Bottom friction and bed forms on the southern flank of Georges Bank

Transcription

1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 108, NO. C11, 8004, doi: /2000jc000692, 2003 Bottom friction and bed forms on the southern flank of Georges Bank Sandra R. Werner, Robert C. Beardsley, and Albert J. Williams III Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA Received 20 October 2000; revised 25 June 2001; accepted 25 June 2001; published 18 October [1] We present observational estimates of bottom stress, bottom roughness, and quadratic drag coefficient for the southern flank of Georges Bank, a shallow region dominated by strong semidiurnal tides. Our estimates are based on near-bottom velocity measurements from two Benthic Acoustic Stress Sensor (BASS) tripod deployments at a depth of 76 m taken during winter (February April 1995) and summer (July August 1995). The dominant tidal constituent is the M 2, which provides 76 89% of the total kinetic energy. Typical bottom friction velocities are cm s 1 in the time mean, with standard deviations of cm s 1. A representative drag coefficient is c D = 3.0 ± at 1.2 m above the bottom. Our drag coefficient estimates at this and other elevations correspond to the apparent bottom roughness range z 0 = cm. Bottom photographs (taken February June) show that from February until April the sea floor was covered by sand ripples 1 2 cm in height. These ripples were approximately aligned with the local isobath; their flanks were perpendicular to the major axis of the tidal ellipse. Ripples exhibited stationary behavior; that is, they did not migrate with the flow. Intermittent modification of the ripple pattern happened during times of strong wavecurrent interaction, but along-isobath ripples formed again after each wave event. Observations of bed forms affected by surface waves coincided with periods of waveenhanced bottom stress predicted by the Grant and Madsen [1979] combined wavecurrent interaction model. Model results and bottom photographs indicate that sediments were kept in suspension only under extreme conditions, here the February 1995 storm. Sand ripples gradually eroded during late March and April until the sea floor was almost flat. Coinciding with slow ripple erosion, bioturbation levels increased. Our estimates of bottom stress do not exhibit seasonal changes, so that seasonal variation of bottom shear stress cannot explain the transition from a rippled to a nearly flat bottom. We conclude that biogenic modification of bottom sediment enhanced the critical shear stress for initiation of sediment movement, thereby prohibiting ripple formation in spring and summer. INDEX TERMS: 4219 Oceanography: General: Continental shelf processes; 4211 Oceanography: General: Benthic boundary layers; 4558 Oceanography: Physical: Sediment transport; 4560 Oceanography: Physical: Surface waves and tides (1255); KEYWORDS: bottom stress, bottom roughness, bottom drag coefficient, wave-current interaction, sand waves, Georges Bank Citation: Werner, S. R., R. C. Beardsley, and A. J. Williams III, Bottom friction and bed forms on the southern flank of Georges Bank, J. Geophys. Res., 108(C11), 8004, doi: /2000jc000692, Introduction [2] Tidal currents represent an important component of the near-bottom flow in many continental shelf and estuarine regions. Strong tidal flow over a hydrodynamically fully rough bottom can generate much of the turbulence found near the lower boundary, with direct consequences for the structure and evolution of the bottom boundary layer and the effective momentum flux (shear stress) felt by the surfacial bottom sediment and overlying fluid. The log layer found in the immediate vicinity of the bottom is Copyright 2003 by the American Geophysical Union /03/2000JC characterized by the bottom shear stress and bottom roughness. Depending on the bottom sediment composition and geometry of bottom bed forms, the flow-induced shear stress can be large enough to induce sediment motion and keep sediment in resuspension. The potential for feedback between bed form evolution and near-bottom flow plus the difficulty in making detailed measurements very close to the bottom make the boundary layer at the sediment/water interface a topic of ongoing research. [3] Bottom bed forms, sediment motion, shear stress, and bottom roughness can vary on a wide range of temporal scales under typical shelf conditions. Tidal variability of current strength introduces temporally varying bottom stress, which may alter bed form geometry throughout the GLO 5-1

2 GLO 5-2 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS course of a tidal cycle. Shelf bottom storm events are frequently characterized by strong wave-induced oscillatory currents, which can enhance the bottom shear stress and apparent roughness and contribute to sediment motion [Grant and Madsen, 1979]. On longer timescales, especially seasonal, changes in ambient stratification, subtidal circulation, or the level of biogenic activity at the sediment/water interface can all influence the bottom boundary dynamics and bed form structure. [4] As part of the U.S. GLOBEC Northwest Atlantic/ Georges Bank 1995 Stratification Study, we examine frictional processes and bottom bed forms at a shallow site on the southern flank of Georges Bank (Figure 1). For our analysis we use bottom tripod data (including bottom photographs) taken on the 76-m isobath during winter (February April 1995) and summer (July August 1995), plus additional measurements collected on nearby surface and subsurface moorings. Among the objectives of this study are the derivation of bottom roughness and bottom stress estimates which can be used as input parameters in numerical models, and which aid the understanding of the overlying flow dynamics. Questions of scientific interest concern the seasonal dependence of bottom roughness and bottom stress, the physical processes leading to bed form formation and the temporal variability of bed form geometry on seasonal, tidal, and event scales. To answer these questions, we use observations as well as the numerical Grant and Madsen [1979] combined wave-current interaction model. Taken together, we attempt to present here a conceptual picture of the near-bottom flow and bed form variability at a typical midshelf site on the southern flank of Georges Bank. 2. Measurement Site and Instrumentation 2.1. Location [5] Data were collected at N, Wonthe 76-m isobath of the southern flank of Georges Bank (Figure 1). The site-label Stratification Site 1 (ST1) refers to the fact that the water column was stratified during spring and summer. ST1 was located approximately at mid-distance between two fronts: the tidal mixing front (TMF) and the shelf slope front (SSF). The TMF surrounds the Bank near the 60-m isobath about 25 km north of ST1 and encircles a water mass that is distinguished in all seasons by its vertical and horizontal homogeneity [Hopkins and Garfield, 1981]. Near the 100-m isobath about 25 km south of ST1 lies the base of the SSF, which marks the boundary from fresh shelf water to saline upper slope water. The southern flank constitutes the region between the TMF and SSF and is nearly unstratified in winter and stratified in summer. A characteristic bottom slope is Surfacial Geology [6] Due to large spatial gradients of current strength and associated bottom stress, the surfacial geology of Georges Bank varies strongly over the Bank. Inside the TMF, tidal currents on the order of 100 cm s 1 create movable sand waves that reach heights of 3 25 m [Twichell et al., 1987]. Bottom sediments on the crest of the Bank are mostly coarse sand; the finer sediments have been eroded and moved offshore. Outside the 60-m isobath on the southern flank, Figure 1. Bathymetry (in meters) of Georges Bank and adjacent region, approximate location of the tidal mixing front and shelf-slope front and the GLOBEC Stratification Study mooring site ST1. The on-bank (+x) and along-bank (+y) axes are oriented toward 30 and 120 (counterclockwise) from true north, respectively, in agreement with Brown [1984]. ST1 is located about 25 km to the north of the shelf-slope front intersecting Georges Bank near the 100-m isobath and about 25 km to the south of the tidal mixing front surrounding the bank near the 60-m isobath. the offshore decrease in tidal current amplitude and bottom stress limits sediment movement. Butman [1987] showed bottom photographs from the 64-m isobath featuring sand ripples of a few centimeters height. [7] Surfacial sediment maps indicate the median sand grain size decreases across the southern flank from north to south and along the flank from east to west [Twichell et al., 1987]. ST1 was located in the transition zone between predominantly medium-to-coarse sand to the east and fine sand to the west. Sediment samples were not available for the deployment site, so that we rely on sediment analysis from nearby locations to infer grain diameter and sediment composition. Maciolek and Grassle [1987] analyzed sediment composition at several locations on the southern flank, including two positions on the 80-m isobath about 40 km east and west of ST1. Their samples from 80-m water depth consisted of less than 1% clay and silt, about 8 17% fine sand, 30 63% medium sand, 19 44% coarse sand and 1 17% very coarse sand and gravel. These results indicate that surface sediments at ST1 were mostly mediumto-coarse sand with medium grain diameter mm Instrumentation [8] In 1995, two Benthic Acoustic Stress Sensor (BASS) tripods were deployed at ST1 during conditions representative of winter and summer. Both tripods featured five acoustic travel time current meters [Williams et al., 1987] and eight temperature sensors (Figure 2). The first deployment (BASS 1) returned good velocity and temperature data for 3 February to 14 April 1995, and the second (BASS 2) for 11 July to 14 August The exact positions of the tripod instrumentation and the time and duration of the measurement and analysis periods are listed in Table 1. Temperature and acoustic travel time were recorded at 2 Hz during 7.5-min long bursts which occurred every half hour.

3 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS GLO 5-3 Figure 2. Schematic of the GLOBEC BASS winter tripod. Camera, light strobe, and battery case were not mounted on the summer tripod. Data from optical backscatter sensors (OBS) are not used in this study due to sensor calibration problems. In this study, burst-averaged data are used unless mentioned otherwise. Processing of BASS velocity data followed the procedure described in Appendix A. [9] A camera was mounted on the winter tripod which took one picture of the seafloor every eight hours and delivered good quality bottom photographs through the end of May. No camera was attached to the summer tripod. [10] Also used here are temperature and current measurements from Vector Measuring Current Meters (VMCMs) at 6, 12, 30.5, 39, 45, 57, 62, 66, 68.5, and 71 m above the bottom (sample rate 7.5 min), and SeaCAT temperature and conductivity measurements at 11 m above the bottom (sample rate 1.5 min). These measurements were taken 1 February to 23 August 1995, at the ST1 surface and subsurface moorings located within 100 m and 350 m of the BASS deployment site, respectively. A Vector-Averaging Wind Recorder (VAWR, sample rate 15 min) and an Improved Meteorological Recorder (IMET, sample rate 1.0 min) recorded meteorological data; both instruments were mounted on the 3-m discus buoy of the surface mooring. See Alessi et al. [2001] for additional information about the 1995 stratification study moored array instrumentation. 3. Flow Field [11] BASS current measurements indicate that the semidiurnal M 2 tide (period h) carried 76 89% of the kinetic energy near the bottom (Table 2). Tidal decomposi- Table 1. Measurement and Analysis Periods for the Winter (BASS 1) and Summer (BASS 2) Tripod Deployments and Sensor Elevations Above the Bottom a BASS 1 BASS 2 Velocity Temperature Velocity Temperature Measurement period 3 Feb. to 25 Apr. 3 Feb. to 25 Apr. 11 Jul. to 14 Aug. 11 Jul. to 24 Aug. Analysis period 3 Feb. to 14 Apr. 3 Feb. to 14 Apr. 11 Jul. to 14 Aug. 11 Jul. to 14 Aug. Sensor height, 0.22, 0.58, 1.18, (0.24), 0.62, 1.22, 0.26, 0.61, 1.21, 0.32, 0.58, 1.18, meters above 2.53, , 2.53, 3.24, 2.56, , 2.53, 3.48, bottom 4.43, , 6.02 a Sensor elevations correspond to the heights above ship deck minus 2 cm to account for sinking of the tripod legs into the sea floor. The last 10 days of the winter deployment (15-25 April) were excluded from analysis since remnants of a plastic bag had wrapped around the lowest two current sensors. Numbers in parentheses indicate instruments with no data return.

4 GLO 5-4 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS Table 2. Kinetic Energy Contribution of the Largest Diurnal and Semidiurnal Tidal Constituents, High-frequency Components, and Subtidal Flow to the Total Kinetic Energy at ST1 a Tidal Constituent BASS 1 3 Feb. to 14 Apr. (71 d) Percent of Total Kinetic Energy BASS 2 11 Jul. to 14 Aug. (35 d) VMCM 1 Feb. to 22 Aug. (203 d) S 2 (12.00 h) M 2 (12.42 h) N 2 (12.66 h) K 1 (23.93 h) O 1 (25.82 h) High frequency (<12 h) Subtidal (>33 h) Total KE, cm 2 s a Results are shown for depth-averaged BASS measurements at current sensors 1 5 during winter (BASS1) and summer (BASS2), and the continuous (winter through summer) time series of depth-averaged VMCM measurements. Note that column 3 is based on currents averages over the entire water depth. tion was performed on hourly-averaged velocity data using Godin s harmonic method [Foreman, 1978]. Differences in kinetic energy contribution of the semidiurnal constituents between BASS 1 and BASS 2 arise from two factors: (i) lowfrequency modulation of the S 2 (12.00 h) and N 2 (12.66 h) by unresolved semidiurnal constituents such as K 2 (11.97 h) and NU 2 (12.62 h), respectively; and (ii) superposition of internal and barotropic tides during the stratified season. The effects of internal tide motion on current amplitudes vanish if tidal decomposition is performed using depth-averaged VMCM measurements (Table 2, last column). [12] The orientation of the major axis of the M 2 tidal ellipse was approximately across-isobath (Figure 3; Appendix A). Typical current amplitudes along the major and minor axes Figure 3. (top) M 2 current ellipse from tidal decomposition of the depth-averaged currents measured by BASS sensors 1 5 and (bottom) daily averaged subtidal velocities (timescales > 33 h) for the (left) BASS 1 and (right) BASS 2 analysis periods (Table 1). Dashed lines give uncertainties of ellipse inclination, and arrows mark the direction of current rotation. Stippled ellipses correspond to spring and neap tide conditions. The x and y axes are the on-bank and along-bank orientations defined in Figure 1, respectively.

5 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS GLO 5-5 Table 3. Ellipse Parameters of the Largest Diurnal and Semidiurnal Tidal Constituents With (±) 95% Confidence Limits a BASS 1 3 Feb. to 14 Apr. (71 d) BASS 2 11 July to 14 Aug. (35 d) VMCM 1 Feb. to 22 Aug. (203 d) Tidal constituent U maj (U min ), cm s 1 q, degrees U maj (U min ), cm s 1 q, degrees U maj (U min ), cm s 1 q, degrees S 2 (12.00 h) 4.7 ± 0.3 (2.6 ± 0.3) 16 ± ± 0.5 (0.7 ± 0.5) 21 ± ± 0.3 (3.5 ± 0.3) 4 ± 5 M 2 (12.42 h) 25.7 ± 0.3 (15.8 ± 0.3) 20 ± ± 0.5 (13.1 ± 0.5) 16 ± ± 0.3 (25.4 ± 0.3) 2 ±5 N 2 (12.66 h) 4.4 ± 0.3 (3.0 ± 0.3) 12 ± ± 0.5 (3.7 ± 0.5) 14 ± ± 0.3 (5.7 ± 0.3) 3 ±5 K 1 (23.93 h) 2.0 ± 0.2 (1.5 ± 0.2) 84 ± ± 0.2 (1.8 ± 0.2) 72 ± ± 0.1 (3.3 ± 0.1) 83 ± 10 O 1 (25.82 h) 1.0 ± 0.2 (0.7 ± 0.2) 81 ± ± 0.2 (0.5 ± 0.2) 22 ± ± 0.1 (1.1 ± 0.1) 75 ± 19 a Results are shown for depth-averaged BASS measurements at current sensors 1 5 during winter (BASS 1) and summer (BASS 2), and the continuous (winter through summer) depth-averaged VMCM measurements. See Appendix A about orientation uncertainty. were U Maj = 25.7 ± 0.3 cm s 1 (23.7 ± 0.5 cm s 1 ) and U Min = 15.8 ± 0.3 cm s 1 (13.1 ± 0.5 cm s 1 ) for BASS 1 (BASS 2), respectively, where values denoted by ± are 95% confidence limits (Table 3). Low-frequency modulation of the M 2 currents by the N 2 and S 2 is on the order of several cm s 1 and takes place on timescales of 27.6 days (large spring neap cycle) and 14.8 days (small spring neap cycle), respectively (Figures 3 and 4). Subtidal velocities with periods >33 h reached magnitudes comparable to the along-bank ( y) M 2 amplitude in winter, and were smaller than the M 2 currents in summer (Figure 3). [13] The depth-averaged currents measured by BASS were >10 cm s 1 for 95 96% of the BASS 1 and BASS 2 analysis periods. Hence flow speeds rarely approached zero independent of flow direction (Figure 4). The implication is that shear production of turbulence took place during all parts of the tidal cycle. 4. Logarithmic Layer [14] Dimensional analysis shows that in the absence of stratification, the velocity distribution of turbulent flow near a rough surface is logarithmic for steady nonrotational [Clauser, 1956], steady planetary [Tennekes, 1973], and rectilinear oscillating flows [Grant and Madsen, 1986]. In the logarithmic layer, the velocity distribution follows the law of the wall UðÞ¼ z k ln z z 0 ; where U is current speed at elevation z, is bottom shear velocity, k = 0.4 is von Karman s constant, z is height above bottom, and z 0 is apparent bottom roughness as inferred from the height of the zero intercept of U(z). The existence of a logarithmic layer is consistent with the assumption of a constant stress layer in which t b ð1þ j j ¼ u 2 r 0 * ; ð2þ where t b is bottom stress and r 0 is a reference density [e.g., Kundu, 1990; Tennekes, 1973]. Expressions (1) and (2) imply that the direction of the velocity vector in the Figure 4. Time series of current speed U from (top) depth-averaged BASS measurements at current sensors 1 5 and (bottom) VMCM data at 6 m above the bottom. Heavy marks on the horizontal axis correspond to the end of each month, and thin marks denote the end of days 10, 20, and 30.

6 GLO 5-6 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS logarithmic layer is vertically uniform. For BASS data, veering of the velocity vector over the tripod height is indistinguishable from the ±5 sensor alignment uncertainty plus additional error due to flow distortion by the tripod instrumentation. [15] Applicability of (1) to BASS data, particularly during the stratified summer months, depends on the rate of buoyant destruction versus flow-induced shear production of turbulence. Effects of stratification on velocity distribution in the wall layer are described by the log-linear profile [Turner, 1973] UðÞ z ¼ u * k ln z þ 5R f ; ð3þ z 0 where R f ¼ z L M is the flux-richardson number and L M ¼ u3 * kagwt 0 is the Monin-Obukhov length. In (5), wt 0 is vertical heat flux, and a is the coefficient of thermal expansion. For stable stratification, the Monin-Obukhov length is the distance from the boundary at which buoyant destruction approximately balances shear production. In the ocean, awt 0 translates into wr 0 /r 0. With the eddy coefficient assumption wr 0 = Kdr/dz and K = k z, expressions (5) and (4) become and L M ¼ u 2 * 1 kn z R f ¼ ð4þ ð5þ ð6þ! k Nz 2 ; ð7þ where N =( g/r 0 dr/dz) 1/2 is the buoyancy frequency. [16] We computed N 2 and R f based on thermal stratification estimates from BASS thermistor data and salinity stratification estimates from linear T-S fits. Fits were performed using SeaCAT data collected at the nearby subsurface mooring at z = 11m above the bottom. Each fit utilized two days of hourly averaged data, and the center point of the fits was shifted though the time series at one-hour intervals. The fitted intercepts and slopes were then used to compute salinity from BASS temperature measurements. The underlying assumptions are that the T-S relation is depth-independent for z 11m above the bottom and that it varies slowly in time. These assumptions together with the measurement error and uncertainties of the least squares fits imply that N 2 estimates present no more than an order of magnitude approximation. Our results suggest that during 98% of the BASS 1 deployment, temperature was vertically homogeneous within ±0.01 C over the tripod height, and N 2 <10 5 s 2. For BASS 2, thermal stratification was within ±0.01 C about 90% of the time and N 2 <10 5 s 2 about 83% of the time within z 2.5m above the bottom. Taking = 1.0 cm s 1 as a representative shear velocity (see section 5.1), N s 2 and z 2.5m, expression (7) gives R f 0.1. We then performed logarithmic fits to (3) assuming that R f increased linearly from zero at BASS sensor 1 (0.2 m) to 0.1 at sensor 4 (2.6 m). The resulting shear velocities were indistinguishable from fits to (1) within the 95% confidence limits of (Appendix B). Hence we conclude that (1) is a good approximation of the velocity distribution at BASS sensors 1 4 during both winter and summer. [17] In the case of a strong rectilinear tide, Soulsby and Dyer [1981] found expression (1) needs to be modified in order to account for the effects of tidal acceleration and deceleration. Gross and Nowell [1983] presented limited observational evidence supporting Soulsby and Dyer s [1981] results, but concluded uncertainties of turbulence measurements and logarithmic fits were too large to justify refinements in logarithmic layer theory. Given the complexity of BASS data processing (Appendix A) and the uncertainty of the logarithmic fits (Appendix B), we agree with Gross and Nowell [1983] and use (1) without modification. [18] Based on Clauser s [1956] laboratory results, Grant and Madsen [1986] estimated the logarithmic layer height is about 0.1d, where d is the thickness of the vertically sheared bottom boundary layer. Analysis of VMCM data indicates the M 2 currents are strongly sheared in the lowest 40% of the water column, corresponding to d 30 m [Werner et al., 2003]. Hence a logarithmic layer of 3-m thickness is expected. This thickness approximately corresponds to the elevations of the four BASS current meters used for logarithmic fits ( m). 5. Bottom Friction Parameters 5.1. Shear Velocity [19] Logarithmic fits were performed on burst-averaged velocity data at BASS sensors 1 4 using expression (1). From the fits, one estimate of was obtained every half hour. Time-averaged shear velocities are =1.24± 0.02 cm s 1 (BASS 1) and = 1.13 ± 0.02 cm s 1 (BASS 2), where ±0.02 cm s 1 represents the 95% confidence limit of (Appendix B). Standard deviations from are 0.5 cm s 1 (BASS 1) and 0.4 cm s 1 (BASS 2). Figure 5 shows largest shear velocities coincide with flow along the major axis of the tidal ellipse. For BASS 1, estimates of exhibit an apparent bias toward larger values during off-bank compared to on-bank flow. We believe that this bias is the result of flow distortion by the camera and light battery (Figure 2). BASS 1 current meters 1 3 were located in the wake of the battery case for off-bank flow at angles a (clockwise from +x) Bottom Drag [20] Shear velocities from logarithmic fits show a linear increase with current speed, in agreement with the quadratic drag law u 2 * ¼ c D U 2 ; where c D is p the ffiffiffiffiffi bottom drag coefficient (Figure 6). Estimates of c D were obtained from least squares fits of U versus, and results were squared to give c D. For the ð8þ

7 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS GLO 5-7 Figure 5. Half-hourly estimates of shear velocity derived from logarithmic profiles for (top) BASS 1 and (bottom) BASS 2 versus average flow direction a at BASS current sensors 1 3. Flow directions a = are shown schematically at the bottom of the figure. Solid lines mark the orientation of the major axis of the tidal current ellipse, and dashed lines are uncertainties of ellipse inclination. Dotted lines in the top panel indicate flow directions in the wake of the camera and light battery. fits, we used burst-averaged velocity data and half-hourly estimates of. The few profiles with shear velocities <0 were excluded from the least squares fits (Appendix B). Also used were half-hourly averaged VMCM measurements at z = 6 m. Although the VMCM elevation was well above the expected 2 3-m thickness of the logarithmic layer, estimation of bottom stress from VMCM measurements is necessary at times when BASS data are not available. Hourly subsampling of current speed and shear velocity ensured estimation of c D from temporally independent data. The 1-h subsampling interval was chosen based on autocorrelation functions of U and following the procedure outlined in Appendix B for. [21] Table 4 lists the results of the fits for BASS 1, BASS 2, and BASS 1 and BASS 2 combined. The computed winter drag coefficients are about 6 13% larger at the 95% confidence level than the corresponding summer values. Exclusion of profiles affected by the camera and light battery (a = ) reduces the BASS 1 record length by 15% and gives c D estimates that are indistinguishable from the summer results within error limits. For comparison, removing the same flow directions from (not obstructed) BASS 2 data changes the summer drag coefficients by less than 4%. The logical conclusion is that using the complete BASS 1 data record leads to overestimation of c D, and that exclusion of profiles impacted by flow obstruc-

8 GLO 5-8 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS Figure 6. Current speed U at 1.2 m above the bottom versus estimated friction velocity from logarithmic profiles for (left) BASS 1 and (right) BASS 2. Each data ppoint corresponds to one BASS measurement burst. Also shown are (solid) the fitted slope = ¼ ffiffiffiffiffiffiffiffiffi ðc D Þ U and (dashed) the 95% confidence limits of the fits. Fits were performed on hourly subsampled data. tion from the camera and light battery yields more reliable results. [22] Temporal variation of c D on timescales of days to weeks was examined by subdividing the BASS 1 and BASS 2 data records into bins of 4 days length, and by performing least squares fits on hourly subsampled data within each bin. All profiles including those in the wake of the camera and light battery entered the fits. This decision was made based on comparison of least squares fits with and without off-bank flow directions a = removed from BASS 1 and BASS 2 data; for the 4-day time periods investigated here, exclusion of off-bank flow profiles reduced c D by about 20 40% for BASS 1 and BASS 2 compared to incorporation of all available data. No camera and light battery was attached to the BASS 2 tripod, so that the observed reduction in drag coefficient is the result of systematic rejection of large off-bank currents rather than a realistic consequence of avoiding obstructed flow directions. This problem does not apply to the much longer analysis periods listed in Table 4. [23] Since the number of data points entering the 4-day fits is small compared to that of the previously used full BASS 1 and BASS 2 data records, the 95% confidence limits of c D are large (Figure 7). As a result, we cannot conclude whether the observed temporal variation of c D is real for much of the BASS 1 and BASS 2 deployments. An exception is the beginning of February, when c D estimates at 1.2 m above the bottom fall from peak values near 4.9 ± (4 7 February) to 2.9 ± (8 11 February). During that time, a strong winter storm caused surface wind stress magnitudes near 9 dyne cm 2, producing large surface waves and bottom orbital wave velocities (Figure 8). Subsequent storms did not raise the drag coefficient beyond the background level to a degree that can be distinguished from experimental uncertainty Bottom Stress [24] Bottom stress can be estimated from shear velocity by t x b ; ty b ¼ u 2 * ðcos b; sin bþ; ð9þ or using c D and current velocity t x b ; ty b ¼ cd Uu; ð vþ: ð10þ In (9), b is the angle of (t b x, t b y ) with respect to the on-bank (+x) axis and defined as the direction of the vertically averaged velocity vectors measured by BASS current meters 1 3. Since the M 2 tide carries about 76% or more Table 4. Quadratic Drag Coefficients From Least-Squares Fits of U Versus and Upper and Lower Bounds of Bottom Roughness Length z 0 From Inversion of c D at BASS Current Sensors 1 5 and the Lowest VMCM a Sensor Height, Meters Above Bottom b BASS 1 (A) 3 Feb. to 14 Apr. (71 d) BASS 1 (B) 3 Feb. to 14 Apr. (71 d) BASS 2 11 Jul. to 14 Aug. (35 d) BASS 1 (B) and BASS 2 (106 d) c D 10 3 (z 0 cm) 0.22/ ± 0.2 ( ) 5.0 ± 0.2 ( ) 4.6 ± 0.2 ( ) 4.6 ± 0.1 ( ) 0.58/ ± 0.1 ( ) 3.6 ± 0.1 ( ) 3.7 ± 0.1 ( ) 3.6 ± 0.1 ( ) 1.18/ ± 0.1 ( ) 3.0 ± 0.1 ( ) 3.0 ± 0.1 ( ) 3.0 ± 0.1 ( ) 2.53/ ± 0.1 ( ) 2.4 ± 0.1 ( ) 2.4 ± 0.1 ( ) 2.4 ± 0.1 ( ) 4.43/ ± 0.1 ( ) 2.0 ± 0.1 ( ) 2.0 ± 0.1 ( ) 2.0 ± 0.1 ( ) VMCM, 6 m 2.1 ± 0.1 ( ) 1.8 ± 0.1 ( ) 1.7 ± 0.1 ( ) 1.8 ± 0.1 ( ) a Results for BASS 1 are listed for (A) fits using all burst-averaged velocity measurements and for (B) fits with flow profiles taken in the wake of the camera and light battery (a = ) excluded. Drag coefficients in the last column are from combined BASS 1 (B) and BASS 2 data. b Values on the left correspond to BASS 1, while values on the right correspond to BASS 2.

9 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS GLO 5-9 Figure 7. Temporal variation of c D. Drag coefficients are from least squares fits of hourly subsampled data as described in the text. Fits were performed on time windows of 4 days length, with center points of subsequent fits separated by 2 days. Error bars give the 95% confidence limits of the fits. Heavy marks on the horizontal axis correspond to the end of each month, and thin marks denote the end of days 10, 20, and 30. of the kinetic energy (Table 2), most of the variance in bottom stress occurs on tidal timescales. However, comprehensive analysis of the circulation dynamics on Georges Bank requires investigation of tidal as well as subtidal flow processes. Hence we test the predictive capabilities of (9) and (10) for the tidal and subtidal components of bottom stress. [25] Bottom stress estimates were made utilizing halfhourly velocity data and shear velocities. The quadratic drag law was applied to BASS measurements at z = m and VMCM data at z = 6 m, with c D from the last column of Table 4. Results from (9) and (10) were hourly averaged, and a harmonic analysis was performed on time periods of four days length. The length of the analysis periods allows for resolution of the M 2, but not of the N 2 and S 2.Asa result, amplitudes of tidal bottom stress exhibit modulation by the large and small spring neap cycles (Figure 9). A lowpass filter operator with half-power period 33 h was applied to obtain subtidal estimates of bottom stress. [26] M 2 bottom stress amplitudes from (9) and (10) show good agreement for BASS 2, while significant deviations exist for BASS 1. Large differences between the two methods before 9 February can be explained by the failure of the constant-c D approximation to account for enhanced bottom drag during the February storm (Figure 7). After 9 February, deviations between (9) and (10) are obvious for the tidal and subtidal cross-bank components during BASS 1, while both expressions yield similar results for the tidal and subtidal along-bank components. In light of earlier results from logarithmic fits indicating estimates are larger during the off-bank than the on-bank flow phase of BASS 1 (Figure 5), we conclude differences between (9) and (10) are mainly due to flow distortion caused by the camera and light battery. This conclusion is supported by a simple experiment: increasing c D by 100% of its estimated value at times when the flow passes the camera and light battery (a = ), and using this temporally varying drag coefficient in (10) gives standard errors relative to (9) x x that are about 30% (40%) smaller for t bm2 (t bst ) than in the case of constant c D. The implication is that flow obstruction by the camera and light battery increases the drag coefficient experienced by the off-bank flow. Since the enhancement of bottom drag occurred periodically, it particularly affected the cross-bank component of bottom stress. Such behavior is not representative of the flow physics, but merely an artifact of the tripod instrumentation. Hence we conclude (10) gives more realistic bottom stress estimates than (9) during BASS 1. [27] For both BASS deployments, bottom stress estimates from the quadratic drag law at BASS sensors 1 3 are in close agreement with those obtained at the lowest VMCM 6 m above the bottom (Figure 9). This result implies that bottom stress can be estimated from VMCM data at times when bottom tripod measurements are not available Bottom Roughness [28] Combination of (1), (2), and (8) gives the relation between drag coefficient and bottom roughness 2 c D ¼ 4 32 k 5 ; ð11þ ln z z 0 which can be inverted to infer z 0. In (11) and (1), z 0 is the apparent bottom roughness, i.e., the bottom roughness experienced by the flow as it passes over roughness elements of a given shape and size. For hydrodynamically fully rough flow over a flat bottom, z 0 = d/30, where d is the median grain diameter of the bottom sediments [Nikuradse, 1950]. In the presence of larger roughness elements such as sea shells or sand ripples, d is replaced by an equivalent roughness height k b. [29] Using (11) and the drag coefficients derived in section 5.2, representative bottom roughness lengths at the BASS deployment site are cm within error limits (Table 4). Estimates of z 0 are similar for BASS 1 and BASS 2, consistent with the fact that c D displays no seasonal variability. The range of bottom roughness length in Table 4 exceeds d/30 = cm for medium-to-

10 GLO 5-10 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS Figure 8. From top to bottom: time series of wind stress t W inferred from ST1-meteorological data using the Large and Pond [1981] neutral stability algorithm, significant wave height A W reported by the NDBC environmental buoy 44011, wave period T W and wave velocity u b at 1.2 m above the bottom from spectral analysis of BASS velocity data (Appendix C), and friction velocity in the wave boundary layer cw predicted by GM79 for z 0 =0.07cm(k b = 2.1 cm). Horizontal lines in the bottom panel mark the critical shear velocities for bed load transport of medium sand ( cr =1.38cms 1 ), bed load transport of coarse sand ( cr = 1.7 cm s 1 ), and suspension of fine sand ( s = 3.1 cm s 1 ). E1-8 are wave events investigated in this study. coarse sand (d = mm) by at least one order of magnitude. Thus, the bottom roughness experienced by the flow was not defined by the sand grain size. The explanation is that the bottom was not flat, but partly covered by sea shells, benthic grazers and sand ripples (e.g., frame 1 in Figure 10). Taking k b = 1 2 cm as a characteristic vertical scale of the largest roughness elements (sea shells) protruding from the sediment surface yields k b /30 = cm in better agreement with z 0 = cm from Table 4. Alternatively, we can derive a representative range of k b from the geometry of the observed sand ripples. For continuous ripples of infinite length, k b ¼ 30 h h l ; ð12þ where h is ripple height and l is separation distance [Grant and Madsen, 1982]. In the present case, (12) represents

11 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS GLO 5-11 x Figure 9. From top to bottom: time series of tidal (M 2 ) cross-bank bottom stress amplitude t bm2, y x along-bank bottom stress amplitude t bm2, subtidal (timescales > 33 h) cross-bank bottom stress t bst, and y along-bank bottom stress t bst. M 2 amplitudes are from harmonic analysis of bottom stress estimates performed on time windows of 4 days length; the temporal spacing of consecutive time windows is 1 day. Heavy marks on the horizontal axis correspond to the end of each month, and thin marks denote the end of days 10, 20, and 30. Heavy solid lines are average results from the quadratic drag law at BASS sensors 1 3 with c D taken from Table 4. The thin solid line represents estimated bottom stress from the quadratic drag law at 6 m above the bottom using VMCM data and c D from Table 4. Dashed lines are from logarithmic fits to BASS measurements at current sensors 1 4, where the direction of the bottom stress vector corresponds to the average flow direction at BASS sensors 1 3. Shaded lines in the upper two panels give the difference between the M 2 amplitudes of BASS bottom stress estimates from logarithmic fits and the quadratic drag law (i.e., the difference between the dashed and heavy solid line). The line width corresponds to the 95% confidence limits imposed by the tidal decomposition. a rough approximation since ripple crests were discontinuous and not fully unidirectional (e.g., frame 1 in Figure 10). Sand ripples were about 1 2 cm high and spaced at cm distance. Substituting h = 1 2 cm and l = cm in (12) gives k b = cm, and hence k b /30 = cm. The lower range of these estimates is consistent with z 0 = cm from Table 4. [30] In summary, the consistency of z 0 from inversion of c D with estimates z 0 = k b /30 from bottom photographs increases our confidence in the results listed in Table Bottom Bed Forms [31] Bed form dynamics take place on timescales of hours to months and develop as a result of winter storms, tidal

12 GLO 5-12 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS Figure 10. Sequence of bottom photographs taken (frame 1) immediately before, (frames 2 5) during, and (frame 6) immediately after the February storm (E1 in Figure 8). The black arrow marks the on-bank (+x) orientation (Figure 1). Yellow and red arrows give the direction of the instantaneous total (subtidal and tidal combined) and tidal flow vectors, respectively. Blue arrows correspond to (frame 3 5) instantaneous wave direction and (frame 6) wave direction at the end of the storm, i.e., during the resuspension event preceding frame 6 (see section 6.1). variation, spring neap cycling and biological activity. The following sections describe the main aspects of bed form evolution at ST1 through winter and spring 1995, corresponding to the period for which bottom photographs are available. We present the results in chronological order, starting with the February storm. Quantitative examination of events featuring strong wave-current interaction (including the February storm) is reserved for section 7.

13 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS GLO 5-13 Figure 11. Bottom photographs taken (frame 1) 4 days and (frame 2) 5 days after the February storm around the peak of ebb. Frame 2 shows ripple amplification as described in section 6.1. The black arrow marks the on-bank (+x) orientation (Figure 1). Yellow and red arrows give the direction of the instantaneous total (subtidal and tidal combined) and tidal flow vectors, respectively February Storm [32] A sequence of bottom photographs taken 4 6 February illustrates the effects of large surface waves on bottom bed forms (Figure 10). Immediately before the storm, the nearbottom waters were clear (frame 1). The sea floor was covered by east-west sand ripples oriented in the approximate direction of the local isobath, with flanks perpendicular to the major axis of the tidal current ellipse. Ripple crests were discontinuous and about 1 2 cm high, their horizontal spacing was cm. Grainy substances in the ripple troughs represent aggregations of biomass such as detritus from sinking diatoms and fecal material from benthic organisms. Also visible are shell hash and benthic grazers. Photographs taken 8, 16, and 24 h previously to frame 1 indicate that ripples were stationary, i.e., they did not migrate with the flow (not shown). Bed forms in frame 1 resemble observations by Butman [1987] made in March 1978 at a 64-m deep study site on the southern flank of Georges Bank near ST1. [33] In the early hours of 5 February, large surface winds produced a steep increase in significant surface wave height and wave-induced bottom orbital velocity (Figure 8). This marked the beginning of the February storm. Bottom photographs show intermittent events of sediment resuspension (frames 2 and 5) as well as bed forms of changing geometry (frames 3 6). Bifurcated ripples in frame 4 developed in response to the large angle (30 50 ) between the instantaneous waves and currents (Appendix C), as opposed to the more uniformly oriented ripples for wave and currents at small angle (frame 3). This result is in agreement with earlier observations from the Scotian shelf [Li and Amos, 1998]. Sediment resuspension has recurred in frame 5. Bed forms are difficult to identify, but appear to be irregular north-south ripples that are roughly aligned with the tidal velocity vector. This suggests that ripples in frame 5 are wave-dominant, i.e., they formed in response to the surface-wave-induced velocities with principal axis perpendicular to the tidal flow and main ripple orientation. Observations of wave-dominant ripples were also made by Li and Amos [1998] on the Scotian shelf, where they occurred during events of strong sediment resuspension. Several other bottom photographs taken during the February storm show wave-dominant ripples; these photographs either coincided with or immediately followed intermittent resuspension events (not shown). Frame 6 displays conditions immediately after the storm when surface waves barely penetrated to the bottom (Figure 8). Suspended material was clearly visible on the photograph preceding frame 6 by 8 h (not shown). Waveinduced velocities during the suspension event were approximately perpendicular to the predominant orientation of the roughness elements in frame 6, indicating the observed bed forms are relics of earlier ripples formed by waves. [34] Bed forms did not change significantly during the first four days after the storm, a period which coincided with neap tide conditions. An example is shown for 12 February in Figure 11 (frame 1). BASS bottom stress estimates from the quadratic drag law (averaged over sensors 1 3) explain the absence of bed form variation: from 8 to 12 February, bottom stress estimates were below the critical stress for initiation ofmovement ofmedium sand (t cr =1.94 dynecm 2 ) about 96% of the time and did not exceed the threshold for movement of coarse sand (t cr = 2.95 dyne cm 2 ). Critical stress values were taken from Shield s curve for sand grain diameters 0.25 mm (medium sand) and 0.50 mm (coarse sand) [Mantz, 1977; Yalin and Karahan, 1979; Butman, 1987]. After 12 February, the bottom stress amplitude had increased sufficiently toward spring tide conditions to exceed the thresholds for fine and medium sand movement near the peaks of ebb and flood. This lead to gradual amplification of those roughness elements with crests that were predominantly aligned along-bank, i.e., perpendicular to the major axis of the tidal current ellipse. Frame 2 in Figure 11 shows the beginning stages of this amplification. Over the next days, bed forms evolved into discontinuous east-west ripples with similar geometry to the prestorm conditions in frame 1 of Figure February-April: East-West (Along-Bank) Sand Ripples [35] East-west sand ripples similar to Figure 11 (frame 2) were observed from February to April. Their geometry was

14 GLO 5-14 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS Figure 12. Sequence of bottom photographs representative of spring tide conditions in early March. The black arrow marks the on-bank (+x) orientation (Figure 1). Yellow and red arrows give the direction of the instantaneous total (subtidal and tidal combined) and tidal flow vectors, respectively. subject to tidal variation as well as low-frequency modulation over a spring neap cycle. Intermittent modification of the east-west ripples happened during a few wave events, but visible effects of surface waves on bottom bed forms were short-lived (hours) and not nearly as pronounced as during the February storm. In the following, we discuss typical bed form geometry during spring and neap tide conditions Spring Tide [36] During spring tide, ripple geometry varied on tidal timescales as a result of the reversing tidal currents. This is illustrated in Figure 12, which shows four bottom photographs taken at 8-h intervals on 1 2 March around the peak of the spring neap cycle (Figure 4). During this period, significant surface wave heights were <3m, and wave-induced near-bottom velocities were negligible (<3 cm s 1 ). Frame 1 was taken about 2 h after the tidal (M 2 ) velocity vector had passed the off-bank ( x) axis, corresponding to the fifth hour of ebb. Ripples are aligned east-west, their crests are clearly defined by sharp lines. Elongated, smooth flanks face on-bank, while rough, steep flanks face off-bank. Grainy substances representing biomass are mostly clustered in the ripple troughs. Ripple geometry and biomass distribution indicate that sediment transport occurred as bed load, leaving a smooth surface on the upstream facing ripple flanks where organic deposits were covered by sand. Sediment movement from troughs to crests caused elongation of the upstream ripple flanks, accumulation of sediment on the crests and steepening of the downstream ripple flanks. Frame 2 corresponds to flow conditions approximately 1 h after the beginning of ebb. Ripples have reversed, i.e., smooth, elongated flanks face off bank, while rough, steep flanks face on-bank. The elongation and plane surface of the off-bank-facing ripple flanks are remnants of the preceding flood. Frame 3 shows similar ripple geometry as frame 2, but with shorter flanks on the upstream (off-bank) side of the sand ripples. The photograph was taken about 3 h after the onset of flood, as opposed to briefly after completion of a flood cycle in frame 2. Hence downstream bed load transport has not taken place sufficiently long to elongate upstream ripple flanks to the same extent as in frame 2. The last frame in the series, frame 4, was taken 24 h (about 2 tidal cycles) after frame 1. Frame 1 and frame 4 represent similar flow conditions and ripple geometry. [37] BASS data give quantitative support for the conclusions drawn from bottom photographs. This is illustrated in Figure 13a, which shows BASS velocity measurements and bottom stress estimates for the spring tide conditions pre-

15 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS GLO 5-15 Figure 13. (top) Half-hourly, depth-averaged velocities at BASS sensors 1 3, and (bottom) half-hourly bottom stress estimates for (a) 2000 UT 1 March to 2000 UT 2 March (spring tide conditions), and (b) 0400 UT 13 March to 0400 UT 14 March (neap tide conditions). Bottom stress was computed using c D from Table 4 and BASS measurements at sensors 1 3; estimates in the Figure represent the average results for sensors 1 3. Solid and dashed lines give the direction of the instantaneous total (subtidal and tidal combined) and tidal flow, respectively, at the times of frames 1 4 (F1-4) in Figures 12 and 14. The two concentric circles in the lower panels mark the thresholds for initiation of sediment movement of (inner circle) medium sand and (outer circle) coarse sand. sented in Figure 12. Asymmetries between flood and ebb reflect the subtidal currents and their nonlinear effect on bottom stress. For the 2-day period investigated here, mean currents reached 8 cm s 1 in the along-bank ( y) direction. About 55% (16%) of the stress estimates for 1 2 March are >t cr = 1.94 dyne cm 2 (>t cr = 2.95 dyne cm 2 ), corresponding to the critical shear stress for initiation of movement of medium (coarse) sand. Shear stress estimates are >t cr most frequently during ebb and flood, and least frequently around the reversal from flood to ebb/ebb to flood. Hence bottom stress was sufficiently large to induce bed load transport during on- and off-bank flow, in agreement with the observed elongation of the upstream ripple flanks on alternating sides of the ripple crests (Figure 12). On the other hand, bottom stress was not sufficiently large to cause significant sediment motion during slack tide (along-bank flow), thus prohibiting ripple migration perpendicular to the rotating tidal flow Neap Tide [38] Variation of ripple geometry was not recognizable during neap tide. This is illustrated in Figure 14 for March corresponding to the time of smallest current amplitude within the spring neap cycle (Figure 4). Ripple flanks show little asymmetry between the upstream and downstream sides, and deposits of biomass are visible anywhere on the ripple surface and in the ripple troughs. These observations suggest that the bottom shear stress was too small to initiate sediment movement from the troughs to the crests, thus no longer elongating the upstream ripple flanks and covering biomass in the process. Rounded ripple crests are indicative of erosion. The implication is that the shear stress fell below the critical value for sediment movement first in the troughs and then over the crests [e.g., Nielsen, 1986], hence allowing for gradual erosion of the ripple crests as the current amplitude decreased from spring to neap tide. Similar erosion was prohibited during spring tide when surfacial material on the crests was replenished by sediment from the troughs. [39] Estimates of bottom stress are <t cr = 1.94 dyne cm 2 (<t cr = 2.95 dyne cm 2 ) about 95% (100%) of the 2-day investigation period in Figure 14 (Figure 13b). Hence BASS

16 GLO 5-16 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS Figure 14. Sequence of bottom photographs representative of neap tide conditions in mid-march. The black arrow marks the on-bank (+x) orientation (Figure 1). Yellow and red arrows give the direction of the instantaneous total (subtidal and tidal combined) and tidal flow vectors, respectively. data support the conclusion that sediment movement was not initiated most of the time during neap tide April June: Flat Bottom [40] Sand ripples gradually vanished during late March and April as shown in Figure 15. Frame 1 (28 March) is representative of spring tide conditions and exhibits ripples that are less clearly defined than in Figure 12 as well as higher bioturbation levels. The observed increase of bioturbation is in agreement with studies showing rising phytoplankton levels from March to May on the southern flank of Georges Bank, which contrast the sharply defined spring bloom observed over the rest of the Bank [O Reilly and Zetlin, 1998]. Evidence is mounting that much of the phytoplankton is not assimilated by zooplankton, but instead sinks to the lower water column where it triggers enhanced abundance on higher trophic levels. As a result, the density of benthic organisms with size >0.5 mm nearly doubles from February to May [Maciolek and Grassle, 1987]. Frame 2 (17 April) features a nearly flat bottom covered by a dense biological mat. With the exception of small wave-induced ripples during a storm event in May, bottom photographs are similar to frame 2 until our series ends in June. [41] Continuous bottom stress estimates were greater than the earlier defined thresholds for initiation of fine and medium sand movement during large parts of the spring neap cycle in winter as well as in summer (Figure 16). Hence changes in near-bottom flow dynamics cannot explain the disappearance of sand ripples. An alternative explanation is biogenic modification of bottom sediment during the biologically active spring and summer months. This conclusion is supported by results from Grant et al. [1982], who found that biological adhesion, i.e., binding of bottom sediment by bacterial mucus, enhanced t cr by as much as 100% compared to predictions from Shield s curve in a region north of Cape Cod. Bottom sediments were fine sand mixed with less than 2% silt and clay, similar to the mostly sandy bottom at the BASS deployment site. Based on Grant et al. s [1982] results, we infer that cementing of the sediment surface by organic material gradually enhanced t cr at ST1 beyond the bottom stress imposed by the tidal flow. Gradual disappearance of bottom bed forms can be explained by slow erosion of the ripple crests: the shear stress fell below the increasing threshold for sediment movement first in the troughs and then over the crests [e.g., Nielsen, 1986], thus subjecting the crests to gradual erosion. In contrast, recurring sediment

17 WERNER ET AL.: BOTTOM FRICTION AND BED FORMS GLO 5-17 Figure 15. Bottom photographs showing increasing bioturbation levels and gradual disappearance of sand ripples during late March and April. The black arrow marks the on-bank (+x) orientation (Figure 1). Yellow and red arrows give the direction of the instantaneous total (subtidal and tidal combined) and tidal flow vectors, respectively. transport from troughs to crests during spring tide maintained ripples in winter when t cr was smaller. 7. Wave-Current Interaction [42] When surface wave induced motion penetrates to the bottom, a thin region of wave-induced turbulence may develop immediately above the sea floor. This region is called the wave boundary layer. Inside the wave boundary layer, wave-induced turbulence enhances the apparent bottom roughness felt by the flow above it and increases the effective shear stress acting upon the sediment surface. A characteristic wave boundary layer thickness is d cw ¼ k cw, w where cw is the friction velocity in the wave boundary layer due to waves and currents and w is wave frequency [Grant and Madsen, 1986]. For typical wave periods near 10 s and friction velocities around 1 cm s 1, the wave boundary layer is several centimeters thick. [43] Surface waves penetrated to the bottom on several occasions during the BASS 1 deployment, either because wave amplitudes were particularly large or wave periods were particularly long (Figure 8). For some events, the resulting increase in bottom shear stress was significant enough to have a visible effect on bottom bed forms. The most obvious case was the February storm (Figure 10). [44] Predictions of sediment movement and resuspension by combined waves and currents require knowledge of the wave-enhanced bottom stress at the sediment/water interface. Our measurements do not allow for direct estimation of cw since BASS velocity sensors were positioned at heights >0.2m well above the wave boundary layer. We used the Grant and Madsen [1979] combined wave current interaction model (hereafter GM79) to estimate bottom stress during the February storm and other events, and to relate predictions of sediment movement and resuspension to observations. All events investigated here occurred Figure 16. Continuous time series of bottom stress estimated from the quadratic drag law using VMCM measurements at z = 6 m and c D from Table 4. Horizontal lines mark the critical shear stress for initiation of motion of medium and coarse sand. Heavy marks on the x-axis correspond to the end of each month, and thin marks denote the end of days 10, 20, and 30.