives

The overall goal of this program, as defined in the original BOEM statement of work, is to provide an 

improved understanding of the physical oceanography of the northeastern Chukchi shelf and exchanges 

between the Chukchi/Beaufort shelves and the adjacent basin. The questions outlined above guided the 

development of several specific objectives designed to meet the overall goal of the study. These include:

1. Conditions, locations, and frequency under which the surface (upper 1 m) circulation does/does not 

coincide with the sub-surface flow as measured by moored current meters;

2. Strength, persistence, and variability of the model-predicted clockwise circulation around Hanna 

Shoal, including the southwestward flow over the shelf between the Barrow Canyon and Hanna 

Shoal (Winsor and Chapman 2004; Spall 2007);

3. Conditions that govern, and the frequency of occurrence of, the along-shore and divergent modes 

that occur in the Wainwright-Point Lay region;

4. Mass transport contributions of the flow around Hanna Shoal and the eastward (shoreward) flow 

from the central Chukchi Sea shelf to the Barrow Canyon transport;

5. Mesoscale dynamics driving the shelf-basin exchange of heat, salt, momentum, and nitrate, and 

quantify the magnitude of the exchange;

6. Spatial and temporal variability in the shelf-basin exchange processes (including eddies) and isolate

the causes of this variability in terms of local wind forcing, boundary current stability, and/or shelf 

wave phenomena.

7. Connections between the circulation on the northeast Chukchi shelf, the flow of water into/out of 

Barrow Canyon, and the shelf-basin exchange processes on either side of the canyon.
A4

8. Local momentum balances, and, where and when possible, how the pressure field adjusts for given 

wind conditions and/or mesoscale flows.

In aggregate, the results can be used to guide the development and evaluation of oil spill trajectory models 

likely to be used for this region and can provide insights on processes relevant to the functioning of this 

marine ecosystem.

The project included three field years (2012 – 2014) and involved measurements along the Chukchi and 

Beaufort shelfbreaks, around Hanna Shoal, and over the northeast Chukchi Sea including Barrow Canyon.

The field program involved high-frequency, shore-based, surface current mapping radars (HFR), 

autonomous underwater vehicles (gliders), satellite-tracked drifters, year-round moored oceanographic 

instruments, and shipboard measurements of hydrography (standard Conductivity-Temperature-Depth 

(CTD) profiles and/or towed CTD packages) and, in some cases, the current structure based on vesselmounted acoustic Doppler current profilers (VMADCP).

A.3 Partnerships

A unique aspect of this program is that it involved a very extensive set of partnerships with industry, other 

agencies, and our own academic institutions. These partnerships included both direct and/or in-kind 

support. The following partners were involved in this program:

1. The State of Alaska’s Coastal Impact Assistance Program (CIAP) provided support for a HFR site 

at Cape Simpson in the western Beaufort Sea. This support enabled purchase of an HFR system, a 

remote power module (RPM) (Statscewich et al. 2011), and the logistics to operate the site for the

measurement years of 2013 – 2014. The Cape Simpson site was paired with a similar HFR-RPM 

installation at Point Barrow to enable HFR measurements over the western Beaufort Sea and the 

mouth of Barrow Canyon.

2. The autonomous underwater vehicle glider program was augmented with a Teledyne Webb glider 

purchased with AOOS funding. This enabled us to carry out additional glider surveys and collect 

extensive sections over the ice-free season in the northeastern Chukchi Sea.

3. Pickart maintained a subsurface mooring in the western Beaufort Sea, as part of the Arctic 

Observing Network (AON) funded by the National Science Foundation (NSF). The mooring was in 

place during the BOEM fieldwork and the data made available to the project. In addition, shipboard 

CTD/ADCP surveys were carried out in October 2013 (13 transects) and July 2014 (10 transects) on 

the USCGC Healy. This included occupations along the Chukchi slope mooring line as well as 

transects to the west and east (extending into the Beaufort Sea). 

4. Japanese colleagues, funded by the Japan Agency for Marine Earth Science and Technology

maintained moorings at the mouth of Barrow Canyon during the study period. 

5. The BOEM-funded Chukchi Sea Offshore Monitoring in Drilling Area (COMIDA) Hanna Shoal 

Ecosystem Study commenced in August 2012 and concluded fieldwork in September 2014. This 

was an interdisciplinary program that included hydrographic surveys over and around Hanna Shoal 

and the deployment of 6 oceanographic, subsurface moorings per year northwest and northeast of 

the Shoal.

6. The industry-sponsored Chukchi Sea Environmental Studies Program (CSESP) provided extensive 

hydrographic data from 2008 – 2014 over the northeastern Chukchi shelf. These data are used in the 

synthesis component of this project.

7. The Alaska Ocean Observing System (AOOS) as a part of the NOAA-supported Integrated Ocean 

Observing System (IOOS) program has provided support for the operations and maintenance of the 

HFR network involved in this program.
A5

8. NSF supported S. Okkonen (UAF) and C. Ashjian (WHOI) in a biophysical process study relevant 

to marine mammal foraging in the Point Barrow region. As part of this effort Okkonen deployed 

two oceanographic moorings midway along the western and eastern walls of Barrow Canyon in 

2011 – 2012. Drs. Okkonen and Ashjian have shared these data with us.

9. Industry (Conoco-Phillips, Shell, and Statoil) have recently provided their ADCP and IPS (iceprofiling sonar) data from oceanographic moorings deployed in the Chukchi Sea since 2008.

10. This BOEM program has supported Lagrangian surface current measurements using satellitetracked drifters. This sampling effort was substantially augmented from 2012 – 2014 by the North 

Slope Borough (NSB)/Shell Baseline Studies Program. This program supports research efforts that 

address concerns raised by NSB residents. In collaboration with residents from the communities of 

Point Hope, Point Lay, Wainwright, Barrow, and Kaktovik, the NSB executed an extensive set of 

near-shore drifter deployments in collaboration with this BOEM study.

11. Pickart was a PI in the NASA-supported program “Impacts of Climate on EcoSystems and 

Chemistry of the Arctic Pacific Environment” (ICESCAPE). This entailed two extensive 

physical/biological summertime shipboard surveys of the Chukchi Sea in 2010 and 2011, including 

vessel-mounted ADCP data. These data were provided to the synthesis component of this report.

12. Pickart is a Co-PI in the NSF-funded program “Study of Under Ice Blooms in the Chukchi 

Ecosystem” (SUBICE). This included an extensive shipboard hydrographic survey on Healy in May 

to June 2014, from which the data were made available to the project. In addition, a mooring was 

deployed from Oct 2013 to July 2014 north of the Central Channel to augment the BOEM array. 

The data from the mooring were provided to the project.

13. Pickart is a Co-PI of the Russian-American Long Term Census of the Arctic (RUSALCA) program, 

funded by NOAA. In summer 2012 a physical/biological survey was carried out in the western 

Chukchi Sea in Russian waters aboard the Professor Khromov. This also included direct velocity 

measurements from the ship using a lowered ADCP, as well as a detailed survey of Herald Canyon. 

These data are available to the synthesis component of this report.

14. The University of Alaska School of Fisheries and Ocean Sciences provided support for the 

continuation of oceanographic measurements in Barrow Canyon from August 2014 – August 2015. 

This extra year has permitted the establishment of a 5-year transport time series in Barrow Canyon.

15. The Woods Hole Oceanographic Institution provided funding for four additional moorings to 

augment the BOEM-funded moorings on the Chukchi slope. These were in place from October 

2013 to September 2014, and consisted of profiling CTD measurements and ADCP measurements. 

The shiptime to deploy and recover these moorings (and the BOEM moorings) was provided by 

NSF. The additional mooring data will be available to the synthesis component of this report.

16. Pickart is a Co-PI of the Distributed Biological Observatory (DBO) project funded by NSF as part 

of AON. This is an ongoing program that includes a collection of opportunistic measurements at 

critical locations in the Chukchi Sea (depending on ship schedules and availability). The DBO data 

are available to the synthesis component of this report.

17. Each HFR field site was made possible due to partnerships with several villages (Point Lay, 

Wainwright, and Barrow) as well as the Cully, Olgoonik, and Ukpeagvik Iñupiat Corporations, who 

allowed use of their lands and often provided logistic support.

18. Winsor is the Co-PI of the BOEM sponsored program "Arctic Tracer Release Experiment 

(ARCTREX): Applications for Mapping Spilled Oil in Arctic Waters" conducted Autonomous 

Underwater Vehicle (AUV) and Acrobat surveys in the northeastern Chukchi Sea.

19. Industry provided shiptime support for the summer deployment and fall recovery of the 

meteorological buoy and shared the data from their additional buoys in the Chukchi and Beaufort 

Seas.
A6

20. Winsor is a Co-PI on the AOOS funded project “Autonomous real-time detection of marine 

mammals in the Arctic” which utilizes glider technologies developed within the project reported on 

here. We used data from those surveys in this report.

The report proceeds through a number of sections which have been and/or are being prepared for 

submission in peer-reviewed journals. Section B consists of a retrospective analysis of the hydrographic 

and velocity structure of the shelfbreak in the northeastern Chukchi Sea published in Progress in 

Oceanography. Section C outlines the main findings from the HFR measurements, in press with the 

Journal of Geophysical Research-Oceans. Section D examines transport variability at the head of Barrow 

Canyon based on data obtained from five years of oceanographic moorings, published in the Journal of 

Geophysical Research-Oceans. Section E augments previous analyses (Weingartner et al. 2015) of the 

surface circulation as captured by satellite-tracked drifters. Section F provides a description of the 

hydrography of the Hanna Shoal region. Section G addresses the Chukchi Sea shelbreak and slope, and 

Section H consists of a preliminary synthesis of these and other data sets from the Chukchi Sea. (Please 

note that the circulation maps given in each section vary slightly from each other to better address their

subject matter, with the overall average circulation derived from measurements in this study given in 

Section H.) The last two sections consist of a suite of recommended future studies and references, 

respectively.
From Corlett and Pickart 2017

B1

B. The Chukchi Slope Current

Abstract

Using a collection of 46 shipboard hydrographic/velocity transects occupied across the shelfbreak and 

slope of the Chukchi Sea between 2002 – 2014, we have quantified the existence of a current transporting 

Pacific-origin water westward over the upper continental slope. It has been named the Chukchi slope 

current, which is believed to emanate from Barrow Canyon. The current is surface-intensified, order 50 

km wide, and advects both summer and winter waters. It is not trapped to a particular isobath, but instead 

is reminiscent of a free jet. There is no significant variation in Pacific water transport with distance from 

Barrow Canyon. A potential vorticity analysis suggests that the flow is baroclinically unstable, consistent 

with the notion that it meanders. The current is present during all synoptic wind conditions, but increases 

in strength from summer to fall presumably due to the seasonal enhancement of the easterly winds in the 

region. Its transport increased over the 12-year period of data coverage, also likely in response to wind 

forcing. In the mean, the slope current transports 0.50 ± 0.07 Sv of Pacific water. This estimate allows us 

to construct a balanced mass budget of the Chukchi shelf inflows and outflows. Our study also confirms 

the existence of an eastward-flowing Chukchi shelfbreak jet transporting 0.10 ± 0.03 Sv of Pacific water

towards Barrow Canyon.

B.1 Introduction

The means by which Pacific water crosses the Chukchi Sea, and the locations where the water exits the 

shelf into the interior Arctic Ocean, directly impacts various aspects of the Arctic ecosystem. The three 

main pathways of Pacific water on the shelf are the western branch into Herald Canyon, the Central 

Channel branch which flows northward between Herald and Hanna Shoals, and the coastal pathway 

(known as the Alaskan coastal current in summer and fall, Figure B1). During winter and spring, cold 

Pacific water flows through Bering Strait into the Chukchi Sea (Woodgate et al. 2005). This water is high 

in nutrients (Lowry et al. 2015), and consequently helps spur primary production on the shelf (Hill et al. 

2005; Arrigo et al. 2014). In late spring and summer, warmer and fresher Pacific waters enter the Chukchi 

Sea, which are believed to play a significant role in both melting and delaying the formation of pack-ice 

both on the shelf (e.g., Weingartner et al. 2005) and in the basin (Steele et al. 2010; Woodgate et al. 2012; 

Brugler et al. 2014). The Pacific water transported across the shelf is also believed to contribute 

significantly to the reservoir of freshwater offshore in the Beaufort Gyre (e.g., Pickart et al. 2013).

The coldest type of Pacific water is known as newly-ventilated winter water (WW), which is near the 

freezing point. This originates from the northern Bering Sea (e.g., Muench et al. 1988) but can also be 

formed, or further transformed, locally on the Chukchi shelf in polynyas and leads (e.g., Weingartner et 

al. 1998; Itoh et al. 2012; Pickart et al. 2016; Pacini et al. submitted). As the season progresses, the 

temperature of this water moderates via solar heating and/or mixing with warmer ambient waters, at 

which point it is referred to as remnant winter water (RWW). (In the case of extreme warming, the WW 

can be converted to a weakly stratified summer water mass; Gong and Pickart 2016.) During summer and 

early fall, the Chukchi Sea contains two different types of warm Pacific water masses. The first is 

Alaskan coastal water (ACW) which stems largely from fluvial runoff in the Gulf of Alaska. The second 

is a combination of Anadyr water and central Bering shelf water, which mix north of Bering Strait 

(Coachman et al. 1975) to form a water mass known as Bering Sea summer water (BSSW).

Recent studies have revised our understanding of the circulation of Pacific-origin water on the Chukchi 

shelf, including the partitioning of transport between the different flow pathways. While the notion of 
From Corlett and Pickart 2017

B2

Figure B1. Revised schematic of the circulation of the Chukchi Sea and western Beaufort Sea from 

Brugler et al. (2014), including an extended Chukchi shelfbreak jet and the newly described Chukchi 

slope current based on the results of this study.

three main branches remains intact, it is now believed that, as the Pacific water progresses across the 

shelf, it divides into a number of smaller branches or filaments on the northeast part of the shelf (Pickart 

et al. 2016). Among other things, this impacts the timing of the advection of the high-nutrient WW across 

the shelf, which in turn has ramifications for the primary production (Lowry et al. 2015). With regard to 

transport, the yearly averaged volume flux in each of the three main flow branches is thought to be 

comparable (Woodgate et al. 2005). However, recent data suggest that, at least during the summer 

months, much of the Pacific water entering Bering Strait drains into Barrow Canyon in the northeast part 

of the shelf (Itoh et al. 2013; Gong and Pickart 2016; Pickart et al. 2016).

Presently, the mechanisms by which Pacific water exits the Chukchi shelf into the Canada basin — and 

the geographical locations where this occurs — are not fully understood. It is known that some portion of 

the outflowing Pacific water ends up as a shelfbreak jet that, in the mean, flows eastward along the edge 

of the Beaufort Sea (Nikolopoulos et al. 2009). While the configuration of this jet changes seasonally 

(surface-intensified during late-summer/early-fall, bottom-intensified over the remainder of the year), the 

current is a year-round feature. There is also evidence of a shelfbreak jet along the edge of the Chukchi 

Sea that, in the absence of wind forcing, flows to the east (Pickart et al. 2005; Mathis et al. 2007; Llinás et 

al. 2009; Pickart et al. 2016). The source of this is believed to be the outflow from Herald Canyon 
From Corlett and Pickart 2017

B3

(Pickart et al. 2010). However, the data are largely anecdotal, and there are no published estimates of the 

transport of this shelfbreak flow.

Notably, the volume transport of Pacific water in the Beaufort shelfbreak jet is only a small fraction of 

what enters the Chukchi Sea through Bering Strait. Using data from seven moorings deployed across the 

current from 2002 – 2003, Nikolopoulos et al. (2009) calculated the mean volume flux of Pacific water to 

be 0.13 ± 0.08 Sv, which is only about 15% of the long-term transport through Bering Strait (0.83 Sv;

Roach et al. 1995). Furthermore, while the northward volume flux through Bering Strait has increased in 

recent years to just over 1 Sv (Woodgate et al. 2012), the eastward transport of Pacific water in the 

Beaufort shelfbreak jet has decreased to 0.021 – 0.041 Sv (Brugler et al. 2014). This implies that the jet 

now only accounts for less than 5% of the Pacific water that enters the Chukchi Sea. (The summertime 

transport of the Beaufort shelfbreak jet increases to approximately 0.25 Sv (Brugler et al. 2014), still far 

less than the transport through Bering Strait.)

This begs the question, where and how does the bulk of the Pacific water exit the Chukchi Shelf? As 

noted above, Woodgate et al. (2005) argued that each of the three main branches transports a similar 

amount of Pacific water. However, a significant portion of the water in the western branch is diverted to 

the east just north of Herald Shoal (Pickart et al. 2010), progressing along the northern Chukchi shelf and 

joining the central branch (Spall 2007; Pickart et al. 2016) (Figure B1). This combined central/western 

branch is then believed to flow into Barrow Canyon, adding to the puzzle regarding the small transport of 

the Beaufort shelfbreak jet. Brugler et al. (2014) attributed the recent decrease in strength of the Beaufort 

shelfbreak jet to enhanced easterly winds. This is consistent with the larger number of wind-driven 

upwelling events over the last decade noted by Pickart et al. (2013), during which the flow of the jet is 

reversed to the west (Schulze and Pickart 2012). Brugler et al. (2014) presented data from a shipboard 

section occupied across the Chukchi slope just west of Barrow Canyon, which, together with a series of 

satellite images, revealed that the outflow from the canyon veered westward during such an event. This 

may provide a clue as to the fate of the Pacific water exiting the canyon and hence the “missing” transport 

in the Beaufort shelfbreak jet, but it should be remembered that this was only a synoptic occurrence.

Unlike the Beaufort Sea, where the shelfbreak jet is robustly documented using both mooring and 

shipboard data, the circulation along the Chukchi shelfbreak and slope has yet to be quantified. In light of 

the unaccounted transport leaving Barrow Canyon, and the observation of a jet of water veering to the 

west out of the canyon during a wind event, this motivates a study of the flow along the outer edge of the 

Chukchi Sea. Here we use a collection of historical shipboard sections occupied across the Chukchi 

shelfbreak/slope to investigate the circulation during the warm months of the year. The data reveal the 

existence of a heretofore unquantified current over the Chukchi slope that advects a significant amount of 

Pacific-origin water westward. The data also robustly document the presence of an eastward-flowing 

shelfbreak jet transporting a small amount of Pacific water toward Barrow Canyon. The paper is 

organized as follows. We start with a presentation of the shipboard data and an explanation of the 

technique used to create a mean section. We then describe the mean characteristics of the Chukchi slope 

current and shelfbreak jet, including the water masses they advect. This is followed by an investigation of 

the seasonal and interannual variation of the slope current. Next we explore the nature of the slope current 

and some aspects of its dynamics. Finally, we discuss the ramifications of these new components of the 

circulation in light of the mass budget of the Chukchi shelf.

B.2 Data

B.2.1 Shipboard Hydrographic and Velocity Data

We have compiled all publicly available shipboard sections occupied across the Chukchi shelfbreak/slope 

with the criteria that (1) they have high spatial resolution (typical station spacing of approximately 10 
From Corlett and Pickart 2017

B4

km), and (2) direct velocity measurements were made in addition to the hydrographic measurements. This 

resulted in 46 sections carried out during the period 2002 – 2014 (Table B1), spanning geographically 

from west of Barrow Canyon to 168°W (Figure B2). All but two of the cruises were done on the USCGC 

Healy. The exceptions were the July – August 2002 cruise on the USCGC Polar Star and the July –

August 2003 cruise on the R/V Nathaniel B. Palmer. In each case a Sea-Bird 911+ conductivitytemperature-depth (CTD) instrument was used, with dual T/C sensors that were calibrated pre- and postcruise. Most of the velocity measurements were made with a hull-mounted acoustic Doppler current 

profiler (ADCP), except the Polar Star cruise which used a lowered ADCP.

The data extend seasonally from May through October, although all of the May and June occupations 

occurred during 2002 – 2004. Therefore, the seasonal analysis is restricted to the months of July through 

October. In addition, there were no sections occupied between 2005 and 2008, so for the interannual 

analysis we consider two time periods: 2002 – 2004, which is referred to as the early regime, and 2009 –

2014, which is referred to as the recent regime. All of the cruises in the early time period were conducted 

as part of the Western Arctic Shelf-Basin Interactions (SBI) program. To avoid confounding spatial and 

interannual trends, both the spatial and interannual analyses have been restricted to a region within 300 

km of Barrow Canyon.

B.2.2 Surface Winds

To assess the impact of winds on the hydrographic and velocity structure in our domain, we used the 

North American Regional Reanalysis (NARR) 10 m wind field (Mesinger et al. 2006), which has a spatial 

and temporal resolution of 32 km and 6 hours, respectively. The wind conditions for each shipboard 

transect were characterized by averaging over the 36 hours prior to the midpoint time of the section. We 

considered the along-shelfbreak component of the wind, where the orientation of the shelfbreak (119.6°T) 

was determined using a linear regression of the 90 m isobath between 166°W and the mouth of Barrow 

Canyon. Along-shelfbreak winds are hereafter referred to as easterly (negative) and westerly (positive) 

winds.

Table B1. Hydrographic and velocity data sources, delineated between cruise (by row) and section (by

semicolons). VMADCP = vessel-mounted ADCP; LADCP = lowered ADCP.

Cruise Ship Year Month Dates Velocity

HLY02 USCGC Healy 2002 May 17–22; 23–30 VMADCP

PLS0201 USCGC Polar 

Star

2002 Jul./Aug. 21–22; 4–6 LADCP

HLY0203 USCGC Healy 2002 Aug. 6–12; 13–18 VMADCP

NBP03 R/V N.B. Palmer 2003 Jul./Aug. 23–26; 27–31; 3–4; 5–7; 7–8; 8–9; 10–

11

VMADCP

HLY0303 USCGC Healy 2003 Sep./Oct. 16–17; 17–19; 11–14; 15–16 VMADCP

HLY0402 USCGC Healy 2004 May/Jun. 24–4 VMADCP

HLY0403 USCGC Healy 2004 Aug. 10–17; 18–24 VMADCP

HLY0404 USCGC Healy 2004 Sep. 15–16; 17–19 VMADCP

HLY09 USCGC Healy 2009 Aug. 6 VMADCP

HLY10 USCGC Healy 2010 Jul. 11; 13–14; 14–15 VMADCP

HLY1003 USCGC Healy 2010 Sep. 11; 12 VMADCP

HLY11 USCGC Healy 2011 Jul. 7–8; 9–10; 10–12; 14–15 VMADCP

HLY1103 USCGC Healy 2011 Oct. 8–9; 22 VMADCP

HLY12 USCGC Healy 2012 Oct. 9–11 VMADCP

HLY1301 USCGC Healy 2013 Aug. 13–14 VMADCP

HLY 1303 USCGC Healy 2013 Oct. 10–11; 13; 13–14; 15 VMADCP

HLY 1402 USCGC Healy 2014 Jul. 10–11; 13; 13–14; 20–22; 22–23; 24–25 VMADCP
From Corlett and Pickart 2017

B5

Figure B2. Locations of the 46 shipboard sections used in the study, color coded by cruise (see the 

legend and Table B1). The symbols denote the CTD stations comprising each section. The 90 m isobath 

is highlighted bold.

B.3 Methods

B.3.1 Gridding and Averaging

Vertical sections of hydrographic variables (potential temperature, salinity, and potential density) were 

constructed for each transect using a Laplacian-spline interpolator. This was also done for the crosstransect component of the ADCP velocity. The gridded velocity was then used to reference the gridded 

sections of thermal wind shear to compute sections of absolute geostrophic velocity, where the matching 

was done at each grid point over the common depth range of the thermal wind shear and directlymeasured velocity.

It is of interest to compute mean vertical sections of the various properties. In order to do this, it was 

necessary to construct an average cross-slope bottom profile and to locate each CTD station along this 


From Corlett and Pickart 2017

B8

Figure B3. Temperature-Salinity diagram of all of the hydrographic data offshore of the 90 m isobath and 

above 300 m depth. Each point denotes a single measurement, and the color represents the percentage 

of data within each grid cell of 0.1°C temperature by 0.1 salinity. The red lines indicate the different water 

mass boundaries. ACW = Alaskan coastal water; BSSW = Bering summer water; WW = newly-ventilated 

Pacific winter water; RWW = remnant Pacific winter water; MW = meltwater; AW = Atlantic water.

The mean section of absolute geostrophic velocity (Figure B4b) reveals two distinct circulation 

components. The first is a bottom-intensified, eastward-flowing shelfbreak jet. While previous studies 

have suggested the existence of this current using synoptic data (e.g., Pickart et al. 2005; Mathis et al. 

2007), our mean section robustly documents that such a feature exists along the edge of the Chukchi Sea. 

Dynamically this makes sense, as there is northward flow of Pacific water along the eastern flank of 

Herald Canyon (Woodgate et al. 2005; Pickart et al. 2010) some of which should turn eastward, following 

isobaths, to form a shelfbreak jet. This is analogous to the formation of the Beaufort shelfbreak jet via the 

outflow from Barrow Canyon (Pickart et al. 2005). A notable feature of the shelfbreak jet in Figure B4b is 

the enhanced near-bottom flow of AW associated with the strong isopycnal tilt noted above. Interestingly, 

the Beaufort shelfbreak jet also has a deep tail of eastward-flowing AW (Nikolopoulos et al. 2009; 

Brugler et al. 2014), which is believed to be associated with the spin-down phase of upwelling events that 

are common throughout the year (see Pickart et al. 2011). Further investigation is necessary to determine 

if the same mechanism is at work on the Chukchi slope.

While the existence of an eastward-flowing shelfbreak jet was anticipated, it was less clear what to expect 

on the continental slope. Our mean velocity section reveals the presence of a surface-intensified westward 

current seaward of the shelfbreak that is O(50 km) wide (Figure B4b), significantly broader than the

shelfbreak jet. The strongest velocities are in the upper 150 m (i.e., in the Pacific layer), although 

thecurrent extends deeper than this into the Atlantic layer. Laterally, the strongest mean flow is within 25 

km of the shelfbreak, although there is significant flow out to 60 km (the mean section suggests a double 

core structure, but this is due to the relatively small sample size as the synoptic realizations generally do 

not indicate such a feature).
From Corlett and Pickart 2017

B9

Figure B4. Mean vertical sections constructed from the 46 shipboard transects. (a) Potential temperature, 

and (b) Absolute geostrophic velocity overlain by potential density (contours, kg m-3). The black dashed 

lines denote the 5 cm s-1 velocity anomaly contour outlining the shelfbreak jet and slope current (see text). 

Positive velocities are eastward (the viewer is looking westward). The number of sections contributing to 

the mean at each cross-slope location is plotted in the top panel.

We refer to this current as the Chukchi slope current. To our knowledge this is the first robust 

documentation of the current, which we believe emanates (at least in part) from the outflow from Barrow 

Canyon. As noted in the introduction, the synoptic measurements reported in Brugler et al. (2014) are 

consistent with this notion, as are the sea surface temperature data and water column velocity 

measurements reported by Okkonen et al. (2009). Notably, the largest signature of Pacific summer water 

in the mean section of Figure B4a occurs within the current. The mass budget constructed below (Section 

B7) also suggests that the Chukchi slope current is fed by outflow from Barrow Canyon. Seaward of the 

slope current, near the offshore end of the mean section, there is another region of surface-intensified 

westward flow. This is weaker and less baroclinic and is presumed to be the southward edge of the 

Beaufort Gyre. Such an interpretation is in line with unpublished data from the Beaufort Sea, which 

indicates that the edge of the gyre is roughly 100 km offshore of the shelfbreak. This is considered below 

in more detail in the Discussion section.
From Corlett and Pickart 2017

B10

The distribution of water masses associated with the mean hydrographic sections supports our claim that 

the Chukchi slope current advects Pacific water out of Barrow Canyon. In particular, we computed the 

percent occurrence across the shelf/slope of each of the water masses present in the T/S diagram of Figure

B3. These are shown in the different panels of Figure B5. (Note that there is a steady decrease in 

occurrence near the seaward edge of the section due to data coverage, but this has been accounted for by 

normalizing by the number of realizations available at each given location.) MW is present across the 

entire section in the surface layer, though most prevalent on the outer shelf, while AW is found across the 

entire section in the deep layer. Notably, the highest percentages of BSSW, ACW, and RWW are found 

within the slope current. By contrast, the largest amount of WW occurs within the shelfbreak jet. These 

results suggest that the slope current is a fast-track for advecting Pacific water out of Barrow Canyon (in 

Figure B5. Relative occurrence of the different water masses in the mean hydrographic section. The 

black dashed lines denote the outline the shelfbreak jet and slope current as in Figure B4.
From Corlett and Pickart 2017

B11

the warm months this would consist of RWW, BSSW, and ACW), while the shelfbreak jet is a slower, 

longer route for Pacific water emanating from Herald Canyon (hence WW is only showing up along the 

Chukchi shelfbreak at this time of year).

B.4.2 Transport

Based on the average absolute geostrophic velocity section (Figure B4b), the mean westward volume 

transport of Pacific water in the Chukchi slope current, for the months of July – October, is -0.50 ± 0.07 

Sv (all error estimates in the paper are standard errors). This excludes MW (see Section B3.3 for how we 

distinguished between Pacific water and Atlantic water). The corresponding eastward transport of the 

shelfbreak jet is much smaller than this, 0.10 ± 0.03 Sv. These values are put into context of the mass 

budget of the Chukchi shelf in Section B7. The mean heat flux of the Pacific water in the slope current is 

-1.4 TW, which is roughly half of the summertime estimate of Brugler et al. (2014) for the Beaufort 

shelfbreak jet (~3 TW). The mean Pacific freshwater transport of the slope current is -33 mSv, which is 

roughly 50% greater than the summertime value of the Beaufort shelfbreak jet (~20 mSv). Both the 

corresponding heat and freshwater transports of the Chukchi shelfbreak jet are far smaller than these 

values (0.18 TW and 5.3 mSv, respectively).

The volume, heat, and freshwater transports of the Chukchi slope current are broken down by water 

masses in Figure B6. For this calculation we computed the transports for each realization that measured 

the slope current and then computed the mean, in order to circumvent the smoothing out of individual 

water masses in the mean section. (Note that this results in a 2% increase in volume flux, 22% increase in 

heat flux, and 3% increase in freshwater flux because of the selective sampling.) The largest volume 

transport component is associated with the RWW, followed by the MW contribution. The other 

components are comparable with the exception of the ACW, which is quite small. As is true for the 

volume flux, the heat flux is dominated by RWW and MW. However, the two Pacific summer waters 

(particularly BSSW) have fractionally greater contributions due to their warm temperatures, as does the 

AW. Finally, the same two water masses dominate the freshwater flux of the slope current — RWW and 

MW — but the latter contribution is greater due to its lower salinity.

It is worth noting that AW contributes to the full volume transport and heat transport of the slope current 

by roughly the same amount as BSSW (roughly 10% and 16% for volume transport and heat transport, 

respectively; Fig B6a, b). The fact that AW is flowing westward on the Chukchi slope is curious — this is 

Figure B6. Transport of the Chukchi slope current broken down by water mass (see Figure B3 for the 

water masses). Standard errors are indicated by the lines. (a) Volume transport; (b) Heat transport; (c) 

Freshwater transport.
From Corlett and Pickart 2017

B12

the opposite direction of the circumpolar boundary current that transports this water mass cyclonically 

around the sub-basins of the Arctic Ocean (e.g., Rudels et al. 1994; Aksenov et al. 2011). We have no 

explanation for this, and it warrants further investigation (beyond the scope of the present study).

B.5 Variability

B.5.1 Seasonality

There are pronounced seasonal changes in the presence of the different water masses on the Chukchi 

slope, as well as variations in the fluxes of mass, heat, and freshwater, from July to October. Starting with 

the water masses, we quantified the percent occurrence of each water type for each month seaward of the 

90 m isobath (Figure B7). Both of the Pacific summer waters, BSSW and ACW, were barely present (or 

completely absent) in July and August, but increased in abundance in September. This trend continued for 

the BSSW in October, but the ACW content dropped again that month. Appen and Pickart (2012) noted 

that BSSW also remained on the Beaufort slope later in the warm season than ACW. The Pacific winter 

waters, RWW, and WW varied out of phase with each other on the Chukchi slope, with RWW increasing 

in October and WW decreasing markedly at that time. This is not surprising, since heating/mixing 

through the summer transforms WW into RWW. Finally, the AW content remained steady from July–

September and then increased in October, while the amount of MW decreased steadily through the fourmonth period.

The most conspicuous seasonal change in Pacific water volume flux of the Chukchi slope current is the 

increase in westward transport in October (Figure B8). A possible explanation for this is the seasonality 

of the winds. We computed the monthly climatological along-shelfbreak 10 m wind speed over the 

Chukchi slope for the period 2002 – 2014. October has the strongest easterly winds during any month of 

the year, which would enhance the westward flow of the slope jet. A similar increase in heat flux and 

freshwater flux occurs during October as well. This is related not only to the wind, but is due in part to a 

substantial increase in the heat content and freshwater content of the water column (not shown). We note 

that the freshwater increase in October is not in line with the MW percentage (which is a minimum that 

month, Figure B7); it is related instead to the increase in the percentage of the RWW and the relatively

Figure B7. Mean monthly occurrence of water masses on the Chukchi slope offshore of the 90 m isobath 

and shallower than 300 m. Standard errors are indicated. (See Figure B3 for the water masses.)
From Corlett and Pickart 2017

B13

Figure B8. Mean monthly Pacific water transport of the Chukchi slope current with standard errors; N = 

11 for July, 10 for August, 5 for September, and 8 for October. (a) Volume transport; (b) Heat transport; 

(c) Freshwater transport.

fresh BSSW (note in Figure B3 that most of the BSSW is on the fresh end of that water mass range). In 

contrast to the Chukchi slope current, there is no pronounced seasonal variation in volume flux and 

freshwater flux of the Chukchi shelfbreak jet. While the shelfbreak jet reverses direction from eastward to 

westward in October — likely due to the increased easterly winds that month — the change is not 

statistically significant.

B.5.2 Interannual Variability

As noted in Section B2.1, the temporal coverage of shipboard sections is such that we can only contrast 

the early regime (2002 – 2004) versus the recent regime (2009 – 2014) for July through October. In 

addition, to avoid aliasing spatial trends into interannual variability, we have restricted the spatial extent 

for interannual analyses to be within 300 km of Barrow Canyon. In terms of the water masses on the 

slope, there was no significant change in the three most prevalent water masses, the RWW, AW, and MW 

(Figure B9). However, the percent contribution of all of the other water masses did change significantly. 

Both of the Pacific summer waters increased in occurrence, particularly the BSSW, while the WW 

contribution decreased.

The volume, heat, and freshwater transport of Pacific water in the Chukchi slope current all increased 

significantly from the early regime to the recent regime (Figure B10). As was the case with the seasonal 

changes in the fluxes of the current, it seems likely that this was the result of the wind. We calculated the 

mean along-shelfbreak 10 m wind speed for July – October for each year that we have data. This reveals a 

noticeable change between the two regimes (Figure B11). In the early regime, two out of the three years 

had very weak average winds during this four-month period. By contrast, in the recent regime five out of 

the six years were characterized by easterly winds. This would favor a stronger westward-flowing slope 

current in the latter period. This is consistent the results of Brugler et al. (2014), who demonstrated that 

the eastward-flowing Beaufort shelfbreak jet was retarded by the enhanced easterly winds along the north 

slope of Alaska over a similar time period. It is also in line with the observed increase in the presence of 

ACW and BSSW on the Chukchi slope during the later regime (Figure B9). (Our data indicate a 

decreased transport of the Chukchi shelfbreak jet from the early regime to the later regime, but the change 

is not statistically meaningful.)
From Corlett and Pickart 2017

B14

Figure B9. Interannual change in occurrence of water masses on the Chukchi slope from the early 

regime to the recent regime, offshore of the 90 m isobath and shallower than 300 m. Standard errors are 

indicated. (See Figure B3 for the water masses.)

Figure B10. Interannual change in Pacific water transport of the Chukchi slope current from the early

regime to the recent regime, with standard errors; N = 18 for the early regime, and 16 for the recent 

regime. (a) Volume transport; (b) Heat transport; (c) Freshwater transport.

B.6 Nature of the Chukchi Slope Current

B.6.1 Wind Influence

The apparent sensitivity of the Chukchi slope current to easterly winds on seasonal to interannual 

timescales makes one wonder if the current is present only during periods of substantial winds. For 

instance, it might be analogous to the Beaufort shelfbreak jet, which only flows to the west during 

upwelling-favorable winds. To investigate this, for each of the synoptic realizations we compared the 

transport of the current with the strength of the along-shelfbreak 10 m wind speed during/prior to the 

occupation of the transect (see Section B2.2). This reveals that, while the slope current is enhanced under 

synoptic wind forcing, it also exists when the winds are weak or even westerly. This relationship is 
From Corlett and Pickart 2017

B15

Figure B11. Mean along-shelfbreak 10 m wind speed from July – August each year, averaged over the 

Chukchi slope region (155 – 162°W and 72 – 74°N). The standard errors are indicated. Negative wind 

speeds are easterly.

significant at the 90% confidence interval and reveals a wind-dependent trend of -0.04 Sv per m s-1 of 

easterly wind speed with a no-wind volume transport of -0.45 Sv.

The influence of the wind on the Chukchi slope current is nicely visualized by constructing composite 

mean vertical sections under calm conditions versus times when there are moderate-to-strong easterly 

winds (Figure B12). In the weak wind composite the slope current is clearly evident, but it is relatively 

weak and the largest velocities are confined to the upper 40 m. By contrast, in the strong easterly wind 

composite the current is substantially stronger and the large velocities extend to 125 m. One also sees that 

the shelfbreak jet is considerably weakened under these latter conditions (the shelfbreak jet is reversed to 

the west in some of the synoptic easterly wind realizations). The overall conclusion is that, while the 

Chukchi slope current is sensitive to wind forcing, it still transports a notable amount of Pacific water to 

the west regardless of the synoptic winds.

B.6.2 Path and Downstream Evolution

Additional questions regarding the slope current are: What is the path of the current? How does it evolve 

going downstream? What is its ultimate fate? While definitive answers to these questions will require 

further observations and modeling, we can address some aspects here with the data in hand. As noted 

above, 37 of the 46 transects measured the presence of the slope current, and these realizations are 

depicted on a lateral map in Figure B13a. In the figure, the width of the arrows indicates the lateral scale 

of the current for the realization in question, while the length of the arrows denotes the mean speed of the 

current averaged over its width and depth (keep in mind that these are not true vectors; they are 

constrained to be normal to the sections).

One sees that the current is consistently located over the Chukchi continental slope throughout the entire 

domain, a distance of approximately 500 km westward from the mouth of Barrow Canyon. The median 

isobath over which the current resides (i.e., the region of strongest flow) is 238 m. In all but two of the 

realizations the current was shoreward of the 800 m isobath, and in 70% of the realizations it was inshore 

of the 400 m isobath. This implies that the current mainly resides on the upper continental slope but is not


From Corlett and Pickart 2017

B18

it flows westward from Barrow Canyon, this suggests that cross-shelfbreak subduction of mass is 

minimal.

An updated version of the Chukchi Sea circulation diagram shown by Brugler et al. (2014) is displayed in 

Figure B1, based on the new information presented in this study. In particular, it depicts Pacific water 

leaving Barrow Canyon in the slope current, as well as Pacific water flowing towards Barrow Canyon in 

the shelfbreak jet (where it is assumed to enter the western side of the canyon, then recirculate and exit 

the canyon). Using the known entrances and exits of Pacific water, a mass budget for the Chukchi shelf 

can be written as follows: 

0 = [Bering Strait Inflow] − [Barrow Canyon Outflow]

− Long Strait Outflow − Herald Canyon Outflow

Barrow Canyon Outflow = [Beaufort Shelfbreak Jet] + [Chukchi Slope Current]

−[Chukchi Shelfbreak Jet]

Herald Canyon Outflow = [Chukchi Shelfbreak Jet] + [HCwest]

(4)

As a simplification, we assume that any mass loss directly north into the basin from Barrow Canyon or 

Herald Canyon is minimal. The idea here is that such a flux should occur via turbulent processes (e.g.,

eddy formation as depicted in Figure B1), which are sporadic. We assume instead that the main transport 

occurs via the major advective outflows. For Herald Canyon it is unknown if a portion of the outflow 

veers to the west (akin the Chukchi slope current), so we include an unknown contribution from such a 

current which is referred to in (4) as HCwest. The mean transport in the Beaufort shelfbreak jet is taken 

from Nikolopoulos et al. (2009), and the mean transports through Bering Strait and Long Strait are taken 

from Woodgate et al. (2005). The other two transport values — the Chukchi slope current and Chukchi 

shelfbreak jet — come from the present study.

The magnitudes of the known terms in (4) are shown graphically in Figure B14. One sees that the residual 

is very small, implying that mass is balanced on the Chukchi shelf within the accuracy of the different 

transport estimates. It further suggests that the outflow from Herald Canyon is small and that most/all of it 

feeds the Chukchi shelfbreak jet (i.e., HCwest is essentially zero). This seems to be at odds with the 

results of Woodgate et al. (2005) who presented a year-long transport value in Herald Canyon of 0.2 – 0.3 

Sv, based mainly on a mooring on the eastern side of the canyon. However, the mooring in question was 

situated at the head of the canyon, and Pickart et al. (2010) present evidence, both from observations and 

numerical modeling, that most of the inflow on the eastern side of the canyon gets diverted to the east on 

the Chukchi shelf just north of Herald Shoal. This water then joins the Central Channel pathway and 

flows into Barrow Canyon (Pickart et al. 2016; Figure B1). In addition, data from multiple cruises suggest 

that there is no outflow from the western side of the mouth of Herald Canyon (Pickart et al. 2010; Linders 

et al. 2017).

Recently it has been argued that the transport of Pacific water through Bering Strait has increased over the 

past decade to 1.1 Sv (Woodgate et al. 2012). The interannual increase in transport of the Chukchi slope 

current, discussed in Section B5.2, could account for most of this increase. However, some of the 

enhancement of the slope current could also be due to diversion of Pacific water away from the Beaufort 

shelfbreak jet due to the stronger easterly winds of late (Brugler et al. 2014). The balance in Figure B14

implies that the mean outflow from Barrow Canyon is 0.53 ± 0.11 Sv. Using 9 years of mooring data

from the mouth of Barrow Canyon, Itoh et al. (2013) calculate a mean transport of 0.45 ± 0.07 Sv. The 

two values agree with each other within the error bars, although it should be kept in mind that the former 

estimate does not include the cold season for the Chukchi slope current and Chukchi shelfbreak jet 

portions.
From Corlett and Pickart 2017

B19

Figure B14. Mass budget of the Chukchi shelf inflows and outflows as discussed in the text, including the 

residual. Positive transports are inflows and negative transports are outflows. Standard errors are 

indicated.

Despite the uncertainty in the transport estimates of Figure B14, and keeping in mind the disparity in 

measurement techniques and spatial and temporal coverage of the different studies, the small residual 

nonetheless suggests that we have accurately accounted for most of the transport of Pacific water onto and 

off of the Chukchi shelf.

B.8 Summary and Discussion

The analysis presented here has elucidated our understanding of the water masses and circulation along 

the shelfbreak and continental slope of the Chukchi Sea. The collection of historical shipboard transects 

confirmed the existence of a shelfbreak jet during the warm season that, in the mean, flows to the east as a 

bottom-intensified current. The jet is weakened under easterly wind forcing, and, synoptically, it can 

reverse and flow towards the west. Interestingly, the eastward mean flow extends into the Atlantic layer, 

associated with strongly sloped isopycnals becoming shallower onshore. Such a deep tail of the 

shelfbreak jet transporting Atlantic water is also observed in the Beaufort shelfbreak jet, but the 

mechanism for this in the Chukchi domain remains unexplored.

Offshore of the shelfbreak our data indicated the presence of a current flowing to the west, which we have 

named the Chukchi slope current. The current is surface-intensified, O(50 km) wide, located mainly on 

the upper continental slope, and extends into the Atlantic layer (although the flow at depth is weak). In the 

mean (July – October) it transports 0.50 ± 0.07 Sv of Pacific water westward. In October, the Pacific 

water transport increases likely due to the stronger autumn easterly winds. The transport increased from 

the early regime (2002 – 2004) to the recent regime (2009 – 2014) presumably due to the enhanced zonal 

winds over this time frame. There is no alongstream trend in the cross-slope position of the current nor in 

the volume flux of the current. However, the current appears to meander, which implies that it is 

hydrodynamically unstable. Indeed, the cross-stream gradient of the Ertel potential vorticity changes sign 

with depth, which satisfies the necessary condition for baroclinic instability.
From Corlett and Pickart 2017

B20

We believe that the slope current emanates from Barrow Canyon. This is consistent with the fact that the 

highest concentrations of the Pacific summer waters and remnant Pacific winter water were found in the 

current (the highest concentrations of newly ventilated Pacific winter water were observed in the 

shelfbreak jet, which is expected based on the seasonal timing of water exiting the Chukchi shelf). 

Furthermore, taking into account the transport of the Chukchi slope current and the Chukchi shelfbreak 

jet, we were able to construct a balanced mass budget for the different Chukchi shelf inflows and 

outflows. We note that several recent studies present schematic representations of the regional circulation 

that indicate flow emanating from Barrow Canyon and turning westward (e.g., Day et al. 2013; Dunton et 

al. 2014; Wood et al. 2015; Moore et al. 2016). However, none of these studies invoke data to support this 

notion.

It remains to be determined why such a large portion of the Pacific water exiting Barrow Canyon turns to 

the left instead of the right. Pickart et al. (2005) observed that newly ventilated Pacific winter water 

flowing northward through Barrow Canyon transposed from the western side of the canyon to the eastern 

side, suggesting that most of the flow should progress eastward in the Beaufort shelfbreak jet. However, 

this was a synoptic survey which may not be representative of the mean, or it may be that the densest 

Pacific water does in fact progress eastward out of the canyon. Keep in mind that most of the newly 

ventilated Pacific winter water in our mean section was found flowing eastward in the Chukchi shelfbreak 

jet. It also should be stressed that our data were confined to the warm season; the situation could be 

different during the other months of the year, although the mass balance in Figure B14 suggests that this 

is not generally the case.

It also needs to be explored what role the Beaufort Gyre might play in the existence and path of the 

Chukchi slope current. We have argued here that the slope current is not the southern arm of the gyre. In 

addition to the above evidence that the slope current originates from Barrow Canyon, it is difficult to 

believe that the gyre would be situated at such a shallow isobath on the Chukchi slope (the median isobath 

corresponding to the center of the slope current is 238 m). If this were the case, then one would think the 

same should be true on the Beaufort slope. A mooring array deployed across the Beaufort shelfbreak and 

slope from 2002 – 2004 at 152°W showed no evidence of this (Nikolopoulos et al. 2009). In particular, 

the flow is eastward over this part of the continental slope. We note however, that the bathymetry of the 

Beaufort slope is different than that of the Chukchi Slope, and the presence of the Northwind Ridge could 

influence the location of the gyre west of Barrow Canyon.

On the other hand, it is quite possible that the presence of the Beaufort Gyre is dynamically tied to the 

existence of the slope current such that the two features are not entirely distinct. For instance, ice velocity 

maps show westward flow at the outskirts of the Beaufort Gyre near the edge of the Chukchi Sea. This is 

shown clearly in Steele et al. (2004), who argue that, under negative Arctic Oscillation conditions, the 

gyre expands towards the Beaufort and Chukchi shelves (although Steele et al. (2004) state that ACW 

emanates from Barrow Canyon as eddies that subsequently get entrained into the gyre under these 

conditions). The winter mean sea surface height and surface geostrophic velocity maps of Mizobata et al.

(2016) show evidence of a westward flow over the Chukchi slope that at times appears to be a separate 

current and at other times seems to be part of the gyre (see their Figure 4). Furthermore, a 10 year mean 

model simulation presented by Zhang et al. (2016) shows water exiting Barrow Canyon and turning 

westward along the southern edge of the Beaufort Gyre. Neither of these studies discuss this aspect of 

their results.

The role of wind in the formation of the Chukchi slope current needs to be investigated further as well. 

Okkonen et al. (2009) noted that, when the winds are out of the east, the Alaska coastal current can be 

displaced to the western side of Barrow Canyon. They also presented a sea surface temperature map 

during such conditions that showed a plume of warm water extending to the west out of the canyon. This 

is consistent with Brugler et al. (2014) who showed similar satellite evidence, along with velocity data 
From Corlett and Pickart 2017

B21

revealing the beginnings of the slope current. Surface drifters also depict flow heading west out of the 

canyon in response to wind events (Weingartner et al. 2015). Using a Pacific water tracer in a model 

simulation, Watanabe and Hasumi (2009) demonstrated that much of the water veers westward from 

Barrow Canyon during periods of strong easterly winds. It must be remembered, however, that while the 

Chukchi slope current is enhanced due to wind forcing, it is also present outside of such events, as 

demonstrated in the present study.

There are numerous ramifications associated with the notion that most of the Pacific water exiting Barrow 

Canyon progresses westward instead of eastward. Based on our data, the current has the potential to melt 

approximately 64,000 km2 of 1 m thick ice over the period July – October, which is equivalent to a region 

extending 170 km seaward from the shelf edge between the mouth of Barrow Canyon and 166°W. The 

freshwater transport of the slope current is also substantial and could contribute to the freshwater 

reservoir of the Beaufort Gyre. The Pacific water needs to be fluxed offshore for this to happen, and it is 

unknown by what means this might occur, e.g., via eddy formation and/or wind-driven exchange (both of 

which occur in the Beaufort shelfbreak jet). It also remains to be determined what the ultimate fate of the 

Chukchi slope current is. Perhaps the current becomes ill-defined farther to the west or becomes entrained 

into the Beaufort Gyre. Further investigation is necessary to explore these and other aspects of the current, 

including its role in the regional ecosystem.
From Fang et al. 2017

C1

C. Surface Current Patterns in the Northeastern Chukchi Sea and 

Their Response to Wind Forcing

Abstract

We measured northeastern Chukchi Sea surface currents using high-frequency radar systems (HFR) 

during the ice-free periods of August to October from 2010 – 2014. We analyzed these data, along with 

regional winds, using Self-Organizing Maps (SOM) to develop a set of surface current-wind patterns. 

Temporal changes in the SOM patterns consist predominantly of two patterns comprising northeastward 

and southwestward surface currents. A third pattern represents a transitional stage established during the 

onset of strong northeasterly winds. These patterns are analogous to the first two eigenmodes of an 

empirical orthogonal function analysis of the HFR data. The first principal component (PC1) is 

significantly correlated (~0.8) to that of the winds and is directly related to the time series of SOMderived patterns. The sign of PC1 changes when the speed of local northeasterly winds exceeds ~6 m s

-1

, 

at which point the northeastward surface currents reverse to the southwest. This finding agrees with 

previous models and observations that suggest this wind threshold is needed to overcome the pressure 

gradient between the Pacific and Arctic Oceans. The transitional stage is characterized by alongshore 

currents bifurcating in the vicinity of Icy Cape and wind-driven Ekman currents north of 71.5°N. Its 

development is a manifestation of interactions amongst the poleward pressure gradient, wind stress, and 

geostrophic flow due to the coastal setdown. 

C.1 Introduction

The Chukchi Sea is the gateway between the Pacific and Arctic Oceans. This vast marginal sea, though 

shallow in depth (~50 m), is prominent in shaping the thermohaline structure and freshwater budget of the 

western Arctic Ocean. The annual mean transport of Pacific water into the Chukchi Sea through Bering 

Strait is ~0.8 Sv (Roach et al. 1995; Woodgate et al. 2005) and was more recently reported as ~1.1 Sv 

(Woodgate et al. 2012), resulting in a poleward flux of heat and freshwater that affects sea ice 

distributions (Shimada et al. 2006; Woodgate et al. 2006; Steele et al. 2008; Wood et al. 2015) and 

supplies ~30% of the freshwater input to the Arctic Ocean (Serreze et al. 2006). Although the northward 

transport is forced by the steric height difference between the Bering Sea and Arctic Ocean due to 

interbasin salinity differences (Stigebrandt 1984; Aagaard et al. 2006), the flow field varies due to 

atmospheric forcing (Woodgate et al. 2012; Danielson et al. 2014). 

As Pacific waters flow northward through Bering Strait and across the Chukchi Sea, observations 

(Paquette and Bourke 1981; Weingartner et al. 1998, 2005, 2013a; Woodgate et al. 2005; Pickart et al. 

2010, 2016) and numerical models (Winsor and Chapman 2004; Spall 2007) indicate that the throughflow 

is bathymetrically steered along three major pathways (Figure C1). One branch follows Hope Valley and 

flows northward through Herald Canyon; another flows through the Central Channel between Herald and 

Hanna Shoals; and the third flows along the Alaskan coast and into Barrow Canyon where it becomes 

swift (~50 cm s-1) and narrow (~40 km) (Winsor and Chapman 2004; Spall 2007; Itoh et al. 2013; Gong 

and Pickart 2015; Weingartner et al. 2017a). During summer and early fall, the coastal branch includes 

buoyant, nutrient poor Alaskan Coastal Water, carried by the Alaskan Coastal Current (ACC) (Paquette 

and Bourke 1974). Itoh et al. (2013) used long-term mooring observations at the mouth of Barrow 

Canyon and found the greatest heat and freshwater fluxes occur from August to October.

Owing to the shallowness of the Chukchi shelf, wind forcing substantially influences the circulation 

pathways. On average, each branch flows poleward, whereas the mean wind field over the Chukchi shelf 
From Fang et al. 2017

C2

Figure C1. (a) Map of the Chukchi Sea with place names. Blue dots indicate 2012 HFR network grid 

points with more than 60% temporal coverage. Red squares show locations of HFR field sites. The red 

outlined area represents the NARR wind domain used in the SOM and EOF analyses, with the red 

triangle the location representative of the shelf wind time series. Legs 1 and 2 are hydrographic transects 

conducted in September 2013. Black dot north of Wainwright denotes the BC2 mooring. Gray arrows 

show schematic pathways of Pacific waters. Bathymetric contours are drawn from 10 – 100 m at 10 m 

intervals. Place names include HV = Hope Valley, HC = Herald Canyon, HeS = Herald Shoal, CC = 

Central Channel, and HaS = Hanna Shoal. (b) Grid points within the red rectangle are used to determine 

upcanyon and downcanyon flow conditions.

is from the east-northeast (Weingartner et al. 2013a) and opposes the mean flow. Local winds are highly 

correlated with the flow field on the Chukchi shelf (Weingartner et al. 2005; Itoh et al. 2013), except in

Herald Canyon (Woodgate et al. 2005). Winsor and Chapman (2004) used a barotropic model to examine

the sensitivity of shelf flow under changing winds and found that when northeasterly winds exceeded ~6 

m s-1, the coastal current near Barrow reversed to the southwest. Similar reversals were observed using 

subsurface moorings and ship-borne surveys (Mountain et al. 1976; Johnson 1989; Aagaard and Roach 

1990; Weingartner et al. 1998; Okkonen et al. 2009; Hirano et al. 2016); however subsurface 

measurements do not capture the upper few meters and so may not be representative of the surface current 

field. Therefore, surface current observations are essential to evaluate the flow imposed by wind forcing, 

especially when considering the potential transport of surface-trapped dissolved and suspended materials. 

In particular, surface currents may be important in the distribution of fish larvae (Wyllie-Echeverria et al. 

1992; Geoffroy et al. 2016) and zooplankton (Questel et al. 2013). 



From Fang et al. 2017

C6

Figure C2. Representative circulation regimes categorized from twelve SOM-derived patterns of surface 

currents (blue vectors) for 2012: (a) northeastward-flowing regime and (b) reversal regime. The frequency 

of regime occurrence is included in each panel, and the 80 m isobath is thicker to define Barrow Canyon. 

The inserted scatter plot denotes vertically-averaged velocities from mooring BC2 (black dot) when the 

flow regime occurred. Polar histograms on the right denote wind velocities associated with the flow 

regime (red triangle). The direction follows oceanographic convention and speed is shaded. The 

percentage indicates frequencies of winds blowing toward that direction.

this regime the flow is coastally intensified with speeds >30 cm s-1 in the canyon and within ~70 km of the 

coast in the southern mask. Currents are weak over the central shelf (<20 cm s-1) and even weaker (~5 cm 

s

-1

) north of 71.5oN. The northeastward-flowing regime occurred ~53% of the time, and ~76% of the 

winds associated with this pattern have a southerly component. Less than 10% of the winds are 

northeasterly with wind speeds <~6 m s-1, suggesting that under these wind conditions the flow is forced 

primarily by the poleward pressure gradient. Velocity measurements from the BC2 mooring indicate the 

subsurface flow is also northeastward under these conditions. Coincident southwestward subsurface 

currents are remnant from flow transitions that have not been fully completed and lag the surface flow. 

Weingartner et al. (2017a) suggest an adjustment time scale of ~1 day for the vertically-averaged flow in 

Barrow Canyon. 

The second most common circulation feature is the reversal regime (Figure C2b), which occurred ~11% 

of the time and whose structure is nearly opposite the northeastward-flowing regime. It consists of 

southwestward flow nearshore and westward flow farther offshore. The wind histogram shows ~60% of 

the associated winds are northeasterly at >6 m s-1 with ~10% of occurrences associated with northerly 
From Fang et al. 2017

C7

winds with speeds >6 m s-1. Reversal regime currents are stronger than those for the other circulation 

patterns, with current speeds >30 cm s-1 in Barrow Canyon and inshore of the 40 m isobath offshore of 

Point Lay. Currents north of 71.5oN and east of Hanna Shoal are also swifter (~15 cm s-1) for this regime 

compared to the northeastward-flowing regime. Southwestward surface flow at BC2 predominates during 

the reversal regime. The occurrence of subsurface northeastward currents during the reversal regime is 

again due to the adjustment time from surface to depth. This result indicates that, at the head of Barrow

Canyon, the flow structure is coherent vertically during the two most common surface circulation patterns 

but that the vertical shear may be substantial during flow transitions.

The third regime occurred ~16% of the time and is denoted as the northwesterly wind regime (Figure 

C2c). For this pattern ~90% of the winds are northwesterly, with more than half of the wind events having 

speeds >6 m s-1. This regime appears to result from interactions between the poleward pressure gradient 

(northeastward flow) and wind-induced Ekman transport (southward flow). The flow field appears to be 

spatially variable depending on which driving force dominates. For example, when winds initially shift to 

the northwesterly quadrant, the resultant Ekman transport may not be large enough to overcome the 

background pressure gradient. We have found cases where the currents south of 71.5°N are eastnortheastward (similar to those seen in the northeastward-flowing regime), but southward north of 

71.5°N. This flow pattern develops under evolving northwesterly winds and may last longer than a day. 

As northwesterly winds persist, the Ekman currents gradually overwhelm the poleward pressure gradient, 

so that currents are southward at ~10 – 15 cm s-1 over much of the radar mask, with stronger flows (~20 

cm s-1) inshore of the 40 m isobath and southeast of Hanna Shoal. 

Figure C2. (continued) (c) northwesterly wind regime and (d) divergent mode. Note that the scales of 

current vectors in (c) and (d) are different than those in (a) and (b).
From Fang et al. 2017

C8

Subsurface flows observed at BC2 indicate predominately northeastward currents, indicating that the flow 

was vertically sheared over much of the canyon, except near Point Franklin where the surface flow was 

still northeastward. This finding implies that the subsurface flow over the central shelf must have been 

onshore in order to feed the transport in Barrow Canyon. This current pattern differs from the reversal 

regime, during which both the subsurface and surface flows were southwestward. 

The fourth regime is the divergent mode (Figure C2d). North of 71.5°N, currents are ~10 cm s-1 and 

westward. South of this latitude, a recirculation is suggested, which includes cyclonic flow near the head 

of Barrow Canyon and an anticyclonic circulation at ~164°W, 70.5°N. The recirculation includes 

northeastward currents in Barrow Canyon and southwestward currents in the southern mask between Icy 

Cape and Point Lay. Approximately 80% of the winds concurrent with the divergent mode are 

northeasterly, and ~70% of these winds have speeds >6 m s-1, similar to those of the reversal regime; 

however, the subsurface and surface flow at BC2 is still northeastward, in contrast to the reversal regime. 

We will demonstrate that the divergent mode is a transitional stage between the northeastward-flowing 

and reversal regimes, as northeasterly winds begin to overcome the poleward pressure gradient. 

As previously mentioned in Section C2.3, SOM patterns vary slightly over the years; however, the regime 

descriptions above apply to all years. For example, the reversal regimes of 2010, 2011, 2013, and 2014 

are all analogous to our description for the reversal regime in 2012. We find that the location of 

southwestward flow, offshore westward flow, and subsurface currents observed at BC2 approximate to 

the 2012 results. Similar agreements apply to the other three flow regimes. 

Histograms show the monthly distribution of the four flow regimes for all years (Figure C3). The results 

indicate that the northwesterly wind regime was rare, consistent with the regional mean winds being 

predominantly from the east-northeast. The northeastward-flowing and reversal regimes occurred ~43% 

and ~27% of the time, respectively. In general, as the frequency of the northeastward-flowing regime 

increases the frequency of the reversal pattern decreases, and vice-versa. The divergent mode appeared 

~7% of the time overall with durations varying from several hours to a maximum of ~5 days.

The remaining pattern, not included in our categorization, consists of variable and weak (~1 cm s-1) 

currents which are unreliable and occurred during periods of sparse data returns due to equipment 

maintenance, ice, ionospheric interference, and/or low winds. Collectively these conditions occurred 

~14% of the time during 2012 and were aggregated into a pattern for data gaps referred to as G (see 

Appendix A Figure A-1e). 

C.3.2 The Role of Winds: SOM Perspective

To illustrate how surface currents change under varying wind conditions, time series of the SOM regimes 

from August through October 2012, along with wind vectors, are presented in Figure C4. The gray line in 

Figure C4 tracks the normalized data return and facilitates interpretation of regimes associated with data 

gaps. For example, Pattern G occurred from 11 to 14 August when data returns were very low. Low data 

returns also occurred from 10 to 14 October due to diurnal ionospheric interference (Teague 2001). 

The wind field from 4 – 28 August was mainly southwesterly-southerly, and the northeastward-flowing 

regime persisted through most of the month. It was also present during the southerly wind events of 1 – 3, 

6 – 7, and 20 – 28 October and occurred from 1 – 3 August and from 18 – 22 September when weak (≤4 

m s-1) winds from varying directions prevailed, consistent with our notion that weak winds are insufficient 

to overcome the poleward pressure gradient force.

After 28 August, southerly winds relaxed, and the winds became northeasterly and increased to a 

maximum of 10 m s-1 on 31 August. During this wind transition, the circulation shifted briefly into the 
From Fang et al. 2017

C9

Figure C3. Monthly occurrences of SOM-derived circulation regimes for 2010 – 2014 with different 

hatching denoting different years. (a)-(c) Northeastward-flowing regime (NE). (d)-(f) Reversal regime (R). 

(g)-(i) Divergent mode (D). (j)-(l) Northwesterly wind regime (NW). A 15-day data gap in August 2013 is 

the reason for few regime estimates in that month. 

divergent mode before the reversal regime was established. Reversal regimes consistently occurred during 

and/or shortly after pulses of strong (~10 m s-1) northerly and/or northeasterly winds (e.g., 16 – 17, 22 –

23, and 27 – 29 September and 30 – 31 October). The divergent mode accompanied each transition from 

the northeastward-flowing regime to the reversal pattern, appearing as the northeasterly winds increased 

to relatively high magnitudes. Based on the wind time series and the EOF analysis of Section C3.3, we 

find that the divergent mode occurred when northeasterly winds reached ~6 m s-1. Hence, the divergent 

mode appears to be a transition between these two regimes; a point we will return to in Section C4. 

The northwesterly wind regime (Figure C2c) occurs primarily during periods when winds are veering 

from northwesterly to westerly or vice versa (e.g., 8 – 18 October) and/or during wind transitions 

involving westerlies (14 – 15 September). These results suggest that strong (>6 m s-1) and sustained 

northwesterly winds (>1 day) are required to initiate this regime. We did not observe this pattern in 2011 

and 2013, however, because northwesterly winds seldom occurred. When present, these events were 

short-lived (<1 day) in comparison to the >3 day events registered in 2012. 
From Fang et al. 2017

C10

Figure C4. (a) Upper panel: time series of wind vectors in August 2012. The vector direction follows 

oceanographic convention. Lower panel: SOM-derived circulation regimes (black dots) in August 2012 

and normalized data returns (gray line). The abbreviations are: R, reversal regime; D, divergent mode; G, 

pattern G; NW, northwesterly wind regime; NE, northeastward-flowing regime. (b) As in (a), but for 

September 2012. (c) As in (a), but for October 2012.
From Fang et al. 2017

C11

C.3.3 The Role of Winds: EOF Perspective

The bulk of the SOM patterns are the northeastward-flowing and reversal regimes, which suggests that 

these should be linked to the leading EOFs (Mau et al. 2007). In this section, we use EOFs and SOM 

patterns to corroborate one another. Table C1 summarizes the EOF results in terms of the variance 

explained by the first (Mode 1) and second eigenmodes (Mode 2), the number of significant eigenmodes, 

and the correlation between PC1 of the currents and the winds. Mode 1 of the currents and winds account 

for at least 50% of the total variance, and the correlation between each variable’s PC1 is significant. 

Mode 1 of the 2012 surface currents (Figure C5a) shows a pattern similar to the reversal regime derived 

from SOM (Figure C2b), which, when multiplied by its negative weight in the time domain, is analogous 

to the northeastward-flowing regime (Figure C2a). Mode 1 of the winds depicts northeasterlies (Figure 

C5c), and its negative weight describes southwesterly winds. The structures of Mode 1 in other years 

were similar to those in 2012, except in 2010 when Mode 1 for the winds was aligned in the east-west 

direction. 

Mode 2 of the currents, which accounts for 9% to 17% (depending upon year) of the surface current 

variance, is characterized by southward flows and onshore currents that appear to bifurcate somewhere 

between Icy Cape and Wainwright (Figure C5b). Meanwhile, Mode 2 of the winds portrays 

northwesterlies (Figure C5d) and explains 19% to 34% of the variance across years. The structure of 

Mode 2 of the currents and winds in other years is comparable to those from 2012. The correlation 

between the principal components of Mode 2 (PC2) of the currents and the winds is only statistically 

significant for 2012 and 2014; however, as discussed later, we suspect that this relationship may not be 

meaningful.

Figure C6 shows the relationship between currents and winds in terms of wind speed and direction, PC1 

of currents and winds, and the circulation regimes for the 2012 data. PC1 of currents and winds are 

normalized individually, and both are referenced to the reversal regime and northeasterly winds. PC1 

values for the currents (winds) approaching unity indicate a flow field similar to that of Mode 1 (Figure 

C5a). In August, PC1 of the currents and winds were both generally negative (i.e., northeastward flow), 

consistent with the SOM results.

Northeasterly and northerly winds, as well as the divergent mode and reversal regime, were more frequent 

in September and October than in August. As a consequence positive PC1 values for both the currents 

and winds were also more frequent. For example, the winds were northeasterly at ~6 m s-1 throughout 10

Table C1. Summary of the EOF correlation analysis from the reconstructed fields. The correlations are all 

significant at the 95% significance level using the effective number of degrees of freedom derived from 

the integral time scale.

Year Months

Correlation 

coefficient 

between PC1 of 

currents and 

winds

# of 

significant 

eigenmodes 

(currents)

% variance explained 

by Mode 1

(currents, winds)

% variance explained 

by Mode 2

(currents, winds)

2010 9 – 10 0.36a 2 50, 63 17, 30

2011 8 – 10 0.56 1 65, 65 9, 27

2012 8 – 10 0.74 2 57, 57 11, 34

2013 9 – 10 0.75 3 66, 60 15, 29

2014 8 – 10 0.83 3 64, 75 14, 19 a

Low correlation coefficient is because wind Mode 1 in 2010 aligns nearly east-west, but Mode 1 of the currents is in 

the northeast-southwest direction.
From Fang et al. 2017

C12

Figure C5. (a) EOF Mode 1 of 2012 reconstructed HFR data (blue vectors) and its explained variance. 

For clarity, only subsampled vectors are shown. (b) As in (a), but for Mode 2. (c) EOF Mode 1 of 2012 

NARR winds (black vectors), also subsampled. (d) As in (c), but for Mode 2.

– 13 September, when the pattern corresponded to the divergent mode (indicated by a red arrow, Figure 

C6). During this period, the current PC1 was ~0.1, indicating the flow field had not yet fully reversed, 

consistent with the SOM analysis. Therefore, the EOF analysis corroborates the SOM conclusion that 

northeasterly winds ≥6 m s-1 are critical in shifting the surface circulation through the divergent mode and 

toward the reversal regime.

We examined four selected periods (labeled T1 – T4, Figure C6) in September that consisted of reversal 

regimes. Periods T1 – T3 had northeasterly winds >6 m s-1 (i.e., the wind PC1 had values >0.5). Each 

event was preceded by periods of weaker winds from varying directions. The current PC1 evolved 

similarly as its values changed from negative to positive. The temporal evolution of PC1 during each of 

these periods suggests that the reversals lagged the winds by 6 –9 hours. Period T4 also corresponded to 

the reversal regime, when winds were northerly with speeds >~6 m s-1. Each of these reversals coincided 

with southwestward or upcanyon transports of ~1 Sv as measured by BC2 (Weingartner et al. 2017a).

The PC1 values for the currents fluctuate when the northwesterly wind regime is present (e.g., 9 – 18 

October) but are comparatively small (≤ ± 0.2). The northwesterly wind regime events are not described 

well by the evolution of PC1 because their overall flow behaviors are different from those associated with 

Mode 1. Unlike the good correspondence between the evolution of PC1 and the SOM-flow regimes, we 

do not find a direct relationship between PC2 and the flow regimes. For example, high positive PC2 
From Fang et al. 2017

C13

Figure C6. (a) Time series of wind speed with wind speeds ≥6 m s-1 highlighted in gray. (b) Time series 

of wind direction with gray shading corresponding to winds blowing from the northeast quadrant. (c) Time 

series of PC1 of surface currents (black) and PC1 of winds (gray) with gray shaded areas highlighting 

positive PC1 values. PC1 values approaching one indicate currents or winds approximating the Mode 1 

structure. (d) Time series of representative circulation regimes derived from SOM (black dots): R, reversal 

(highlighted in gray); D, divergent mode; G, pattern G; NW, northwesterly wind; NE, northeastwardflowing. The gray shaded area highlights the reversal regime. Vertical red lines in each panel denote 

periods of selected reversal regimes (T1, T2, T3, and T4). Red arrow and bar denote a northeasterly wind 

event described in the text.
From Fang et al. 2017

C14

(~0.8) for winds and currents were found in August, but the observed winds and flow fields were 

predominately southerly and northeastward, respectively. We suspect that EOF Mode 2 of winds and 

currents is a consequence of the orthogonality requirement of the EOFs computation and thus a limitation 

of EOFs. Our results suggest that higher order eigenmodes should be interpreted cautiously for data sets 

containing large variability. 

C.3.4 Mean Circulation Patterns

We formed composite circulation maps for periods when the HFR data recorded downcanyon 

(northeastward) and upcanyon (southwestward) flow events for 2010 – 2014 (Figure C7). The 

classifications were defined by average flow conditions derived from twelve HFR grids (Figure C1b) near 

the head of Barrow Canyon. For the downcanyon condition flow is defined to be toward 56°T ± 4°, while 

for the upcanyon condition it is toward 236°T ± 4°, with 56°T the approximate axis of Barrow Canyon. 

Mean downcanyon and upcanyon patterns observed by HFR are comparable with the SOM-derived 

northeastward-flowing and reversal regimes, respectively. 

For the 5-year period, we found 868 hourly surface current observations satisfying our downcanyon 

criterion. The polar histogram indicates that ~75% of the time the downcanyon pattern occurs under 

variable wind directions at moderate wind speeds (≤6 m s-1). Approximately 25% of the winds were 

northeasterly (203 – 246°T), with only 5% of those >6 m s-1. These strong northeasterlies are the catalyst 

for a reversal to upcanyon flow, and thus indicate when the flow field lagged the wind. The downcanyon 

average suggests four distinct flow regions (Figure C7a; labeled 1 – 4) under mean winds of ~1 m s-1

westward (~264°T). Region 1 lies north of 71.5°N on the eastern flank of Hanna Shoal, and here the mean 

currents are weak (a few cm s-1) and directionally variable. The flow in Region 2 is generally eastward 

with mean currents 5 – 10 cm s-1, consistent with results from moorings, ship-borne surveys, and models 

(Winsor and Chapman 2004; Weingartner et al. 2005; Spall 2007; Gong and Pickart 2015). In the 

northwestern part of Region 2 the flow is southeastward, suggestive of flow moving eastward from the 

Central Channel across the shelf south of Hanna Shoal. Region 3 encompasses the head of Barrow 

Canyon, where mean currents are northeastward and swift (~30 cm s-1). The mean currents in Region 4, 

which covers the southern portion of the southern radar mask, are northeastward at ~10 cm s-1. The 

downcanyon average suggests convergence of the nearshore flow with that from the central shelf near 

70.8°N, 162.5°W, with current speeds increasing as these flows converge. 

Observations consistent with the upcanyon criterion were fewer, with only 368 hourly values. Mean 

winds for the upcanyon composite (Figure C7b) were ~7 m s-1 toward ~236°T, consistent with the 

reversal regime. Winds were ~70% northeasterly, ~6% northerly, and ~11% easterly. Overall the currents 

are stronger than those of the downcanyon average. Surface currents in Region 1 are westward at ~9 cm s1

, diminishing northward, while currents in Region 2 are westward or northwestward at 15 – 20 cm s-1, 

which could carry canyon waters toward the south side of Hanna Shoal and across the central shelf. 

Currents near Barrow Canyon (Region 3) are particularly strong, with a mean of ~50 cm s-1 to the 

southwest (~240°T). The flow in Region 4 is also southwestward (~240°T) but with speeds of ~20 cm s-1. 

The composites indicate that the down- and upcanyon flow structures in Regions 2 and 3 are mirror 

images of each other; however, Regions 1 and 4 are quite different from one another. For the upcanyon 

case, in Region 1 the flow is westward and oriented ~35° to the right of the wind. For the downcanyon 

case the surface currents are weak and variable. We show later that the surface flow in Region 1 is 

consistent with Ekman dynamics. 

For the Region 4 downcanyon case, the mean current speed is maximum (~25 cm s-1) over the 40 m 

isobath and not statistically different (p<0.05) from the speed over the 30 m isobaths, but in the upcanyon 

case, the mean speeds along these isobaths are significantly different from one another. It is maximum 
From Fang et al. 2017

C15

Figure C7. Mean (a) downcanyon and (b) upcanyon surface currents (blue vectors) for 2010 – 2014. 

Bathymetric contours (gray lines) are drawn within 200 m at 40 m intervals, with depths less than 40 m 

drawn at 10 m intervals. The black vector denotes the mean winds, and the polar histogram along 72oN 

denotes wind velocities. Areas circumscribed by dashed lines denote regions (labeled 1 – 4) of distinct 

flow behaviors (see text). For clarity, the vectors were subsampled.
From Fang et al. 2017

C16

(~38 cm s-1) over the 30 m isobath and monotonically decreases offshore with the flow being ~32 cm s-1

over the 40 m isobath. These differences imply an asymmetry in the alongshore transport, with more 

transport carried by the inner shelf flow during upcanyon events compared to downcanyon cases. In the 

upcanyon case, the alongshore winds cause a sea level setdown and an alongshore southwestward flow

established by the cross-shore pressure gradient. This gradient should be greatest within ~50 km of the 

coast, which is the e-folding scale for the 140 km barotropic radius of deformation (for a shelf depth of 40 

m). In the downcanyon mean the winds are weak and directed offshore, in which case the mean flow is 

largely forced by the poleward pressure field and appears strongest over the 40 m isobath. This suggestion 

is consistent with the convergence in modeled streamlines shown by Winsor and Chapman (2004) and 

Spall (2007) for the same region. 

C.3.5 Spatial Correlation Structure

In this section we use the complex correlation function (Kundu and Allen 1976) to examine the spatial 

correlation structure of the down- and upcanyon flows using the DINEOF reconstructed data. The 

calculations use two reference grid points. The first, at 71.2°N, 160°W, is chosen to examine relationships 

between flow near Barrow Canyon and those elsewhere. The second grid point, at 72°N, 160°W, is over 

Hanna Shoal. 

Based on the sample number (1148 for downcanyon; 249 for upcanyon) and estimated integral time scale 

(~40 hour) from the 2012 DINEOF field, the 95% significance levels for the down- and upcanyon 

correlations are 0.36 and 0.71, respectively. The correlation field using the Barrow Canyon grid point is 

similar for both flows (Figures C8a-b), with correlations high in the canyon and across the southern mask 

but much smaller north of 71.5°N. The correlated regions are quite large. For example, the largest 

distance between grid points with correlations ≥0.8 is ~170 km for the downcanyon case and ~300 km for 

the upcanyon. 

Using the Hanna Shoal grid point, the well-correlated region is limited to the area north of 71.5°N for the 

downcanyon case (Figure C8c). The length scale of the strongest correlation is only ~30 km, as expected 

given that the contributing currents are weak and variable in this region (Figure C7a). For the upcanyon 

case (Figure C8d), the spatial correlation structure is again significant and mainly confined to the region 

north of 71.5°N over a length scale of ~120 km. In aggregate, the SOM and the mean and correlation 

fields indicate that the shelf circulation north of 71.5°N is dynamically different from the shelf to the 

south. 

C.4 Discussion

Our analyses have revealed two major surface circulation regimes related to the local winds. The 

northeastward-flowing regime transports water from along the coast and the central shelf into Barrow 

Canyon and toward the Arctic Ocean. During summer and early fall, this surface transport involves warm, 

low-salinity coastal waters and somewhat more saline waters from the central shelf. The reversal regime 

transports waters up the canyon and across the Chukchi shelf, suggesting upwelling within the canyon 

and along the west coast of Alaska. We also identified the divergent mode, a transitional flow field 

established as northeasterly winds intensify, which eventually evolves into the reversal regime. This mode 

includes divergence in the nearshore currents between Icy Cape and Wainwright, with one branch leading 

into Barrow Canyon and the other proceeding southwestward. In addition, we found that the currents 

north of 71.5°N (Region 1 in Figure C7) are distinctly different from the currents south of this latitude. 

We explore these findings in greater detail in the following paragraphs.

In summer and fall, a portion of the Bering Strait inflow is carried eastward from the Central Channel and 

merges with ACC waters as they approach Barrow Canyon. This eastward flow is represented in 
From Fang et al. 2017

C17

Figure C8. Spatial distribution of complex correlations of surface currents with respect to a reference grid 

point (large black dot) for (a, c) downcanyon and (b, d) upcanyon flow. Black contour denotes the 0.8 

correlation level. Correlation below the e-folding scale is shaded in cool colors. The 95% significance 

level is estimated from the effective number of degrees of freedom using the integral time scale.

circulation models forced solely by the poleward pressure gradient (Winsor and Chapman 2004; Spall

2007), in prior observations (Weingartner et al. 2005, 2013a), and is largely consistent with the 

northeastward-flowing regime (Figure C2a) and observed downcanyon mean (Figure C7a). The models 

also indicate that some of the Central Channel transport flows geostrophically around the north side of 

Hanna Shoal. Along the east side of the Shoal the modeled flow is southward before turning eastward into 

the head of Barrow Canyon. This southward flow is much weaker than that on the western side because 

the isobaths diverge on the northeast side of the Shoal. However, the modeled southward flow is not 

captured by the SOM, nor is it evident in the vertically-averaged mean flows estimated from moorings on 

the shelf east of Hanna Shoal (Weingartner et al. 2017a).

We believe that one cause in the differences in surface circulation north and south of 71.5°N lies with the 

shelf hydrographic structure. Figure C9 shows vertical sections of potential temperature (θ) and salinity 

(S) along Legs 1 and 2 (Figure C1) from September 2013. In both sections waters deeper than ~20 m 

were cold (<-1°C) and salty (≥32.5) winter waters. North of 71.5°N a ~15 m deep surface layer contained 
From Fang et al. 2017

C18

Figure C9. Vertical sections from September 2013 of potential temperature and salinity overlain with 

isopycnals (white contours) for (a, c) Leg 1 and (b, d) Leg 2. Black arrow denotes the location of 71.5°N. 

Note that the shading scales in (a) and (b) and the horizontal scales for Legs 1 and 2 are different.

cool (0°C) and fresh (salinity ~26 – 29) meltwaters (MW) underlain by a thin but strong pycnocline. 

South of 71.5°N the stratification was weaker and the surface waters were warmer (3 – 5°C), moderately 

salty (salinity ~31), and derived from Bering Sea summer waters. The front delineating these surface 

water masses was centered near 71.5°N. Similar fronts around this latitude are apparent in sections shown 

by Lu et al. (2015) and Weingartner et al. (2017b). The latter find that the shelf north of 71.5°N and east 

of Hanna Shoal supports a counterclockwise baroclinic geostrophic flow component that opposes the 

modeled barotropic flow around the Shoal. These opposing pressure tendencies result in weak geostrophic 

flow. 

We suggest that the strong stratification north of 71.5°N in conjunction with weak geostrophic flow 

suggests that the surface currents in this region are largely governed by Ekman dynamics, with the Ekman 

layer presumably confined to the upper 15 m. For the upcanyon case, the mean winds were ~7 m s-1

toward ~236°T, which would impel a mean flow in the upper 15 m of ~6 cm s-1 toward 326°T. The 

observed flow (within the upper 2 m) averaged 9 cm s-1 toward 270°T (i.e., ~34° to the right of the wind). 

As discussed in Dzwonkowski et al. (2011), baroclinic shear and stratification inhibit deepening of the 

surface boundary layer and reduce its interaction with the bottom boundary layer, resulting in 

enhancement of surface transport. South of 71.5°N the stratification is weaker and the flow is primarily 

barotropic and geostrophic (Weingartner et al. 2013a), so that wind momentum is diffused over much of 

the water column.
From Fang et al. 2017

C19

We next examine the dynamics associated with the divergent mode aided by snapshots of wind and 

surface current maps (Figure C10) from September 2011. The maps encompass a period when the winds 

transitioned from southerly (12 September; Figure C10a) to strong northeasterly (15 – 19 September; 

Figure C10b-d). Each map includes a vector showing the wind at the measurement time (black) along 

with the wind vectors for the preceding 12 (dark gray) and 24 hours (light gray). 

Prior to 12 September the winds were southwesterly, and the northeastward-flowing regime held. At this 

time both the poleward pressure gradient and the coastal sea level setup acted in concert to propel the 

flow. On 13 September (not shown), the winds became east- northeasterly, and the divergent mode 

appeared for the next few days (Figure C10b). This transitional mode consisted of westward Ekman drift 

north of ~71.5°N and southwestward alongshore flow inshore of the 30 m isobath offshore of Point Lay. 

The time scale for the appearance of these two circulation features is plausible. The Ekman adjustment 

time scale is only a few hours (Allen 1973), whereas the spinup time scale for the alongshore flow is 

proportional to h/r, where h is the water depth and r is the linear bottom friction coefficient. The

adjustment to southwestward flow in the shallow (<30 m) nearshore waters should be ≤0.5 day for 

Figure C10. HFR observations (blue vectors) during 12 – 19 September 2011 at indicated UTC time. The 

black vector denotes the wind at the cardinal hour of the snapshot, with the magnitude and direction 

shown in the legend, and the gray shaded vectors are winds for the preceding 12 and 24 hours. 

Bathymetric contours (gray lines) are drawn within 80 m at 10 m intervals. For clarity, the current vectors 

were subsampled.
From Fang et al. 2017

C20

r ~5 x 10-4m s-1 (e.g., Brink 1998) but longer in deeper waters. Apparently there is a ridge in the crossshore pressure gradient northwest of Icy Cape near ~70.5°N, manifested in the velocity field as the region 

of negligible flow along ~70.5°N on 15 September (Figures C10b). Note that the width of the 

southwestward flow is ~90 km and less than the local barotropic radius of deformation (~130 km). The

implication is that although the cross-shore sea level gradient sloped downward onshore, it was of 

insufficient magnitude to reverse the pressure gradient farther offshore or in Barrow Canyon. By 17 

September (Figure C10c) the winds had veered more towards the southwest. In response, canyon waters 

offshore of Point Franklin began veering northward, while the area of southwestward flow between Icy 

Cape and Point Lay expanded offshore. The winds continued to veer toward the southwest and intensified 

such that by 19 September the reversal regime was fully established, with the swiftest flows in the canyon 

and along the coast (Figure C10d). The divergent mode, although generally short-lived, thus has the 

potential to disperse materials advected from the central shelf along the coast of western Alaska and into 

Barrow Canyon. 

We expect that the SOM-derived surface circulation patterns occur throughout the year and that the 

divergent mode should occur more frequently through late fall and winter when transitions between 

downcanyon and upcanyon flow events are more common (Weingartner et al. 2017a). However, threshold 

wind speeds that catalyze the changes amongst the various modes may differ given that the surface stress 

will be modulated by sea ice. 

C.5 Summary

We used SOM and EOFs in an examination of HFR-derived surface current patterns in the northeastern 

Chukchi Sea during the open-water seasons (nominally August – October) from 2010 – 2014. We 

identified major surface current circulation patterns and their dynamics in conjunction with regional 

winds and showed how this portion of the shelf surface circulation responds to the wind field and a 

background poleward pressure gradient. 

We found that surface currents south of ~71.5°N and in Barrow Canyon flow northeastward except when 

northeasterly winds exceed ~6 m s-1. The northeastward flow is primarily a manifestation of the largescale pressure gradient between the Pacific and Arctic Oceans. When northeasterly wind speeds are >6 m 

s

-1

, the coastal sea level setdown is large enough to cause southwestward surface flow inshore and 

westward flow over the central shelf. Less frequently southward surface currents occur over portions of 

the domain in response to northwesterly or westerly winds exceeding ~6 m s-1. 

We also detected the divergent mode, a transitional circulation feature of 1 to 5 days duration, which 

develops as the flow field adjusts from northeastward to southwestward in response to intensifying 

northeasterly winds. The divergent mode is characterized by eastward (onshore) flow across the central 

shelf, which diverges upon approaching the coast. A portion of this onshore flow continues into the head 

of Barrow Canyon while the remainder turns southwestward inshore of the 40 m isobath; however, 

sampling constraints prevented us from a precise determination of the nature of the divergence between 

these two regions. The divergent mode is a consequence of interaction amongst the poleward pressure 

gradient, increasing northeasterly winds (>~6 m s-1), local bathymetry, and the regional hydrographic 

structure. 

Along ~71.5°N there is an extensive zonally-oriented front. North of this front, surface pressure gradients 

are weak, and the surface currents are Ekman-like. The strong Ekman response occurs because the water 

column is strongly stratified; fresh, cool MW occupies the upper 15 m, and cold, salty winter waters 

occupy the bottom layer. South of the front the manifestation of the poleward pressure gradient is 

stronger, and the water column is less stratified, with the surface layer containing moderately salty and 

warmer Bering Sea summer waters.
From Weingartner et al. 2017a

D1

D. Transport and Thermohaline Variability in Barrow Canyon on the 

Northeastern Chukchi Sea Shelf

Abstract

We used a 5-year time series of transport, temperature, and salinity from moorings at the head of Barrow 

Canyon to describe seasonal variations and construct a 37-year transport hindcast. The latter was 

developed from summer/winter regressions of transport against Bering-Chukchi winds. Seasonally, the 

regressions differ due to baroclinicity, stratification, spatial and seasonal variations in winds and/or the 

surface drag coefficients. The climatological annual cycle consists of summer downcanyon (toward the 

Arctic Ocean) transport of ~0.45 Sv of warm, fresh waters; fall (October – December) upcanyon transport 

of ~-0.1 Sv of cooler, saltier waters; and negligible net winter (January – April) mass transport when shelf 

waters are saline and near-freezing. Fall upcanyon transports may modulate shelf freeze-up, and 

negligible winter transports could influence winter water properties. Transport variability is largest in fall 

and winter. Daily transport probability density functions are negatively skewed in all seasons, leptokurtic

in summer and platykurtic in fall/winter. Kurtosis differences are manifested by transport intermittency 

with possible consequences for shelf-basin exchanges. The climatology implies that the Chukchi shelf 

circulation reorganizes annually: in summer ~40% of the summer Bering Strait inflow leaves the shelf via 

Barrow Canyon, but from fall through winter all of it exits via the western Chukchi or Central Channel. 

We estimate a mean transport of ~0.2 Sv; ~50% less than estimates at the mouth of the canyon. Transport 

discrepancies may be due to inflows from the Beaufort shelf and from the Chukchi shelfbreak, with the 

latter entering the western side of the canyon. 

D.1 Introduction

Pacific-derived waters flowing northward through Bering Strait cross the Chukchi Sea along three 

principal pathways: Barrow Canyon on the northeastern shelf, Herald Valley in the west, and the Central 

Channel over the central shelf. The Pacific inflow is a substantial source of nutrients and carbon (Walsh et 

al. 1989; Codispoti et al. 2005; Nishino et al. 2011), heat (Steele et al. 2004; Shimada et al. 2006; 

Woodgate et al. 2010; Corlett and Pickart 2017), and freshwater (Yamamoto-Kawai et al. 2008) to the 

polar basin. The partitioning of the inflow amongst these channels is poorly understood, although it is 

clear that the transport variations are largely associated with winds over the Pacific sector of the Arctic 

(Winsor and Chapman 2004; Spall 2007; Danielson et al. 2014). Changes in Pacific water fluxes through 

these channels can affect rates and patterns of ice processes, biological production, and the potential 

vorticity (PV) structure of the adjacent shelfbreak and slope. The latter has consequences on shelf-basin 

exchange processes (Spall et al. 2008) and the circumpolar slope circulation (Spall 2013). Herein, we use 

a 5-year time series of moored measurements to examine hydrographic and transport variability at the 

head of Barrow Canyon and to construct a transport climatology using 37 years of regional winds.

Barrow Canyon extends ~220 km from its head, north of Wainwright, to its mouth on the continental 

slope (Figure D1). Bottom depths increase linearly along its length from 50 m to 300 m. Between 

Wainwright and Barrow (~140 km) the canyon is bounded by the Alaskan coast; whereas north of Point 

Barrow, the canyon delimits the western end of the Beaufort Sea shelf. The western wall of the canyon 

adjoins a broad, shallow (40 – 60 m) bench that extends ~80 km westward to Hanna Shoal. In crosssection, the canyon is nearly U-shaped, with the width varying from ~40 km at its head to ~60 km at its 

mouth (based on the distance between the 70 m isobaths on either side of the canyon mouth). The aspect 

ratio (d), the ratio of the cross- to along-canyon length scales, is ~0.3.
From Weingartner et al. 2017a

D2

Figure D1. Bathymetric map of the northeastern Chukchi Sea showing mooring locations. The 2010 –

2012 arrays of 6 moorings are shown with black dots and consist of moorings BC1 (closest to the coast) 

through BC6. Additional moorings, summarized in the discussion, are shown using colored symbols. 

HFRs were located in Point Lay, Wainwright, and Barrow. The striped regions encapsulate that portion of 

the HFR mask containing 50% or more good data.

Most submarine canyons that incise shelfbreaks lie athwart the prevailing along-shelf currents that control 

the canyon circulation (e.g., Allen and Durrieu de Madron 2009). Barrow Canyon shares this feature 

insofar as the flow over the Chukchi/Beaufort shelfbreak is transverse to the canyon axis (e.g., Mountain 

et al. 1976; Pickart 2004; Nikolopoulos et al. 2009). In contrast to most canyons, currents along the 

longitudinal axis of Barrow Canyon are swifter than the transverse flows nearer its mouth and are strongly 

influenced by Bering and Chukchi shelf processes operating to the south. These shelf regions are also the 

source of summer and fall waters comprising the horizontally- and vertically-structured along-canyon 

flow (Pickart et al. 2005; Shroyer and Plueddemann 2012). The contributing water masses include 

moderately salty Bering Sea Water, fresh, warm Alaskan Coastal Water (ACW; Coachman, et al. 1975), 

near-freezing, saline winter-formed waters (WW) derived from ice production, and cool, dilute icemeltwaters (MW). 

The most comprehensive set of long-term transport measurements in Barrow Canyon is by Itoh et al. 

(2012) and Itoh et al. (2013). They used 3 moorings at the mouth of the canyon to estimate seasonal and 

interannual variations in canyon transports of mass, heat, and freshwater from 2000 – 2008. Their 

moorings were deployed at 10 km intervals, from east to west, on the 80, 250, and 150 m isobaths, 

respectively. They estimated a mean average transport of 0.45 Sv northeastward (which with the 

convention adopted in this paper is the downcanyon or positive flow direction) and into the Arctic Ocean. 

This flow opposes the mean northeasterly winds and is a consequence of the Pacific-Arctic pressure 

gradient that is also responsible for the mean northward transport in Bering Strait. Itoh et al. (2013) find 

that the flow varies seasonally and in-phase with the annual cycle of Bering Strait transport; i.e., 

maximum downcanyon transport occurs in summer, and minimum transport occurs in winter. In 
From Weingartner et al. 2017a

D3

accordance with previous studies (Aagaard and Roach 1990; Münchow and Carmack 1997; Weingartner 

et al. 1998; Pickart et al. 2005; Woodgate et al. 2005a; Itoh et al. 2012), they find that transport variations 

are primarily wind-forced, include reversals, have magnitudes of ~0.5 – 1.0 Sv, and durations of one to 

many days.

Our measurements, made at the head of Barrow Canyon ~190 km southwest (upcanyon) of the arrays 

discussed by Itoh et al. (2013), were part of a larger program to examine connections between Barrow 

Canyon, the northeastern Chukchi shelf, and the western Beaufort shelf in a region contemplated for 

offshore oil development. Section 2 describes the data sets. Section 3 examines kinematic features and 

transport variability derived from the observations. These data and regional winds are used to hindcast a 

37-year (1979 – 2015) transport record. This section also examines the hydrographic properties in relation 

to circulation variability. Section 4 examines why our transports differ from Itoh et al.’s (2013) values and 

discusses potential implications of our findings on the Chukchi and Beaufort shelves and slopes. Section 

5 summarizes the paper.

D.2 Methods

D.2.1 Moorings

From August 2010 – September 2012, six moorings (BC1 inshore to BC6 offshore), spaced ~13 km apart, 

spanned the head of Barrow Canyon (Figure D1). Each mooring contained an upward looking 300 or 600 

kHz Teledyne ADCP and a temperature/conductivity/pressure (TCP) recorder (Sea-Bird 37 or Sea-Bird 

19). Moorings BC2 – BC5 included bottom-tracking for measuring ice keel depths and ice drift. Bottomtracking was inactive on BC1 and BC6 because these included directional wave firmware (wave data are 

not discussed here). The TCP and ADCP were housed in a float situated ~4 m above bottom and sampled 

hourly. BC1, nearest the coast, was deployed in a water depth of 31 m with its instruments mounted 1 m 

above the bottom in a Sea Spider frame. This design minimized damage from nearshore ice keels that can 

exceed 25 m depth. After recovering the array in September 2012, we deployed a single mooring (BC2) 

in succeeding years. We will show that BC2 provides statistically valid estimates of along-canyon 

transport, so BC2 data served as a proxy for transport from September 2012 to early August 2015. Table 

D1 lists positions, depths, and period of record for each mooring. There were several gaps in the mooring 

data due to premature failure of the ADCP on mooring BC2. As explained below, we filled some of these 

gaps with other data.

To provide context over a wider area of the shelf, we obtained mooring data supported by other programs. 

The red diamonds in Figure D1 indicate the locations of oil industry-sponsored moorings described by 

Mudge et al. (2015). The yellow circles are three moorings deployed northeast of Hanna Shoal (supported 

by the Bureau of Ocean Energy Management), and the two orange triangles are moorings (courtesy of S. 

Okkonen) spanning the canyon east of Barrow.

D.2.2 High-Frequency Radars (HFR)

During the open water seasons of 2010 – 2014 we measured hourly surface currents (upper 2 m) at 6 km 

horizontal resolution in the northeast Chukchi Sea using ~5 MHz HFRs (CODAR Ocean Sensors). 

Subsets of these data fill gaps in the transport time series. Field sites were located in the villages of Point 

Lay, Barrow, and Wainwright (Figure D1), which allowed collecting surface current data nominally 

within 180 km of the coast. An additional system was deployed at Cape Simpson (Figure D15) in 2013

and paired with the Barrow system to map currents in the western Beaufort Sea and over the mouth of the

canyon. HFR calibrations were applied following (Barrick and Lipa 1986; Kohut and Glenn 2003), and 

HFR data were processed according to Weingartner et al. (2013a).
From Weingartner et al. 2017a

D4

Table D1. Mooring positions and depths for the Barrow Canyon (BC) array and locations of NARR wind 

gridpoints.

2010 – 2011

Mooring

Name

Latitude

(°N)

Longitude

(°W)

Bottom 

Depth (m)

ADCP 

Depth (m)

ADCP

Period of Record

BC1 70.85 -159.67 30.1 29 18 Aug. – 26 Aug.

BC2 70.92 -159.94 52.3 49 17 Aug. – 20 Aug.

BC3 71.00 -160.21 53.3 48.5 17 Aug. – 26 Aug.

BC4 71.06 -160.49 49.3 45 17 Aug. – 26 Aug.

BC5 71.13 -160.79 49.5 45.5 17 Aug. – 26 Aug.

BC6 71.17 -161.07 46.6 43 18 Aug. – 26 Aug.

2011 – 2012

BC1 70.85 -159.67 30.1 29 30 Aug. – 7 Jul.

BC2 70.92 -159.94 52.3 49 30 Aug. – 4 May

BC3 71.00 -160.21 53.3 48.5 30 Aug. – 4 Jul.

BC4 71.06 -160.49 49.3 45 30 Aug. – 19 Jul.

BC5 71.13 -160.79 49.5 45.5 30 Aug. – 3 May

BC6 71.17 -161.07 46.6 43 30 Aug. – 31 Oct.

(2011)

2012 – 2013

BC2 70.92 159.94 52.3 49 9 Sep. – 13 Aug.

2013 – 2014

BC2 70.92 159.94 52.3 49 11 Sep. – 24 Sep.

2014 – 2015

BC2 70.92 159.94 52.3 49 24 Sep. – 10 Aug.

NARR Wind Gridpoints

Geographic Area Latitude (°N) Longitude (°W)

Barrow Canyon (BC2) 70.92 159.94

Bering Strait (A3) 66.33 168.97

Northern Bering Sea (C40) 60.34 169.02

D.2.3 Winds

Wind and sea level pressure (SLP) estimates were obtained from NOAA’s North American Regional 

Reanalysis (NARR) models (Mesinger et al. 2006) at gridpoints near the Barrow Canyon mooring array 

(BC2), Bering Strait, and the northern Bering Sea (Table D1). Both variables have space and time 

resolutions of 32 km and 3 hours, respectively.

D.3 Results

D.3.1 Kinematic Description

All six moorings from the 2010 – 2011 array provided year-long data sets. These data, lowpass-filtered to 

remove diurnal and shorter period fluctuations, are used to describe kinematic features of the flow at the 

head of the canyon. The record-length mean vertically-averaged velocity vectors and variance ellipses 

(Figure D2a) show that between moorings BC1 and BC3, the mean flow was northeastward at between 

0.1 and 0.2 m s-1, with maximum values at BC2. Farther offshore (moorings BC4 – BC6), the mean 
From Weingartner et al. 2017a

D5

Figure D2. A) Vertically-averaged velocity vectors and variance ellipses for the moorings at the head of 

Barrow Canyon. Note the velocity scale change between the mean vector and the variance ellipse. B) 

Time series of the along-canyon currents at 20 m depth at mooring BC2. Both figures are based on the 

August 2010 – August 2011 mooring records.

vectors were east-southeastward and onshore at ~0.1 m s-1. The vectors imply that upon entering the 

western side of the canyon head the mean flow veered counterclockwise by ~25° over a radius of 13 km 

(between BC4 and BC3), so the bulk of the along-shore flow occurred within ~35 km of the coast. Ellipse 

orientations were along-shore polarized at the four inshore moorings (BC1 – BC4; where the topographic 

slopes are greatest) and cross-shore polarized at the two outer moorings (BC5 and BC6). Standard 

deviations of the along-shore flow were three times the mean values at BC1 – BC3 and twice the means at 

BC4 – BC6. On a monthly basis, the principal axes at each site varied by less than 15°. Currents along 

these axes explained >95% of the current variance at moorings BC1 – BC3, >90% at BC4, and >75% at 

mooring BC5 and BC6.

Another perspective on the velocity structure is provided by vertical sections of velocity using an alongand cross-shore coordinate system based on the approximate orientation of the coast between Wainwright 

and Barrow. In this system, positive along-shore or downcanyon velocities, ur, are toward 56°T, and ur < 

0 is upcanyon toward 236°T. Positive cross-shore velocities, vr, are offshore toward 326°T, and vr <0 is 

onshore toward 146°T. We classified the flow as being down- or upcanyon according to the sign of ur at 


From Weingartner et al. 2017a

D8

Table D2. Row 1 is the linear regression model used to compute Barrow Canyon transports (T) from 

mooring BC2. Rows 2 and 3 are the winter and summer generalized linear models (GLM) used to 

hindcast Barrow Canyon transport based on winds listed in Table D1 over the Bering Sea shelf (C40), 

Bering Strait (A3) and Barrow Canyon (BC2). Terms in parentheses are the 95% confidence limits on the 

model coefficients using an integral time scale of 4 days. The wind components (V, U) and lags at each 

site are defined in the footnote.

r

2

Linear regression: transport vs. vertically-averaged along-canyon current at BC2

T(Sv) = 0.019(±0.001)  BC2(cm s-1) – 0.007(±0.04) 0.96

GLM: Mean daily winter (October – April) transport from along- and cross-shore winds1

T = 0.04(±0.038) + 0.021(±0.007)  C40V + 0.067(±0.011)  A3V

 + 0.084(±0.015)  BC2U – 0.038(±0.013)  BC2V

0.59

GLM: Mean daily summer (May – September) transports from winds1

T = 0.56(±0.031) + 0.025(±0.007)  C40V + 0.029(±0.007)  A3V

 + 0.062(±0.013)  BC2U – 0.018(±0.013)  BC2V

0.55

1

BC2U: 67°T; BC2V: 157°T both lead transport by 1 day; A3V: 15°T leads transport by 1.5 days; C40V: 315°T leads 

transport by 2.5 days.

Figure D4. Mean daily along-canyon winds (top) and transport (bottom). The dashed red line on the wind 

and transport time series denotes the record length means of -1.7 m s-1 and 0.19 Sv, respectively.
From Weingartner et al. 2017a

D9

durations varied from 1 – 5 days, in which fluctuation amplitudes were frequently 0.5 Sv or greater, to 

episodes lasting a week or more.

Transport and along-canyon wind variations at monthly and longer timescales are shown with mean 

monthly values (Figure D5) along with their 95% confidence limits. (For both variables these limits were 

constructed by bootstrapping (Efron and Tibshirani 1993) because both have integral time scales of ~4 

days, and neither are normally distributed.) Typical monthly transport amplitudes are ~0.5 Sv, ~2.5 times 

the record length mean.

In addition to month-to-month variability, there is considerable interannual variability. For example, 

downcanyon transport predominated from November 2010 – March 2011, while upcanyon transport 

dominated over the same period in 2012. Although the seasonal variability is largely masked by monthto-month variations, the mean monthly transports of Figure D5 describe an annual cycle (Figure D6a).

The mean summer (May – September) transport was 0.45 Sv downcanyon, with the annual maximum of 

~0.7 Sv occurring in July, whereas the mean transport from October through April was not significantly 

different from zero.

Seasonal transport variability is also reflected in the higher order statistics, as summarized in the summer 

and combined fall and winter probability density functions (pdfs; Figure D7; fall and winter were 

combined because their pdfs are similar.) For example, the summer variance (~0.29 Sv2) is about half the 

fall/winter variance (~0.54 Sv2). In both seasons the transport pdfs are negatively skewed; extreme flow 

events are primarily upcanyon. Skewness is more negative in summer than in fall/winter even though 

upcanyon flows are more frequent and stronger in fall and winter (i.e., 21% of all fall/winter transports 

are ≤-0.6 Sv, but only ~7% of summer upcanyon transports are ≤-0.6 Sv). The seasonal skewness 

differences reflect seasonal changes in the kurtosis. The summer pdfs are leptokurtic; ~76% of all mean 

daily transports are within one standard deviation of the mean, implying that summer transport changes 

are small. The fall/winter (and record length) pdfs are platykurtic, consistent with upcanyon flow events 

being intermittent but more frequent and larger than those of summer. Neither pdf is normally distributed; 

each fails the Kolmogorov-Smirnov goodness of fit test at the 95% confidence level.

Variability in the along-canyon transport is largely wind-forced (Figure D4). The maximum correlation 

accounts for ~47% of the transport variance, using the wind-component projected along 67°T and leading 

the transport by ~1 day. This result is similar to that found by Itoh et al. (2013) at the mouth of the canyon 

and by Gong and Pickart (2015) from an assemblage of synoptic cross-canyon velocity sections. The 

lagged response of the transport may reflect the spin-up time (h/r) for the bottom boundary layer; where h

is the water depth (~45 m) and r the linear resistance coefficient. For a frictional time scale of 1 day, r is ~ 

5 x 10-4m s-1, consistent with estimates from other continental shelves (Brink 1998).

Although the local winds explain nearly half the transport variance, there are prominent deviations from 

this relationship. For example, there is a distinct annual transport cycle (Figure D6a), but the mean 

monthly along-canyon winds for the 2011 – 2015 period showed no seasonality (Figure D6b). Moreover, 

from June through August 2011 the mean flow was ~0.5 Sv downcanyon while the winds averaged ~3 m 

s

-1 upcanyon (Figure D5). In contrast, the average transport was upcanyon at 0.5 Sv from November 2012 

through March 2013 when the mean winds were upcanyon at ~4 m s -1. Such large transport differences 

cannot be ascribed to the small changes in local winds.

One clue to the differential transport responses to the winds is apparent in comparing the regional wind 

and sea level pressure maps for June, July, August, and September 2011 (Figures D8a – d), December 

2012, and January, February, and March 2013 (Figures D8e – h). From June through September 2011, 

winds over the northeastern Chukchi shelf were easterly or northeasterly at 3 – 6 m s-1. In all months 

except September, the flow was downcanyon at rates of 0.2 to 0.7 Sv. Only in September did the flow 
From Weingartner et al. 2017a

D10

Figure D5. Mean monthly along-canyon winds (top) and transport (bottom) are denoted by black lines 

and dots. Crosses and gray lines delineate the 95% confidence limits. Dashed red lines on show the 

record-length means in transport (0.19 Sv) and along-canyon wind velocity (1.7 m s-1).

Figure D6. Panels A and B are the means (filled circles) and standard deviations (error bars) of the 

monthly transport and along-canyon wind velocity, respectively, for 2010 – 2015. Panel C is the mean 

monthly transport (black line and circles) and 95% confidence limits (error bars) based on the January 

1979 – July 2015 GLM hindcast. The gray lines and circles indicate the maximum and minimum values in 

each month. Panel D displays the monthly means (circles) and standard deviations (squares) of the 

pseudostress for the January 1979 – July 2015 period at the three locations used in the GLM.
From Weingartner et al. 2017a

D11

Figure D7. Probability density functions of the mean daily transport in Barrow Canyon for winter (October 

– April; top) and summer (May – September; bottom). Each panel includes the value of the mean, sample 

variance (s

2

), and the skewness (Skw) and kurtosis (Kur) coefficients, computed per King and Julstrom 

(1982). The vertical dashed line delineates the upcanyon from the downcanyon class intervals.

become upcanyon, coincident with the onset of strong northeasterly winds over the northeastern Chukchi 

shelf and northerly winds over the northern Bering Sea. The transport was upcanyon at ~-0.2 Sv in 

December 2012 when the along-canyon winds were northeasterly at ~2 m s-1, but the winds over the 

northern Bering shelf were northerly at 5 – 7 m s-1. Upcanyon transports of ~-1 Sv occurred in January 

and February 2013 as well, when strong northeasterly (northerly) winds of 4 – 6 m s-1 blew over the 

northeastern Chukchi (Bering Strait). The mean transport was nil in March when winds over the northern 

Bering Sea collapsed, although moderate northeasterly winds of 3 – 4 m s-1 persisted over the 

northeastern Chukchi. These results recall those of Danielson et al. (2014), who showed that circulation 

variability in the eastern Chukchi Sea was mediated by northward propagating, down- and/or upwelling, 

long, continental shelf waves generated on the Bering Sea shelf. In addition, other adjustments in the shelf 

wide pressure field arise from coastal convergences and divergences over the entire Chukchi (Winsor and 

Chapman 2004; Spall 2007) and northern Bering shelves (Danielson et al. 2014).

D.3.3 Transport: Hindcast and Climatology

Based on the preceding results and following Danielson et al. (2014), we applied general linearized 

models (GLM) in an effort to improve transport predictability in Barrow Canyon. This approach 

recognizes that it is the along-shore integral of the shelf winds upstream of the point of interest that 

affects the local circulation (e.g., Gill and Schumann 1974). Our GLM uses NARR winds at locations 

listed in Table D1 in Barrow Canyon (BC2), Bering Strait (A3), and the northern Bering Sea shelf (C40). 

We proceeded by first finding the highest correlation for both the rotation angle and time lag by 

regressing the transport against each of the six individual wind components over the entire transport 

record and then applying these lags and angles to the GLM regression (and given in Table D2). Separate 

regressions were made for summer (May – September) and fall/winter (October – April).
From Weingartner et al. 2017a

D12

Figure D8. Mean monthly maps of sea level pressure (contours), wind speed (color), and direction 

(arrows) over the northern Bering and Chukchi seas. The 2011 maps are for: A) June, B) July, C) August, 

and D) September and the 2012 – 2013 maps are for: E) December, F) January, G) February, and H) 

March.
From Weingartner et al. 2017a

D13

The summer and fall/winter regressions (rows 2 and 3 of Table D2) explain 55 and 59%, respectively, of 

the transport variance. Although these seasonal differences are not substantially different from one 

another, the GLM results are a significant improvement over models using only Barrow Canyon winds by 

explaining ~10% or more of the variance. All but the C40 (Bering shelf) model coefficients differ 

significantly from their seasonal counterparts. For example, the summer intercept is almost 30% larger 

than the winter intercept, and both seasons’ intercepts are positive and consistent with downcanyon 

transport being maintained by an along-shore pressure gradient in the absence of wind-forcing. In 

contrast, the magnitudes of all the winter wind coefficients (except C40) are significantly larger than the 

summer coefficients.

We applied the GLM to hindcast mean daily Barrow Canyon transports using NARR winds from January 

1979 through July 2015 and then formed monthly means to construct a climatological annual transport 

cycle (along with 95% confidence limits and the range in mean monthly transports, Figure D6c). The 

climatology suggests that three seasons comprise the annual transport cycle: 1) winter (January – April), 

when the net transport is zero; 2) summer (May –September), when the net transport is ~0.45 Sv 

downcanyon; and 3) fall (October – December), when the net transport is ~-0.1 Sv upcanyon. The ranges 

for summer months imply that individual monthly transports are always downcanyon, whereas the 

fall/winter ranges indicate that individual monthly transports can be upcanyon or downcanyon. The 

hindcast cycle is similar in phasing to the annual cycle computed from the five years of measurements 

(Figure D6a), and the 95% confidence limits on the observed monthly means overlap those of the 

hindcast values.

As a further illustration of the seasonality in wind forcing over the Bering and Chukchi seas, we show the 

seasonal variation in the along-shore pseudostress (the product of wind speed and the along-shore wind 

component) from the northern Bering Sea, Bering Strait, and Barrow Canyon derived from the 37-year 

NARR record (Figure D6d). Although the annual pseudostress cycles at each location are nearly in-phase 

everywhere, the annual amplitudes of the pseudostress at the Bering locations are ~30 – 45 m2s

-2 while 

that for Barrow Canyon is ~5 m2s

-2

. The monthly standard deviations are nearly equal at all locations 

from December through July, and over Barrow Canyon they are virtually constant throughout the whole 

year. At the Bering locations pseudostress variability from September through November is 2 – 3 times 

greater in comparison to other months and/or to Barrow Canyon. These differences suggest that the 

influence of remote winds on Barrow Canyon transports should vary considerably from year-to-year, 

especially in autumn.

Variability in the entire monthly hindcast record (not shown) is dominated by the annual signal, which 

oscillates about a long-term mean value of 0.16 Sv. Monthly transport anomalies derived from the 

hindcast climatology cycle (Figure D9) show that the anomaly magnitudes are large and exceed 0.2 Sv in 

about one-third of the months. Fall and winter anomaly magnitudes are ~0.2 Sv on average, twice those of 

summer. The maximum mean monthly transport is ~1.0 Sv (February 1989), and the minimum is ~-0.7 

Sv (January 2001). Month-to-month changes in anomaly sign show little persistence, although upon 

smoothing with a 13-month running mean, there is a suggestion of anomaly magnitudes of ~0.1 Sv 

persisting for 3 to 10 years. For example, transports between 1979 – 1988 and 1999 – 2006 were ~0.1 Sv 

below the long-term average, while those from 1989 – 1998 and since 2006 were above the mean by a 

similar amount.

As discussed later, monthly transports that are either up- or downcanyon in fall or winter likely have a 

profound influence on the circulation and water masses on the northeastern Chukchi Sea. We thus 

consider the larger-scale atmospheric patterns associated with monthly transport anomalies. Figure D10 

shows maps of atmospheric pressures and surface winds composited using all March and October months 

(representative of the winter and fall, respectively) for which the transport anomalies exceed one standard 

deviation from the climatological monthly mean. The maps in the top row indicate that abnormally large 
From Weingartner et al. 2017a

D14

Figure D9. Monthly transport anomalies (gray) for the 1979 – 2015 period based on the GLM hindcast. 

The thick red curve is the 13-month running mean of the anomalies.

Figure D10. Sea level pressure (contours), surface wind vectors (arrows), and wind speeds (colors) for 

Marches (left) and Octobers (right) in which there was anomalously large upcanyon (top row; contour 

interval of 5 mb) and downcanyon transports (bottom row; contour interval of 2 mb). The maps are 

composites based on the month and years shown to the right of each map, which includes the mean 

monthly transport for those years.
From Weingartner et al. 2017a

D15

winter and fall upcanyon transports are associated with a deep Aleutian Low, centered along 52°N 

between 150°W and 170°W, and a strong Beaufort High. The juxtaposition of these pressure systems 

impels northeasterly (northerly) winds over the Chukchi (Bering) shelves. For months when the 

downcanyon transport is anomalously large, winds are feeble, the Aleutian Low is weak and displaced 

into the Gulf of Alaska, and the meridional pressure gradients over northern Alaska are diminished.

The observed transport pdfs indicated large seasonal differences in the kurtosis, which we further explore 

using the mean daily hindcast transports to compute cumulative probabilities of flow event duration and 

the total number of events for each season for both the up- and downcanyon cases (Figures D11a and b, 

respectively). (Transport events that extend across seasons were assigned to the season in which most of 

the event occurred.)

From the climatological perspective, there are nearly equal probabilities of up- and downcanyon events 

within any given season. Consequently, seasonal differences in kurtosis are associated with seasonal 

differences in event durations and/or transport magnitudes. Differences in event durations are most 

striking in summer, when 75% of upcanyon events are ≤2 days, and only 10% are 4 – 10 days in length 

(with 10 days being the maximum length). By contrast, only ~20% of summer downcanyon events are ≤2 

days duration, and more than 40% exceed 10 days (with the maximum duration being 120 days). Short 

(≤2 days) duration events have smaller mean daily transport magnitudes than events >2 days. Hence in 

summer, the magnitude of the mean daily downcanyon transport is ~2.5 times larger than that for the 

upcanyon transport. Differences in event durations are smaller in fall and winter, when ~75% of up- and 

downcanyon event durations are ≤5 – 6 days, and mean daily transport magnitudes for each event are 

more equal.

D.3.4 Thermohaline Variability

Our description of temperature and salinity variations at the head of the canyon is based on near-bottom 

measurements (within ~4 m of the seabed). Although these may not be representative of the water 

column, they provide a sense of the spatial and temporal scales in water mass variability. We first address 

these scales by examining the potential temperature (θ) and salinity (S) records from moorings BC1, BC2, 

and BC5 for the August 2010 – August 2012 period (Figure D12; we omit the other records to enhance 

clarity). BC5 data are representative of θ and S at BC4 and BC6. The BC3 θ and S records are similar to 

those at BC1 and BC2, but the magnitude of the variations at BC3 is smaller.

Figure D11. The seasonal cumulative probabilities for a) upcanyon and b) downcanyon events based on 

the hindcast transport record. The numbers associated with the seasons in the legend on each graph 

refer to the number of events over the 37-year record.
From Weingartner et al. 2017a

D16

Figure D12. Time series of salinity (top) and potential temperature, θ, (bottom) at moorings BC1 (black), 

BC2 (blue), and BC5 (red) for the August 2010 – August 2012 period.

There is little seasonal variability in salinity at BC5 where the average and standard deviation of S = 32.6 

± 0.3. BC5 temperatures are θ = -1.4 ± 0.7°C, with q near-freezing in all months except between June 

and November. In contrast BC1 and BC2 have distinct seasonal variations. From late November/early 

December through April/early May, θ is at or near the freezing point, and the salinity average and 

standard deviation are ~33 and 0.5, respectively. Consequently, the density of near-bottom waters in 

winter decreases moving offshore. At all locations, but most noticeably at BC1 and BC2, there are 

episodes of warm (-1.5°C<θ<~0°C), salty (≥33.1) intrusions in fall and winter, which are associated with 

upcanyon flows of slope waters (discussed later). There were also periods when very salty (>~33), nearfreezing waters were present at BC1 and BC2 (and BC3) but not at BC5, e.g., February – March 2011 and 

January – February 2012. In fact, the saltiest fractions occur more frequently at BC1 than at BC2, in spite 

of BC1 measuring at a shallower depth. These are cold, salty, recently formed winter waters, most likely 

produced in the latent heat polynyas along Alaska’s northwestern coast (Cavalieri and Martin 1994). They 

are subsequently transported into the head of the canyon (Weingartner et al. 1998) where, as expected, 

they are trapped to the coast by the sloping topography (Chapman 2000).

The cross-shore density gradient reverses sign from June through September, when, in general, salinity 

(and density) increases and temperature decreases moving offshore. This reversal has little to do with 

changes in water properties at BC5 but is instead due to the impressive changes in temperature and 

salinity at BC1 and BC2. Warming and freshening begins in May and proceeds gradually via a series of 

pulses coincident with energetic downcanyon transport events. In 2012 (and perhaps 2010), these pulses 
From Weingartner et al. 2017a

D17

gave way to a sustained period of warm (8°C), fresh (29 – 30) ACW, which persisted at BC1 through 

August and September. The salinity records imply that the cross-shore salinity gradients within the 

canyon attain their maximum values in these months. In August and September 2010, the mean density 

difference between BC1 and BC3 was ~1.4 kg m-3 implying a vertical shear in the along-canyon velocity 

of ~5 x 10-3 s

-1

, assuming thermal wind balance and that the gradient holds over the entire water column. 

Under these assumptions the velocity difference over the uppermost 30 m of the water column between 

BC1 and BC3 would be ~20 cm s-1. If we further assume that this shear was typical for August and 

September 2010, the baroclinic geostrophic contribution to the total along-canyon transport was ~0.15 Sv, 

which is in line with the synoptic observations of Münchow and Carmack (1997), Pickart et al. (2005), 

Gong and Pickart (2015), and the model results of Signorini et al. (1997). By comparison, the measured 

transport for this period was ~0.8 Sv, suggesting the barotropic component accounted for ~80% of the 

canyon transport.

Figure D13 shows the 5-year BC2 temperature and salinity time series along with the mean daily 

transport. Clearly the inferences drawn from the two years shown in Figure D12 regarding the seasonal 

variability in temperature and salinity at BC2 hold over the entire record. The longer record underscores 

Figure D13. Time series of temperature (bottom) and salinity (middle) from mooring BC2 and the alongcanyon transport (smoothed with a 7-day running mean; top). The horizontal line on the salinity plot 

delineates the 33.1 value. The light grey shading encompasses the periods of upcanyon transports that 

contained lower halocline and Atlantic Water signatures (e.g., S>33.1 and θ>-1.6°C). The light yellow 

shading captures periods of dense water (S>33.1 and θ = freezing point).
From Weingartner et al. 2017a

D18

the tremendous interannual variability in water mass properties. For example, the maximum water 

temperature in summer and fall 2013 was ~3 – 4°C cooler than in the other years, and while winter 

bottom temperatures were always close to the freezing point, winter-to-winter variations were primarily 

associated with upcanyon incursions of warmer slope waters. There is also variability in the timing of the 

spring and/or summer arrival of above-freezing temperatures. We assessed this by examining the dates 

when BC2 temperatures first reached 0 and 2°C. (The arrival of water of -1°C is followed within a few 

days by water at 0°C, and the 1.5 and 2°C arrival dates are similarly close). The “early” summers of 2011, 

2012, and 2015 all detected 0°C between 20th and 22nd of June, with temperatures rising quite rapidly 

afterwards. In 2011 and 2012, 2°C water occurred during the first week of August, and in 2015 it arrived 

a full month earlier. In the “late” summers of 2013 and 2014, 0°C water appeared on the 29th and 18th of 

July, respectively. In 2013, temperatures reached 2°C in late September, and in 2014 this benchmark 

occurred in late August. Variations in the spring/summer arrival of warm bottom waters are not obviously 

related to the spring or early summer transports. The average transports from April – June in the “late” 

summers of 2013 and 2014 were ~0.38 Sv and more than double the corresponding averages for the 

“early” summers. Nor did we find a compelling relationship between the May and June ice extent over the 

Chukchi Sea shelf and the timing of warm water arrival. Such a relationship would suggest that solar 

penetration and heating of sub-surface waters was responsible for the difference in arrival times. The lack 

of convincing connections between the timing of the summer arrival of “warm” water and summer 

transport magnitudes or ice extent suggest that the summer warm water signature is established by 

conditions well south of Barrow Canyon and most likely on the Bering Sea shelf.

The fresh, warm waters of late summer/fall routinely terminate with the upcanyon transport of salty 

(>33.1) water with θ between -1.6°C and 0.5°C. Waters with these characteristics are found over the 

continental slope within the lower halocline (140 – 180 m depth) and/or even deeper (~200 m) within the 

Atlantic Layer (Shimada et al. 2005). The grey shaded regions in Figure D13 highlight the more 

prominent packets of upcanyon events, in which lower halocline or Atlantic Water was transported to the 

head of the canyon. Most of these (indeed four of the five highlighted events) occur in the October –

December period. Following these events, temperatures at the head of the canyon decreased to ≤-1°C.

Interestingly, the depth of the slope water source (based on the θ/S characteristics described earlier) does 

not appear to be a function of the magnitude or duration of the upcanyon flow events as suggested by 

Signorini et al. (1997). The most obvious example of this was the prolonged and strong upcanyon flows 

that occurred between December 2012 and February 2013. Throughout this time the upcanyon transport 

had salinities<33.1 and θ≤-1°C. In contrast, the warmest (θ ~-1°C) and most saline (~34.5) waters 

occurred during the sequence of up- and downcanyon flows of October 2014. Hence, the magnitude and 

persistence of upcanyon flow events do not solely determine the source depth of slope waters fluxed into 

the canyon. Other mechanisms, likely occurring over the shelfbreak and slope, must be important in this 

regard.

D.4 Discussion

D.4.1 Transport Comparisons

Our average transport estimate of ~0.2 Sv at the head of Barrow Canyon is in reasonable agreement with 

the modeled mean transport of Spall (2007) of ~0.25 Sv, but it is ~50% smaller than the ~0.45 Sv average 

of Itoh et al. (2013) determined at the mouth of the canyon. We examine the possibility that these 

differences are a result of sampling bias and/or unaccounted transports into the canyon between its head 

and mouth. We first address these possibilities using vertically-averaged means and variances from 

various moorings located in Barrow Canyon and over the northeastern shelf for the 2010 – 2011, 2011 –

2012, and 2012 – 2013 periods shown in Figures D14a, b, and c, respectively. (The plot contains two 
From Weingartner et al. 2017a

D19

Figure D14. Vertically-averaged mean velocity vectors and variance ellipses for currents and winds in A) 

2010 – 2011, B) 2011 – 2012, and C) 2012 – 2013. Red symbols and text are for winds, black are for 

currents outside of the blue box, and blue are for currents within the blue boxes. Gray (black) vectors and 

ellipses in each figure denote record lengths of from 6 to 10 (11 to 12) months duration.
From Weingartner et al. 2017a

D20

different scales for the currents; scales within the blue box are one-half those outside of the box to 

enhance viewing the smaller currents farther from the canyon axis.)

Our array may be biased if there is substantial flow moving northeastward between the offshore end of 

the BC mooring array and the southeast side of Hanna Shoal. This undetected flow would enter the 

canyon north of the BC array. We examine this possibility using results from moorings A, B, and C 

(Figure D14b) that extend northwestward from the end of the BC array to the southern side of Hanna 

Shoal. The means and standard deviations of these records were ~2 cm s

-1 and ~10 cm s-1

, respectively. 

The vectors and the ellipses are oriented toward the head of Barrow Canyon (similar to the statistics from 

moorings BC4 and BC5) at A and C, although the flow direction at B suggests that some of this flow 

might enter the canyon north of the BC array.

At each mooring the mean flow along the minor axis was <1 cm s-1 southwestward but not significantly 

different from zero at the 95% confidence level. This finding is consistent with that of Pickart et al.

(2016), who observed southeastward flow along the south side of Hanna Shoal from synoptic summer 

velocity and hydrographic sections. Collectively these results suggest that: 1) flow over the northeastern 

Chukchi Sea shelf south of Hanna Shoal was funneled toward the head of the canyon (consistent with 

inferences by Weingartner et al. 2005) and 2) that there was little northeastward transport from the central 

Chukchi to the shelf between Hanna Shoal and Barrow Canyon. Both findings comport with the model 

results of Winsor and Chapman (2004) and Spall (2007).

We next examine current data for the shelf east and northeast of Hanna Shoal. At mooring D (Figures 

D14b and c) the mean vertically-averaged flow was not statistically different from zero. There appears to 

be negligible flow overly the weakly sloping shelf between the 40 and 55 m isobaths as measured by 

moorings E, F, and G northeast of the Shoal in both 2012 – 2013 (Figure D14c) and in 2013 – 2014 (not 

shown). Although somewhat ambiguous, these results suggest that this portion of the shelf is not a major 

feed for Barrow Canyon. There is the possibility, however, of net eastward transport between the 55 m 

isobath and the shelfbreak.

A second potential source of bias is in the mooring array at the mouth of the canyon. Recall that Itoh et 

al.’s (2013) transport estimates derive from 3 moorings, deployed at 10 km intervals, from east to west, 

on the 80, 250, and 150 m isobaths, respectively (Figure D15). The gap between their westernmost 

mooring and the 70 m isobath on the west side of the canyon is ~30 km (~5 times the baroclinic radius of 

deformation). This gap is wide enough to permit the ~15 km wide, subsurface, eastward flowing Chukchi 

shelfbreak current carrying winter waters (Corlett and Pickart 2017) to enter the western side of the 

canyon. (Their estimate is tentative insofar as it is based on a compilation of synoptic sections collected 

between the months of May and October in different years.) We expect that the eastward shelfbreak flow 

follows the isobaths around the western lip of Barrow Canyon and continues upcanyon along the western 

wall. Proceeding upcanyon, this transport should diminish as it becomes entrained into the prevailing 

downcanyon flow along the eastern wall. This additional mass transport would be captured by the 

mooring array at the mouth of the canyon, but not the array at the head of the canyon. This interpretation 

is reflected in Shroyer’s (2012) synoptic sections and in the vertically-averaged current statistics from two 

moorings deployed west of Barrow on the 70 m isobath on the eastern (EBC) and western (WBC) sides of 

the canyon (Figures D14a, b, and c). In all 3 years, the vertically-averaged flow at WBC was statistically 

significant and southwestward (or westward in 2012 – 2013) at ~1 – 3 cm s-1. By contrast, along the 

eastern wall of the canyon the mean flow was ~15 cm s-1 northeastward and statistically significant.

Another potential mass source is associated with westward flow over the Beaufort shelf. As seen in 

Figure D15, the HFR-averaged surface currents for September – October 2013 between the coast and the 

40 m isobath (~50 km offshore) was ~7 cm s-1. If this estimate is comparable to the vertically and 

annually-averaged currents, then the western Beaufort shelf contributes ~0.07 Sv into Barrow Canyon
From Weingartner et al. 2017a

D21

Figure D15. Mean September – October 2013 surface current map of the western Beaufort Sea shelf and 

the northern end of Barrow Canyon. The mean winds at Barrow were from the east-northeast (as 

indicated by the arrow) at 3 m s-1. The red squares at the mouth of Barrow Canyon are the locations of 

the moorings discussed by Itoh et al. (2013).

upstream of the canyon mouth. The mean winds on the Beaufort Sea shelf are westward year-round 

suggesting that this shelf transports water into the canyon.

Taken together the subsurface eastward flow along the Chukchi shelfbreak and the westward flow from 

the Beaufort shelf may add ~0.15 Sv to the ~0.2 Sv canyon outflow we measured. If correct, these 

contributions substantially reduce the 0.25 Sv transport deficit between the head and the mouth of the 

canyon. We emphasize, however, that the estimates of the additional mass sources are tentative at best 

and await better resolution.

D.4.2 Consequences of Transport Seasonality

When compared to the transport climatology in Bering Strait (Woodgate et al. 2005b), we find that, on 

average in summer, ~40% of the Bering Strait summer transport is carried into the head of Barrow 

Canyon, but none of the Strait’s transport enters the canyon in fall and winter. This finding implies 

fall/winter flow convergence over the shelf south of Barrow Canyon because the net northward Strait 

transport must be diverted elsewhere, with the Central Channel, Herald Valley, and Long Strait being the 

likely avenues of egress. Although diversion has been inferred to happen in some winters (e.g., 
From Weingartner et al. 2017a

D22

Weingartner et al. 1998) and in numerical models subject to synoptic wind events (Winsor and Chapman 

2004; Spall 2007), our results imply that it is an annual occurrence. If so, seasonal switching of the points 

of entry into the Arctic Ocean of Pacific waters should affect the shelfbreak circulation structure and the 

seasonality of slope/basin exchange. We next consider other implications of this annual cycle.

D.4.3 Hydrographic Implications

The downcanyon advection of warm waters characteristic of summer effectively concludes with the 

upcanyon transport of colder waters in fall. The duration and intensity of these upcanyon flow events may 

be critical in initiating freeze-up on the northeastern Chukchi shelf because the oceanic heat flux 

convergences associated with these events can be substantially greater than the heat loss to the 

atmosphere (Weingartner et al. 2013a). The interplay between up- and downcanyon transports in fall may 

also affect the properties of dense winter waters produced on the shelf. Large volumes of dense water can 

be produced in the extensive latent heat polynyas along the northwest coast of Alaska, which are most 

prominent in December and January (Cavalieri and Martin 1994; Weingartner et al. 1998; Ladd et al. 

2016). The initial polynya salinity depends upon advection from either the south (low to moderately 

saline Bering waters) or from the north (MW and WW) and is a key ingredient in establishing the 

properties of the dense water formed (Winsor and Chapman 2002). Winter circulation can also affect 

dense water production. If the along-shelf flow collapses, dense water residence time in the polynyas can 

be prolonged and thus become denser over the duration of the polynya event. In addition, the absence of 

an along-shelf flow implies that a greater fraction of the dense water is carried cross-shelf by eddies 

(Gawarkiewicz and Chapman 1995) rather than into the canyon by the along-shelf flow. Finally, 

negligible winter canyon transport implies sluggish flow over the northeastern Chukchi shelf, thus 

expanding the time for exchanges between the seabed and bottom waters. Each of these processes affects 

the properties of the dense winter water that ultimately enters the Arctic Ocean.

D.4.4 Shelf Response to Wind-Forcing

Transport variations are largely controlled by local and remote winds. The latter includes Bering Sea shelf 

winds, where the seasonal variations in wind stress are much larger than over the northeastern Chukchi 

Sea. The summer and fall/winter wind-transport regression coefficients were significantly different from 

each other with the summer intercept being ~30% greater than the winter intercept. One reason for this 

difference might be due to the baroclinic contribution to the downcanyon transport, which amounted to 

~20% of the total August and September transport. The baroclinic tendency should weaken through fall, 

and in winter it should either vanish entirely or reverse, as suggested by our observations.

Seasonal differences in regression coefficients may arise from oceanic processes that mediate vertical 

stress divergence in the along-shore momentum balance. Chapman and Lentz (2005) demonstrated that 

for stratified shelves the frictional stress exerted by the bottom boundary layer (bbl) is diminished 

compared to the unstratified case so that stratification enhances the along-shore transport forced by an 

along-shore pressure gradient. The response depends upon the magnitude of the slope Burger number (Bu

= αN/f, where N is the buoyancy frequency and α is the bottom slope). Bottom slopes are ~5 x 10-4 over 

the Bering and Chukchi shelves (and ~10-3 in Barrow Canyon). N varies from ~0.1 s-1 in spring and early 

summer due to heavy stratification associated with ice melt (Weingartner et al. 2013b) and/or runoff over 

portions of the Bering Sea shelf (Danielson et al. 2012) to ~0 s-1 in winter (Aagaard et al. 1981; Aagaard 

et al. 1985). As a result, Bu may be ≥ 0.2 in summer, and Bu ~0 in winter on the shelf. Complete bbl 

shutdown occurs for Bu >1 (MacCready and Rhines 1993), but Chapman and Lentz (2005) find an 

appreciable increase in the along-shore transport for Bu ≥ 0.2. Spatial variations in Bu render 

comparisons with theory difficult, but the seasonal differences in Bu are consistent with the summer 

intercept being larger than that of winter.
From Weingartner et al. 2017a

D23

Chapman and Lentz’s (2005) arguments imply that the summer wind coefficients should be greater than 

those of winter, but instead the winter coefficients are ~1.5 times greater than those of summer. Three 

possible reasons, operating alone or in concert, may account for this difference. First, fall and winter wind 

along-shore pseudostresses over the Bering Sea are greater than those of summer, and so the differences 

may simply reflect our choice of seasons in constructing the statistical model. Second, mobile sea ice 

covers all or portions of these shelves in fall and winter, and the ice-ocean drag should enhance the 

transfer of momentum from the atmosphere to the ocean (Lu et al. 2011; Schulze and Pickart 2012). 

Third, the winds force rapidly propagating, long continental shelf waves in fall/winter when the 

stratification is weak. In summer, a greater fraction of the wave energy should be carried by smaller, more 

slowly propagating coastal-trapped waves. These are more susceptible to coastline scattering as they 

propagate northward, reducing the response of the transport to the winds.

D.4.5 Transport Intermittency 

The intermittent nature of upcanyon and downcanyon transport events is an important feature of the flow. 

Intermittency is seasonally-modulated, as reflected in the seasonal pdfs (Figure D7) and in the seasonal 

flow event duration pdfs (Figure D11). The differences, which arise because summer downcanyon 

(upcanyon) events are of greater (shorter) duration than those of fall and winter, may have consequences 

for shelf-basin exchange.

Upcanyon transport events have a maximum velocity scale of ~0.4 m s-1 (Figure D3e), suggesting that the 

advective time scale for a parcel to travel from the mouth to the head of the canyon is ~6 days. In 

summer, ~5% of upcanyon flow events are ≥6 days, whereas in winter and fall, nearly 20% of upcanyon 

flows exceed 6 days. Thus, the transport of basin waters onto the shallower reaches of the Chukchi shelf 

is relatively rare and largely confined to fall and winter. Based on a comparison of the temperature and 

salinity records from moorings BC2 and EBC (where EBC is ~110 km northeast of BC2, Figures D15a 

and b), we find that only ~1/3 of the upcanyon events that transport slope waters as far as Barrow carry 

them to the head of the canyon.

Intermittency in downcanyon transports may have consequences on the formation of Arctic Ocean eddies. 

Nof (1991) argues that intermittent outflows can lead to unbalanced fluid patches that eventually break up 

into anticyclonic lenses. This mechanism is independent of the stability properties that evolve from quasigeostrophic flows, such as that modeled by Spall et al. (2008) for a steady eastward shelfbreak current 

over the Chukchi/Beaufort slope. Presumably, short duration outflows from Barrow Canyon are more 

prone to generating eddies through imbalance, whereas longer, steadier duration outflows preferentially 

lead to quasi-geostrophic instabilities. Given the marked seasonal asymmetry in downcanyon event 

duration, Nof’s arguments would suggest that current instabilities predominate in summer, and 

unbalanced motions prevail in fall and winter. Conceivably eddy generation sites might also change 

seasonally; unbalanced motions predominate at the canyon mouth while instabilities occur farther 

eastward along the Beaufort shelfbreak. As a final point, the halocline cold-core anticyclones generated 

along the Beaufort shelfbreak contain ~5x1010 m3 Pacific-derived waters (based on a diameter of ~20 km 

and a thickness of ~150 m (Zhao et al. 2014)). A similar volume is transported during fall/winter 

downcanyon events of ~2 days duration, which have a mean daily transport of ~0.3 Sv. This agreement 

may be fortuitous but it suggests consideration be given to the connection between intermittency and 

basin eddies containing Pacific-derived waters.

As noted by Itoh et al. (2014) the summer downcanyon transport is a significant source of heat and 

freshwater to the Arctic basin by feeding the shelfbreak currents that flow eastward along the Beaufort 

shelfbreak and, within the upper ocean, westward along the Chukchi slope (Corlett and Pickart 2017). 

Upwelling and reversals from east to west of the Beaufort shelfbreak flow depend upon the magnitude of 

the easterly along-shelf winds (Schulze and Pickart 2012); however, these winds may not cause a 
From Weingartner et al. 2017a

D24

simultaneous reversal in the canyon transport. For example, in July 2011 strong northeasterly winds 

(Figure D8b) resulted in westward flow and upwelling along the entire Chukchi and Beaufort shelfbreak 

(Spall et al. 2014). The Barrow Canyon transport in July 2011 was downcanyon (Figures D4 and D5), so 

the warm, fresh outflow was presumably carried westward along the Chukchi shelfbreak. Similarly, the 

mean westward surface flow (Figure D15) at the mouth of Barrow Canyon and over the Chukchi and 

Beaufort shelves and shelfbreaks was westward in September and October 2013 when mean winds were 

from the east-northeast at ~3 m s-1 (Figure D14). These winds likely caused weak shelfbreak upwelling 

but not a complete reversal of the subsurface eastward shelfbreak jet (Schulze and Pickart 2012). 

However, the winds were not strong enough to reverse the transport at the mouth (Figure D15) or head of 

the canyon (Figure D5). Schulze and Pickart (2012) find that >65% of all westward wind events having 

wind speeds of 4 – 7 m s-1 result in significant Beaufort shelfbreak upwelling. Using the hindcast record 

we computed the probability of a downcanyon event versus easterly wind speed over the Beaufort slope 

(at the NARR gridpoint located at ~72°N, 152°W). We find that ~35% of all downcanyon transports 

occur under easterly winds when these wind speeds are within the 4 – 7 m s-1range.

The confluence of the westward-flowing shelfbreak current and the downcanyon flow at the mouth of the 

canyon is presumably complicated. The PV structures of these flows differ; the shelfbreak flow has an 

upwelled frontal structure whereas the downcanyon flow is in the downwelled configuration. Upon 

converging at the mouth of the canyon, the different PV fields of both currents must adjust to one another. 

How this adjustment evolves remains to be determined but may involve intense mixing and a variety of 

sub-mesoscale motions. In both the numerical model of Watanabe (2011) and Nof’s (1988) inviscid, 

analytical model the fate of the outflow is sensitive to the PV structure of both the outflow and the 

ambient current field. Interestingly, Nof finds that the outflow could be deflected to the east or to the west 

or even switch back and forth depending on PV structures.

D.5 Summary

We used a 5-year time series of transport, temperature, and salinity measurements obtained at the head of 

Barrow Canyon to describe the seasonality of these variables and to construct a 37-year hindcast of the 

mean daily transport. The annual cycle is in-phase with that in Bering Strait insofar as the maximum 

occurs in summer and the minimum in winter. The climatological Bering Strait transport is northward 

year-round (Woodgate et al. 2005b), but this is not the case in Barrow Canyon. Here the annual transport 

cycle consists of three seasons: a) the May through September or “summer” season in which the transport 

is downcanyon (northeastward and into the Arctic Ocean) and averages 0.45 Sv, b) the October through 

December or “fall” season in which the average transport is upcanyon (southwestward and onto the 

Chukchi Shelf) at 0.1 Sv and; c) the January through April or “winter” season when the along-canyon 

transport is not significantly different from zero.
E1

E. A Quasi-Lagrangian Perspective on the Surface Circulation of the 

Chukchi Sea

E.1 Introduction

This section of the report builds on a number of satellite-tracked drifter studies undertaken in 

collaboration with BOEM and included as a part of this project. These studies were initiated in 2011 

under a Coastal Marine Institute program in partnership with Shell, and from 2012 – 2014 in partnership 

with the North Slope Borough-Shell Baseline Studies Program. The 2011 – 2013 work and analysis 

details were summarized by Weingartner et al. (2015). Additional deployments were made in 2014 by the 

North Slope Borough-Shell Baseline Studies Program with the analysis conducted in partnership with this 

BOEM program and through BSEE.

The 2014 deployments complement the previous efforts in several ways. First, the 2014 deployments 

provide a unique perspective of a broad-scale coastal upwelling event along the west coast of Alaska. 

Such events are likely common in the Chukchi Sea, especially in fall and winter when there are few 

surface current observations.

Second, the drifters were deployed in large clusters (similar to the deployments made in 2012 and 2013), 

which allows us to improve climatological estimates of relative dispersion during the open water season 

on the Chukchi shelf. Dispersion measures how a patch of fluid containing suspended and/or dissolved 

tracers (such as pollutants, chemicals, plankton, and sediments) evolves; i.e., where the “tracer cloud” 

goes and how these particles are mixed into the environment. Relative dispersion quantifies the spread of 

particles about the center of mass of the “patch” (or cluster in our case). This measure, which depends 

upon the initial size of the patch, is due to sub-mesoscale (defined here as being <1 baroclinic radius of 

deformation, Rbc, which is ~6 km in the Chukchi Sea in summer and fall), mesoscale (1 – 4 Rbc) and larger 

scale processes (LaCasce 2008). Although in general we may not be able to identify specific processes 

associated with dispersive “events”, the climatological construct allows a quantifiable estimate of these 

processes. These estimates may then be compared with estimates derived from numerical models, which 

require some parameterization of sub-grid scale processes.

Third, the clustered sampling approach also allows us to examine key kinematical and dynamical aspects 

of the surface circulation. The former includes velocity gradients and the terms comprising the 

deformation tensor. Our dynamical exploration focuses on the relative roles of surface gravity waves, 

winds, and horizontal pressure gradients in driving the near-surface circulation. Understanding the 

relative importance of these three influences should help guide the development of oil spill and/or 

ecosystem models that examine how particles and/or pollutants are advected. The dynamical examination 

is of particular importance in the Chukchi Sea where the geostrophic flow is controlled by both the mean 

pressure gradient between the Pacific and Arctic oceans (Woodgate et al. 2006) and more rapidly varying 

pressure gradients associated with wind-forced coastal convergences/divergences (Weingartner et al.

2013a). The surface circulation results from the interaction of these pressure gradients, the local Ekman 

balance, and the wave field, none of which have been separately quantified to our knowledge. The 

individual contributions may be quite important and may conflict with conclusions drawn from simple 

two-dimensional wind-driven models of continental shelf circulation. These models find that at distances 

>1 barotropic radius of deformation (Rbt, ~150 km in the Chukchi Sea) the surface circulation is 

controlled solely by Ekman dynamics (Csanady 1982). The momentum balance estimates made by 

Weingartner et al. (2013a) indicate that that this is not the case in the Chukchi Sea where the horizontal 

pressure gradients, which appear to be largely barotropic, may be quite large at distances >1 Rbt.
E2

This section will also address two important questions pertinent to understanding mesoscale circulation 

phenomena in the lower reaches of Barrow Canyon (e.g., north of Point Barrow). The first question we 

ask bears on the fate of the outflow from Barrow Canyon. As evident in the drifter records of Weingartner 

et al. (2015) and those shown herein, some of the drifters turn east and some turn west at the mouth or in 

the lower portion of the canyon. In general, the drifters that turn at the mouth of the canyon then move 

along the shelfbreak of either the Beaufort or the Chukchi seas, whereas the others are generally confined 

to the shelf. We hypothesize that winds play a central role in this turning process, which in turn governs 

the fate of the outflow from Barrow Canyon; mesoscale processes such as current instability or current 

separation from curving isobaths could also be involved. Our second question asks if there are differences 

in the level of mesoscale (eddy) energy between the mouth of Barrow Canyon and that of the continental 

slope to the east and west of the canyon. This question is motivated by the fact that the shelf topography 

and width changes moving eastward from the Chukchi Sea, across Barrow Canyon, and onto the Beaufort 

shelf. To the west of the canyon, the shelf is very broad, whereas the Beaufort shelf to the east of the 

canyon is substantially narrower. These transitions in shelfbreak width, coupled with the abrupt 

topographic transition associated with the canyon itself, suggest that eastward-propagating continental 

shelf or coastally-trapped waves may be scattered by these abrupt topographic changes. Although the 

scattering mechanisms are complex (Yankovsky and Chapman 1995; Yankovsky and Chapman 1996), 

they effectively transfer incoming wave energy into other waves, some of which have considerably 

smaller scales than the incident wave. Shorter waves are associated with current variations at similarly 

small scales and this variability may be reflected in the drifter statistics. The lower magnitude group 

velocity of short waves means that these waves are unlikely to propagate far from the canyon’s mouth 

before dissipating. Wave scattering and its attendant mesoscale motions may be one source of the 

enhanced vertical mixing observed at the mouth of Barrow Canyon (Shroyer 2012). We emphasize that 

our data cannot resolve waves or wave-related processes, but spatial differences in current variability 

would be consistent with wave-scattering. The large number of drifter observations collected in the 

vicinity of Barrow Canyon over the several years of this program, provides the ability to address these 

questions.

E.2 Methods

E.2.1 Drifter Deployments

The satellite-tracked drifters were CODE-type drifters (Pacific Gyre Microstar) that measure currents in 

the upper 1 m of the water column. Three of 2014’s clusters were deployed in nearshore waters by NSB 

residents from local villages, including a cluster of 27 drifters deployed offshore of Point Hope on 23 

July, a cluster of 23 drifters deployed offshore of Point Lay on 26 July, and a cluster of 27 drifters 

deployed offshore of Wainwright on 29 July (Figure E1). Additionally, a cluster of 28 drifters was 

deployed on 15 September 2012 near the Burger prospect. Most of these drifters entered Barrow Canyon 

and contributed to the canyon analyses described below. Data processing procedures of all drifters follow 

the procedure described by Weingartner et al. (2015). The drifters in each cluster were deployed rapidly 

so that all were within a few 10s of meters of one another initially. These clusters are used to update the 

relative dispersion relationships for the Chukchi Sea shelf given in Weingartner et al. (2015). In 

evaluating the surface dynamics, we use these clusters in addition to the 2013 deployments made offshore 

of Point Lay and to the northeast and northwest of Hanna Shoal. We note that a limited number of drifters 

deployed over the lifetime of this program included sensors that detected when the drogue separated from 

the surface float. We find that this occurred on a limited number of drifters 60 days after deployment. 

Consequently, calculations are limited to within 60 days of deployment.


E5

the error in our residual velocities. For winds speeds <10 m s-1 the uncertainty in the residual velocity is 3 

cm s-1 and at greater winds speeds the uncertainty increases to ~5 cm s-1. We emphasize that we are not 

precisely calculating the contribution due to various sources of motion. Instead, our calculations are 

meant to provide a qualitative sense of the importance of one source of motion relative to another over the 

course of the cluster’s trajectory.

E.2.3 Wind and Current Variability

To examine the influence of the winds on the waters moving through Barrow Canyon, the downcanyon 

(northeastward) trajectories were visually inspected and classified into 5 groups based on the location 

where the trajectory diverged from its initial downcanyon path (Figure E2). We selected a point of 

interest (POI) along the east side of the mouth of the canyon on the 100 m isobath and located at 71.75°N, 

154°W. This point was chosen because it is the northernmost location of the 100 m isobath as it bends 

from trending northeastward along the canyon to the southeast along the Beaufort shelfbreak. Group I 

consists of trajectories that turn to the west or northwest after passing north of this point. Group II 

trajectories are those that move east of 154°W after passing to the north of the POI at 71.75°N. Group III 

trajectories are those that exit the east side of the canyon south of the POI and remain on the Beaufort 

shelf for 24 hours or more before returning to the canyon (if they do). Group IV turns westward or 

northwestward before reaching 71.75°N. Finally, Group V consists of drifters that turn around and return 

up-canyon. In order for a drifter to be considered in this classification, it must cross the “Start Line” 

(Figure E2), which extends westward across the canyon from Point Barrow. Only drifters that move from 

the southwest to the northeast are considered. A drifter that meets these criteria on more than one 

occasion is considered a valid candidate for the analysis. Approximate dates of the turns made by these 

drifters (or the dates of passing the POI in the case of Group II) are recorded; winds interpolated to each 

drifter track on those dates are used to determine the mean and standard deviation of the wind velocity at 

the time of the turn for each group. 

We compared the current variability at the mouth of the canyon to current variations over the shelfbreak 

to the east and west of the canyon by dividing the region into three geographic boxes. The mouth of the 

canyon is defined to be between 157° and 153°W and from 71.5°N northward to the 3,000 m isobath 

Figure E2. Delineation of the five different classification groups used to examine the fates of drifters 

moving northeastward through Barrow Canyon. A drifter must pass northeastward across the “Start Line” 

in order to be considered in the analysis. The yellow star indicates the location of the point of interest 

(POI) on the 100 m isobath.

I

II

III

IV

V

Start 

Line
E6

(green, Figure E3). The Chukchi slope is defined as between 157° and 162°W and between the 100 and 

3,000 m isobaths (blue, Figure E3). The Beaufort slope is defined as between 153° and 145°W and 

between the 100 and 3,000 m isobaths (yellow, Figure E3). In this analysis, multiple reentries by a drifter 

into the same box are included in the analyses. Drifter and wind velocities were recorded at each time step 

and assigned to one of the three appropriate boxes. We computed the drifter and wind Mean Kinetic 

Energy (MKE) and Eddy Kinetic Energy (EKE) as:

MKE = 0.5[Ui

2 + Vi2

] and EKE = 0.5[ui

2 + vi2

],

where Ui, Vi are the zonal and meridional velocity component means, respectively, in the “ith” box. The 

variables ui, vi are the corresponding zonal and meridional deviations from the “ith” box mean. The 

Barrow Canyon analyses also included a number of 10 m drogued SVP drifters that were deployed in 

2011. These were analyzed separately from the 1 m drogued drifters.

E.3 Results

E.3.1 Hydrographic Background

We begin with a brief description of the hydrography from 2013, because this illustrates the different 

types of stratification likely encountered by the drifters. Figure E4 shows a vertical section orthogonal to 

the coast occupied by glider 191, which transited ~120 km eastward from the Central Channel toward the 

Alaskan coast along 70.5°N between 11 and 14 August 2013.

Figure E3. Map showing the shelfbreak and slope “boxes” used to evaluate mean and eddy kinetic 

energies. These regions are the Chukchi (blue), Barrow Canyon (green), and Beaufort (yellow).
E7

Figure E4. Top panel: map showing location (indicated by red line) of the west-to-east transect of glider 

191 from 10 to 14 August 2013. The bottom three panels are the vertical sections of temperature (top), 

salinity (middle), and density (expressed as sigma-t; bottom) along the transect. The westernmost end of 

the transect is on the left of the figure.
E8

For reasons pertaining to glider navigability, it did not enter very shallow waters, but only came within 

~40 km of the coast. Close to the coast (between 100 – 120 km), the waters were warm (8 – 9oC), fresh 

(<29.5) Alaskan Coastal Water in the uppermost 20 m. Beneath this layer, the waters were cooler (~4oC) 

and saltier (31). A surface front, located at km 100, separated the nearshore surface waters from cooler (7 

– 8oC), saltier (>~30) water. The offshore waters below 20 m depth had temperatures of 5 – 6oC and 

salinities of ~31. The stratification across the section was moderately strong, with the strongest 

stratification inshore. The pycnocline was thin ~10 m and centered at ~20 m depth. Our Ekman 

calculations for the region of the Chukchi shelf south of 70.5°N used a stratification that was similar to 

the offshore region sampled by glider 191.

Figures E5 – E7 show sections taken from transects to the northeast, over the top of, and to the northwest 

side of Hanna Shoal. All sections indicate a strongly stratified water column, which is a consequence of 

fresh, cold meltwaters (MW) in the upper 15 m and saline, near-freezing winter waters below 25 m depth. 

The strong stratification is prevalent throughout this region even across the top of the Shoal (Figure E6).

There are two other noteworthy features of these sections. The first is the suggestion of small-scale 

variability in temperature and salinity within the surface layer. These variations occur over a variety of 

spatial scales and range in size from a few to 10 km or so. They may be MW fronts and/or small eddies 

confined to the mixed layer and are thus capable of affecting the surface circulation. The other 

Figure E5. Vertical sections of temperature, salinity, and σt from an Acrobat transect along the 

northeastern side of Hanna Shoal on 11 September 2013. The western part of the section is on the lefthand side of the figure.
E9

Figure E6. Vertical sections of temperature, salinity, and σt from an Acrobat transect across the top of 

Hanna Shoal on 10 September 2013. The northern part of the section is on the left-hand side of the 

figure.

Figure E7. Vertical sections of temperature, salinity, and σt from an Acrobat transect on the northwest 

side of Hanna Shoal on 10 September 2013. The southeast end of the section is on the left-hand side of 

the figure.
E10

feature of note is that the isopycnals on the northeast and northwest sides of Hanna Shoal (Figures E5 and 

E7, respectively) slope downward away from the Shoal. These slopes are indications of baroclinic 

pressure gradients, which would tend to accelerate the currents counterclockwise around the Shoal and 

opposite to the barotropic pressure gradients that tend to accelerate the flow clockwise around Hanna 

Shoal (Winsor and Chapman 2004; Spall 2007). The model results and the observations suggest that these 

pressure gradients act in opposition to one another as discussed in Section H.

E.3.2 Surface Circulation

We begin this section by examining the cluster of drifters deployed ~20 km offshore of Point Lay on 19 

August 2013. Figure E8 shows the trajectory of the cluster centroid and the individual drifters along with 

time series of the centroid and wind velocities and the mean sea surface temperature. Figure E9 consists 

of time series of the zonal and meridional components of the velocities. Labels along the track and 

associated with the time series in both Figures E8 and E9 refer to distinct events that are discussed. After 

deployment (Period A) the cluster moved to the southwest offshore under the influence of northeasterly 

winds until 25 August. Note that as the drifter began moving offshore, SSTs increased from 6 to 8°C. 

Recall that ~5 days earlier the inshore glider data indicated 6°C water near the bottom and >8°C at the 

surface. One interpretation of the drifter SST measurements is that the near-bottom waters of the previous 

week had upwelled to the surface and the warmer surface water was spreading offshore. Throughout this 

initial movement to the southwest, the net motion was governed by Ekman dynamics zonally and 

geostrophic meridionally. Typical velocities for both of these components were ~10 – 20 cm s-1. The 

winds shifted and became more variable through the end of August. The circulation response included a 

rapid reversal to the northeast due primarily to geostrophy. This response is consistent with the 

northeastward flowing Alaskan Coastal Current (ACC) in this region. The motive force for this flow is

the Pacific-Arctic pressure gradient and the zonal density gradient evident in Figure E4. During Period B 

(9/1 – 9/6), the winds were northeasterly at 5 – 10 m s

-1 and the cluster drifted ~120 km rapidly westward 

and entered the Central Channel southeast of Herald Shoal. The zonal Ekman drift, at 10 – 30 cm s-1, was 

principally responsible for this westward transit. Note that the meridional geostrophic component 

remained northward throughout this period, and was slightly larger in magnitude and tended to cancel the 

southward Ekman component. Once in the Central Channel, the cluster was subjected to weak and 

variable winds. During Periods CC1 (9/7 – 9/15) and CC3 (9/22 – 10/6), northward geostrophic motion 

dominated. The zonal motion was eastward and geostrophic during CC1 while both Ekman and 

geostrophic dynamics governed motion during CC3. Strong (~10 m s-1) northeasterly winds prevailed 

during Periods CC2 (9/16 – 9/21) and CC4 (10/7 – 10/10). During CC2, the net southwestward zonal 

motion was dominated by Ekman dynamics; however, the geostrophic motion was northeastward and 

thus tended to offset the Ekman displacement. In Period CC4, the net westward displacement was due to 

both Ekman and geostrophic dynamics. Over all periods, the Ekman and geostrophic velocity components 

have the largest magnitudes with most falling within the 5 – 20 cm s-1 range. In contrast, the Stokes’ 

current speeds are ≤5 cm s-1. Although these speeds are comparatively small, the Stokes’ drift, when 

integrated over time can lead to substantial displacements. For example, during Period B, the cluster 

moved ~120 km westward in 6 days. The Stokes’ drift accounted for ~20% of this westward 

displacement.

We next consider the drift of the cluster of the drifter cluster deployed offshore of Point Lay on 28 July 

2014. Figure E10 shows the trajectory of the cluster centroid and the individual drifters along with time 

series of the centroid and wind velocities and the mean sea surface temperature. Figure E11 contains the 

time series of the zonal and meridional components of the velocities. The initial stages of this drift are 

quite similar to the 2013 Point Lay deployment insofar as this includes upwelling-favorable winds that 

moved the cluster offshore to the Central Channel followed by northward displacement in the Channel. 

Unlike 2013, the cluster turned eastward toward the coast upon reaching 71°N in the Channel. The other 
E11

Figure E8. Trajectory of the Point Lay 2013 cluster centroid (color-coded by SST) and the 13 individual 

drifters within this cluster (top panel). Additional plots are the time series of the centroid (2nd from top) and 

wind (3rd from top) velocities and SST (bottom). The colored symbols along each trajectory are at 5-day 

intervals and correspond to the date axis on the SST time series plot. Events A, B, and CC1 – 4 are 

discussed in the text.

Herald

Shoal
E12

Figure E9. Time series of the zonal (top) and meridional (bottom) velocity components associated with 

the Stokes (blue), Ekman (red), and geostrophic (yellow) contributions to the surface velocity. The letters 

and vertical lines correspond to the periods in Figure E8 and discussed in the text.

obvious difference between these two clusters is the remarkable degree to which the 2014 cluster 

remained intact compared to the much larger dispersal in 2013.

The cluster headed northeastward at ~10 cm s-1 for the first week following deployment with winds being 

both up- and downwelling-favorable but at speeds of ~5 m s-1. This drift was primarily geostrophic; the 

meridional Ekman currents were weak and the zonal Ekman currents opposed geostrophy initially, but 

then enhanced it during the latter half of the week. The Stokes drift was negligible. Strong northeasterly 

winds of ~10 m s

-1 prevailed throughout Period D (8/3 – 8/15) when the cluster moved ~200 km 

westward in 12 days. The net westward motion was due to a mean Ekman drift of ~20 cm s-1. In contrast, 

the zonal geostrophic flow was eastward at 5 – 10 cm s-1. In the meridional direction, the northward 

Ekman component was largely compensated for by the southward geostrophic component. The Stokes’ 

drift averaged ~4 cm s-1 and was consistently southwestward throughout Period D. Given that the Ekman 

and geostrophic components tend to buffer one another, the Stokes’ drift alone can result in substantial 

displacement of the surface waters over a prolonged wind event.

The westward motion halted on 15 August as the cluster reached the east side of Herald Shoal and the 

winds veered to being southerly at ~6 m s-1 over the next 3 days. The cluster moved eastward over these 3 

days and then it moved ~100 km northward in the Central Channel throughout the 6-day long Period E 

(8/18 – 8/24). Geostrophy dominated during this period with the geostrophic currents being ~15 – 20 cm 

s

-1

. When the drifters reached 71°N on 24 August, the winds switched and became west-northwesterly at

~5 m s

-1

. During Period F (8/26 – 9/5) the cluster moved eastward towards Barrow Canyon and eventually 
E13

Figure E10. As in Figure E8 but for the Point Lay 2014 cluster. Segments D, E, and F are discussed in 

the text.
E14

Figure E11. Time series of the zonal (top) and meridional (bottom) velocity components associated with 

the Stokes (blue), Ekman (red), and geostrophic (yellow) contributions to the surface velocity for the Point 

Lay 2014 deployment. The letters and vertical lines correspond to the periods in Figure E10 and are 

discussed in the text.

through the canyon (not shown here). During this transit from the Central Channel to Barrow Canyon, the 

Ekman and geostrophic components were both eastward and contributed to the zonal velocity. In contrast, 

the meridional Ekman and geostrophic components were equal in magnitude but opposite in direction, 

with the geostrophic velocity being northward. After 1 September, the cluster approached Barrow Canyon 

and the zonal, meridional geostrophic components rapidly increased in magnitude, and the cluster began 

to disperse.

In another deployment, a cluster of 27 drifters was deployed a few kilometers offshore of Point Hope on 

24 July 2014. All the drifters began heading to the northwest but within about 5 days, the cluster split. All 

of the drifters continued west, but about half took a northerly route across Hope Sea Valley and proceeded 

westward along the south side of Herald Shoal (Figure E12). The other half took a more southerly route 

and drifted westward across Hope Sea Valley (Figure E13). Because of this bifurcation, we shall analyze 

the momentum balances of these two “sub-clusters” separately. Before proceeding with that analysis 

though, we consider some of the kinematic aspects of each cluster. The two groups separated from one 

another at about 68.8°N, 169°W. We applied the approach developed by Okubo and Ebbesmeyer (1976)

and used by Weingartner et al. (2015) for clusters from previous years to compute the terms in the 
E15

Figure E12. As in Figure E8 but for the Point Hope 2014 northerly cluster. Segments G, H, and I are 

discussed in the text.

G H I

G

H

I
E16

Figure E13. As in Figure E8 but for the Point Hope 2014 southerly cluster. Segments J, K, L, and M are 

discussed in the text.

J

K

L

M

J K L M

E18

Figure E14. Time series of the four terms comprising the deformation tensor (div = divergence, RV = 

relative vorticity, STR DEF = stretching deformation, and SHR DEF = shearing deformation. The Point 

Hope North cluster terms are in blue and those for the Point Hope South cluster are red. Units are 10-5 s

-1

.

Figure E15. Time series of the terms comprising the absolute vorticity tendencies for the Point Hope 

North (blue) and Point Hope South (red) clusters. From top to bottom these are the rate of change of 

absolute vorticity, stretching vorticity, and the sum of the first two terms. Units are10-9 s

-2

.
E19

feature of interest shared by both clusters is that the SSTs remained remarkably uniform throughout the 

~45 day record. For the most part SSTs were ~6°C. The most prominent temperature changes occurred 

early in the records and towards the end. The first major change occurred after the two clusters separated 

on 29 July, when the SST increased from 6 to 8°C for several days and then decreased to 6°C again as the 

clusters moved westward across Hope Sea Valley. Temperatures also decreased noticeably in the northern 

cluster as it moved northward through Herald Canyon.

The momentum balance for the Point Hope North cluster is shown in Figure E16. Initially, geostrophy 

dominates the northwestward drift of ~10 – 15 cm s-1 during Period G (7/25 – 8/3). The flow continued to 

the northwest during Period H (8/4 – 8/13) with the surface Ekman drift dominating the balance from 6 –

13 August when strong northeasterlies prevailed. The geostrophic currents were also northwestward at 

this time, although substantially smaller than the zonal Ekman flow. These winds abated on 16 August 

and during Period I (8/13 – 8/29), the meridional geostrophic flow moved the cluster northward into 

Herald Canyon; by 29 August, the cluster began dispersing rapidly. The dominant balances for the Point 

Hope South cluster (Figure E17) also included geostrophic flow initially (Period J; 7/25 – 8/3) followed 

by westward Ekman drift during Period K (8/3 – 8/17). During Period L (8/17 – 9/5), the cluster remained 

largely in place, with the geostrophic and Ekman terms variable and largely balancing. On approximately 

9/3, the winds relaxed and the geostrophic flow carried the cluster northward. Period M (9/5 – 9/12) was 

dominated by the northward geostrophic flow. 

Figure E16. Time series of the zonal (top) and meridional (bottom) velocity components associated with 

the Stokes (blue), Ekman (red), and geostrophic (yellow) contributions to the surface velocity for the Point 

Hope North deployment. The letters and vertical lines correspond to the periods in Figure E12 and are 

discussed in the text.

G H I
E20

Figure E17. Time series of the zonal (top) and meridional (bottom) velocity components associated with 

the Stokes (blue), Ekman (red), and geostrophic (yellow) contributions to the surface velocity for the Point 

Hope South deployment. The letters and vertical lines correspond to the periods in Figure E13 and are 

discussed in the text.

Each of the previous clusters discussed evinced impressive coastal upwelling events in which warm, 

dilute nearshore Alaskan Coastal Water was transported substantial distances across the shelf. In the case 

of the Point Lay clusters, this water was brought as far west as the Central Channel; the Point Hope 

cluster suggests that these waters were carried to the western extreme of the Chukchi Sea and into Herald 

Canyon. We next consider the two Hanna Shoal clusters, which were deployed at the northern end of the 

shelf and far from the coast.

The NE Hanna Shoal cluster of 12 drifters was deployed ~50 km east of Hanna Shoal on 10 September 

2013. The trajectory and inferred dynamic balances are shown in Figures E18 and E19, respectively. 

Winds were variable but mainly westward at 5 m s-1 over the first few days as the drifters moved 

northward at 5 – 10 cm s-1 driven by Ekman drift. Period N (9/14 – 10/13) encompasses the time when the 

cluster first began moving onto Hanna Shoal and when it eventually moved off. The movement onto the 

Shoal occurred primarily via the zonal component of the Ekman drift from 9/14 – 9/23, when the winds 

were northeasterly and strong (7 – 10 m s

-1

). As the cluster moved onto the Shoal, the SST decreased by 

~2°C, a change consistent with the temperature section shown in Figure E5. From 9/24 – 10/14, the winds 

were weaker, and more variable. The cluster meandered around the top of the Shoal but slowly drifted 

westward. The circulation dynamics over the Shoal are complicated. Over most of this portion of the 

record, the meridional balance indicates weak and opposing Ekman and geostrophic contributions. In the 

zonal direction, Ekman forcing was the dominant cause of the westward motion, whereas the zonal 

geostrophic tendency was weak and variable and generally opposed this drift. The eastward geostrophic 

tendency is consistent with the mean clockwise flow around the Shoal associated with barotropic forcing 

J K L M
E21

Figure E18. As in Figure E8 but for the NE Hanna Shoal cluster. Segments N and O are discussed in the 

text.
E22

Figure E19. Time series of the zonal (top) and meridional (bottom) velocity components associated with 

the Stokes (blue), Ekman (red), and geostrophic (yellow) contributions to the surface velocity for the NE 

Hanna Shoal deployment. The letters and vertical lines correspond to the periods in Figure E18 and are 

discussed in the text.

(e.g., Spall 2007; Fang et al. in prep) even though this forcing should be very weak atop the Shoal. 

Clearly, there are baroclinic pressure gradients associated with the numerous, shallow frontal features that 

range in horizontal size from the mesoscale (~20 km) to the sub-mesoscale (~1 km). These features are 

evident in the various CTD sections (Figures E5 – E7) around the Shoal and are also documented by 

Timmermans and Winsor (2013) elsewhere in this region of the Chukchi Sea. These shallow fronts are 

expected to respond very quickly to changes in the wind field and some of the noise in our data may 

reflect these variations. By 14 October, the cluster had drifted to the southwest corner of Hanna Shoal. 

Period O (10/15 – 11/1) consists of a generally southwestward drift into the Central Channel, during 

which winds were primarily from the northeast at speeds ranging from 8 – 12 m s-1. That movement was 

largely accomplished by both the Ekman and Stokes drifts. From 10/15 – 10/23, the geostrophic tendency 

was northeastward and afterwards it was southwestward.

The NW Hanna Shoal cluster, deployed 11 September 2013 is the final cluster examined; these results are 

displayed in Figures E20 and E21. During Period P (9/11 – 9/25), the winds were northeasterly from 5 to 

10 m s

-1

, and the cluster moved to the west-northwest with the zonal motion primarily due to Ekman 

forcing and the meridional motion due to geostrophy. Beginning on 23 September and continuing through 

13 October, the winds were weak and variable in direction. For the most part, the cluster continued 

moving westward at speeds of ~10 cm s-1, with this motion being primarily geostrophic. There was a 

particularly strong west-southwest geostrophic flow of ~25 cm s-1 centered around 26 September. The 

coincidence of this event with a rapid change in SST suggests that the cluster may have been entrained in 

a mesoscale feature, perhaps associated with a front. Overall, the meridional displacement during Period 
E23

Figure E20. As in Figure E8 but for the NW Hanna Shoal cluster. Segments P, Q, and R are discussed in 

the text.
E24

Figure E21. Time series of the zonal (top) and meridional (bottom) velocity components associated with 

the Stokes (blue), Ekman (red), and geostrophic (yellow) contributions to the surface velocity for the NW 

Hanna Shoal 2013 deployment. The letters and vertical lines correspond to the periods in Figure E20 and 

are discussed in the text.

P was small and governed primarily by both northward and southward geostrophic motion. During Period 

Q (9/27 – 10/12), the wind speeds ranged from 5 to 7 m s-1 and directions were variable. Over the first 

half of the period, the net drift was toward the northwest at a comparatively sluggish rate. From 3 – 13 

October, the cluster moved rapidly (10 – 15 cm s-1) westward along 73.33°N and over the 100 m isobath. 

Over most of this period, the motion was primarily geostrophic. We offer two possible sources for the 

pressure gradient driving this geostrophic flow. The first is an ice-edge front. Satellite-imagery indicates 

that the ice-edge was zonally oriented in this region, advancing southward and within 50 km of the 

cluster. The other possibility is that the westward flow was part of the southern limb of the clockwiseflowing Beaufort Gyre. Strong northeasterly winds of 10 m s

-1 dominated during Period R (10/13 – 10/19) 

and the cluster moved southwestward and ultimately entered Herald Canyon. The Stokes, Ekman, and 

geostrophic zonal contributions were all westward and of comparable magnitude. In the meridional 

direction, the Stokes and Ekman contributions were dominant and southward.

E.3.3 Dispersion

Following Weingartner et al. (2015), the relative dispersion results are presented in terms of the zonal 

(Dx

2

), meridional (Dy

2

), cross-product or cross-correlation (Dxy) and total horizontal dispersion (D2). The 

results for the Point Lay 2014 cluster (Figure E22) are remarkable insofar as the dispersion variables are 

all small until late August, after the cluster had moved northward through the Central Channel and begun 

moving eastward toward the coast. Zonal dispersion was the dominant term and accounted for ~80% of 

the total horizontal dispersion.

Before examining the dispersion properties for the entire Point Hope deployment, we first compare the 

dispersion characteristics of the northern (Figure E23) and southern (Figure E24) clusters. For the 
E25

Figure E22. Time series of relative zonal (Dx

2

; top left) and meridional (Dy

2

; top right) dispersion, the 

cross-correlation term (Dxy) bottom left), and the relative dispersion (D2; bottom right) for the Point Lay 

2014 cluster. As indicated in the upper left, each panel includes the mean (μ) value and the mean value ±

one standard deviation (s).

Figure E23. Time series of relative zonal (Dx

2

; top left) and meridional (Dy

2

; top right) dispersion, the 

cross-correlation term (Dxy; bottom left), and the relative dispersion (D2; bottom right) for the Point Hope 

northern cluster. Each panel includes the mean (μ) value and the mean value ± one standard deviation 

(s) per Figure E22.

μ+s

μ-s

μ
E26

Figure E24. Time series of relative zonal (Dx

2

; top left) and meridional (Dy

2

; top right) dispersion, the 

cross-correlation term (Dxy; bottom left), and the relative dispersion (D2; bottom right) for the Point Hope 

southern cluster. Each panel includes the mean (μ) value and the mean value ± one standard deviation 

(s) per Figure E22.

northern cluster, all components of the dispersion are very small (≤1,000 km2) until 28 August, which was 

after the cluster had entered Herald Canyon. The dispersion accelerated as part of the cluster continued 

northward in the Valley and the other drifters turned anticyclonically around the north side of Herald 

Shoal. Dispersion accelerated rapidly through 4 September and then reached a plateau of ~15,000 km2

over the last few days of the record. Very little dispersion occurred for the southern cluster, with the zonal 

component being <1,000 km2 throughout the record and the meridional component being ~3,000 km2 at a 

maximum. The cross-correlation term is effectively zero as there was no major eddying or turning 

motions associated with this cluster.

The dispersion calculation for the entire Point Hope deployment (Figure E25) indicates that all the 

dispersion terms began to increase rapidly on 20 August, i.e., when the northern cluster began its 

northward turn into Herald Canyon. What is remarkable about the entire Point Hope cluster is that there 

were only two key transitions in its entire ~45 day life history. The first event was its breakup into two 

sub-clusters as it rounded Point Hope. The dispersion characteristics of the sub-clusters were then very 

similar over the next month. The next major transition was when the northern cluster turned into Herald 

Canyon on 20 August. It is noteworthy that the duration of each of these events was only a couple of days 

and that both events occurred under comparatively mild wind conditions!

E.3.4 Barrow Canyon Mesoscale Analyses

One of our stated goals is to determine if there is a discernible difference in EKE, effectively a measure of 

mesoscale current variance, over the shelfbreak and slope at the mouth of Barrow Canyon compared to 

the EKE levels along the Chukchi and Beaufort slopes and shelfbreaks. The results for the 1 m drogued
E27

Figure E25. Time series of relative zonal (Dx

2

; top left) and meridional (Dy

2

; top right) dispersion, the 

cross-correlation term (Dxy; bottom left), and the relative dispersion (D2; bottom right) for the entire Point 

Hope cluster. Each panel includes the mean (μ) value and the mean value ± one standard deviation (s)

per Figure E22.

Microstar drifters and the 10 m drogued SVP drifters are listed in Tables E1 and E2, respectively. The 

tables also contain the EKE and MKE values for the winds based on simultaneous observations of the 

drifter velocities and the ratio of the ocean EKE to the wind EKE. Before discussing the results, it is 

worth noting that the number of observations in the Barrow Canyon and Chukchi shelfbreak boxes is 

more than twice that for the Beaufort shelfbreak. Also note that the wind MKE levels over all regions are 

quite dissimilar. There is no physical reason why these wind MKE levels should differ and we conclude 

that our Beaufort and Chukchi sample sizes are probably too small or weighted heavily towards a 

particular year or set of events. Bearing this in mind, we find that the current MKE levels are smallest at 

the mouth of Barrow Canyon, four times larger over the Chukchi region, but only twice as large for the 

Beaufort shelfbreak. The EKE levels are quite similar across all locations and differ only by a factor of

~1.5 at most. The EKE/MKE ratio at the mouth of Barrow Canyon is 1.5 – 5 times greater than at the 

other locations. Although these differences are suggestive, the ratio of the ocean EKE to the wind EKE is 

virtually identical at all locations. We would expect that if the oceanic EKE was generated by spatiallyvarying internal processes, then these ratio might vary spatially as well. The SVP results (Table E2) also 

show no compelling differences between the three regions with respect to MKE and EKE levels. Note, 

however, that the sample sizes for the SVP analyses are nearly an order of magnitude smaller than for the 

Microstar drifters so the results are likely biased by the small sample size.

In summary, we find suggestions, but no compelling evidence, that the EKE levels are higher at the 

mouth of Barrow Canyon than elsewhere along the Chukchi and Beaufort slopes and shelfbreaks. We 

emphasize that our analysis has concentrated on surface measurements and that these may not be 

representative of mesoscale variances at greater depths. It would not be surprising if the amplitudes of 

mesoscale currents were much higher at depth than at the surface. This would be expected if, as 
E28

Table E1. Mean and Eddy Kinetic Energy comparisons for 1 m drogued Microstar drifters.

Analysis

Location Currents Winds EKEOceanEKEWind

Number 

of 

Points

MKE

(m2 s-2)

EKE

(m2 s-2)

EKE / MKE

(x104)

MKE

(m2 s-2)

EKE

(m2 s-2) EKE / MKE

Barrow

Canyon 0.004 139 3.6 1 12 13 11 3824

Chukchi 

Slope 0.016 115 0.7 0.3 12 45 9.5 2435

Beaufort 

Slope 0.008 179 2.2 5 15 3 12 1476

Table E2. Mean and Eddy Kinetic Energy comparisons for 10 m drogued SVP drifters.

Analysis

Location Currents Winds EKEOceanEKEWind

Number 

of 

Points

MKE

(m2 s-2)

EKE

(m2 s-2)

EKE / MKE

(x104)

MKE

(m2 s-2)

EKE

(m2 s-2) EKE / MKE

Barrow

Canyon 0.006 202 3.1 2 13 6 15 442

Chukchi 

Slope 0.012 173 1.4 0.6 12 19 14 339

Beaufort 

Slope 0.006 206 3.3 3 9 3 23 258

postulated earlier, the source of the mescoscale variability is associated with baroclinic waves, in which 

case the waves’ amplitudes will be larger beneath the upper layers of the ocean.

Figure E26 summarizes our analysis of the turning of surface water outflow from Barrow Canyon. Recall 

that the turning criteria were based on where a drifter turned with respect to a point of interest (POI), 

which we established at 71.75°N, 154°W, the point where the 100 m isobaths angles southward (Figure 

E2). We find that ~80% of all of the drifters initially turned eastward crossing the meridian of the POI.

Group III (turning eastward to pass the south of the POI) contained 46% of all observations and Group II 

(turning eastward after passing the POI on the north) contained 33% of the observations. Of the remaining 

drifters, 11% passed the POI on the west and then continued westward (Group I), 4% turned westward 

before the POI (Group IV), and 6% reversed their motion and were swept upcanyon (Group V). The 

numbers in this analysis are small and it seems quite likely that these percentages would change 

considerably with a larger sample size that included deployments in different years and throughout the 

fall when the flow in Barrow Canyon is often reversed. Nevertheless, there are two intriguing aspects of 

the results. The first is that even though most of the drifters initially turned eastward and passed the POI, 

they subsequently reversed direction and were transported westward along the Chukchi slope. The other 

point of interest is that Groups II and III segregated according to the magnitude of the zonal component of 

the wind. Group II turns occurred under mild easterly winds (-2.4 m s-1), whereas Group III turns 

occurred under westerlies of ~2 m s-1. Groups I and IV turned westward under easterly winds of 4 m s-1or 

greater. The meridional wind component was weak in all of these cases and appears to have little 

influence on turning. Group V turned under strong northeasterly winds of 9 m s-1. With the possible 
E29

Figure E26. Trajectories and summary statistics for all 1 m drogued drifters according to their turning 

classification. 
E30

exception of Groups II and III, the mean wind vectors cited here should not be used as a turning criterion 

because they derive from a very small set of measurements.

E.4 Discussion and Summary

The addition of the 2014 clusters and re-analysis of some of the earlier Chukchi clusters have revealed 

aspects of the surface circulation dynamics expressed in terms of the relative importance of the Ekman, 

geostrophic, and Stokes’ drift contributions to the velocity field. In general, the Stokes’ drift was always 

the smallest contributor to the surface velocities, although we believe that it cannot be regarded as 

negligible in terms of its potential in transporting pollutants and/or other dissolved and suspended 

materials in the upper meter of the Chukchi Sea. For example, the Point Lay and Point Hope 2014 Stokes’ 

drift estimates suggests that this term alone would have caused a westward drift of 30 – 50 km between 3 

and 15 August. That drift is ~20% of the estimated drift due to the Ekman current at 1 m depth.

Our dynamical analysis provides some insight on why the linear regressions we have applied to the 

drifters in the past have yielded uneven results (Weingartner et al. 2015) including varying from one 

cluster to another or, more specifically, from one portion of the shelf to another. To some degree the 

uneven results are a consequence of spatially varying geostrophic flows. There are several examples of 

these in our data set. For example, in the Central Channel and Herald Canyon the geostrophic flows are 

strong and northward on average (Weingartner et al. 2005; Woodgate et al. 2005; Pickart et al. 2010; 

Weingartner et al. 2013a) and likely include a cross-channel velocity shear. We suggested that submesoscale and mesoscale geostrophic variability might be associated with the patchy remnants of MW

fronts over Hanna Shoal. Similarly, the divergence of the Point Hope cluster shortly after deployment was 

probably associated with mesoscale geostrophic variability such as an eddy or front to the west of Point 

Hope. In the western Chukchi Sea, we noted that the disparate behaviors in the southern and northern 

Point Hope clusters had to be associated with spatial variations in the geostrophic flow. And finally, we 

ascribe the swift westward flow observed along 73°N and over the 100 m isobath as being associated with 

a portion of the Beaufort Gyre. In addition to the spatially-varying pressure gradients, the shelf-wide 

geostrophic field will change in response to local and remote wind forcing (Winsor and Chapman 2004; 

Spall 2007; Danielson et al. 2014). Consequently, these background geostrophic motions, which are 

neither constant in time or space, can be as large as the Ekman drift and lead to degradation of statistical 

relationships between the local currents and the winds.

Mesoscale variability may also affect the relative dispersion characteristics. For example, the total relative 

dispersion of the NE Hanna Shoal cluster reported by Weingartner et al. (2015) reaches 1,000 km2 and 

2,000 km2 by 7 and 9 days, respectively, after deployment. For the clusters reported here and the others 

examined by Weingartner et al. (2015), the 1,000 km2 dispersion value was attained 14 – 31 days after 

deployment, while the 2,000 km2 limit was reached 17 – 36 days after deployment. These differences 

cannot be attributed to the wind variability, as this was similar in all of the cases. Instead, we hypothesize 

that these differences are a function of the sub-mesoscale and mesoscale structure at the ocean surface. 

Although our hydrographic data is limited, it suggests that this variability is greater over Hanna Shoal 

than over the Chukchi shelf south of ~71°N. We will explore this hypothesis in the future by examining 

the spatial variability in sea surface temperature and salinity data collected by underway shipboard 

systems and in the Acrobat and glider CTD data.

The Point Hope and Point Lay 2014 drifter deployments have yielded more than 2,000 drifter days of data 

from the Chukchi Sea, which allows us to upgrade the relative dispersion climatology presented by 

Weingartner et al. (2015). The new climatology, based on clustered deployments from 2012, 2013, and 

2014, is summarized in Figure E27 on log-log plots of the relative meridional, zonal and horizontal 

dispersion. The improved climatology suggests that the dispersion is isotropic insofar as both the zonal 
E31

Figure E27. Log-log plots of the relative meridional (Dy

2

; top), zonal (Dx

2

; middle), and horizontal (D2; 

bottom) dispersion based on the 2012, 2013, and 2014 clusters. The black dashed lines are the 95% 

confidence interval on the mean (blue curve). On the bottom plot, the Lagrangian time scale (TL) is ≤2 

days and denoted by the red, dotted line and the red solid line denotes a dispersion growth rate

proportional to t

2

.
E32

and meridional dispersion values are ~10,000 km2 by day 40. The zonal relative diffusivity is given by Kxx

= 0.5[dDx

2 / dt] (similar to the meridional relative diffusivity). Our new estimates are based on values 

between day 2 and 40, where 2 days is the approximate Lagrangian decorrelation time scale (TL) for the 

drifters. The results suggest relative diffusivities of ~3,300 m2s

-1 over these time scales. The new 

climatology suggests that all components of dispersion grow at a rate proportional to the square of time 

(again after ~2 days) implying that the area occupied by a dissolved or suspended substance is doubling 

roughly every 1.5 days. 

We examined the drifter data at the mouth of Barrow Canyon to test the hypotheses that current variance 

here may be higher than elsewhere along the Chukchi and Beaufort slopes due to internal oceanic 

processes rather than the winds. The results did not support this hypothesis, although this test may have 

failed for two reasons related to sample design. First, mesoscale current variances may have larger 

amplitudes at depth, not at the surface where our measurements were made. Second, our drifter 

deployments were performed in clusters. This “clumped” sampling approach is probably not ideal for 

undertaking the type of hypothesis test that we conducted. A sampling approach that was more broadly 

distributed in time may have been better suited to detecting different levels of mesoscale variability in this 

region.

Finally, we examined the conditions under which surface drifters turned east or west at the mouth of 

Barrow Canyon. We found that the vast majority (80%) of the drifters initially turned eastward upon 

exiting the canyon, although virtually all of these eventually drifted westward. Under westerly winds, the 

drifters turned eastward onto the Beaufort Sea shelf south of the 100 m isobath. Under mild easterly 

winds, the drifters turn eastward north of the 100 m isobath. Future work will examine the wind 

conditions that force the drifters to reverse their eastward motion.
From Weingartner et al. 2017b

F1

F. The Summer Hydrographic Structure of the Hanna Shoal Region on 

the Northeastern Chukchi Sea Shelf: 2011 – 2013

Abstract

We used shipboard and towed CTD, current meter, and satellite-tracked drifter data to examine the 

hydrographic structure in the northeastern Chukchi Sea in August – September of 2011, 2012, and 2013. 

In all years the densest winter water was around and east of Hanna Shoal. In 2012 and 2013, a ~15 m 

deep layer of cold, dilute meltwater overlaid the dense water north of the shelf region between ~71.2 and 

~71.5°N. A front extends from the southwest side of Hanna Shoal toward the head of Barrow Canyon, 

separated meltwaters from warmer, saltier Bering Sea Summer Waters to the south. Stratification was 

stronger and the surface density variances in the meso- and sub-mesoscale range were higher north of the 

front than to the south. No meltwater or surface fronts were present in 2011 due to a very early ice retreat. 

Differences in summer ice cover may be due to differences in the amount of grounded ice atop Hanna 

Shoal associated with the previous winter’s regional ice drift.

Along the north side of Hanna Shoal the model-predicted clockwise barotropic flow carrying waters from 

the western side of the Shoal appears to converge with a counterclockwise, baroclinic flow on the 

northeast side. The baroclinic tendency is confined to the upper 30 m and can include waters transported 

from the shelfbreak. The inferred zonal convergence implies that north of the Shoal: a) near-surface 

waters are a mixture of waters from the western and eastern Chukchi Sea and b) the cross-isobath 

pressure gradient collapses thereby facilitating leakage of upper layer waters northward across the shelf.

F.1 Introduction

Pacific waters flowing northward through Bering Strait and across the Chukchi Sea spread poleward 

across three principal bathymetric depressions: Herald Valley in the west, the Central Channel in midshelf, and Barrow Canyon along the Alaskan coast (Figure F1a). Hanna Shoal lies between the latter two 

features (Figure F1b). Its western side abuts the northern end of the Central Channel, while its eastern 

flank yields to an 80 km wide bench along the western wall of Barrow Canyon. The Shoal is a west-east 

oriented oval ~150 km long and ~55 km wide (based on the 40 m isobath). The shelfbreak (~100 m 

isobath) lies 75 km to the north, and the broad, gently sloping central shelf of 40 – 45 m depth is to the 

south. Minimum depths atop the Shoal are ~20 m and shallow enough to ground sea ice with deep keels, 

as evidenced by the heavily scoured gravelly seabed (Grantz and Eittreim 1979).

Circulation models (Winsor and Chapman 2004; Spall 2007) depict the average flow as northward in the 

Central Channel and Herald Valley (Figure F1a), with the outflows from both proceeding eastward over 

the outer shelf and shelfbreak. Some of this eastward flow is predicted to continue clockwise around the

southeast side of Hanna Shoal before retroflecting eastward. Upon retroflecting, this flow merges at the 

head of Barrow Canyon with eastward flow over the central shelf and northeastward flow along the coast 

(Figure F1a) (Weingartner et al. 2017a; Fang et al. in prep). The models suggest that southward flow on 

the east side and over Hanna Shoal is weaker than the flow along its northern and southern flanks, which 

is consistent with the spreading of the isobaths between the western and eastern sides of the Shoal. Martin 

and Drucker (1997) attributed the weak flow over the Shoal to Taylor column formation and suggested 

that this feature was one reason why ice persists here well after it has retreated elsewhere on the shelf. 

The summertime persistence of sea ice over Hanna Shoal is of considerable ecological importance as the 

area supports a large number of ice-obligate and ice-associated species (Moore and Huntington 2008; 

Moore et al. 2016). These species are sustained by the high benthic biomass found throughout the 
From Weingartner et al. 2017b

F2

Figure F1. A) Bathymetric map of the Chukchi Sea showing principal flow pathways northward from 

Bering Strait. The pathway denoted by the red arrow denotes the Alaskan Coastal Current, which carries 

the freshest and warmest fraction of the Bering Strait outflow and the blue arrows denote the pathways 

for the cooler and saltier fractions. B) Map of Hanna Shoal and the surrounding shelf showing locations of 

2012 (blue) and 2013 (red) CTD stations. C) Location of various current meter and ice profiling sonar 

(IPS) moorings (symbols) and Acrobat towed-CTD transects (blue lines).

region, with particularly large numbers of benthic organisms found on the southeast side of the Shoal

(Blanchard et al. 2013).

This study was motivated by the need to provide observations for evaluating ocean circulation models in a 

region of the Chukchi shelf susceptible to potential impacts from offshore oil development. Moreover, 

Pacific-derived waters undergo substantial modification while crossing this shelf and not all the flow 

pathways and modification sites are understood. In particular, the region around Hanna Shoal has 

received limited attention because heavy ice has often impeded sampling here. We describe interannual 

variations in the regional hydrographic structure of the northeastern shelf in summer (August –

September) and examine the connectivity in the circulation between the west and east sides of the Shoal. 

Section F2 outlines the data sets used, which span the years from 2011 – 2013. In Section F3.1, we 

summarize the summer wind and ice conditions in these years. We next provide an overview of the 

hydrography of the central shelf and Hanna Shoal region with an emphasis on differences between 

summers with light and heavy ice concentration (Section F3.2). Section F3.3 examines in greater detail 

the differences in hydrographic and circulation structure between the western and eastern sides of the 
From Weingartner et al. 2017b

F3

Shoal. Section F4 discusses the results and suggests reasons for the marked differences in summer ice 

concentrations between 2011 and the later years.

F.2 Data and Methods

F.2.1 Hydrography

Three different hydrographic data sets, each averaged into 1 db vertical bins, are used. The first consists 

of shipboard CTDs collected during the Chukchi Offshore Monitoring in Drilling Area (COMIDA) 

program from the USCGC Healy in 2012 and 2013 (Figure F1b). The station spacing varied but was 

typically ~10 km. Time constraints prevented us from conducting extensive CTD surveys south of Hanna 

Shoal during both COMIDA cruises. To place the COMIDA CTD data in a broader spatial and temporal 

context, we use CTD data collected on separate cruises under the auspices of the Chukchi Sea 

Environmental Studies Program (CSESP; sponsored by the oil industry) in August and September of 

2011, 2012, and 2013. The August cruises were confined to small study areas over the shelf south of 

Hanna Shoal, whereas the September cruises enabled broader surveys that extended to the north and east 

of Hanna Shoal. CTD processing followed the procedures of Weingartner et al. (2013a). The August and 

September cruises occurred at identical times in each year, with the sampling progressing from south to 

north across the domain. Although not synoptic, the data underscore the large interannual differences in 

the seasonal hydrography of this region. The third data set consists of several high-resolution (~250 m) 

Acrobat towed CTD and fluorometric sections (Figure F1c) supported by the Bureau of Ocean Energy 

Management (BOEM) and processed per Martini et al. (2016). The sensors were calibrated at the factory 

prior to the cruise and the chlorophyll values are estimated based on this calibration because no 

chlorophyll samples were collected on these cruises. The chlorophyll a concentrations are to be regarded 

as relative, not absolute, values.

F.2.2 Meteorology; Sea Ice Concentration, Drift, and Thickness; and Satellite-Tracked 

Drifters

We used winds and sea level pressures forecast at 3-hour intervals from NOAA’s North American 

Regional Reanalysis (NARR) models (Mesinger et al. 2006). Sea ice concentration maps for June and 

July, from 2011 to 2013, were constructed from data obtained by the Advanced Microwave Scanning 

Radiometer (AMSR-E), Special Sensor Microwave Imager (SSMI) satellite sensor, and the Advanced 

Microwave Scanning Radiometer-2 (AMSR2), respectively and processed according to Spreen et al.

(2008). More detailed maps of the ice edge for the August – September periods of each year were 

prepared based on National Ice Center (NIC) analyses. Through CSESP, the oil industry supported paired 

ice-profiling sonar (IPS) and ADCP moorings to measure ice thickness and drift respectively during the 

winters of 2011 – 2013 at several locations south of Hanna Shoals (Figure F1c). The IPS processing 

follows Melling et al. (1995), and a detailed description of the processing procedures for the IPS and 

ADCP data is given by Mudge et al. (2015). In Section F4 we use the mean ice keel values computed over 

the upper 25% of the keel depth frequency distribution. Following Melling et al. (1995), the latter were 

formed from 50 km long sections based on the ice drift. We then computed monthly means based on these 

values assigning segments that spanned across months to the month in which the majority of the segment 

occurred. The COMIDA program also included current data from five moored ADCPs positioned on the 

northwest and northeast sides of Hanna Shoal, 2012 – 2014 (Figure F1c). We use subsets of these data 

coincident in time with the COMIDA CTD sections. A more in-depth analysis of these and the CSESP 

moorings will be the focus of a future paper.

In addition, we show trajectories from a subset of 13 satellite-tracked Microstar drifters (manufactured by 

Pacific Gyre) equipped with a CODE-type drogue (Davis 1985) at 1 m depth. The drifters were deployed 
From Weingartner et al. 2017b

F4

concurrently near Icy Cape (Figure F1c) and within ~15 km of the Alaskan coast on 12 August 2012. 

Processing details and additional descriptions are given by Weingartner et al. (2015).

F.3 Results

F.3.1 Sea Ice Conditions

Figure F2 shows maps of the shelf-wide distribution of sea-ice concentrations in mid-June and mid-July 

of 2011 – 2013 and time series of the mean daily winds for June and July. In 2011, the ice retreat 

occurred early insofar as the Chukchi shelf south of 71°N was largely ice-free by mid-June (Figure F2a). 

The remainder of the Chukchi shelf and the western Beaufort Sea were ice free by mid-July (Figure F2b). 

Northeasterly winds prevailed in both months (Figure F2c). In 2012, winds varied from northeasterly to 

southwesterly but were from the southeast on average (Figure F2f). By mid-June 2012, the western 

Chukchi had heavy ice concentrations, but a broad swath of the northeast shelf was ice-free (Figure 2d). 

Figure F2. Sea ice concentration maps on 15 June (top) 15 July (middle) and mean daily wind vectors 

(bottom). Panels a, b and c are for 2011, d, e, and f are for 2012, and g, h, and i are for 2013. The 

features labeled in map E are: Herald Valley (HV), Herald Shoal (HeS), Central Channel (CC), Hanna 

Shoal (HaS), and Barrow Canyon (BC). Wind vector plots include mean monthly wind velocity and wind 

speed in parentheses.
From Weingartner et al. 2017b

F5

By mid-July, the ice had retreated northward over the western shelf but advanced southward over portions 

of the northeast shelf (Figure F2e). In addition, distinct embayments formed in Herald Valley, the Central 

Channel, and Barrow Canyon with the ice edge oriented from west to east between the Channel and the 

Canyon along ~71°N. The embayments, discussed by Paquette and Bourke (1981) and Martin and 

Drucker (1997), correspond to the principal flow pathways along which warm Bering Sea Summer 

Waters (BSSW) cross the shelf. These indentations are annually recurring features and were well-known 

to 19th century whalers (Bockstoce 1986). They vary in extent from year-to-year and are separated from 

one another by heavier ice concentrations over Herald and Hanna Shoals (Figure F2e). In mid-June 2013 

(Fig 2g), the ice-edge had retreated to ~70°N over the central shelf, but heavy concentrations remained 

elsewhere. By mid-July 2013 (Figure F2h), the western Chukchi was largely ice-free with open water 

present at the northern end of the Central Channel (along 73°N, between 165° and 170°W) and over 

Barrow Canyon, although heavy ice concentrations remained over Hanna Shoal. In both months winds 

were variable but primarily southerly (Figure F2i).

The subsequent evolution of ice concentrations in the Hanna Shoal region is given by biweekly maps of 

the 15% ice concentration isopleth (as estimated by NIC) for the August-September period of each year 

along with the corresponding mean daily NARR winds for the same period (Figure F3). In 2011, the 

region was virtually ice-free and remained so through September. In 2012, ice remained over Hanna 

Shoal through mid-September. The same basic pattern persisted in 2013, although the ice rapidly retreated 

from south of Hanna Shoal on 1 September to the northwest of the Shoal by mid-September. The ice 

conditions of Figure F3 were not obviously related to the concurrent winds. For example, in both 2011 

(ice-free year) and 2013 (moderate to heavy ice year) the winds were persistently from the northeast and 

moderately strong (~5 – 10 m s-1). In contrast, in 2012 (heavy ice year) the winds were from the south 

through August and then northeasterly through most of September. During the 2012 and 2013, COMIDA 

cruises observations from the Healy bridge suggested extensive grounding of very thick ice atop Hanna 

Shoal, which limited sampling here. Possible causes for these interannual variations in the Hanna Shoal 

ice cover will be considered in Section F4.

F.3.2 Hydrography: Broad-Scale Perspective

We begin by examining plan views of the vertically-averaged upper and bottom 10 m of the water column 

in both August and September 2012 (Figure F4) because we believe that these conditions are typical of 

the Hanna Shoal region (Weingartner et al. 2013a). There are several water mass classifications for the 

Chukchi Sea shelf (e.g., Coachman et al. 1975; Gong and Pickart 2015). For our purposes we define 

BSSW as consisting of a warm (≥4°C) and dilute (~31 – 32) fraction and a colder (2 – 4°C) and saltier 

(32 – 32.5) fraction. The warmer waters are typically carried northward in the Alaskan Coastal Current 

(red arrow in Figure F1a), while the cooler and saltier fraction is transported through the Central Channel 

and Herald Valley. In addition, the shelf contains ice meltwaters (MW), which are cold (≤2°C) and fresh 

(24 – ~30) and winter waters (WW), which are ≤-1.0 °C and salty (>32.5). WW is near freezing when 

produced in winter on the Bering and Chukchi shelves as a result of ice formation and can warm slightly 

in summer months (Gong and Pickart 2015). The disposition of these water masses varies seasonally. 

WW is pervasive in winter but is gradually displaced northward with the springtime increase in transport 

through Bering Strait (Weingartner et al. 2005; Woodgate et al. 2005) and by late summer occurs over the 

northern half of the shelf underlying MW and/or BSSW (Weingartner et al. 2013a; Pickart et al. 2016).

In August, the surface waters from ~70.5° to 71°N and between 164° and 165°W were warm (~6°C) and 

moderately salty (31 – 32). Bottom waters were slightly colder with both water masses consistent with 

BSSW advected northward from Bering Strait in summer. North of this location and south of Hanna 

Shoal, surface waters were cool (2 – 4°C) and fresh (~30), suggestive of MW, perhaps warmed by solar 

radiation and/or mixed with BSSW. North of ~71.25°N bottom waters were the very cold (≤-1°C) and 

salty (≥32.8) WW. The more complete coverage of September 2012 illustrates the enormous spatial 
From Weingartner et al. 2017b

F6

Figure F3. Maps of the 15% ice concentration isopleth at biweekly intervals for the August – September 

period (top row) and mean daily wind vectors (bottom row). Panels a and b are for 2011, c and d are for 

2012, and e and f are for 2013. The mean monthly wind velocity and wind speed (in parentheses) are 

listed below the x-axis for each year.

extent of the surface MW, which occupied the entire region north of ~71.5°N. South of this latitude, the 

shelf contained BSSW, with this water mass extending northward in the Central Channel along 166°W as 

well as eastward toward the coast near Wainwright. Overall these data suggest a front which extends 

southward from 72°N, 166°W (approximately parallel to the Central Channel) and then trending eastward 

along 71.5°N towards Barrow Canyon. Very likely this MW front is part of the same frontal system that 

extends along the length of Barrow Canyon (Pickart et al. 2005; Gong and Pickart 2015). The spatial 

distribution of WW along the bottom in September was similar in extent to the surface MW, although 

WW extended farther inshore than MW. Note also that the coldest WW formed an arc along the southern 

and eastern sides of Hanna Shoal. The 2013 CSESP hydrography (Weingartner et al. 2014) was broadly 

similar to that of 2012 in terms of the areal distributions of the major water masses and location of the 

MW/BSSW frontal system along ~71.5°N. The major difference was that bottom salinities were slightly 

fresher (32.5 – 32.8) compared to 2012.

Water mass distributions in 2011 were quite different in several important respects (Figure F5). Most 

obvious was the complete absence of MW, which is not surprising given the early retreat of sea ice in 

2011 (Figures F2 and F3). The surface layer in both August and September consisted entirely of BSSW, 

was nearly homogeneous in salinity, lacked surface fronts, and had only weak meridional thermal 

gradients. In August 2011, the bottom waters between 70° – 71°N and 164° – 165°W were BSSW, while 

WW occupied the shelf north of this location and along the southern side of Hanna Shoal. In September, 

BSSW was prominent at the surface and bottom in the Central Channel, over the shelf south of ~71°N, 

and was even found over the shallowest portions of Hanna Shoal, although the low salinities here could 

have been a remnant of MW that had either mixed with the BSSW or had warmed. Except as noted with 

respect to the Central Channel, WW surrounded most of Hanna Shoal. It extended as far south as 71°N 
From Weingartner et al. 2017b

F7

Figure F4. Plan views of temperature (left) and salinity (right) in August and September of 2012 based on 

averages of the upper and bottom 10 m of the water column. The village of Wainwright is designated with 

a “W”. Filled circles show CTD station locations.
From Weingartner et al. 2017b

F8

Figure F5. Plan views of temperature (left) and salinity (right) from August and September 2011 based on 

averages of the upper and bottom 10 m of the water column. The village of Wainwright is designated with 

a “W”. Filled circles show CTD station locations. The black line in panel c shows the location of the 

vertical sections from 2011 and 2102 shown in Figure F6.
From Weingartner et al. 2017b

F9

with the coldest waters forming a prominent lobe on the southern side of Hanna Shoal. A thermohaline 

bottom front separated this WW from the adjacent BSSW. The WW properties in 2011 were fresher 

(~32.5) than the corresponding WW salinities ≥32.8 of 2012 (Figure F4h).

An alternative perspective of the contrasting thermohaline structure between these two years is given by 

Figure F6, which consists of two, 300 km long vertical sections occupied in September 2011 and 2012. 

The transect (shown in Figure F5c) extended from the southwest (~70.5°N, 166°W) to the northeast 

(~72.5°N, 160°W) with the northern portion of the transect along the eastern side of Hanna Shoal. In 

2011, warm, moderately salty BSSW entirely encompassed the upper 20 m of the water column and along 

the bottom over the first 100 km. WW was present below ~20 m on the northern and southern sides of 

Hanna Shoal and was separated from BSSW in the south by a bottom temperature front. In contrast, the 

2012 section consisted of a strong frontal system associated with MW and BSSW at km 140, in addition 

to weaker MW fronts at km 175 and km 200. The sections also differ with respect to the vertical 

stratification, which in both years was dominated by the vertical salinity gradient. The stratification in 

2012 (and 2013) was nearly twice as strong as in 2011, with these differences primarily due to the 

absence of MW in 2011 and its presence in 2012 (and 2013).

Another notable feature in the 2012 section was the presence, at km 225, of a “lens” of warm (~2°C) 

water, signified by an upward distension of the 29 isohaline above the lens and a less prominent, 

downward-bowing of the 32 isohaline below. This feature is likely an anticyclonic, intrapycnocline eddy 

Figure F6. Vertical sections of potential temperature (top row) and salinity (bottom row) collected from 

early to late of September 2011 (a, b) and 2012 (c, d), respectively. The location of both sections is 

shown in Figure F5c. Hanna Shoal is denoted as “HS”. Inverted triangles indicate station locations. 
From Weingartner et al. 2017b

F10

formed via a baroclinic instability of the MW/BSSW front. These features are often observed when 

MW/BSSW fronts are present and are probably important agents for lateral mixing (Lu et al. 2015).

The coarse resolution of the CSESP CTD sampling did not permit adequate horizontal resolution of the 

MW/BSSW fronts but Acrobat sections collected in September 2013 along Legs Q and M (Figure F1b) 

resolve the frontal width scale. In both cases, the front was ~10 km wide and located at ~71.5°N on Leg Q 

and 71.2°N on Leg M (Figure F7a, b, c, and d). We examined the stability of these fronts using the 

balanced Richardson number criterion (Thomas et al. 2016) and found that portions of the MW/BSSW 

front (including some sections not presented here) met the necessary conditions for symmetric instability, 

whereas other segments did not. The results are equivocal and, at best, suggest that there was spatiotemporal variability along the front with respect to this criterion.

The chlorophyll distributions (Figure F7e, f) along these sections are interesting in two regards. First, the 

densest chlorophyll concentrations occur in the highly-stratified region north of the front on the 26 σθ

isopycnals at ~25 m depth, and below the pycnocline. Second, the chlorophyll distribution appears very 

patchy with the patch scales ranging from ~1 to ~20 km based upon visual inspection. Nutrient 

concentrations, reported by Codispoti et al. (2005), Questel et al. (2013), and Danielson et al. (2016)

indicate that WW bottom waters have sufficiently high nutrient levels to sustain primary production, 

while MW is nutrient deficient. Although high in nutrients upon leaving Bering Strait, BSSW nutrient 

levels are depleted by the time these waters reach the northeastern Chukchi shelf in August and 

September. Consequently, neither the absence of chlorophyll in the surface waters or the sub-surface 

chlorophyll maxima in the heavily stratified MW/WW is surprising. There is, however, no indication of 

chlorophyll patchiness or of well-defined sub-surface chlorophyll maximum in the less-stratified regions 

south of the front where BSSW overlies WW.

These (and other) Acrobat sections suggest that the surface layer containing BSSW south of the front is 

horizontally homogenous in contrast to the more heterogeneous surface layer north of the front. Indeed, 

Figure F7 suggests that north of the main front, there were weaker and equally shallow fronts composed 

solely of MW of varying salinities and temperatures. We quantified these differences by computing

Figure F7. Vertical sections of potential temperature (a, b), salinity (c, d), and fluorescence expressed as 

μg l-1 of chlorophyll a (e, f) in September 2012 along Leg M (left) and Leg Q (right). The white contours 

indicate σθ isolines. The arrows at the top of each figure indicate latitude 71.5°N. The locations of these 

transects are shown in Figure F1c. The front along Leg M is located at 71.2°N as indicated.
From Weingartner et al. 2017b

F11

the wavenumber spectra of density at 10 m depth (similar to Timmermans and Winsor, 2012). The spectra 

were computed north and south of the front from different acrobat sections (Figure F1b) but not across the 

front. We then integrated the spectra over three different spatial bands (1 – 5 km, 5 – 10 km, and 10 – 20 

km) and compared the variance in each band. Regionally the baroclinic radius of deformation is ~5 km, 

so we define the mesoscale portion of the spectrum as having length scales of >5 km and the submesoscale range <5 km. The computations were done separately using sections north and south of the 

front and the results summarized in Figure F8. South of the front, the variance was uniformly low across 

the entire spectral range, with typical values being 1 – 2 kg2m-6. North of the front, the variances 

increase, although not consistently across all the sections. In general, the variances north of the front were 

a factor of 2 to 5 times greater in these bands than the variances south of the front. Quite possibly this 

greater spatial variability includes vertical motions over similar spatial scales that contribute to the 

patchiness in the chlorophyll distributions.

F.3.3 Hydrography: Hanna Shoal

In this section we examine the temperature and salinity structure around the Shoal based on CTD sections 

occupied during the 2012 and 2013 mid-August COMIDA cruises. In both years heavy ice and other time 

constraints limited the spatial coverage. We occupied five sections in 2012 and four sections in 2013

(Figure F1b), with each section radiating outward from the Shoal. The 2012 sections included three 

sections (1-12, 2-12, and 3-12) arrayed from west to east on the northern side of the Shoal (Figure F9) 

where ice concentrations were 60 – 80%, except along the inner portion of section 1-12 where 

concentrations were ~30%. Winds during the 2012 sampling were southerly at ~4 m s-1. Under these 

conditions the flow over the shelf south of the front was likely eastward at 5 – 10 cm s-1, veering 

northeastward and becoming swifter along the coast (Weingartner et al. 2013a) while the surface currents 

east of Hanna Shoal were weakly eastward at 2 – 5 cm s-1 (Weingartner et al. 2017a; Fang et al. in prep). 

Similar sections were occupied in 2013 (Figure F10, but note that section 2-13 differs in location from 

section 2-12, Figure F1a). Winds during the 2013 sampling were northeasterly at ~4 m s-1, in which

Figure F8. Histograms of density variance at 10 m depth as a function of horizontal length scale based 

on 2013 Acrobat CTD tows collected north of the front (gray stippled bars) and south (black bars) of the 

front. The horizontal bars denote the mean density variance in each band for the gray bars.
From Weingartner et al. 2017b

F12

Figure F9. Vertical sections of potential temperature (θ; top) and salinity (S; bottom) along transects 1-12 

(a, b), 2-12 (c, d), and 3-12 (e, f) on the northwestern, northern, and northeastern sides of Hanna Shoal, 

respectively. Figure F1 shows transect locations. Hanna Shoal is on the left of each panel.

Figure F10. As in Figure F9 but for transects 1-13 (a, b), 2-13 (c, d), and 3-13 (e, f) on the northwestern, 

northern, and northeastern sides of Hanna Shoal, respectively. Figure F1 shows transect locations. 

Hanna Shoal is on the left of each panel.
From Weingartner et al. 2017b

F13

currents should also be eastward south of the front at ~5 cm s-1 (Weingartner et al. 2013a), but westward 

at ~5 cm s-1 at the surface east of Hanna Shoal (Fang et al. in prep). Thus we feel that in both years the 

sections were sampled synoptically during the occupation of the COMIDA transects. Ice concentrations 

along these sections were similar to those of 2012, except the concentrations were generally 0 – 30% 

along section 1-13. Figure F11 shows sections (4-12, 5-12, and 4-13) on the southern side of the Shoal, 

where there was heavy ice along the northern part of each section.

Several features are common to all sections. First, below ~20 m depth the water column consists entirely 

of WW with salinities >32.5 and temperatures ≤-1.6°C. In 2012, salinities within 10 – 15 m of the bottom 

exceeded 33, with maximum salinities (>33.4) near the bottom on section 1-12 (Figure F9b). Bottom 

salinities rarely exceeded 33 in 2013. Second, in general the surface layer consisted of MW with low 

salinities (<30) and temperatures between -0.5°C and 0.5°C. Third, along the northern side of the Shoal 

and proceeding from west to east, the salinities in the upper 10 m decreased by ~2. Fourth, the 2-layer 

stratification of the water column was set by a strong halocline centered between 10 and 15 m depth.

There are a number of exceptions to these general features. First, warmer (1 – 2°C), near surface waters 

were observed south of Hanna Shoal at the southern end of sections 4-12, 5-12, and 4-13 (Figure F11). 

Each of these sections extended as far south as ~71.5°N, suggesting that the warm water source was 

BSSW intruding across the front (Figure F4).

Second, warm (≥0°C) surface or near surface waters were found within the first 35 km of sections 1-12 

(Figure F9a) and 1-13 (Figure F10a) along stations nearest to the Shoal. In both cases, the warmest water 

was in a subsurface temperature maximum embedded in the pycnocline. These warm waters were likely 

BSSW, albeit modified by MW, which had flowed northward through the Central Channel and had

Figure F11. As in Figure F9 but for transects 4-12 (a, b), 5-12 (c, d), and 4-13 (e, f). Panels a, b, e, and f 

are on the southern side of Hanna Shoal, and panels c and d are on the southeastern side. Figure F1 

shows transect locations. Hanna Shoal is on the left of each panel. Note the temperature scale differs 

from that of Figures F9 and F10.
From Weingartner et al. 2017b

F14

begun to flow clockwise around the northern side of the Shoal. This interpretation is consistent with 

velocities at moorings HSNW-40 and -50 in 2012 where the mean flow was ~7 cm s-1 toward the eastnortheast during the time that sections 1-12 to 3-12 were occupied. In 2013, this flow averaged ~5 cm s-1

eastward over the time required to occupy sections 1-13 to 3-13. (In both years the flow was largely 

barotropic insofar as the velocity shears were small; the differences between the near-surface and nearbottom ADCP bins were ~3 cm s-1

.) However, subsurface temperature maxima were absent along the 

sections further east. In 2012 the warmest waters along 2-12 (Figure F9c) were at the distal end of the 

section, whereas on 2-13 the warmest water was adjacent to the Shoal. We argue later that these warm 

signals are likely a mixture of the waters from the western and northeastern sides of the Shoal.

Third, surface waters over the outer parts of sections 3-12 and 3-13 were also warm and fresh, suggesting 

that these were either MW that had warmed by solar radiation or were derived from the Alaska Coastal 

Current emanating from Barrow Canyon. Corlett and Pickart (2017) reported the presence of coastal 

waters flowing westward in the Chukchi Slope Current. Here we show that the coastal waters can, at least 

occasionally, be transported from the Chukchi Slope Current onto the shelf northeast of Hanna Shoal and 

thence continue westward on the shelf and/or southward over the Shoal. Figure F12 shows the sea surface 

temperature (SST) color-coded trajectories for 6 of the 13 total drifters between 12 August 2012 (when 

they were deployed in the Alaskan Coastal Current near Icy Cape) and 30 September, along with the 

winds for the same time period. Upon deployment, the winds were southerly and the drifters moved 

rapidly through Barrow Canyon. At the mouth of the canyon, all but one eventually moved westward in 

the slope current. Six of the drifters (those shown in Figure F12) crossed back onto the shelf northeast of 

Hanna Shoal after the winds became northeasterly on ~31 August, while the remaining six continued 

westward along the slope. SSTs decreased from ~10°C at deployment to 2 – 3°C by mid-September on 

the shelf north and east of Hanna Shoal. For those drifters that continued farther south, the SST decreased 

to ~≤0°C by month’s end. The cooling was most likely due to mixing between the coastal water and MW 

along the drifter path and, beginning mid-September, due to heat loss to the atmosphere.

The final noteworthy difference among the COMIDA CTD sections is that the slopes of the subsurface 

isopycnals differ around the Shoal. On the south side of the Shoal the isopycnal slopes were nearly flat 

indicating very little baroclinicity. On the north side and below the pycnocline, the isopycnals on sections 

1-12, 3-12, 1-13, 2-13, and 3-13 slope downward away from the Shoal. The largest slopes were below the 

pycnocline along sections 3-12 and 3-13, both on the northeast side of Hanna Shoal. As shown below 

these slopes imply a baroclinic velocity field that likely plays an important role in the circulation north of 

Hanna Shoal.

F.3.4 The North Side of Hanna Shoal

In this section we seek to understand the source of the warm surface waters on the north side of the Shoal, 

specifically along sections 2-12 and 2-13. Recall that these waters are cooler than those on either the 

northwest or northeast sides of the Shoal. We first consider the possibility that these waters simply 

resulted from the cooling of warm waters from the Central Channel flowing around the northwest side of 

the Shoal across sections 1-12 and 1-13. We examine this possibility by computing the horizontal heat 

flux divergence of the warm water band over the upper 15 m and over a horizontal width of 30 km 

(essentially the width of the warm water flowing from the Central Channel) and assume only vertical heat 

exchanges occur, to either the ice or to the water column below 15 m. The calculation uses the velocity 

averages given previously and also assumes the warm water at the outer end of section 2-12 entirely 

captured the remnants of the warm waters flowing through section 1-12. In 2012, the heat flux 

calculations imply a heat loss to the ice of ~20 W m-2, for an ice melt rate of 0.6 cm day-1. The advective 

time scale between sections 1-12 and 2-12 is ~10 days, which would result in the addition of 6 cm of MW

to the upper 15 m. Assuming a sea ice salinity of 5; this MW flux would reduce the salinity by 0.1 

between the two sections, which is far less than the observed salinity decrease of ~0.8. 
From Weingartner et al. 2017b

F15

Figure F12. The trajectories of 6 satellite-tracked drifters (1 m drogue depth) deployed offshore of Icy 

Cape on 12 August 2012. One of the trajectories is color-coded according to measured SST. The various 

symbols along each trajectory are at 7-day intervals with the dates given on the wind vector time series 

(bottom plot). The 20, 40, 50, 100, and 1,000 m isobaths are labeled. Green triangles signify the 

deployment position and red stars the last good transmission from each drifter.

Moreover, the additional ice melt influx required to satisfy the salinity change implies an atmospheric 

heating rate of ~185 W m-2. For a surface ice albedo of 0.6, the required solar radiative influx amounts to 

~300 W m-2. These values seem unrealistically high even in the absence of the dense fog present in the 

area at the time. For example, on a monthly basis, Maykut (1986) estimates that the net radiation balance 

in mid-summer is ~100 W m-2. The same calculation for the 2013 data yields a similar conclusion, e.g., 

the solar radiative influx is much greater than feasible. We conclude the salinity imbalance implies an 

advective contribution of MW from elsewhere, most likely from the northeast side of the Shoal.

There are two lines of evidence to support this contention. The first follows from the θ/S diagrams 

compiled from sections 1, 2, and 3 from each year (Figure F13) over the upper 30 m of the water column. 
From Weingartner et al. 2017b

F16

Included in this figure are the average θ/S values in the upper 15 m encompassing the warm waters along 

each section. Clearly the surface water properties along section 2-12 could be obtained by cooling and 

mixing of the surface waters from the eastern and western sides of the Shoal. In 2013, the warm band 

along section 2-13 was ~0.3°C cooler than that on the northeast side of the Shoal, and the salinities were 

identical at ~28.7. In contrast the salinity on the northwest side of the Shoal was, at ~30.8, much greater. 

The second line of evidence follows from the baroclinic geostrophic tendencies along sections 3-12, 2-13, 

and 3-13 (Figure F14). (Baroclinicity along section 2-12 is weak and variable and, thus, not shown.) In 

2012, the baroclinic tendency was northwestward along section 3-12, and, over the outer half of the 

section, averaged ~5 cm s-1 in the upper 20 m. Along the 40 m isobath, the tendency was southeastward at 

Figure F13. Scatterplots of potential temperature versus salinity over the upper 30 m of the water column 

for sections 1, 2, and 3 in 2012 (left) and 2013 (right). Only stations along the first 65 km of the transect 

are included. Note the scales change between 2012 and 2013. The black slanted line denotes the 

freezing point curve. The larger symbols outlined in yellow depict the mean values along each section in 

the upper 15 m.

Figure F14. Baroclinic, geostrophic velocity field for transects a) 3-12, b) 2-13, and c) 3-13. The 

corresponding temperature and salinity sections are shown in Figures F9 e, f; F10 b, d; and F10 e, f, 

respectively. Positive velocities are north-northwestward and negative values are south-southeastward.
From Weingartner et al. 2017b

F17

~5 cm s-1. These values are in good agreement with currents measured by the moorings deployed during 

occupation of the section. During the first three days after deployment, the currents in the upper 20 m 

were west-northwest currents at 6 and 13 cm s-1 over the 50 and 56 m isobaths, respectively. There was 

similar agreement between the computed and measured velocity shears as well. The velocity difference 

over the 25 m between 20 and 45 m depth at the two moorings were ~10 cm s-1 compared to ~6 cm s-1

based on the thermal wind estimates. The mooring deployed on the 40 m isobath recorded 3 cm s-1 toward 

the west- southwest. In 2013, the geostrophic computations yielded results similar to those of 2012 

between the 40 and 60 m isobaths. The calculations also indicate that the baroclinic tendency increased to 

~10 cm s-1 over the deeper part of the section. The moored measurements averaged over the upper 25 m 

of the water column during the transit of this section in 2013 section were ~5 cm s-1 to the southwest over 

the 40 m isobath and ~10 cm s-1 northwestward on the 50 m isobath (the ADCP moored on the 56 m 

isobath died a month earlier). The baroclinic contribution was weaker along section 2-13 (north of Hanna 

Shoal) although still westward. In summary our results suggest that there was zonal convergence, at least 

along the 50 m isobath and over the upper portion of the water column, north of Hanna Shoal in August 

of both 2012 and 2013.

F.4 Discussion

Our analyses clearly highlight the importance of sea ice, MW, and dense water to the hydrographic 

structure of the northeastern Chukchi Sea, particularly in the Hanna Shoal region. MW was absent in 

2011, consistent with early ice retreat and the absence of ice over the Shoal in August. The heavy and 

persistent ice concentrations over Hanna Shoal in 2012 and 2013 resulted in a ~15 m thick cap of dilute 

MW sitting above much denser winter water over most of the shelf north of ~71.5°N. As a consequence 

the shelf stratification was substantially greater in 2012 and 2013 than in 2011. A front extending from 

the southwest side of Hanna Shoal (along ~71.5°N) eastward to Barrow Canyon separated the MW region 

from BSSW to the south. The MW pool north of the front supported substantially more mesoscale and 

sub-mesoscale surface layer density variations than the waters south of the front. In the absence of MW in 

2011, we presume that surface density variability was comparatively small during that summer.

We suggest that the enhanced variability in the MW region in 2012 and 2013 may be a result of two 

distinct causes. The first is due to instability of the MW/BSSW front (Lu et al. 2015), which is expected 

to enhance the variance at the longer length scales and subsequently, through forward cascade 

(McWilliams 2008), to smaller length scales. The second may arise as a consequence of the ice floe size 

distribution, which changes seasonally. Winter floes tend to be large and consolidated, whereas in 

summer, floe consolidation decreases, and floe areas and thicknesses span a broader range (Perovich and 

Jones 2014). Spatial heterogeneity in the ice thickness distribution implies variations on similar spatial 

scales in the ice-ocean drag and melt rates. In August of 2012 and 2013, Hanna Shoal was covered by 

consolidated and grounded ice spanning 10s of km2, as well as regions containing smaller, unconsolidated 

floes interspersed with leads. The buoyant MW plumes formed should vary in size accordingly, thus 

contribute to the surface density variance wavenumber spectrum. Different plume sizes are subject to 

different dynamical constraints; those within the mesoscale range are subject to rotation while 

ageostrophic dynamics are important over the sub-mesoscale portion of the spectrum (Yankovsky and 

Yashayaev 2014).

What caused the differences in Hanna Shoal summer ice conditions across the 2011 – 2013 period? While 

we cannot provide a definitive answer, we offer several possibilities. First we note that the northward heat 

transport through Bering Strait was substantially greater in 2011 than in 2012 and 2013 (Woodgate et al. 

2012; Woodgate et al. 2015), which suggests that the ocean heat flux to the ice was greater in 2011. 

Indeed, the 2011 heat flux through Bering Strait was as large as that in 2007, when ice retreat across the 

Chukchi Sea shelf was also unusually early (Woodgate et al. 2010). Persistent northeasterly winds 
From Weingartner et al. 2017b

F18

observed in summer 2011 favors the westward drift of ice (Spall 2007). These winds would also force a 

westward drift (Pisareva et al. 2015) of newly exposed surface waters heated by solar radiation and an 

ever-increasing fetch over which waves would be generated. Both factors would enhance ice-edge retreat. 

We have also examined solar radiation differences amongst these summers and find that this variable 

alone cannot account for the differences among years. For example, NARR incoming solar radiation 

estimates from Bering Strait and the southern and northeastern Chukchi shelf indicate that the mean daily 

June and July values in 2012 exceeded that of the other years by ~40 W m-2, and yet ice concentrations 

were quite high that summer.

While northeasterly winds in the early part of summer 2011 may have played a role in ice retreat, the 

heavy ice concentrations over Hanna Shoal in 2012 and 2013 do not appear to be related to the August 

winds. In August 2012, mean winds were southerly at 1 m s-1, whereas in 2013, the mean August winds 

were northeasterly at 2 m s-1 (Figure F2c – f). Additionally ice conditions on Hanna Shoal do not 

correspond to ice concentrations in the Arctic basin. According to the NIC, the heavy ice concentrations 

in summer 2012 occurred when the arctic-wide sea ice extent was the third lowest over the 1979 to 2013 

period, with only the summers of 2007 and 2011 having lower ice concentrations. Indeed, in 2012, the 

entire Beaufort Sea shelf and slope were ice-free while Hanna Shoal was covered by thick ice at high 

concentrations.

Another factor affecting Hanna Shoal summer ice conditions may be processes that occurred the previous 

winter. In both August 2012 and 2013 thick ice was grounded atop Hanna Shoal. According to A. 

Mahoney (pers. comm.), grounding was evident over Hanna Shoal in late winter of both years. Eicken 

and Mahoney (2015) maintain that the source of this grounded ice is westward-drifting, heavily deformed 

ice from the eastern Beaufort (corroborated by Babb et al. (2013) who deployed ice beacons in 2011)

and/or from a heavily-ridged ice arch that often protrudes north of Point Barrow. Ice keels must be ~20 m 

in order to ground on the shallower portions of the Shoal. Sufficiently thick and deformed ice does not 

typically occur in the Beaufort Sea until January, at the earliest (H. Eicken pers. comm.). Once firmly 

grounded ice is established over Hanna Shoal, additional ice, even if thin, may collide and adhere on the 

windward side of the grounding zone. In this scenario the initial grounding subsequently promotes 

consolidation and additional ice deformation, which leads to heavy concentrations of thick ice that persist 

through summer.

There is indirect support for this hypothesis based on the mean January through April winds and ice drift 

derived from moored ADCP and IPS data (Figure F15). Figures F15b, d, and f show the mean winds and 

sea level pressure over the northern Bering and Chukchi Seas from January through April of 2011, 2012, 

and 2013. In 2011, the mean winds were weak (<1 m s-1) and southerly. In contrast, the mean winds in 

2012 were north-northeasterly at 2 – 3 m s-1, and in 2013 they were northeasterly at 3 – 5 m s-1. These 

differences are reflected in the mean monthly ice velocity and variance ellipses for January through April 

of each year (Figure F15a, c, and e). In the winter of 2011, the ice drift was northeastward in accordance 

with mean currents in these areas (Weingartner et al. 2005). Although the drift east and north of Hanna 

Shoal may have been different, the mean winds were clearly unfavorable for ice advection from the 

Beaufort Sea. In contrast the ice drift in the winter of 2012 was southwesterly at ~5 cm s-1 at all sites and 

in 2013 the drift was westerly at ~5 – 10 cm s-1. Although the interannual differences in drifts corroborate 

this hypothesis, the results based on the mean keel depths using only the upper 25% of the ice keel 

distributions are less supportive. In each year the largest of these means occurred in March and April and 

there was generally little spatial variability amongst sites. In 2011 and 2013 the values were 14.5 m and in 

2012 they were 21 m. On this basis, it appears that 2012 was more favorable for ice grounding on the 

Shoal than in the other years.

Our measurements indicate that in all three years the densest winter water surrounded the Shoal. The 

source of this dense water is somewhat uncertain, however. One conceivable source is local formation 
From Weingartner et al. 2017b

F19

Figure F15. Each row shows the mean January – March ice and wind velocities (left panels) and regional 

wind vectors and sea level pressure (right side) over the northern Bering and Chukchi seas for 2011 (top 

row), 2012 (middle row), and 2013 (bottom row). Black arrows and ellipses denote the mean ice velocity 

and the corresponding velocity variance ellipses, while the red arrows and ellipses are for the winds.

over the Shoal during the preceding winter, as polynyas can form in the lee of grounded ice (Mahoney et 

al. 2012). Alternatively, brine repulsion from leads over the shallow Shoal would enhance local dense 

water formation. Under either circumstance, the dense water would be trapped to the Shoal and, in the 

absence of additional forcing, circulate anticyclonically around the Shoal (Spall 2013). However, the 
From Weingartner et al. 2017b

F20

summer synoptic shipboard measurements of Pickart et al. (2016) and the annually averaged 

measurements from moorings in 2011 (Weingartner et al. 2017a) along the south side of the Shoal 

suggest a mean flow of ~2 cm s-1 to the southeast, which would transport the dense water toward Barrow 

Canyon. On the northwest side of the Shoal, the mean annual flow is northeastward at ~5 cm s-1

(unpublished data). Although the vertically-averaged flow on the northeast side of the Shoal is not 

significantly different from zero (Weingartner et al. 2017a), the mean bottom flow is southward at ~2 cm 

s

-1 (unpublished data) The smaller of these values implies an advective time scale of 75 days for the 

movement of bottom waters from the west to the east side of the Shoal. If all of this dense water resulted 

solely from local ice formation processes, this time scale implies that the dense water surrounding the 

Shoal should have been absent by August assuming freezing ends in mid-May as suggested by 

climatology (Maykut 1986). The presence of dense water well into September suggests another advective 

source that transports dense waters toward the Shoal from either the Central Channel and/or Herald 

Valley (Pickart et al. 2010). Regardless of the source, there were considerable interannual differences in 

dense water salinities with the saltiest (~33) water observed in 2012 and the least salty (~32.5) in 2011. 

These differences may affect the vertical stratification in the following summer and perhaps the strength 

of the baroclinic flow on the northeast side of the Shoal.

F.5 Conclusion

Hydrographic data from the northeastern Chukchi shelf collected in August and September of 2011 –

2013 showed large interannual variations in the hydrographic properties surrounding Hanna Shoal. These 

differences primarily relate to the salinities of the bottom waters and to the presence or absence of surface 

MW. The latter is tied to processes that govern summer ice retreat and, as hypothesized here, to the 

previous winter’s history of ice advection, which may control grounding on the Shoal. We also found a 

northwestward baroclinic flow on the northeast side of the Shoal, which opposed the model-predicted 

clockwise barotropic motion around the northwest side of the Shoal. These opposing flow tendencies 

suggest zonal flow convergence on the north side of the Shoal, which implies that the vertically-integrated 

meridional pressure gradient must vanish or otherwise adjust to these opposing tendencies. If convergence 

is a persistent feature of the circulation here then the region north of Hanna Shoal may be a site of 

enhanced cross-shelf transport.
G1

G. Circulation of the Chukchi Sea Shelfbreak and Slope from Moored 

Timeseries

Abstract

Data from a year-long mooring array, deployed across the shelfbreak and slope of the Chukchi Sea 160

km to the west of Barrow Canyon, are used to describe and quantify the circulation and water masses of 

the region. Two currents are identified: an eastward-flowing shelfbreak jet and a westward-flowing 

current over the continental slope. While both of these features have been noted previously in the 

literature, the timeseries presented here demonstrate their year-round existence. In the mean the Chukchi 

slope current is estimated to transport 0.45 ± 0.03Sv of Pacific water westward, while the shelfbreak jet 

transports 0.049 ± 0.004Sv eastward towards Barrow Canyon. The slope current is surface-intensified in 

summer and fall, but in winter and spring it becomes middepth-intensified, moves shoreward, and 

weakens. Using composite averages, we investigate the two extreme states of the circulation: (i) a strong 

slope current coincident with a reversed (westward-flowing) shelfbreak, and (ii) a strong (eastwardflowing) shelfbreak jet coincident with a weak slope current. Both states occur under varied wind 

conditions, but the wind stress curl patterns are consistent with the anomalous circulation in each case. In 

state (i) the wind stress curl field implies a sea surface height gradient across the Chukchi shelf-slope that 

drives enhanced flow to the west via geostrophic set up, while the opposite is true for state (ii). Upwelling 

at the shelfbreak occurs throughout the year, but these events are not always associated with easterly 

winds as is the case in the Beaufort Sea. Finally, the observed signals and timing of the cold winter water 

within Barrow Canyon and within the Chukchi slope current are consistent with the notion that the 

current is fed by the Pacific water outflow from the canyon.

G.1 Introduction

The Pacific inflow through Bering Strait, driven by the large-scale sea level gradient between the Pacific 

and Arctic Oceans (Coachman and Aagaard 1988), plays a key role in the regional ecosystem of the 

Chukchi Sea and Canada Basin (Aagaard and Carmack 1989; Walsh 1995; Steele et al. 2004; Shimada et 

al. 2006). The Pacific-origin water carries nutrients, heat, and freshwater into the Chukchi Sea which, 

among other things, impacts the circulation and stratification of the shelf, the growth of phytoplankton, 

and the distribution of sea ice (Weingartner et al. 2005; Hill et al. 2005; Yang 2006; Woodgate et al. 2010; 

Spall et al. 2013). After some degree of modification on the Chukchi shelf, the water is then fluxed into 

the Canada Basin via different mechanisms of shelf-basin exchange, where it has a profound effect on the 

chemical and physical properties of the interior halocline (Pickart et al. 2005; Spall et al. 2008; Toole et al.

2010). 

It is generally believed that there are three main, topographically steered pathways by which Pacific water 

flows poleward through the Chukchi Sea (Figure G1) (Weingartner et al. 2005). The western pathway 

progresses through Herald Canyon between Wrangel Island and Herald Shoal; the central pathway flows 

through the Central Channel between Herald and Hanna Shoals; and the eastern pathway parallels the 

Alaskan coast from Cape Lisburne to Barrow Canyon. In summertime this branch is known as the

Alaskan coastal current (ACC; Paquette and Bourke 1974). Recent work has suggested that the central 

branch forms a number of smaller filaments as it flows towards Hanna Shoal (Figure G1) (Pickart et al.

2016). The precise partitioning of transport between the three branches remains uncertain. Woodgate et al.

(2005) suggest that, averaged over the year, the division of transport is roughly equal. However, their 

study was based on a limited number of moorings. On the other hand, various studies have suggested that,
G2

Figure G1. Schematic circulation in the Chukchi Sea (from Corlett and Pickart 2017), showing the three 

pathways by which Pacific water flows poleward through the Chukchi Sea.

at least during the summer months, much of the Pacific water flowing through Bering Strait is eventually 

channeled into Barrow Canyon (Itoh et al. 2013; Gong and Pickart 2015; Pickart et al. 2016). 

There is also uncertainty as to how and where the Pacific water exits the Chukchi shelf into the Canada 

Basin. A portion of the outflow from Barrow Canyon turns eastward along the edge of Beaufort Sea to 

form the Beaufort shelfbreak jet (Pickart 2004; Nikolopoulos et al. 2009). Using data from a highresolution mooring array, the year-long mean transport of the jet from summer 2002 to summer 2003 was 

estimated to be 0.13 ± 0.08Sv (Nikolopoulos et al. 2009). However, Brugler et al. (2014) demonstrated 

that this transport dropped by more than 80% later in the decade, suggesting that the Beaufort shelfbreak 

jet can only account for a small fraction of the Bering Strait inflow. Some of the Pacific water also exits 

the Chukchi shelf through Herald Canyon and forms an eastward-flowing shelfbreak jet along the edge of 

the Chukchi Sea (Mathis et al. 2007; Pickart et al. 2010; Linders et al. 2017). A portion of the water also 

appears to enter the East Siberian Sea through Long Strait (Woodgate et al. 2005), although this has not 

yet been established as a permanent pathway. Recently, Timmermans et al. (2017) argued that some of 

the Pacific water is fluxed into the Canada Basin via subduction along the entire edge of the Chukchi 

shelf. 

The long-term mean northward transport of Pacific water at the mouth of Barrow Canyon has been 

estimated to be 0.44 Sv (Itoh et al. 2013), which is far greater than the eastward transport of the Beaufort 
G3

shelfbreak jet. The obvious question then is: where does the bulk of the Pacific water go upon exiting the 

canyon? A new study has documented the existence of a westward-flowing current along the continental 

slope of the Chukchi Sea. Using hydrographic and velocity data from 46 shipboard transects across the 

shelfbreak/slope of the Chukchi Sea between 2002 and 2014, Corlett and Pickart (2017) revealed the 

presence of the current during the summer months (July – October), which is surface-intensified and order 

50 km wide. The strongest flow occurs within 25 km of the shelfbreak. Corlett and Pickart (2017) named 

the current the Chukchi slope current, and estimated the transport of Pacific water to be 0.50 ± 0.07Sv. It 

was argued that the current is formed from the outflow from Barrow Canyon, and, using these new data 

together with historical measurements, Corlett and Pickart (2017) constructed a mass budget of the 

Chukchi shelf where the inflows and outflows balance each other within the estimated errors. 

In addition to the westward-flowing Chukchi slope current, Corlett and Pickart (2017) also quantified the 

presence of the eastward-flowing Chukchi Shelfbreak Jet (Figure G1), whose existence was implied 

previously from anecdotal evidence only. Using the large number shipboard transects, Corlett and Pickart

(2017) estimated the jet’s mean summertime transport to be 0.10 ± 0.03 Sv. Notably, the flow at the 

shelfbreak can at times be westward. It is thought that the jet gets entrained into the Chukchi slope current 

at the mouth of Barrow Canyon (Figure G1). 

One of the dominant mechanisms of shelf-basin exchange across the edge of the Beaufort Sea is winddriven upwelling (Pickart et al. 2009; Pickart et al. 2011; Lin et al. 2018). Easterly winds, arising from the 

intensification of the Beaufort High and/or passing Aleutian Lows to the south, readily reverse the 

Beaufort shelfbreak jet and drive water from the slope onto the shelf. This occurs during all seasons of the 

year and under different ice conditions (Schulze and Pickart 2012). Evidence of upwelling on the Chukchi 

slope is far less conclusive. Llinás et al. (2009) suggested the occurrence of upwelling based on a single 

shipboard transect north of Hanna Shoal, characterized by the presence of Atlantic water on the upper 

slope as well as surface-intensified westward flow, which they interpreted as a reversed shelfbreak jet. 

Using observations and a simplified numerical model, Spall et al. (2014) argued that upwelling of 

nutrients from the halocline to the outer shelf north of Central Channel contributed to the massive underice phytoplankton bloom reported by Arrigo et al. (2014). Recently, Corlett and Pickart (2017) presented 

evidence that the westward-flowing Chukchi slope current is intensified under enhanced easterly winds. 

However, more extensive measurements are necessary to robustly establish the occurrence of upwelling 

along the Chukchi slope and its forcing mechanisms. 

This study presents results from a mooring array that was deployed across the shelfbreak and slope of 

Chukchi Sea from October 2013 to September 2014 to the northeast of Hanna Shoal. It is the first set of

high spatial resolution timeseries obtained from the region. The primary aim of the study is to elucidate 

the structure and transport of both the Chukchi shelfbreak jet and Chukchi slope current, and to identify 

the nature and causes of the variability of the two currents. We begin with a presentation of the different 

sources of data used in the study in Section G2, followed in Section G3 by an investigation of the mean 

structure and seasonality of the circulation and hydrography. In Section G4 the volume transport of the

shelfbreak jet and slope current, as well as their correlation, are addressed. In Section G5 we consider 

extreme states of the two currents using a composite analysis. The occurrence of upwelling is then 

investigated in Section G6, followed by consideration of the propagation of water mass signals from 

Barrow Canyon into the slope current in Section G7. 

G.2 Data and Methods

The data used in this study were collected as part of a year-long field program funded by the Bureau of 

Ocean and Energy Management (BOEM) entitled “Characterization of the Circulation on the Continental 

Shelf Areas of the Northeast Chukchi and Western Beaufort Seas”. The program employed moorings, 
G4

gliders, drifters, and included multiple shipboard surveys. The present analysis uses primarily the 

mooring data, along with various ancillary data sets. 

G.2.1 Mooring Data

From October 2013 to September 2014, six moorings (CS1-5 and FM1) were deployed across the 

shelfbreak and slope of the Chukchi Sea (Figure G2). All of the moorings were equipped with an upwardfacing acoustic Doppler current profiler (ADCP, 300 KHz or 75 KHz) near the bottom, which provided 

hourly velocity profiles with a vertical resolution of 5 – 10 m. Hydrographic properties were measured by 

MicroCATs situated next to the ADCPs, and with two types of conductivity-temperature-depth (CTD) 

profilers: Coastal Winched Profilers (CWPs) at CS1 and FM1, and Coastal Moored Profilers (CMPs) at 

every site except CS1. The CMPs provided vertical traces of temperature and salinity nominally four 

times per day with a vertical resolution of 2 m, while the CWPs produced profiles once per day with a 

resolution of 1m. A detailed summary of the mooring components is contained in Table G1. 

All of the ADCPs and MicroCATs returned high quality year-long records, with the exception of the 

ADCP at mooring CS1 which experienced compass problems. Consequently, this record was not used in 

the analysis (this has no bearing on our results since CS1 and FM1 were essentially at the same location).

Figure G2. (a) Bathymetric map of the northeastern Chukchi Sea showing mooring locations used in the 

study. The six moorings comprising the shelfbreak/slope array are shown by the yellow stars. The three 

additional moorings east of Hanna Shoal and the mooring at the head of Barrow Canyon are shown by 

the red and blue stars, respectively. The red line and black coordinate frame indicate the rotated 

coordinate system. The bathymetry is from IBCAO v3. (b) Large-scale map showing the entire Chukchi 

Sea. The region in (a) is indicated by the dashed box. The magenta and cyan boxes delineate the domain 

over which the ice concentration is calculated for the shelfbreak/slope array and for the coastal polynya

region south of Barrow Canyon, respectively. The mooring sites are shown by the black dots. (c) 

Configuration of shelfbreak/slope moorings in the vertical plane. The origin of the x-axis is Hanna Shoal.
G5

Table G1. Mooring information

Mooring 

ID Latitude Longitude

Water 

Depth

(m)

Instrument Duration Instrument 

Depth (m)

Range 

Depth

(m)

Sample 

Interval

(h)

Vertical 

resolution

(m)

CS1 72°15.37’N158°04.26’W 66 CWP 10/16/2013-

11/15/2013 - 1-65 24 1

MicroCAT 10/15/2013-

09/21/2014 65 - 0.25 -

FM1 72°15.81’N158°02.46’W 67 ADCP 10/25/2013-

09/21/2014 60 8-53 1 5

MicroCAT 10/25/2013-

09/21/2014 60 - 0.25 -

CS2 72°18.02’N157°43.52’W 102 ADCP 10/12/2013-

09/22/2014 89 11-81 1 5

MicroCAT 10/12/2013-

09/22/2014 89 - 0.25 -

CS3 72°20.18’N157°26.89’W 163 CMP 10/14/2013-

09/21/2014* - 39-146 6 2

ADCP 10/13/2013-

09/22/2014 151 22-132 1 10

MicroCAT 10/12/2013-

09/22/2014 151 - 0.25 -

CS4 72°23.10’N 157°8.76’W 249 CMP 10/15/2013-

09/21/2014 - 50-235 6 2

ADCP 10/13/2013-

09/22/2014 241 22-222 1 10

MicroCAT 10/12/2013-

09/22/2014 241 - 0.25 -

CS5 72°25.82’N156°50.37’W 356 CMP 10/15/2013-

06/21/2014 - 42-340 6 2

ADCP 10/13/2013-

09/22/2014 349 31-331 1 10

MicroCAT 10/13/2013-

09/22/2014 349 - 0.25 -

NE40 72°7.35’N 160°29.68’W 41 ADCP 09/09/2013-

09/18/2014 40 3-37 0.5 1

NE50 72°9.73’N 159°7.52’W 50 ADCP 09/09/2013-

09/18/2014 49 4-46 0.5 1

NE60 72°10.89’N158°33.07’W 57 ADCP 09/09/2013-

09/18/2014 56 5-53 0.5 1

*CMP at CS3 got stuck near the top of the mooring on December 9, 2013.

Unfortunately, the moored profiler coverage was generally poor. The CWP at CS1 collected data for 

about one month and then stopped abruptly, while the CWP at FM1 failed immediately. Of the CMPs, 

only the one at CS4 profiled for the entire duration of the deployment. The instrument at CS5 profiled for 

eight months, the one at CS3 for two months, and the one at FM1 not at all. In the latter two instances, 

however, the CTD sensor on the profiler remained operational at a fixed depth, acting as a de facto

MicroCAT. The CMP at CS2 failed entirely. Details regarding the mooring instrumentation and data 

coverage are found in Table G1
G6

Three additional moorings (NE40, NE50, NE60) were maintained from September 2013 to September 

2014 on the eastern side of Hanna Shoal at roughly the 40 m, 50 m, and 60 m isobaths (Figure G2). 

Together, the two sets of moorings comprise an array extending from the edge of Hanna Shoal across the 

shelfbreak to the upper slope. The shelf moorings were equipped with ADCPs and MicroCATs at the 

bottom, recording velocity twice per hour and hydrographic data four times per hour. The vertical 

resolution of the ADCPs was 1 m. Velocity with same resolution and daily-averaged hydrographic data 

from a mooring at the head of Barrow Canyon (BC2, Figure G2) were also used for part of the analysis. 

The reader should consult Weingartner et al. (2017a) for details about the configuration of this mooring.

All of the velocity data were de-tided using the T_Tide harmonic analysis toolbox (Pawlowicz et al.

2002). This revealed that there was low tidal energy level across the array: the maximum amplitude of the 

eight dominant tidal constituents was found to be less than 2.2 cm s

-1

, which is considerably smaller than 

the sub-tidal signals of interest. The inertial signal was also found to be generally insignificant. A rotated 

coordinate system was used in the analysis. The alongstream direction was determined by averaging the 

year-long mean, depth-integrated velocity vectors at the five outer moorings. The positive x (alongstream) 

direction is defined as southeastward (138°T) and the positive y (cross-stream) direction is northeastward

(48°T, Figure G2a). The associated velocities are referred to as u and v, respectively. Vertical sections of 

the two components of velocity were constructed at each time step using Laplacian-spline interpolation, 

with a horizontal grid spacing of 2 km and vertical gird spacing of 15 m. The domain of the vertical 

sections is limited to the five outer moorings, i.e., the region of the shelfbreak and slope, which is the 

main focus of the paper. 

G.2.2 Wind Data

Wind timeseries from the Barrow, AK meteorological station are used in the study. This site is roughly 

120 km to the southeast of the array. The data were obtained from the National Oceanographic Data 

Center (NODC) of the National Oceanic and Atmospheric Administration (NOAA) and have been quality 

controlled and interpolated to an hourly time base. The reader is referred to Pickart et al. (2013) for details. 

G.2.3 Atmospheric Reanalysis Fields

To assess the effect of the broad-scale atmospheric forcing, we used reanalysis data from the North 

American Regional Reanalysis (NARR) (Mesinger et al. 2006). This includes sea level pressure (SLP) 

and 10 m wind fields with a lateral resolution of 32 km and time resolution of 6 hours. The NARR 

product represents an improvement on the global National Centers for Environmental Prediction (NCEP) 

reanalysis dataset in this region in both accuracy and resolution. The correlation between the Barrow wind 

timeseries and the NARR wind record in the vicinity of moorings is 0.8, at a confidence level of 95%.

G.2.4 Ice Concentration Data

The ice concentration data used in the study are the blended Advanced Very High Resolution Radiometer 

(AVHRR) and the Advanced Microwave Scanning Radiometer (AMSR) product from NODC, NOAA. 

The spatial and temporal resolution of the data are 0.25° and once per day. We constructed a timeseries of 

ice concentration for the location of the array by averaging the data within the cyan box in Figure G2b. To 

assess the polynya activity south of Barrow Canyon we averaged the data within the magenta box in 

Figure G2b. 
G7

G.3 Mean and Seasonality of the Circulation and Hydrography of the 

Shelfbreak and Slope

The year-long, depth-mean vectors (Figure G3) indicate that there is persistent flow to the northwest 

along the Chukchi slope. This confirms that the Chukchi slope current is a year-round feature, i.e., it is 

not only present during the summer months as reported in Corlett and Pickart (2017). Notably, the depthintegrated flow at mooring CS2 is much weaker than at the other sites; in fact, it is not significantly 

different than zero. This can be explained by considering the mean vertical section of alongstream 

velocity (Figure G4a), which reveals the presence of the Chukchi shelfbreak jet at this location. The mean 

velocity section shown here is remarkably consistent with that presented in Corlett and Pickart (2017)

using the summertime shipboard data (their Figure G4b). In particular, both sections (using independent 

data) indicate a bottom-intensified, eastward-flowing shelfbreak jet inshore of the surface-intensified, 

westward-flowing slope current. Furthermore, both data sets show that there is an eastward-flow of 

Atlantic water at depth on the mid-slope. Unfortunately, the mooring array did not extend far enough 

offshore to completely bracket the slope current. The hydrographic data from the mooring array indicates 

that, percentage-wise, the most common water types sampled were remnant winter water and Atlantic 

water (Figure G4b). 

Figure G3. Year-long, depth-mean velocity vectors (blue) at the mooring sites and mean 10 m wind

vector (black) at the Barrow, AK meteorological station. The standard error ellipses are shown (see the 

scales at the lower left). The red line indicates the along-stream direction (see Figure G2).
G8

Figure G4. (a) Year-long mean along-stream velocity section (positive is southeastward). The thick black

line is the zero velocity contour. The dashed black line shows the boundary of shelfbreak and slope. The 

grey shading denotes regions of no data coverage. The mooring sites are shown at the top. (b) Potential 

temperature-Salinity diagram for all of the hydrographic data. The color represents the percentage of data 

within a 0.1°C by 0.1°C grid. The black lines delimit the different water masses: MW = meltwater; ACW = 

Alaskan Coastal Water; BSSW = Bering Sea Summer Water; RWW = Remnant Winter Water; WW = 

newly-ventilated Winter Water; AW = Atlantic Water.
G9

There is pronounced seasonality of the slope current (Figure G5). It is surface-intensified in summer and 

autumn, and middepth-intensified in winter and spring. Furthermore, it migrates onshore during the 

winter and spring. By contrast, there is little seasonal variation of the shelfbreak jet: it is always bottomintensified, although it appears to be a bit stronger in fall and weaker in spring. The hydrographic 

timeseries of temperature and salinity in the slope current (Figure G6) reveal the presence of newlyventilated winter water starting in late-April, lasting until the end of August (although its presence is 

intermittent). There is also evidence of local convection during the winter months. In particular, there are 

numerous instances of newly-ventilated winter water appearing in the upper 50 – 75 m. This is likely the 

signature of convective overturning driven by brine rejection as a result of re-freezing polynyas. The 

nature and frequency of these ventilation events warrants further investigation. 

The mean wind measured by the Barrow meteorological station during 2013 – 2014 was out of the 

east/northeast at 1.6 m s

-1 (Figure G3). The wind rose shown in Figure G7 reveals that, while occasionally 

there were strong southwesterly/westerly winds, these occurred much less frequently. The seasonal cycle 

of winds and ice concentration are shown in Figure G8. Overall the winds were strongest during the fall 

and early winter, and weakest during the spring (Figure G8a). Freeze-up at the array site occurred in late 

November, and the ice cover remained between 90 – 100% until early July (Figure G8b). We also show 

the ice concentration in the region south of Barrow Canyon, where the northeast Chukchi Sea polynya is 

known to form. One sees that the polynya opened up three times during the year – in early-January, lateJanuary/early-February, and again in late-April/early-May. These periods are marked in Figure G8b. 

Interestingly, the ice cover also decreased over the mooring array during these times. Comparing the wind 

record to the ice concentration timeseries, it is seen that in each case the winds were blowing out of the 

northeast prior to the ice opening up. 

Figure G5. Vertical sections of the seasonally averaged along-stream velocity. The presentation is the 

same as in Figure G4a.
G10

Figure G6. Depth-time plot of potential temperature and salinity at CS4 (a, c) and CS5 (b, d), overlain by 

potential density (contours).
G11

Figure G7. Wind rose showing wind speed and direction at Barrow for the duration of the deployment.

Figure G8. (a) Daily-mean wind velocity at the Barrow meteorological station (blue vectors). The gray 

shading and red vectors denote periods of northeasterly wind, and the corresponding mean wind velocity,

preceding the three major occurrences of reduced ice cover at the mooring array site and south of Barrow 

Canyon. (b) Ice concentration timeseries at the array site (cyan curve) and at the location of the polynya 

south of Barrow Canyon (magenta curve). The gray shading indicates the three periods of reduced ice 

concentration.
G12

G.4 Volume Transport of the Slope Current and Shelfbreak Jet

In the volume transport calculations below, we used the value of x = 156 km as the dividing line between 

the shelfbreak jet and slope current (see Figure G4a). In particular, the transport between 140 km < x < 

156 km is taken to be that of the shelfbreak jet, which can be positive (eastward) or negative (westward), 

while the negative (westward) transport seaward of x = 156 km is considered to be the slope current. As 

noted above, the array did not always bracket the slope current, so we invoked a “mirroring” technique to 

estimate the missing transport. Specifically, in instances when the velocity core of the current was close to 

or beyond the edge of the grid, we took the offshore part of the current to be the mirror image of the

inshore part. This was only done using information within 10 km of the edge of the grid, and was also 

limited vertically to the upper 150 m of the water column. This was an attempt to boost the transport 

estimate to be closer to the true value, although this extrapolated estimate is still clearly an underestimate. 

The year-long mean westward transport of the slope current is -0.57 ± 0.04 Sv. Using the potential 

vorticity constraint developed by Nikolopoulos et al. (2009) to identify the boundary between the Pacific 

water and Atlantic water, we find that the transport of Pacific water is -0.45 ± 0.03 Sv. We stress that this 

is an underestimate since the array did not fully capture the current. However, the mean absolute 

geostrophic velocity section of Corlett and Pickart (2017) did bracket the flow, and they found a mean 

Pacific water transport of -0.50 ± 0.07 Sv for the months of July – October, which is fairly close to the 

value computed here (our value for July – October is -0.57 ± 0.05 Sv). Using the mooring data, the yearlong mean transport in the vicinity of the shelfbreak is also westward, -0.014 ± 0.006 Sv. However, as 

seen in Figure G4a, the eastward-flowing shelfbreak jet is bottom intensified; isolating this eastward flow 

gives a transport of 0.049 ± 0.004 Sv. This value is somewhat larger than the transport of the Beaufort 

shelfbreak jet in recent years (mean of 0.023 ± 0.018 Sv from 2008 – 2014; P. Lin pers. comm., 2017). 

The transport of the slope current varies substantially on a variety of time scales, ranging from a value of 

zero to 2 Sv (Figure G9a). The monthly mean timeseries implies that the transport is larger in summer,

with a peak value in September (Figure G9c). We note that this is at odds with the results of Corlett and 

Pickart (2017) who found that the slope water transport was largest in October. The flow at the shelfbreak 

fluctuates between positive and negative throughout the year (Figure G9b), with range of approximately 

0.2 to -0.2 Sv. The monthly mean values are westward from December through June, with eastward flow 

in late-summer and fall (except for September). Notably, removing the high-frequency fluctuations by 

applying a 5-day low-pass filter indicates that the transport of the slope current and that of the shelfbreak 

jet are significantly correlated (r = 0.6, 95% confidence). Possible reasons for this are addressed below.

G.5 Extreme States of the Slope Current and Shelfbreak Jet

Using an extensive collection of shipboard data, Corlett and Pickart (2017) found that the westward flow 

of the Chukchi slope current was enhanced under strong easterly winds. They demonstrated this by 

constructing a composite mean vertical section of absolute geostrophic velocity for calm wind conditions, 

and comparing this to the composite section when the easterly wind exceeded 4 m s-1. In this case easterly 

was taken to be the component of wind directed out of the southeast parallel to the shelfbreak. Their result 

is reproduced in Figure G10 (left hand column). One sees that when the winds are weak there is a stronger 

eastward-flowing shelfbreak jet (and stronger eastward flow of Atlantic water at depth), and when the 

winds are strong the slope current is stronger and wider. When we compute the analogous composite 

sections – using the mooring data for the same seasonal period as Corlett and Pickart (2017) (July –

October) – we obtain a similar result (Figure G10, right hand column). In an effort to shed more light on 

the influence of the wind on both the slope current and the shelfbreak jet, we considered the extreme 

states of the two currents. Specifically, we isolated those times when the slope current was strong while 

the shelfbreak jet was simultaneously reversed to the west, and those times when the shelfbreak jet was 
G13

Figure G9. Volume transport timeseries of (a) the Chukchi slope current (blue curve), and (b) the Chukchi 

shelfbreak jet (yellow curve is the full transport and green dashed curve is the positive transport). (c) 

Monthly-averaged slope current transport with standard errors. (d) Monthly-averaged transport and 

standard errors for the full shelfbreak jet (yellow curve) and for times when the jet transport is positive 

(green dashed curve).

flowing strongly to the east while the slope current was weak. We developed objective criteria for 

identifying these two states of the circulation.

G.5.1 Strong Slope Current and Reversed Shelfbreak Jet

The criteria used for first type of event was that the slope current transport be at least 0.3 standard 

deviations greater than the mean, while the flow at the shelfbreak be at least 0.3 standard deviations 

weaker than the mean (recall that the average flow at the shelfbreak is westward). These conditions were 

met approximately 25% of the time over the course of the year (Table G2). Based on the results of Corlett 

and Pickart (2017), one might expect these conditions to always correspond to an easterly wind (i.e., with 

a component of the wind paralleling the shelfbreak from the southeast). Surprisingly, however, this 

extreme state occurred under various wind conditions, with the wind blowing out of all four quadrants 

depending on the event (Table G2). Here we show the composite fields for the two wind conditions that 

resulted in the most days with a strong slope current and reversed shelfbreak jet. 
G14

Figure G10. Composite velocity sections during (a, b) calm wind conditions and (c, d) times when the 

easterly winds exceeded 4 m s

-1 for the months of July – October. (a) and (c) are reconstructed sections 

of absolute geostrophic velocities (color) from Corlett and Pickart (2017). (b) and (d) are velocity sections 

(color) from mooring data. The thin black lines are velocity contours.

G.5.1.1 Winds from the Southwest

There were 11 instances, totaling 46 days, in which the slope current was anomalously strong and the 

shelfbreak jet was reversed while the wind was out of the southwest. The composite average vertical 

section of alongstream velocity (Figure G11a) shows that there was westward flow throughout the array, 

with the slope current 5 – 10 cm s

-1 stronger than normal (Figure G11b). The atmospheric conditions 

corresponded to low SLP north of the Chukchi Sea and associated cyclonic winds (Figure G11c). The 

wind stress curl was strongly positive over the northern part of the Chukchi shelf, but weakly positive 

offshore of the array on the Chukchi slope and basin (Figure G11d). This implies that there would be a 

larger drop in sea surface height on the shelf than offshore, which would set up a geostrophic response of 

enhanced flow to the west, consistent with the mooring observations. 
G15

Table G2. Statistics for the two types of extreme events considered in the text: (i) strong slope current 

and reversed shelfbreak jet; (ii) strong shelfbreak jet and weak slope current. The first column indicates 

the quadrant from which the wind was blowing. The percentage in parentheses means the fraction of the 

event length relative to the total length in the last row. The underlined percentages represent the fraction 

of total length relative to the entire year-long duration of the record.

Strong Slope Current & Reversed 

Shelfbreak Jet

Strong Shelfbreak Jet & Weak Slope 

Current

Number of 

events

Total length in 

days

Mean event 

length in days 

(range)

Number 

of 

events

Total length in 

days

Mean event 

length in days 

(range)

SW-wind 11 46 (51.7%) 4.1 (0.5-17.1) 7 8 (8.3%) 1.1 (0.6-2.1)

NE-wind 7 17 (19.1%) 2.4 (0.6-5.2) 16 66 (68.8%) 4.1 (0.7-12.3)

SE-wind 5 13 (14.6%) 2.6 (1.0-4.4) 6 12 (12.5%) 2.0 (0.7-6.0)

NW-wind 8 13 (14.6%) 1.6 (0.6-4.0) 5 10 (10.4%) 2.0 (0.6-4.6)

Total 31 89 (25.6%) 34 96 (27.8%)

Figure G11. Composite average fields for the strong slope current and reversed shelfbreak jet events 

with southwesterly wind. (a) Vertical section of alongstream velocity. The number of events and their total 

duration is marked on the lower left. (b) Vertical section of alongstream velocity anomaly (composite 

minus the year-long mean). (c) Sea level pressure (color) and 10 m wind vectors from NARR, along with 

the measured wind from the Barrow meteorological station (purple vector). The location of the 

shelfbreak/slope mooring array is indicated by the purple star. (d) Wind stress curl (color).
G16

G.5.1.2 Winds from the Northeast

The second largest segment of time for this circulation state corresponded to winds out of the northeast 

over the mooring site (17 days, Table G2). Again the alongstream flow was westward everywhere, with 

anomalous velocities in the slope current between 10 – 15 cm s

-1 (the largest increase was at depth, Figure 

G12a, b). In contrast to the previous case, low SLP was present to the southeast of the Chukchi Sea, with 

higher SLP over the basin (Figure G12c). The wind stress curl field was again conducive for increased 

flow to the northwest along the slope via geostrophic set up, with positive curl on the shelf (lowered sea 

surface height) and negative curl offshore of the array (raised sea surface height, Figure G12d). 

G.5.2 Strong Shelfbreak Jet and Weak Slope Current

The other extreme state, that of strong eastward flow along the shelfbreak and a weakened slope current, 

occurred approximately 28% of the time during the year (Table G2). The criteria employed here was that 

the shelfbreak jet be at least 0.3 standard deviations larger than the mean, while the slope current be at 

least 0.3 standard deviations weaker than the mean. In this case there was one dominant scenario in which 

the wind blew from the northeast (totaling 66 days, Table G2). The composite vertical sections 

showeastward flow from the outer shelf to the mid-slope, with a clear signature of the shelfbreak jet 

centered at mooring CS2 and enhanced flow of Atlantic water at depth (Figure G13a, b). The slope 

current is confined to the upper right-hand part of the section. The strongest anomalous flow is between 5 

– 10 cm s

-1

. 

Figure G12. Same as Figure G11, except the strong slope current and reversed shelfbreak jet events 

with northeasterly wind.
G17

Figure G13. Same as Figure G11, except the strong shelfbreak jet and weak slope current events with 

northeasterly wind.

The atmospheric pattern associated with this state corresponds to a strong Beaufort High north of the 

Chukchi Sea and the presence of an Aleutian Low near the southern end of the domain (Figure G13c). 

These conditions typically lead to a reversed (westward-flowing) shelfbreak jet in the Beaufort Sea and 

concomitant upwelling, yet our Chukchi slope array shows a strengthened (eastward-flowing) shelfbreak 

jet. This can be explained by the wind stress curl pattern (Figure G13d). The curl is strongly negative on 

the northeast Chukchi shelf, while offshore of the array it is less negative. The resulting sea level gradient 

(higher sea level on the shelf, lower sea level northeast of the moorings) would promote stronger 

geostrophic flow to the southeast. By contrast, the positive curl on the Beaufort shelf, in conjunction with 

the negative curl offshore, would lead to enhanced westward flow along the Beaufort shelfbreak (i.e., a 

reversed shelfbreak jet). This is the same situation that was studied by Pickart et al. (2011) using in-situ 

data and a numerical model. They showed that, during an upwelling event in the Beaufort Sea, the 

eastward flow along the Chukchi shelfbreak/slope became stronger as a result of a sea level increase on 

the Chukchi shelf. 

G.6 Upwelling

Wind-driven upwelling has been well studied on the Beaufort slope, and is associated with enhanced 

easterly winds and a reversed shelfbreak jet (e.g., Pickart et al. 2009; Pickart et al. 2011; Lin et al. 2018). 
G18

Following Lin et al. (2018), we used the near-bottom potential density anomaly in the vicinity of the 

Chukchi shelfbreak/upper-slope to identify the presence of upwelled water from the basin. In particular, 

we low-passed the deep MicroCAT potential density records at moorings FM1 (shelfbreak) and CS2 

(upper slope) using a filter width of 20 days, and computed the difference between these smoothed values 

and the instantaneous values. Upwelling was deemed to occur when the density anomaly was positive at 

both sites for more than a day. The purpose of the low-pass was to remove the influence of more slowlyvarying hydrographic properties of the shelfbreak jet due to alongstream advection. Lin et al. (2018) 

carried out a similar procedure using monthly averages. 

Using this criterion, 15 upwelling events were identified over the course of the year, ranging in length 

from 1.3 to 6.2 days (Table G3). By comparison, Lin et al. (2018) reported an average of 19 events per 

year on the Beaufort slope over a 6-year period. As a measure of the strength of the upwelling, we 

followed Lin et al. (2018) and defined an upwelling index, UI, as the time integral of the potential density 

anomaly (at CS2) over the duration of the event. This takes into account both the magnitude and duration 

of the density signal. The average value of UI was 5.9 kg m-3 h, with a range of 1.1 – 26.7 kg m-3 h (Table 

G3). On the Beaufort slope Lin et al. (2018) reported an average UI of 49.9 kg m-3 h with a range of 0.9 –

222.1 kg m-3 h. Hence, overall, the upwelling on the Beaufort slope appears to be more intense than that 

on the Chukchi slope.

Here we diagnose the strongest occurrence of upwelling (event #9 in Table G3), which took place in May 

2014. Figures G14 – 16 show composite averages for before, during, and after the event. The timeseries 

of potential density anomaly at the two mooring sites is shown in panel (c) of each figure for reference. 

Prior to the event (Figure G14) the slope current was well established, and the shelfbreak jet was reversed. 

The cross-stream flow varied in magnitude and direction across the array, which can be seen in Figure 

G14 e, f (note that positive cross-stream flow is directed offshore). During the event (Figure G15) the 

westward-flowing shelfbreak jet and the slope current both became stronger and veered shoreward, which 

is especially evident in the vector plot of Figure G15f. The onshore flow extended throughout the water 

column and was especially strong at the shelfbreak (Figure G15e). After the event (Figure G16) the slope 

current became markedly weaker and the shelfbreak jet re-established itself to the east (Figure G16b), 

although there was still weak onshore flow from the mid-slope to the outer-shelf near the bottom (Figure 

G16e).

Table G3. Statistics of all upwelling events. The rows show the event number, length (in days) of each 

event, value of the upwelling index (UI, unit: kg m-3 h), mean wind direction, and sign of the mean wind 

stress curl condition (+ and – mean positive and negative curl, respectively) in the vicinity of mooring 

array (cyan box in Figure G2b) for the day before each event.

Event 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 Mean

Length (d) 1.7 2.5 1.9 6 2.2 6.2 3.3 2 4.5 2.4 1.3 3.6 4.5 5.1 4.6 3.5

UI 3.8 1.8 5.5 11.1 1.1 9.0 6.4 2.4 26.7 2.0 3.5 7.8 1.5 2.4 3.4 5.9

Wind 

direction NE SE SW NE NW W NW SE SW SE NW SW NE SW SW SW

Wind 

stress curl - - + + - + + - + + + + - - + -
G19

Figure G14. Composite average fields prior to the strongest upwelling event (event 9, May 2014). (a) Sea 

level pressure (color) and 10 m wind vectors from NARR, along with the measured wind from the Barrow 

meteorological station (blue vector). The location of the shelfbreak/slope mooring array is indicated by the 

purple star. (b) Along-stream velocity section. (c) Density timeseries at mooring CS2 and FM1 from 3 

days before the event to 3 days after the event, where the bold indicates the time period before the 

upwelling. The dashed lines are the 20-day low-passed curves. (d) Wind stress curl (color). (e) Crossstream velocity section (positive is northeastward). (f) Depth-averaged (0 – 250 m) velocity vectors at the

mooring sites (blue arrows) and the mean velocity vector of all moorings (large arrow).

Figure G15. Same as Figure G14, except for the time period during the upwelling.
G20

Figure G16. Same as Figure G14, except for the time period after the upwelling.

The nature of the upwelling event in May – indeed for all of the upwelling events – remains unclear. It is 

certain, however, that the process is different on the Chukchi slope than on the Beaufort slope. In the 

Beaufort Sea enhanced easterly/northeasterly winds consistently drive coastal upwelling (Lin et al. 2018). 

By contrast, upwelling in the Chukchi Sea seems to occur for all wind directions and for both positive and 

negative local wind stress curl (Table G3). In the example shown, the strength of the wind was 

comparable before and during the event, although the wind direction was more out of the west during the 

upwelling (compare Figures G14a and G15a). This in turn caused a wind stress curl pattern that would 

lead to enhanced westward flow at the array site due to geostrophic set up (Figure G15d). This is 

reminiscent of the extreme state described above in Section G5.1.1 – i.e., a strong slope current and 

reversed shelfbreak jet under southwesterly winds. However, it remains to be determined what caused the 

flow to veer onshore. A more thorough investigation of upwelling, perhaps in conjunction with a 

numerical model, is needed. 

G.7 Propagation of Water Mass Signals from Barrow Canyon to the 

Chukchi Slope

Using the definitions of the different water types in Figure G4b, we constructed timeseries of the water 

masses measured throughout the year at three locations: the head of Barrow Canyon (mooring BC2); in

the Chukchi slope current (mooring CS4), and in the Chukchi shelfbreak jet (mooring CS2, Figure G17). 

The most common water mass passing through the head of Barrow Canyon was newly ventilated winter 

water (WW, keeping in mind that the MicroCAT was located near the bottom). This cold water mass was 

present almost exclusively in the canyon from around the beginning of February to early-July (marked by 

the black triangles in Figure G17a). Comparing this to the site of the Chukchi mooring array, one sees 

that the bulk of the WW appeared in the slope current from early-April to early-September (marked by 

the black triangles in Figure G17b). Hence, the water was present at both locations for roughly five 

months, with an offset on the order of two months. This supports the notion put forth by Corlett and 

Pickart (2017) that the outflow from Barrow Canyon feeds the slope current.
G21

Figure G17. Timeseries of water mass occurrence at the Chukchi shelfbreak/slope moorings (CS2, CS4) 

and the mooring at the head of Barrow Canyon (BC2). See Figures G2 and G3 for mooring locations. 

Note that there is no depth scale for moorings CS2 and BC2 since these sites have a single sensor near 

the bottom. The gray shading indicates times when the flow is in the opposite direction of the predominant 

current (southwestward, southeastward, and northwestward in the three panels, respectively). The black 

triangles denote the time periods when the bulk of Winter Water was present.

To investigate this further, we examined the variation in potential temperature of the WW at the two 

locations (Figure G18). The first thing to note is that the water is systematically warmer on the Chukchi 

slope than in Barrow Canyon, by approximately 0.07°C. This makes sense in that lateral mixing would 

warm the water as it exits Barrow Canyon and then flows westward in the slope current. Furthermore, at 

both sites there is a clear moderation of the WW to warmer temperatures as the season progresses. We 

vertically averaged the moored profiler record at CS4 and compared this to the record at BC2. The 

maximum correlation between the two timeseries (r = 0.6, significant at the 95% confidence level) was 

found for a lag of 60 days (BC2 leading CS4). This is consistent with the offset noted above in the arrival 

times of the WW at the two sites.

The geographical distance from the head of Barrow Canyon to its mouth, plus the distance to the Chukchi 

slope array, is approximately 300 km. For a time lag of 60 days, this implies a mean advective speed of 

5.6 cm s-1. The mean velocity at BC2 during the WW period was 17.9 cm s

-1

, which is considerably larger 

than this. However, it is probably more appropriate to use the velocity at the array site for this comparison. 

This is because the flow at head of the Barrow Canyon is locally convergent and the velocity there is 

stronger than farther down the canyon (Pickart et al. 2005). The velocity at CS4 averaged over the depth 

of the WW layer for the appropriate period is 9.3 cm s

-1

, which is closer to the above estimate deduced 

from the water mass signals. 
G22

Figure G18. Potential temperature of the Winter Water at (a) mooring BC2 at head of Barrow Canyon

(near the bottom) and CS4 on the Chukchi slope. The light gray shading means that there is no WW 

present. The dark gray shading indicates when the flow is in the opposite direction of the primary current 

(southwestward and southeastward in (a) and (b), respectively). The black triangles denote the time 

periods when the bulk of WW appeared.

The WW signal in the Chukchi shelfbreak jet (at mooring CS2) appears in late-May and lasts until midSeptember (Figure G17c). The origin of the water in this current is likely the outflow from Herald 

Canyon. Observations suggest that the Pacific-origin water exiting the canyon turns to the right and forms 

a shelfbreak jet along the northern edge of the Chukchi Sea (Pickart et al. 2010; Linders et al. 2017). 

Models also indicate this (Winsor and Chapman 2004; Spall 2007). The mean velocity at the bottom of 

mooring CS2 (near the MicroCAT) was 4.8 cm s

-1

. Using the distance along the shelfbreak from the 

mouth of Herald Canyon to the array site, this gives an advective time of approximately 6 months for the 

WW to reach the array. This would seem to imply that the cold water finished flushing through Herald 

Canyon in mid-March. Based on the limited observations to date in the canyon, this seems unlikely 

(Woodgate et al. 2005; Pickart et al. 2010; Linders et al. 2017). However, considering the large distance 

between the two sites (order 750 km), it could be that the WW leaving the canyon later in the season 

warms and becomes remnant winter water by the time it reaches the array. Notably, Corlett and Pickart 

(2017) also found that the presence of WW in this region decreased markedly after September.
H1

H. Physical Oceanography of the Northeastern Chukchi Sea: A 

Synthesis

H.1 Introduction

The overall goal of this program is to provide an improved understanding of the physical oceanography of 

the northeastern Chukchi shelf, as well as exchanges between the Chukchi/Beaufort shelves and the 

adjacent basin. In previous sections, we addressed specific objectives of our program in detail. In this 

section, we will attempt to integrate the findings from those previous sections with other data sets, 

including a variety of ice profiling sonar (IPS) and ADCP data sets collected by the oil industry as part of 

the CSESP, the Barrow Canyon and COMIDA Hanna Shoal mooring arrays, NSF-supported moorings 

provided by S. Okkonen and C. Ashjian, and the NSF-sponsored Arctic Observing Network (AON) 

mooring deployed along the Beaufort Sea shelfbreak at 152°W (previously discussed by Nikoloupolus et 

al. 2000).

Assimilating these further datasets allows us to provide a general overview of our emerging 

understanding of the northeastern Chukchi Sea shelf and shelfbreak. We begin with a discussion of the 

seasonal ice cover because ice formation and retreat has a major influence on the hydrography throughout 

the year. We add additional hydrographic data (collected in summer and fall) to that previously discussed 

to underscore how sea ice meltwater (MW) isolates Hanna Shoal from the remainder of the shelf. The 

section concludes with an overview of the circulation connections (or lack thereof) among various 

portions of the northeastern Chukchi Sea shelf and highlights connections with prominent hydrographic 

features. Figure H1 shows the location of moorings used in this section of the report.

H.2 Sea Ice

Our description of sea ice variability on the Chukchi Shelf begins with an overview of 1979 – 2013 and 

then focuses on the northeastern Chukchi Sea shelf for the years 2011 – 2013. We focus on these years for 

two reasons. First, these years illustrate extremes in the interannual variability in summer ice conditions. 

Second, our most comprehensive set of hydrographic data comes from these years, and thus can serve as 

an introduction to how extreme ice conditions affect ocean physics.

Sea ice generally begins forming in the northern Chukchi Sea in late October and effectively covers the 

entire shelf, including Bering Strait, by late December. Ice retreat commences in late April or early May. 

Using satellite data from Scanning Multichannel Microwave Radiometer (SMMR), Special Sensor 

Microwave/Imager (SSM/I), and Special Sensor Microwave Imager/Sounder (SSMIS) acquired form the 

National Snow and Ice Data Center we examined spatial patterns of ice retreat by computing Empirical 

Orthogonal Functions (EOF) of May through August 1979 – 2013 sea ice concentrations. (We performed 

this calculation using both mean daily and monthly ice concentration maps and found virtually identical

results.) 

As a first attempt at understanding the relationship of these patterns to the atmosphere, we correlated the 

EOF time amplitude functions with wind velocity and wind speed over Bering Strait and the Chukchi Sea 

shelf. Winds for all analyses were obtained from the North American Regional Reanalysis (NARR) 

products produced by NOAA’s National Center for Environmental Prediction (NCEP). The NARR 

forecasts are provided at 3-hourly intervals on a 35 km grid (Mesinger et al. 2006).
H2

Figure H1. Map showing the locations of moorings used in this section of the report as a function of year. 

Moorings Cj (Crackerjack), Bu (Burger), Site1 & 2, CPAI 01 & 02, Statoil 3 & 4, and HS 01-06 were all 

deployed and processed under industry support. HSNE and HSNW moorings were supported by the 

BOEM-COMIDA program. WBC and EBC were supported by NSF. Barrow Canyon moorings (BC 1 – 6) 

were supported by BOEM Study 2012-079, and BC2 from 2012 – 2016 was supported by this project.

There are two statistically significant EOF modes of ice-retreat in May (Figure H2). The dominant mode 

(EOF1) accounts for 46% of the variance in ice concentration and corresponds to ice retreat and/or ice 

melt proceeding northward from Bering Strait. The second mode (EOF2) accounts for 27% of the

variance and describes a pattern in which changes in ice concentration along the northwestern coast of 

Alaska are out-of-phase with changes in concentration along the northern coast of Chukotka. Correlations 

between the May EOFs and the monthly wind velocity component anomalies indicate that EOF1 is 

positively correlated with both the east–west (correlation coefficient, r = 0.43) and north–south (r = 0.44) 

wind components in Bering Strait (and also over the central Bering shelf, where the correlation 

coefficients are smaller, but nevertheless still significant). The signs of these correlations are consistent 

with the hypothesis that wind-induced forcing in the Strait and over the Bering shelf affects northward 

transport (and the northward heat flux) through the Strait (Danielson et al. 2014). EOF2 is significantly 

correlated with winds over the northeastern Chukchi Sea, with winds from the east (west) and or south 

(north) leading to reduced (increased) ice concentrations along the northwest Alaskan coast. The 

correlation coefficients are r = 0.68 and r = 0.64 for the zonal and meridional wind components, 

respectively. Although not explored further here, a possibly important consequence of EOF2, when it 

results in open water along the northwest Alaskan coast, is that it exposes the underlying water to more 

intense solar radiation much earlier than would otherwise occur. The increase in solar radiation should 

then lead to earlier pelagic phytoplankton blooms and earlier warming of coastal waters. The structure 

and evolution of EOF2 is analogous to the behavior of a polynya in the absence of freezing air 

temperatures (and an infinitely wide ocean). Under such conditions there is no limit on polynya width as 

long as the winds continue blowing offshore (Pease 1987).
H3

Figure H2. The spatial structure of the most significant EOFs of ice concentration for May (top row), June 

(middle row), July (bottom left), and August (bottom right). Where sketched, the white dotted lines outline 

the region of maximum spatial variability in EOF1.
H4

The June EOFs (Figure H2) have patterns similar to those of May. The June EOF1 describes variability in 

northward ice retreat and accounts for 58% of the variance in ice concentration. The modal weights 

indicate that maximal variability occurs in the area outlined by the dotted white lines (i.e., over Hope Sea 

Valley, with an extension to the northeast in the Central Channel). The spatial structure of EOF1 shows 

that the patterns of ice retreat are beginning to outline the principal flow pathways (in Herald Canyon, the 

Central Channel, and the west coast of Alaska). EOF1 is uncorrelated with the winds over the Bering 

shelf and in Bering Strait, but it is significantly correlated with the zonal and meridional winds in the 

Chukchi Sea (r = 0.37 and r = 0.24, respectively). The decrease in these correlations from May suggests 

that other atmospheric influences (solar radiation) and/or ocean-based processes (heat advection and/or 

mixing) exert more influence on this mode’s variability. EOF2 for June accounts for 17% of the variance, 

and its structure is similar also to the May EOF2, in that it has an anti-phase relationship between the 

eastern and western Chukchi Sea. June EOF2 again shows significant correlations with the zonal and 

meridional components of the wind (r = 0.42 and r = 0.54, respectively) over the Chukchi Sea.

Only the first EOF is statistically significant in July (Figure H2). It accounts for 66% of the variance in 

ice concentration, and the pattern describes meridional variability in ice concentrations. The region of 

maximal variability occupies a zonal belt between 70.5°N and 72°N (outlined by dotted white lines) all 

along the northern perimeter of the Chukchi shelf. There is no significant correlation between EOF1 and 

wind velocities; however, it is significantly correlated (r = 0.43) with wind speed over the Chukchi Sea, 

with the sign implying that a reduction in ice concentration is associated with an increase in wind speed. 

This correlation could arise in two ways. First, mechanical break-up of the ice and the production of 

smaller flows may be enhanced under higher wind speeds. Smaller floes will have higher surface area-tovolume ratios than bigger floes and so melt more rapidly. Second, a significant amount of heat is 

contained in the Pacific-derived waters advected across the Chukchi shelf in June and July (Woodgate et 

al. 2012). Because these waters are generally more dense than the surface waters formed from melting ice 

(Coachman et al. 1975), they reside below the MW layer that surrounds the marginal ice zone and lies 

beneath the melting ice (Paquette and Bourke 1981). Higher wind speeds should enhance turbulent 

mixing, promote the upward flux of oceanic heat from the Pacific waters to the underside of the ice, and 

increase the rate of ice melt.

EOF1 is the only significant mode in August as well (Figure H2), and it accounts for 72% of the variance 

in ice concentration. The region of maximum variability includes Hanna Shoal and lies north of ~72 and 

73°N with north-south meanders that approximately align with the principal flow pathways. The only 

significant correlation (r = 0.42) between this mode and wind velocity is with the east–west component of 

wind over the northeastern Chukchi Sea shelf. The sign of the correlation is such that eastward winds are 

associated with a decrease in ice cover. This would be consistent with an Ekman response of the ice; its 

drift would include a southward component that would make the ice-edge more diffuse and advect floes 

southward toward warmer water. There also is a positive correlation (r = 0.56) between EOF1 and wind 

speed.

Although we have identified a number of significant correlations between EOF modes and the winds, the 

relationships do not account for a large fraction of the ice concentration variance. Indeed the largest 

correlation coefficients were those between EOF2 and the wind velocity components in May and July 

(which account for ~35% of the variance in that mode at most) and the wind speed in August. As 

mentioned, these low correlations may indicate that other physical processes associated with either the 

atmosphere or the ocean are as important, if not more so, than the winds, or it may imply that the wind’s 

effects are manifested in non-linear ways. Bear in mind, however, that our analysis has not been 

exhaustive, and by concentrating on EOF modes as bulk measures of sea ice concentration patterns, our 

statistical approach may have masked the importance of winds in affecting variations in the seasonal 

retreat of sea ice. The results do remind us that ice retreat on this shelf is a complicated process and that 

the role of winds, currents, and air-sea heat exchange play in this retreat have not been adequately 
H5

addressed. As discussed in section F and explored further later, the absence or persistence of sea ice 

critically influences the hydrographic structure of the Chukchi Sea shelf.

H.3 Winter Sea Ice Statistical Properties

Prior to discussing statistical properties of the ice as determined from IPS sensors, we examine some of 

the broader scale aspects of the onset of ice over the northeastern Chukchi Sea for autumn 2011, 2012, 

and 2013. Recall that the main difference among these years was that warm Bering Sea Summer Waters 

pervaded the entire surface layer of the northeastern Chukchi in summer and fall 2011, whereas in 2012 

and 2013 cold MW was present over the Hanna Shoal region between the Central Channel and Barrow 

Canyon. In each year, the Chukchi shelf was ice-free through the end of October, according to SSMIS and 

Advanced Microwave Scanning Radiometer 2 (AMSR2) imagery (Figure H3) (Spreen et al. 2008). In 

2012, the ice advanced rapidly southward with much of the shelf being partially ice-covered by midNovember and completely ice-covered by the end of the month. By contrast, the ice advanced slowly 

southward through mid-November in 2011 and 2013. A direct relationship between wind velocity and the 

rate of advance of sea ice is not readily apparent. For example, moderately strong northeasterly winds 

prevailed through mid-November 2011, while weak and variable winds occurred over the same period in 

2012. One might have expected that the winds in 2011 would have promoted more rapid cooling and 

freezing than in 2012, yet ice developed more rapidly in 2012 than in 2011. Regardless of the exact 

causes, the rate of ice advance in these years increased rapidly after mid-November, although substantial 

areas of open water or low concentrations remained, especially in the western Chukchi Sea, through to the 

end of November in 2011 and 2013. The signatures of the principal current pathways across the shelf are 

also evident in the 15 November ice concentration maps as embayment-like features occurred in all three 

years. Although these embayments have been noted frequently with respect to summer ice retreat, they 

appear to be a regular fall feature of the shelf as well and are an indication of the influence of the 

northward flow from Bering Strait on patterns of ice advance.

Through our partnership with the oil industry, we were provided with the processed IPS data sets from 

Crackerjack, Klondike (Site 1), Burger, CPAI02, Statoil03, and Statoil04 (Figure H1). The processing 

followed the procedure of Melling et al. (1995) and is given in detail by Mudge et al. (2015). Here we 

highlight just a few examples of the types of statistical information available from these data, as the 

analyses are ongoing. The statistical properties discussed here were formed (again following Melling et 

al. 1995) from 50 km long sections of ice drift recorded by ADCP moorings that accompanied each IPS 

mooring. In this procedure, we compute ice displacement as a function of time and record the date 

midway along each segment. (The displacement diagrams are identical to progressive vector diagrams 

often used in descriptive oceanography. The diagrams are qualitatively useful in providing a picture of 

movement, but they should not be interpreted as a Lagrangian measurement. Figure H4 shows 

displacement maps derived from the 2010 – 2013 Burger deployments.) All statistics are then designated 

by this date. Note that the time lapse between dates can be quite variable, with some segments spanning 

only a day while others may span 2 weeks. In most cases the duration of a segment was only a few days. 

These different time lapses simply reflect the fact that ice motion is more rapid during some periods than 

others. As noted by Melling et al. (1995) the reason for adopting this “sampling-through-distance” as 

opposed to “sampling-through-time” is that if the ice does not move then one is effectively sampling the 

same piece of ice repeatedly and generating biased statistics. Although the IPS records keel depth, we 

converted these to thickness by multiplying the keel depth by 1.1 (Bourke and Paquette 1989). For each 

segment we computed the mean thickness (including open water), the first and second modes, and 

examples of ice thickness histograms or probability density functions (pdfs). We show a subset of these 

data and offer general descriptive statements which hold across time and for different sites.
H6

Figure H3. The evolution of sea ice concentrations as seen on 30 October (top row), 15 November 

(middle row), and 30 November (bottom row) for 2011 (left column), 2012 (middle column), and 2013 

(right column).

The displacement figures provide a qualitative sense of the dynamic nature of the ice drift over the 

Chukchi Sea shelf (Figure H4). For example, during January and February 2013 the suggested 

displacement was ~500 km at a mean speed of ~10 cm s-1. In each year, the drifts were southwestward 

beginning in late fall. In early 2011 the drift reversed and was northeastward from January through April

and weak thereafter. For the 2011 – 2012 and 2012 – 2013 records the drifts were primarily toward the 

southwest. By late spring the ice motion was weak, and the drift was variable in direction.

Figure H5 shows the mean ice thickness for each drift segment from the Burger 2010 – 2013 IPS records, 

determined from the 50 km segments of Figure H4. Note that each plot begins on the common date 

October 30, although ice was not always recorded at that time. The last date plotted corresponds to the 

last date in which a 50 km segment was obtained. We have also fitted these data with a power law and 
H7

Figure H4. Ice displacement as a function of time for the Burger deployments from 2010 – 2011 (top), 

2011 – 2012 (middle), and 2012 – 2013 IPS and ADCP deployments. The stars indicate the starting 

position and correspond to the first 50 km segment in which any ice was recorded. Red dots indicate the 

mid-point of each 50 km segment and every other segment is labeled by date.
H8

Figure H5. Mean ice thickness and statistical fits as a function of time (50 km segment) for the Burger 

deployments from 2010 – 2011 (top), 2011 – 2012 (middle), and 2012 – 2013 (bottom). Means encircled 

by a red ellipse suggest periods of new, thinner ice formation, and those with a blue ellipse suggest 

periods when thicker ice was advected over the site.
H9

logarithmic fit between the first date ice was recorded and about the 20th of May. Inspection of these and 

other plots indicated that after this date the mean thickness begins to decrease. Overall, the power law 

and/or logarithmic fits perform equally well (including other years not shown). We ascribe deviations 

from these fits to major advective events in which thicker ice was advected past the mooring site, or 

numerous leads formed in which new and thinner ice developed. In general the mean thickness increases 

to about 2 m by late May and then rapidly declines thereafter due to melting.

Figure H6 shows the seasonal evolution of ice thickness based on subsets of the pdfs of ice thickness as a 

function of selected dates from Burger for the 2010 – 2013 period. These are plotted separately for three 

trimesters: 1) November through mid-January, 2) late January through late March, and 3) from April –

June (or July). To facilitate comparison we prepared the figure by choosing segment dates from each year 

as close to one another as possible. Several features emerge upon inspection. First, the pdfs evolve most 

rapidly (toward progressively thicker ice) during the first trimester as ice first develops. Initially, ice 

thickness increases and the pdfs consist of a single mode which thickens through time. Soon afterwards, 

additional modes develop, as new ice forms and/or thicker ice deforms. However, by mid-January most of 

the ice thicknesses were ≤1 m. During the second trimester, the ice pdfs become more regular and

generally contain 2 modes, one at >~1 m and the second being thinner. These thinner modes presumably 

represent new ice forming in leads. In most years the ice was ~1.5 m thick by late March, although in 

2011 the primary mode was ≤1 m. The thick mode at 3 m on 17 March 2012 is unlikely a result of local 

thermodynamic processes based on the theoretical estimate provided by Maykut (1986). In using his 

formulation we assumed a seasonal average of 3,000 freezing degree-days based upon a freezing season 

of 150 days in length and 20°C as the average difference between the air temperature and freezing point 

of seawater. If we assume a snow cover of 10 cm, then the ice thickness at the end of the freezing season 

would be ~0.8 m. (A thicker snow cover yields thinner ice and vice versa.) Repeating the calculation for a 

temperature difference of 30°C results in a thickness of 1 m of ice at the end of March.

By the third trimester two modes generally occur with the dominant one being at ~1.5 m and the 

secondary mode at ~0.5 m. Figure H6 summarizes the seasonal evolution of the primary and secondary 

modes for each year. These time series are based on all 50 km segments from each deployment and thus 

not sub-sampled as in Figure H5. The overall impression from these pdfs is that the seasonal evolution of 

ice involves the development of numerous ice thickness classes, with a general thickening proceeding 

throughout the season. However, modes can form and disappear rapidly, with ice thickness classes arising 

and disappearing over relatively short time intervals. This conclusion is corroborated in Figure H7, which 

shows the seasonal evolution of the primary (blue) and secondary (red) modes based on all 50 km 

segments in each year. As with the means, we restrict the least squares fits to the period spanning from 

the first detection of ice in the fall through the latter part of May. Although the thickness of both modes 

increases throughout the season the progression is far from regular and includes episodes of rapid 

thinning and/or thickening, which yields poor fits to the data. 

The episodic nature of mode evolution is a consequence of advection in which thicker or thinner floes are 

brought past the site or local convergence or divergence. As floes converge they deform and thicken and 

as they diverge leads develop and new ice forms. New ice formation is critical to the formation of dense 

water that forms across the shelf throughout winter. Upon formation the dense water is at the freezing 

point and is enriched in salinity due to brine repulsion from the forming sea ice. Ice, and in particular, the 

thick floes formed by deformation processes, is a source of cold, dilute surface when they melt the 

following summer. 
H10

Figure H6. Probability density functions (pdfs) for selected dates from the Burger 2010 – 2011 (left), 

Burger 2011 – 2012 (middle), and Burger 2012 – 2013 (right) IPS deployments. The y-axes represent the 

percentage of time a particular ice thickness interval (x-axis) occurred over the 50 km long sampling 

interval. Numbers in parentheses in legend indicate the percentage of all observations less than or equal 

to the maximum value of the ice thickness on the x-axis. Note that the x-axis maximum is 2 m for the top 

row and 4 m for the bottom two rows.

H.4 Hydrography

Summer and fall water masses on the Chukchi Sea shelf consist of an amalgam of different water masses. 

As a prelude to this section we first review these with the aid of the potential temperature-salinity (θ/S) 

diagram in Figure H8. These water masses consist of cold (<4°C), dilute (<30) MW layer, warmer and 

saltier Bering Sea Summer Water (BSSW) waters that have entered the Chukchi Sea through Bering 

Strait during the summer months. BSSW waters include Alaskan Coastal Water (ACW) and Bering Sea 

Water (BSW), where the former is fresher and warmer than the latter. Large portions of the shelf also 

include near-freezing, salty (and dense) bottom winter waters (WW) formed the previous winter on both 

the Chukchi and Bering shelves. This water mass is gradually displaced northward throughout the 

summer. WW can be warmed at depth by penetrating solar radiation, such that it has the very high 

salinities characteristic of WW, but with temperatures elevated above the freezing point. Gong and 

Pickart (2015) termed this water mass Chukchi Summer Water (CSW). As evident in the figure, these 

water masses are not discrete, but instead blend together and tend to overlap in property space as a 

consequence of vertical and horizontal mixing and/or heating by solar radiation. (Finer-scale 

discrimination of water masses is possible based on nutrient and δO18 data, which were not collected as 

part of this program.).
H11

Figure H7. Time series of the primary (blue) and secondary (red) mode ice thickness modes along with 

logarithmic statistical through time. The modes are determined from each 50 km segment for the Burger 

deployments from 2010 – 2011 (top), 2011 – 2012 (middle), and 2012 – 2013 (bottom).
H12

Figure H8. Potential temperature-salinity (θ/S) diagram illustrating various water masses found on the 

northeastern Chukchi Sea shelf in summer and early fall.

As discussed in section F, the rate of disappearance of ice over Hanna Shoal appears to be a controlling 

influence on vertical stratification and the formation of upper ocean fronts on the northeastern Chukchi 

Sea shelf. When large volumes of MW are present, the stratification is strong and shallow, with the 

pycnocline centered between 10 and 15 m depth. The MW region extends to the south of Hanna Shoal 

and is separated by a strong front from warmer and saltier BSSW waters over the shelf to the south and 

west. The western edge of this front appears to extend from the southwest corner of Hanna Shoal 

southward to ~71.5°N and thence southeastward to ~71.2°N toward the head of Barrow Canyon. Here we 

show that this MW layer extends eastward from Hanna Shoal to Barrow Canyon and that the front, on 

approaching the head of Barrow Canyon, curves northeastward along the length of the canyon. This 

conclusion is based on Figures H9 – H14, which are a sequence of Acrobat-towed CTD legs collected in 

2012 that extend from near the head of the canyon (Leg Y; Figure H9) to the mouth of the canyon (Leg 

U; Figure H14). The figures are arranged such that the viewer is looking upcanyon toward the southwest, 

with the coast to the left in the figure and consist of separate vertical sections for potential temperature, 

salinity, and chlorophyll, a map showing the location of the transect, and a θ/S diagram.

On each section ACW occurred closest to the coast and occupied the entire depth of the water column. 

BSW and/or WW was below the ACW along the canyon floor. MW was a consistent feature of the upper 

water column to the west of the ACW with these water masses separated by a strong front, and WW 

occurred beneath the MW. There were indications of horizontal mixing processes evident in sections E, 

Q, S, and U (Figures H10, H12, H13, and H14, respectively) where lenses of warmer ACW and/or BSW 

infiltrated below the MW layer on the west side of the canyon. When present, the chlorophyll distribution 

tended to be dense, but horizontally patchy, and clustered within and/or just below the pycnocline 

separating MW from WW. The reason for this is that WW contains high nutrient concentrations, whereas 

MW and ACW are always nutrient-poor, and the nutrients in BSW are exhausted by the time this water 

mass reaches the northeastern shelf in late summer.

MW

ACW

BSSW

WW

CSW
H13

Figure H9. Summary figures for Acrobat Leg Y. The left hand side consists of vertical sections of 

temperature (top), salinity (middle), and chlorophyll (bottom). The black contours on each section are 

isolines of potential density plotted at a contour interval of 2. The right hand sides of each figure contain a 

transect location map (top right) and a θ/S plot based on the data from that section. The view is toward 

the southwest with land to the left of each panel.

Figure H10. As in Figure H6 but for Acrobat Leg E.
H14

Figure H11. As in Figure H6 but for Acrobat Leg BB.

Figure H12. As in Figure H6 but for Acrobat Leg Q.
H15

Figure H13. As in Figure H6 but for Acrobat Leg S.

Figure H14. As in Figure H6 but for Acrobat Leg U.
H16

Based on the results of Section F and the Acrobat sections presented here, we conclude that when ice 

remains over Hanna Shoal late in summer a front separating MW from BSSW forms an arc extending 

from the southwest side of Hanna Shoal, eastward to Barrow Canyon and thence along the western side of 

the canyon to at least the shelfbreak, circumscribing a distance of ~450 km. Beneath the MW layer, the 

shelf largely consists of WW. Along the south side of Hanna Shoal there is a bottom front separating the 

WW from BSSW, although this front is not necessarily collocated with the surface front. As discussed by 

Lu et al. (2015), these fronts are potentially unstable, resulting in horizontal mixing. In Section C, we 

noted that the within the MW region, the surface circulation tends to be weak and largely a result of 

Ekman dynamics forced by the winds. South of the front the circulation is primarily zonal and forced by a 

combination of geostrophy and Ekman dynamics, while within the canyon the flow is primarily 

geostrophic. In each case the geostrophic pressure field is established by a combination of the windforced sea level setups and the mean pressure gradient between the Pacific and Arctic Oceans. 

The MW/BSSW fronts affect the circulation field in several ways. The cross-frontal density gradient 

forces an along-front jet having the same width as the front and velocities of 10 – 20 cm s-1 (the direction 

of flow is such that the MW is to the right of a viewer looking downstream along the jet’s axis). The jets 

are baroclinically unstable and generate energetic meanders and eddies, which are fundamentally 

important in mixing (Thomas 2008; Thomas et al. 2008). The MW/BSSW front involves horizontal flow 

convergence exemplified by surface current vector maps derived from HFRs averaged during the week of 

5 September 2012 from radars in Barrow, Wainwright, and Point Lay (Figure H15). The vector 

distribution suggests a convergence zone extending from west to east along about 71.5°N. North of this 

zone, but near the front, the surface flow is southward (and downwind based on the mean wind vector at 

Barrow during this time). South of the zone, the flow is northeastward and typical of the shelf flow 

offshore of Wainwright and south of Hanna Shoal under mild wind conditions. The convergence is 

roughly aligned with the position of the fronts discussed above and signatures of it appear to extend 

eastward to Barrow Canyon. Surface current vector maps obtained in August and September 2011 

showed no such features, which is consistent with the hydrography from that year. As might be expected 

the presence of mesoscale motions should 

Figure H15. Mean surface current vectors for the week of 5 September 2012, as estimated from HFRs.
H17

yield differences in surface current variability. In 2011, the variance in both the zonal and meridional 

velocity components across the radar mask was half that for the same velocity components in 2012. These 

differences cannot be attributed to differences in the wind energy between these two years for the 

variances in both wind vector components were identical in both years.

As a second example, we show the trajectory of a 10 m drogued, SVP satellite-tracked drifter that was 

released from the USCG Healy on 13 August 2012 (Figure H16) southwest of Hana Shoal. Over its first 

45 days, the drifter’s trajectory included a number of meanders and large excursions in sea surface 

temperature (SST), few of which were associated with changes in winds. The drifter’s trajectory is 

especially convoluted in the vicinity of 71.3°N, 163.5°W where the SST variations are large and vary 

from ~5°C south of this latitude to ~2°C north of it. This temperature range brackets temperatures typical 

of MW and BSSW. The variations in SST and the drifter’s trajectory path are suggestive (but not 

conclusive) that the drifter was entrained into mesoscale motions associated with frontal instability. The 

unstable motions give rise to current fluctuations across both the sub-mesoscale (<6 km) and the 

mesoscale (6 – 18 km) spatial scales. Such motions should enhance the dispersion of dissolved and 

suspended material.

The strong stratification found over the shelf north of 71.5°N is a perennial feature due to MW and/or 

BSSW in the upper 15 – 20 m and WW at depth. The preceding results indicate that the stratification 

northeast of Hanna Shoal is stronger than that to the northwest of the Shoal. In part this difference is 

because the upper layer salinities west of the Shoal are influenced by the moderately salty BSW flowing 

northward in the Central Channel. East of the Shoal, the BSW influence is very limited and here the 

surface layers may receive contributions from the more dilute ACW. With the onset of freezing and 

increased wind-mixing, the stratification over the shelf should break down. Such a transition appears to 

occur annually on the south of Hanna Shoal (Weingartner et al. 1998; Weingartner et al. 2005).

We next examine the seasonal variations in stratification on the northwest and northeast sides of Hanna 

Shoal by comparing the seasonal changes in vertically-averaged currents and temperature and salinity at 

moorings HSNW50-12 (Figure H17) and HSNE50-12 (Figure H18). The latter figure includes the time

Figure H16. Color-coded (by SST) trajectory of a satellite-tracked drifter released at the location of the 

green triangle on 13 August 2012. The drifter’s last reported position was on 15 November 2012 at the 

location of the red star.

Oct.	16

Nov.	8
H18

Figure H17. Time series of vertically averaged current vectors (top), temperature (middle) and salinity 

(bottom) at 47 m depth from mooring HSNW50-12. The timing of the temperature maximum (Tmax) and 

the descent to the freezing point (Tf) are noted.

Figure H18. Time series of vertically averaged current vectors (top), temperature (middle) and salinity 

(bottom) at 23 m (red) and 47 m (blue) depth from mooring HSNE50-12. The timing of the temperature 

maximum (Tmax) and the descent to the freezing point (Tf) are noted.
H19

series of temperature and salinity from a CTD moored at 23 m below the surface and another moored at 

43 m depth. Except for a southwestward flow reversal in September 2012 and a burst of northward flow 

in March 2013, the vertically-averaged flow at HSNW50-12 was steadily eastward with a mean speed of 

~7 cm s-1. At the beginning of the record in August 2012, the bottom temperatures were near freezing (~-

1.8°C), and aside from some shorter period excursions through fall, bottom temperatures increased 

steadily to a maximum of ~0.5°C by late December. In January the temperature decreased and reached the 

freezing point in early February. Salinities were ~33.5 in September but gradually decreased to a 

minimum of ~31 in mid-January. Over the remainder of the record, the salinity increased gradually and 

reached ~32.5 by September 2013. The gradual fall transition from cold, salty water to warmer, fresher 

conditions reflects advective replacement of WW by BSW. Almost certainly, the stratification eroded 

through fall as well and complete breakdown occurred in January as near-freezing, dilute waters in the 

upper portion of the water column mixed with the warmer, saltier bottom waters. We assume that the 

collapse of bottom temperatures to the freezing point signifies the complete breakdown in stratification.

A very different transition occurred on the shelf east of Hanna Shoal. At mooring HSNE50-12 the 

vertically-averaged velocity was variable throughout the year and not statistically different from zero. In 

fall, the flow was primarily south-southeastward and this was accompanied by a gradual increase in 

temperature and decrease in salinity at the bottom (similar to the trend on the northwest side of Hanna 

Shoal). The velocity record and these trends suggest that bottom waters were being advected from the 

west to the east side of Hanna Shoal through fall. From December through August; however, the flow on 

the east side of the Shoal was weak and variable, and bottom temperatures and salinities evolved quite 

differently from those on the west. Bottom temperatures were ~0.5 – 1°C above freezing until mid-June 

when they reached the freezing point and salinities increased. These changes heralded the arrival of WW. 

This WW was advected into the area because it arrived about two months after the onset of melt. Its 

source was probably the west side of Hanna Shoal because the velocity record shows an increase in the 

number of southeastward flow events beginning in April, with these events increasing in frequency and 

strength through July.

The salinity records at 23 and 47 m provide insight on the vertical stratification changes (because of the 

density dependence upon salinity). In fall, the vertical salinity difference between the two depths was ~1, 

and this difference increased to ~5 by mid-December. These salinity differences are comparable to the 

values observed over the water column in the fall and reflect the gradual deepening of the pycnocline 

through the fall. (Recall that the fall CTD sections suggested that the pycnocline was ~5 m thick and 

centered at ~15 m depth.) The late fall change in the salinity difference was primarily due to freshening at 

23 m even though the bottom salinity had also decreased slightly through fall! The freshening at 23 m 

was accompanied by increasing temperatures, which reached a maximum of ~1.5°C in late November 

before decreasing into January. By mid-January temperatures reached freezing and salinities increased at 

23 m. All of these changes suggest a deepening and weakening of the pycnocline by vertical mixing. 

What is remarkable about this record is that the vertical salinity difference reached a minimum of <1 for 

only a brief period in March but otherwise remained strong throughout the year. The March weakening in 

stratification coincided with northward flow and, at both measurement depths, an increase in salinity and 

temperatures near the freezing point. These changes came about due to advection of a weakly stratified 

water column from the south.

Although these records indicate that vertical stratification over this portion of the Chukchi Sea shelf 

remains strong and intact throughout the year, local vertical mixing does occur and is important. For 

example, the decrease in mid-depth temperature from its maximum in late November to the freezing point 

in early January was almost certainly due to local freezing processes and penetrative convection, which 

was insufficient to entirely erode the pycnocline.
H20

There are two other aspects of the mid-depth temperature and salinity records worth noting. First, middepth temperatures remained at the freezing point through early June but began to increase at about the 

same time that the bottom temperatures collapsed to the freezing point. We speculate that these changes 

are unrelated to each other and that the mid-depth temperature increase was due to warming by solar 

radiation as sea ice began breaking up and meltponds allowed solar radiation to penetrate through the ice 

and into the water column (Light et al. 2008). Second, the 23 m instrument captured the signature of very 

high-frequency oscillations that were most prominent from November through early January. These largeamplitude fluctuations are manifestations of internal waves, evident because the CTD was near the middle 

of the thin, but strong pycnocline.

The bottom temperature and salinity records from the 2013 – 2014 HSNE deployments (Figure H19) 

show the same basic seasonal evolution in temperature and salinity as observed in 2012 – 2013. However, 

the seasonal evolution was quite different at the HSNW moorings. The annual maximum in temperature 

occurred in mid-November 2013, nearly two months earlier than in the previous year. Temperatures then 

dropped to the freezing point in early December and remained there into July. We tentatively ascribe 

these differences to differences in fall ice concentration, as detected by the moored ADCPs, between 2102 

and fall 2013. In 2012, the ADCP records indicated nearly 100% ice coverage over the shelf around 

Hanna Shoal by mid-November, consistent with the satellite estimates of ice coverage shown in Figure 

H3. In 2013, the attainment of the 100% ice cover occurred later and included episodic advances and 

retreats. Moreover, complete ice cover developed earlier on the east side of the Shoal than in the Central 

Channel and northwest side of the Shoal (Figure H3). These differences in ice cover (in conjunction with 

the assumed differences in stratification) on either side of the Shoal would lead to earlier cooling at the 

HSNW moorings compared to the HSNE moorings.

Figure H19. Time series of bottom temperature and salinity from the COMIDA 2013 – 2014 mooring 

array.
H21

H.5 Circulation Connections

H.5.1 Chukchi Sea Shelf

Our description of the circulation field over the northeastern Chukchi Sea is based primarily on the mean 

vectors and variance ellipses of vertically-averaged velocities, and those from the topmost, mid-depth, 

and bottommost ADCP bins. (Typically the topmost bin represents data from ~7 m depth, although 

deeper velocities are used if the ice thickness extends below this depth.) Maps of these quantities are 

shown for the mooring years 2010 – 2011, 2011 – 2012, 2012 – 2013, and 2013 – 2014 (Figures H20 –

H23, respectively). We selected each of these years for discussion because: 1) they contain the largest 

numbers of moorings deployed, 2) the mooring locations changed somewhat from year-to-year, and 3) not 

all moorings yielded full-length records.

The 2010 – 2011 records (Figure H20) indicate that the mean flow within the Central Channel is northnortheastward at ~5 cm s-1

, and over the shelf south of Hanna Shoal the flow is eastward at similar speeds. 

At these locations, the standard deviations are 3 – 4 times the magnitude of the means, and the ellipses are 

aligned in approximately the same direction as the means. The eastward flow transports waters from the 

Central Channel to the coast and accelerates upon nearing Barrow Canyon, where, as discussed in section 

D, it turns northeastward into the canyon within about 40 km of the coast. The vertical current shears in 

the Central Channel, at the head of Barrow Canyon, and over the central shelf are weak. The lack of shear 

over the central shelf is noteworthy given that the flow is comparatively weak and nearly opposite the 

surface wind stress. The only locations where there was substantial vertical shear were the two moorings 

about midway down Barrow Canyon and west of Point Barrow (WBC and EBC; Figure H1). Here the 

Figure H20. Mean vectors and variance ellipses for the 2010 – 2011 moorings. Full-length records are in 

black, partial-length records are in gray, and NARR wind records are in red.
H22

surface flow was westward. Closest to the coast, the vertically-averaged and deeper velocity vectors were 

northeastward, or downcanyon, and over the western side of the canyon the vertically-averaged and 

deeper velocity vectors were weak and oriented upcanyon. The westward drift at the surface at these 

locations implies the advection of fresh surface waters onto the shelf east of Hanna Shoal. This advective 

flux would have carried ACW onto the shelf east of Hanna Shoal and thus contributed to the warm waters 

observed there in 2011.

The shelf circulation portrayed by the 2011 – 2012 moored array (Figure H21) was similar to that 

described above insofar as the flow was north-northeastward in the Central Channel, eastward over the 

shelf south of Hanna Shoal, and northeastward at the head of Barrow Canyon and near Point Barrow. This 

deployment included three moorings along the northwest side of Hanna Shoal (HS01 – HS03; Figure H1).

Closest to the Shoal, the mean flow was negligible at all depths. The two northernmost moorings 

measured a mean drift to the northeast of 5 – 10 cm s-1 and detected little vertical shear. There were three 

other moorings along the south and southeast side of the Shoal (HS04 – HS06), which connected to the 

western end of the Barrow Canyon array (BC1 – BC6). The flow at each of these three moorings was 

southeastward and similar in direction to the flow at the outer end of the Barrow Canyon array. This is the 

same southeastward flow that Pickart et al. (2016) identified from synoptic sections in summer, and the 

moorings indicate that this is a persistent feature of the central shelf’s circulation field. It further 

substantiates previous suggestions that waters from the Central Channel leak eastward along the entire 

length of the Channel (Weingartner et al. 2013a). The data from these three moorings and from the outer 

three moorings of the Barrow Canyon array demonstrate that there is no mean flow orthogonal to this 

mooring line. This implies that waters on the shelf east of Hanna Shoal do not communicate directly with 

waters to the south of the Shoal. The one other mooring deployed in this year was in ~40 meters water 

depth and east of Hanna Shoal at ~71.7°N, 161°W (CPAI02; Figure H1). The flow there was negligible at 

Figure H21. Mean vectors and variance ellipses for the 2011 – 2012 moorings. Full-length records are in 

black, partial-length records are in gray, and NARR wind records are in red.
H23

all depths, except at the bottom where it was ~1 cm s-1 southeastward, i.e., directed toward Barrow 

Canyon.

The 2012 – 2013 and 2013 – 2014 deployments (Figures H22 and H23) also showed that the flow was 

north-northeastward in the Central Channel, eastward over the shelf south of Hanna Shoal, and 

northeastward at the head of Barrow Canyon. For both deployments, the vertically-averaged flow 

northwest of Hanna Shoal at 72.5°N (in 50 m water depth) was northeastward at ~5 cm s-1, and the 

vertical shear was small. Nearer to the Shoal and in shallower water, the vertically-averaged flow was 

weakly east-southeastward and sheared, with the flow strongest at the bottom. Northeast of Hanna Shoal, 

the vertically-averaged flow was negligible in the mean, although the surface flow was northwestward, 

and the bottom flow was southeastward. Finally, the mooring east of Hanna Shoal (deployed only in 2012 

– 2013) had westward surface flow, eastward bottom flow, and negligible vertically-averaged flow.

Based on these reviews of hydrography and the mooring deployments made throughout the years, we 

tentatively offer the following description of the mean flow over the northeastern Chukchi Sea shelf. It 

includes a remarkably consistent north-northeastward flow in Central Channel. West of 162°W, a portion 

of the Central Channel flow is carried eastward between 71 – 72°N (south of Hanna Shoal) across the 

central shelf. In late summer, this flow typically carries BSW, ACW, and WW, and possibly MW. There 

appears to be no net flow between the central shelf and the shelf east of Hanna Shoal. Instead the central 

shelf flow converges on approaching the Alaskan coast, accelerates to the northeast, and enters Barrow 

Canyon within 40 km of the coast offshore of Wainwright. Another portion of the northward flow in the 

Central Channel continues eastward around the northwestern flank of Hanna Shoal. Here it likely merges 

with an eastward-flowing waters derived from Herald Canyon. This description is consistent with results 

from the ocean circulation models of Winsor and Chapman (2004) and Spall (2007).

Figure H22. Mean vectors and variance ellipses for the 2012 – 2013 moorings. Full-length records are in 

black, partial-length records are in gray, and NARR wind records are in red.
H24

Figure H23. Mean vectors and variance ellipses for the 2013 – 2014 moorings. Full-length records are in 

black, partial-length records are in gray, and NARR wind records are in red.

The flow over the shelf between the eastern side of Hanna Shoal and the western wall of Barrow Canyon 

is extremely weak on vertical average. However, the surface waters (and ice) drift to the west or 

northwest, while the bottom waters flow east or southeast. These results imply that the flow field north of 

Hanna Shoal is convergent, at least within the surface layer (Section F) but quite possibly over much of 

the water column. The converging fluid presumably moves northward toward the shelfbreak, where, 

depending upon its density, it will either feed the dilute waters of the westward, surface-flowing Chukchi 

Slope Current or the deeper and denser waters engaged in the eastward-flowing shelfbreak flow discussed 

in Section B. It seems very likely that some of this eastward shelfbreak flow enters Barrow Canyon on the 

west side of the canyon. As it proceeds upcanyon, this flow weakens by losing mass to the much larger 

and swifter downcanyon transport exiting the Chukchi shelf.

The observed flow over the shelf east of Hanna Shoal is not in complete agreement with the models that 

predict a mean clockwise flow around the south side of Hanna Shoal which then retroflects to merge with 

the eastward flow over the central shelf and enter Barrow Canyon. Our observations suggest that the 

bottom flow is clockwise around the Shoal but that it does not continue south of Hanna Shoal. It appears, 

instead, that the bottom waters flow eastward and spill into the canyon. Quite possibly this spillage occurs 

over the entire western side of Barrow Canyon, although this cannot be determined from the observations.

The discrepancy between the models and observations pertains to the surface and vertically-averaged 

circulation on the east side of Hanna Shoal. The counterclockwise circulation in the surface layer 

northeast of Hanna Shoal is a consequence of the west to east divergence of the isobaths and the 

baroclinic pressure gradients over the shelf east of the Shoal. As the isobaths diverge the eastward 

geostrophic flow and the barotropic pressure gradient that forces this flow must weaken. We suggest that 

the weaker barotropic pressure gradient is opposed, and likely overwhelmed, in the upper ocean by 

baroclinic pressure gradients. These gradients arise because ocean densities decrease radially outward 
H25

from Hanna Shoal, as shown in hydrographic sections in Section F. The baroclinic pressure (and density) 

gradients are larger to the east of the Shoal than to its west because dilute surface waters are fed onto the 

eastern shelf occasionally. As noted, this fresher water may be surface waters from the shelfbreak, or it 

may be transported via the mean westward Ekman transport of dilute waters from Barrow Canyon and/or 

the Beaufort Sea shelf displayed in the drifter analysis (Section E). The latter mechanism is also 

illustrated by the mean monthly surface velocities for May, June, July, and August 2011 (Figure H24). 

Recall that in 2011, MW appeared absent in August and September from the shelf around Hanna Shoal. 

The data indicate that the surface flow along the western side of the canyon (mooring WBC indicated by 

the orange box) was westward while the other moorings in the canyon measured northeastward flow and 

the moorings over the central shelf indicated eastward flow toward Barrow Canyon. The surface waters 

advected westward from the canyon in 2011 were ACW (for the most part) and so were much fresher than 

the BSW waters transported through the Central Channel, which bathe the west side of Hanna Shoal in 

summer.

As a final note, the baroclinic pressure gradients weaken with depth and nearly vanish within the bottom 

layer of WW, so that the circulation below the pycnocline is primarily due to the barotropic pressure field 

and thus clockwise. The opposing barotropic and baroclinic pressure gradients are the reason why the 

flow field east of Hanna Shoal is vertically-sheared. The disposition of the barotropic pressure field 

around Hanna Shoal precludes transport between the central shelf and the shelf east of Hanna Shoal. In

Spall’s (2007) model this blockage occurs south of Hanna Shoal. Our data suggests that it occurs farther 

north and to the east, not to the south of the Shoal. Regardless of location, it effectively isolates the shelf 

east of Hanna Shoal from the central shelf. Our results also imply that the bottom waters on the shelf east 

of Hanna Shoal are renewed slowly. We estimate that the area between the 40 and 50 m isobaths on the

Figure H24. Mean monthly vectors and current ellipses (black) in the topmost ADCP bin (~7 m depth) for 

the months of May, June, July, and August 2011. Comparable quantities are shown for the winds in red. 

Mooring WBC is highlighted by the orange box in each figure. Gray ellipses and vectors indicate less than 

a full month’s worth of data was available.
H26

east side of the Shoal is ~1010 m

2

, and we assume that the mean thickness of the WW layer in this region 

is 20 m. The distance between moorings HSNE40 and HSNE50 is ~50 km, and the mean bottom flow is 

~1 cm s1, which implies that the time scale for renewal of the deep water east of Hanna Shoal is ~8 

months. This is somewhat longer than Spall’s (2007) estimate of 6 months for the time it takes water to 

flow from Bering Strait to the Chukchi shelfbreak.

We have examined the coherence scales associated with the circulation over the northeastern shelf by 

calculating the EOFs of the vertically-averaged velocity projected along the principal axis of variance at 

each mooring. We computed annual EOFs for three deployments: 2011 – 2012, 2012 – 2013, and 2013 –

2014. We chose these three years because they have reasonably good coverage around Hanna Shoal, the 

Central Channel, the central shelf, and Barrow Canyon. We have not used all the moorings in each year 

because EOFs are dependent upon the choice of inputs, in that the results can be biased by groups of 

similar moorings that are not present in all years. Interannual comparisons across years are made difficult 

if the sampling sites are not reasonably consistent from year-to-year. For this reason, we have eliminated 

all but one of the Barrow Canyon moorings and all of the industry moorings that extend westward from 

the Barrow Canyon array. We also eliminated the shallower moorings on the east and west sides of Hanna 

Shoal. Because of the irregular shape of the variance ellipses, EOF results at these moorings sites tend to 

be too noisy to add to our results. We conducted the EOF calculations at all depths and locations and find 

that the results are in basic agreement with the following depiction. Figures H25 – H27 show maps of the 

eigenvectors and the amplitude time series for the first three EOFs (which are statistically significant). 

Note that the length of the vector on each map is proportional to the eigenvector’s weight at that location, 

and the direction is that of the major principal axis.

Figure H25. The first three EOF modes of the principal axis velocity component for the 2011 – 2012 

mooring array.
H27

Figure H26. The first three EOF modes of the principal axis velocity component for the 2012 – 2013 

mooring array.

Figure H27. The first three EOF modes of the principal axis velocity component for the 2013 – 2014 

mooring array.
H28

In all years, the first mode explains 50 – 55% of the variance. The velocity fluctuations are coherent and 

in-phase everywhere except on the northeast side of Hanna Shoal. The eigenvector weights are largest 

south of Hanna Shoal, e.g., in the Central Channel, central shelf, and Barrow Canyon. The second mode 

explains 20 – 30 % of the variance. It captures out-of-phase motions between the north and south sides of 

Hanna Shoal, and its weights are largest on the northwest and northeast sides of Hanna Shoal. 

Interestingly, the eigenvector weight for the mooring directly east of Hanna Shoal (CPAI02) is small for 

this mode. Mode 3 captures ~10% of the variance. Current fluctuations between the central shelf and 

Barrow Canyon are in-phase with one another but out-of-phase with those in the Central Channel. The 

relationship between these fluctuations and those around Hanna Shoal are a bit confusing, although it 

appears that the fluctuations on the shelf east of Hanna Shoal are also out-of-phase with those over the 

central shelf and Barrow Canyon.

The corresponding modal amplitude time series indicate that the largest fluctuations are in the fall and 

winter months. As it turns out, all modes are significantly correlated with the along-canyon component of 

the winds discussed in Section D with respect to the transport in Barrow Canyon. The correlation 

coefficients range from 0.51 (2011 – 2012) to 0.64 (2013 – 2014) between this wind component and 

Mode 1. The most straightforward explanation for this relationship is that the currents south of Hanna 

Shoal fluctuate in response to variations in the strength of northeasterly winds, while there is a more 

limited response to these winds around Hanna Shoal. For Mode 3, the correlation coefficients are all 

negative and range from -0.37 (2013 – 2014) to -0.58 (2012 – 2013). We suspect that this relationship 

indicates that there are periods when the winds reverse the currents in Barrow Canyon and the central 

shelf but not those within the Central Channel. We do not have satisfactory explanations for the Mode 3 

response around Hanna Shoal to the winds. The wind Mode 2 correlations were quite variable: -0.38 

(2011 – 2012), +0.18 (2012 – 2013), and -0.30 (2013 – 2014). We cannot provide a satisfactory 

explanation for this correlation but note that it explains less than 20% of the variance in Mode 2.

We have also extended the EOF analyses to include data from Bering Strait collected by R. Woodgate (U. 

Washington) between 2011 and 2014. These results are shown in Figures H28 – H30 and summarized as 

follows. In all years, Mode 1 explains about 50% of the current variance and suggests coherence and inphase fluctuations in Bering Strait, the Central Channel, Barrow Canyon, and the northwest side of Hanna 

Shoal. The structure of Mode 1 is very similar to that of Mode 1 using only the moorings in the northeast 

Chukchi Sea and is significantly correlated with the winds over the northeast Chukchi Sea. The 

correlations range from 0.4 (2011 – 2012) to 0.6 (2012 – 2013) and thus explain from 16 to 36% of the 

total flow variance. Mode 2 explains from 18 – 27% of the variance, and its structure is very similar to 

that using only the moorings over the northeast shelf. In fact, it appears to be confined to the northeastern 

shelf insofar as the modal weights in Bering Strait are very small. Mode 2 correlations with the winds are 

nearly identical to the Mode 2 results using only the northeastern Chukchi moorings. Finally Mode 3 

explains from 10 – 12% of the total variance. The modal structure varies from year-to-year, but generally 

shows, weak northward or strong southward flow in Bering Strait and the Central Channel and either outof-phase or in-phase flow regimes in Barrow Canyon and the northwest side of Hanna Shoal. The 

correlations with the winds vary from year-to-year and are 0 – 0.4 in 2011 – 2012, and 0.46 and 0.4 in 

2012 – 2013 and 2013 – 2014, respectively.

Figures H31 to H33 explore in more detail the correlations between the northward currents in Bering 

Strait and flow along the principal axes in the Central Channel (using the Crackerjack mooring), Burger, 

and in Barrow Canyon, respectively. The correlations are expressed via the coherence and phase spectra, 

which quantify the strength of the correlation and phase differences as a function of period. Currents in 

Bering Strait are generally coherent at all periods with those in the Central Channel (Figure H31), with 

the coherence generally stronger at longer periods (>10 days) than shorter periods. Coherence is generally 

stronger in the October – April (winter) period than for the (May – September) (summer) period. At 
H29

Figure H28. The first three EOF modes of the principal axis velocity component for the 2011 – 2012 

mooring array, which includes Bering Strait.

Figure H29. The first three EOF modes of the principal axis velocity component for the 2012 – 2013 

mooring array, which includes Bering Strait.
H30

Figure H30. The first three EOF modes of the principal axis velocity component for the 2013 – 2014 

mooring array, which includes Bering Strait.

Figure H31. Cross-spectral results between vertically averaged currents in Bering Strait and the Central 

Channel for the entire 2008 – 2014 record (left), the individual May – September periods (middle), and the 

October – April periods (right). Top panel is coherence-squared and bottom panel is the phase.
H31

longer periods the currents are in-phase, whereas at shorter periods, Bering Strait leads the Central 

Channel by about 1 day. Similar results were obtained for the Bering Strait and Burger records (Figure 

H32), although coherences between Burger and Bering Strait are slightly lower overall. Among these 

mooring pairs the strongest coherence was between Bering Strait and Barrow Canyon at all periods >2 

days, with the coherence stronger in winter than summer. We attribute the stronger coherence between the 

Bering Strait and Barrow Canyon pair compared to the other pairs to shelf wave dynamics associated with 

the coastal waveguide. Burger and the Central Channel are outside or at best, at the edge of the 

waveguide, whereas the moorings in both Bering Strait and Barrow Canyon are well within one 

barotropic radius of deformation. The coherence results for Bering Strait and Barrow Canyon are very 

consistent with the conclusions of Section D regarding the influence of continental shelf waves 

propagating northward from the Bering Sea and influencing transport fluctuations in Barrow Canyon.

All the coherence spectra indicate that there is considerable interannnual variability in the coherence 

structure of the shelf circulation field. For example, for the May – September period in 2013, current 

fluctuations in Bering Strait at ~10 day periods were incoherent with those in the Central Channel, while 

in 2011 the currents were significantly coherent with one another for this period. The reasons for this are 

not apparent and would take additional effort to understand. Most likely they are associated with the 

interplay amongst atmospheric forcing, ice patterns, and stratification.

We have sufficiently long and concurrent data at Barrow Canyon, Crackerjack, and the two central shelf 

sites CPAI01 (and Site 1) and Burger (Bu) to construct a tentative climatological annual cycle based on 

the monthly means (Figure H34). The figure contains the annual cycle in transport based on the 2010 –

2015 data from Barrow Canyon discussed in Section D and, from the 2010 – 2014 period, the mean 

monthly velocities projected along the principal axes at the remaining sites. Each site manifests an annual 

cycle. At Barrow Canyon and the central shelf sites (CPAI01 and Burger), the maximum is on July and 

August and the minimum is December and January. These minima are negative, implying a reversal in the 

direction of the flow. The Central Channel differs from these sites in two important ways. First, its annual 

maximum occurs from May through July, and its minimum occurs from September through December. 

Consequently, the annual cycle at all sites except Crackerjack is in-phase with the annual cycle in Bering 

Strait transport (where the monthly averages are all northward). Second, the mean flow in each month is 

positive (e.g., toward the north-northeast) so that the mean monthly flow never reverses, similar to Bering 

Figure H32. Cross-spectral results between vertically averaged currents in Bering Strait and the Burger 

for the entire 2010 – 2013 record (left), the individual May – September periods (middle), and the October 

– April periods (right). Top panel is coherence-squared and bottom panel is the phase.
H32

Figure H33. Cross-spectral results between vertically averaged currents in Bering Strait and Barrow 

Canyon for the entire 2011 – 2013 record (left), the individual May – September periods (middle), and the 

October – April periods (right). Top panel is coherence-squared and bottom panel is the phase.

Figure H34. Climatological annual cycle based on mean monthly values of Barrow Canyon transport (top 

left) and vertically-averaged velocities projected along the principal axis at moorings over the central 

shelf: Burger (Bu), CPAI01 (previously called Site 1), and in the Central Channel at mooring Crackerjack 

(Cj). Dotted horizontal lines on the velocity figures denote the annual mean.
H33

Strait but unlike the other sites. This implies that the entire Chukchi shelf transport must reorganize itself 

seasonally. Reversals in Barrow Canyon imply westward flow over the central shelf and, based on HFR 

measurements, southwestward flow along the Alaskan coast. Because the transport remains northward in 

Bering Strait, reversals in Barrow Canyon and the central shelf imply that this water must escape the 

Chukchi Sea through either the Central Channel and/or Herald Canyon. In a climatological sense, these 

seasonal events must mirror the shorter-lived and massive coastal upwelling events documented in 

Section E based on surface drifters in August 2013 and 2014 (and in Section D based on Barrow Canyon 

transports). During those events, nearshore waters were transported into the Central Channel and/or into 

the western Chukchi Sea and ultimately Herald Canyon.

H.5.2 Chukchi Sea Shelf and Alaskan Beaufort Sea Shelfbreak

In this section we examine the connection between Barrow Canyon transport as determined from BC2 

with the long-term mooring supported by NSF’s Arctic Observing Network Program at 152°W (mooring 

BS3) on the Beaufort Sea shelfbreak (Nikolopoulos et al. 2009). Data from the latter are available from 

2008 – 2014 and from the former from 2010 – 2015. Thus the datasets overlap from August 2010 through 

August 2014. Figure H35 shows the mean monthly transports from both moorings are correlated with one 

another.

The coherence and phase spectra (Figure H36) show that BC2 and BS3 are significantly coherent, 

especially at periods from 2 to 35 days. Overall transport fluctuations in Barrow Canyon explain ~65% of 

the transport variance at BS3 for periods of 2 to 30 days, with BC2 fluctuations leading those at BS3 by 

<0.5 day. As found for the shelf sites, the coherence is poorer in summer than in winter, and there is 

interannual variability in the coherence spectra. 

The coherence calculations suggest that BS3 transports can be predicted from the Barrow Canyon 

transport as illustrated by the linear regressions of the transport between BS3 and BC2 on both a monthly 

(Figure H37) and daily (Figure H38) basis. As expected the predictability is better in winter than in 

summer and when using the mean monthly values. The regression using the monthly means is significant 

at the 95% confidence level with BC2 explaining ~66% of the transport variance at BS3 (Figure H1). 

Note also that BC2 transports are, on average ~5 times greater than the BS3 transports. The regression 

intercept indicates that the transport at BS3 is weakly westward when there is no transport in the canyon. 

The fits are noisier, although still significant, when using the daily values. Seasonally the regressions are 

better in the fall and winter months than for the May – September period. These differences are consistent 

Figure H35. Mean monthly transports at BC2 (blue) and BS3 (red) in Sverdrups.
H34

Figure H36. Coherence squared and phase spectra of mean daily transports at Barrow Canyon and 

along the Beaufort shelfbreak for the entire record (left), May – September (middle) and October – April 

(right).

Figure H37. Linear regressions of BS3 against BC2 mean monthly transports for all months (upper left), 

January – April (upper right), May – September (lower left), and October –December (lower right).
H35

Figure H38. Linear regressions of BS3 against BC2 daily mean transport for all months (upper left), 

January – April (upper right), May – September (lower left), and October – December (lower right).

with Weingartner et al. (2017a) who found that wind-Barrow Canyon transport regressions are better in 

fall and winter than in summer. Weingartner et al. (2017a) argued that the seasonal differences are likely 

associated with the more complicated dynamics associated with enhanced stratification in the summer 

months. 

We have also quantified the frequency distribution of four different flow configurations, or regimes. 

These are defined as follows based on the mean daily transports at each site:

1. Regime 1, northeastward (or downcanyon) at BC2 and eastward at BC3;

2. Regime 2, northeastward at BC2 and westward at BS3;

3. Regime 3, southwestward (upcanyon) at BC2 and westward at BS3;

4. Regime 4, southwestward at BC2 and eastward at BS3. 

The results, summarized in Table H1, indicate that about half of the time Regime 1 holds, e.g., 

downcanyon flow coincides with eastward transport along the Beaufort shelfbreak. However, nearly 20% 

of the time the flow along the Beaufort shelfbreak is westward while the flow is northeastward in the 

canyon. This implies a convergence in transports at the mouth of the canyon. Regime 3 occurs 27% of the 

time and suggests some degree of continuity in the transport with Beaufort shelf waters possibly feeding
H36

Table H1. Number of days observed and the frequency of occurrence of different flow regimes in Barrow 

Canyon (BC2) and the Beaufort shelfbreak (BS3).

Regime BC2 BS3 Days Percent

1 Northeastward Eastward 666 47%

2 Northeastward Westward 297 21%

3 Southwestward Westward 390 27%

4 Southwestward Eastward 69 5%

upcanyon flows. Regime 4 occurs only 5% of the time and implies discontinuous or divergent flow 

between the canyon and Beaufort shelfbreak.

The frequency of occurrence of the flow regimes varies from month-to-month and interannually as shown 

in Figure H39. The figure suggests that Regime 1 dominates in summer months and occurs less frequently 

in the winter months when Regimes 2 and 3 are more abundant. There is considerable variability from 

month-to-month in the regime occurrences. There are several months when the transport is almost entirely 

positive at both locations (Regime 1: June, July, and October 2012) or entirely negative at both (Regime 

3: October 2010 and January and February 2013). Regime occurrence can vary by >50% from month-tomonth.

We investigated the annual cycle in regime occurrences by preparing bar charts that summarize the 

percentage of days that each regime occurs in each month for the whole record and then combine these 

into mean monthly histograms to form a climatology (Figure H40). Regime 1 dominates during the May –

September period, and Regimes 3 and 4 occur most frequently in the winter months, while Regime 2 is 

more equitably distributed throughout the year.

Figure H41 shows the frequency of occurrence of regimes as a function of the direction toward which the 

wind is blowing at the National Weather station in Barrow (PABR). Regime 1 occurs more than 50% of 

the time when winds are blowing from the southern sectors. Regime 2 can occur under any wind direction 

but mostly when the winds are blowing toward the west. Regime 3 dominates when the winds are from 

the north and northeast. A similar exercise was conducted for 2013 HFR surface currents (Figure H42). 

The surface flow was found to be westward nearshore for nearly all wind directions, other than when 

winds were from the southwest, when the nearshore surface flow was generally eastward.

Our results indicate that only a fraction of the mean transport of 0.2 Sv flowing downcanyon in Barrow 

Canyon feeds the eastward shelfbreak current. Indeed the mean transport computed at BS3 over these 

years was ~0.025 Sv eastward or <15% of the mean canyon transport. (The transport along the Beaufort 

shelfbreak has apparently decreased substantially since 2003 due to an increase in easterly winds over the 

Beaufort Sea (Brugler et al. 2014)). This result is surprising insofar as geostrophic dynamics require that 

the canyon transport should exit the canyon and proceed eastward along the shelfbreak in its entirety. Our 

results imply that other dynamics are dominating, which prevent the canyon outflow from proceeding 

eastward. These dynamics must cause the canyon outflow to either turn westward or to enter the basin 

shortly after exiting the canyon, before the outflow reaches 152°W. We cannot address the fate of the 

canyon outflow with existing data, but Nof (1988, 1991) and Watanabe et al. (2011) showed that the fate 

of the outflow can be quite variable and is dependent upon the potential vorticity structures of the 

outflow, the receiving basin, and the continental slope. The PV structures within each region most 

certainly change on synoptic and seasonal time scales in response to the winds and in conjunction with 

the vertical and horizontal density structures in each region.
H37

Figure H39. Monthly distribution of flow regimes.

Figure H40. Annual cycle of the four flow regimes characterizing transports in Barrow Canyon and the 

Beaufort shelfbreak.
H38

Figure H41. Frequency of occurrence of regimes as a function of the direction toward which the wind at 

Barrow (PABR) is blowing.

The coherence calculations between Bering Strait and Barrow Canyon and between Barrow Canyon and 

the Beaufort shelfbreak suggest that Bering Strait and the Beaufort shelfbreak might also reflect coherent 

motions. Dynamically, we expect that wavelike motions excited in the Bering Sea or Bering Strait 

shouldpropagate along the shelfbreak as well. Consequently, we computed the coherence and phase 

spectra between Bering Strait and the Beaufort shelfbreak (Figure H43). These spectra indicate that there 

is some coherence between these two regions at periods less than 30 days, with fluctuations in Bering 

Strait leading those along the Beaufort shelfbreak by ~0.5 days. As with the other pairwise coherence 

tests the results vary seasonally, with coherence being better in the fall and winter months than in 

summer. There is also considerable interannual variability evident in the results. Understanding this 

variability will require additional analyses using models.

We conclude this section by briefly characterizing the tides over the Chukchi Sea shelf. We computed the 

tidal ellipse properties and phases of four semidiurnal (M2, S2, N2, and K2) and four diurnal (O1, K1, P1, 

and Q1) tidal constituents following Foreman (1978). The results are listed in Appendix B, which includes

the number of deployments used in the estimations. Figure H44 displays a map of the tidal ellipses. In 

some locations the constituents were unresolvable because either the record length or the signal to noise 

ratio was too small to yield reliable results. 
H39

Figure H42. Mean surface currents estimated from the Point Barrow and Cape Simpson HFR systesm 

under different wind conditions in 2013. The number of hours under which a given wind condition 

prevailed is indicated in the upper right hand corner of each map.
H40

Figure H43. Coherence squared and phase spectra of mean daily transports at Bering Strait and along 

the Beaufort shelfbreak for the entire record (left), May – September (middle), and October – April (right).

The M2 tide was the dominant tidal constituent over most of the shelf, although the associated currents 

were small (≤4 cm s-1). M2 current magnitudes were largest west and south of Hanna Shoal and 

diminished to ~1 cm s-1 or less in Barrow Canyon. The S2 currents were ≤2 cm s-1; however, this

constituent was largest on the shelf northeast of Hanna Shoal and, in fact, was larger than the M2 currents 

in this location. The reason for this is unclear but it may be related to resonant interaction of the subinertial S2 wave with the sloping bottom in the vicinity of the shelfbreak as found, for example, by

Pnyushkov and Polyakov (2012) in the Laptev Sea. The O1 tide is the dominant diurnal constituent, but 

again, its magnitudes were small (≤1 cm s-1) and largest on the shelf northeast of Hanna Shoal.

H.6 Summary

This program utilized a wide variety of observations obtained from moored oceanographic instruments, 

satellite-tracked drifters, hydrography (both recent from autonomous underwater vehicles and towed- and 

shipboard CTD, as well as historical shipboard measurements), high-frequency radars, meteorological 

buoys, and satellite imagery to develop an improved understanding of the physical oceanography of the 

northeastern Chukchi shelf, as well as exchanges between the Chukchi/Beaufort shelves and the adjacent 

basin. Our measurements spanned the shelf region from Point Hope northward to the Chukchi and

Beaufort continental slopes and from the Alaskan coast to the US-Russian Convention Line. Our primary 

focus area was the shelf and shelfbreak region between the Central Channel and the western Beaufort Sea

and north to the shelfbreak. This is an enormous data set that will continue to be analyzed in the future. In 

this section we summarize our conclusions using Figure H45, which is based largely on the moorings 

discussed herein and from data published by Itoh et al. (2014) for the mouth of Barrow Canyon, as a 

guide.

The figure shows two surface fronts: the MW/BSSW front (denoted by the blue line), which surrounds 

most of Hanna Shoal, and the ACW front (red line) that parallels the Alaskan coast from south of Bering 

Strait to over the Beaufort Sea shelfbreak. This front is primarily a summer and fall phenomenon, as it is 

much weaker or even absent in winter and early spring. We are less certain about the seasonality of the

MW/BSSW front, although we believe that it persists to some degree along the southeast of Hanna Shoal 

year-round except in those years (such as 2011), when MW is absent from the Hanna Shoal region.
H41

Figure H44. Tidal ellipses for the principal semi-diurnal (upper four panels) and diurnal tides on the 

Chukchi Sea shelf.
H42

Figure H45. Schematic of the mean circulation of the northeastern Chukchi Sea. The red line denotes the 

Alaskan Coastal Water (ACW) front bordering the offshore side of the Alaskan Coastal Current. The blue 

line indicates the meltwater/Bering Sea Summer Water (MW/BSSW) front surrounding Hanna Shoal. 

Black arrows indicate regions where the surface and sub-surface currents are similar in direction and 

magnitude. Green arrows indicate near-bottom currents, and purple arrows indicate near-surface 

currents.

Throughout the year and over most of the shelf south of 71.5°N, including within the Central Channel and 

along the northwest side of Hanna Shoal (as far north as the 50 m isobath), the flow is weakly sheared

vertically, with surface and sub-surface flows being within ~30 degrees of one another and all having 

similar magnitudes. Exceptions to this may involve large shears between drifting ice and the sub-surface 

flow and/or large shears during the melt season. During melt, the near-surface stratification is very strong 

(similar to the stratification found on the shelf east of Hanna Shoal), such that the sub-surface and surface 

circulations may be uncoupled from one another. This is the situation encountered on the East Hanna 

Shoal Shelf (EHSS) where vertical shears are large year-round. Other regions of strong vertical shear 

include the Beaufort and Chukchi shelfbreaks and slopes and portions of Barrow Canyon in summer and 

fall.

We have examined the strength, persistence, and variability of the model-predicted clockwise circulation 

around Hanna Shoal, including southwestward flow over the shelf between the Barrow Canyon and 

Hanna Shoal. Substantial elements of the models are largely correct, but there are important differences 
H43

between the models and observations. We found that the component of the clockwise circulation on the 

northwest side of Hanna Shoal extends through the water column as predicted by the models and includes 

cold Bering Sea Water (BSW) and Winter Water (WW) along the bottom. Both of these water masses are 

transported from the Central Channel and Herald Canyon. However, on the EHSS, only the bottom 

waters, consisting of BSW and WW, engage in the clockwise circulation. The upper layer flows westward 

and/or northwestward over the EHSS. This flow is a consequence of horizontal density gradients that 

extend northeastward and eastward from Hanna Shoal, which drive a counterclockwise baroclinic, 

geostrophic circulation in the upper layer that transports MW and/or modified ACW. The latter is 

transported onto the EHSS via Ekman transport from Barrow Canyon and/or by onshelf transport from 

the Chukchi shelfbreak. The different flow regimes on either side of Hanna Shoal have two effects. First, 

these result in a zonal exchange of water masses that maintains strong stratification year-round on the 

EHSS. Second, we infer that the shelf north of Hanna Shoal is a region of zonal flow convergence,

wherein waters must be transported toward the shelfbreak. This transport, which certainly involves the 

surface layers and possibly the bottom waters, should effect the magnitude and structure of the flow over 

the Chukchi shelfbreak.

Models also suggest that the clockwise flow around Hanna Shoal extends south of Hanna Shoal in the 

form of a southwestward drift that eventually turns eastward across the central shelf and continues toward 

the head of Barrow Canyon. The observations indicate that this is not the case. Indeed, we found that 

there was no net transport between the EHSS and the central shelf. Instead there is a southeastward flow 

along the south side of Hanna Shoal that transports water (usually WW) toward the head of Barrow 

Canyon. This flow essentially represents a dynamic boundary that blocks communication between the 

EHSS and the central shelf. As a consequence of this boundary and the weak mean flow on the northern 

side of the Shoal, the residence time of bottom waters on the shelf east of the Shoal is ~9 months, much 

longer than elsewhere in the Chukchi Sea.

As noted, the inferred zonal convergence north of the Shoal provides a possible mechanism for 

shelf/slope exchange in which MW, WW, and BSW are transported offshelf in this area. There are 

however other mechanisms that foster exchange. Our retrospective analysis of hydrographic data from the 

Chukchi shelfbreak indicates that this consists of a surface flow that carries Chukchi Sea water masses 

(the Pacific-derived summer waters, WW, and MW) westward and an eastward sub-surface flow that is 

confined to the shelfbreak and which transports WW eastward. The waters flowing westward are largely 

derived from the Barrow Canyon outflow. The WW that comprises the eastward subsurface flow is 

mainly derived from northward flows through the Central Channel and Herald Canyon, although some 

probably forms in winter around the periphery of Hanna Shoal. The potential vorticity structure of the 

shelfbreak circulation satisfies the conditions for baroclinic instability, which should lead to eddy-driven 

exchanges between the shelfbreak and basin as is known to occur frequently along the Beaufort Sea 

shelfbreak.

Our results suggest that most of the Chukchi shelf water masses that enter the head of Barrow Canyon are 

drawn from the shelf south of Hanna Shoal. Of the 0.2 Sv average transport we estimate to be flowing 

into the head of the canyon, at least ~0.1 Sv flows eastward across the central shelf from the Central 

Channel. The remainder presumably flows northeastward along the coast. Although we are unable to 

estimate the transports involved, we infer that additional waters enter the canyon along its length 

(downcanyon from the head) due to spillage of bottom waters from the EHSS or from westward-flowing 

waters from the Beaufort Sea. There also appears to be, on average, an upcanyon flow along the west side 

of the canyon. This flow probably transports waters from the subsurface shelfbreak current flowing 

eastward along the Chukchi slope. Most likely this inflow proceeds only part way up the canyon because 

it continuously loses mass to the downcanyon transport along the east side of the canyon. The WW 

comprise the bulk of the water masses spilling into the canyon and/or transported upcanyon from the 

shelfbreak. 
H44

The circulation and transports in the Central Channel, across the central shelf, and in Barrow Canyon vary 

annually and are (mostly) in-phase with the annual transport cycle in Bering Strait. For the 2010 – 2015 

period, we found that the downcanyon transport is large (0.4 – 0.6 Sv) from May through August, weak 

(≤0.2 Sv) from September through November and from February through April, and reversed (e.g., 

upcanyon at ~0.1 Sv) in December and January. The annual cycle over the central shelf is in-phase with 

that in Barrow Canyon and includes a reversed (westward) flow in December and January. In contrast, the 

flow in the Central Channel is northward in all months. These results suggest that in early winter the 

circulation over the northeastern Chukchi Sea shelf undergoes a major reorganization: waters transported 

upcanyon may extend onto the central shelf, and nearshore waters are transported westward into the 

Central Channel and even the western Chukchi Sea. Once in the Central Channel, these nearshore waters 

would be transported northward. This reorganization provides a means by which WW, formed in the 

latent heat polynyas along the northwest coast of Alaska in December and January, can be carried far 

offshore and eventually to Hanna Shoal.

This seasonal reorganization of the flow was likely reflected during synoptic coastal upwelling events 

such as the one captured by satellite-tracked drifters August 2014. During that event, the cross-shelf 

transport of surface coastal waters was a response to various competing mechanisms including the 

geostrophic transport associated with the mean Pacific-Arctic pressure gradient and in response to the 

winds, surface Ekman drift, shelf-wide pressure gradient adjustments, and Stokes’ drift, which although 

small is not negligible. The drifter observations showed that surface nearshore waters can be transported 

from the coast to either the Central Channel and/or the head of Herald Canyon in less than a month. 

Similarly, self-organized mapping analysis applied to the HFR observations clearly indicate that waters 

transported up Barrow Canyon can easily be carried westward onto the central shelf and/or southward 

along the Alaskan coast. In fall, the coastal upwelling events and Barrow Canyon reversals likely lead to 

cooling of nearshore waters and, thus, enhance ice formation. 

The shelf circulation also includes a complex variety of mesoscale circulations. One example (observed 

about 10% of the time) is the “divergent” mode, which occurs during transitions from northeasterly to 

southeasterly winds. In such events, a pair of counter-rotating vortices form in which the eastward 

onshore flow from the central shelf splits upon approaching the coast; some water enters a cyclonic vortex 

to the north of Wainwright and some enters an anticyclonic vortex offshore of Point Lay. The water 

entering the anticyclonic vortex recirculates over the central shelf. Two other examples are the inferred 

shedding or generation of eddies as the Alaskan Coastal Current rounds Point Hope and the collision of 

the westward-flowing Beaufort shelf waters with the along-canyon flow in Barrow Canyon. There are 

also numerous MW fronts that occur on this shelf, which we suggest are of two types (at least). The first 

type consists of what appears to be numerous small sub-mesoscale (≤6 km) fronts that are shallow (10 –

15 m thick) and found over Hanna Shoal and the EHSS. There is another, more prominent, type of 

mesoscale front, which in August and September extends from the western end of Hanna Shoal eastward 

to the head of Barrow Canyon along the south side of the Shoal and then northward along the western half 

of Barrow Canyon. These fronts separate moderately-stratified BSW to the south (and east in Barrow 

Canyon) from the strongly-stratified, two-layered MW/WW structure to the north. The fronts are ~15 km 

wide, baroclinically unstable, and generate both anticyclonic and cyclonic eddies at the surface and 

anticyclonic eddies primarily within the pycnocline north of the front.

These mesoscale and sub-mesoscale motions play an important role in the lateral and vertical mixing of 

water masses and in the dispersion of dissolved and suspended materials. To a large degree, the mesoscale 

phenomena that we have identified on the shelf are linked to the distribution of MW. Over the 

northeastern shelf, the summer MW pool is tied to ice conditions over Hanna Shoal, which can vary 

considerably from year-to-year. We suggest that the rate of ice retreat over the Shoal in summer may 

depend upon a number of variables. In summer, early ice retreat seems to be favored when winds are 

strong and persistent from the northeast. However, the degree to which ice retreats over the Shoal depends 
H45

upon the ice thickness distribution, in particular, the presence or absence of thick, grounded ice. We 

suggest that the summer thickness distribution on the Shoal is likely a consequence of ice advection onto 

the Shoal during the previous winter. If the advected ice is thick enough to ground on the Shoal in winter 

then it may encourage additional deformation and accretion of ice throughout the winter and early spring.
I1

I. Recommendations for Future Study

1. There is a strong eastward flow along the northwest side of Hanna Shoal, which likely extends 

farther to the northwest than the 60 m isobath. This flow includes waters emanating from both the 

Central Channel and Herald Canyon farther to the west. Information is needed on the circulation 

between Herald Canyon and Hanna Shoal as this transport has an important influence on the 

northeastern Chukchi Sea. This can be determined by deploying an array of moorings from south 

to north slightly east of the Russian-US Convention Line. We recommend that a meridional array 

of moorings be deployed between the 50 and 100 m isobaths to assess this transport.

2. An effort should be made to quantify the transport in the Central Channel. We believe that this 

could be done be deploying 2 – 3 moorings at the southern end of the Channel and 2 – 3 at the 

northern end to the west of Hanna Shoal. The moorings in recommendation 1 and 2 should 

include T/C/P recorders at multiple depths in order to assess seasonality in stratification.

3. We hypothesized that flow convergence occurs over the shelf to the north of Hanna Shoal. This 

hypothesis should be explored by conducting a zonal sequence of cross-isobath synoptic sections 

to examine water properties and the circulation here. This survey should include VMADCP data 

and consideration should be given to deploying shallow- and deep-drogued (i.e., 10 and 30 m) 

satellite tracked drifters to assess the surface and deep circulation on the north side of the Shoal.

4. Two surprising and related observations were the year-round presence of heavy stratification and 

the inferred year-round counterclockwise baroclinic pressure tendency on the shelf east of Hanna 

Shoal. The baroclinic pressure field is largest in the upper half of the water column where it 

opposes the barotropic pressure field. These competing pressure fields are critical in establishing 

the inferred convergence north of the Shoal and the weak circulation and water properties east of 

the Shoal. Verification of this baroclinicity and its seasonality can be accomplished with a few 

moorings using ISCATs (used successfully in COMIDA) and/or a string of T/C/Ps in conjunction 

with bottom pressure sensors. This deployment could be done in conjunction with the Chukchi 

Environmental Observatory (Danielson pers comm.).

5. The persistence of sea ice (or lack thereof) in summer on Hanna Shoal has a profound influence 

on the regional circulation and hydrography. In addition, it also affects the regional biology 

through control of stratification and as habitat for walruses and other marine mammals. It may 

also play a key role in governing sources and sinks of carbon and the structure of the benthos. A 

prudent area of investigation would be to determine the sources of ice on Hanna Shoal, the 

mechanisms controlling these sources, and those that affect its persistence or its ablation.

6. A systematic effort is required to determine the relative roles of winds, air-sea heat exchange and 

ocean heat flux convergences, on the seasonal retreat and advance of sea ice on the Chukchi Sea 

shelf. Such a study will require careful experimental work that addresses horizontal advection and 

vertical and horizontal mixing processes. 

7. It is becoming increasingly clear that not all flow field fluctuations are forced locally over the 

Chukchi shelf and that some but not all northeastern Chukchi shelf variations are subsequently 

transmitted to the Beaufort shelf. We recommend studies that will reveal the role of remote 

processes impacting advection of ice and water locally. Such fluctuations may be important for 

certain landfast ice breakout events and synoptic-scale flow field variations. Ocean-atmosphere 

couplings and oceanic transmissions may vary seasonally so investigations would need to assess 
I2

the roles of landfast ice, river plumes, and stratification, and their interactions with the seafloor 

topography. 

8. The accuracy of satellite-derived ice motions using the ADCP bottom-track data should be 

assessed. Comparisons from other sites around the Arctic suggest that derived ice motion 

products perform reasonably well at some locations and some times of year but the performance 

of these products have not been well evaluated in the study region.

The outcomes of each of these recommended studies are relevant to an understanding of ecosystem 

processes as well as providing an understanding of, and an ability to predict, the fate of pollutants.
J1

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Appendix A - 1

Appendix A: Surface Current Patterns in the Northeastern Chukchi 

Sea and Their Response to Wind Forcing

A.1 Introduction

SOM derived surface current and wind patterns from 2012 were used to define circulation features in the 

northeastern Chukchi Sea. Generating 12 patterns allowed us to isolate data gaps and unique circulation 

features; however, it also derived many patterns that were similar to one another. To simplify presented 

results in the main manuscript, similar patterns were grouped together (Table A-1). We provide all 12 

original patterns here for those interested in more detail. Figure A-1 corresponds to the simplified Figures 

C2a-d, while Figure A-2 corresponds to Figure C4.

Please note that the number of the pattern is not a ranking, but rather an arbitrarily assigned identification 

number. This should be evident from the percentage of occurrence recorded in each figure. 

A.2 SOM Grouping Explained

Figures A-1a and A-1b were grouped together as the reversal regime (Figure C2b) as these patterns both 

depict overall flow in the domain to the southwest/west under northeasterly winds. Figures A-1c and A-1f 

are classified as the northwesterly wind regime (Figure C2c). Figure A-1d stands alone as the divergent 

mode (Figure C2d). Figure A-1e represents Pattern G, indicative of times when data returns were sparse. 

Figures A-1g through A-1l were grouped together as the northeastward-flowing regime (Figure C2a). All 

of these patterns contained predominantly northeastward-flowing currents along the coastline. The main 

differences result from winds and data gap distributions. Patterns 7 (Figure A-1g) and 10 (Figure A-1j) 

depict discontinuities between the northern and southern masks, with sparse, variable currents in the

Table A-1. A summary of the original SOM patterns and how they are grouped as circulation regimes in 

the main manuscript.

SOM Pattern Number Regime Grouping

1 Reversal

2 Reversal

3 Northwesterly Wind

4 Divergent Mode

5 Data Gap

6 Northwesterly Wind

7 Northeastward-Flowing

8 Northeastward-Flowing

9 Northeastward-Flowing

10 Northeastward-Flowing

11 Northeastward-Flowing

12 Northeastward-Flowing
Appendix A - 2

Figure A-1. (a)-(h), SOM-derived patterns No. 1–8 of surface currents (blue vectors) and paired winds 

(black vectors) for 2012 data. Pattern number and its relative frequency of occurrence are shown in the 

lower right corner. The 80 m isobath is thicker to define Barrow Canyon. Note that the scales of current 

vectors in Patterns No. 3–8 are different.
Appendix A - 3

Figure A-1. (continued) (i)-(l), SOM-derived patterns No. 9–12 of surface currents (blue vectors) and 

paired winds (black vectors) for 2012 data.

southern region near Point Lay and Icy Cape. A review of the data confirmed that the Point Lay HFR was 

experiencing maintenance issues in August 2012, when patterns 7 and 10 occurred most often. Given that 

the northern region of currents is derived from the Wainwright and Barrow HFR, which were operational, 

it follows that the currents in the northern region are accurate while the currents in the southern region are 

not. Therefore it is still appropriate to categorize patterns 7 and 10 as the northeastward-flowing regime, 

although they depict currents when only the northern mask is operational.
Appendix A - 4

Figure A-2. (a) Upper panel: time series of wind vectors in August 2012. The vector direction follows 

oceanographic convention. Lower panel: SOM-derived patterns (black dots) in August 2012 and 

normalized data returns (gray line). Pattern numbers correspond to Figure A-1. Pattern 0 denotes times 

when data return is zero. (b) As in (a), but for September 2012. (c) As in (a), but for October 2012.
Appendix B - 1

Appendix B: Tidal Ellipse Properties for the Chukchi Sea Shelf

Table B-1. M2 Tidal Ellipse Parameters

Mooring 

Id

Major 

Axis 

(cm/s)

Minor 

Axis 

(cm/s)

Rotation 

Direction

Inclination 

(°)

Greenwich 

Phase (°)

# of 

years

BC1 0.8 ± 0.2 0.1 ± 0.2 CCW 30 ± 13 -155 ± 15 0.19 2

BC2 0.9 ± 0.2 0.1 ± 0.2 22 ± 15 -152 ± 18 0.15 5

BC3 0.8 ± 0.1 0.1 ± 0.1 CW 1 ± 11 -87 ± 12 0.14 2

BC4 0.9 ± 0.1 0.3 ± 0.1 CW 2 ± 10 176 ± 9 0.28 2

BC5 1.0 ± 0.2 0.4 ± 0.1 CW 2 ± 11 73 ± 12 0.38 2

BC6 1.4 ± 0.3 0.6 ± 0.3 CW 11 ± 16 165 ± 13 0.42 2

Bu 2.1 ± 0.1 1.5 ± 0.1 CW 23 ± 9 136 ± 8 0.71 6

CPAI01 2.8 ± 0.1 2.1 ± 0.2 CW 35 ± 9 125 ± 9 0.76 5

CPAI02 0.9 ± 0.1 0.6 ± 0.1 CW 163 ± 18 -15 ± 16 0.59 2

Cj 3.0 ± 0.1 2.3 ± 0.1 CW 50 ± 9 111 ± 8 0.78 6

EBC 0.5 ± 0.2 0.0 ± 0.2 CW 51 ± 31 -28 ± 29 0.04 3

HS01 2.4 ± 0.1 2.2 ± 0.1 CW 59 ± 31 103 ± 32 0.92 1

HS02 2.7 ± 0.2 2.4 ± 0.2 CW 43 ± 30 104 ± 30 0.92 1

HS03 2.7 ± 0.1 2.4 ± 0.1 CW 59 ± 25 109 ± 25 0.9 1

HS04 2.5 ± 0.1 2.2 ± 0.1 CW 14 ± 28 138 ± 27 0.91 1

HS05 1.9 ± 0.1 1.5 ± 0.1 CW 9 ± 12 147 ± 11 0.77 1

HS06 1.4 ± 0.1 0.8 ± 0.1 CW 4 ± 9 155 ± 9 0.58 1

HSNE40 1.6 ± 0.2 1.4 ± 0.2 CW -8 ± 55 72 ± 56 0.91 2

HSNE50 0.7 ± 0.2 0.5 ± 0.2 CW 139 ± 57 0 ± 59 0.78 2

HSNE60 0.5 ± 0.2 0.4 ± 0.2 CW 130 ± 60 -15 ± 57 0.7 2

HSNW40 3.1 ± 0.2 3.0 ± 0.2 CW 32 ± 43 129 ± 44 0.95 2

HSNW50 2.8 ± 0.2 2.6 ± 0.2 CW 47 ± 36 115 ± 35 0.91 2

Site2 1.1 ± 0.2 0.3 ± 0.3 CW 10 ± 17 170 ± 15 0.28 4

Statoil3 2.7 ± 0.1 2.3 ± 0.1 CW 24 ± 18 138 ± 18 0.88 1

Statoil4 3.3 ± 0.1 3.0 ± 0.1 CW 34 ± 19 125 ± 20 0.91 1

WBC 0.4 ± 0.2 0.1 ± 0.2 CW 76 ± 38 -38 ± 30 0.26 3

Minor

Major
Appendix B - 2

Table B-2. S2 Tidal Ellipse Parameters

Mooring 

Id

Major 

Axis 

(cm/s)

Minor 

Axis 

(cm/s)

Rotation 

Direction

Inclination 

(°)

Greenwich 

Phase (°)

# of 

Months

BC1 0.3 ± 0.2 0.1 ± 0.2 64 ± 46 70 ± 39 0.26 2

BC2 0.3 ± 0.2 0.1 ± 0.2 CW 83 ± 52 83 ± 47 0.27 5

BC3 0.3 ± 0.1 0.2 ± 0.1 CW 53 ± 59 104 ± 52 0.58 2

BC4 0.4 ± 0.1 0.1 ± 0.1 CW 89 ± 24 74 ± 25 0.33 2

BC5 0.4 ± 0.1 0.3 ± 0.1 CW 28 ± 90 138 ± 90 0.87 2

BC6 0.7 ± 0.3 0.3 ± 0.3 CW 118 ± 38 60 ± 35 0.49 2

Bu 0.8 ± 0.1 0.6 ± 0.1 CW 134 ± 40 41 ± 41 0.82 6

CPAI01 1.1 ± 0.1 0.9 ± 0.1 CW 10 ± 31 51 ± 31 0.82 5

CPAI02 0.7 ± 0.1 0.7 ± 0.1 CW 52 ± 74 97 ± 75 0.91 2

Cj 1.2 ± 0.1 1.0 ± 0.1 CW 29 ± 41 169 ± 40 0.87 6

EBC 0.5 ± 0.2 0.1 ± 0.3 CW 51 ± 33 104 ± 33 0.16 3

HS01 1.1 ± 0.1 0.9 ± 0.1 CW 75 ± 43 103 ± 44 0.88 1

HS02 1.4 ± 0.1 1.2 ± 0.2 CW 70 ± 58 92 ± 57 0.91 1

HS03 1.4 ± 0.1 1.2 ± 0.1 CW 79 ± 31 102 ± 31 0.85 1

HS04 1.1 ± 0.1 1.0 ± 0.1 CW 123 ± 49 49 ± 51 0.91 1

HS05 0.9 ± 0.1 0.8 ± 0.1 CW 154 ± 44 12 ± 42 0.86 1

HS06 0.6 ± 0.1 0.6 ± 0.1 CW 116 ± 98 47 ± 104 0.95 1

HSNE40 1.4 ± 0.2 1.2 ± 0.2 CW 58 ± 55 84 ± 56 0.88 2

HSNE50 1.3 ± 0.2 1.1 ± 0.2 CW 58 ± 59 80 ± 59 0.87 2

HSNE60 1.1 ± 0.2 1.1 ± 0.2 CW 52 ± 87 77 ± 85 0.93 2

HSNW40 1.4 ± 0.2 1.3 ± 0.2 CW 87 ± 63 83 ± 64 0.92 2

HSNW50 1.3 ± 0.2 1.2 ± 0.2 CW 67 ± 69 114 ± 68 0.93 2

Site2 0.4 ± 0.2 0.2 ± 0.2 CW 107 ± 53 35 ± 49 0.57 4

Statoil3 1.1 ± 0.1 1.0 ± 0.1 CW 145 ± 74 32 ± 74 0.94 1

Statoil4 1.4 ± 0.1 1.3 ± 0.1 CW 120 ± 109 60 ± 111 0.99 1

WBC 0.6 ± 0.2 0.3 ± 0.2 CW 44 ± 36 90 ± 37 0.6 3

Minor

Major
Appendix B - 3

Table B-3. N2 Tidal Ellipse Parameters

Mooring 

Id

Major 

Axis 

(cm/s)

Minor 

Axis 

(cm/s)

Rotation 

Direction

Inclination 

(°)

Greenwich 

Phase (°)

# of 

Months

BC1 0.4 ± 0.2 0.0 ± 0.2 25 ± 23 97 ± 32 0.04 2

BC2 0.4 ± 0.2 0.1 ± 0.2 30 ± 40 140 ± 41 0.27 5

BC3 0.3 ± 0.1 0.1 ± 0.1 42 ± 36 99 ± 39 0.21 2

BC4 0.3 ± 0.1 0.1 ± 0.1 CW 6 ± 25 17 ± 26 0.15 2

BC5 0.3 ± 0.2 0.1 ± 0.2 CW 33 ± 42 104 ± 47 0.26 2

BC6 0.5 ± 0.2 0.1 ± 0.3 CW 13 ± 36 86 ± 35 0.15 2

Bu 0.6 ± 0.1 0.4 ± 0.1 CW 48 ± 31 75 ± 32 0.71 6

CPAI01 0.6 ± 0.1 0.4 ± 0.1 CW 56 ± 43 64 ± 46 0.75 5

CPAI02 0.5 ± 0.1 0.3 ± 0.1 CW 19 ± 39 111 ± 37 0.65 2

Cj 0.6 ± 0.1 0.5 ± 0.1 CW 66 ± 61 55 ± 61 0.83 6

EBC 0.3 ± 0.2 0.1 ± 0.2 64 ± 78 -176 ± 80 0.39 3

HS01 0.7 ± 0.1 0.6 ± 0.2 CW 65 ± 45 57 ± 42 0.79 1

HS02 0.9 ± 0.1 0.9 ± 0.1 CW 164 ± 124 319 ± 129 0.95 1

HS03 0.8 ± 0.1 0.7 ± 0.1 CW 59 ± 52 85 ± 53 0.86 1

HS04 0.8 ± 0.1 0.6 ± 0.1 CW 43 ± 38 72 ± 38 0.84 1

HS05 0.7 ± 0.1 0.5 ± 0.1 CW 27 ± 40 98 ± 36 0.8 1

HS06 0.4 ± 0.1 0.3 ± 0.1 CW 15 ± 47 104 ± 43 0.7 1

HSNE40 0.9 ± 0.2 0.8 ± 0.2 CW 26 ± 43 114 ± 45 0.81 2

HSNE50 0.7 ± 0.2 0.6 ± 0.2 CW 38 ± 95 120 ± 89 0.91 2

HSNE60 0.5 ± 0.2 0.5 ± 0.2 CW 15 ± 91 145 ± 94 0.89 2

HSNW40 1.0 ± 0.2 1.0 ± 0.1 CW 50 ± 90 80 ± 89 0.93 2

HSNW50 0.8 ± 0.2 0.7 ± 0.2 CW 56 ± 86 51 ± 84 0.91 2

Site2 0.5 ± 0.2 0.0 ± 0.3 CW 35 ± 32 98 ± 31 0.11 4

Statoil3 0.7 ± 0.1 0.7 ± 0.1 CW 93 ± 72 46 ± 72 0.93 1

Statoil4 1.0 ± 0.1 0.8 ± 0.1 CW 40 ± 24 88 ± 24 0.75 1

WBC 0.3 ± 0.1 0.2 ± 0.2 CW 53 ± 75 140 ± 76 0.66 3

Minor

Major
Appendix B - 4

Table B-4. K2 Tidal Ellipse Parameters

Mooring 

Id

Major 

Axis 

(cm/s)

Minor 

Axis 

(cm/s)

Rotation 

Direction

Inclination 

(°)

Greenwich 

Phase (°)

# of 

Months

BC1 0.1 ± 0.2 0.0 ± 0.1 CW 41 ± 81 162 ± 99 0.11 2

BC2 0.3 ± 0.2 0.0 ± 0.2 CW 16 ± 60 171 ± 66 0.09 5

BC3 0.2 ± 0.1 0.1 ± 0.1 CW 9 ± 55 -140 ± 49 0.5 2

BC4 0.3 ± 0.1 0.1 ± 0.1 CW 109 ± 38 101 ± 38 0.36 2

BC5 0.2 ± 0.2 0.1 ± 0.1 CW -5 ± 56 125 ± 61 0.49 2

BC6 0.1 ± 0.1 0.0 ± 0.1 CW 40 ± 102 210 ± 112 0.2 1

Bu 0.3 ± 0.1 0.2 ± 0.1 CW 63 ± 66 128 ± 66 0.74 6

CPAI01 0.3 ± 0.1 0.3 ± 0.1 CW 34 ± 77 165 ± 78 0.81 5

CPAI02 0.2 ± 0.1 0.1 ± 0.1 CW 143 ± 72 30 ± 99 0.31 2

Cj 0.4 ± 0.1 0.3 ± 0.1 CW 55 ± 64 148 ± 65 0.76 6

EBC 0.2 ± 0.2 0.1 ± 0.2 77 ± 77 124 ± 81 0.33 3

HS01 0.4 ± 0.1 0.4 ± 0.1 CW 29 ± 98 164 ± 101 0.92 1

HS02 0.4 ± 0.2 0.3 ± 0.2 CW 144 ± 110 35 ± 98 0.92 1

HS03 0.4 ± 0.1 0.3 ± 0.1 CW 50 ± 93 130 ± 90 0.83 1

HS04 0.4 ± 0.1 0.3 ± 0.1 CW 132 ± 68 54 ± 70 0.79 1

HS05 0.2 ± 0.1 0.1 ± 0.1 CW 170 ± 52 19 ± 50 0.48 1

HS06 0.3 ± 0.1 0.2 ± 0.1 CW 18 ± 61 155 ± 54 0.63 1

HSNE40 0.3 ± 0.2 0.3 ± 0.2 CW -18 ± 103 132 ± 108 0.85 2

HSNE50 0.2 ± 0.2 0.1 ± 0.2 CW 138 ± 105 59 ± 113 0.48 2

HSNE60 0.2 ± 0.2 0.1 ± 0.2 CW -8 ± 101 99 ± 106 0.44 2

HSNW40 0.6 ± 0.2 0.5 ± 0.2 CW 126 ± 112 78 ± 117 0.95 2

HSNW50 0.5 ± 0.2 0.4 ± 0.2 CW 106 ± 80 123 ± 77 0.79 2

Site2 0.1 ± 0.1 0.1 ± 0.1 73 ± 88 128 ± 84 0.49 2

Statoil3 0.4 ± 0.1 0.4 ± 0.1 CW 170 ± 92 21 ± 92 0.93 1

Statoil4 0.5 ± 0.1 0.4 ± 0.1 CW 36 ± 97 163 ± 93 0.93 1

WBC 0.1 ± 0.1 0.0 ± 0.1 38 ± 112 -50 ± 137 0.22 2

Minor

Major
Appendix B - 5

Table B-5. O1 Tidal Ellipse Parameters

Mooring 

Id

Major 

Axis 

(cm/s)

Minor 

Axis 

(cm/s)

Rotation 

Direction

Inclination 

(°)

Greenwich 

Phase (°)

# of 

Months

BC1 0.9 ± 0.4 0.2 ± 0.3 CCW 38 ± 32 49 ± 31 0.24 2

BC2 0.8 ± 0.5 0.1 ± 0.4 29 ± 32 42 ± 40 0.12 5

BC3 0.8 ± 0.3 0.1 ± 0.2 CW 15 ± 19 54 ± 28 0.18 2

BC4 0.8 ± 0.3 0.2 ± 0.2 CW 10 ± 20 55 ± 22 0.27 2

BC5 0.8 ± 0.3 0.1 ± 0.2 CW 171 ± 16 -131 ± 24 0.16 2

BC6 1.1 ± 0.4 0.3 ± 0.5 CW 160 ± 33 -136 ± 27 0.33 2

Bu 0.6 ± 0.2 0.3 ± 0.2 CW -17 ± 27 -144 ± 30 0.49 6

CPAI01 0.6 ± 0.2 0.3 ± 0.2 CW 2 ± 33 -32 ± 31 0.48 5

CPAI02 1.2 ± 0.2 0.5 ± 0.2 CW 158 ± 14 -138 ± 16 0.42 2

Cj 0.5 ± 0.1 0.3 ± 0.1 CW 12 ± 38 -51 ± 39 0.57 6

EBC 1.4 ± 0.6 0.1 ± 0.6 41 ± 27 69 ± 30 0.1 3

HS01 1.1 ± 0.2 0.6 ± 0.2 CW 179 ± 18 200 ± 17 0.56 1

HS02 1.2 ± 0.2 0.6 ± 0.2 CW 171 ± 14 181 ± 13 0.53 1

HS03 1.3 ± 0.2 0.8 ± 0.1 CW 7 ± 14 5 ± 17 0.63 1

HS04 0.8 ± 0.1 0.6 ± 0.1 CW 163 ± 24 195 ± 22 0.68 1

HS05 0.9 ± 0.2 0.5 ± 0.1 CW 161 ± 15 211 ± 17 0.55 1

HS06 0.8 ± 0.3 0.1 ± 0.2 CW 177 ± 15 218 ± 18 0.12 1

HSNE40 1.6 ± 0.3 0.9 ± 0.2 CW 162 ± 12 -172 ± 14 0.56 2

HSNE50 1.5 ± 0.2 0.9 ± 0.2 CW 151 ± 17 -169 ± 19 0.61 2

HSNE60 1.7 ± 0.3 1.0 ± 0.3 CW 141 ± 17 -171 ± 17 0.59 2

HSNW40 1.2 ± 0.2 0.8 ± 0.2 CW -7 ± 14 -85 ± 15 0.62 2

HSNW50 1.1 ± 0.3 0.6 ± 0.2 CW 168 ± 17 -167 ± 19 0.51 2

Site2 0.7 ± 0.4 0.1 ± 0.4 -9 ± 33 -144 ± 36 0.13 4

Statoil3 0.9 ± 0.2 0.6 ± 0.2 CW 1 ± 27 3 ± 29 0.64 1

Statoil4 0.9 ± 0.2 0.6 ± 0.2 CW 1 ± 32 351 ± 32 0.7 1

WBC 1.3 ± 0.7 0.2 ± 0.5 CW -5 ± 22 37 ± 29 0.15 3

Minor

Major
Appendix B - 6

Table B-6. P1 Tidal Ellipse Parameters

Mooring 

Id

Major 

Axis 

(cm/s)

Minor 

Axis 

(cm/s)

Rotation 

Direction

Inclination 

(°)

Greenwich 

Phase (°)

# of 

Months

BC1 0.3 ± 0.3 0.0 ± 0.3 CW 53 ± 69 107 ± 75 0.15 2

BC2 0.4 ± 0.3 0.1 ± 0.3 45 ± 83 115 ± 82 0.29 5

BC3 0.2 ± 0.2 0.0 ± 0.2 CW 81 ± 92 74 ± 111 0.3 2

BC4 0.2 ± 0.2 0.1 ± 0.2 38 ± 96 74 ± 107 0.46 2

BC5 0.2 ± 0.2 0.1 ± 0.2 0 ± 64 -27 ± 100 0.28 2

BC6 0.1 ± 0.1 0.1 ± 0.2 CW 134 ± 119 270 ± 120 0.87 1

Bu 0.2 ± 0.2 0.0 ± 0.1 3 ± 59 -167 ± 72 0.19 6

CPAI01 0.2 ± 0.2 0.0 ± 0.2 132 ± 70 -168 ± 75 0.2 5

CPAI02 0.3 ± 0.2 0.1 ± 0.2 CW 166 ± 48 -132 ± 68 0.27 2

Cj 0.1 ± 0.1 0.0 ± 0.1 160 ± 79 -137 ± 91 0.17 6

EBC 0.5 ± 0.5 0.3 ± 0.4 55 ± 89 145 ± 88 0.54 3

HS01 0.2 ± 0.1 0.2 ± 0.1 CW 165 ± 82 219 ± 81 0.75 1

HS02 0.3 ± 0.1 0.2 ± 0.1 CW 91 ± 70 230 ± 61 0.62 1

HS03 0.3 ± 0.2 0.1 ± 0.1 CW 1 ± 29 18 ± 39 0.43 1

HS04 0.2 ± 0.1 0.1 ± 0.1 CW 102 ± 49 231 ± 49 0.38 1

HS05 0.2 ± 0.1 0.0 ± 0.1 135 ± 51 224 ± 51 0 1

HS06 0.3 ± 0.2 0.0 ± 0.2 CW 116 ± 51 245 ± 50 0.15 1

HSNE40 0.3 ± 0.2 0.2 ± 0.2 CW 11 ± 54 -108 ± 62 0.52 2

HSNE50 0.3 ± 0.2 0.2 ± 0.2 CW 44 ± 89 -63 ± 105 0.81 2

HSNE60 0.3 ± 0.2 0.2 ± 0.2 CW 168 ± 90 156 ± 106 0.71 2

HSNW40 0.2 ± 0.1 0.2 ± 0.1 CW 148 ± 70 -153 ± 71 0.68 2

HSNW50 0.3 ± 0.2 0.1 ± 0.2 CW 122 ± 58 -138 ± 53 0.36 2

Site2 0.3 ± 0.2 0.1 ± 0.2 5 ± 77 18 ± 95 0.22 2

Statoil3 0.2 ± 0.2 0.1 ± 0.2 CW 116 ± 54 233 ± 55 0.3 1

Statoil4 0.1 ± 0.2 0.1 ± 0.2 CW 49 ± 119 269 ± 119 0.86 1

WBC 0.2 ± 0.3 0.0 ± 0.2 CCW 74 ± 76 124 ± 107 0.08 2

Minor

Major
Appendix B - 7

Table B-7. K1 Tidal Ellipse Parameters

Mooring 

Id

Major 

Axis 

(cm/s)

Minor 

Axis 

(cm/s)

Rotation 

Direction

Inclination 

(°)

Greenwich 

Phase (°)

# of 

Months

BC1 0.6 ± 0.3 0.1 ± 0.3 CCW 42 ± 40 125 ± 43 0.19 2

BC2 0.6 ± 0.4 0.1 ± 0.4 CCW 33 ± 43 70 ± 62 0.22 5

BC3 0.6 ± 0.3 0.2 ± 0.2 CCW 13 ± 38 175 ± 36 0.36 2

BC4 0.5 ± 0.3 0.2 ± 0.2 CCW -1 ± 40 174 ± 48 0.42 2

BC5 0.6 ± 0.3 0.1 ± 0.2 158 ± 22 -114 ± 31 0.13 2

BC6 0.6 ± 0.4 0.1 ± 0.4 CW 5 ± 64 -164 ± 50 0.18 2

Bu 0.3 ± 0.2 0.1 ± 0.2 146 ± 43 -145 ± 46 0.22 6

CPAI01 0.4 ± 0.2 0.1 ± 0.2 123 ± 48 -140 ± 47 0.31 5

CPAI02 0.6 ± 0.2 0.1 ± 0.2 CW 146 ± 22 -150 ± 24 0.2 2

Cj 0.3 ± 0.1 0.1 ± 0.1 CW 138 ± 38 -155 ± 36 0.42 6

EBC 0.8 ± 0.5 0.1 ± 0.5 CCW 42 ± 46 91 ± 47 0.08 3

HS01 0.8 ± 0.2 0.3 ± 0.2 CW 159 ± 20 187 ± 18 0.45 1

HS02 0.8 ± 0.1 0.6 ± 0.1 CW 148 ± 36 186 ± 37 0.78 1

HS03 0.8 ± 0.2 0.7 ± 0.1 CW 4 ± 37 351 ± 39 0.8 1

HS04 0.6 ± 0.1 0.4 ± 0.1 CW 157 ± 32 189 ± 31 0.69 1

HS05 0.5 ± 0.2 0.2 ± 0.1 CW 145 ± 20 205 ± 19 0.32 1

HS06 0.6 ± 0.2 0.2 ± 0.2 CW 164 ± 22 228 ± 24 0.31 1

HSNE40 0.8 ± 0.2 0.5 ± 0.2 CW 137 ± 31 -163 ± 34 0.64 2

HSNE50 0.8 ± 0.2 0.5 ± 0.2 CW 148 ± 28 -170 ± 29 0.61 2

HSNE60 0.9 ± 0.2 0.5 ± 0.2 CW 141 ± 26 -173 ± 27 0.52 2

HSNW40 0.6 ± 0.2 0.4 ± 0.1 CW 149 ± 35 -175 ± 35 0.69 2

HSNW50 0.6 ± 0.2 0.3 ± 0.2 CW 146 ± 45 -173 ± 46 0.6 2

Site2 0.5 ± 0.3 0.1 ± 0.4 20 ± 55 28 ± 74 0.34 4

Statoil3 0.4 ± 0.1 0.4 ± 0.1 CW 146 ± 108 191 ± 107 0.93 1

Statoil4 0.6 ± 0.2 0.4 ± 0.2 CW 113 ± 51 216 ± 45 0.67 1

WBC 0.8 ± 0.5 0.2 ± 0.5 -25 ± 38 160 ± 42 0.24 3

Minor

Major
Appendix B - 8

Table B-8. Q1 Tidal Ellipse Parameters

Mooring 

Id

Major 

Axis 

(cm/s)

Minor 

Axis 

(cm/s)

Rotation 

Direction

Inclination 

(°)

Greenwich 

Phase (°)

# of 

Months

BC1 0.4 ± 0.4 0.0 ± 0.3 CW 121 ± 50 156 ± 62 0.03 2

BC2 0.4 ± 0.4 0.1 ± 0.4 31 ± 76 9 ± 93 0.35 5

BC3 0.2 ± 0.2 0.0 ± 0.2 CCW 8 ± 67 -53 ± 100 0.2 2

BC4 0.4 ± 0.3 0.1 ± 0.2 18 ± 45 -76 ± 50 0.24 2

BC5 0.3 ± 0.3 0.0 ± 0.2 CW 165 ± 53 -140 ± 74 0.23 2

BC6 0.3 ± 0.4 0.1 ± 0.3 75 ± 64 -98 ± 109 0.34 2

Bu 0.2 ± 0.2 0.1 ± 0.2 19 ± 68 -93 ± 91 0.35 6

CPAI01 0.2 ± 0.2 0.1 ± 0.2 106 ± 71 -126 ± 69 0.29 5

CPAI02 0.3 ± 0.2 0.2 ± 0.2 CW 166 ± 71 -160 ± 82 0.63 2

Cj 0.1 ± 0.1 0.0 ± 0.1 70 ± 80 -70 ± 96 0.21 6

EBC 0.7 ± 0.6 0.2 ± 0.5 44 ± 61 57 ± 64 0.38 3

HS01 0.3 ± 0.2 0.0 ± 0.2 CW 168 ± 39 160 ± 45 0.04 1

HS02 0.3 ± 0.2 0.1 ± 0.1 CCW 178 ± 41 150 ± 36 0.2 1

HS03 0.3 ± 0.2 0.1 ± 0.2 CW 19 ± 47 3 ± 56 0.46 1

HS04 0.3 ± 0.1 0.1 ± 0.1 CW 156 ± 26 188 ± 27 0.23 1

HS05 0.3 ± 0.1 0.1 ± 0.2 CW 138 ± 38 215 ± 39 0.3 1

HS06 0.3 ± 0.2 0.1 ± 0.2 CCW 2 ± 43 43 ± 57 0.19 1

HSNE40 0.6 ± 0.2 0.2 ± 0.2 CW -7 ± 24 69 ± 28 0.32 2

HSNE50 0.4 ± 0.2 0.3 ± 0.2 CW -8 ± 57 57 ± 68 0.59 2

HSNE60 0.4 ± 0.2 0.2 ± 0.2 CW 133 ± 58 173 ± 61 0.57 2

HSNW40 0.3 ± 0.2 0.2 ± 0.1 CW 7 ± 62 -30 ± 75 0.58 2

HSNW50 0.3 ± 0.2 0.2 ± 0.2 CW 17 ± 57 -133 ± 66 0.48 2

Site2 0.4 ± 0.3 0.1 ± 0.3 32 ± 64 -68 ± 74 0.26 4

Statoil3 0.2 ± 0.2 0.1 ± 0.2 CW 128 ± 66 195 ± 62 0.42 1

Statoil4 0.3 ± 0.2 0.2 ± 0.2 CW 125 ± 88 207 ± 93 0.7 1

WBC 0.5 ± 0.5 0.1 ± 0.4 CW -9 ± 50 -172 ± 91 0.26 3

Minor

Major
Department of the Interior (DOI)

The Department of the Interior protects and manages the Nation's natural 

resources and cultural heritage; provides scientific and other information about 

those resources; and honors the Nation’s trust responsibilities or special 

commitments to American Indians, Alaska Natives, and affiliated island 

communities.

Bureau of Ocean Energy Management (BOEM)

The mission of the Bureau of Ocean Energy Management is to manage 

development of U.S. Outer Continental Shelf energy and mineral resources in an 

environmentally and economically responsible way.

BOEM Environmental Studies Program

The mission of the Environmental Studies Program is to provide the information 

needed to predict, assess, and manage impacts from offshore energy and marine 

mineral exploration, development, and production activities on human, marine, 

and coastal environments. The proposal, selection, research, review, 

collaboration, production, and dissemination of each of BOEM’s Environmental 

Studies follows the DOI Code of Scientific and Scholarly Conduct, in support of a 

culture of scientific and professional integrity, as set out in the DOI Departmental 

Manual (305 DM 3).