OCS Study MMS 2005-028 

CIRCULATION AND WATER PROPERTY VARIATIONS IN THE NEARSHORE 

ALASKAN BEAUFORT SEA

Thomas J. Weingartner 

Stephen R. Okkonen 

and 

Seth L. Danielson 

Institute of Marine Science 

University of Alaska 

Fairbanks, AK 99775 

June 2005 

Final Report 

MMS Contract 1435-01-00-CA-31083 

 

US Department of Interior 

Minerals Management Service 

Alaska Outer Continental Shelf Region 
i 

CIRCULATION AND WATER PROPERTY VARIATIONS IN THE NEARSHORE 

ALASKAN BEAUFORT SEA

Thomas J. Weingartner 

Stephen R. Okkonen 

and 

Seth L. Danielson 

Institute of Marine Science 

University of Alaska 

Fairbanks, AK 99775 

This was funded by the U.S. Department of the Interior Minerals Management Service (MMS), 

Alaska Outer Continental Shelf Region, Anchorage Alaska, under Contract No 1435-01-00-CA31083, as part of the MMS Alaska Environmental Studies Program. 

May 2005 

The opinions, findings, conclusions, or recommendations expressed in this report or product are 

those of the authors and do not necessarily reflect the views of the U.S. Department of the 

Interior, nor does mention of trade names or commercial products constitute endorsement or 

recommendation for use by the Federal Government. 
ii 

TABLE OF CONTENTS

ABSTRACT 1

I. INTRODUCTION 4

II. REGIONAL SETTING 6

III. METHODS 11 

IV. RESULTS 18

1. Sea ice, currents, and winds 18 

 2. Tides 40

 3. Sea Level 60 

 4. Correlations 60 

 5. Dynamics 65 

 6. Temperature, salinity, transmissivity, fluorescence 73 

V. FRESHWATER INFLUENCE 84 

VI. CONCLUSIONS AND RECOMMENDATIONS 95 

VII. ACKNOWLEDGMENTS 99 

VII. REFERENCES 100 
iii 

LIST OF TABLES

Table 1. Mooring specifics for data collection period 2200/13/AUG/1999 - 

2100/30/AUG/2000 15 

Table 2. Mooring Specifics for data collection period 2200/13/AUG/2000 - 

 0000/29/AUG/2001 16 

Table 3. Mooring Specifics for data collection period 2200/13/AUG/2001 - 

 0000/29/AUG/2002 17 

Table 4. Percent variance explained by the first empirical orthogonal function 

for velocities projected along the principal axis of variance. 23 

Table 5. Landfast ice set-up and breakup dates and times for each year and mooring. 

Table 6. Current and wind statistics 1999 – 2000. 34 

Table 7. Current and wind statistics 1999 – 2000. . 35

Table 8. Current and wind statistics 1999 – 2000. . 36

Table 9. Tidal ellipse parameters (based on vertically averaged current component) 

for major tidal constituents at each site and year. . 43 

Table 10. Wind-current correlations open-water period. 62 

Table 11. Wind-sea level correlations open-water period. 62 

Table 12. Wind – sea level correlations for all stations. 62 

Table 13. Wind-current correlations landfast ice period. 63 

Table 14. Sea-level-current correlations landfast ice period. 63
iv 

LIST OF FIGURES

Figure 1. Map of the Alaskan Beaufort Sea and North Slope with place names and 

subdivisions indicated. 7 

Figure 2. Daily runoff for a) Kuparuk and b) Sagavanirktok rivers in 2000, 2001, 

and 2002. Note that the spring freshet in 2002 occurred prior to the 

installation of the gauges. The Sagavanirktok River gauge is located about 

140 km inland from the coast so the discharge at the coast is probably greater 

than that measured by the gauge. 11 

Figure 3. Location map showing current meter moorings and the PASC weather 

station in Deadhorse. Moorings ARGO, DINKUM, and McCLURE (red 

squares) were deployed in each year from summer 1999 – summer 2002, 

while REINDEER (blue square) was deployed from summer 2001 – summer 

2002. Northstar is a production island constructed in 2000 and 2001, while 

the McCLURE mooring lies within the undeveloped Liberty field. The solid 

line through DINKUM indicates the approximate location of the June 2000 

CTD transect. Depth contours are in meters. 13 

Figure 4. Photograph of the PVC mooring frames with various instruments used 

in the study. 14 

Figure 5. Time series of (from top to bottom) ice thickness, bottom track speed, 

shear, and velocity along the principal axis of variance at DINKUM for the 

first deployment period. 20 

Figure 6. Time series of (from top to bottom) ice thickness, bottom track speed, 

shear, and velocity along the principal axis of variance at DINKUM for the 

second deployment period. (A programming error voided measurements of 

the bottom track velocity.) 21 

Figure 7. Time series of (from top to bottom) ice thickness, bottom track speed, 

shear, and velocity along the principal axis of variance at DINKUM for the 

third deployment period. (A programming error voided measurements of the 

bottom track velocity.) 22 

Figure 8. Time series of currents, sea level, and along- and cross-shore component 

of wind stress at ARGO for all landfast ice periods. 24 

Figure 9.. Time series of currents, sea level, and along- and cross-shore component 

of wind stress at DINKUM for all landfast ice periods. 25 

Figure 10.. Time series of currents, sea level, and along- and cross-shore component 

of wind stress at McCLURE for the 1999-2000 and 2000 – 2001 landfast iceperiods. 26 

Figure 11.. Time series of currents, sea level, and along- and cross-shore component 

of wind stress at REINDEER for the 2001 – 2002 landfast ice period. 27 

Figure 12. Time series of currents, sea level, and along- and cross-shore component 

of wind stress at ARGO for all open water periods. 28 
v 

Figure 13. Time series of currents, sea level, and along- and cross-shore component 

of wind stress at DINKUM for all open water periods. 29 

Figure 14. Time series of currents, sea level, and along- and cross-shore component 

of wind stress at MCCLURE for the open water periods of 2000 and 2001. 30 

Figure 15. Time series of currents, sea level, and along- and cross-shore component 

of wind stress at REINDEER for the open water period of 2002. 31 

Figure 16. Mean velocity profiles for the cross-shore (left) and alongshore (right) velocity 

components during eastward flow conditions at REINDEER for the landfast ice season. 38 

Figure 17. Mean velocity profiles for the cross-shore (left) and alongshore (right) velocity 

components during westward flow conditions at REINDEER for the landfast ice season. 38 

Figure 18. Histograms of unfiltered current speeds for the landfast ice and open 

water period at all moorings. The solid horizontal line indicates the 1% level. 39 

Figure 19. Illustration of the terms associated with the tidal ellipse (or hodograph) 

parameters: M, the semi-major axis, m, the semi-minor axis, and θ, the angle of 

inclination (from east). The tidal velocity vector rotates cyclonically 

(counterclockwise if m>0) and anticyclonically (clockwise) if m<0. 41 

Figure 20. M2 tidal properties at ARGO based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 44 

Figure 21. S2 tidal properties at ARGO based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 45 

Figure 22. K1 tidal properties at ARGO based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 46 

Figure 23. O1 tidal properties at ARGO based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 47 

Figure 24. M2 tidal properties at DINKUM based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 48 

Figure 25. S2 tidal properties at DINKUM based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 49 

Figure 26. K1 tidal properties at DINKUM based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 50 
vi 

Figure 27. O1 tidal properties at DINKUM based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 51 

Figure 28. M2 tidal properties at MCCLURE based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 52 

Figure 29. S2 tidal properties at MCCLURE based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 53 

Figure 30. K1 tidal properties at MCCLURE based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, and 

phase. Solid black line indicates landfast ice season. 54 

Figure 31. O1 tidal properties at MCCLURE based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 55 

Figure 32. M2 tidal properties at REINDEER based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 56 

Figure 33. S2 tidal properties at REINDEER based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 57 

Figure 34. K1 tidal properties at REINDEER based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 58 

Figure 35. O1 tidal properties at REINDEER based on 29-day overlapping tidal analyses. 

The parameters are (from top to bottom) the major axis, minor axis, inclination, 

and phase. Solid black line indicates landfast ice season. 59 

Figure 36. Time series of sea level corrected for the inverted barometer effect from 

September 1999 – August 2002 (blue line) and the least square harmonic fit to the annual 

and semi-annual periods (red line). 61 

Figure 37. Time series of the terms in the vertically integrated along-shelf momentum 

equation for the 2000 open water season using data from DINKUM. 68 

Figure 38. Time series of the terms in the vertically integrated along-shelf momentum 

equation for the 2001 open water season using data from DINKUM. 69 

Figure 39. Time series of the terms in the vertically integrated along-shelf momentum 

equation for the 2000 – 2001 landfast ice season using data from DINKUM. 70 

Figure 40. Time series of the terms in the vertically integrated along-shelf momentum 

equation for the 2001 – 2002 landfast ice season using data from DINKUM. 71 

Figure 41. Time series of the terms in the vertically integrated along-shelf momentum 

equation for the 2001 – 2002 landfast ice season using data from DINKUM. 72 
vii 

Figure 42. Ice thickness and bottom track velocity (upper panel) and temperature and 

salinity (lower panel) at ARGO during the 1999 – 2000 deployment. 75 

Figure 43. Ice thickness and bottom track velocity (upper panel) and temperature and 

salinity (lower panel) at ARGO during the 2000 – 2001 deployment. 76 

Figure 44. Ice thickness (upper panel), temperature and salinity (middle panel) and 

fluorescence (lower panel) at ARGO during the 2001 – 2002 deployment. 77 

Figure 45. Ice thickness (upper panel), temperature and salinity (middle panel) and 

transmissivity (lower panel) at DINKUM during the 1999 - 2000 deployment. 78 

Figure 46. Ice thickness (upper panel) and temperature and salinity (lower panel) 

at DINKUM during the 2000 - 2001 deployment. 79 

Figure 47. Ice thickness (upper panel), temperature and salinity (middle panel) and 

transmissivity and flourescence (lower panel) at DINKUM during the 2001 - 2000 

deployment. 80 

Figure 48. Ice thickness and bottom-track velocity (upper panel) and temperature 

and salinity (lower panel) at MCCLURE during the 2000 - 2001 deployment. 81 

Figure 49. Ice thickness and bottom-track velocity (upper panel), temperature and salinity 

(middle panel) and transmissivity at MCCLURE during the 2001 - 2002 deployment. 82 

Figure 50. May through June time series of (from top to bottom) Sagavanirktok River 

discharge, ice thickness (black), transmissivity (green), cross- and along-shore velocity 

shear, and cross- and along-shore velocities. Along-shore (cross-shore) components are 

red (blue). 83 

Figure 51. Salinity along the DINKUM 2001 transect (see Figure 3 for location of this 

transect). 85 

Figure 52. Temperature along the DINKUM 2001 transect (see Figure 3 for location of 

this transect). 86 

Figure 53. Transmissivity along the DINKUM 2001 transect (see Figure 3 for location of 

this transect). 87 

Figure 54. Mean velocity profiles for the cross-shore (left) and alongshore (right) velocity 

components during eastward flow conditions at REINDEER for the open water season. 91 

Figure 55. Mean velocity profiles for the cross-shore (left) and alongshore (right) velocity 

components during westward flow conditions at REINDEER for the open water season. 91 

Figure 56. Schematic of the freshwater plume types likely to form in summer and fall in 

the nearshore region of the Beaufort Sea. 92 

Figure 57. Beaufort Sea SeaWIFS imagery from August 25, 2002 (upper panel) and 

September 5, 2002 (lower panel). Winds were weak and variable for the week 

preceding these images. 93 

Figure 58. Visible SeaWIFS image from September 4, 2004 showing turbid plumes 

extending seaward from the Alaskan Beaufort Sea shelf across the continental slope. 94 
- 1 - 

ABSTRACT 

Three years of current meter and water property data were collected year-round (1999 – 

2002) from the landfast ice zone of the nearshore Alaskan Beaufort Sea. The data show large 

seasonal differences in the circulation that is defined by the set-up and breakup of the landfast 

ice. During the open water season (July – mid-October) mid-depth currents often exceed 20 cms

-1, whereas during the landfast ice season (mid-October – June) these currents are generally <10 

cm-s-1. Tidal currents are feeble (<3 cm-s-1) year-round and probably do not play a dynamically 

significant role on the inner shelf. 

Most (>90%) of the current variability is in the along-shore direction year-round. For the 

most part, mean currents are not statistically different from zero over the whole record or in 

individual seasons. Open water currents are significantly correlated with the local winds, but 

currents beneath the landfast ice are not. Calculations conducted over both seasons suggest 

along-shore sea-level gradients are about 10-6, with the magnitude of these gradients being only 

slightly larger during the open water season than during the landfast ice season. These gradients 

are presumably set-up by the winds during the open water season, but their origin during the 

landfast ice season is unknown. During the open water season upwelling-favorable winds force 

westward flows that are strongly sheared in the vertical and with maximum currents at the 

surface. In contrast, downwelling favorable winds are weakly sheared in the vertical. The 

symmetry is presumably to the strong stratification associated with upwelling winds. 

Cross-shore flows are generally small (~3 cm s-1) compared to along-shore currents. 

However, cross-shore flows of ~10 cm-s-1 were observed during the landfast ice season when the 

spring freshet resulted in an offshore spreading of a buoyant plume beneath the landfast ice. 

Although measured cross-shore flows are generally small, satellite imagery suggests that frontal 
- 2 - 

instabilities associated with low-salinity nearshore plumes can transport inner shelf waters 

offshore to the Beaufort shelfbreak during the open water season. Observations from elsewhere 

in the Arctic suggest that cross-shore current speeds associated with instabilities can be as large 

as 30 cm s-1. 

The results suggest that oil spilled beneath the landfast ice will stay within the vicinity of 

the oil spill source as current speeds will rarely exceed the threshold velocity required to 

transport an oil slick once it has attained its equilibrium thickness. Because of the broad spatial 

coherence in the flow field (at least 30 km in along-shore extent), underice currents could be 

monitored at one point and transmitted real-time to cleanup crews in the event of an underice 

spill. This information would verify the current speeds and whether oil would stay in the vicinity 

of the spill. Oil spilled during the open water season could be rapidly dispersed over great 

distance in both the along- and cross-shore directions, however. 

Water properties also vary seasonally in response to ice formation and melting, the spring 

freshet, and wind-mixing. Salinities increase and temperatures decrease throughout the winter 

due to freezing and brine expulsion from sea-ice. During the spring freshet, the inner shelf is 

strongly stratified and remains so until the ice retreats and winds mix the water column. The 

annual suspended sediment cycle, based on transmissivity measurements, suggests rapid 

deposition of river borne sediments beneath the landfast ice during the spring freshet, with resuspension and transport occurring throughout the open water season depending upon storm 

frequency. Re-suspension and transport is also vigorous during the formation of landfast ice and 

we conclude that much sediment is incorporated into the ice matrix at this time of the year. Iceincorporated sediments are either transported with the ice or returned to the water column during 

melting the following summer. 
- 3 - 

There are several important issues that we believe need to be addressed in the future. 

Modeling of the landfast ice zone requires an understanding of the role that ice-water friction 

plays in this region. This is likely influenced by the unknown, but complicated, underice 

topography. Second, the source and magnitude of the along-shore pressure gradients responsible 

for the underice currents needs to be determined. Third, it is not clear if the findings based on 

current measurements made in water depths <12 m apply to deeper portions of the landfast ice 

zone. Hence the cross-shore coherence in the underice circulation field needs to be determined. 

Fourth, the introduction of freshwater creates stratification that can lead to an asymmetric current 

response to wind-forcing during the open water season. Observations on the thermohaline 

structure of the Beaufort shelf are needed in order to understand and model the circulation field 

during the open water season. Cross-shore salinity fronts, established by river runoff, can 

become unstable and cause energetic cross-shelf flows capable of carrying pollutants far 

offshore. The dynamics and kinematics of these features need to be explored to properly include 

them into ocean circulation models. Fifth, sediments can adsorb pollutants; hence we 

recommend that consideration be given to the potential role that ice plays in the transport of 

sediment on this shelf 
- 4 - 

I. INTRODUCTION 

The nearshore zone (here defined as the portion of the shelf between the coast and the 

~20 m isobath) of arctic shelves differs substantially from their mid- and low-latitude 

counterparts because of unique geomorphologic and climatological attributes. Arctic shelf tides 

are generally weak and rivers empty directly onto the shelf rather than into bays. Thus, the 

nearshore effectively functions as an estuary, insofar as it serves as the transition and mixing 

zone between the Arctic’s terrestrial and marine environments. Climatologically, the large 

annual cycle in air-sea heat exchange and freshwater runoff cause enormous seasonal variations 

in the buoyancy and momentum fluxes to the shelf. For example, more than 90% of the annual 

river discharge from arctic rivers (defined here as rivers whose watersheds lie entirely north of 

the Arctic Circle) occurs over a brief period in early summer. (For larger rivers, such as the 

Mackenzie or Lena, approximately 90% of the annual discharge occurs between June and 

September.) Much of this discharge occurs before the landfast ice has completely retreated. 

Landfast ice is a unique characteristic of most arctic shelves. During the winter months it is 

anchored to the seafloor out to the 2 m isobath [Reimnitz, 2000] and can extend offshore to 

between the 20 and 40 m isobath [Reimnitz and Kempema, 1984; Macdonald and Carmack, 

1991]. In contrast to drifting pack ice found over the outer shelf and basin, landfast ice is 

virtually immobile. On windward shelves, such as the Beaufort, its offshore boundary is highly 

deformed when pack ice collides with the landfast ice edge. On lee shelves, such as the Laptev 

Sea, the offshore edge of the landfast ice is typically smooth and terminates at the edge of a 

polynya. The landfast ice extent varies among arctic shelves; it extends ~100 km offshore in the 

East Siberian Sea [Morris et al., 1999] but is of much smaller extent, and in a few places even 

absent, on the Chukchi shelf. Because it is effectively immobile, landfast ice inhibits the transfer 

of momentum from the wind to the ocean and therefore drastically influences arctic shelf 
- 5 - 

dynamics. Moreover, it probably reduces communication between nearshore and outer shelf 

waters where the wind stress is transmitted efficiently to the ocean. As will be seen the landfast 

ice effectively determines seasonal conditions on arctic shelves. 

Herein we report on year-round measurements of currents, temperature, salinity, 

transmissivity and fluorescence (which provide qualitative measures of suspended load and 

chlorophyll, respectively) obtained from moored instruments in the Alaskan Beaufort Sea. The 

primary goal of the program was to assess the magnitude of underice currents in the vicinity of 

Prudhoe Bay where offshore petroleum development activities were planned or underway. 

These measurements were needed to guide response procedures in the event of an accidental oil 

spill beneath the landfast ice. Laboratory studies indicate that oil slicks at equilibrium thickness 

and in contact with immobile, smooth ice begin to move at current speeds ranging from 3 cm-s-1

for low viscosity oils to 7 cm-s-1 for high viscosity oils [Cox and Schultz, 1980]. Threshold 

velocities increase to 15 - 25 cm-s-1 as the underice roughness increases. 

Previous current measurements in this region yielded contradictory results. Aagaard’s

[1984] measurements suggested that current speeds seldom exceed 10 cm-s-1 in the landfast ice 

portion of the Beaufort Sea, while Matthews [1981] inferred that speeds of up to 35 cm-s-1 were 

possible at least occasionally. Both studies were of relatively short-term duration, however with 

instruments moored close to the seabed. The velocity measurements reported here were 

collected by bottom-mounted acoustic Doppler current profilers (ADCPs) over a three-year 

period (August 1999 – September 2002) and represent 9 instrument years of data from beneath 

the landfast ice zone. Although the geographic scope of the program was limited, the results 

provide a firm basis for regional oil spill response planning. In addition, they provide a look at 
- 6 - 

this unique marine environment and raise a number of questions pertinent to the physical 

oceanography of nearshore arctic shelves influenced by landfast ice. 

The outline of the report is a follows. Section II provides background information on the 

Alaskan Beaufort Sea shelf. The field program is described in Section III and the results 

presented in Section IV. Section V explores some of the possible influences of freshwater runoff 

on the oceanic dispersal of any contaminants introduced at the coast. Conclusions and 

recommendations are summarized in section VI. 

II. THE REGIONAL SETTING 

The Alaskan Beaufort Sea shelf (Figure 1) extends ~500 km eastward from Point Barrow 

to the Mackenzie portion of the Beaufort Sea shelf in Canadian waters. The shelf width is ~80 

km as measured from the coast to the 200 m isobath. Shelf depths grade smoothly offshore with 

bottom slopes typically being ~10-3 inshore of the 100 m isobath. 

Sea ice can cover the shelf year-round, although more typically the inner shelf (and in 

recent years the entire shelf) is ice-free during the summer months. Landfast ice begins to form 

in October and extends 20 – 40 km offshore through mid-June. It covers approximately 25% of 

the shelf area [Barnes et al., 1984]. The landfast ice is relatively smooth adjacent to the coast, 

but is increasingly deformed offshore. Maximum ridge intensity and height increases moving 

seaward, and the magnitudes of both variables increase through winter [Tucker et al., 1979]. 

These parameters can also vary along-shelf and it appears that the landfast ice zone on the 

Mackenzie shelf is much less deformed than on the Alaskan Beaufort shelf [Tucker et al. 1979]. 

Ice keels form beneath the ridges and can gouge the seafloor [Barnes et al., 1984] and form piles 

of grounded ice, stamukhi, along the seaward edge of the landfast ice. The stamukhi appears to 

be important in protecting the inner shelf (and landfast ice) from pack ice forces [Reimnitz and 
- 7 - 

Figure 1. Map of the Alaskan Beaufort Sea and North Slope with place names and subdivisions 

indicated. 

Kempena, 1984] and on the Mackenzie shelf provides an effective barrier to the exchange of 

dilute nearshore waters with saline offshore waters [Macdonald and Carmack, 1991]. 

Ice deformation and stamukhi formation is related to the northeasterly winds over the 

outer shelf. Northeasterly winds prevail from fall through spring and force pack ice into the 

landfast ice edge. In summer, weak westerly winds blow over the southern Canada Basin 

[Furey, 1998] although several strong westerly wind events typically occur each summer and fall 

along the Beaufort coast as low-pressure systems propagate from the Bering Sea [Maslanik et al., 

1999]. Seasonally varying mesoscale winds may substantially alter the synoptic wind field over 

the nearshore zone, however. For example, a persistent summer sea breeze results in mean 

westward winds within ~25 km of the coast [Kozo, 1982a, b]. Brower et al. [1988] indicate that 

mean summer winds are easterly, which suggests there is a reversal in wind direction on crossing 

Kuparuk 

River 
- 8 - 

the shelf. From October through April mountain barrier baroclinicity [Kozo, 1980; 1984] can 

produce along-shore divergence in the wind field. This effect occurs when the southward flow 

of low-level cold air from the Arctic Ocean is blocked along the northern flank of the Brooks 

Range. The resulting isopycnal slopes induce eastward surface winds of O(15 m s-1) over a 

horizontal width scale of 200 – 300 km. The western Beaufort coast is rarely influenced by the 

mountain barrier effect because it lies more than 300 km north of the Brooks Range, but the 

eastern Beaufort coast lies within 60 km of the mountains. Consequently, winds can be 

westward over the western Beaufort coast but eastward along the eastern coast. Kozo [1984] 

estimated that the mountain barrier baroclinicity effect occurs ~20% of the time during winter. 

Three distinct oceanic regimes bound the Alaskan Beaufort Sea. To the west, waters of 

Pacific Ocean origin flow northward from Bering Strait and across the Chukchi shelf. While this 

flow divides along three main branches across the Chukchi shelf, the one most relevant for the 

Beaufort shelf is the outflow through Barrow Canyon in the northeast Chukchi Sea [Mountain et 

al., 1976; Aagaard and Roach, 1990; Weingartner et al., 1998; accepted]. Variability in the 

canyon outflow flow is large, especially in fall and winter, and mainly due to fluctuations in the 

regional winds [Weingartner et al., 1998; Weingartner et al., accepted; Woodgate et al., 

accepted]. Some of the Barrow Canyon outflow continues eastward as a subsurface current (or 

slope undercurrent) along the Beaufort shelfbreak and slope where it forms the upper halocline 

waters of the Canada Basin [Mountain et al., 1976; Aagaard, 1984; Pickart, 2004; Pickart et al., 

accepted]. Presumably some of the water exiting Barrow Canyon rounds Pt. Barrow and 

continues onto the inner portion of the western Beaufort shelf, although no measurements have 

been made here to examine this issue. The outer shelf and continental slope provide the offshore 

boundary for the Alaskan Beaufort Sea. In the upper 50 m or so the flow is westward and part of 
- 9 - 

the southern limb of the wind-driven Beaufort Gyre. This flow can occasionally be reversed by 

strong westerly winds and/or by occasional shelfbreak upwelling that advects eastward 

momentum from the slope undercurrent onto the shelf at least as far inshore as the 50 m isobath 

[Aagaard, 1984; Pickart 2004]. 

The Mackenzie shelf joins the Alaskan Beaufort shelf to the east and it is likely that, at 

the very least, the eastern Beaufort shelf is influenced by the year-round discharge from the 

Mackenzie River [Carmack et al., 1989; Macdonald et al., 1989; Macdonald and Carmack,

1991]. Mackenzie shelf water has been detected throughout much of the Canada basin, including 

the continental slope of the Chukchi and western Beaufort Sea as far as 160o

W longitude [Guay 

and Falkner, 1998; MacDonald et al., 1999a]. It thus seems likely that wind- driven currents 

transport Mackenzie shelf waters onto the Alaskan Beaufort shelf as well. In this regard, we note 

that the migratory behavior of arctic cisco provide indirect evidence for the intrusion of 

Mackenzie River waters onto the inner shelf of the Alaskan Beaufort Sea. These fish apparently 

require a nearshore band of low-salinity water in order to complete their annual migration 

between the Mackenzie and Colville rivers each summer [Colonell and Galloway, 1997]. The 

migratory corridor is presumably maintained by the westward drift of low-salinity water from the 

Mackenzie shelf. In addition to the Mackenzie River, a large number of smaller rivers discharge 

into the Alaskan Beaufort Sea (Figure 1). These are asymmetrically distributed with most of 

them discharging into the central and eastern portions of the shelf. This asymmetric discharge, 

along with the influence of the Mackenzie, might establish an along-shelf density gradient that 

gives rise to an along-shelf baroclinic pressure gradient. 

Our measurements were made in the vicinity of Prudhoe Bay, which lies about midway 

along the Alaskan Beaufort coast. Although a number of smaller streams empty into this area 
- 10 - 

the three major rivers (and their watershed areas) that discharge into the study region are: the 

Sagavanirktok River (4800 km2

), the Kuparuk River (8100 km2), and the Colville River (53,500 

km2). Only the first two of these are routinely gauged, and only seasonally, because the gauges 

are installed at breakup and removed in fall prior to freeze-up. The discharge time series for the 

two USGS gauged rivers are shown in Figure 2 for each summer during which ocean 

measurements were made. (The spring freshet in summer 2002 was not detected as this occurred 

earlier than normal and prior to the installation of the gauge.) The annual discharge cycle is 

characterized by a rapid initiation and increase in runoff in late May or June that lasts about 2 

weeks during which time nearly 90% of the annual discharge occurs. Following this spring 

freshet, the discharge is small and gradually decays to negligible values by October, although the 

decay can be punctuated by smaller, shorter-lived and sporadic discharge events following 

summer rain storms. The Colville has a similar seasonal cycle, although the discharge is 

substantially larger because of its larger drainage area. Unlike the Mackenzie and other large 

Arctic rivers, there is no measurable winter discharge from any North Slope rivers as most freeze 

to the bottom and all have watersheds lying entirely within drainages underlain by permafrost.. 
- 11 - 

Figure 2. Daily runoff for a) Kuparuk and b) Sagavanirktok rivers in 2000, 2001, and 2002. 

Note that the spring freshet in 2002 occurred prior to the installation of the gauges. The 

Sagavanirktok River gauge is located about 140 km inland from the coast so the discharge at the 

coast is probably greater than that measured by the gauge. 

III. METHODS 

The data set described below comes from four oceanographic moorings deployed (Figure 3) in 

Stefansson Sound and adjacent waters where present (Northstar Island) and future oil 
- 12 - 

development efforts are targeted. The area provides important habitat or is a migratory route for 

a variety of fish, birds, and marine mammals, including seals and bowhead whales. Of particular 

concern is the potential influence on the unique Boulder Patch kelp community in Stefansson 

Sound [Dunton et al., 1982], which lies near the Liberty oil development prospect in the eastern 

portion of Stefannsson Sound. Moorings ARGO, DINKUM, and McCLURE were deployed in 

late summer or fall of each year beginning August 1999. Because we were uncertain if moorings 

would survive drifting ice in this shallow (<10 m) environment the mooring deployments in 

these years took advantage of partial protection afforded by a widely-separated chain of barrier 

islands and shoal, which lie 15 – 20 km offshore. In the final year (Sept. 2001 – Sept. 2002), we 

felt confident that our mooring design minimized ice damage risk and so we deployed the 

REINDEER mooring in 12 m of water offshore of these islands. All moorings survived except 

the 2001 deployment of McCLURE, which was demolished by drifting ice 8 days after 

deployment. (Although all of the instruments were recovered, the data are unusable for our 

purposes, except for the pressure record.) To the best of our knowledge the data presented here 

are the first year-round records from the nearshore zone of any arctic shelf. 

Velocities were measured from either a 1200 or 600 kHz ADCP set in a gimbaled collar 

and mounted onto a mooring frame constructed from plastic angle stock (Figure 4). The 

gimbaled mount insured that the ADCP remained vertical even if the frame tilted from the 

horizontal by 20o

 or less after deployment. (Observations made during recovery dives indicated 

that frame tilts were always negligible. Moreover, the ADCP tilt sensor indicates that the 

instruments remained level throughout the deployment.) Ancillary instruments (Seabird, Inc., 
- 13 - 

Figure 3. Location map showing current meter moorings and the PASC weather station in 

Deadhorse. Moorings ARGO, DINKUM, and McCLURE (red squares) were deployed in each 

year from summer 1999 – summer 2002, while REINDEER (blue square) was deployed from 

summer 2001 – summer 2002. Northstar is a production island constructed in 2000 and 2001, 

while the McCLURE mooring lies within the undeveloped Liberty field. The solid line through 

DINKUM indicates the approximate location of the June 2000 CTD transect. Depth contours are 

in meters. 

MicroCats and/or SeaCats, with at least one mooring containing a strain-gauge pressure sensor) 

were fastened to the legs of each frame (Figure 4) and inclined to the vertical. At some 

locations the SeaCats included a transmissometer and in the last year a fluorometer. The inclined 

mount has no effect on the temperature and salinity measurements, however, the transmissometer 

lens might have collected settled sediment on occasion. We will note these possible biases when 

discussing the data. 
- 14 - 

The mooring positions, sensor packages, and bottom depths at each mooring site and for 

each year are listed in Tables 1 – 3. The MicroCats and SeaCats were re-deployed after the first 

year without a post-calibration (performed by the manufacturer) while freshly calibrated 

instruments were used in the final year. Based on pre- and post-calibrations and comparisons 

with the winter T/S relationship along the freezing point curve we estimate that the salinity 

values are better than 0.1. (Salinity is evaluated using the practical salinity scale and is thus 

unitless.) The salinity record at mooring ARGO was erroneously high throughout the 2000 – 

2001 winter and is not used. Hourly wind speed and direction were obtained from the Deadhorse 

airport, denoted as PASC in Figure 3. 

Figure 4. Photograph of the PVC mooring frames with various instruments used in the study. 
15 

Table 1. Mooring Specifics for Data Analysis Period 2200/14/AUG/1999 - 2100/31/AUG/2000 

Mooring Name Latitude (oN) Longitude (oW) Instruments Variables Measured 

ARGO 70o 27.177’ 148o 12.722’ 1200 kHz ADCP MicroCat 

Water Velocity 

Temperature 

Salinity 

Water Pressure 

DINKUM 70o 24.352’ 147o 53.656’ 1200 kHz ADCP SeaCat 

Water Velocity 

Temperature 

Salinity 

Transmissivity 

MCCLURE 70o 20.204’ 147o 32.701’ 1200 kHz ADCP SeaCat Water Velocity Temperature* 

Salinity* 

PASC 70o 11.7’ 148o 27.7’ Anemometer Wind Speed and Directio

• McClure SeaCat data of 1999-2000 unusable because of instrument setup problem. 
16 

•

Table 2. Mooring Specifics for Data Analysis Period 2200/3/SEP/2001 - 1700/19/AUG/2002 

Mooring Name Latitude (oN) Longitude (oW) Instruments Variables Measured 

ARGO 70o 27.172’ 148o 12.666’ 1200 kHz ADCP MicroCat 

Water Velocity 

Temperature 

 Salinity* 

 Water Pressure 

DINKUM 70o 24.371’ 147o 53.632’ 1200 kHz ADCP MicroCat 

Water Velocity 

Temperature 

 Salinity 

 Water Pressure 

MCCLURE 70o 20.164’ 147o 32.700’ 

1200 kHz ADCP 

 SeaCat 

MicroCat 

Water Velocity 

Temperature 

 Salinity 

 Water Pressure 

Transmissivity 

PASC 70o 11.7’ 148o 27.7’ Anemometer Wind Speed and Directio

*ARGO salinity record in 2000-2001 not used because of sediment/biofouling and calibration problems. 
17 

Table 3. Mooring Specifics for Data Analysis Period 2200/2/SEPT/2001 - 1700/18/AUG/2002 

Mooring Name Latitude (oN) Longitude (oW) Instruments Variables Measured 

ARGO 70o 27.168’ 148o 12.654’ 1200 kHz ADCP SeaCat 

Water Velocity* 

Temperature 

 Salinity 

 Fluorescence 

DINKUM 70o 24.371’ 147o 53.632’ 600 kHz ADCP SeaCat 

Water Velocity* 

Temperature 

 Salinity 

 Transmissivity 

Fluorescence 

MCCLURE 70o 20.164’ 147o 32.700’ 1200 kHz ADCP SeaCat 

Water Velocity** 

Temperature** 

 Salinity** 

 Water Pressure 

 Transmissivity** 

Fluorescence** 

REINDEER 70o 30.539’ 148o 19.212’ 1200 kHz ADCP Water Velocity* 

PASC 70o 11.7’ 148o 27.7’ Anemometer Wind Speed and Directio

* Bottom Track data on all ADCP instruments for this 2001-2002 was not collected. Water velocity data is good

** McClure 2001-2002 mooring hit by ice 8 days into record and most data not usable. Pressure data is good. 
18 

IV. RESULTS 

1. Sea ice, currents, and winds 

Before presenting the record length and seasonal statistics, we describe the seasonal cycle 

of currents and ice thickness for each year from the DINKUM mooring (Figures 4 -6). Sea ice 

thickness, Hice, was estimated from ADCP data as the difference between the bottom depth, Zb

(Table 1-3), and the sum of the height of the transducer above the bottom, HT, and the range 

from the transducer at which the intensity of the reflected acoustic signal was maximum, R, (e.g., 

Hice = {Zb – [HT + R]}*1.12), where the factor 1.12 is the ratio of the sea water density to the sea 

ice density. There is an uncertainty in the thickness estimates of 0.5 m because the ADCP 

integrates return signals over a 0.5 m depth bin. This error occasionally introduces highfrequency noise in the data which was eliminated by smoothing the ice thickness time series with 

a 5-day running mean. Sea ice forms in early October, gradually thickens to ~1.7 – 2.2 m by 

mid-March, and then remains constant through mid-June. Ice then melts rapidly and disappears 

by mid-July. The seasonal asymmetry in ice growth and ablation rates occurs because new ice 

forms on the underside of the ice with the freezing rate decreasing exponentially with Hice

[Maykut, 1986]. By contrast, ablation primarily occurs at the surface and proceeds rapidly once 

the surface snow cover melts or the surface is covered by low-albedo, silt-laden waters from 

river overflow [Dean et al., 1994; Searcy et al., 1996]. 

Unfiltered velocity time series from the uppermost bin (bottom-track velocity; UBT in the 

figures) are large and noisy when the surface is ice-free, but negligible once landfast ice is 

established. (A programming error in the final year’s deployment voided the ice-tracking 

capability on the ADCPs. Instead of using UBT to indicate the onset and end of the landfast ice 

period we used the large seasonal reduction in current speeds to approximate the dates of 
19 

landfast ice formation and retreat in this year.) The abrupt transition in UBT allows unambiguous 

definition of the open water and landfast ice periods in each record. The data suggest that in 

each year landfast ice is present from mid-October through the end of June or the beginning of 

July, while the open water season (including drifting ice) spans the July through mid-October 

period. Thus landfast ice is present about eight and one-half months. The figures also show time 

series of the current shear, d|V|/dz, (where |V| is the speed and z is depth) between the shallowest 

depth bin beneath the ice or sea surface and the deepest depth bin. At the DINKUM mooring the 

depth range over which these differences are computed varies ~4 m during the open water period 

to ~2 m during maximum ice thickness. Current speeds typically differ in the vertical by <4 cm 

s

-1 when ice is present, but can exceed 40 cm s-1 during the open water period. The bottom panel 

in each plot shows the unfiltered time series of the velocity component projected onto the 

principal axis of variance (listed in Tables 6 - 8). These indicate that the largest current speeds 

and variance occurs when landfast ice is absent. This seasonal scenario is repeated in all three 

years among all moorings. Table 5 lists the dates of landfast ice set-up and break-up for each 

year and mooring based on analysis of the ADCP data sets. 

Table 5. Landfast ice set-up and breakup dates and times for each year and mooring. 

Landfast Ice Setup Date 

 Argo Dinkum McClure Reindeer 

Year 1999 2000 2001 1999 2000 2001 1999 2000 2001 1999 2000 2001 

Month 10 10 10 10 10 10 10 10 10 

Day 15 20 13 15 22 13 15 22 10 

Hour 0 3 12 0 16 10 0 16 15 

 

Landfast Ice Breakup Date 

 Argo Dinkum McClure Reindeer 

Year 2000 2001 2002 2000 2001 2002 2000 2001 2002 2000 2001 2002 

Month 6 7 6 6 7 6 6 7 6 

Day 30 4 30 30 2 23 30 2 27 

Hour 23 18 1 23 5 23 23 0 14 
20 

Figures 8 - 15 are time series of the low-pass filtered (35-hr cutoff) major and minor 

velocity components, demeaned (over the deployment period) sea-level, and the wind stress 

components based on Deadhorse winds for each mooring are shown for the landfast ice (Figures 

8 – 11)and open water (Figures 12 – 15) seasons. Visual inspection suggests that the current 

components along the principal axes are highly correlated with one another and this is confirmed 

through empirical orthogonal function (EOF) analysis (based on the correlation matrix) whether 

calculated over the whole record or separately for each season. The results for each deployment 

(Table 5) show that the first mode accounts for ~90% of the variance, while the higher modes 

Figure 5. Time series of (from top to bottom) ice thickness, bottom track speed, shear, and 

velocity along the principal axis of variance at DINKUM for the first deployment period (1999 – 

2000). 
21 

Figure 6. Time series of (from top to bottom) ice thickness, bottom track speed, shear, and 

velocity along the principal axis of variance at DINKUM for the second deployment period 

(2000 – 2001). 
22 

Figure 7. Time series of (from top to bottom) ice thickness, bottom track speed, shear, and 

velocity along the principal axis of variance at DINKUM for the third deployment period (2001 

– 2002). A programming error voided measurements of the bottom track velocity. 
23 

are not significant based on North et al.’s [1982] criterion. The dominance of the first EOF 

implies that flow variations in this region are coherent over spatial scales of at least 30 km so that 

a single mooring could be used to capture the local along-shore flow variability. Although this 

finding is not surprising given the simple bathymetry of the area, it was not clear at the outset of 

the study that these scales should hold in the presence of landfast ice because complex underice 

topography could generate velocity variability over small horizontal scales and thus degrade 

spatial coherence. There might in fact be seasonal differences in the velocity de-correlation 

scales because of the landfast ice, but if so these scales are larger than 30 km. Indeed the 

reduction in the variance explained by the first EOF mode during the 1999-2000 landfast ice 

season relative to later years could be due to differences in underice topography, although we 

have no means to examine this possibility. 

Table 5. Percent variance explained by the first empirical orthogonal function for velocities 

projected along the principal axis of variance.

Period Analyzed 1999-2000 2000-2001 2001 – 2002 

Landfast 77 90 94 

Open Water 93 94 93 

Whole Record 95 93 93 
24 

Figure 8. Time series of currents, demeaned sea level, and along- and cross-shore component of 

wind stress at ARGO for all landfast ice periods. 
25 

Figure 9. Time series of currents, demeaned sea level, and along- and cross-shore component of 

wind stress at DINKUM for all landfast ice periods. 
26 

Figure 10. Time series of currents, demeaned sea level, and along- and cross-shore component 

of wind stress at McCLURE for the 1999-2000 and 2000 – 2001 landfast ice periods. 
27 

Figure 11. Time series of currents, demeaned sea level, and along- and cross-shore component 

of wind stress at REINDEER for the 2001 – 2002 landfast ice period 
28 

Figure 12. Time series of currents, demeaned sea level, and along- and cross-shore component 

of wind stress at ARGO for all open water periods. 
29 

Figure 13. Time series of currents, demeaned sea level, and along- and cross-shore component 

of wind stress at DINKUM for all open water periods. 
30 

Figure 14. Time series of currents, demeaned sea level, and along- and cross-shore component 

of wind stress at MCCLURE for the open water periods of 2000 and 2001. 
31 

Figure 15. Time series of currents, demeaned sea level, and along- and cross-shore component 

of wind stress at REINDEER for the open water period of 2002. 
32 

Summary statistics for the currents and winds for each year and season are listed in 

Tables 6 – 8 as a function of record length and season (landfast versus open water) as defined by 

the ADCP parameters. The current meter data were filtered prior to computing the statistics in 

order to focus on the subtidal properties of the flow field. The tidal properties are discussed 

separately later. Statistical significance is based on 95% confidence limits (listed in the tables) 

using the effective number of degrees of freedom (Neff), computed as Neff = N∆t/τ. Here ∆t is the 

sampling interval, N is the number of data points in the time series, and τ is the integral time 

scale of the currents or winds. The integral time scale is the decay time scale for a current or 

wind event and is the time required to obtain a new independent observation in an autocorrelated time series. 

The winds are weakly westward with record-length averages ranging from 1 – 3 m s-1

(depending upon deployment period) and statistically different from zero at the 95% confidence 

level. However, mean westward winds do not prevail throughout the year. For example, the 

mean wind was westward during the open water period of 1999-2000, but not significantly 

different from zero during the landfast ice period of that deployment. The opposite situation 

occurred in 2000-2001 and 2001-2002, when mean winds were westward during the landfast ice 

period, but not significantly different from zero during the open water seasons of these two 

years. The open water seasonal mean along-shelf wind, averaged over all years, is ~1m-s-1 and 

thus substantially smaller than the ~3.5 m-s-1 mean wind speed indicated for summer along the 

Beaufort coast in the climatology of Brower et al. [1988]. There is also very little difference in 

the wind variances among the deployment years and between seasons. The lack of a distinct 

annual cycle in Deadhorse winds contrasts with the prominent annual cycle in winds computed 

over the southern Canada Basin using synoptic forecast fields [Furey, 1998]. He found that the 
33 

winds are westward on average from fall through spring, while in summer they are weakly 

eastward. He found a distinct annual cycle in wind variance, especially in the zonal velocity 

component, with the variance being three times greater in winter than in summer. Although the 

data sets and time periods used in comparing these annual cycles differ, this discussion suggests 

that there might be a cross-shelf gradient in wind stress. Mesoscale effects could be responsible 

for cross-shelf shears in the wind field. For example, sea-breezes (open water season only) 

promote westward coastal winds with this influence limited to the innermost 25 km of the shelf 

according to Kozo’s [1982b] model results. In winter, the mountain barrier effect might 

influence winds at Deadhorse [Kozo, 1984] with this phenomenon occurring when atmospheric 

pressure systems force cold, polar air masses southward towards the Brooks Range. Kozo [1984] 

estimated that this baroclinic effect occurs up to 20% of the time during winter. However, his 

estimate was based on limited observations conducted during the 1970s. The frequency in which 

the mountain barrier baroclinic effect occurs should vary interannually in association with yearto-year differences in the large scale atmospheric pressure field over the arctic. 

For most moorings and deployment periods the mean current vectors are not significantly 

different from zero, and when significant, they are small with mean alongshore velocities being 

<3 cm-s-1. In fact, the largest mean values occurred during the open water period in 1999 when a 

statistically significant mean westward flow was observed at all sites except MCCLURE. The 

relatively large mean values for this portion of the record are biased, however, by the strong 

easterly storm event that occurred in early October 1999. Our mean values are somewhat 

surprising given the general belief that the mean flow is westward during the open water season 

[Barnes and Reimnitz, 1974; Wiseman et al., 1974]. 
34 

Table 6. Current and wind statistics 1999 – 2000. The direction toward which the mean velocity vector points is axis, % refers to the percentage of the velocity variance accounted for by the component projected onto the axis wThe mean north-south <V> and east-west <U> velocity components and + the 95% confidence limits are listed. T

computed using Neff based on the integral time scale (τ) determined from the first zero crossing of the autocorrelsignificantly different from zero are italicized. The variances (s2) of the V and U are given along with the ratio of 

to the total variance, which is the variance of the filtered data divided by the variance of the unfiltered data. 

Mooring Mean Velocity Principal Axis

Speed 

(cm s-1) ΘM % ΘP

Max 

Speed 

(cm s-1) 

τ 

 (days)

<V> 

(cm s-1) s2V

<U> 

(cm s-1) 

Full Record

Argo 3.4 277 99 96 111 4.5 0.4 + 0.3 4 -3.4 + 2.0 

Dinkum 2 307 98 119 110 4 1 + 1 49 -2 + 1.4 

McClure 2.2 3 92 151 68 5.5 2.2 + 1.6 87 0.1 + 1.0 

Winds (m s-1) 1.3 250 90 77 25 3 -0.5 + 0.3 6 -1.3 + 1.1 

Landfast Ice 

Period

Argo 1.6 283 98 99 24 2.5 0.4 + 0.1 1 -1.6 + 0.7 

Dinkum 0 - 94 141 20 2 0 + 0.3 8 0 + 0.3 

McClure 1.1 29 89 176 14 3 1 + 0.6 9 0.6 + 0.2 

Winds (m s-1) 0.8 249 91 75 25 3.5 -0.3 + 0.4 6 -0.8 + 1.3 

Open Water 

Period

Argo 7.3 273 99 96 111 4.5 0.4 + 0.8 5 -7.3 + 5.9 

Dinkum 6 306 99 117 110 4 4 + 1.4 126 -5 + 3.0 

McClure 4.8 350 93 149 68 5.5 4.7 + 4.6 244 -0.8 + 2.9 

Winds (m s-1) 2.8 251 89 83 20 3 -0.9 + 0.4 5 -2.6 + 2.1 
35 

Table 7. Current and wind statistics 2000 - 2001. 

Mooring Mean Velocity Principal Axis

Speed 

(cm s-1) ΘM % ΘP

Max 

Speed 

(cm s-1) 

τ 

 (days)

<V> 

(cm s-1) s2V

<U> 

(cm s-1) 

Full Record

Argo 0.4 277 99 98 51 4 0.0 + 0.2 3 -0.4 + 1.4 

Dinkum 0.5 50 98 120 58 4.5 0.3 + 1 30 0.4 + 1.0 

McClure 0.8 11 81 136 60 4.5 0.8 + 0.7 24 0.2 + 0.8 

Winds (m s-1) 2.0 250 87 74 24 2.5 -0.7 +0.2 6 -1.9 + 0.8 

Landfast Ice 

Period

Argo 0.2 303 99 98 22 3 0.1 + 0.3 2 -0.2 + 1.4 

Dinkum 0.5 47 98 125 15 5 0.3 + 1 10 0.3 + 1.0 

McClure 1.0 4 94 154 11 4 1.0 + 0.8 9 0.1 + 0.3 

Winds (m s-1) -2.6 250 88 73 24 1 -0.9 + 0.4 6 -2.5 + 1.5 

Open Water 

Period

Argo 0.9 261 99 97 51 4 -0.1 + 0.5 5 -0.9 + 4.0 

Dinkum 1 57 98 119 58 4.5 0.4 + 1.5 78 0.6 + 4.5 

McClure 0.5 43 80 131 60 3.5 0.4 + 1.9 60 0.4 + 2.5 

Winds (m s-1) 0.9 249 84 76 17 2.5 -0.3 + 0.6 6 -0.8 + 1.3 
36 

Table 8. Current and wind statistics 2001 - 2002. 

Mooring Mean Velocity Principal Axis

Speed 

(cm s-1) ΘM % ΘP

Max 

Speed 

(cm s-1) 

τ 

 (days)

<V> 

(cm s-1) s2V

<U> 

(cm s-1) 

Full Record

Argo 0.4 115 98 101 71 2.5 -0.2 + 0.2 4 0.4 + 1.0 

Dinkum 0.9 115 97 126 102 3 -0.4 + 0.7 39 0.8 + 1.0 

Reindeer 1 120 96 104 79 3 -0.7 + 0.3 4 1.2 + 0.7 

Winds (m s-1) 0.8 245 85 73 25 1.5 -0.3 + 0.3 7 -0.7 + 0.6 

Landfast Ice 

Period

Argo 0.4 110 98 106 23 2.5 -0.1 + 0.2 3 0.3 + 0.6 

Dinkum 0.7 136 97 132 14 2 -0.5 + 0.3 10 0.5 + 0.4 

Reindeer 0.9 112 98 109 20 2 -0.3 + 0.2 3 0.8 + 0.5 

Winds (m s-1) 0.9 268 88 72 25 1.5 0.0 + 0.4 7 -0.9 + 0.8 

Open Water 

Period

Argo 0.6 124 99 100 71 2.5 -0.3 + 0.6 9 0.5 + 3.6 

Dinkum 1.6 90 98 125 102 3 0.0 + 2.5 113 1.6 + 3.6 

Reindeer 2.7 127 95 99 79 2.5 -1.6 + 0.6 7 2.2 + 2.2 

Winds (m s-1) 0.9 197 80 76 20 1.5 -0.9 + 0.4 6 -0.3 + 0.9 
37 

Although the mean currents are small or negligible, the variability is large. Most of the 

current variance is contained in the along-shore velocity component where >~90% of the current 

variability is associated with this principal axis. For each mooring these axes are oriented 

approximately along-shore (e.g., northwest to southeast). Similarly, the winds blow primarily 

along-shore such that >85% of the wind variance is aligned along the east-west axis. As 

suggested by the current time series in the lowest panels of Figures 5-7, the current velocity 

variance changes seasonally, with the variance during the landfast ice period being roughly an 

order of magnitude smaller than the variance of the open water season. Along-shore current 

variances do not change significantly throughout the landfast ice season. This contrasts with 

findings on the Chukchi Sea shelf [Aagaard and Roach, 1990; Weingartner et al., 1998; 

accepted] where current variance is generally much higher from November through January than 

from February through April. 

Figures 16 and 17 show the mean velocity profiles projected along the major and minor 

principal axis for the landfast ice period at Reindeer for all eastward and westward flow events. 

The depth axis is the scaled depth which varies from 0 for the bin closest to the ADCP transducer 

to 1 for the bin nearest to the ice. The scaling takes into account the changing depth of the water 

column due either to ice growth or sea level fluctuations. In both cases the mean velocity varies 

by ~1 cm s-1 over the depth of the water column. The open water season velocity profiles are 

considerably different under eastward and westward flows due to the stratifying influence of 

freshwater runoff. These profiles will be discussed in Section V. 
38 

Figure 16. Mean velocity profiles for the cross-shore (left) and alongshore (right) velocity 

components during eastward flow conditions at REINDEER for the landfast ice season. 

Figure 17. Mean velocity profiles for the cross-shore (left) and alongshore (right) velocity 

components during westward flow conditions at REINDEER for the landfast ice season. 

These seasonal current speed differences are also reflected in the summary speed 

histograms constructed from the unfiltered data for all years (Figure 18). Less than 2% of all 

current speeds exceed 15 cm s-1 during the landfast ice period, while more than 50% of the 

observations exceed this speed during the open water season. Our direct measurements during 
39 

the landfast ice period are thus more in line with Aagaard’s [1984] finding than with Matthews’

[1981] inferences that current speeds could be as large as 35 cm s-1. Speeds of this magnitude 

were never observed under the landfast ice except once at ARGO where the maximum speed was 

25 cm-s-1. In general we find that ~90% of the underice current speeds are <10 cm s-1. 

Figure 18. Histograms of unfiltered current speeds for the landfast ice and open water period at 

all moorings. The solid horizontal line indicates the 1% level. 

The ratio of the subtidal to the total variance was calculated by comparing the variance 

from the filtered data to the unfiltered data. Subtidal variance accounts for from 50 – 80% of the 

total variance during the land fast ice period and for more than ~80% of the total variance during 

the open water period. There are differences in these ratios among years and sites, however. For 

example, the ratios at McCLURE are generally smaller than elsewhere, although not because of 
40 

smaller tidal current amplitudes, which as shown later, are similar at all locations. More likely 

these differences reflect the geometry of Stefansson Sound, which broadens from west to east. 

Current variations are consequently larger in the western and central portions of the sound (e.g., 

DINKUM, ARGO) than in the eastern sound (e.g., McCLURE). There are also interannual 

differences in this ratio. For example, at DINKUM the subtidal variance amounts to 54%, 74%, 

and 63% of the total variance for each landfast ice period. Interannual variations in these ratios 

reflect differences in the subtidal current variance and the differences between years are 

statistically significant at the 95% significance level based on the F-statistic. Although the 

subtidal variances differ significantly among years, there are no corresponding significant yearto-year differences in wind variances. This suggests little coupling between local winds and the 

underice currents, a point which we will return to later. 

2. Tidal properties 

The tides were estimated for each current meter record following Foreman [1978] and 

Pawlowicz et al. [2002]. The four largest constituents and periods (Tp) are the semi-diurnal, M2, 

(lunar, TM2 = 12.42) and S2 (solar TS2 = 12.00) and the diurnal O1 (luni-solar, TO1 = 25.82 hrs) 

and K1 (lunar, TK1 = 23.93 hrs). Tidal hodograph analyses are summarized in Table 9. Figure 

19 illustrates the various meaning of the terms that describe the hodograph. The M2 constituent 

has the largest current amplitudes (~2 –3 cm s-1), while the other species have amplitudes of <1 

cm s-1. The semi-diurnal tides execute nearly circular tidal motions and rotate anticyclonically, 

while the diurnal tides are nearly rectilinear (flattened ellipses) and oriented alongshore. Their 

rotational sense is ill-defined because the semi-minor axis is not significantly different from zero. 

The semi-major axis, semi-minor axis, ellipse orientation, and Greenwich phase lag, 

calculated from successive 29-day overlapping segments of the current records at each mooring 
41 

and for all years are shown in Figures 20 – 35 for the four largest tidal constituents. Hodograph 

parameters vary throughout the year, particularly for the semi-diurnal species because of 

seasonal changes in ice thickness and extent, possibly larger-scale underice topography that 

affects frictional stresses between the propagating tidal wave and sea-ice, and stratification. 

Although the seasonal variations are generally small, those associated with the spring 

freshet are substantial. Consider for example, the variation in the hodograph parameters of the 

M2 tide at DINKUM (Figure 24). The semi-major axis varies seasonally and with depth but is a 

maximum of 5 cm s-1 in July 2001 beneath the ice during river breakup and a minimum of about 

1.5 cm s-1 shortly thereafter near the bottom. Depth variations in tidal properties indicate either 

generation of an internal tide associated with stratification or vertical variations in eddy viscosity 

[Danielson and Kowalik, in press] associated with the strong pycnocline established as the spring 

Figure 19. Illustration of the terms associated with the tidal ellipse (or hodograph) parameters: 

M, the semi-major axis, m, the semi-minor axis, and θ, the angle of inclination (from east). The 

tidal velocity vector rotates cyclonically (counterclockwise if m>0) and anticyclonically 

(clockwise) if m<0. 
42 

freshet flows beneath the landfast ice as discussed later. Tidal velocity shears are also small 

being ~0.003 s-1 at most and are not an important source of kinetic energy for mixing (as shown 

later). The tidal analyses suggest that hodograph properties vary seasonally throughout the 

landfast portions of all arctic shelves. While these variations are unlikely to be dynamically 

significant in the Alaskan Beaufort Sea, they might be important in the landfast ice zones of the 

Laptev and Kara seas, where tides are much more energetic [Kowalik and Proshutinsky, 1994]. 

Furthermore, the results suggest caution should be exercised when comparing tidal hodographs 

from short-term observations in the landfast ice zone with those from numerical tidal models. 

Differences between the two might not necessarily indicate model deficiencies but instead be due 

to seasonal variations in the tidal structure masked by short-duration measurements. We also 

note that the effect of landfast ice should have a strong influence on the structure of the superinertial tidal (M2 and S2) currents. The boundary layers associated with these constituents might 

overlap in shallow water and promote mixing (assuming sufficient tidal energy dissipation). As 

the water depth increases the tidal boundary layers will separate, possibly allowing a stratified 

layer to be maintained in the middle of the water column. Conceivably this might lead to the 

development of a tidal front further offshore, but inshore of the landfast ice edge. Such fronts 

are capable of supporting an along-shore flow, and if present, likely to be found in the Laptev or 

Kara Sea where tidal velocities are substantially greater. 
43 

Table 9. Tidal ellipse parameters (based on the vertically averaged current component) for major 

tidal constituents at each site and year. Major (M) is the semi-major axis of the current ellipse 

and minor (m) is the semi-minor axis (negative values imply clockwise rotation of the velocity 

vectors and positive values imply counterclockwise rotation), The inclination (θ) is the angle M

makes from the east. (The figure following the table illustrates the geometrical meaning of these 

terms.) The Greenwich phase lag is the time of the maximum constituent velocity relative to 

Greenwich. Values in parenthesis are the 95% confidence limits for each estimated parameter. 

Luni-solar diurnal K1 (Period = 23.93 hours) 

Name Year Major 

(M; cm s-1) 

Minor 

(m; cm s-1) 

Inclination, θ 

degrees from east) 

Greenwich Phase 

Lag 

Argo 1999 0.9 (0.2) -0.1 (0.1) 7.1 (4.6) 273.2 (11.8) 

Argo 2000 0.7 (0.2) -0.1 (0.1) 0.6 (6.4) 271.6 (19.5) 

Argo 2001 1.0 (0.3) -0.1 (0.1) 173.8 (5.6) 83.6 (14.2) 

Dinkum 1999 0.5 (0.1) -0.1 (0.1) 19.6 (15.7) 239.3 (17.7) 

Dinkum 2000 0.6 (0.2) -0.2 (0.2) 170.1 (19.9) 79.5 (27.0) 

Dinkum 2001 0.7 (0.2) -0.3 (0.1) 170.1 (14.5) 83.3 (18.2) 

McClure 1999 0.3 (0.1) -0.1 (0.1) 57.2 (20.6) 238.0 (22.9) 

McClure 2000 0.5 (0.1) -0.1 (0.1) 80.7 (13.4) 254.6 (12.8) 

Reindeer 2001 0.8 (0.2) -0.0 (0.1) 162.1 (8.7) 101.4 (14.7) 

Principal lunar diurnal O1 (Period = 25.82 hours) 

Argo 1999 1.1 (0.2) -0.2 (0.1) 2.1 (4.4) 310.4 (10.6) 

Argo 2000 1.0 (0.2) 0.1 (0.1) 179.6 (4.4) 121.8 (14.4) 

Argo 2001 1.2 (0.2) 0.1 (0.1) 177.6 (4.2) 82.3 (10.6) 

Dinkum 1999 0.5 (0.2) -0.1 (0.1) 19.7 (16.5) 304.6 (20.5) 

Dinkum 2000 0.8 (0.2) 0.1 (0.2) 170.8 (11.5) 99.9 (14.1) 

Dinkum 2001 0.9 (0.2) 0.2 (0.1) 167.1 (11.1) 80.1 (12.3) 

McClure 1999 0.4 (0.1) 0.0 (0.1) 49.4 (16.8) 297.7 (16.5) 

McClure 2000 0.4 (0.1) 0.2 (0.1) 38.9 (24.5) 314.0 (22.8) 

Reindeer 2001 1.0 (0.2) 0.0 (0.1) 163.2 (7.6) 80.5 (11.8) 

Principal lunar semi-diurnal M2 (Period = 12.42 hours) 

Argo 1999 1.8 (0.1) -1.3 (0.1) 28.4 (6.2) 62.4 (6.9) 

Argo 2000 1.7 (0.1) -1.3 (0.1) 27.0 (8.4) 81.5 (8.8) 

Argo 2001 1.6 (0.1) -1.3 (0.1) 35.4 (11.7) 55.0 (13.1) 

Dinkum 1999 2.0 (0.1) -0.6 (0.1) 93.0 (3.1) 18.5 (2.5) 

Dinkum 2000 1.7 (0.1) -1.2 (0.1) 78.0 (10.3) 34.8 (10.0) 

Dinkum 2001 2.0 (0.1) -1.0 (0.1) 81.6 (5.5) 17.8 (5.6) 

McClure 1999 2.2 (0.1) -0.2 (0.1) 78.9 (1.9) 28.5 (1.9) 

McClure 2000 2.2 (0.2) -0.4 (0.1) 94.8 (3.2) 53.4 (4.0) 

Reindeer 2001 1.4 (0.1) -0.8 (0.1) 139.4 (5.3) 321.1 (5.8) 

Principal solar semi-diurnal S2 (Period = 12.00 hours) 

Argo 1999 0.8 (0.1) -0.6 (0.1) 39.5 (17.7) 70.0 (17.1) 

Argo 2000 1.0 (0.1) -0.6 (0.1) 23.5 (10.8) 78.6 (10.6) 

Argo 2001 0.8 (0.1) -0.6 (0.1) 25.5 (16.6) 76.2 (16.0) 

Dinkum 1999 0.9 (0.1) -0.3 (0.1) 92.9 (7.0) 31.9 (6.6) 

Dinkum 2000 0.8 (0.1) -0.5 (0.1) 68.3 (18.9) 44.8 (18.0) 

Dinkum 2001 0.8 (0.1) -0.5 (0.1) 80.4 (16.7) 33.0 (16.3) 

McClure 1999 1.0 (0.1) -0.1 (0.1) 76.0 (4.0) 44.3 (4.7) 

McClure 2000 1.0 (0.1) -0.2 (0.1) 92.8 (6.8) 53.6 (10.7) 

Reindeer 2001 0.7 (0.1) -0.4 (0.1) 149.7 (10.6) 322.6 (11.1) 
44 

Figure 20. M2 tidal properties at ARGO based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
45 

Figure 21. S2 tidal properties at ARGO based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
46 

Figure 22. K1 tidal properties at ARGO based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
47 

Figure 23. O1 tidal properties at ARGO based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
48 

Figure 24. M2 tidal properties at DINKUM based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
49 

Figure 25. S2 tidal properties at DINKUM based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
50 

Figure 26. K1 tidal properties at DINKUM based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
51 

Figure 27. O1 tidal properties at DINKUM based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
52 

Figure 28. M2 tidal properties at MCCLURE based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
53 

Figure 29. S2 tidal properties at MCCLURE based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
54 

Figure 30. K1 tidal properties at MCCLURE based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
55 

Figure 31. O1 tidal properties at MCCLURE based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
56 

Figure 32. M2 tidal properties at REINDEER based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
57 

Figure 33. S2 tidal properties at REINDEER based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
58 

Figure 34. K1 tidal properties at REINDEER based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
59 

Figure 35. O1 tidal properties at REINDEER based on 29-day overlapping tidal analyses. The 

parameters are (from top to bottom) the major axis, minor axis, inclination, and phase. Solid 

black line indicates landfast ice season. 
60 

3. Sea level

Time series of the unfiltered and de-meaned sea level (after correcting for the inverted 

barometer effect) for all years are shown in Figure 36. Sea-level fluctuations generally range 

between + 0.5 m with a minimum of -1.1m (October 1999) to a maximum of 1.2 m (August 

2000). Sea-level fluctuations vary seasonally with rms values being about 0.18 m during the 

open water season and 0.15 m during the landfast ice period. There are also significant annual 

and semi-annual sea-level variations, which together account for ~30% of the sea-level variance. 

Sea-level is a maximum (+0.2 m) in August and a minimum in April-May (-0.2 m) with subminima (sub-maxima) occurring in October-December (January-February). The annual range 

(~0.4 m) cannot be explained by thermosteric (sea-level changes associated with changes in seawater density) effects, which we estimate are ~0.07 cm. Sea-level fluctuations can be quite large 

at shorter periods. For example, in April 2000 and in late April 2003 sea level rose by ~1 m and 

then rapidly decreased by more than 1 m over a 10 day period, with both events coincident with 

rapid reversals in the along-shore wind stress. 

4. Correlations 

We turn now to an examination of the seasonal relationship among wind stress, currents, 

and sea level (η) using both correlations (Tables 10 – 14) and dynamical estimates of the terms 

in linearized along-shelf momentum equation. All correlations are computed after low-pass 

filtering the currents and sea level time series. Wind stress (rather than wind) is the dynamically 

important variable and the stresses are computed as: 

 x y τ =ρ τ =ρ aD W aD W C WU ; C WV ;
61 

where τ

x (y)

 is the east-west (x) or north-south (y) component of the stress, ρa is the air density, CD

is an open water value for the drag coefficient (2.5 x 10-3), W is the wind speed, and UW (VW) is 

the east-west (north-south) component of the wind velocity. 

Figure 36. Time series of sea level corrected for the inverted barometer effect from September 

1999 – August 2002 (blue line) and the least square harmonic fit to the annual and semi-annual 

periods (red line). 

Both τ

x

 (Table 10) and η (Table 11) are significantly correlated with the along-shore 

currents at all moorings for all open water periods, although these variables explain generally 

less than 20 – 50% of the current variance. Correlations between τ

y

 and along-shelf currents are 

small and, in general, not statistically significant. Statistically significant positive correlations 

are found between τ

x

 and η during the open water period (first row Table 12) with the sign of the 
62 

correlation consistent with eastward winds inducing a sea-level set-up and westward winds 

leading to a sea level decrease during the open water season. The correlation is smaller during 

the 2002 open water season than in the other years, perhaps because of the shorter record length 

and the absence of large storms during that summer. 

Table 10. Open-water period wind-current correlations. (Italicized entries indicate 

statistical significance at the α < 0.05 level.). 

Mooring Time Period R: (τ

x

 vs Ur) R: (τ

y

 vs Ur) 

ARGO 2000: (6/30 – 10/20) 0..68 0.13 

ARGO 2001 (7/4 – 10/13) 0.50 0.33 

ARGO 2002 (6/30 – 8/19) 0.50 0.16 

DINKUM 2000: (6/30 – 10/22) 0.71 0.16 

DINKUM 2001: (7/2 – 10/13) 0.54 0.31 

DINKUM 2002 (6/23 – 8/19) 0.52 0.09 

MCCLURE 2000: (6/30 – 10/22) 0.65 0.16 

MCCLURE 2001: 7/2 – 8/31 0.58 0.34 

REINDEER 2002: (6/27 – 8/19) 0.53 0.15 

Table 11. Open-water period wind-sea level correlations. 

(Italicized entries indicate statistical significance at the 

α < 0.05 level.) 

Mooring Time Period R: (η vs Ur) 

ARGO 2000: (6/30 – 10/20) 0.66 

ARGO 2001 (7/4 – 10/13) 0.45 

ARGO 2002 (6/30 – 8/19) 0.73 

DINKUM 2000: (6/30 – 10/22) 0.67 

DINKUM 2001: (7/2 – 10/13) 0.49 

DINKUM 2002 (6/23 – 8/19 0.69 

MCCLURE 2000: (6/30 – 10/22) 0.61 

MCCLURE 2001: 7/2 – 8/31 0.46 

REINDEER 2002: (6/27 – 8/19) 0.71 

Table 12. Wind – sea level correlations for all stations (Italicized entries indicate statistical 

significance at the α < 0.05 level.) 

OPEN WATER 2000: 6/30 – 10/22 2001: 7/2 – 10/13 2002: 6/27 – 8/19 

0.61 0.53 0.49 

LANDFAST ICE 10/15/99 – 6/30/00 10/22/00 – 7/2/01 10/13/01 – 6/23/02 

0.25 0.28 0.32 

In contrast to the open water season, there is no statistically significant correlation 

between along-shelf currents and τ

x

, τ

y

 (Table 13), or η (Table 14) during the landfast ice 
63 

period. On the other hand, τ

x

 and η are both significantly correlated at this time (second row of 

Table 12), although the magnitude of these correlations is about one-half that of the open water 

season. Correlations between along-shore currents and the zonal component of the geostrophic 

wind at 72.5o

N, 147.5oW from National Center for Environmental Prediction (NCEP) forecast 

models were also computed. These were non-significant during the landfast ice period, but 

significant (r = 0.7) during the open water period. The correlation between along-shore currents 

and the Deadhorse-Pt. Barrow atmospheric pressure difference was significant (r ~ .26) in both 

seasons, although the correlation explains <10% of the current variance. 

Table 13. Landfast ice season wind-current correlations. (None are statistically 

significant at the α < 0.05 level.) 

Mooring Time Period R: (τ

x

 vs Ur) R: (τ

y

 vs Ur) 

ARGO 99/10/15 – 00/6/30 -0.01 0.09 

ARGO 00/10/20 – 01/7/4 -0.06 0.21 

ARGO 01/10/13 – 02/6/30 0.11 0.15 

DINKUM 99/10/15 – 00/6/30 0.07 0.11 

DINKUM 00/10/20 – 01/7/4 -0.04 .21 

DINKUM 01/10/13 – 02/06/23 0.10 0.14 

MCCLURE 99/10/15 – 00/6/30 -0.06 0.05 

MCCLURE 00/10/20 – 01/7/4 -0.05 0.17 

REINDEER 01/10/10 – 02/6/27 0.12 0.14 

Table 14. Landfast ice season sea-level-current correlations. (None are statistically 

significant at the α < 0.05 level.) 

Mooring Time Period R 

ARGO 99/10/15 – 00/6/30 0.07 

ARGO 00/10/20 – 01/7/4 0.12 

ARGO 01/10/13 – 02/6/30 0.12 

DINKUM 99/10/15 – 00/6/30 0.15 

DINKUM 00/10/20 – 01/7/4 0.13 

DINKUM 01/10/13 – 

02/06/23 

0.09 

MCCLURE 99/10/15 – 00/6/30 -0.01 

MCCLURE 00/10/20 – 01/7/4 0.09 

REINDEER 01/10/10 – 02/6/27 0.07 
64 

The absence of a significant wind-current correlation contrasts with Aagaard’s [1984] 

suggestion that currents on the inner shelf were correlated with the winds during the landfast ice 

period. His conclusion, however, was based on visual inspection of wind and current records 

from two 3-week current meter records collected under the landfast ice in water depths of 27 and 

38 m, approximately 35 km northeast of the DINKUM mooring. There are two possible reasons 

for the difference. First, his measurements might have coincided with a period when winds and 

currents were roughly correlated and such periods times are evident upon inspection of our data 

set. On the other hand, Aagaard’s measurements were made much closer to the edge of the 

landfast ice, perhaps within one or two deformation radii (say within ~10 km), so that the 

offshore winds influenced the currents near but inshore of the ice edge at his locations, but 

perhaps not at locations further inshore. The latter is plausible based upon results from simple 

idealized numerical circulation models we have begun conducting of the landfast ice zone and 

offshore region for the Beaufort Sea. We find that the imposition of an alongshore wind stress 

over the region seaward of the landfast ice produces an along-ice edge jet whose inshore extent is 

<~10 km. We also find that the flow inshore of the jet reverses, although the nearshore flow is 

small (~1cm-s-1). The flow reversal occurs because the sea-level slope reverses at the ice edge 

due to the large wind stress curl imposed at the offshore boundary of the ice edge. This curl 

induces Ekman suction, which lowers or raises the sea level at the ice edge relative to the sealevel beneath the ice. Hence a current reversal develops beneath the landfast ice and inshore of 

the ice edge. Although our model results are preliminary (and assume no frictional coupling 

between the ice and the under-ice flow at this stage), they suggest a cross-shore reversal in the 

along-shelf flow across the landfast ice boundary could develop. 
65 

5. Dynamics 

We now examine the vertically integrated along-shore momentum balance at subtidal and 

seasonal time scales using data from the moorings and wind stresses as calculated above. The 

vertically integrated, linearized, along-shore momentum equation is: 

 + 

 (a) (b) (c) (d)

x x U s b g t xH H

∂ ∂η τ τ =− − ∂ ∂ρ ρ (1) 

where U is the vertically averaged along-shore velocity component, g the gravitational 

acceleration, H the water depth, ρ the density, and τ

x

 the stress in the along-shore direction. The 

subscripts “s” and “b” refer to the stress at the surface and bottom, respectively. The equation 

states that local accelerations in along-shore velocities (term a) are balanced by the along-shore 

pressure gradient arising from the along-shore sea level slope (term b) and the vertically 

integrated effects of surface (term c) and bottom stresses (term d). The bottom stress ( x

b τ ) is 

often parameterized by a linear stress law, x

b τ = ρrU, with r, the resistance coefficient being~10-4

m-s-1 [Csanady, 1981; Lentz, 1994]. We assumed that the non-linear terms in this force balance 

are small and we have neglected the effects of wave radiation stresses. The latter arise due to 

surface gravity waves propagating obliquely onshore and these stresses can be a substantial 

momentum source in shallow water (<5m) during storms [Lentz et al., 1999]. Although wave 

radiation stresses are absent during the landfast ice period, they might occasionally be important 

during the open water period. We believe that these stresses are likely small because the energy 

of onshore propagating waves is probably dissipated (to some extent) along the barrier islands 

and shoals girding the offshore boundary of Stefannsson Sound. In summer, x

s τ is the surface 

wind stress, whereas in winter it arises due to frictional coupling between the ocean and the 
66 

underside of the immobile landfast ice. Consequently, both the sea bottom and the underside of 

the ice exert a frictional stress against the along-shore flow during the landfast ice season, 

whereas during the open water season bottom and surface stresses oppose one another. For 

present purposes we assume that the underice stress can be parameterized in the same manner as 

the bottom stress, e.g. ( x x

b s τ =τ = ρrU) with r ~ 10-4 m-s-1 for both stress terms. For reasons to be 

discussed later, our choice of r for the resistance coefficient between ice and water might not be 

valid. However, in the absence of the necessary observations, we assume that it is of the same 

magnitude as the bottom resistance coefficient. With this assumption, equation 1 becomes: 

2 U rU

g t xH

∂ ∂η =− − ∂ ∂ (2)

We computed all terms in equations 1 and 2 except the along-shore sea level slope, which 

is estimated as the residual. The results are shown using data from DINKUM for the 2000 and 

2001 open water seasons (Figures 37 and 38) and for each landfast ice season (Figures 39 - 41). 

Results using other moorings are similar and not shown. 

In both seasons balance is primarily achieved among the stresses and along-shelf pressure 

gradient, with the local acceleration term ( U

t

∂

∂ ) typically being 10-6 or smaller and an order of 

magnitude smaller than the other terms. In both seasons the along-shelf pressure gradients imply 

along-shore sea level slopes of +10 cm/100 km. If these slopes extend the entire length of the 

Alaskan Beaufort coast then the sea level difference between Barrow and the US-Canada 

boundary can be as large as 0.5 m, which is easily detected with pressure gauges. The average 

along-shore velocity over the landfast ice period is virtually negligible implying that the there is 

no mean sea level slope during this season. In general, the along-shelf sea level gradients are 

about twice as large during the open water season than during the landfast ice period. We 
67 

emphasize that the estimates of the winter along-shore sea-level gradients are uncertain, 

however, because of uncertainty in the ice-water frictional coupling. 

Nevertheless, the presence of sub-tidal underice flows requires an along-shore sea level 

slope. How these slopes originate is not clear, however, although several possibilities exist. 

First, sea level fluctuations might be induced on the western shelf near Barrow due to variations 

in the coastal current that drains the Chukchi Sea through Barrow Canyon [Weingartner et al., 

1998; accepted]. If this were the primary forcing for the underice flows, we would expect a 

stronger correlation between the underice currents and the NCEP wind fields, since currents in 

the Chukchi Sea are strongly correlated with these winds [Weingartner et al., 1998; accepted]. 

We would also expect that the variance in current and along-shore sea level would be greater in 

late fall-early winter than in late winter as these authors found for the flow in Barrow Canyon. 

Second, fluctuating along-shore coastal sea-level gradients can be established by time-varying 

divergences in the along-shelf wind field. Third, spatial variations in the underice friction 

associated with the complex deformation field of the sea ice might establish cross-shelf flows 

that alter the along-shelf pressure field [Trowbridge et al., 1998]. Finally, remotely forced 

continental shelf waves could also affect the along-shore pressure gradient. These waves would 

be forced by wind stress variations to the west (as far west as the East Siberian Sea, for example) 

of our measurement site. This possibility cannot be discounted, but is beyond the scope of the 

present study. 
68 

Figure 37. Time series of the terms in the vertically integrated along-shelf momentum equation 

for the 2000 open water season using data from DINKUM. 
69 

Figure 38. Time series of the terms in the vertically integrated along-shelf momentum equation 

for the 2001 open water season using data from DINKUM. 
70 

Figure 39. Time series of the terms in the vertically integrated along-shelf momentum equation 

for the 2000 – 2001 landfast ice season using data from DINKUM. 
71 

Figure 40. Time series of the terms in the vertically integrated along-shelf momentum equation 

for the 2001 – 2002 landfast ice season using data from DINKUM. 
72 

Figure 41. Time series of the terms in the vertically integrated along-shelf momentum equation 

for the 2001 – 2002 landfast ice season using data from DINKUM. 
73 

6. Temperature, salinity, transmissivity, fluorescence 

Water property parameters also vary throughout the year in accordance with the 

development and retreat of sea ice. The annual cycles of temperature (T) and salinity (S) are 

shown for each mooring (with high quality data) in Figures 42 – 49. At each location the annual 

temperature cycle consists of maximum near-bottom temperatures that range from 5 – 2o

C in 

summer and fall, followed by a rapid (1 – 2 week) collapse to the freezing point (~-1.7o

C) 

(usually in early October), after which temperatures remain near-freezing until late June or early 

July. At that time, temperatures slowly increase and reach about 0o

C by late July after the ice 

melts. Salinity varies from 14 through 32 through the open water season, with the lowest 

salinities observed immediately following the decay of the landfast ice and in the aftermath of 

sufficiently strong winds that mix the water column. After the ice forms in October, salinities 

increase and attain values of 34 – 35 by January due to the expulsion of salt from growing sea 

ice. Thereafter, salinities remain relatively constant through winter and spring before slowly 

starting to decrease in June. Following the removal of ice and the first significant wind-mixing 

event, salinities rapidly decrease as a consequence of mixing of low-salinity ice meltwater and 

the river plume (discussed in Section V). 

Transmissivity also shows a strong seasonal cycle. During late summer and fall it varies 

due to stirring by the winds and currents, but generally decreases to very low values at about the 

time of rapid ice formation in early October. The low values are due to turbulent mixing of the 

water column brought about by winds and cooling and freezing that extend over the entire water 

column. The high suspended sediment levels in the water column during freeze-up suggest that 

much sediment is included into the landfast ice matrix during the vigorous production period. 

The incorporation of sediments into sea-ice has bearing on oil dispersal because the sea ice 
74 

represents a potential vehicle by which oiled sediments can be subsequently transported out of 

the region. However, that transport would be delayed until the following summer during and 

after break-up. Either the sediment-laden ice will be advected out of the region or will melt insitu and release its oiled sediments back into the water. Transmissivity values remain relatively 

high through winter although the data suggests periods of moderate suspended sediments 

followed by clearer water masses. Winter periods of moderately low transmissivity might, in 

fact, be a sampling artifact because the transmissometer was inclined to the horizontal on the 

mooring frame. Thus, sediment might have accumulated on the lens during periods of low flow 

and then be cleansed when current speeds increased. Although there is some uncertainty in the 

winter transmissivity record, there is a consistently large decrease in transmissivity in early June 

of each year, coincident with the spring freshet when rivers are carrying a heavy suspended 

sediment load. This load sinks to the seabed under the landfast ice because there is little energy 

available for mixing in the water column. Once the ice retreats, the near-bottom suspended 

sediment load remains relatively low until vigorous storms mix sediment back into the water 

column where it is advected by wind-driven currents. 

Fluorescence provides a qualitative measure of the chlorophyll content in the water. 

These values are always low except in late July or August after the ice has disappeared and 

transmissivity levels have risen above the post freshet minimum. 
75 

Figure 42. Ice thickness and bottom track velocity (upper panel) and temperature and salinity 

(lower panel) at ARGO during the 1999 – 2000 deployment. 
76 

Figure 43. Ice thickness and bottom track velocity (upper panel) and temperature and salinity 

(lower panel) at ARGO during the 2000 – 2001 deployment. 
77 

Figure 44. Ice thickness (upper panel), temperature and salinity (middle panel) and fluorescence 

(lower panel) at ARGO during the 2001 – 2002 deployment. 
78 

Figure 45. Ice thickness (upper panel), temperature and salinity (middle panel) and 

transmissivity (lower panel) at DINKUM during the 1999 - 2000 deployment. 
79 

Figure 46. Ice thickness (upper panel) and temperature and salinity (lower panel) at DINKUM 

during the 2000 - 20001 deployment. 
80 

Figure 47. Ice thickness (upper panel), temperature and salinity (middle panel) and 

transmissivity and flourescence (lower panel) at DINKUM during the 2001 - 2000 deployment. 
81 

Figure 48. Ice thickness and bottom-track velocity (upper panel) and temperature and salinity 

(lower panel) at MCCLURE during the 2000 - 2001 deployment. 
82 

Figure 49. Ice thickness and bottom-track velocity (upper panel), temperature and salinity 

(middle panel) and transmissivity at MCCLURE during the 2001 - 2002 deployment. 
83 

Figure 50 shows in detail the relationship between Sagavanirktok River discharge, 

transmissivity, ice thickness, currents and current shears from May through June from mooring 

DINKUM in 2000. Prior to the onset of discharge, all parameters are relatively constant and 

typical of winter conditions. However, once the freshet reaches its maximum on June 9, the 

transmissivity and ice thickness rapidly decrease, relatively large cross-shore flows (of up to 10 

cm-s-1) are initiated, and vertical shears increase. Rapid sea-ice ablation occurs due both to 

strong solar heating and to the rapid decrease in albedo of the ice surface by melt water ponds 

and/or the spreading of turbid river water over the surface of the ice Searcy et al. [1996]. 

Figure 50. May through June time series of (from top to bottom) Sagavanirktok River 

discharge, ice thickness (black), transmissivity (green), cross- and along-shore velocity shear, 

and cross- and along-shore velocities. Along-shore (cross-shore) components are ref (blue). 
84 

V. FRESHWATER INFLUENCE 

In this section we further explore the influence that freshwater runoff exerts on the inner 

shelf’s density structure and circulation field. During river breakup a significant fraction of the 

river runoff flows beneath the landfast ice where it establishes a strongly stratified water column, 

with salinity accounting for most of the stratification (Figure 51). Salinities increase by ~25 m-1

across the halocline, while temperature decreases by 1.5 o

C m-1 (Figure 52) and transmissivity 

by 80% m-1 (Figure 53). The inner shelf will remain highly stratified until sufficient turbulent 

energy is supplied to the water column to mix it vertically. This is evident from the bulk 

Richardson number: 

 ( ) ( ) ( ) 2 2 g Ri u v

zzz

ρ

ρ

∂ ⎡ ∂ ∂ ⎤ = + ∂∂∂ ⎢ ⎥ ⎣ ⎦

 (3) 

where 

z

∂ρ

∂ is the vertical density gradient and the denominator is the square of the vertical 

gradient of horizontal velocity. Ri ~ 325 for typical values of the gradients of underice shear and 

density observed during river runoff in summer. These values are substantially greater than the 

range of 1 - 10 typical of weak and moderately stratified shelf conditions and much larger than 

the value of 0.25 where mixing occurs. The large Richardson number also implies that 

interfacial stresses are small so that most of the momentum imparted by the wind stress, once the 

ice has receded will be confined to the plume. This has important implications for the 

subsequent spreading of the plume after the landfast ice retreats. For example, if we assume that 

the wind stress is confined to a strongly stratified 2 m thick plume, a weak but upwelling 

favorable wind speed of 2 m s-1 would transport the plume seaward at ~5 cm s-1 (~5 km day-1). 

Such winds are not unrealistic for the nearshore Beaufort Sea in summer where the mean winds 

are westward (and upwelling favorable) due to the sea breeze effect. 
85 

Figure 51. Salinity along the DINKUM 2001 transect (see Figure 3 for location of this transect). 

AN2 AN3 AN4 AN5 AN6 AN7

Ice Canopy

Sea Bottom

Salinity (psu) - (6/11/01)
86 

Figure 52. Temperature along the DINKUM 2001 transect (see Figure 3 for location of this 

transect). 

0 1.5 3 4.5 6 7.5

Distance (km)

-7

-6

-5

-4

-3

-2

-1

0

De

pth (

m)

-1.8

-1.6

-1.4

-1.2

-1.0

-0.8

-0.6

-0.4

-0.2

AN2 AN3 AN4 AN5 AN6 AN7

Ice Canopy

Sea Bottom

Temperature (deg. C) - (6/11/01)
87 

Figure 53. Transmissivity along the DINKUM 2001 transect (see Figure 3 for location of this 

transect). 

The offshore velocity of the plume is comparable to the 6 cm s-1 alongshore speed a surface oil 

slick would have assuming it moves at 2% of the wind speed. Although simplified, this example 

suggests that the effects of stratification and vertical mixing cannot be ignored in oil spill 

trajectory models for the nearshore Beaufort Sea in summer. 

How rapidly the stratification erodes depends upon the strength of the winds once the ice 

has retreated or becomes mobile, since wind is the primary agent for supplying mixing energy 

for the water column. We determine the work required to mix from the potential energy of the 

water column: 

0 1

h

M ( ( z ) )gzdz h −= ρ −ρ ∫ (4) 

0 1.5 3 4.5 6 7.5

Distance (km)

-7

-6

-5

-4

-3

-2

-1

0

Dept

h (

m)

-0

10

20

30

40

50

60

70

80

90

AN2 AN3 AN4 AN5 AN6 AN7

Ice Canopy

Sea Bottom

Transmissivity (%) - (6/11/01)
88 

where g is the gravitational acceleration, ρ(z) is the density at depth z, and ρ is the mean density 

of the water column. The winds are the sole source of mixing energy because the currents (tidal 

and subtidal) have insufficient shear to overcome the large density gradient. We estimate the 

time required to vertical mix the water column from Denman and Miyake’s [1973] relationship: 

 

3 E a D mC U10

t h

∂ ρ ≅

∂ (5) 

where ( E

t

∂

∂ ) is the rate of working by the wind on the water column, ρa is the air density (1.29 

kg m-3), m is an efficiency factor and CD is the drag coefficient (both ~ 10-3), h is the water depth 

(5.6 m for the CTD profiles shown in Figure 51) and 3 U10 is the cube of the wind speed at 10 m 

elevation. Dividing eq. 4 by eq. 5 gives the time required for a steady wind to mix the water 

column. Complete mixing is achieved in 3 (28) days for constant winds of 6 (3) m s-1. Thus the 

structure of the nearshore water column can vary considerably from year to year depending upon 

the seasonal evolution of the wind field. 

The cross-shore velocities associated with the spreading of the river plume shown in 

Figure 50 suggest that the plume spreads rapidly offshore. How far offshore does the plume 

propagate under the ice? Yankovsky and Chapman [1997] developed a scale-length for the 

offshore extent of a buoyant surface-advected plume for a steady outflow in the absence of 

surface friction: 

 ( )

( )

2

1 2 2

2 3

2

o i

s /

o i

g'h v

y

f g'h v

+ =

+

where s y is the offshore extent of the plume, g' g = ∆ρ ρo and i v is the inflow velocity, ho is the 

inflow depth, and f is the Coriolis parameter (1.37 x 10-4 s-1 at 71o

N). Although the steadystate, inviscid assumptions do not strictly apply to the impulsive-type discharge characteristic of 
89 

arctic rivers in early summer or for plumes under landfast ice (where friction might be 

important), we nevertheless apply this theory to early June when the peak outflow of the 

Kuparuk River is about 2000 m3

-s-1. We take the effective width of the river mouth to be 1 km 

and the inflow depth to be 2 m, so that vi ~ 1.1 m-s-1. From the salinity cross-section, we 

estimate ∆ρ ~ 22 kg-m-3, ρo ~ 1025 kg-m-3 so that g’ ~ 0.21 m2

- s-2. For these values the 

offshore extent of the plume is ~25 km and well beyond the region of our measurements. The 

plume might in fact extend further offshore because of frictional coupling between the plume and 

ice. Nevertheless, for the case considered, the propagation speed of a gravity current is 

2 o c g'h = [Benjamin, 1968], so that c ~ 0.9 m-s-1 for the assumed parameters. Hence the 

plume should propagate the distance ys in less than a day. Note also that Yankovsky and 

Chapman’s scaling assume no ambient mean flow. Our observations indicate that the spring 

freshet can occur during periods of variable along-shelf flow and these might alter the offshore 

distance that the buoyant river plume can spread [Yankovsky, in press]. 

Stratification also affects the velocity profile as shown by the mean open water season 

profiles from REINDEER (Figure 54 and 55). Similar to Figures 16 and 17, the means are 

computed separately for westward and eastward flow events and plotted versus the scaled depth 

to account for changes in water column depth. 

When the along-shore flow is eastward (Figure 54), there is little shear in the along- and 

cross-shore velocities. Eastward currents occur under eastward or downwelling favorable winds, 

which (at steady state and under an idealized two-dimensional case) cause offshore transport of 

low density water in the bottom boundary layer. As this water flows seaward it convectively 

mixes with denser surface water to rapidly erode stratification. The breakdown in stratification 

allows the efficient vertical transport of vertical momentum from the wind throughout the water 
90 

column leading to small velocity shears and, if the water column is sufficiently shallow, 

overlapping of the surface and bottom Ekman layers. Interestingly, we do not observe onshore 

transport in the surface layer and offshore transport in the bottom for eastward flow (left panel of 

Figure 54) for this downwelling case. There are several possible reasons for this. First, our 

simple averaging approach might mask bottom Ekman layers. Second, a sufficiently strong 

alongshore pressure gradient would force an onshore geostrophic flow that swamps the Ekman 

transport. Lastly, the shelf flow field is likely three-dimensional, so that our simple twodimensional analysis cannot be applied to this setting. 

For westward or upwelling favorable winds the mean flow is westward (right panel of 

Figure 55), but the velocity profile is highly sheared. The sheared structure is consistent with 

surface offshore transport of low-density water and onshore transport of denser water beneath the 

surface layer. In aggregate these effects tend to enhance stratification, which in turn, inhibits 

vertical mixing of momentum. Thus a shallow surface Ekman layer forms wherein the wind 

momentum is confined to the surface layer (e.g., above the pycnocline depth). Hence the alongshore flow is substantially greater at the surface than below the pycnocline. While we do not 

have measurements of the stratification throughout the open water period, the mean velocity 

shear in Figure 55 suggests that the depth of the pycnocline during upwelling conditions in July 

and August 2002 is at about 1 – 2 m. We note that the difference in speeds and velocity structure 

are not due to the winds being stronger to the west than to the east during the 2002 open water 

season. In fact, the mean wind stress is strongly eastward during this season (Figure 8) although 

these did not cause swifter surface velocities than those observed under the weaker westward 

wind stress. An understanding of the asymmetric response to upwelling and downwelling 
91 

favorable winds in the presence of stratification is extremely important for understanding the 

regional circulation field. 

Figure 54. Mean velocity profiles for the cross-shore (left) and alongshore (right) velocity 

components during eastward flow conditions at REINDEER for the open water season. 

Figure 55. Mean velocity profiles for the cross-shore (left) and alongshore (right) velocity 

components during westward flow conditions at REINDEER for the open water season. 

The character of the plume is expected to change seasonally, as shown schematically in 

Figure 56, from the shallow, strongly-stratified and surface-advected plume (red) of early 

summer to the bottom-advected plume type of late summer and fall (blue). The latter develops 
92 

in late summer and fall after mixing has destroyed the stratification and likely includes a swift 

alongshore geostrophic flow embedded within the front that can carry materials eastward along 

the Alaskan coast. Moreover the front effectively blocks cross-shelf transport and so materials 

shoreward of the front tend to remain trapped there unless the flow is interrupted by upwelling 

winds and/or frontal instabilities. 

COAST OFFSHORE 

Figure 56. Schematic of the freshwater plume types likely to form in summer and fall in the 

nearshore region of the Beaufort Sea. 

Instabilities have been observed in satellite images of the Alaskan Beaufort Sea (Figures 

57 and 58). Although not directly wind-driven, the frontal structure and hence the 

characteristics of the instability do depend upon the seasonal wind history as these affect the 

frontal structure through mixing. The images show tongues of turbid water extending from near 

the coast to the shelfbreak and beyond (in some cases more than 100 km offshore). The flow in 

these unstable plumes has not been measured in the Alaskan Beaufort Sea. However, 

Weingartner et al. [1999] sampled similar features associated with the buoyancy-forced Siberian 

Coastal Current in the Chukchi Sea and found cross-shore velocities of up to 30 cm-s-1. Thus, 

instabilities could rapidly transport materials and pollutants from the nearshore Beaufort Sea 

Low-salinity water 

High-salinity water 
93 

offshore and possibly, based on these images, across the shelfbreak and along the major fall 

migration corridor for bowhead whales. 

Figure 57. Beaufort Sea SeaWIFS imagery from August 25, 2002 (upper panel) and September 

5, 2002 (lower panel). Winds were weak and variable for the week preceding these images. 
94 

Figure 58. Visible SeaWIFS image from September 4, 2004 showing turbid plumes extending 

seaward from the Alaskan Beaufort Sea shelf across the continental slope. 

150 km

Mackenzie River Delta 

Turbid Plumes 
95 

VI. CONCLUSIONS AND RECOMMENDATIONS 

Three years of current meter and water property measurements were made in the 

nearshore region of the Alaskan Beaufort Sea to assess circulation dynamics of the region 

beneath the landfast ice. The measurements, made within Stefannsson Sound show a distinct 

seasonal cycle associated with the formation and ablation of the landfast ice. 

The mean flow, whether averaged over the entire record or by season, is small and in 

seldom significantly different from zero. It is, however, highly variable in time with the 

dominant mode of variability being in the along-shore direction. During the open water season 

the currents are swift (10 – 50 cm-s-1), strongly sheared, especially when the flow is westward, 

and significantly correlated with the winds. During the landfast ice season currents are small 

(generally ~5 cm-s-1), weakly sheared, and uncorrelated with winds. Nevertheless, the winter 

sub-tidal flow variance implies fluctuating along-shore sea level gradients of O(10-6). The origin 

of these sea-level slopes is not known. Measurements now being made under MMS support are 

focused on directly measuring the along-shore pressure and velocity fields in order to identify the 

possible origins of these gradients and to investigate the regional momentum balance. We 

further recommend that simple idealized numerical model studies be conducted to examine the 

origins of the underice forcing and its affects on the underice flow field. 

Freshwater discharge associated with the springtime freshet creates shallow, strongly 

stratified, underice plumes that likely spread up to 20 km or more offshore. The cross-shore 

flows associated with these plume can be as large as 10 cm-s-1 and are much larger than the 

cross-shore flows generally observed beneath the landfast ice in winter. Little is known about 

these plumes, although they provide a vehicle by which nearshore suspended and dissolved 

materials can be carried offshore by spreading of the plume beneath the ice, offshore Ekman 
96 

transport (once the ice breaks up), or through frontal instabilities. The latter generate large crossshelf plumes that can extend across the Beaufort shelf and slope. Measurements obtained from 

similar features in the Chukchi Sea indicate that cross-shelf velocities associated with these 

instabilities can be ~25 cm-s-1. We recommend that measurements be conducted to quantify the 

flow within these cross-shore jets and to determine the mechanisms by which they are generated. 

MMS is currently supporting nearshore measurements using shore-based, high-frequency radar. 

We urge that shipboard sampling be made in conjunction with the radar measurements to 

identify the characteristics (stratification, plume thickness) of these jets and the mechanisms that 

generate them. 

Stratification due to freshwater inflow leads to a strong asymmetric response in the 

velocity shear between upwelling (westward winds) and downwelling (eastward) conditions on 

the inner shelf. Westward surface flows are intensified during upwelling events because the 

wind’s momentum is trapped to a strongly stratified surface layer, while eastward surface 

currents are weaker during downwelling because stratification is weaker and the momentum 

from the surface stress is mixed over a deeper layer. The results imply that circulation models 

must correctly incorporate runoff and stratification in order to reproduce the surface circulation 

field correctly. Observations on the seasonal evolution of the open water stratification structure 

are needed to better understand this asymmetry and for model evaluations. 

The seasonal cycle in sediment transport likely consists of rapid deposition from the 

freshwater plume as it spreads beneath the sea ice followed by several re-suspension events 

during the open water season. Upon re-suspension, sediments can be advected offshore within 

the coastal flows or carried offshore due to instabilities. Re-suspension also appears to be 

prominent during freeze-up and we presume that much sediment is incorporated into the landfast 
97 

ice at this time. This sediment will then remain in the landfast ice matrix until the ice break-ups 

and drifts away the following summer or will drop out during ice melt. 

Our data suggest that oil is unlikely to be carried far in the event of a spill beneath the 

landfast ice. Instead the oil will be transported back and forth along the coast. If the Cox and 

Schmidt [1980] laboratory measurements apply to this region of the Beaufort Sea, then oil will 

not move far once it has attained its equilibrium slick thickness. In the event of an underice spill 

here, the rate and direction of the spreading oil can be monitored because the current field is 

spatially coherent over at least 30 km in the alongshore direction. Thus a single current meter 

can be lowered through a hole in the ice and transmit current data in real-time to the spill 

recovery team. Direct measurements are required because of the absence of a significant windcurrent correlation in winter. 

Oil spilled beneath the ice during the spring freshet could be carried offshore of 

Stefansson Sound within the underice river plume. Application of theories developed for midlatitude settings on the offshore extent of a river plume discharged into the sea suggests that oil 

might be carried at least 20 km offshore during the spring freshet. These theories do not consider 

the possibly complex frictional coupling between the ice and flow field or the impulsive nature 

of arctic river discharges. We recommend that theoretical and observational studies be 

conducted in the Alaskan Beaufort Sea directed at understanding the underice spreading of river 

plumes. We believe that the current measuring techniques developed in this project can be 

applied to such a study. 

Oil adsorbed onto sediments provides another mechanism for the dispersal of this 

pollutant. Some of the sediments are transported by the currents and some by the sea ice, 

although it is not known how the load is partitioned between the two. The sediment load in ice 
98 

can be easily determined from ice cores collected in spring prior to break-up. The fate of the 

landfast ice can then be determined by deploying satellite tracked drifters in the ice prior to 

break-up. Once the ice begins to break-up its drift and subsequent melt can be ascertained with 

the aid of the drifters. 

The complex deformation field of the landfast ice regime suggests that frictional coupling 

between ice and currents will vary substantially over the shelf covered by landfast ice due to both 

skin and form drag, which are functions of both the underice topography and flow field 

[McPhee, 1990]. Measurements and models are required to quantify this frictional coupling. 

Mapping of underice topography should be conducted several times per winter and cover a 

variety of along- and cross-shore spatial scales. The larger scales can be mapped efficiently 

using airborne electromagnetic sensors or laser profilometry, whereas smaller scales will require 

ground-based measurements. 

Based upon preliminary numerical modeling activities, it appears that there is little 

exchange between waters beneath the landfast ice and those offshore. This topic needs further 

exploration using models and observations. One integrated observational approach that is 

relatively simple to undertake is to measure the δ18O fraction in ice cores in spring after the ice 

reaches its maximum thickness. Since this isotope ratio is substantially different between seawater and river water, spring ice cores will precisely record when the river water was exhausted 

from the nearshore region [Macdonald et al., 1999b]. This will lead to a distinct horizon in the 

ice core of the isotope ratio transition. The timing of the transition can be determined by 

calculating ice growth rate from coastal meteorological data. This is a relatively inexpensive 

procedure that would provide a measure of year-to-year differences in the rate of freshwater 

depletion from the nearshore Beaufort Sea after the landfast ice forms. 
99 

ACKNOWLEDGEMENTS 

We are grateful for the financial support provided by the Minerals Management Service, the 

equipment provided by the Alaska Department of Environmental Conservation, and logistical 

support from BP, Inc. We thank Dave Leech for his expertise in designing and deploying the 

moorings and for his cheerful approach throughout the project. John Trefey assisted with the 

2001 CTD measurements. G. Michael Schmidt kindly processed the SeaWIFS images. 
100 

REFERENCES 

Aagaard, K., The Beaufort Undercurrent. IN: The Alaskan Beaufort Sea: Ecosystems and 

Environment, edited by P. W. Barnes, D. M. Schell, and E. Reimnitz, pp. 47-71. Academic 

Press, New York, 1984. 

Aagaard, K., and A.T. Roach, Arctic ocean-shelf exchange: Measurements in Barrow Canyon, J. 

Geophys. Res., 95, 18163-18175, 1990. 

Barnes, P. W., D. M. Rearic, and E. Reimnitz, Ice gouging characteristics and processes IN: The 

Alaskan Beaufort Sea: Ecosystems and Environment, edited by P. W. Barnes, D. M. Schell, 

and E. Reimnitz, pp. 185 - 212. Academic Press, New York, 1984. 

Barnes and Reimnitz, 1974 IN The Coast and Shelf of the Beaufort Sea (JC Reed and J E Sater, 

eds.) p. 439, Arctic Institute of North America, Arlington, VA. 1974 

Benjamin, T. B., Gravity currents and related phenomena, J. Fluid Mech., 31, (2), 209 – 248, 

1968 

Brower, W. A., Jr., R. G. Baldwin, C. N. Williams, Jr., J. L .Wise, and L.D. Leslie, Climate atlas 

of the outer continental shelf waters and coastal regions of Alaska, volume III, ChukchiBeaufort Sea, 497 pp., National Climatic Data Center, Asheville, NC, 28801, 1988. 

Carmack, E. C., R. W. Macdonald, and J. E. Papdakis, Water mass structure and boundaries in 

the Mackenzie Shelf Estuary, J. Geophys. Res., 94, 18043-18055, 1989. 

Chapman, D. C. and S. J. Lentz, Trapping of a coastal density front by the bottom boundary 

layer, J. Phys. Oceanogr., 24, 1464-1479, 1994. 

Colonell, J. M. and B. J. Galloway, Wind-driven transport and dispersion of age-0 arctic ciscoes 

along the Alaska Beaufort coast. IN: Fish Ecology in Arctic North America, American 

Fisheries Society Symposium 19, Bethesda, MD., p – 90 – 103, 1997. 

Cox, J. C. and I. A. Schultz, The transport and behavior of spilled oil under ice. Proceedings of 

the Third Annual Technical Seminar, Arctic Marine Oil Spill Program, Edmonton, June 

1979, p. 45 – 61, 1980.

Csanady, G. T. Circulation in the Coastal Ocean, D. Reidel, Boston, p. 279, 1982. 

Danielson, S. L. and Z. Kowalik, Tidal currents in the St. Lawrence Island region. J. Geophys. 

Res., in press. 

Dean, K. G., W. Stringer, K. Ahlnaes, S.C. Searcy, and T. Weingartner, The influence of river 

discharge on the thawing of sea ice: Mackenzie River Delta: Albedo and temperature 

analysis, Polar Res., 13, 83-94, 1994. 

Dunton, K. H., E. Reimnitz, and S. V. Schonberg, An arctic kelp community in the Alaskan 

Beaufort Sea, Arctic, 35: 465 – 484., 1982. 

Denman, K. L. and M. Miyake, Upper layer modification at ocean station Papa: observations 

and simulation, J. Phys. Oceanogr., 3, 185 – 196, 1973. 

Guay, C. K. and K. K. Falkner, A survey of dissolved barium in the estuaries of major Arctic 

rivers and adjacent seas. Cont. Shelf Res., 8: 859 – 882, 1998. 
101 

Foreman, M. G.G. Manual for Tidal Currents Analysis and Prediction. Pacific Marine Science 

Report 78-6, Institute of Ocean Sciences, Patricia Bay, Sidney, B.C., 1978, 70 pp. 

Furey, P., The large-scale surface wind field over the western Arctic Ocean, 1981 – 1993, MS 

thesis, 121 pp. Univ. of Alaska, Fairbanks, 1998 

Kowalik, Z. and A. Y. Proshutinsky, The Arctic Ocean Tides, in The Polar Oceans and Their 

Role in Shaping the Global Environment; Geophys. Monogr. 85, pp. 137-158, edited by O. 

M. Johannessen, R. D. Muench and J. E. Overland, Amer. Geophys. Union, Washington, 

DC, 1994. 

Kozo, T. L., An observational study of sea breezes along the Alaska Beaufort Sea coast, J. Appl. 

Met. 21: 891- 905, 1982a. 

Kozo, T. L., An mathematical model of sea breezes along the Alaska Beaufort Sea coast, J. 

Appl. Met. 21: 906 - 924, 1982b. 

Kozo, T. L. Mesoscale wind phenomena along the Alaskan Beaufort Sea coast. IN: The Alaskan 

Beaufort Sea: Ecosystems and Environment, edited by P. Barnes and E. Reimnitz, pp. 23 - 

45. Academic Press, New York, 1984. 

Kozo, T. L. Mountain barrier baroclinicity effects on surface winds along the Alaskan Arctic 

coast. Geophys. Res. Lett., 7, 377 – 380, 1980. 

Lentz, S. J. Current dynamics over the northern California inner shelf, J. Phys. Oceanogr., 24, 

2461 – 2478, 1994. 

Lentz, S., R. T. Guza, S. Elgar, F. Fedderesen, and T.H. C. Herbers, Momentum balances on the 

North Carolina inner shelf, J. Geophys. Res., 104, 18205 – 18226, 1999. 

Macdonald, R.W., E.C. Carmack, F. A. McLaughlin, K. K. Falkner, and J. H. Swift, Connections 

among ice, runoff, and atmospheric forcing in the Beaufort Gyre, Geophys. Res. Lett., 

26:15:2223 – 2226, 1999a. 

MacDonald, R. W., E. C. Carmack, and D. W. Paton. Using the delta 18O composition in 

landfast ice as a record of arctic estuarine processes. Mar. Chem. 65: 3 – 24, 1999b. 

Macdonald, R.W., E.C. Carmack, F. A. McLaughlin, K. Iseki, D.M. Macdonald, and M. C. 

O’Brien, Composition and Modification of water masses in the Mackenzie Shelf Estuary, J. 

Geophys. Res., 94, 18057-18070, 1989. 

Macdonald, R. W. and E. C. Carmack, The role of large-scale under-ice topography in separating 

estuary and ocean on an arctic shelf. Atmosphere-Ocean, 29: 37 -51, 1991. 

Matthews, J.B., Observations of Under-ice circulation in a shallow lagoon in the Alaskan 

Beaufort Sea, Ocean Management, 6, 223-234, 1981. 

Maslanik, J. A., M. A. Serreze, and T. Agnew, On the record reduction in 1998 Western Arctic 

Sea ice cover, Geophys. Res. Lett., 26, 1905 – 1908, 1999. 

Maykut, G. A. The surface heat and mass balance, IN: The Geophysics of Sea Ice, edited by N. 

Untersteiner, pp. 395 -464, Plenum Press, New York, NY, 1986. 

McPhee, M. G., Small-Scale Processes. IN: Polar Oceaongraphy Part A: Physical Science, 

edited by W. O. Smith, Jr., pp. 287 - 280. Academic Press, New York, 1990. 
102 

Morris, K., S. Li, and M. Jefferies, Meso- and microscale sea-ice motion in the East Siberian Sea 

as determined from ERS-1 SAR data. J. Glaciol., 45(15): 370 – 383, 1999. 

Mountain, D. G., Coachman, L. K., and K. Aagaard, On the flow through Barrow Canyon, J. 

Phys. Oceanogr. 6, 461 – 470, 1976. 

North, G. R., T. L. Bell, and R. F. Cahalan, Sampling Errors in the estimation of empirical 

orthogonal functions, Monthly Weath. Rev., 110, 699-706, 1982 

Pawlowicz, R., B. Beardsley, and S. Lentz. Classical tidal harmonic analysis including error 

estimates in MATLAB using T.TIDE, Comp. and Geosci., 28, 929- 937, 2002. 

Pickart, R. S. Shelbreak circulation in the Alaskan Beaufort Sea: Mean structure and variability, 

J. Geophys. Res., 2004. 

Pickart, R.S., T. Weingartner, L.J. Pratt, S. Zimmermann, and D. J. Torres, Flow of wintertransformed Pacific water into the western Arctic (accepted, Deep-Sea Research) 

Reimnitz, E., Interaction of river discharge with sea ice in proximity of Arctic deltas: a review, 

Polarforschung, 70: 123 – 134, 2000. 

Reimnitz, E. and E. W. Kempema, Pack ice interaction with stamukhi shoal. IN: The Alaskan 

Beaufort Sea: Ecosystems and Environment, edited by P. W. Barnes, D. M. Schell, and E. 

Reimnitz, pp. 159 - 184. Academic Press, New York, 1984. 

Searcy, C., K. Dean, and W. Stringer, A river-coastal sea ice interaction model: Mackenzie River 

Delta, J. Geophys. Res., 101, 8885-8894, 1996. 

Tucker III, W. B., W. F. Weeks, and M. Frank, Sea Ice ridging over the Alaskan continental 

shelf, J. Geophys. Res., 84, 4885 – 4897, 1979. 

Weingartner, T., K. Aagaard, R. Woodgate, S. Danielson, Y. Sasaki, D. Cavalieri, Circulation on 

the North Central Chukchi Sea Shelf (accepted, Deep-Sea Research) 

Weingartner, T.J., S. Danielson, Y. Sasaki, V. Pavlov, and M. Kulakov, The Siberian Coastal 

Current: A wind- and buoyancy-forced arctic coastal current, J. Geophys. Res., 104, 29697-

29713, 1999. 

Weingartner, T.J., D.J. Cavalieri, K. Aagaard, and Y. Sasaki, Circulation, dense water formation, 

and outflow on the northeast Chukchi shelf, J. Geophys. Res., 103, 7647-7661, 1998. 

Wiseman et al., W.J., Suhayda, J.N., Hus, S.A., and Walters, C.D., IN The Coast and Shelf of the 

Beaufort Sea (JC Reed and J E Sater, eds.) p. 49, Arctic Institute of North America, 

Arlington, VA. 1974 

Woodgate, R. A., K. Aagaard, and T. Weingartner. A year in the physical oceanography of the 

Chukchi Sea: Moored measurements from autumn 1990-91. (accepted, Deep-Sea Research). 

Yankovsky, A. Interaction of transient shelf currents with a buoyancy-driven coastal current. J. 

Geophys. Res., (in press). 

Yankovsky, A. E. and D. C. Chapman, A simple theory for the fate of buoyant coastal 

discharges, J. Phys. Oceanogr., 27, 1386-1401, 1997. 
103 

The Department of the Interior Mission

As the Nation's principal conservation agency, the Department of the Interior has responsibility for most of our 

nationally owned public lands and natural resources. This includes fostering sound use of our land and water 

resources; protecting our fish, wildlife, and biological diversity; preserving the environmental and cultural values of 

our national parks and historical places; and providing for the enjoyment of life through outdoor recreation. The 

Department assesses our energy and mineral resources and works to ensure that their development is in the best 

interests of all our people by encouraging stewardship and citizen participation in their care. The Department also 

has a major responsibility for American Indian reservation communities and for people who live in island territories 

under U.S. administration. 

The Minerals Management Service Mission 

As a bureau of the Department of the Interior, the Minerals Management Service's (MMS) primary responsibilities 

are to manage the mineral resources located on the Nation's Outer Continental Shelf (OCS), collect revenue from 

the Federal OCS and onshore Federal and Indian lands, and distribute those revenues. 

Moreover, in working to meet its responsibilities, the Offshore Minerals Management Program administers the 

OCS competitive leasing program and oversees the safe and environmentally sound exploration and production of 

our Nation's offshore natural gas, oil and other mineral resources. The MMS Royalty Management Program

meets its responsibilities by ensuring the efficient, timely and accurate collection and disbursement of revenue from 

mineral leasing and production due to Indian tribes and allottees, States and the U.S. Treasury. 

The MMS strives to fulfill its responsibilities through the general guiding principles of: (1) being responsive to the 

public's concerns and interests by maintaining a dialogue with all potentially affected parties and (2) carrying out its 

programs with an emphasis on working to enhance the quality of life for all Americans by lending MMS assistance 

and expertise to economic development and environmental protection.