Page | 1

The origin of along-shelf pressure gradient in the Middle Atlantic Bight

F.-H. Xu* and L.-Y. Oey

Princeton University

*Corresponding Author:

September08, 2010

7915 words.

Abstract

It has been known for some time that the along-shelf pressure gradient (ASPG) contributes to driving the time-mean, equatorwarddepth-averaged currents on the Middle Atlantic Bight (MAB) shelves off the U.S. northeastern coast. The origin of ASPG and its seasonal and inter-annual variations remain to be explained, however. Possible contributors to ASPG are: wind and wind stress curl, Gulf Stream’s path, warm-core rings, Coastal Labrador Sea Water(CLSW) transport, and river discharge. In this work, sixteen years (1993-2008) of satellite data, data-assimilated model reanalysis, and tide-gauge sea level data are analyzed. It is shown thatthe mean ASPG isforced by river discharge and the CLSW transport. On the other hand, the seasonal and inter-annual variations in ASPG are caused by southwestward propagation of warm-core rings along the MAB slope, and also by impingement of these rings upon the shelf break north of the Gulf Stream near Cape Hatteras.It is shown that rings change the ASPG by forcing cross-shelfbreak transports that modify the sea surface heightover the outer shelf, and by producinga sea surface sloping down to the north when they are near Cape Hatteras. It is also shown that the large-scale wind cannot directly affect the ASPG. However, north of the Gulf Stream, there exist seasonal and inter-annual variations of eddy kinetic energy which affect the ASPG, and which may be in part forced by the wind.

  1. Introduction

The Middle Atlantic Bight (MAB) is the continental shelf region off the northeastern coast of the United States stretching between Nantucket Shoalsto the northeastand Cape Hatteras to the south (Beardsley and Boicourt, 1981). The MAB is a dynamically complex region where cooler and fresher shelf water is separated from warmer and saltier slope water by a shelf break front (e.g. Flagg et al. 2006). Understanding the water properties and currents in MABis important for navigation, fisheries, and coastal ecosystems.

The MAB circulation has been investigated through observational and modeling studies over several decades (Beardsley and Boicourt 1981; Chapman1986; Linder and Gawarkievicz 1998; Flagg et al. 2006; Lentz 2008a). Depth-averaged mean currents are mainly along-isobath directed equatorward, with speeds of 0.03-0.1 m s-1 (Lentz 2008a). The mean currents are westward on the New England shelf,and southwestward in the middle of MAB. South of the Chesapeake Bay, the currents veer offshore. The mean along-shelf currents also increase with distance offshore (Lentz 2010).

An important driving force for the depth-averaged mean currents in the MAB is the along-shelf pressure gradient (ASPG) (Beardsley and Boicourt 1981). Stommel and Leetmaa (1972) modeled the steady-statewintertime circulation, and concluded that an ASPG of order of 10-7 is required to drive the southwestward flow. Csanady (1976) argued that the cross-shelf density gradients are important and concluded also that an ASPG must exist to account for the observed circulation within the MAB. Lentz (2008a) extended Csanady’s model, analyzed observations, and showedquite convincingly that the southwestward along-shelf current is consistent with an along-shelf sea surface slope; he estimated an ASPG value of approximately 3.710-8. Lentz (2008a) discussed the possibility of other types of forcing, but the hypothesis that ASPG exists seems reasonable.

What drives the ASPG? Lentz (2008a) showed that along-shelf density gradients are negligible, so that ASPG is mainly due to the sea surface slope. The Gulf Stream and Slope Sea gyre (Csanady and Hamilton, 1988) may drive an ASPG at the shelf break, but the penetration of the pressure field onto the shelf is limited (Wang 1982; Csanady and Shaw, 1983; Chapman 1986). This work will further investigate the large-scale contributions to ASPG.

Observations also show seasonal variations in the depth-averaged along-shelf currents which are different in different sub-regions of the MAB (Lentz 2008b). Over the southern flank of Georges Bank, the along-shelf flow is maximum southwestward in September (Butman and Beardsley 1987; Brink et al. 2003; Flagg and Dunn 2003; Shearman and Lentz 2003). Further west and south in the MAB, the seasonalvariation is less clear (Mayer et al. 1979; Beardsley et al. 1985; Aikman et al. 1988). Along the Oleander line, Flagg et al. (2006) observed a shelfbreak jet (offshore of 100m-isobath)which was stronger southwestward in fall and winter and weaker in spring and summer. ADCP measurements at station 5 of the Coastal Ocean Bio-optical Buoy (COBY) transect (75.029W, 37.833N) show maximum southwestward currents in spring, and weak currents in summer and fall (Xu et al., manuscript in preparation). From analyses of 27 long-termmeasurements, many of which were taken in the New England Shelf, Lentz (2008b) found that the alongshore currents have amplitudes of a few cm s-1.The residual alongshore flow after the removal of wind-driven component is maximum southwestward in spring onshore of the 60m isobath. He suggested that the seasonality of the along shelf currents is primarily driven by the cross-shelf density gradient induced by freshwater discharge. Does ASPG also have seasonal and inter-annual variations, and if it does, how are they produced?

In this study, we carry out a set of model experiments and analyze them in conjunction with satellite and tide-gauge observations. We attempt to provide some answers to the origin of ASPG: its mean as well as seasonal and inter-annual variability. Although the focus is on the shelf, it seems reasonable (from the literature) to suggest that the ASPG can be due to larger-scale process(es) that requires careful considerations of forcing outside the MAB. We will examine mean, seasonal, and inter-annual variability. We examine the driving mechanisms, including the wind stress curl over the North Atlantic, Gulf Stream’s latitudinal shifts, warm-core rings, Coastal Labrador Sea Water(CLSW; Csanady and Hamilton, 1988) transport, and river discharge.

The paper is organized as follows: Section 2 presents a description of theobservation datasets used in the study. Section 3 describes the numerical model and experiments. In section 4, we analyze the mean, seasonal, and inter-annual variations in the along-shelf flow and ASPG in the MAB. The influence of wind, river, Gulf Stream, warm-core rings and CLSW transport are discussed in section 5. The paper concludes in section 6.

  1. Data

Sixteenyears (1993-2008) of quality-controlled sea level data off the eastern coast of the United States are obtained from the University of Hawaii Sea Level Center (UHSLC, at 12 stations (excluding Bermuda and Wilmington NC, Figure 1) are monthly running-averaged and analyzed to study seasonal and inter-annual variations in sea surface slope.

The gridded sea surface heights (SSH) and the corresponding geostrophic velocities for the period 1993 to 2008are from AVISO ( This dataset has a temporal resolution of 7 days and spatial resolution of . A detail description of the dataset is in Le Traon et al. (1998).

The Cross-Calibrated, Multi-Platform (CCMP) level 3 ocean surface wind velocity data is used to force the numerical ocean model (below). This is an ERA-40 Re-analysis, 6-hourly gridded () productthat incorporates satellite surface winds from Seawinds on QuikSCAT, Seawinds on ADEOS-II, AMSR-E, TRMM TMI and SSM/I, as well as ships and buoys measurements.

  1. The Numerical Model and the Experiment with Data Assimilation

The terrain-following (i.e. sigma) coordinate and time-dependent numerical model for this study is based on the Princeton Ocean Model (Mellor, 2002). The Mellor and Yamada’s (1982) turbulence closure scheme modified by Craig and Banner (1994) to effect wave-enhanced turbulence near the surface is used (Mellor and Blumberg, 2004). A fourth-order scheme is used to evaluate the pressure-gradient terms (Berntsen and Oey, 2010) and, in combination with high resolution and subtraction of the mean -profile, guarantees small pressure-gradient errors of O(mm/s) (c.f. Oey et al. 2003). The Smagorinsky’s (1963) shear and grid-dependent horizontal viscosity is used with coefficient = 0.1, and the corresponding diffusivity is set 5 times smaller (c.f. Mellor et al., 1994). The northwestern Atlantic Ocean model (NWAOM; Oey et al. 2003) uses an orthogonal curvilinear grid to cover the region 98W-55W and 6N-50N (Fig.2). In a process study in which forcing and sensitivity are to be explored, such a regional model offers some advantages compared to, e.g. a basin-scale North Atlantic model. The improved efficiency allows multiple long-term experiments to be conducted at a reasonably high resolution. Moreover, the NWAOM is used here to test the sensitivity of the modeled dynamics in MAB to the CLSWtransport, which is specified as a boundary inflow (see below). The drawback is that larger, basin or even global-scale variability are excluded. The assumption is then that these variability are of secondary importance in the interpretations of the MAB circulation processes.

The NWAOM has 25 vertical sigma levelsandhorizontal grid sizes 8~12 km in the MAB and the Slope Sea. The World Ocean Atlas data (“Climatological” data) from NODC ( pr_woa05.html) was used for initial condition as well as boundary condition along the eastern open boundary at 55oW. Across 55W, a steady transport combined with radiation using also the geostrophically-balanced surface elevation g (Oey and Chen, 1992a) specifies the Gulf Stream exiting near the Grand Banks south of Newfoundland with a magnitude of 93 Sv following Schmitz (personal correspondence, see also Schmitz, 1996; Hendry 1982; Hogg 1992; Hogg and Johns 1995). This is balanced by transports specified as broad return flows south (the “Worthington Gyre” - Worthington, 1976) and north (the “North Recirculation gyre” - Hogg et al. 1986) of the jet. The CLSW inflow is then specified (and adjusted; see below) as the northern portion of this “North Recirculation gyre.”The CLSW is identified here as being the same as Csanady and Hamilton’s (1988; their Fig.20b) “offshoot of the Labrador Current” that “turns the other way and intrudes into the Slope Sea.”Much of this west/southwestward transport turns eastward to join the Gulf Stream, however, near 60oW. In addition to these steady transports, a combination of flow-relaxation and radiation conditions (Oey and Chen, 1992a,b) are used to also specify the WOAT/S and the M2-tideinterpolated from the Oregon State University’s tidal data [ at the open boundary at 55oW.Thus the vertical structures of the currents (i.e. after a depth-averaged value is removed) are specified using radiation conditions. The velocity component tangential to the boundary, as well as turbulence kinetic energy and length scale, are specified using one-sided advection scheme at outflow grids and are set zero at inflow. The (potential) temperature (T) and salinity (S) are similarly advected during outflow, but are specified using climatological values at inflow grids. Radiation is used for the surface elevation , but since POM uses a staggered C-grid, and because transports are specified, the boundary  plays only a minor role and a zero-gradient condition on it works well also. Sea surface fluxes are specified as detailed below. To prevent temperature and salinity drift in deep layers in long-term integration, theT and S for z < 1000 m are (weakly) restored to annual-mean climatological values with a time scale of 600 days; this does not impede short-period mesoscale variability. More details are in Oey et al. (2003). The NWAOM has been used for research primarily in the Gulf of Mexico where we have also extensively compared the results against observations both in the surface and subsurface (Oey and Lee, 2002;Ezer et al. 2003;Wang et al. 2003;Fan et al. 2004; Oey et al. 2005a,b,2006,2007;Lin et al. 2007;Yin and Oey, 2007;Oey, 2008; Wang and Oey, 2008;Mellor et al. 2008;Oey et al. 2009;Chang and Oey, 2010a,b).

For the present application to the MAB, the NWAOM is first run for 15 years, forced by monthly climatological NCEP surface fluxes. This 15-year run establishes a statistically equilibrium ocean field, as verified by examining the domain-averaged kinetic energy and eddy potential energy time series (not shown). This run is then continued by applying the CCMP six-hourly winds from Jan/01/1993 through 2008. Surface heat and evaporative fluxes are relaxed to monthly climatological values with a time scale of 100 days.

To calculate wind stresses, we use a bulk formula with a high wind-speed limited drag coefficient that curve-fits data for low-to-moderate winds (Large and Pond, 1981) and data for high wind speeds (Powell et al. 2003):

Cd103= 1.2, |ua|  11 m s-1;

= 0.49 + 0.065 |ua|, 11 < |ua|  19 m s-1;

= 1.364 + 0.0234 |ua|  0.00023158 |ua| 2, 19 < |ua|  100 m s-1

(2)

where |ua| is the wind speed.[1] According to this formula, Cd is constant at low winds, linearly increases for moderate winds, reaches a broad maximum for hurricane-force winds, |ua|  30~50 m s-1, and then decreases slightly for extreme winds. It is necessary to use a Cd formula that accounts for high winds since within the NWAOM domain, becausethe study period (1993-2008) includes a few hurricanes. Donelan et al. (2004) suggest that the Cd-leveling at high wind may be caused by flow separation from steep waves. Moon et al. (2004) found that Cd decreases for younger waves that predominate in hurricane-forced wave fields. Bye and Jenkins (2006) attribute the broad Cd-maximum to the effect of spray, which flattens the sea surface by transferring energy to longer wavelengths.

Daily discharges from 17 major sources in the MAB (and also from 33 sources in the Gulf of Mexico) are specified. These are specified as point sources at the “heads” of major bays or rivers using the method described in Oey (1996).Although broad bathymetric outlines and dimensions of bays and rivers are included (Fig.2), detailedestuarine circulation within them is not of interest for the purpose of this work. Their function is to allowa gradual transition of brackish waters onto the shelves. In other words, instead of inputting fresh river waters directly at the coast, they are allowed to mix (by tides and winds) with saline sea water within the bays or rivers before flowing out onto the continental shelves.

The northeastern corner of NWAOM domain is where the CLSW transport flows west-southwestward. We find that, for this model, a CLSW transport = 1.5 Sv is the minimum inflow that can give a “reasonable-looking,” time-mean separation and path of the Gulf Stream near Cape Hatteras (Mellor and Ezer, 1991; Ezer and Mellor, 1994a). The (mean) Gulf Stream tends to separate from the coast too far to the north for a CLSW inflow below this minimum. It is convenient to non-dimensionalize various values of the CLSW transport discussed below by this “critical” transport of 1.5 Sv, which we will refer to as “1UA” or UA=1. In the “standard experiment” (below), aCLSW transport of 4.5 Sv (i.e. 3UA or UA=3) is used, and this value will be adjusted (below) in sensitivity tests.The 3UA value may be compared with Csanady and Hamilton’s (1988) estimate of approximately 4 Svfor the CLSW transport.

Various experiments are carried out as discussed separately below (Table 1). Here we describe the standard experiment (Ex.DA) which consists of all the forcing and specifications described above. Additionally, satellite SSH anomaly data from AVISO ( also assimilated into the model. The purpose of this data-assimilative (DA) analysis is to provide a realistic open-ocean state – the Gulf Stream, rings and the Slope Sea gyre – to which the shelf then can respond. The Gulf Stream and eddies are assimilated in deep ocean regions only (water depth H1000 m) using the Mellor and Ezer’s [1991; see also Ezer and Mellor, 1994b] scheme. In this scheme, the SSH anomaly is projected into the subsurface density field using correlation functions pre-computed from the model’s eddy statistics derived from a non-assimilated 15-year model run. The method is simple, yet it yields fairly accurate upper-layer structures (z = 0 to approximately 800 m) of mesoscale currents and eddies [Oey et al. 2005a; Lin et al. 2007; Yin and Oey, 2007]. No assimilation is done in deep layers for z <800 mand as mentioned above in regions where the topography is shallower than 1000 m. In these regions, the simulated currents rely entirely on the model’s dynamics.

  1. Results

Model Sea Surface Height over the MAB shelf

Figure 2 and Figure 3a show the 16-year (1993-2008) mean SSH from the Ex.DA simulation. A cyclonic gyre is seen in the Gulf of Maine (e.g. Pettigrew et al., 2005). The cyclonic flow branches eastward and south-southwestward offCape Cod. The eastward branch flows anticyclonically over Georges Bank and then merges with the weaker south-southwestward branch over the shelf off Cape Cod. Figure 3a thus shows two local high pressure cells on the shelf, one over the Georges Bank, and a weaker one directly south of Cape Cod.It is clear that south of Cape Cod is where the sea level begins to slope down westward and southwestward along the entire length of the MAB shelf to Cape Hatteras. The SSH-contours tend to be across-shelf for water depths shallower than about 100 m, and over the shelf break and slope they are aligned along the isobaths.