1

 OCS Study 

 MMS 2002-053 

Northeastern Gulf of Mexico 

Circulation Modeling Study 

Final Report 

Authors 

Y. Hsueh 

Florida State University 

Robert Weisberg 

University of South Florida 

Prepared under MMS Cooperative Agreement 

14-35-0001-30804 

by 

Department of Oceanography 

Florida State University 

Tallahassee, FL 32306-4320 

Published by 

U. S. Department of the Interior 

Minerals Management Service Herndon, VA 

Gulf of Mexico OCS Region September 2002 
2

DISCLAIMER 

This report was prepared under a cooperative agreement between the Minerals 

Management Service (MMS) and the Florida State University. This report has 

been technically reviewed by the MMS and approved for publication. Approval 

does not signify that the contents necessarily reflect the views and policies of the 

Service, nor does mention of trade names or commercial products constitute 

endorsement or recommendation for use. It is, however, exempt from review 

and compliance with the MMS editorial standards. 

REPORT AVAILABILITY 

While supplies last, copies of the report may be obtained from: 

U.S. Department of the Interior 

Minerals Management Service 

Environmental Division 

381 Elden Street 

Herndon, VA 20170 

 Telephone Number: (703) 787-1642 

 

CITATION 

Suggested Citation: 

Hsueh, Y., and Robert H. Weisberg. Northeastern Gulf of Mexico 

Circulation Modeling Study: Final Report, MMS Cooperative 

Agreement 14-35-0001-30804. OCS Study MMS 2002-053, 

U.S. Dept of the Interior, Minerals Management Service, 

Herndon, VA. 187pp. 

ACKNOWLEDGEMENTS 

The Program Manager wishes to extend thanks to Prof. W. Sturges for numerous 

fruitful discussions, Prof. P. Niiler for providing surface drifter data, and Drs. P. 

Hamilton, and T. Berger for making available moored current meter data in 

DeSoto Canyon. Ms. Paula Tamaddoni Jahromi helped with preparing the pdf 

version and Dr. Reinard Harkema helped with the CD production. The 

continuing, enthusiastic and timely support of Dr. Walter Johnson, the MMS 

Contract Officer’s Technical Representative, during all phases of this project is 

gratefully acknowledged. 
3

Table of Contents 

Acknowledgements 2 

List of Figures 4 

Executive Summary 5 

1. Introduction 6 

2. FSU Modeling Effort 

 2a. Model Configuration 6 

 2b. Major Findings 8 

2c. Accomplishments 13 

3. USF Modeling Effort 

 3a. Introduction 14 

 3b. Major Findings 17 

 

3c. Accomplishments 20 

4. References 22 

5. Appendix 

 

 5a. A numerical model calculation of the flow in DeSoto 25 

 Canyon in response to northerly wind bursts in winter 

 5b. Temperature variability in the northeastern Gulf of 52 

 Mexico in 1997-1998 

 5c. Low-frequency variability of surface currents on the 78 

 Northeastern Gulf of Mexico (NEGOM) shelf 

 5d. West Florida shelf response to local wind forcing: 94 

April 1998 

 

 5e. West Florida shelf circulation and temperature budget 140 

 for the 1999 spring transition 
4

LIST OF FIGURES 

Figure No. Caption Page

1. (a) Topography of the Gulf-of-Mexico model. Contours 

shown are 100, 500, 1000, 2000, 3000m. Solid circles, 

F and G, mark locations where inflow-outflow model 

experiment time series are extracted. The solid square 

at the top marks Pensacola. (b) SAIC moorings. 

2. Evolution of temperature at the 5 m depth in the 

inflow-outflow experiment. The 200 m and 1000 m 

isobaths are plotted to identify the shelf area. The 

letters F and G mark the positions where velocity time 

series are extracted. In Panel 1, the anti-cyclonic 

and cyclonic eddies are marked by the letters A and 

C, respectively. 

3. Time series of north-south velocity component at 5 

m depth at G (84o

 W, 28o N) on the west Florida shelf in 

the inflow-outflow experiment. The upper panel shows the 

time series during more than five years of integration. The 

lower panel shows the time series between the two vertical 

lines in the upper panel. The expanded portion covers the 

time window of Figure 2. Numbers in the lower panel mark 

the time of the panels in Figure 2. The horizontal lines are 

zero velocity lines. 

4. Time series plot same as in Figure 3, but at F (87o W, 

29.5o

 N) inside DeSoto Canyon. Solid curve represents the 

east-west velocity component and dash curve represents the 

north-south component. The horizontal lines are zero 

velocity lines. 

5. USF model domain of DeSoto Canyon to Florida Keys. 

6. USF model domain of Mississippi River to Florida Keys. 
5

Executive Summary 

The project examines the flow on the continental shelf in the northeastern 

Gulf of Mexico (NEGOM) on the basis of two numerical models, each for a 

distinct portion of the shelf. The Princeton Ocean Model (POM) deals with the 

flow near the coast driven by local forcing and the Geophysical Fluid Dynamics 

Laboratory Model (GFDL) deals with that near the shelf-break driven by local and 

open-ocean forcing. 

 

The spin-up on the inner shelf follows the fast Ekman-geostrophic track 

and the presence of the Big Bend and the Florida Keys accentuates the 

response of the central west Florida coast water to local wind forcing. Baroclinic 

effects as a result of surface heating are important for seasonal variations in 

inner shelf circulation. 

The flow on the outer shelf is strongly influenced by the Loop Current. A 

southward shelf break jet is discovered, driven by pressure imposed by the Loop 

Current. The flow in DeSoto Canyon in winter is found to be an extension of the 

southward flow on the west Florida shelf set up by the passage of continental 

shelf waves. Monthly mean temperature variation in DeSoto Canyon features 

large cross-shelf gradients in May-June and November-December, following 

semi-annual peaking of strong anti-cyclonic wind-stress curl in NEGOM. 

Apparently, the wind-stress curl drives downward the isotherms in the deep water 

and the associated geostrophic currents to the east along the shelf-break give 

rise to bottom Ekman flow onshore, raising the isotherms in the shallow region. 

Thus, the large monthly mean temperature contrast in DeSoto Canyon. 
6

1. Introduction 

 The purpose of the cooperative program is to enhance the present 

understanding of the circulation in the Northeastern Gulf of Mexico through the 

use of computer models. The focus is on the wind-driven flows near the coast 

and the influence on them by the Gulf of Mexico Loop Current. The program 

involves a group of investigators from the Florida State University and the 

University of South Florida and parallels the field program directed by W. 

Sturges, as well as that conducted by the Science Applications International 

Corporation (SAIC). 

 There is a strong tie-in of the numerical modeling effort with the 

observational program. The observations in various forms, including that from 

satellite-tracked surface drifters, from moored current meters, from remote 

sensing, and from hydrographic surveys, collectively provide a basis for 

calibrating the performance of the models. Conversely, a well-calibrated model 

provides outputs that form a basis for the interpretation of the observations in 

terms of the dynamics. 

 The program has been a success in that useful models are developed and 

model/observation comparisons are concluded with positive results. 

2. 2. FSU Modeling Effort 

Prof. Y. Hsueh 

Department of Oceanography 

with 

 Scientific Associates: 

Drs. Dongliang Yuan and Yury Golubev 

Graduate Students 

Dr. Robert H. hetland and Howard Huang 

2.a Model Configuration 

 

 The FSU model is based on the Bryan-Cox code with z-levels in the 

vertical. In order to represent properly the Loop Current and its rings, the model 

encompasses the whole Gulf of Mexico, from 18o

 N to 31o N from 100oW to 

80oW. The horizontal grid size is one-sixth of a degree in both longitude and 

latitude. The vertical span is covered with a variable grid that is stretched in 

favor of a higher resolution in the surface layer than in the deep portion. 

 The geometry and topography are extracted from ETOPO5 and are shown 

in Figure 1. 
7

The circulation is initially driven with an inflow of 30 Sv across the 

southern open boundary in a 100 km wide boundary current next to the Yucatan 

Peninsula. An outflow of the same transport is implemented across the opening 

in the eastern boundary in the Florida Straits. 

 The model is initialized with the climatological temperature and salinity 

fields (Levitus et al., 1994). The initial velocity field is obtained from adding the 

thermal-wind velocity to the barotropic velocity forced by the imposed transports. 

Across the inflow and outflow ports, the tracer fields and the baroclinic velocity 

are relaxed to those from the Levitus climatology. At the sea surface, the first 

level temperature and salinity are relaxed to their climatological values. 

 Since the pioneering numerical work of Hurlburt and Thompson (1980) in 

which Loop Current ring shedding was reproduced for the first time, there have 

been several more models of the Gulf-of-Mexico basin-wide circulation (Dietrich 

et al., 1993, Oey, 1996; Welsh, 1996; Frantantoni, 1998). Except for the high 

resolution (7 km) model of Frantantoni (1998) used to study the Dry Tortugas 

eddies, these models, as well as that of Hurlburt and Thompson (1980), are 

designed to focus on the circulation in the western Gulf of Mexico and the 

evolution and role of the anti-cyclonic rings generated from the Loop Current. As 

Figure 1. (a). Topography of the 

Gulf-of-Mexico model. Contours shown are 

100, 500, 1000, 2000, 3000m. Solid circles, 

F and G, mark locations where inflowoutflow model experiment time series are 

extracted. The solid square at the top 

marks Pensacola. (b). SAIC moorings. 
8

such, they lack the vertical resolution necessary to resolve the continental shelf 

topography west of Florida. Although Oey (1996) uses improved vertical 

resolution to include a continental shelf, the simulated Loop Current bends 

unrealistically far away from the west Florida shelf. The effect of this bending is 

to decouple the Loop-Current forced circulation from the topographic influence of 

the west Florida shelf, a key factor in the formation of the southward flow on the 

west Florida shelf (see later). The present model avoids this excessive bending 

with the use of a biharmonic horizontal mixing formulation, which helps preserve 

the inertial current characteristics of the Loop Current. 

2b. Major Findings 

 

 A basic model run with no wind forcing is first made to establish the 

fundamentals of the Loop Current dominated circulation. The inflow-outflow 

condition produces the basic flow patterns with nearly periodic Loop Current ring 

separation at a period of around 13 months. Figure 2 shows, in eight panels, the 

evolution of surface temperature through a typical ring-shedding cycle of the 10-

year long model integration. The 200 m and 1000 m isobaths are also drawn to 

delineate the shelf area. On Day 1230 of the run (Panel 1), ring shedding is 

imminent. This is the time when the Loop Current has nearly achieved its 

maximum penetration into the Gulf of Mexico and the Loop-Current front is near 

its closest to the northwest Florida shelf (the Big Bend). A noticeable cyclonic 

eddy (marked by the letter C in Panel 1) is gaining in strength at this time 

southwest of Key West. This cyclonic eddy grows and moves to the west and 

causes a "necking-down" of the Loop Current, leading to the eventual separation 

of an anti-cyclonic ring (Vukovich et al., 1979; Pratt and Stern, 1986; Vukovich, 

1986). Panel 2 shows the temperature distribution two months later. A warmcore ring is clearly seen separated from the main stream of the Loop Current and 

the Loop Current has retreated to a minimum penetration state. In the meantime, 

the isolated ring moves to the west and southwest at a speed of about 4cm/s and 

is eventually dissipated off the Mexico coast (Panels 3-6) as is found elsewhere 

(Sturges et al., 1993). It takes about a year for the Loop Current to regain its full 

penetration and start another round of ring shedding (Panels 7-8). The ring 

shedding cycle repeats itself and constitutes the anticipated periodic solution for 

the Gulf of Mexico circulation (Pichevin and Nof, 1997). 
9

 Figure 2 Evolution of temperature at the 5 m depth in the inflow-outflow 

experiment. The 200 m and 1000 m isobaths are plotted to identify the shelf area. The 

letters F and G mark the positions where velocity time series are extracted. In Panel 1, 

the anti-cyclonic and cyclonic eddies are marked by the letters A and C, respectively.
10

The influence of this Loop Current evolution on the west Florida shelf 

circulation is shown in Figure 3. The upper panel shows the evolution of the 

north-south velocity component for more than 5 years at 5 m depth in 40 m of 

water at the solid circle marked G in Figure 1 (84o W, 28o N) off Tampa (similarly 

marked in Figure 2). The velocity at this location is qualitatively representative 

of the velocity on the west Florida shelf. The mean current over the time span is 

southward, suggesting that, on the average, the Loop Current generates a shelf 

flow to the south. The oscillation on the ring-shedding period appears phaselocked to the ring-shedding cycle. The time span of each trough (large 

southward velocity) roughly begins with the time when a ring has just separated 

(Panels 2 and 3 in Fig. 2). 

Figure 3. Time series of north-south velocity component at 5m depth at G (84° W, 28° N) 

on the west Florida shelf in the inflow-outflow experiment. The upper panel shows the time series 

during more than five years of integration. The lower panel shows the time series between the two 

vertical lines in the upper panel. The expanded portion covers the time window of Figure 2. 

Numbers in the lower panel mark the time of the panels in Figure 2. The horizontal lines are zero 

velocity lines. 
11

 The lower panel shows an expanded portion of the north-south velocity 

time series. The expanded portion corresponds to the time window spanned by 

Figure 2. Notice that, at the time (marked “1”) when the Loop-Current front is the 

closest to the Big Bend shelf, the southward drift over the shelf is at its minimum. 

This suggests that the flow over the shelf is not simply frictionally dragged by the 

Loop Current. The forcing of the Loop Current on the shelf circulation must be 

through other mechanisms. It turns out that the drop in dynamic pressure 

brought about by the Bernoulli effect as the Loop Current careens along the 

continental slope/shelf edge effects just such a flow in accordance with the 

arrested topographic wave (ATW) theory (Csanady, 1978; Csanady and Shaw, 

1983). The manner with which the pressure is imposed varies with the Loop 

Current position. The variation created in the shelf flow appears to explain the 

phase lead over the northward position of the Loop Current of southward flows 

inferred from the coastal sea-level fluctuations found in Sturges and Evans 

(1983). The detailed quasi-steady-state physics of this pressure forcing process 

forms the theme for a Ph. D. dissertation and is discussed in two publications 

(Hetland, 1999; Hetland et al., 1999; Hetland et al. 2001). Based on the results 

so far presented, it suffices to say that the Loop Current generally forces a 

southward flow over the west Florida shelf and the low-frequency variations of 

the flow occur primarily at the Loop Current ring-shedding frequency. 

 The influence of the Loop Current variation on the circulation inside 

DeSoto Canyon is shown in Figure 4, in which are plotted, over the same time 

period as in Figure 3, the velocity time series at 5 m depth in 324 m of water at 

the solid circle marked F in Figure 1 (87o

 W, 29.5o N) at the center of DeSoto 

Canyon (also marked similarly in Figure 2). The solid curve represents the eastwest velocity component and the dashed curve represents the north-south 

velocity component. The east-west component is particularly comparable to that 

shown in Figure 3, both being roughly alongshore. The east-west component 

time series in Figure 4 shows more irregularities, but the phase-locking with the 

Loop Current variations emerges again. Again, the maximum occur at the time 

when the Loop Current is farthest away to the south. The alongshore flow in 

DeSoto Canyon appears simply to respond to the southward flow over the west 

Florida shelf. This is again consistent with the ATW theory and attributable to the 

insulating effect of the continental slope topography, isolating the shelf flow, 

which roughly follows the isobaths. The east-west velocity component at F thus 

varies similarly with the north-south velocity component at G, 180o

 out of phase 

with the extent of the northward penetration of the Loop Current. A contour plot 

(not shown) of east-west velocity against time for the meridian through F from 26o

N to 30o

 N shows that not until the latitude of 27o N does the east-west velocity 

begin to vary in phase with the northward penetration of the Loop Current. 
12

 On the basis of the validity of the inflow-outflow experiment, a hindcast for 

the circulation is then conducted for the period of March 1997-April 1998 for 

which direct current measurements in DeSoto Canyon are available from a 

Science Applications International Corporation (SAIC) mooring experiment 

(SAIC, 1998). Use is made of six-hourly wind stresses derived from the 

NCEP/NCAR 40-year reanalysis winds (Kalnay et al., 1996). The hindcast 

results form the basis for two manuscripts. One manuscript, now in press, 

demonstrates that the flow in winter in DeSoto Canyon is but a part of the 

NEGOM continental shelf flow in response to the wind-driven southward current 

on the west Florida shelf (Hsueh and Golubev, 2001). The other, in review, 

Figure 4. Time series plot same as in Figure 3, but at F (87o W, 29.5o N) inside 

DeSoto Canyon. Solid curve represents the east-west velocity component and dash curve 

represents the north-south component. The horizontal lines are zero velocity lines. 
13

shows that the observed large DeSoto Canyon cross-shelf gradients in monthly 

mean temperature are the manifestation of shelf response to semi-annual 

peaking (in April and October) in the negative wind-stress curl field over NEGOM 

(Golubev and Hsueh, 2001). A third one has just been prepared and deals with 

the variability in surface currents surmised from the trajectory of the drifters. All 

three manuscript are attached in the Appendix. 

 In addition, for comparison purposes, the output is obtained for the Gulf of 

Mexico region from a similar model (Carton and Chao, 1999). An examination of 

the output has led to the discovery of a mechanism by which upwelled water from 

Campech Banks can propagate along the Loop Current to the west Florida shelf. 

The mechanism is one of frontal dynamics involving the Loop Current and 

imbedded eddies. The theoretical work form the basis of a Ph. D. dissertation 

research currently underway (Huang, 2001). 

2c. Accomplishments 

 The following is a list of papers and dissertations published, submitted, or 

in preparation. 

1). Hetland, Robert, Y. Hsueh, Bob Leben, and P. P. Niiler, 1999: A Loop 

Current-induced jet along the edge of the West Florida Shelf. Geophys. Res. 

Letters. 26, 2239-2242. 

 

2). Robert D., Y. Hsueh, and Dongliang Yuan, 2001: On the decay of a 

baroclinic jet flowing along a continental slope. J. Geophys. Res. 106, 19,797-

19807. 

3). Hsueh, Y., and Yury Golubev, 2002: A numerical model calculation of the 

flow in DeSoto Canyon in response to northerly wind bursts in winter. Gulf of 

Mexico Science. 20 (1), 44-59. 

4). Golubev, Yury, and Y. Hsueh, 2002: Temperature variability in the 

northeastern Gulf of Mexico in 1997-1998. Submitted to Gulf of Mexico Science. 

6). Golubev, Yury, and Y. Hsueh, 2002: Low-frequency variability of surface 

currents on the NEGOM shelf. In preparation. 

5). Hetland, Robert D., 1999: The dynamics of a Loop Current forced shelf 

break jet on the west Florida shelf. Dissertation. Department of Oceanography, 

Florida State University. 68 pp. 

6). Huang, Haosheng, 2002: Embedded eddies in the Loop Current. 

Dissertation. Department of Oceanography, Florida State University. In 

progress. 
14

3. USF Modeling Efforts 

Robert H. Weisberg 

Professor, Physical Oceanography 

With 

Scientific Associates: Drs. M. Luther, H. Yang, and Z. Li, 

Graduate Student: R. He, and 

Technical Staff: R. Cole, J. Donovan, C. Merz, and P. Smith. 

College of Marine Science 

University of South Florida 

140 7th Ave. S. 

St. Petersburg, FL. 33701 

3a. Introduction 

 Continental shelves may be discussed in terms of three regions, each 

controlled by different physical processes. The inner-shelf is where divergent 

surface and bottom frictional boundary layers play a major role; the outer-shelf 

(shelf-break) is where strong interactions occur with the adjacent ocean; and the 

middle-shelf is a transition region. Of these regions, the inner-shelf is the least 

studied, either through in-situ observations, or through the application of 

analytical and numerical models. Yet, in many respects, depending upon the 

shelf width, this may be the scientifically most interesting region because of its 

physical, biological, chemical, and geological interactions in response to synoptic 

and seasonal forcing. 

 The WFS is broad enough so that its inner, middle, and outer-shelf 

regions are relatively distinct, allowing the dominant processes within each 

region to be viewed somewhat independently from one another. Important innershelf scientific questions can therefore be addressed through the combination of 

in-situ measurements and models. These questions include the nature of: (i) the 

coastal jet's three-dimensional structure, time dependence, and turbulence 

evolution; (ii) its rectification by stratified boundary layer effects and how this 

impacts along and across-shelf transports; and (iii) the seasonal modulation of 

these processes by local and offshore forcing. The gently sloping WFS therefore 

provides an excellent natural laboratory for studying the physical processes that 

set the stage for biological (blooms), chemical (nutrient concentrations), and 

geological (sediment resuspension) events. 
15

 

At USF, we initiated a WFS circulation study in 1993 with USGS support 

that subsequently evolved with support from the MMS, the State of Florida, 

NOAA, and ONR. Presently ongoing is an Ecology of Harmful Algal Blooms 

(ECOHAB) regional field study to advance our understanding and forecasting 

ability for red-tides; a State of Florida supported Coastal Ocean Monitoring and 

Prediction System (COMPS) to provide real-time data for emergency 

management use; and ONR supported in-situ measurements and modeling. We 

adhere to the tenet that the in-situ measurements and modeling must be 

interactive. Measurements are sparse and require a multiple scale context. 

Models are critically dependent upon their parameterizations, initial state, and 

boundary conditions, so they must be pegged to measurements to be useful. 

Our evolving program of study was aided substantially by a cooperative 

agreement with the Minerals Management Service coordinated through the 

Florida State University, MMS 14-35-0001-30804. Over the four-year period: 

9/30/95-9/29/99, plus a one-year no cost extension, USF personnel, under the 

direction of R.H. Weisberg, engaged in numerical model experiments on the 

WFS circulation. We recognize that the shelf responds to both local and offshore 

forcing. Local forcing is by the combined effects of momentum and buoyancy 

inputs at the sea surface and buoyancy inputs at the coast. Offshore forcing is 

by momentum and buoyancy inputs from the deep ocean at the shelf-break. By 

virtue of its width, the WFS is an end member in the spectrum of continental 

shelves for which local forcing should be important. Thus our emphasis has 

been on local forcing and its effects on the inner-shelf. 

 

Our model domains have been regional ones, and the model that we 

chose to use is the Princeton Ocean Model (POM). Two different versions were 

employed. The first includes the domain between the DeSoto Canyon in the 

northwest and the Florida Keys in the southeast. As shown in Figure 5 it has 

three open boundaries for which radiation conditions are applied. Horizontal 

resolution for this curvilinear coordinate system model over the inner-shelf 

averages at about 9 km, and the model employs 16 sigma layers. The second 

encompasses the region from west of the Mississippi River to the Florida Keys. 

As shown in Figure 6 it has one open boundary governed by a radiation 

condition. Horizontal resolution for this curvilinear coordinate system model over 

the inner-shelf averages at about 2 km, and the model employs 21 sigma layers. 

Both model configurations employ realistic topography (ETOPO5 with some 

smoothing in the southeast where a spurious shoal is encountered). As an 

evolving modeling project we employed different forcing functions and 

initializations depending on the scientific questioned asked. The simplest density 

fields and forcing functions consisted of constant density and switched-on, 

spatially uniform winds. 
16

The most elaborate experiments used fully baroclinic density fields driven 

by spatially and temporally varying winds, surface heat fluxes, and river inflows. 

We also have an automated experimental nowcast/forecast system (available at 

http://ocg6.marine.usf.edu) using NCEP Eta model winds. 

Figure 5. USF model domain of DeSoto Canyon to Florida 

Keys. 

Figure 6. USF model domain of Mississippi 

River to Florida Keys.
17

3b. Major Findings 

Our approach has been to apply various versions of these POM-based 

models in a stepwise manner to develop an understanding on how the WFS 

responds to tidal, synoptic, and seasonal scale forcing. The first and simplest 

sets of experiments were designed to determine how the WFS responds to 

upwelling favorable winds directed either along-shore or offshore (relative to the 

west Florida shoreline). The Figure 5 model grid was used where the outermost 

open boundary is sufficiently offshore so that it does not interfere with the model 

responses over the continental shelf. While constant density and barotropic, the 

model runs are fully three-dimensional and directed toward understanding the 

response to synoptic scale winds (previous model experiments were either twodimensional or employed unrealistic geometries). Li and Weisberg (1999a) 

describes the model response kinematics and the companion paper, Li and 

Weisberg (1999b) describes the model response dynamics. Relevant findings 

are as follows. The model spin up follows an Ekman-geostrophic route resulting 

in a rapid adjustment over the course of a pendulum day. Surface Ekman layer 

divergence effects a sea surface slope that causes a geostrophic coastal jet that, 

in turn, causes a bottom Ekman layer response. While the WFS is gently 

sloping, subtle variations in coastline geometry have major effects on the fully 

three-dimensional circulation fields. In particular, partial closure in the south by 

the Florida Keys and the coastline curvatures of the Florida Big Bend conspire to 

yield strong responses to uniform winds directed either along-shore or offshore 

with respect to the central west Florida coastline. This may seem surprising 

based on two-dimensional arguments that offshore winds should not effect a 

coastal jet/upwelling response. Nevertheless, these findings are borne out by 

observations, and they are attributable to the coastline and isobath geometries. 

These geometry effects result in fully three-dimensional circulation patterns that 

are relevant to material property transports. The dynamics analyses define the 

inner-shelf as the region over which the sea surface slope sets up. We describe, 

in a vertically integrated sense, the interactions that occur between the alongshelf and across-shelf momentum balances in establishing the flow fields. The 

offshore scale of the inner-shelf is a frictional one associated with the divergence 

of the surface and bottom Ekman layers. This varies with bottom slope as shown 

for different regions along the coast. Offshore of the west Florida coastline, 

where we have corroborating data, the inner-shelf under a constant density 

setting extends out to about the 50-75m isobath. Fully three-dimensional 

momentum balance analyses provide further diagnoses of the model responses 

and how the Ekman and geostrophic regions interact. 

 Another preliminary examination was on the seasonal cycle. Yang and 

Weisberg (1999) asked the question of whether or not the seasonal variations in 

the circulation can be accounted for by seasonally varying winds alone. 

Climatological monthly mean winds were used to drive monthly model 

simulations. While interesting patterns are found suggesting that the WFS has 
18

wind-driven seasons of predominantly upwelling and downwelling circulations, 

monthly mean winds alone can not account fully for the seasonal variations in the 

circulation. Baroclinic effects must also be important, along with the surface heat 

fluxes that facilitate these. 

 Given the simultaneous surface drifter program we examined whether or 

not the model could be used to simulate the drifter tracks. Yang et al. (1999) 

compares observed Lagrangian surface drifter tracks with numerical model 

simulations. We demonstrate that model drifters can mimic in-situ drifters over 

time scales of the synoptic weather events, at least on the shelf where the model 

responses to these synoptic weather events exceed the effects of the Loop 

Current or eddies positioned on or offshore of the shelf-break. We also attempt 

to explain the curious finding that no drifters deployed to the north of the west 

central Florida shelf tracked toward the shoreline – hence the title of the paper. 

 A specific case study of an upwelling event recorded both in satellite 

AVHRR imagery and in-situ data was examined by Weisberg et al. (2000). While 

upwelling (and downwelling) occur regularly on continental shelves with the 

passage of each synoptic weather front, rarely do all the necessary ingredients of 

nature coincide to make the upwelling response visible in satellite imagery. In 

this case, the winds were light for several days prior to a strong, upwelling 

favorable wind event, and we had in-situ velocity measurements for comparison. 

Through a combination of data and model analyses we clearly demonstrate the 

Ekman-geostrophic spin-up route described earlier. Using both constant density 

and stratified runs (stratification estimated from velocity shear by thermal wind) 

we describe the relative effects and explain the finding of maximum upwelling 

just offshore and to the south of Tampa Bay. We also explain other regions of 

local upwelling maxima south of Cape San Blas. 

 The circulation physics highly influences the biology of the WFS. Walsh et 

al. (2001) is our first attempt at combining the physical models developed under 

this project with the ecological models developed by Dr. J. Walsh. Here we 

attempt to simulate the evolution of the 1979 red tide, providing some insights 

into the joint workings of the biology and physics that culminate in near shore red 

tide effects. This work is continuing under separate NOAA ECOHAB funding, 

and we appreciate the influence played by the MMS in helping to facilitate this 

continuing research. 

 Spring and summer of 1998 exhibited large stratification over the innershelf, even right up to the shoreline. In-situ data shows that the circulation is 

very sensitive to stratification. By separating the surface and bottom Ekman 

layers, stratification in increases across-shelf transports. Weisberg et al. (2001) 

describes the in-situ data and the numerical model simulation for the month of 

April, 1998 and, buoyed by the fidelity between the two, uses the model to 

analyze the dynamics. A new finding is identified. The data show a rectification 

of the inner-shelf responses to synoptic wind forcing wherein upwelling favorable 
19

winds produce disproportionately larger responses in both sea level and currents 

than downwelling favorable winds. Stratification accounts for the rectification, 

and this is most readily understood in terms of the streamwise component of 

vorticity. For downwelling favorable winds, the buoyancy torque due to 

isopycnals bending into the sloping bottom opposes the tendency by planetary 

vorticity tilting due to the vertically sheared coastal jet. This thermal wind effect 

negates the need for large relative vorticity dissipation by the across-shelf flow in 

the bottom Ekman layer. The opposite occurs for upwelling favorable winds. 

Buoyancy torque adds constructively with planetary vorticity tilting requiring 

larger dissipation of relative vorticity by the bottom Ekman layer. By enhancing 

(upwelling) or suppressing (downwelling) the bottom Ekman layer the entire 

response is reduced or increased, respectively. Such rectification is evident on 

the WFS because the WFS is wide enough to distinguish the frictional inner-shelf 

(which this paper also helps to better define) from the shelf break. 

 We are also applying the POM to three-dimensional analyses of hurricane 

storm surges along Florida's west coast. Yang and Weisberg (2000) reports on 

the effects of prototypical Saffir-Simpson Categories 2 and 4 hurricanes that 

approach Florida's west coast from several different directions and make landfall 

at several different locations. We show the space-time evolution of sea level 

along with sea level time series at specific points along the coast. Surge 

magnitude depends on local geometry and storm track, but for a given track it 

appears to scale approximately as the wind stress. Thus, for the recent 

"hurricane" Gordon we were able to surmise, from our real-time buoy winds (over 

which the storm closely passed) and model, that while the storm may have had 

hurricane winds aloft, it was not a hurricane on the ground. The ratio of the 

observed surge to that modeled (for a category 2 hurricane) scaled as the square 

of the observed and modeled wind speeds. Our results were disseminated to the 

NWS, emergency managers, and local weather reporters. 

 After laying the groundwork for interpreting our model results we are now 

using more realistic forcing functions to simulate longer-term WFS responses to 

local forcing. Our overall goal is to determine quantitatively the relative 

influences of local and deep ocean forcing on the west Florida shelf circulation on 

time scales from tidal to inter-annual and such forcing impacts material property 

distributions on the shelf. He and Weisberg (2001a) is a spring 1999 simulation 

in which the model, extending from west of the Mississippi River to the Florida 

Keys, is forced by NCEP reanalysis winds, surface heat fluxes (with flux 

correction), and river inflows (climatological). We show that local forcing, 

independent of the Loop Current, can account for the spring season transition 

and much of the synoptic scale variability. In preparation are similar manuscripts 

on other seasons (for which we have data) and on the inter-annual variations 

among these. This work will form the basis for a Ph.D. dissertation by R. He. 

 He and Weisberg (2001b) is an analysis of the four principal tidal 

constituents in the northeast Gulf of Mexico: M2, S2, O1, and K1. By forcing our 
20

regional model at its open boundary by a global tidal model, we produce tidal 

current ellipses and sea level variations for comparison with in-situ observations. 

On the basis of these comparisons we describe the tides from the Mississippi 

River to the Florida Keys. We also analyze the tidal residual flows. 

 Finally, MMS support facilitated the Ph.D. dissertation of Dr. Zhenjiang Li. 

Despite our inner-shelf emphasis we understand that offshore effects can 

permeate the inner-shelf in a variety of ways. We have not published specific 

results on this topic, but we did initiate collaborations with other groups that have 

larger scale models inclusive of the Gulf of Mexico Loop Current, and these 

collaborations are continuing. One way to approach this problem is to use the 

larger scale model fields to provide open boundary conditions for the regional 

scale model. Another way is to assimilate enough data directly into a regional 

model to maintain the baroclinic structure imposed by the deep ocean at the shelf 

break. We are exploring both approaches in our continuing research. Adequate 

initial fields and assimilative fields are required as are sufficient and accurate 

surface fluxes, by both in-situ measurements and atmosphere models, and we 

are progressing in each of these areas. 

3c. Accomplishments 

Refereed Publications 

1) Li, Z. and R.H. Weisberg (1999a). West Florida Shelf response to 

upwelling favorable wind forcing, Part 1: Kinematics. J. Geophys. Res., 104, 

13,507-13,527. 

2) Li, Z. and R. H. Weisberg (1999b). West Florida Shelf response to upwelling 

favorable wind forcing, Part 2: Dynamics. J. Geophys. Res., 104, 23427-23442. 

3) Yang, H. and R. H. Weisberg (1999). West Florida continental shelf circulation 

response to climatological wind forcing, J. Geophys. Res., 104, 5301-5320. 

4) Yang, H., R.H. Weisberg, P.P. Niiler, W. Sturges, and W. Johnson (1999). 

Forbidden zone over the West Florida Shelf, Cont. Shelf. Res., 19, 1221-1245. 

5) Weisberg, R.H., B. Black, Z. Li (2000). An upwelling case study on Florida’s 

west coast, J. Geophys. Res., 105, 11459-11469 

6) Weisberg, R.H., Z. Li, and F.E. Muller-Karger (2001). West Florida shelf 

response to local wind forcing: April 1998. J. Geophys. Res., 106, 31239-31262. 

7) Walsh, J.J. K.D. Haddad, D.A. Dieterle, R.H. Weisberg, Z. Li, H. Yang, F.E. 

Muller-Karger, C.A. Heil, and W.P. Bissett (2002). A numerical analysis of the 

landfall of the 1979 red tide of Karenia brevis along the west coast of Florida - the 

fall/winter upwelling mode. Cont. Shelf Res., 22, 15-38 
21

8) He, R. and R.H. Weisberg (2002a). West Florida Shelf circulation and 

temperature budget for the 1999spring transition. Cont. Shelf Res., in press. 

9) He, R. and R.H. Weisberg (2002b). Tides on the West Florida shelf. J. Phys. 

Oceanogr., submitted 

10) He, R. and R.H. Weisberg (2002c). A Loop current intrusion case study on 

the West Florida Shelf. J. Phys. Oceanogr., submitted. 

11) Weisberg, R.H. and R. He (2002). Local and deep ocean forcing 

contributions to anomalous water properties on the West Florida Continental 

Shelf. J. Geophys. Res. Submitted. 

12) Walsh, J.J., R.H. Weisberg, D.A. Dieterle, R. He, B.P. Darrow, J.K. Jolliff, 

K.M. Lester, G.A. Vargo, G.J. Kirkpatrick, K.A. Fanning, T.T. Sutton, A E. 

Jochens, D.C. Briggs, B. Nababan, C. Hu, and F. Muller-Karger (2002). The 

phytoplankton response to intrusions of slope water on the West Florida Shelf. J. 

Geophys. Res., submitted. 

Technical Reports and Dissertations 

1) Li, Z. (1998). Upwelling circulation on the west Florida continental shelf. 

Ph.D. dissertation, Department of Marine Science, University of South Florida, 

St. Petersburg, FL., 33701. 

2) Yang, H. and R.H. Weisberg (2000). A three dimensional numerical study of 

storm surges along the west Florida coast. COMPS Technical report, December 

2000. USF College of Marine Science, St. Petersburg, FL. 33701, 54pp. 

Representative Papers Appended 

1) Weisberg, R.H., Z. Li, and F.E. Muller-Karger (2001). West Florida shelf 

response to local wind forcing: April 1998. J. Geophys. Res., 106, 31239-31262. 

2) He, R. and R.H. Weisberg (2002a). West Florida Shelf circulation and 

temperature budget for the 1999spring transition. Cont. Shelf Res., in press. 

 In summary, we achieved our program objectives of developing a better 

understanding on how the WFS inner-shelf circulation responds to various 

forcing functions and the role of the seasonally varying stratification in modulating 

these responses. We published five papers in refereed scientific journals, three 

more are in press, and one is ready to be submitted. Additionally, we published 

a technical report on hurricane storm surge simulation and produced one Ph.D. 

The modeling work, greatly assisted by this MMS support, is presently continuing 

under NOAA/ECOHAB and ONR support. A preliminary version of a 

nowcast/forecast product is available at http://ocg6.marine.usf.edu. 
22

 The dissertation work of R. He (in progress) will further clarify the relative 

importance of local versus offshore forcing in determining the circulation of the 

NEGOM region. Our preliminary results show that much of the variability on both 

seasonal and inter-annual time scales can be accounted for by local forcing. We 

close this final report by noting that we will make additional important NEGOM 

contributions over the next year that had their beginnings with the funding herein. 

4. References 

Carton, James A., and Yi Chao, 1999: Caribbean Sea eddies inferred from 

TOPEX/ Poseidon altimetry and a 1/6o

 Atlantic Ocean model simulation. J. 

Geophys. Res., 104, 7743-7752. 

Csanady, G. T., 1978: The arrested topographic wave. J. Phys. Oceanogr. 8,47-

62. 

Csanady, G. T., and Ping Tung Shaw, 1983: The “insulating” effect of a steep 

continental slope. J. Geophys. Res., 88, 7519-7524. 

Dietrich, D. E., D.-S. Ko, and L. A. Yeske, 1993: On the application and 

evaluation of the relocatable Diecast ocean circulation model in coastal and 

semi-enclosed seas. Technical Report 93-1, Mississippi State University, 71pp. 

Frantantoni, Paula Sue, 1998: The formation and evolution of Tortugas Eddies in 

the southern Straits of Florida and Gulf of Mexico. Ph. D. Dissertation, University 

of Miami, Coral Gables, Florida. 

Golubev, Yury, and Y. Hsueh, 2001: Temperature variability in the northeastern 

Gulf of Mexico in 1997-1998. Submitted to Gulf of Mexico Science. 

He, R. and R.H. Weisberg, 2001a: The circulation and temperature budget on 

the west Florida continental shelf during the 1999spring transition. Cont. Shelf 

Res., in press. 

He, R. and R.H. Weisberg, 2001b: Tides on the west Florida shelf observed and 

modeled. Draft complete for submittal to J. Phys. Oceanogr. 

Hetland, R. D., 1999: The dynamics of a Loop Current forced shelf break jet on 

the west Florida shelf. Ph. D. dissertation, 69pp., Florida State University. 

Hetland, Robert, Y. Hsueh, Bob Leben, and P. P. Niiler, 1999: A Loop Currentinduced jet along the edge of the West Florida Shelf. Geophys. Res. Letters. 26, 

2239-2242. 
23

Hetland, Robert D., Ya Hsueh, and Dongliang Yuan, 2001: On the decay of a 

baroclinic jet flowing along a continental slope. J. Geophys. Res. 106, 19,797-

19807. 

Hsueh, Y., and Yury Golubev, 2001: A numerical model calculation of the flow in 

DeSoto Canyon in response to northerly wind bursts in winter. Gulf of Mexico 

Science., 20 (1), 44-59. 

Huang, Haosheng, 2001: Embedded eddies in the Loop Current. Dissertation. 

Department of Oceanography, Forida State University. In preparation. 

Hurlburt, H. E., and J. Dana Thompson, 1980: A numerical study of loop current 

intrusions and eddy-shedding. J. Phys. Oceanogr., 10, 1611-1651. 

Kalnay, E., M. Kanamitsu, R. Kistler, W. Collins, D. Deaven, L. Gandin, M. 

Iredell, S. Saha, G. White, J. Woollen, Y. Zhu, A. Leetmaa, R. Reynolds, M. 

Chelliah, W. Ebisuzaki, W.Higgins, J. Janowiak, K. C. Mo, 

C. Ropelewski, J. Wang, Roy Jenne, Dennis Joseph, 1996: The NCEP/NCAR 

40-year reanalysis project. Bulletin of the American Meteorological Society, 77, 

437-471. 

Li, Z. and R.H. Weisberg, 1999a. West Florida Shelf response to upwelling 

favorable wind forcing, Part 1: Kinematics. J. Geophys. Res., 104, 13,507-

13,527. 

Li, Z. and R. H. Weisberg, 1999b. West Florida Shelf response to upwelling 

favorable wind forcing, Part 2: Dynamics. J. Geophys. Res., 104, 23427-23442. 

Oey, L.-Y., 1996: Simulation of meso-scale variability in the Gulf of Mexico: 

sensitivity studies, comparison with observations, and trapped wave propagation. 

J. Phys. Oceanogr., 26, 145-175. 

Pichevin, T., and D. Nof, 1997: The momentum imbalance paradox. Tellus, 

49A, 298-319. 

Pratt, L. J., and Melvin E. Stern, 1986: Dynamics of potential vorticity fronts and 

eddy detachment. J. Phys. Oceanogr., 16, 1101-1120. 

Science Applications International Corporation, 1998: DeSoto Canyon eddy 

intrusion study. Data Products: Set 2 and Set 3. Science Applications 

International Corporation, Raleigh, North Carolina. 

Sturges, W., and J. C. Evans, 1983: On the variability of the Loop Current in the 

Gulf of Mexico. J. Mar. Res., 41, 639-653. 
24

Sturges, W., J. C. Evans, S. Welsh, and W. Holland, 1993: Separation of warmcore rings in the Gulf of Mexico. J. Phys. Oceanogr., 23, 250-268. 

Sturges, W., 1994: The frequency of ring separation from the Loop Current. J. 

Phys. Oceanogr.., 24, 1648-1651. 

Vukovich, F. M., B. W. Crissman, M. Bushnell, and W. J. King, 1979: Some 

aspects of the oceanography of the Gulf of Mexico via satellite and in situ data, J. 

Geophys. Res., 84,7749-7769. 

Vukovich, F. M., 1986: Aspects of the behavior of cold perturbations in the 

Eastern Gulf of Mexico: a case study. J. Phys. Oceanogr., 16, 175-188. 

Walsh, J.J. K.D. Haddad, D.A. Dieterle, R.H. Weisberg, Z. Li, H. Yang, F.E. 

Muller-Karger, C.A. Heil, and W.P. Bissett, 2001. A numerical analysis of the 

landfall of 1979 red tide of Gymnodinium breve along the west coast of Florida - 

the fall/winter upwelling mode. Cont. Shelf Res., in press. 

Weisberg, R.H., B. Black, Z. Li, 2000. An upwelling case study on Florida’s west 

coast, J. Geophys. Res., 105, 11459-11469. 

Weisberg, R.H., Z. Li, and F.E. Muller-Karger, 2001: West Florida shelf response 

to local wind forcing: April 1998. J. Geophys. Res., in press. 

Welsh, S. E., 1996: Numerical simulation of Gulf of Mexico circulation under 

present and glacial climate conditions. Ph. D. dissertation, Louisiana State 

University, 146pp. 

Yang, H., R.H. Weisberg, P.P. Niiler, W. Sturges, and W. Johnson, 1999. 

Forbidden zone over the West Florida Shelf, Cont. Shelf. Res., 19, 1221-1245. 

Yang, H. and R.H. Weisberg, 2000: A three dimensional numerical study of 

storm surges along the west Florida coast. COMPS Technical report, December 

2000. USF College of Marine Science, St. Petersburg, FL. 33701, 54pp.. 
25

5. Appendix 

5a. A numerical model calculation of the flow in DeSoto Canyon in response 

to northerly wind bursts in winter 
26

A Numerical Model Calculation of the Flow in DeSoto Canyon 

In Response to Northerly Wind Bursts in Winter 

by 

Y. Hsueh and Yury Golubev 

Department of Oceanography 

Florida State University 

Tallahassee, FL 32306 

Abstract

The continental shelf currents observed at the head of DeSoto Canyon 

offshore of Pensacola, Florida, during November 1997-February 1998 are 

studied with the aid of a Bryan-Cox model of the entire Gulf of Mexico. The basic 

model circulation with no winds features a Loop Current and a weak flow in 

DeSoto Canyon. In contrast, the model response to northerly wind bursts 

produces strong Canyon currents. Wind-driven southward currents on the West 

Florida Shelf (WFS) appear to pull a strong flow around DeSoto Canyon, giving 

rise to a strong up-canyon flow on the west flank. The wind-enhanced flow in 

DeSoto Canyon is reproduced with a model run that incorporates wind stresses 

derived from the NCEP/NCAR reanalysis winds for March 1997-April 1998. The 

model flow for the winter period (November 1997-February 1998) is compared to 

observed currents at three canyon-head locations in 100m of water available 

from a Science Applications International Corporation (SAIC) mooring 

experiment. Model and observed velocities in the general around-Canyon 

direction track each other reasonably well with peaks synchronized roughly with 

but lagging slightly northerly wind bursts. The lag time appears consistent with 

the generation at Key West of first-mode continental shelf waves responsible for 

wind-driven alongshore currents on the WFS. Pressure gradient slightly in 

excess of that necessary for a geostrophic balance with the (eastward) 

alongshore flow appears to generate onshore flows, giving rise to coastal 

upwelling following northerly wind bursts. 
27

1. Introduction 

 The wintertime, wind-driven continental shelf currents in DeSoto Canyon 

are studied using a Bryan-Cox model of the entire Gulf of Mexico. DeSoto 

Canyon is a broad, northward tending submarine canyon south of the Florida 

Panhandle coast in the northeastern Gulf of Mexico (NEGOM) with the 200m 

isobath just 72km away from the coast (at approximately 30o

24’N) (Figure 1a). 

The width of the canyon measured at the 100m isobath along 29o 30’N is about 

120km, much greater than the 24km internal deformation radius based on 

climatology. It is known that warm Gulf-of-Mexico water can move shoreward 

along the western rim of the canyon to within 8 km of the shore during northerly 

wind bursts in winter (Huh et al., 1981). Additional current measurements have 

since become available that bear upon the intrusion issue. The additional data 

come from the Science Applications International Corporation (SAIC) mooring 

experiment of March 1997 - April 1998, which includes three head-of-the-canyon 

moorings along the 100m isobath, C1 (87o21.0’W, 29o35.15’N), D1 (86o 50.5’W, 

30o04’N), and E1 (86o19.73’W, 29o42.0’N) (see Figure 1b). The purpose of the 

present paper is to present a comparison of the model current with these 

measurements for the wintertime intrusion flow in DeSoto Canyon and to 

suggest, on the basis of model dynamics, that the intrusion flow may be a winddriven response of continental shelf water in NEGOM as a whole. This basic 

response seems to explain the long suspected episodic intrusion of deep Gulf 

water along the western rim of the Canyon (Schroeder et al., 1994). 

 

Figure 1. (a). Topography of the 

Gulf-of-Mexico model. Contours shown are 

100, 500, 1000, 2000, 3000m. Solid circles, 

F and G, mark locations where inflowoutflow model experiment time series are 

extracted. The solid square at the top 

marks Pensacola where the model windstress time series are shown. (b). SAIC 

moorings. Data from C1, D1, and E1 form 

the basis for model verification. 
28

 Submarine canyon flows have been studied on both coasts of the United 

States mostly with canyon width less than or on the order of the internal 

deformation radius (Klink, 1996; Hickey, 1997; She and Klinck, 2000). The winddriven continental shelf flow in the vicinity of such deep and narrow canyons 

exhibits characteristics consistent with intense generation of relative vorticity due 

to vortex stretching (Hickey, 1997). In the case of broad canyons, as is the case 

at hand where the deformation radius is much less than the radius of curvature of 

the isobaths, the dynamics are predominantly linear (Klinck, 1996). In this case, 

the flow is expected to go approximately along isobaths and to exhibit the familiar 

linear dynamics (Hsueh, 1980). In the concept of such a flow, DeSoto Canyon 

thus becomes an extension of the WFS. It has been shown that wind-driven sealevel heights over the WFS are reasonably explained by the linear long-wave 

theory (Mitchum and Clarke, 1986). It stands to reason that the flow in DeSoto 

Canyon is just a part of this wave response. In addition, on the long-term, 

Sturges and Evans (1983) find that the 8-month band sea-level fluctuations along 

the west Florida coast interpreted as southward currents are coherent with, but 

lead, the northward penetration positions of the Loop Current, indicating Loop 

Current influence. 

 To study the intrusion in DeSoto Canyon with both these forcing 

possibilities, a model of the entire Gulf of Mexico is constructed (see Figure 1a 

for the model domain). The model is of a Bryan-Cox type (Cox, 1984) with a 

rigid-lid with increased vertical resolution in the upper water column to resolve 

the topography of the continental shelf and slope west of the Florida peninsula. 

(For future reference, the rigid-lid assumption produces a slight error in the phase 

speed of the continental shelf wave. The error is on the order of the square of 

the ratio of shelf width to the barotropic deformation radius, thus quite negligible 

even for a shelf as wide as the WFS (Clarke, 1977).) The forcing in the model 

includes wind stress at the surface and open boundary conditions in the 

Caribbean Sea and in the Straits of Florida. The southern boundary of the model 

cuts through the Caribbean Sea so that the structure of the inflow through the 

Yucatan Straits is not specified as in Hurlburt and Thompson (1980). A volume 

transport of 30 Sv is imposed across the southern open boundary and through 

the Straits of Florida to represent the mean Loop Current transport. No attempt 

has been made to study the model sensitivity to the flow specification at these 

open boundaries. The purpose is to achieve a Gulf of Mexico model with a 

broad-scale representation of the circulation to serve as background for a few 

process studies in continental shelf waters. 

 

 Since the pioneering numerical work of Hurlburt and Thompson (1980) in 

which Loop Current ring shedding was reproduced for the first time, there have 

been several more models of the Gulf-of-Mexico basin-wide circulation (Dietrich 

et al., 1993, Oey, 1996; Welsh, 1996; Frantantoni, 1998). Except for the high 

resolution (7 km) model of Frantantoni (1998) used to study Dry Tortugas eddies, 

these models, as well as that of Hurlburt and Thompson (1980), are designed to 
29

focus on the circulation in the western Gulf of Mexico and the evolution and role 

of the anti-cyclonic rings generated from the Loop-Current. As such, they lack 

the vertical resolution necessary to resolve the continental shelf topography west 

of Florida. Oey (1996) uses improved vertical resolution to include a continental 

shelf and successfully simulates the irregular shedding of Loop Current rings. 

The Loop Current and the migration path of its rings in Oey (1996), as in the 

present model, bend unrealistically to the west, hinting at an overly diffusive 

representation due perhaps in part to a lack of horizontal resolution. Even 

though the present model uses a biharmonic formulation, the flow is still too 

frictional so that there is a complete absence of filaments and parasite eddies 

often observed around the Loop Current (Perez et al., 1999) and the model Loop 

Current never reaches so high a latitude as the observed 28o

N (Sturges, 1992). 

 The outline of the paper is as follows. In the next section, a description is 

given of the model configuration and then in section 3 are reported model flows 

driven by the Loop Current. Section 4 considers the response of the WFS and 

the DeSoto Canyon waters to winds during November 1997- February 1998, the 

SAIC experiment winter period. The dynamics of the flow are discussed in 

section 5. A final section 6 contains the conclusions. 

2. The Model 

 As mentioned above, the model used in this study is the z-level model of 

Bryan-Cox’s (Cox, 1984). The model domain covers the entire Gulf of Mexico, 

from 18o N to 31oN and from 100o W to 80o W, with topography obtained from the 

5-minute ETOPO global relief model (see Figure 1a). The horizontal grid 

resolution is 1/6o by 1/6o. The total number of levels in the vertical is 30, with 

larger vertical spacings in the deeper ocean. The smallest vertical spacing is 10 

m at the sea surface, small enough only to give the Ekman veering and transport 

crucial for the Ekman-geostrophic response in coastal waters, which is more 

important here than the detailed Ekman layer structure not resolved. The 

number of levels above the 200 m depth is 8 and the number of levels above the 

1000 m depth is 16. With this grid, both the continental shelf and the continental 

slope are reasonably resolved. 

 At the southern open boundary across the Caribbean Sea, a 30 Sv (1 

Sv=106

m3s-1) volume transport is specified across a 100 km section next to the 

Yucatan Peninsula to represent nominally the western boundary current of the 

Caribbean Sea. The same volume transport leaves the model domain through a 

100 km section in the northern Straits of Florida at the eastern open boundary. A 

five-grid-interval buffer layer is used along the open boundaries to relax salinity 

and temperature fields to their climatological values (Levitus et. al., 1994; Levitus 

and Boyer, 1994) and the velocity to the geostrophic velocity. The geostrophic 

velocity at the open boundaries is obtained by combining the barotropic velocity 

based on the specified transport with the thermal-wind velocity calculated from 

the climatological tracers fields. The relaxation coefficients at the open 

boundaries correspond to time scales of a few hours. Weak relaxation with a 
30

time scale of 50 days is used at the sea surface for temperature and salinity. 

Radiation boundary conditions for temperature, salinity, and velocity similar to 

that of Carmelengo and O'Brien (1980) are used on the eastern open boundary. 

The strategy of the open-boundary condition specification has been explained in 

detail in Schultz (1994) and in Hsueh et al. (1997). It must be noted that the 

inflow and outflow conditions are probably far too simple to represent the full 

reality. Again, the aim here is to achieve in a broad sense an inflow-outflow 

configuration for the Gulf of Mexico for process studies. The effect of open 

boundary conditions on the flow is obviously a worthy study subject on its own 

right but is outside the present scope. 

 At closed boundaries, the no-slip velocity and no-normal-flux temperature 

and salinity conditions are employed. A biharmonic horizontal mixing formula is 

used in the momentum and tracer equations with a biharmonic coefficient of –

4x1018cm4

s-1. The use of a biharmonic formulation is intended to reduce 

damping and preserves the nonlinear characteristics of inertial currents (Holland, 

1978). (The model, however, still proves too diffusive (see later).) Vertical 

mixing and diffusion are embodied in a Laplacian with a coefficient of 1cm2

s-1. 

The quadratic-law bottom friction is used in the momentum equations with a 

friction coefficient of 10-3. The model is initialized with climatological temperature 

and salinity fields and with a geostrophic velocity field obtained from adding the 

climatological thermal-wind velocity to the barotropic velocity forced by the 

imposed transports (Hsueh et al., 1997). 

3. An Inflow-Outflow Experiment 

 In order to assess the influence of the Loop Current on the circulation over 

the WFS and inside DeSoto Canyon, a run (referred to as GOM-01) is first made 

of the model without the wind stress. Under the forcing of the open-boundary 

conditions, the inflow-outflow combination gives rise to a Loop Current in the Gulf 

of Mexico, which sheds anti-cyclonic rings of 300-400 km radius at a period of 

around 13 months. The lack of a spread in separation periods points to the fact 

that the model leaves out important factors which contribute to the Loop Current 

behavior and ring separation (Sturges, 1994; Sturges and Leben, 2000). It does 

however conform to what is anticipated in an analytical theory (Pichevin and Nof, 

1997). 
31

(a) 

(b) 

Figure 2 Evolution of velocity at the 5 m depth in the inflow-outflow 

experiment in eight panels (panels 1-4 in Figure 2a and panels 5-8 in 

Figure 2b) dated in integration days and numbered in the upper right 

corner. The 200 m and 1000 m isobaths are plotted to identify the shelf 

area.
32

 Figures 2a and 2b show, in eight panels, the evolution of the 5m velocity 

over NEGOM through a typical ring-shedding cycle of the model integration. On 

Day 2070 of the run (Panel 1), an anti-cyclonic ring is clearly seen separated 

from the main stream of the Loop Current and the Loop Current has retreated to 

near its southernmost position. The isolated current ring proceeds to move to the 

west and southwest at a speed of about 4cms-1 as is found elsewhere (Sturges 

et al., 1993). In the meantime, the Loop Current begins to grow northward. In 

early stages in this growth (Panels 3-4), a shelf-break jet is seen formed, 

apparently from the impingement of the Loop Current on the WFS (see later). As 

the Loop Current reaches nearly its maximum northward penetration with its axis 

at its closest to the WFS (Panels 5-6), a weak cyclonic eddy is gaining in strength 

southwest of Key West. This cyclonic eddy grows and moves to the west and 

causes a "necking-down" of the Loop Current, leading to the eventual separation 

of an anti-cyclonic current ring (Panels 7-8) (Vukovich et al., 1979; Pratt and 

Stern, 1985; Vukovich, 1986). The spawning of this cyclonic circulation cell and 

its possible triggering effect for ring separation have been widely acknowledged 

(Hurlburt and Thompson, 1982; Dietrich and Lin, 1994; Frantantoni, 1998). It 

takes about a year for the model Loop Current to regain its full penetration and 

start another round of ring shedding. The model ring shedding cycle repeats 

itself and constitutes the anticipated simplified periodic solution found in 

theoretical studies (Pichevin and Nof, 1997). In reality, the ring shedding is quite 

aperiodic and the appearance of the Loop Current is irregular with filaments and 

parasite eddies, exposing the overly diffusive nature of the present model run 

(see Oey, 1995 for a successful model in this regard). 

 
33

 The influence of this simple model Loop Current evolution on the WFS 

circulation is shown in Figure 3. The upper panel shows the evolution of the 

north-south velocity component (dashed) (the solid curve is for the east-west 

velocity which is much weaker) for more than 5 years at the 5 m depth in 40 m of 

water at the solid circle marked G in Figure 1a (84o

 W, 28o N) off Tampa. The 

velocity at this location is qualitatively representative of the velocity on the WFS. 

The mean current over the time span is southward, suggesting that, on the 

average, the Loop Current generates a shelf flow to the south. The oscillation on 

the ring-shedding period appears phase-locked to the ring-shedding cycle. The 

time span of each trough (large southward velocity) roughly begins with the time 

when a ring has just separated (Panels 2 and 3 in Fig. 2). 

 The lower panel shows an expanded portion of the north-south velocity 

time series. The expanded portion corresponds to the time window spanned by 

Figure 2. Notice that, at the time (marked 2370) when the Loop-Current front is 

closer to the WFS, the southward drift over the shelf is weaker than, say, that on 

Day 2250 when the front is farther away. This suggests that the flow over the 

Figure 3 Time series of north-south velocity component (dashed) at 5 m 

depth at 84o

 W, 28o N (G in Figure 1a) on the WFS in the inflow-outflow 

experiment. The upper panel shows the time series during more than five years 

of integration. The lower panel shows the boxed portion of time series in the 

upper panel. The expanded portion covers the time window of Figure 2. The 

horizontal lines are zero velocity lines. 
34

shelf is not simply generated frictionally by the Loop Current. The forcing of the 

Loop Current on the shelf circulation must be through other mechanisms. It turns 

out that the drop in dynamic pressure brought about by the Bernoulli effect as the 

Loop Current careens along the continental slope/shelf edge after it collides with 

the continental mass effects just such a flow in accordance with the arrested 

topographic wave (ATW) theory (Csanady, 1978; Csanady and Shaw, 1983). 

The manner with which this along-shelf-edge pressure change is imposed varies 

with the Loop Current position. This change in shelf-edge pressure as a function 

of the Loop Current position induces a corresponding change in shelf flow 

making the southward flow to appear to lead the northward position of the Loop 

Current (Sturges and Evans,1983). The detailed steady-state physics of this 

pressure forcing process is discussed elsewhere (Hetland et al., 1999). Based 

on the evidence so far presented, it suffices to say that the model Loop Current 

generally forces a southward flow over the WFS and the low-frequency variations 

of the flow occur primarily at the Loop Current ring-shedding frequency. 

 

Figure 4 Time series plot same as in Figure 3, but at 87oW, 29.5oN (F in 

Figure 1a) inside DeSoto Canyon. Solid curve represents the east-west velocity 

component and dash curve represents the north-south component. The horizontal 

lines are zero velocity lines. 
35

The influence of the model Loop Current variation on the circulation inside 

DeSoto Canyon is shown in Figure 4, in which are plotted, over the same time 

period as in Figure 3, the velocity time series at 5 m depth in 324 m of water at 

the solid circle marked F in Figure 1a (87o W, 29.5o N) at the center of DeSoto 

Canyon. The solid curve represents the east-west velocity component and the 

dashed curve represents the north-south velocity component. The east-west 

component is particularly comparable to the north-south component shown in 

Figure 3, both being roughly alongshore. The east-west component time series 

in Figure 4 shows more irregularities, but the phase-lock with the Loop Current 

variation emerges again. Again, larger velocity tends to occur at a time when the 

Loop Current is farther away to the south (Day 2250 vs. Day 2310). The 

alongshore flow in DeSoto Canyon appears simply to respond to the southward 

flow over the WFS. This is again consistent with the ATW theory and attributable 

to the insulating effect of the continental slope topography, confining the shelf 

flow to follow roughly the isobaths. The east-west velocity component at the 

solid circle F thus varies similarly with the north-south velocity component at G, 

180o

 out of phase with the extent of the northward penetration of the Loop 

Current. A time-distance contour plot (not shown but see 

www.ocean.fsu.edu/oce/gom) of east-west velocity at 5m along the meridian 

through F from 26o N to 30o N shows that not until the latitude of 27o N is reached 

does the east-west velocity begin to vary in phase with the northward penetration 

of the Loop Current. 

 

Figure 5 Horizontal velocity charts at 16.4m in the 

DeSoto Canyon area in the inflow-outflow experiment. The 

times of the panels correspond to those in Figure 2. The 100 

m isobath is drawn to identify DeSoto Canyon. 
36

 Figure 5 shows the velocity vector plots at 16.4m in the DeSoto Canyon 

area at the time of the eight panels in Figure 2. The Loop Current-driven flow 

inside DeSoto Canyon remains weak and basically follows the isobaths. Since 

the observed intrusion in DeSoto Canyon suggests strong currents, causes in 

addition to the low-frequency influence of the Loop-Current must be found to 

explain the intrusion. It must be stressed that the results of weak direct Loop 

Current influence in both the WFS and DeSoto Canyon are probably a model 

artifact due to the inability of the model in resolving Loop Current filaments and 

parasite eddies. Model runs with better resolution and mixing representation 

than used in the present run are needed to address the influence of these smallscale Loop Current structures. With their influence suppressed, the model 

presents an opportunity for studying, in almost a default isolation, other factors 

influencing continental shelf flows in NEGOM, such as the wind. 

4. A Hindcast for November 1997 – February 1998 

 Direct current measurements made by Science Applications International 

Corporation (SAIC) in NEGOM, particularly at the head of DeSoto Canyon for 

March 1997 - April 1998 become recently available. To take advantage of the 

opportunity thus provided, a hindcast (GOM-09) for currents is conducted with 

the imposition at the model sea surface of 6-hourly wind stresses converted from 

the 10 m surface winds from the NCEP/NCAR 40-year reanalysis data (Kalnay et 

al., 1996), using a drag coefficient of 1.3x10-3. (Time series plots (not shown but 

see www.ocean.fsu.edu.edu/oce/gom) of NCEP/NCAR winds together with those 

observed at Pensacola and at a number of buoys in the northeastern Gulf of 

Mexico show that the two track each other closely most of the time.) The initial 

condition is the GOM-01 output at Day 2070 (Panel 1 in Figure 2). A 

TOPEX/Poseidon sea-surface height anomaly map for March 1, 1997 (not shown 

but see www.ocean.fsu.edu.edu/oce/gom) shows approximately the same Loop 

Current position as in the chosen frame of the GOM-01 output. 
37

 Figure 6 shows the model wind-stress time series at Pensacola for the 

selected hindcast period. The middle panel shows the model north-south wind 

stress at a point south of G at (84.45o

W, 27.07oN). A number of northerly wind 

bursts are clearly identifiable (November 7, 15, 23, December 5,12-15, 27, and 

February 4, for example). Each burst lasts for several days and then collapses. 

Figure 6 Plot of the model-used 6-hourly wind stress at Pensacola in November 

1997 – February 1998. The upper panel shows the time series of the east-west (thin) 

and north-south (heavy) components of the wind stress. The lower panel shows the stick 

plot of the daily averaged Pensacola wind stress. Times of northerly wind bursts 

discussed are marked with heavy tick marks. In the middle panel, the model north-south 

wind stress component time series at a point (84.45o

W, 27.07oN) on the WFS south of G 

is shown. 
38
39

Figure 7 presents comparison between the hindcast and the inflow-outflow 

experiment (GOM-01) in terms of DeSoto Canyon area velocity vector plots at 

the time about one to two days later of several of these wind burst events (with 

Day 2070 in GOM-01 as March 1, 1997). The hindcast clearly produces strong 

onshore flows on the west flank of DeSoto Canyon not otherwise present. To 

verify this hindcast model response, use is made of SAIC direct current velocity 

measurements during November 1997 – February 1998 at C1 (87o

21.0’W, 

29o

35.15’N), D1 (86o 50.5’W, 30o 04’N), and E1 (86o19.73’W, 29o42.0’N) (see 

Figure 1b). 

Figure 7 Model velocity vector plots at 16m in the DeSoto Canyon area 

during northerly wind bursts. Plots are for November 8, 17 (Figure 7a), 

December 7, 16 (Figure 7b), December 31 and February 5 (Figure 7c). For 

each case, the corresponding plot from the inflow-outflow experiment is shown 

in the top panel. The plot from the hindcast is in the middle. The plot of the 

vertical velocity in contours of 2x10-3cms-1 appears in the bottom panel, where 

the zero contour is dashed and negative values are contoured in dashed 

curves. The 100m isobath is drawn to identify DeSoto Canyon. 
40

(a)
41

Figure 8 shows time series comparisons, at two interior depths 16 and 

32m, of the around-canyon and crosswise velocity components from data with 

the corresponding calculation from the model. (Because of the bottom slope, the 

depth of 48m represents the last depth for which model velocities are calculated 

at D1.) The local around-canyon directions (positive y-axis) are 45o

, 75o, and 

170o

 clockwise from north for C1, D1, and E1, respectively. The crosswise 

direction (positive x-axis) is generally offshore. At all locations, there is a 

measure of similarity in the around-canyon component between the observation 

and the hindcast. In particular, the strong model flow as indicated by the aroundcanyon velocity peaks shortly after northerly wind bursts of November 7, 15, 23, 

December 5, 12-15, 27, and February 4 seems well substantiated by the direct 

Figure 8 Time series plots of (a) around-canyon and (b) crosswise velocity 

components at 16m and 32m at C1, D1, and E1 for November 1997 – February 

1998. The model results are dashed and observations are solid. The wind stress 

component time series shown in Figure 6 is reproduced in the upper panel for easy 

comparison. 

(b)
42

measurements. In all events, the current lags slightly the wind. This is 

consistent with the notion that the DeSoto Canyon flow simply responds to the 

southward flow on the WFS which is established by the passage of continental 

shelf waves originated at the end of the WFS following the onset of each wind 

burst (Mitchum and Clarke, 1986). 

Figure 9. Coherence squared ( 2 γ ) and phase (θ ) with cycles 

per day (cpd) between the 6-hourly north-south wind component and 6-

hourly around-canyon velocity at station D1 at the depth of 16m for (a) 

modeled currents and (b) observed currents. The dashed line in the 

coherence squared plots indicates the 95% confidence level. Negative 

phase indicates the wind leads. Phase is calculated only for the 

frequency band in which the coherence squared is above the 95% 

confidence level. 
43

 Figure 9 shows the coherence squared and phase between the northsouth wind component and the around-canyon current at station D1 at the depth 

of 16m. In the three to ten-day band, a lag of a couple of days is detected for the 

around-canyon currents, both observed and modeled, consistent with the first 

mode continental shelf wave phase speed for the WFS (Mitchum and Clarke, 

1986)). There is little similarity in the crosswise component, which is much 

weaker. The difficulty is typical of that encountered in modeling cross-shore 

flows on the continental shelf (Hsueh, 1989). 

 There remains the question of the degree to which the around-canyon flow 

is locally driven by the east-west wind bursts of which there are several at the 

time of north-south wind bursts (for example, December 31 and February 4). In 

order to address this issue, three model runs with a constant wind stress to the 

southeast (at 1 dyne/cm2

 for the components) are conducted from a common 

initial condition: one with the wind over the entire Gulf of Mexico, one with the 

wind covering only the domain east of 87oW, and one with the wind covering the 

domain east of 88oW with the wind stress ramping from zero at 88oW to the 

constant value of 1dyne/cm2

 at 87oW for both components. Against these runs, a 

control run is made with no wind forcing. In each case the model is run for 15 

days. (The continental shelf flow pattern becomes steady at about the 10-day 

mark.) The around-canyon velocity at station D1 at 32m from the run with no 

wind is then subtracted from that from the three experimental runs. The winddriven around-canyon velocity thus obtained differs very little (3 cm/s in 40 cm/s), 

with the one from the experiment with the forcing ramp the greatest. This result 

indicates empirically that the local wind-driven around-canyon velocity is small. 

Theoretically, one could estimate the amplitude of the first-mode continental shelf 

wave at D1 by integrating along the characteristics (in this case, the 100m 

isobath) the along-characteristics wind-stress from Key West to D1 (see, for 

example, Mitchum and Clarke, 1986). Since the east-west running portion of the 

isobath is relatively short compared with the north-south running portion on the 

WFS, it stands for reason that the amplitude contribution from east-west wind 

forcing in the neighborhood of D1 will be small. A few words for the event around 

February 4 are in order. Although, the winds at Pensacola are northerly, those 

on WFS are actually southerly a couple of days earlier (see Figure 6, middle 

panel). This southerly wind drives a northward coastal flow, leading to a negative 

around-canyon velocity on February 4 (Figure 8a). 

 

The idea that the DeSoto Canyon flow is but a part of the continental shelf 

flow over the NEGOM is demonstrable with the hindcast model output. Figure 10 

shows a time-distance plot of the north-south velocity at 48m for the hindcast 

period along the parallel at 29o

40’N for the segment spanned by C1, D9, and E1 

(see Figure 1b). (No data return at shallower depth than 48m at D9.) During this 

winter period, it is clear that the flow is predominantly northward (red) on the west 

half and southward (blue) on the east. The contrast is particularly strong for the 

times of northerly wind bursts. The plot from interpolation from observed 

currents at C1, D9, and E1 shows a similar tendency. Thus the flow around 
44

DeSoto Canyon is clockwise under these northerly wind bursts, connecting the 

flow off the Florida Panhandle with that of the WFS. 

 An important consequence of strong currents in DeSoto Canyon following 

northerly wind bursts is that it leads to coastal upwelling. In the bottom panels of 

Figure 7 are found model vertical velocity contour plots in the DeSoto Canyon 

area during the six intrusion events in response to the northerly-wind bursts. 

Clearly, there is a strong upwelling center at the head of DeSoto Canyon due to 

onshore flows associated with strong around-canyon currents. It has been 

suggested that upwelling here could be due to the bottom Ekman flow under a 

strong clockwise current around DeSoto Canyon. Unfortunately, Ekman layer 

over a sloping terrain in a z-level model is decimated and not well resolved 

(Winton et al., 1998). As a consequence, north-south velocity at a nearboundary grid-point like the one at D1 is set to zero at the next level (72m) below 

49m to satisfy the no-normal flow condition, leading to an exaggerated frictional 

effect at mid-column depth. An effect that is represented reasonably correctly in 

the model that also leads to upwelling is the production of onshore flow due to a 

Figure 10. Time-distance plot of 16m north-south velocity components 

along the segment of the parallel at 29o40’N spanned by C1, D9, and E1. 

Model results are in the middle panel and observations are in the lower panel. 

The wind stress component time series shown in Figure 6 is reproduced in the 

top are contoured dashed. The contour interval is 4.5cms-1. 
45

slight excess in pressure gradient in the geostrophic balance associated with a 

clockwise around-canyon currents (see later in Discussion). The onshore flow 

over the shoaling topography then leads to upwelling. Because of this upwelling 

associated with strong around-canyon currents and flow to the southwest on the 

east flank, it is perhaps not surprising that peak monthly mean pigment 

concentration along the 40m isobath over the WFS should be found consistently 

for November through April (Gilbe et al., 1996). 

5. Discussion 

 It is argued here that strong currents in the DeSoto Canyon region result 

from the establishment of coastal currents over the WFS by the passage of 

continental shelf waves originating from the end of the WFS following each 

northerly wind burst. 

To substantiate this argument the model term balance (with all terms 

moved to the right side of the equation) at 16m in both the around-canyon and 

crosswise momentum equations at D1 in Figure 1b during the hindcast period of 

Figure 11. Around-canyon (left panel) and crosswise (right panel) momentum balance are 

shown at the 16.4 m depth at D1 at the head of DeSoto Canyon. Here, positive y direction is 

eastward and positive x direction is offshore. From top to bottom in the left panel, the curves 

represent, − fv , −∂ ∂ p / y , −∂ ∂ v t / , nonlinear term, and 2 2 / Avz v∂ ∂ . From top to bottom in the 

right panel, the curves represent, fv , −∂ ∂ p / x , −∂ ∂ u t / , nonlinear term, and 2 2 / Auz v∂ ∂ . The 

unit is 10-4cms-2. 
46

November 1997 – February 1998 is shown in Figure 11 (nonlinear terms are 

negligibly small and not shown). The crosswise momentum is nearly in 

geostrophic balance, with the crosswise pressure gradient nearly balancing the 

Coriolis term arising from the around-canyon flow. In the around-canyon 

direction, the balance must include the local time derivative of the around-canyon 

velocity, suggesting a continental shelf wave-like dynamics (Mitchum and Clarke, 

1986). 

 Figure 12 shows at 16m at D1 (in the lower panel) the time series of the 

excess crosswise pressure gradient (sum of the first two curves in the crosswise 

momentum balance in Figure 11), negative of local derivative, and the 

(exaggerated) vertical friction. The north-south component of wind-stress at 

Figure 12. Bottom Panel. Crosswise pressure gradient excess (sum of top 

curves in right panel of Figure 11, fv px − ∂ ∂/ ) at 16m at D1 for the hindcast 

period. The lower two curves in Figure 11, −∂ ∂ u t / and 2 2 / Auz v∂ ∂ , are 

reproduced in the same order here as the lower curves. Upper Panel. North-south 

wind stress component at Pensacola. Middle Panel. Around-canyon velocity 

component at16m at D1 reproduced from Figure 8. 
47

Pensacola (upper panel) and the model around-canyon velocity at 16m at D1 

from Figure 8 (middle panel) are reproduced here for easy reference. It is 

interesting to note that after northerly wind bursts when the around-canyon 

velocity is generally eastward, the excess pressure gradient force is onshore, 

balancing a crosswise local onshore acceleration. This is particularly clear 

around December 7,31, and February 4. The excess is perhaps due to the pileup of surface water under the northerly wind. (Plots (not shown) in the DeSoto 

Canyon area of model surface pressure gradient diagnosed for the November 7 

northerly wind event seems to confirm this in that the greatest offshore gradient 

(high pressure offshore) is found just after the passage of strong northerly 

winds.) It is the crosswise acceleration due to this pressure gradient that gives 

rise to a flow across the shoaling topography, leading to coastal upwelling in 

Figure 7. The contribution from vertical friction is small but not always negligible. 

A model with terrain-following coordinates is needed to properly assess the effect 

of the bottom Ekman flow, which can also produce upwelling under an eastward 

around-canyon current. 

6. Conclusions 

 A numerical study is conducted of both the Loop Current-driven and winddriven circulation in the northeastern Gulf of Mexico. The Loop Current of the 

Gulf of Mexico generally imposes an eastward drift in DeSoto Canyon and a 

southward drift over the WFS. The Loop Current-driven drift inside DeSoto 

Canyon produces only a weak flow. But, this picture of a suppressed role of the 

Loop Current could be due to the fact that the model is deficient in representing 

filaments and parasite eddies of the Loop Current that could extend the Loop 

Current influence to the continental shelf. Given this model deficiency, northerly 

winds are found to generate large around-canyon water movements in DeSoto 

Canyon. The dynamics of the wind-driven flow appears primarily to be that of the 

setup of coastal currents by continental shelf waves all along the NEGOM shelf. 

 An important aside is the discovery that strong coastal upwelling is found 

in the model at the head of DeSoto Canyon during the northerly wind bursts. 

This is apparently, in part, the result of excess in pressure gradient in the 

crosswise momentum balance brought by the northerly wind pile-up. The 

pressure gradient excess accelerates flow onshore across the shoaling 

topography, leading to coastal upwelling. Upwelling could also be due to bottom 

Ekman flow under the strong clockwise circulation around DeSoto Canyon 

following northerly wind bursts. Unfortunately the present model does not 

resolve the Ekman layer structure over the sloping shelf. Further clarification of 

the contribution of bottom friction to upwelling in DeSoto Canyon must await 

future studies using a model with terrain-following coordinates. 
48

ACKNOWLEDGEMENTS 

 This research is supported by the Minerals Management Service 

Cooperative Agreement 14-35-0001-30804 and Contract 1435-01-99-CT-31027. 

We thank Thomas Berger and Peter Hamilton of Science Applications 

International Corporation who kindly supplied the current meter data. Dr. 

Dongliang Yuan helped with model development. 

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52

5b. Temperature variability in the northeastern Gulf of Mexico in 1997-1998 
53

Temperature variability in the northeastern Gulf of Mexico in 1997-1998 

by 

Yury Golubev and Y. Hsueh 

Department of Oceanography 

Florida State University 

Tallahassee, FL 32306-4320 

Abstract

 

Observations made during 1997-1998 in the northeastern Gulf of Mexico 

(NEGOM) show that the monthly mean temperature at stations located near the 

shelf break behaves differently from that at stations in deep water 30-70 km 

away. Semi-annual variability is found in monthly mean temperatures at stations 

in both groups. But, there is a phase difference of about 180 degrees between 

the two groups. This phase difference makes for large monthly mean 

temperature contrasts at neighboring stations across the shelf break from each 

other and, accordingly, large cross-shelf temperature gradients in NEGOM. A 

basin-wide circulation model of the Gulf of Mexico run for the time period of the 

observation produces variability of monthly mean temperature along sections 

across the shelf break qualitatively similar to that observed. The cause is traced 

to the wind stress curl (mostly negative) in NEGOM, which has a strong semiannual variability. Following periods of particularly intense negative wind-stress 

curl in April and October, currents directed eastward along the shelf break in 

NEGOM associated with the wind-stress curl-driven anti-cyclonic deep water 

circulation intensify and generate onshore flow due to bottom friction, causing the 

isotherms near the shelf break to rise. At the same time, isotherms at deepwater stations are depressed on account of the vorticity balance forced by the 

negative wind-stress curl. Model balances lend support to this scenario. Thus, 

the observed phase difference in the monthly mean temperature at the semiannual frequency. Two individual events of anomalously intense cooling 

observed in NEGOM are addressed separately. It is shown that the dynamics 

leading to these events is probably associated with short-term responses to 

northerly wind bursts which produce strong eastward alongshore currents and 

super-geostrophic pressure gradients. Both are capable of uplifting isotherms 

over shoaling topography. 
54

1. Introduction 

 It is recognized that the circulation in the northeastern Gulf of Mexico 

(NEGOM) is dominated in the deep basin by the Loop Current (LC) and, on the 

continental shelf, by local winds (Hsueh and Golubev, 2001). Irregular northward 

penetration of the LC into NEGOM and shedding of anti-cyclonic rings from the 

LC occur at periods that fall into two ranges, 8-9 and 11-14 months (Sturges, 

1992,1994; Vukovich, 1995). The penetration of the LC itself in NEGOM 

apparently does not reach much beyond 27oN (Hsueh and Golubev, 2001). 

Thus, it is not expected that the LC heat and momentum have much of a direct 

effect in determining long-term variability (on the order of season) in the northern 

part of the NEGOM deep region. There is ample evidence of wind-driven 

currents on the continental shelf around NEGOM on the seasonal time scale 

(Blaha and Sturges, 1981; Mitchem and Sturges, 1982; Weisberg et al, 1996). 

Hsueh and Golubev (2001) show that on short time scales (3-10 days) the 

southward wind-driven current over the West Florida Shelf (WFS) pulls a strong 

flow along the continental shelf in NEGOM, leading to open ocean water intrusion 

and coastal upwelling in DeSoto Canyon. A still unanswered question remains, 

however, about the extent and manner in which the temperature field in the 

intervening region between the LC dominated deep Gulf and the wind-driven 

continental shelf is affected by wind-induced circulation on the seasonal time 

scale. The availability of current and temperature data from a recent mooring 

experiment carried out by Science Applications International Cooperation (SAIC) 

(SAIC, 1998) in the DeSoto Canyon area in NEGOM provides an opportunity for 

an examination of this question. 

 The goal of this paper is thus primarily to study processes governing the 

low-frequency temperature variability in the vicinity of the continental shelf break 

in NEGOM. Special attention will also be given to two individual events when the 

observed temperature at about the 100 m depth contour is anomalously low. 

Temperatures and currents from the SAIC observations and from a Bryan-Cox 

model will be compared to show and explain the dominant features in the shelf 

break area around DeSoto Canyon. 

 The paper is organized as follows. In the next section, a description is 

given of the data used and a presentation of the data is given. The configuration 

of the Bryan-Cox model of the basin-wide Gulf of Mexico circulation and results 

of the simulation are presented in Section 3. The nature of the two individual 

cooling events in December 1997 and January 1998 is discussed in Section 4. 

Concluding remarks are found in Section 5. 
55

2. Data 

a. Data description 

 There are thirteen current-meter moorings with various types of 

instruments in the year long (March, 1997 - April, 1998) SAIC measurement 

program that covers the Desoto Canyon area shown in Figure 1. 

The array is bracketed by two approximately along-shore lines. The line strung 

with moorings indicated with the index "1" in Figure 1(b), where velocity and near 

bottom temperature are measured, runs in about 100m of water. The deep water 

line (strung with moorings with index "3") are located at depths greater than 

1000m. Between these two lines the depth changes sharply, especially between 

stations B1 and B3 where depth goes approximately from 100m to 1100m in a 

distance of about 20 km. A catalogue of current and temperature data is given in 

Table 1. 

Figure 1. (a) The Gulf of Mexico model 

domain with the DeSoto Canyon study area in a 

rectangle. (b) SAIC mooring locations (solid circles) 

for the April 1997-March 1998 experiment with the 

solid square indicating the location of NDBC buoy 

42040. 
56

Table 1. SAIC Station number, Lat. and Long., Depth and Period of 

Temperature and Current Measurements 

NEGOM March 1997-April 1998. 

Station Lat Long, W Depth, m Period 

A1 88.485 29.241 80 97/03/21 - 98/04/01* 

A2 88.390 29.059 90 97/03/21 - 98/04/01 

A3 88.265 28.766 80 97/03/20 - 98/04/01 

B1 87.908 29.342 80 97/03/28 - 98/04/08 

B2 87.872 29.212 90 97/03/28 - 98/04/03 

B3 87.856 29.070 80 97/03/22 - 98/04/03 

C1 87.349 29.586 82 97/03/23 - 98/04/08 

C2 87.356 29.371 90 97/03/24 - 98/04/08 

C3 87.353 29.003 80 97/03/24 - 98/04/04 

D1 86.841 30.068 95 97/03/26 - 98/03/28 

D9 86.852 29.334 No Temperature Obs. 

D2 86.852 29.700 90 97/03/25 - 98/04/06 

E1 86.328 29.700 80 97/03/27 - 98/04/07 

* No observations from 11/13/97 to 11/21/97. 

 

 Hourly mean temperature at 80-90 m for all but two moorings (D9 with no 

temperature observations and E1 with monthly mean values very close to that at 

D1) are used to calculate daily and monthly mean values. Velocity data at the 

16m and 72m depths are also used. The location and measurement duration of 

each mooring are presented in Table 1. The level of 80-90m corresponds 

approximately to the lower boundary of the upper mixed layer. Hourly values of 

temperature and velocity components are arithmetically averaged for each of the 

month from April 1997 through March 1998. It will be convenient for subsequent 

presentation to divide the moorings from which data are used into four sets: set 

"A" (moorings A1, A2, and A3), set "B" (moorings B1, B2, and B3), set "C" 

(moorings C1, C2, and C3), and set "D" (moorings D1 and D2). 

 Three sets of wind data are used. The first one consists of winds 

observed at the meteorological buoy 42040 (29.2o

N, 88.25oW) located in the 

northern Gulf of Mexico (marked by the solid square in Figure 1(b)), from which 

data are available at the National Data Buoy Center. The second is the 

Hellerman and Rosenstein (1983) winds. The third is the 6-hourly, 10m surface 

winds from the NCEP/NCAR 40-year reanalysis data set (Kalnay et al., 1996). 
57

Similar to the Hellerman and Rosenstein winds, the NCEP winds are arranged in 

a grid with a 2-degree resolution. The NCEP winds are converted into wind 

stresses for the period from April 1997 through March 1998 using bulk formula 

and a drag coefficient of 1.3x10-3. 

b. Climatic features of temperature in NEGOM. 

 It would be useful to first give a brief description of the annual variability of 

temperature in NEGOM from climatic data. Monthly mean temperatures at 50, 

75, and 100m from Levitus (1982) are shown at two points in NEGOM (Figure 

2(a)). 

 

The northern point (solid circle near top) is located at the shelf break in 

100m of water and the southern one (solid circle near bottom) is located in water 

2800m deep. The temperature variability during the year differs from one point to 

the other. The difference in the shape of the curve becomes particularly 

noticeable at the point with large depth (Figures 2(b)). For both locations, the 

temperature at 100m shows a maximum in October of 21o

C. Most likely, this 

maximum represents what remains of the downward propagating temperature 

Figure 2. (a) Solid circles mark the locations in NEGOM where the 

climatological monthly mean temperatures from Levitus (1982) at 50, 75, and 100m 

are shown in (b) (the southern location at 27.5oN) and (c) (the northern location at 

29.5o

N). (d) Climatological monthly mean meridional (dashed) and zonal (solid) 

component of windstress for the region in (a) from Hellerman and Rosenstein (1983). 
58

signal arising from intense heating at the sea surface in July-August. In months 

other than September-October temperatures at 75 and 100m are much lower at 

the northern (shelf break) point than at the southern (deep water) point. Because 

of the location, it seems that the lower temperature at the northern point might 

indicate coastal upwelling. The lack of a favorable zonal wind stress (solid curve 

in Figure 2(d)) from the Hellerman and Rosenstein (1983) data set discounts this. 

The zonal wind-stress component is mostly directed toward the west during the 

entire year, favoring down-welling instead. This wind generates shoreward 

Ekman transport in the upper layer, opposite to what is necessary for the 

appearance of cold water near the shelf break in July and December (Figure 

2(c), 100m). So, the reason for the appearance of low monthly-mean 

temperature near the shelf break has to be other than a direct result of the local 

along-shelf wind-stress and its associated Ekman transport. Strong currents of 

short duration in DeSoto Canyon can, however, arise from the establishment of 

southward wind-driven currents on the WFS, giving rise to an around-canyon 

flow conducive to an onshore bottom Ekman transport (Hsueh and Golubev, 

2001). The onshore transport across the shoaling coastal topography then leads 

to upwelling. The efficacy of this dynamics on the monthly mean time scale is, 

however, doubtful. (The monthly mean wind stress is generally an order of 

magnitude smaller and the north-south component (not shown but see 

www.ocean.fsu.edu/oce/gom) over the WFS is northward for July, leading to a 

counterclockwise around-canyon flow.) 

c. Variability of monthly mean temperature and velocity during 1997-1998. 

 Figure 3 presents monthly mean temperature at 80-90m at the shelfbreak and deep water stations in each of the four mooring sets obtained from 

observations made during 1997-1998. The common features of temperature 

variation for all four mooring sets are: 

I. Temperature at the shelf-break stations (lower curves) is generally 

lower and has two maxima in May and October of 1997 and two minima in July 

1997 and January 1998. The second minimum is lower than the first one at 

stations A1, B1, and C1. At station D1 both minima have nearly equal values. 

These variations at the shelf-break moorings are qualitatively similar to those 

found in the climatic distribution at 100m at the northern point in Figure 2(c). 

2. Temperature at the deep-water stations (upper curves) also has 

two highs and two lows but time-shifted so that the highs line up with the lows 

and the lows line up with the highs at the shelf-break. Furthermore, no clear 

resemblance is found in the variation at stations A3, B3, C3, and D2 with the 

climatic distribution at 100m (Figure 2(b)). 
59

 

 

3. Temperature variations at A2, B2, and C2 between the shelf break 

and the deep water feature what is the average for the shelf-break and deepwater stations and the temperature curves (not shown) generally lie in between 

the two curves presented in Figure 3. 

 The vertical bars in Figure 3 represent the root-mean-square (rms) of 

temperature fluctuations for each month of the year. These values can be 

treated as a measure of temperature variability during each individual month. 

There are instances of particularly large rms' in December 1997 and January 

1998. These reflect the presence of anomalous short-term cases, which will be 

discussed in Section 4. 

 

Figure 3. Monthly mean temperatures obtained from observations during March 1997-

April 1998 at four pairs of mooring stations across the continental margin shown in Figure 

1(b). The mooring pair is shown in the lower left-hand corner of each of the four panels. In 

each panel, the top curve is from the deep mooring and the lower curve from the shelf-break 

mooring. The vertical bars indicate the root-mean-square values for each month.
60

 

 Monthly mean values of the along-shelf component of observed currents 

at 16m at the four shelf-break moorings are presented in four panels in Figure 4. 

(The directions in angles of rotation clockwise from the north are indicated in 

each panel.) Curves (not shown but see www.ocean.fsu.edu/oce/gom) for 72m 

are qualitatively similar. There are two-peaks with the largest velocities in JuneJuly and December-January. The velocity reaches a minimum in SeptemberOctober. The peaks in velocity occur roughly at the time of temperature minima 

in Figure 3. This is particularly clear for the June-July peak. 

 It is particularly interesting to underline the large temperature contrasts 

found between shelf-break and deep-water moorings (temperature being higher 

at deep-water moorings). Monthly mean temperature at D2 exceeds that at D1 in 

July 1997 by more than 5o

C. It is 4oC between C3 and C1, and almost 5oC in 

December 1997-January 1998. Table 2 lists temperature contrasts and 

horizontal temperature gradients between pairs of moorings with indices 1 and 2 

in each mooring set. Horizontal temperature gradients of 0.15oC/km are large for 

the seasonal time scale and may have important impact, in particular, on the 

biology in NEGOM. 

Figure 4. Monthly mean along-shelf component of velocity at 16m at the shelfbreak moorings of the mooring pairs shown in Figure 3. The direction in angle of 

rotation clockwise from the north is given in each of the four panels. The vertical 

bars represent the root-mean-square values for each month. 
61

Table 2. Temperature Difference (∆ oC) and Gradients (°C/km) for 

Stations with Indices 1 and 2 (Figure 1(b)). 

July, 1997 January, 1998 

∆ Gradient ∆ Gradient 

A2 – A1 1.46 0.072 2.63 0.131 

B2 – B1 1.38 0.095 2.24 0.154 

C2 – C1 3.39 0.141 2.45 0.102 

D2 – D1 5.15 0.127 3.18 0.078 

3. Model 

 

a. Model design and configuration. 

A numerical model for the Gulf of Mexico basin-wide circulation driven by 

both the Loop Current through the Yucatan Strait and the wind-stress field must 

be developed that is able to resolve dynamic processes on the continental shelf 

and over the sharply changing upper continental slope topography in NEGOM. 

The eddy-resolved circulation model employed to meet this requirement is the 

rigid-lid, primitive-equation model of Bryan and Cox (Bryan, 1969; Cox, 1984). A 

detailed description of the model in a hindcast for the observation period can be 

found elsewhere (Hsueh and Golubev, 2001). It suffices to say that the model 

reproduces the Loop Current, but is deficient in resolving small-scale flow 

structures such as Loop Current filaments and parasite eddies which may extend 

into the continental shelf region. The model thus accentuates the wind influence. 

Care has been taken to make sure that the initial condition for the integration 

features a Loop Current with a configuration similar to that found for March 1, 

1997 from the TOPEX/Poseidon altimeter measurement (Hsueh and Golubev, 

2001). 

 b. Temperature variability. 

 Monthly mean model temperature and velocity are examined for a number 

of north-south (nearly cross-shelf) sections over the shelf break in the area of the 

mooring experiment. Monthly mean temperature at each grid-point along one of 

these sections (Figure 5(a)) is shown in Figure 5(b). 
62

The section is located near mooring set "A" (see Figure 1(b)). The 

northernmost grid-point (70, 68) is located on the 100m isobath. The 

southernmost point (70, 64) is in 1500m of water. So the section spans sharp 

changes in topography (with distance between (70, 64) and (70, 68) equal to 72 

km). The monthly mean temperature at 115m at each grid-point is shown in 

Figure 5(b). 

Two impressions immediately present themselves. First, the shape of the 

curves at both shallow and deep-water points is qualitatively similar to that of the 

observed (Figure 3). Quantitatively, the model cross-shelf gradients are much 

weaker than those found in observation. Be it as may, it is still remarkable that 

the gradient in May or June is almost twice as large as that in April or August and 

four time as large as that in March 1998. Second, in the model, the northernmost 

location of the northern boundary of the model LC during the year is around 

27.5o

N (Hsueh and Golubev, 2001). So it must not be the model LC that is 

directly responsible for this variation in temperature. This leaves the wind as the 

only factor in this part of the Gulf that lies north of 28o

N. Since, as is allude to 

earlier, climatically, alongshore winds are mostly easterly, local winds are not 

able to lower temperatures in shallow water areas. Might it not be that it is 

actually the wind-stress curl field that is known to have semi-annual variability in 

NEGOM that is responsible for this large temperature contrasts from the shelf 

break to the deep? 

Figure 5. (a) A north-south section of model 

grid-points across the continental margin in the 

DeSoto Canyon area of the SAIC mooring 

experiment. (b) Monthly mean model temperature at 

115m during April 1997-March 1998 at each of the 

grid-points in (a). The curves sort themselves out 

with decreasing depth from the top one to the one at 

the bottom. 
63

c. Winds and currents in NEGOM and their influence on the temperature field. 

Monthly mean values (averaged for the two marked locations in Figure 

2(a)) of zonal component of model wind-stress for the period March 1997-April 

1998 for the area in Figure 2(a) are presented in Figure 6(a). 

To the surprise of no one, the variation has many features common with 

that from Hellerman and Rosenstein (1983) presented in Figure 2(d). Most 

importantly, as in the case of the Hellerman and Rosenstein wind-stress, the 

wind-stress during 1997-1998 is predominant directed to the west (negative 

values) and so is the annual mean. This means, for April 1997-March 1998, with 

local winds alone the water temperature at the shelf-break stations would be 

higher than that at the deep water stations, not the other way around. It is thus 

safe to discard the possibility that local wind-stress-produced coastal upwelling 

would lead to the observed cross slope temperature difference. The appearance 

of cold water near the shelf break and warm water at the deep water stations 

must be rooted in the wind-driven circulation in other ways. 

 

In the northern part of Gulf of Mexico (north of 24oN), the wind field shows 

an anti-cyclonic rotation (Elliott, 1979; Krishnamurti and Krishnamurti, 1975) with 

a curl maximum in the western part of the Gulf (Blaha and Sturges, 1978, 1981). 

Monthly mean values of wind-stress curl for 1997-1998 derived from NCEP 

winds averaged for the area in Figure 2(a) is shown as Figure 6(b). The windFigure 6. (a) Monthly mean zonal windstress component obtained from averaging the 

values used in the model at the two locations in 

Figure 2(a) and (b) monthly mean model windstress curl in the area in Figure 2(a) during April 

1997-March 1998 calculated using NCEP/NCAR 

reanalysis data.
64

stress curl has two peaks in the anti-cyclonic sense, one in April and the other in 

October associated with the weakening of easterly and northeasterly winds in the 

northern part of the Gulf during mid-spring and mid-fall. Model results (see later 

Figure 8) show that this wind-stress curl produces anti-cyclonic circulations in 

NEGOM. Along-shelf component (directed in a direction obtained by rotating 

clockwise 70o from north) of currents averaged for the five grid-points of the 

section in Figure 6(a) is shown in Figure 7. 

Similar to the shape of the wind-stress curl, currents have two peaks 

during the year: May-June and November-December. These months all follow 

periods with strongest anti-cyclonic wind-stress curl in April and October (with a 

1-1.5 months delay) (see Figure 6(b)). The time delay amounts to a phase shift 

of about one-quarter of a period for the semi-annual variability, reasonable for an 

inviscid forcing-response relationship (see later (1)). 

Pictures of monthly mean horizontal model circulation averaged for the 

upper 100m in NEGOM is presented in Figure 8 for April and June (two upper 

panels). 

Figure 7. Monthly mean model velocities in 

the along-shelf direction (70o

 clockwise from 

north) at three depths averaged over the five 

grid points in Figure 5(a).
65

There is a clear distinction between the two in terms of weak currents in 

April and strong circulation in June. (A similar contrast is also found for October 

and December, another month-pair when the currents change form being weak 

to strong.) All monthly mean pictures of such horizontal circulation share one 

common feature. The along-shelf currents have cyclonic vorticity to their 

northeastern side that is bounded by the shelf and anti-cyclonic vorticity on the 

seaward side (Figure 8 lower panels). In months when the circulation is strong 

(June and December) relative vorticity is also strong: negative relative vorticity is 

strong offshore because the anti-cyclonic circulation is more intense; positive 

relative vorticity is strong because the velocity at the shelf break is increased due 

to the intensification of anti-cyclonic circulation. The process is similar to the 

case with the Gulf Stream. In the coastal side of the Gulf Stream, large cyclonic 

vorticity values exceed the local parameter of Coriolis due to the intensification of 

the currents when the Gulf Stream axis shifts shoreward (Richardson et al., 

1969; Gill, 1982). 

 

Figure 8. Vector plots of the average model currents in the upper 100m in April 

(upper left panel) and June (upper right panel). The corresponding average relative 

vorticity distributions for the same layer are shown in the lower panels, left for April 

and right for June. Contours of relative vorticity, at interval of 2x10-9s-1, are dashed 

for negative values and solid for positive values. The depth contours are 100, 200, 

and 1000m in the upper panels and 100m in the lower panels.
66

The strong eastward currents along the shelf break associated with the 

anti-cyclonic circulation then drive an onshore Ekman transport over the shoaling 

topography, leading to cooling and raised isotherms at the shelf break. Figure 9 

presents monthly mean along and cross-shelf model velocities at 49m at the gridpoints that make up the section in Figure 5(a). 

 

The cross-shelf velocities at (70, 68), where 49m being the deepest level 

of non-zero velocity, and (70, 67) are onshore (positive) when the along-shelf 

current is to the east (positive), indicating an Ekman flow. The reason that 

horizontal velocity is set to zero at a level (72m) below 49m is because the step 

representation of the slope bottom imposes a solid northern boundary at this 

depth for (70, 68). As a consequence, bottom friction is present at 49m, a midcolumn depth. This is an aberration attributable to the nature of z-level models in 

dealing with a sloping bottom (Hsueh and Golubev, 2001). But, the qualitative 

effect of bottom friction in generating an onshore flow is probably still valid (see 

later Figure 12). In deep water off the shelf break, but still beyond the reach of 

the model LC, at grid-points (70, 64) and (70, 65), deep isotherms is depressed 

due to the presence of the anti-cyclonic wind-stress curl. So, the monthly mean 

temperature at the deep-water end of the section will show warming. Figure 10 

summarizes the cooling and warming combination with a plot of the monthly 

mean depth of the 17.5o

 C isotherm along the cross-shelf section in Figure 5(a) 

as a function of time and cross-shelf distance. 

Figure 9. Model velocity components at 49m at the grid points in Figure 

5(a) in a coordinate frame in which the positive along-shelf direction is 70o

clockwise from north. (a) Along-shelf components (positive to the east). (b) 

Cross-shelf components (positive onshore). The numbers on the curves give 

the north-south grid-points. 
67

 

In June, a month or so after the occurrence of a strong negative windstress curl (see Figure 6(b)) the isotherm deepens at the deep end. At almost 

the same time, the same isotherm at the shallow end moves up. The same outof-phase oscillation shows up in November, again about one month after the 

occurrence of a strong negative wind-stress curl in Figure 6(b). The linear 

vorticity equation that might help to explain this is 

0

1 0 h curl

t f

τ

ρ

∂

+ =

∂ (1) 

Here, h is the monthly mean thickness of the layer between the sea surface and 

the 17.5o

 C isotherm, f is the Coriolis parameter, and τ is the monthly mean 

wind-stress vector. Figure 11 presents the balance in (1) from the model output 

at grid-point (70, 64). 

Figure 10. The depth of the 17.5o C isotherm in 

the model along the cross-shelf section in Figure 

5(a) is contoured in a time-latitude plot.
68

 Other than the last four months or so where wind-stress curl becomes 

more variable, the qualitative impression is that, to the first order, the two terms 

tend to cancel each other. 

So, the negative wind-stress curl generates an anti-cyclonic circulation in 

the deep part of NEGOM in a month or so and gives rise to a strong shelf-break 

current at 100m to the east on the west flank of DeSoto Canyon. The strong 

current to the east along the shelf break creates an Ekman onshore flow due to 

bottom friction that, when coupled with the shoaling topography, lifts the 

isotherms upward, causing cooling. Figure 12 presents the monthly mean 

balance of the cross-shelf momentum at 49m at grid-point (70,68). 

 

Figure 11. Plot of the monthly mean balance in (1) at 

model grid-point (70, 64), the southernmost point of the 

cross-shelf section in Figure 5(a). The solid curve 

represents the model calculation of

h

t

∂

∂

. The dashed curve 

represents the model wind-stress curl divided by the 

product of density (1.023g/cm3

) and the Coriolis parameter 

at (70, 64). Here, the Coriolis parameter is taken as a 

constant of 10-5 s-1.

Figure 12. The monthly mean cross-shelf momentum 

balance is plotted at 49m at (70, 68), the northernmost gridpoint on the cross-shelf section in Figure 5(a), in a 

coordinate frame in which the positive along-shelf (x-) 

direction is 70o

 clockwise from north and the positive crossshelf (y-) direction is onshore. All the terms in the crossshelf momentum equation are moved to the right-hand side. 

The local derivative, 

v

t

∂ − ∂ , is dotted, advection dot-dashed, 

and ageostrophic term, p fu

y

∂ − − ∂ , dashed, and vertical 

friction solid. 
69

The cross-shelf pressure gradient and the Coriolis term arising from the 

strong along-shelf current is combined to give the ageostrophic term so that it 

may be compared with friction to show the Ekman balance. The plot shows 

clearly that the (vertical) friction term nearly balances the ageostrophic term, 

giving support to the contention that on the monthly mean the cross-shelf flow is 

Ekman. (The onshore flow at the deep-water points (Figure 9(b)) is just a part of 

the geostrophic flow around the anti-cyclonic gyre.) 

 In the deep water, the deepening of the isotherms is argued to be due to 

the forcing of the anti-cyclonic wind-stress curl. The linear vorticity equation (1) 

shows a rough balance between the variation of the 17.5oC isotherm depth and 

wind-stress curl. The deepening of isotherms in the deep water region due to the 

anti-cyclonic wind-stress curl and the raising of isotherms at the shelf break due 

to Ekman flow under eastward currents over the shelf break associated with the 

anti-cyclonic circulation in the deep water region thus combine to create the out 

of phase variation in monthly mean temperatures presented in Figure 3. 

4. COOLING EVENTS IN NEGOM. 

The main goal of this section is to examine the role of wind-driven currents 

in strong short-term cooling events in the northeastern Gulf of Mexico. First, 

temperature and current data in DeSoto Canyon will be examined to show two 

strong cooling events observed in December 1997-January 1998. Second, the 

Bryan-Cox model output will be analyzed to explain the dominant mechanisms 

that lead to these events. 

Two strong cooling events are found in daily mean temperature time 

series at 85-90m at stations C1 and D1 (Figure 13). The first event is observed 

on December 17-18, 1997. During the 3-5 days temperature drops from 20o

C on 

December 13-14 to 12oC at station C1. The drop is to 11.5oC at station D1. The 

temperature does not recover fully after the event. The second event is found on 

January 1-2, 1998. The temperature drop from December 26 to January 1-2 is 7-

8o

C and again the temperature does not recover completely after the event. It 

can also be noted that, in general, the temperature drops precipitously and 

recovers relatively gradually. Second, as Figure 13 shows, there are only two 

stations among the twelve-station set where these features in temperature are 

observed. There is no evidence that these cooling events occur at off-shelf 

stations. 

The cooling events are detected in the observation at 85-90m below the 

sea surface. But, both leave a mark, especially the second one, in sea surface 

temperatures. Figure 14 shows two aerial AVHRR sea-surface temperature 

images at about the time of these events. 
70

Figure 13. Daily mean temperature at the 85-90m depth at all mooring stations 

in Figure 1(b) except D9 during December 1, 1997 - January 14, 1998. Station 

designation is given in the lower left corner of each panel.
71

As can be seen, stations C1 and D1 are indeed located in the area of low 

water temperatures. Strong temperature contrasts, particularly in the January 

case, are found across the shelf. To the east of 87.5oW, the isotherms basically 

follow the shelf break. 

An examination of wind data at NDBC buoy 42040 during December 1997 

- January 1998 uncovers two cases of strong winds in NEGOM during this 

period. The first case occurs on December 12, 1997 when the wind is northerly 

with a daily mean velocity of about 10 m/s, translating into a wind-stress of more 

than 1 dyne/cm2

 (Figure 15(a)). 

Figure 14. AVHRR sea surface temperature 

images for the region of observation from the 

Pathfinder 9km SST Climatology 

(http://podaac.jpl.nasa.gov). (a) December 17-23, 

1997. (b) January 1-6, 1998. White dots on both 

plates show the mooring locations (Figure .1(b)).
72

The second strong wind event with west-northwesterly winds occurs on 

December 29, 1997. As can be seen from Figure 15(b)), after each wind event, 

the kinetic energy of model horizontal circulation peaks on December 17 and 

December 31, 1997, just before the occurrence of the cooling events. 

There are interesting differences in pictures of currents before and after 

wind action (Figure 16). 

Figure 15. (a) Daily mean meridional wind 

velocity at NDBC buoy 42040 during December 

1, 1997 - January 14, 1998. The two days with 

the strongest northerly winds are indicated. (b) 

Model current kinetic energy at D1 for the 

considered period of time. Solid curve 

represents the kinetic energy at 16m and the 

dashed curve for 72m. Days with maximum of 

current intensity are indicated. 
73

The pattern features meandering along-shelf currents with approximately 

equal intensity for all cases. In days with relatively weak winds (December 10 

and 24 - two upper panels) the current near the shelf break is either weak or 

cyclonic. After the wind action (two lower panels), strong and narrow along-shelf 

currents emerge particularly along the west side of DeSoto Canyon with 

maximum onshore velocities at the head of the canyon (Figure 17a). 

Figure 16. Vector plots of model currents at 72m in NEGOM. Two upper panels 

represent currents on December 10 and 24 before strong winds. Two lower panels 

show currents after strong of winds of the two days indicated in Figure 15. 
74

Apparently, the strong along-shelf current to the east is able to produce, 

near the bottom, an Ekman veering to transport cold water from the deep to the 

shelf. As mentioned early on, because of the step-wise representation of the 

sloping bottom in a z-level model, the bottom friction effect appears at shallower 

depth than in reality (Hsueh and Golubev, 2001). In fact, in both these cases, a 

viable mechanism may be that of excess onshore pressure gradient forcing an 

onshore flow, leading to upwelling at the canyon head. The case on December 

17 is indeed one of the cases discussed in such light in Hsueh and Golubev 

(2001). Even with exaggerated bottom friction effect, the term balance for crossshelf (north-south) momentum at 72m (Figure 17b) shows the Ekman dynamics 

compensates only 35-40% of ageostrophic term. The rest is balanced by the 

local acceleration (the one term stressed in the super-geostrophic pressure 

gradient argument (Hsueh and Golubev (2001)) and advection terms. In 

addition, unlike the case with monthly mean balance (Figure 12), sometimes 

(December 31) a significant part of the ageostrophic term is balanced by 

advection. Thus, some of the cross-shelf flow may be a result of flow overshooting along the west flank of DeSoto Canyon, an aspect not discussed in 

Hsueh and Golubev (2001). Further work is clearly necessary to address the 

inertial nature of the along-shelf currents in NEGOM following strong northerly 

wind bursts. 

Figure 17. (a) Daily mean model cross-shelf velocity at 72m at (79, 72), a 

grid-point just south of D1. (b) Plot of the daily mean cross-shelf 

momentum balance at 72m at the same grid-point. All the terms in the 

momentum equation are moved to the right-hand side. Local acceleration 

is dotted, advection is dot-dashed, ageostrophic term is dased, and vertical 

friction term is solid. 
75

5. Conclusions. 

 It is argued that the out-of-phase oscillation in the monthly mean 

temperature at 100m across the continental margin from the deep NEGOM to the 

shelf break observed in March 1997-April 1998 stems from the difference in the 

dynamics that operate in the deep Gulf of Mexico and over the shelf break, 

although all due to the prevalent anti-cyclonic wind-stress curl in NEGOM. The 

negative wind-stress curl generates an anti-cyclonic circulation in the deep part 

of NEGOM and gives rise to a strong shelf-break current at 100m to the east on 

the west flank of DeSoto Canyon. The strong current to the east along the shelf 

break creates an onshore flow from Ekman veering that, when coupled with the 

shoaling topography, lifts the isotherms upward, causing cooling. 

 

 During the occurrences of intense anti-cyclonic wind-stress curl, the deep 

isotherms at the deep-water moorings are lowered according to the linear 

vorticity balance on the month mean and the strong along-shelf current near the 

shelf break associated with the anti-cyclonic circulation in the deep Gulf 

generates a cross-shelf flow in the bottom Ekman layer, leading to elevated 

isotherms found at moorings over the shelf break. Both the bottom Ekman layer 

balance at the shelf-break stations and the vorticity balance at the deep-water 

stations are demonstrated using the numerical model as a proxy. The two 

mechanisms are consistent with each other and combine to produce the 

observed large temperature contrast across the continental margin in the DeSoto 

Canyon. 

Two strong short-term cooling events are found in temperature series from 

stations C1 and D1 on December 17-18, 1997 and January 1-2, 1998 when 

temperatures at 85-90m drop by 7-8o

C during 3-5 days. An examination of 

model output shows the presence of strong eastward narrow along-shelf current 

during the days of the temperature drop, indicating perhaps a measure of Ekman 

veering in velocity in the lower water column. The term-balance for cross-shelf 

momentum shows the term responsible for the Ekman transport compensates 

only 35-40% of the ageostrophic term. The rest is balanced by the local 

acceleration and advection terms. The local derivative appears to come from an 

excess onshore pressure gradient found elsewhere (Hsueh and Golubev, 2001). 

The substantial amount of advection required for the second event indicates that 

the inertia of the along-shelf flowing current may sometimes lead to overshooting that can produce additional cross-isobath flow, leading to upwelling and 

the appearance of cooling inshore. 
76

Acknowledgements 

 Appreciation is noted of helpful discussions with W. Sturges and G. 

Weatherly. Thomas Berger and Peter Hamilton of Science Application 

International Corporation kindly supplied the data. Assistance rendered by Dr. 

Dongliang Yuan is also acknowledged. This research is supported by the U.S 

Minerals Management Service Cooperative Agreement 14-35-0001-30804 and 

Contract 1435-01-99-CT-31027. 

REFERENCES 

Blaha J., and W. Sturges, 1978: Evidence for wind forced circulation in the Gulf 

of Mexico, Technical Report, Department of Oceanography, Florida State 

University, 134pp. 

Blaha J., and W. Sturges, 1981: Evidence for wind-forced circulation in the Gulf 

of Mexico, J.Mar.Res., 39(4),711-734. 

Blanton,J., L.Atkinson, L.,Pietrafesa and T.Lee, 1981: The intrusion of Gulf 

Stream water arross the continental shelf due to topographically induced 

upwelling. Deep-Sea Res., 28A, 393-405. 

Bryan, Kirk, 1969: A numerical method for the study of the circulation of the 

world ocean. Journal of Computational Physics, 4, 347-376. 

Elliott, B.A., 1982: Anticyclonic rings in the Gulf of Mexico, J. Phys. Oceanogr., 

12(11), 1292-1309. 

 

Garrett, C., 1979: Topographic Rossby waves off East Australia: Identification 

and Role in Shelf Circulation, J. Phys. Oceanogr., 9, 244-253. 

Gill A., 1982: Atmosphere-Ocean Dynamics, Academic Press, 876p. 

Hellerman S, and M. Rosenstein, 1983: Normal monthly wind stress over the 

world ocean with error estimates, J. Phys. Oceanogr., 13, 1093-1104. 

Hsueh, Y., and Yury Golubev, 2001: A numerical model calculation of the flow in 

DeSoto Canyon in response to northerly wind bursts in winter. Submitted to Gulf 

of Mexico Science. 
77

Kalnay, E., M. Kanamitsu, R. Kistler, W. Collins, D. Deaven, L. Gandin, M. 

Iredell, S. Saha, G. White, J. Woollen, Y. Zhu, A. Leetmaa, R. Reynolds 

M. Chelliah, W. Ebisuzaki, W.Higgins, J. Janowiak, K. C. Mo, 

C. Ropelewski, J. Wang, Roy Jenne, Dennis Joseph, 1996: The NCEP/NCAR 

40-year reanalysis project. Bulletin of the American Meteorological Society, 77, 

437-471.

Krishnamurti, T.N. and R. Krishnamurti, 1979: Surface meteorology of the Ascale during one-hundred days of GATE. Deep-Sea Res., GATE Suppl. 26, 29-

26. 

Levitus, S., 1982: Climatological atlas of the world ocean. NOAA Prof. Pap.,13, 

US Government Printing Office, Washington DC, 172. 

Mitchum, G., and W. Sturges, 1982: Wind-driven currents on the west Florida 

shelf. J. Phys. Oceanogr., 12, 1310-1317. 

Science Application International Corporation: DeSoto Canyon Eddy Entrusion 

Study. MMS Contract 1435-01-96-CT-30825. Data Products: Set 1-5, SAIC, 

Raleigh, North Carolina, 1998. 

Richardson W.S., Schmitz W.S., and Niiler P. 1969: The velocity structure of the 

Florida Currents from the Florida Strate to Cape Fear. Deep-Sea Res. 16, Suppl., 

225-331. 

Sturges, W., 1992: The spectrum of Loop Current variability from gappy data. J. 

Phys. Oceanogr., 22, 1245-1256. 

Sturges, W., 1994: The frequency of ring separation from the Loop Current. J. 

Phys. Oceanogr., 24, 1647-1651. 

Vukovich, Fred M., 1995: An updated evaluation of the Loop Current’s eddyshedding frequency. J. Geophys. Res.,100, 8655-8659. 

Weisberg, R., B. D. Black, and H. Yang, 1996: Seasonal modulation of the west 

Florida continental shelf circulation. Geophys. Res. Lett., 23, 2247-2250. 
78

5c. Low-frequency variability of surface currents on the northeastern Gulf of 

Mexico (NEGOM) shelf 
79

LOW-FREQUENCY VARIABILITY OF SURFACE CURRENTS ON THE 

NORTHEASTERN GULF OF MEXICO (NEGOM) SHELF 

by 

Yury Golubev* and Ya Hsueh 

Department of Oceanography 

Florida State University 

Tallahassee, FL 32306 

Abstract 

 An examination of the paths of satellite-tracked drifters uncovers 

differences in the temporal variability of low-frequency surface currents on 

NEGOM shelf areas east and west of 85W. The monthly mean currents obtained 

from drifter paths east of this longitude are well correlated with sea surface 

heights (SSH) in the Straits of Florida, reflecting the offshore pressure forcing 

that drives a barotropic flow on the west Florida shelf (WFS) on a nearly annual 

frequency. In contrast, monthly mean surface currents on the shelf west of 85W 

are found to correlate highly with the zonal wind component, indicating the 

efficacy of alongshore winds in driving the alongshore shelf currents. 

1. Introduction. 

The study area of the northeastern shelf of the Gulf of Mexico (Figure 1) 

comprises of the area between the Mississippi River mouth and Cape San Blas 

in the Florida Panhandle and the northernmost part of the broad west Florida 

shelf (WFS) with a width on the order of 200km. 

Figure 1. The map depicts the continental 

shelf in the northeastern Gulf of Mexico. 
80

There have been a number of studies of the continental shelf circulation 

and its variability in the northeastern and eastern Gulf of Mexico. The general 

hydrography and current regime in this region are summarized in early literature 

(Ichiye et al., 1973; Jones, 1973). Direct current measurements are discussed in 

several more recent articles (Mitchum and Sturges, 1982; Marmorino, 1982; 

Weisberg et al., 1996; Yang and Weisberg, 1999). Drift bottle studies (Tolbert 

and Salsman, 1964) show a measure of seasonality in the surface circulation. 

Released in summer, drift bottles tend to return to the north and those released 

in winter tend to return to the south. Seasonal surface-current variability on the 

WFS has also been reported on the basis of data from satellite-tracked drifters 

(P.P. Niiler at al., personal communication, 1996). 

As is well known, the Loop Current (LC) dominates the circulation in the 

deep part of the eastern Gulf of Mexico. Irregular northward penetrations of the 

LC into the eastern Gulf occur with periods between 8-9 and 13-14 months 

(Sturges, 1992,1994; Vukovich, 1995). The northward penetration usually 

culminates with the separation of an anti-cyclonic ring, followed by the collapse of 

the loop to its starting position around Cuba. Vukovich (1995) finds a 

predominant period of the LC northward penetration of 11.1 months. Sturges 

and Leben (2000) give two main frequencies of ring detachment, one at 9 and 

the other at 11.6 months. So, the northward penetration of the LC occurs at a 

nearly annual frequency. This wax and wane of the LC imparts variability on the 

WFS. An oddity actually is noted in the influence on the shelf flow of the LC and 

its variability. Sturges and Evans (1983) finds that on the low frequency, the 

southward flow inferred from coastal sea-surface elevations over the WFS does 

not reach its maximum strength at the time when the LC is fully extended into the 

Gulf, with the LC front in close proximity of the WFS. This phase miss-match 

indicates the LC influence on the WFS comes about not through lateral friction, 

but through pressure imposed by the LC along the shelf break (Hsueh and 

Golubev, 2001). 

On the other hand, there is ample evidence of direct wind influence on a 

seasonal time scale in shelf currents in the Gulf of Mexico (Jones, 1973; Blaha 

and Sturges, 1981; Mitchum and Sturges, 1982; Weisberg et al, 1996). Thus, 

the above-mentioned seasonality in surface velocity data from the drifters in the 

study area can be either (locally) wind-driven or remotely induced by the LC 

careening around the south end of the WFS with a variability approximately at 

the annual period. The purpose of the present paper is to discuss, on the basis 

of a variety of available data, the role of these two factors, the wind and LCrelated pressure, in the seasonal variability of drifter-based surface currents on 

the continental shelf in NEGOM. Section 2 describes the data source and data 

statistics. Data analysis is reported in section 3. A final section 4 contains the 

summary and conclusions. 
81

2. Data description. 

Three sets of data are used in the analysis. The first one is a surface 

velocity data set based on satellite-tracked drifters deployed from February 1996 

through April 1997 in the eastern Gulf of Mexico (P. P. Niiler at al., personal 

communication, 1996) as a part of Mineral Management Service NEGOM 

Project. The second set of data comprises the sea surface height measurements 

obtained by the TOPEX/Poseidon satellite altimeter for the same time period. 

The third set of data used is wind velocity data from three NOAA stationary 

buoys in the northeastern Gulf of Mexico (see later Figure 3). 

The following steps are taken to treat the velocity data for the study. First 

of all, drifter velocity components are binned into daily mean series with a 0.25 

(longitude) by 0.17 (latitude) degree bin size in the area 27N-30.5N and 83W89W from February 7,1996 through April 30, 1997. If there is only one drifter 

velocity in a particular bin square, it is accepted as the daily mean for that 

particular day. If there are several drifter velocities during a day in the square, 

the velocity components are averaged to give the daily mean. The number of 

days with non-zero daily mean velocity over the study area in the time period 

between February 7, 1996 and April 30, 1997 is shown at Fig. 2a. 

Figure 2. (a) The field of number-of-day with non-zero daily mean velocity from the 

drifter trajectories between February 7, 1996 and April 30, 1997 is contoured at an 

interval of 15. The total number of days with non-zero values for the entire study 

area is 448. (b) Bar-charted is the number-of-square with a non-zero monthly mean 

drifter velocity between February 1996 and April 1997. Each square measures 0.25 

by 0.17 degrees. (c) The twelve meridional strips of NEGOM shelf area in which 

daily mean and 10-day mean drifter velocities are calculated. 
82

As can be seen, the shelf area between 29N and 30.5N has the highest 

density of observations. To the south and southeast, the study area is bounded 

by the 500m isobath. 

The next step is to obtain monthly mean drifter velocities at the center of 

the square. There are 15 months from February 1996 till April 1997. To obtain a 

monthly mean in a given square, the rule is that there must be at least two daily 

mean values in that square for that particular month. So, the smallest number of 

daily mean values is 30 for any area to be considered during the entire 

observational period. Thus the study area is further restricted to the area 

enclosed by the 30-day contour in Figure 2(a). Of course, obtaining a monthly 

mean by averaging two daily values is a very crude procedure and the calculated 

monthly means are subject to large errors near the boundary of the study area. 

One has also to be very careful with the interpretation of the monthly mean 

values. Figure 2(b) gives the number of squares with non-zero monthly mean for 

each of the month from February 1996 till April 1997. From Figure 2(b), the area 

covered by monthly mean data in a particular month can be estimated. Such 

area is the smallest in February 1996 and April 1997 when the area is 40% of the 

area enclosed by the 30-day contour in Figure 2(a). The area for the two months 

of August and September is about 70% of the full area, which contains 375 

squares. 

Finally, to compare drifter velocity with altimeter data and surface winds a 

daily mean and a 10-day mean of velocity components are calculated for the 12 

meridional strips shown in Figure 2 (c). Each strip spans 0.5 degree in longitude. 

The area covered is bounded by 89W to the west, 83W to the east, the coastline 

to the north, and the 500m isobath and 28N to the south. 

The next set of data used is the TOPEX/Poseidon altimeter 

measurements available every 10 days at points shown in Figure 3 (numbered 

north to south). 

 

Figure 3. Seven points (black dots) on the 

TOPEX/Poseidon ground track where altimeter 

measurements are available every 10 days on the 

WFS. Open squares mark the location of NOAA 

buoys (42040, 42039, and 42036) where winds at 

10m above the sea surface are measured.
83

In addition, monthly mean two-dimensional fields of sea level 

reconstructed for the entire Gulf of Mexico from the TOPEX/Poseidon altimeter 

data are used (Leben R., 1998; personal communication). The third set of data 

used contains the two components of wind velocity measured at three NOAA 

stationary buoys in NEGOM (Figure 3). Daily mean wind data are available from 

February 1, 1996 through April 30, 1997. 

3. Data Analysis. 

Monthly mean fields of surface velocity obtained from positioning satellitetracked drifters are presented in Figures 4(a)-4(c). 

Figure 4. Monthly mean surface currents vectors from satellite-tracked 

drifters are shown for the study area during February 1996 – April 1997 
84

Figure 4. Continued. 
85

The currents vary significantly during the year. They are the strongest in 

October 1996 when westward flow with speed up to 25 cm/s are obtained and 

they are the weak in September 1996 and in March-April 1997 when the speed 

does not exceed 10 cm/s. The meridional structure of currents also varies from 

month to month. In May 1996, currents in the northern portion of the area are 

westward, while those in the south are eastward or southeastward. In October, 

the strong westward flow is found nearly everywhere west of 85W. The motion 

appears disorganized in September 1996 and April 1997. A close look reveals 

that the monthly mean currents behave differently east and west of 85W (dashed 

line in Figure 4). The difference is particularly noticeable in June-July 1996 and 

in October and November 1996. Shown in Figure 5 are the kinetic energy of 

mean currents averaged over the squares for the area to the west of 85W (first 

eight strips in Figure 2(c)) and for the area to the east of 85W (strips 9 to 12 in 

Figure 2 (c)). To produce the curves in Figure 5, kinetic energy values are 

obtained for each 10-day period between February 9, 1996 and April 30, 1997 

and then smoothed with a sliding mean. 

Figure 4. Continued. 
86

Figure 5 shows that the currents to the west of 85W are generally more 

intense and particularly so in October and also in March 1996 and January 1997. 

Currents to the east of 85W are generally less energetic and show peaks in May 

1996 and November 1996. It appears that the currents on the shelf to the west 

of 85W and to the east of 85W have different origins. To expose the difference, 

two sets of correlation functions are calculated. The first comprises of correlation 

functions between wind and current velocity components (daily mean data are 

used). The second set comprises those between the sea level heights (SSH) at 

the seven points shown in Figure 3 and current velocity components (10-days 

mean data are used). Calculations are done for each of the twelve strips in 

Figure 2(c). It should be mentioned that in the case of wind and current velocity, 

the lag is 1-2 days (current velocity lagging the wind). In the case of SSH and 

current velocity, the lag is 40-50 days (current lagging behind SSH). 

Distributions of maximum of correlation function (MCF) with longitude for 

both sets are shown in Figure 6. 

Figure 5. Kinetic energy of mean currents 

averaged over the squares to the west of 

85W (solid) and to the east of 85W (dashed).
87

The upper panel links the wind and current velocity. For the zonal 

components, high values of MCF are found between 85.5W and 88.5W and the 

MCF values are lower outside this range of longitudes. The MCF values are 

generally small for the meridional component, particularly to the west. In the 

case of SSH and zonal component of current velocity, high values of MCF are 

found to the east of 85W and negative values between 85W and 88W. There is 

no marked difference in magnitude between the MCF’s for the two components. 

Both components of current velocity are correlated with SSH, but the MCF 

variation in longitude shows a 180-degree phase shift. It should be noted that 

the results in the lower panel is for point 7 in Figure 3, in the Straits of Florida. 

 

Figure 6. (a) Correlation function 

maximum (CFM) between wind stress 

and surface currents for the zonal 

(squares) and meridional (circles) 

components. (b) CFM between sealevel-height (SSH) at point 7 in Figure 3 

and zonal (squares) and meridional 

(dots) surface current velocity 

component. The correlation functions 

are calculated for each of the twelve 

strips in Figure 2(c). Only the maxima 

are shown. 
88

At other TOPEX/Poseidon track points in Figure 3, the correlation is not as 

good. Figure 7 (a) shows correlation functions between SSH at points 1, 4, and 

7 in Figure 3 and velocity components. As can be seen, the best result is found 

at point 7 despite the fact that this point is the farthest away from the study area. 

It should be noted that the variation of SSH at this point is the largest (see later 

Figure 9(b)). The origin of the oscillation seems obviously related to the Loop 

Current variability, leading to similar variability on the shelf the north. 

The linkage between surface currents on the shelf to the west of 85W and 

local winds on a lag of 1-2 days seems obvious too: the currents here basically 

are wind-driven. 

Figure 7. Correlation function between SSH 

at points 1, 4, and 7 in Figure 3 and zonal (u) 

and meridional (v) surface current velocity 

components in strip 11 in Figure 2(c). 
89

 

Figure 8 shows the variability of monthly mean values of zonal and 

meridional component of surface current velocity averaged for the area in Figure 

2 (c) between 86W and 88W and the variability of the wind components there. It 

is clear that the zonal components of the current and wind are highly coherent. 

The remaining questions are: What is responsible for the observed 

surface current variability in area to the east of 85W; how SSH oscillations in the 

Straits of Florida could affect surface currents on the northeastern shelf and what 

produces the SSH oscillations at the Straits of Florida? 

Figure 8. (a) Monthly mean surface 

velocity components from drifters 

averaged over the shelf area between 

88W and 86W (strip 3 through 6 in Figure 

2 (c)). (b) Monthly mean wind 

components averaged for the three NOAA 

buoys in Figure 3. 
90

 

Drifter-based monthly mean surface current velocity components 

averaged in the area between 85W and 83W are presented at Figure 9 (a). 

Compared to similar plots in Figure 8 (a) for the western portion of the study 

area, it is clear that the surface current here is no longer simply in a east-west 

orientation. The north-south component appears equally important. The monthly 

mean SSH at point 7 in the Straits of Florida is shown in Figure 9 (b). It appears 

that with a shift of 40-50 days the SSH curve is quite similar to the zonal 

component of the surface current velocity. Also, the meridional component 

seems 180-degree phase-shifted from the SSH. The flow is strongly southward 

when the SSH is high around November and December of 1996. Point 7 

possesses the largest amplitude of SSH oscillation among all the points in Figure 

3. The difference in SSH between April and September-October 1996 and 

between September-October 1996 and April 1997 is almost 1m. Figure 9(c) 

shows the approximate location of the LC northern boundary in latitude. This 

picture is obtained using the two-dimensional SSH fields from the Historical 

Altimeter Data Viewer web site (Colorado Center for Astrodynamics Research). 

The northernmost location of zero value isoline contouring anti-cyclonic eddy 

formed by the Loop Current is fixed. Period of February-August represents a 

phase of the mature LC with a well-developed anti-cyclonic ring which is about to 

separate from the LC. According to the two-dimensional SSH distribution there 

Figure 9. (a) Monthly mean surface current 

velocity components averaged over the area 

between 85W and 83W (strips 8 through 13 in 

Figure 2 (c). (b) Monthly mean SSH at point 7 

in Figure 3. (c) Latitudinal location of the 

northern edge of the Loop Current during 1996-

97. 
91

are two imminent occurrence of ring separattion during this phase: in the mid of 

February and in the end of July - August 1996. Final detachment occurs in 

September; after that the LC collapse to the south and began to grow (a young 

phase). The SSH variability is consistent with the procession of the LC phases. 

The SSH is low during mature phase, when the LC follows the shelf break. The 

high values occur soon after the ring separation, when the LC runs into the shelf 

break near the location of point 7, giving rise to a high SSH. The step function 

structure of the imposed pressure along shelf-break leads to a flow to the south 

(Hetland et al., 1999). 

4. Summary and conclusions. 

Numerical simulation of the Gulf of Mexico with circulation driven solely by 

the LC inflow (Hetland et al., 1999) shows that the southward flow on the WFS 

varies out of phase with the northward penetration of the LC. The southward 

current on the WFS is caused by the north-to-south drop in dynamic pressure 

imposed on the shelf by the LC. When the LC is in its young phase with small 

northward penetration, it collides with the shelf break, creating a pressure drop to 

the east as the LC speeds up due to the constriction of the Straits of Florida. 

This pressure structure forces a southward flow that extends to the northern 

reaches of the WFS. When the LC is in its mature phase and its penetration to 

the north is at the fullest extent, pressure drop is much weaker, leading to a 

southward flow that may not extend too far to the north (Hetland, 1999; Hetland 

et al., 1999). The results presented here appear consistent with this, providing a 

link between surface current variability on the northeastern WFS and SSH 

variation in the Straits of Florida - the place where the SSH responds to the LC 

development. It is in some sense a surprise to find low correlation between 

currents and the wind to the east of 85.5W (Figure 6(a)). This result appears in 

contradiction with findings elsewhere (Yang et al., 1999), where the conclusion 

the seasonal winds may play a dominant role in the seasonal variations of 

surface currents at the WFs is made. Another important thing is the following. 

The wind in the northern and northeastern part of the Gulf of Mexico has 

predominant semiannual variability (Blaha and Sturges, 1978; Golubev and 

Hsueh, 2000; see also Fig.8) which can even be easily seen in the climatology 

wind data (Hellerman and Rosenstein, 1983). Different predominant harmonics 

in the wind-driven circulation (semi-annual) and the LC induced currents (almost 

annual) may be a point for separation of currents on the shelf forced by these two 

mechanisms simultaneously. 
92

Acknowledgement 

Dr. P. P. Niiler provides the drifters data. Mrs. Debria Hayes and other 

people from NODC (NOAA) have been very generous in sharing 

TOPEX/Poseidon altimetry data. The National Data Buoy Center staff helps with 

providing the wind data. The help is also acknowledged of the people at 

Colorado Center for Astrodynamics Research (CCAR) for the use the Gulf of 

Mexico Historical Altimeter Data Viewer at CCAR web site. 

REFERENCES 

Blaha, J., and W. Sturges. 1981. Evidence for wind-forced circulation in the Gulf 

of Mexico. J. Mar. Res., 39(4), 711-734. 

Golubev, Y., and Y. Hsueh. 2001. Temperature variability in the northeastern 

Gulf of Mexico in 1997-1998. Submitted to Gulf of Mexico Science. 

Hellerman, S,, and M. Rosenstein. 1983. Normal monthly wind stress over the 

world ocean with error estimates. J. Phys. Oceanogr., 13, 1093-1104. 

Hetland, R. D., Y. Hsueh, R .R. Leben, and P. P. Niiler. 1999. A Loop CurrentInduced Jet Along the Edge of the West Florida Shelf. Geophys. Res. Lett., 26, 

2239-2242. 

Hsueh, Y., and Yury Golubev. 2001. A numerical model calculation of the flow 

in DeSoto Canyon in response to northerly wind bursts in winter. Submitted to 

Gulf of Mexico Science. 

Jones,J. I. 1973. Physical oceanography of the northeast Gulf of Mexico and 

Florida continental shelf area. A Summary of Knowledge of the Eastern Gulf of 

Mexico, J.I. Jones, R.E.Ring, M.O.Rinkel and R.E.Smith, Eds., Florida Institute of 

Oceanography, St. Petersburg, 28-96. 

Ichiye, T., H.-H. Kuo, and M. R.Carnes. 1973. Assessment of currents and 

hydrography of the eastern Gulf of Mexico. Tech. Rep., Dept Oceanogr., Texas 

A&M University, College Station. 

Marmorino, G. O.. 1983. Summertime coastal currents in the Northeastern Gulf 

of Mexico. J. Phys. Oceanogr., 13, 65-77, 1983. 

Mitchum, G. T., and W. Sturges. 1982. Wind-driven currents on the West 

Florida Shelf. J. Phys. Oceanogr. 12, 1310-1317. 
93

Sturges, W. 1992. The spectrum of Loop Current variability from gappy data. J. 

Phys. Oceanogr., 22, 1245-1256. 

Sturges W. 1994. The frequency of ring separations from the Loop Current. J. 

Phys. Oceanogr., 24, 1647-1651. 

Sturges, W., and J. C. Evans. 1983. On the variability of the Loop Current in 

the Gulf of Mexico. J. Mar. Res. 41, 639-653. 

Sturges, W., and R. Leben. 2000. Frequency of ring separation from the Loop 

Current in the Gulf of Mexico: A revised estimate. J. Phys. Oceanogr., 30, 1814-

1819. 

Tolbert, W. H., and G. G. Salsman. 1964. Surface circulation of the eastern Gulf 

of Mexico as determined by drift-bottled studies. J. Geophys. Res., 69, 223-230. 

Vukovich, F. M. 1988. Loop Current boundary variations. J. Geophys. Res., 93, 

15,585-15,591. 

Vukovich, F. M. 1995. An updated evaluation of the Loop Currents eddyshedding frequency. J. Geophys. Res., 100, 8655-8659. 

Weisberg, R. H., B. D., Black, and H. J., Yang. 1996. Seasonal modulation of 

the west Florida continental shelf circulation. Geophys. Res. Lett., 23, 2247 -

2250, 1996. 

Yang, H. J., and R. H., Weisberg. 1999. Response of the West Florida Shelf 

circulation to climatological wind stress forcing. J. Geophys. Res., 104, 5301-

5320, 1999. 

Yang, H. J., R. H. Weisberg, and P. P., Niiler. 1999. Lagrangian circulation and 

forbidden zone on the West Florida Shelf Cont. Shelf Res. 19, 1221-1245. 
94

5d. West Florida shelf response to local wind forcing April 1998 
95

West Florida shelf response to local wind forcing: April 1998 

Robert H. Weisberg* 

Zhenjiang Li 

And 

Frank Muller-Karger 

College of Marine Science 

University of South Florida 

St. Petersburg, Florida 33701 

Final, April 2001 

_______________________ 

• Corresponding author address: Dr. Robert H. Weisberg, College of Marine Science, University of 

South Florida, 140 Seventh Avenue South, St. Petersburg, Florida 33701. Email: 

weisberg@marine.usf.edu
96

Abstract 

 We compare in-situ velocity and sea level data with a model 

simulation for the west Florida continental shelf during April 1998. Responses for 

three periods of downwelling and upwelling favorable winds are documented. 

Velocity profiles show coastal jets in the along-shelf direction that are 

accompanied by oppositely directed upper and lower layer flows in the acrossshelf direction. These along- and across-shelf currents and their associated 

vertical velocity comprise the fully three-dimensional inner-shelf responses to 

local wind forcing. The responses are sensitive to stratification. Large velocity 

vector turning occurs under stratified conditions, whereas less turning occurs 

under well-mixed conditions. With an initial density field representative of April 

1998 the model simulates velocity and sea level variations that agree with the 

observations. Without stratification substantial mismatches occur, both for 

currents and sea level. Local momentum input is a primary forcing agent for the 

inner-shelf circulation when the shelf is wide enough to distinguish the inner-shelf 

from the shelf-break. Inner-shelf responses, however, depend on the density 

field, which depends on local and offshore buoyancy fluxes. Modeling the innershelf responses to wind forcing requires maintenance of the density field through 

data assimilation and adequate surface, offshore, and land-derived buoyancy 

inputs. 

Dynamical analyses define the inner-shelf as the region where the surface 

and bottom boundary layers are important in the momentum balance. 

Kinematically, this is where the surface Ekman layer divergence, fed by the 

bottom Ekman layer convergence, or conversely, sets up the across-shelf 

pressure gradient. With respect to vertically integrated vorticity, the inner-shelf is 

where bottom stress torque tends to balance bottom pressure torque, as 

contrasted with the shelf-break where the primary balance is between bottom 

pressure torque and the rate of change of relative vorticity. 

Under stratified conditions the inner-shelf responds asymmetrically to 

upwelling and downwelling favorable winds, with the offshore scale and 

magnitude of the responses for upwelling favorable winds both being larger than 

for downwelling favorable winds. This asymmetry is a consequence of the 

bottom Ekman layer. With respect to the streamwise vorticity component, 

buoyancy torque by isopycnals bending into the bottom adds constructively 

(destructively) with planetary vorticity tilting by the sheared coastal jet under 

upwelling (downwelling) favorable winds causing the asymmetry. Such 

asymmetry may have important ramifications for material property distributions 

on the shelf. 
97

1. Introduction 

 Continental shelf circulation results from both local and offshore 

forcing. Local forcing is by momentum and buoyancy input at the sea surface 

and buoyancy input at the land boundary. Offshore forcing is by momentum and 

buoyancy input at the shelf-break. The partition of influence between these 

contributing factors varies with shelf location and geometry, and nowhere is this 

partition quantitatively accounted for. How continental shelves are forced thus 

remains a fundamental question for both the circulation and for the biological, 

chemical, and geological consequences of the circulation. 

The west Florida shelf (WFS) is an end member in the ensemble of 

continental shelves for which local forcing may be presumed to be of primary 

importance. This is because the WFS is wide enough for its inner-shelf to be 

distinguished from the shelf-break. Given the complex nature of continental 

shelves and the limitations of both in-situ data and models, it is useful to simplify 

the circulation problem by isolating factors and questioning how well a given data 

set can be explained by these factors. Along this vein, we ask how well coastal 

sea level and inner-shelf currents can be explained by the influence of local wind 

forcing alone. We compare in-situ data from April 1998 with a numerical model 

driven by observed time varying, but spatially uniform, winds. Our results 

suggest that inner-shelf currents may be accounted for by local wind forcing, 

given an appropriately specified initial density field. The density field, however, is 

shown to play an important role, and since the density field is determined by a 

combination of offshore and local buoyancy forcing, all such factors must be 

treated in order to produce useful model products. 

 The paper is organized as follows. Section 2 reviews background 

materials on the response of a continental shelf to synoptic wind forcing. Section 

3 introduces the in-situ data and the model used to simulate these data. 

Comparisons between model results and data are given in section 4. Motivated 

by these comparisons, albeit at limited observational locations, we expand upon 

the model flow field kinematics in section 5, where fully three-dimensional 

structures are shown. For brevity, we emphasize one complete cycle of 

responses to successive downwelling and upwelling favorable winds, and an 

asymmetry between upwelling and downwelling responses is noted. Section 6 

then exploits the model fields to develop the momentum and vorticity balances 

that help to define the inner-shelf separate from the shelf break. The results are 

summarized in section 7, where, along with discussing the asymmetry, it is 

recognized that the synoptic and seasonal scale variability can not be fully 

separated since responses of one depends upon the density field provided by the 

other. This is true for sea level, and more so for currents. 

2. Background 

 Continental shelves are dynamically complex. In relatively deep 

water the surface and bottom boundary layers are separated by a bottom slope 

constrained geostrophic flow. As the water depth decreases towards the coast 

the surface and bottom boundary layers tend to merge across the geostrophic 

interior. Here, we define the inner-shelf as the near-coastal region where the 
98

surface and bottom boundary layers are dynamically important. Our definition is 

somewhat different from that given by Lentz (1994), and this difference will be 

clarified in section 6. These boundary layer influences, coupled with the 

coastline constraint, result in kinematical and dynamical characteristics of the 

inner-shelf that may differ markedly from those occurring farther offshore. 

Generally, for synoptic scale variability where wind stress is the principal motive 

agent, the inner-shelf reacts to the presence of the coast through the 

establishment of a surface pressure gradient. This occurs through divergence 

brought about by surface and bottom Ekman layer transports (e.g., Gill, 1982, 

p394), and recent evidence for such adjustment on the WFS is given by the 

upwelling case study of Weisberg et al. (2000). Regularly occurring sub-tidal sea 

level variations show this adjustment to be a common feature of all shelves. 

 Descriptions of inner-shelf variability have emerged largely from 

observations on the continental shelves and Great Lakes of North America. 

Csanady (1982), Walsh (1988), and Brink et al. (1998) provide overviews, and 

recent representative works from different regions include those of Lentz (1994), 

Smith (1995), Weisberg et al. (1996), Weatherly and Thistle, (1997), Yankovsky 

and Garvine (1998), Lentz et al. (1999), and Munchow and Chant (2000). 

Insights on the dynamics of the inner-shelf are found in the analytical treatment 

of Mitchum and Clarke (1986). The complex nature of turbulent boundary layers, 

however, limits the use of analytical techniques, and this has led to numerical 

model studies of inner-shelf turbulence in simulations of one or two dimensions 

(e.g., Weatherly and Martin, 1978; Allen et al., 1995; Lentz, 1995). Two 

dimensions, however, exclude the interactions that may occur between local and 

large-scale processes and the possibility that the dynamical interactions 

controlling divergence and vertical motion may vary in the along-shelf direction. 

For example, in contrast with the two-dimensional circulation pattern of equal and 

opposite upper and lower layer flows reported by Lentz (1994) for the northern 

California coast, Munchow and Chant (2000) report a fully three-dimensional flow 

with a net across-shelf transport for the New Jersey coast. Three-dimensionality 

there is attributed to an along-shelf pressure gradient force due to coastal 

geometry and baroclinicity. Long-wave models that assume a boundary 

condition of no net across-shelf flow at some distance from the coast (e.g., 

Clarke and Van Gorder, 1986; Lopez and Clarke, 1989) are to some extent 

contrary to these findings, and additional sensitivity studies such as Samelson 

(1997) are warranted. While the general physical concepts espoused in all of 

these studies apply, it is becoming increasingly clear that each individual 

continental shelf environment is somewhat unique, and that the region of the 

inner-shelf, where important across-shelf transports occur, requires better 

understanding. 

3. The Observations and the Model 

3.1 Observations 

 Exploratory measurements of inner-shelf velocity profiles on the 

WFS were initiated in November 1996 with the deployment on the 20 m isobath 

offshore of Sarasota, FL of a bottom-mounted 300 kHz acoustic Doppler current 
99

profiler (ADCP) manufactured by RD Instruments, Inc. Using 0.5 m bin spacing, 

and after editing surface effects, the data set yielded horizontal velocity profiles 

between depths of 2.5 m to 16 m. Ancillary data sets include sea level from the 

NOAA tide gauge at St. Petersburg, FL, surface winds from the NOAA buoy 

number 42036 located at mid-shelf, and velocity profile data from a buoy located 

at the 50 m isobath. These measurement locations are shown in Figure 1. The 

data collection effort was initially unsupported by hydrographic data. Through 

efforts by the Mote Marine Laboratory (G. Kirkpatrick, personal communication, 

1998) and the University of South Florida, monthly hydrographic sections were 

initiated in March 1998. 

 Co-variations in winds, sea level, and currents occur throughout the 

record, as expected from previous WFS shelf observations (e.g., Mitchum and 

Sturges, 1982). A new finding is the variability in the relative turning of the 

velocity vectors with depth for each synoptic weather event. Responses to 

upwelling favorable winds generally show onshore flow near the bottom and 

offshore flow near the surface, and conversely for responses to downwelling 

favorable winds. These responses are oftentimes accompanied by bottom 

temperature changes. Once the supportive hydrographic sections commenced it 

became clear that the velocity vector turning is related to the density 

stratification. April 1998 is a particularly interesting month with a succession of 

synoptic weather events and large velocity vector turning. We focus our attention 

on this month for a comparison between the in-situ data and a numerical model 

simulation. 

 The April 1998 data are shown in Figure 2. With emphasis on the 

time scales of the synoptic weather variations, all of the data are low-pass filtered 

to remove oscillations at time scales shorter than 36 hours (subsequent model 

results are also low-pass filtered). Provided are velocity component isotachs for 

the along-shelf and across-shelf directions, coastal sea level, wind velocity 

vectors, and wind stress vectors. Along-shelf is defined as 3330

 relative to the 

true north. Clear contours are in this up-coast direction and shaded contours are 

in the reciprocal down-coast direction. For the across-shelf component clear 

(shaded) contours are directed onshore (offshore). The wind velocity (stress) 

variations for April 1998 are primarily in the along-shelf direction. Upwelling and 

downwelling favorable winds are observed over three distinctive periods, and sea 

level responds to each of these. In the along-shelf direction, each wind change 

gives rise to a coastal jet of varying vertical structure, i.e., some show subsurface 

maxima while others show surface maxima. Accompanying the coastal jets are 

opposing onshore and offshore flows over the upper and lower portions of the 

water column. The most pronounced response occurs on April 12 when the 

upwelling favorable winds are largest causing a coastal jet of magnitude 0.6 ms-1

with onshore (offshore) flow in the bottom (surface) Ekman layer exceeding 0.2 

ms-1. Although the wind stress magnitudes do not differ much from downwelling 

to upwelling events (April 12 being an exception), the responses in the currents 

to the upwelling winds are larger in all three cases than the responses to the 

downwelling winds. A response asymmetry on the WFS is thus observed. 

Asymmetric behavior in bottom boundary layer thickness has been reported 
100

elsewhere (e.g., Trowbridge and Lentz, 1991 and Lentz and Trowbridge, 1991 for 

the northern California coast), and the destabilizing or stabilizing influences of 

either down or up slope density transports, respectively, has been offered in 

explanation. As we will discuss, however, these boundary layer turbulence 

arguments do not explain the response asymmetry reported here. 

3.2 Model 

We employ the Princeton Ocean Model (Blumberg and Mellor, 1987) 

based on the following attributes. As a primitive equation model, it allows us to 

diagnose dynamical balances. Its vertical sigma coordinate helps to resolve the 

flow structures on a gently sloping shelf, and its horizontal orthogonal curvilinear 

coordinates allows the model grid to conform to topography. By parameterizing 

friction with an embedded turbulence closure sub-model (Mellor and Yamada, 

1982) the frictional forces evolve with the flow and density fields. 

The model domain (Figure 1) extends about 750 km along-shelf from the 

Florida Panhandle to the Florida Keys and about 400 km offshore from the coast. 

It covers the entire WFS and part of the eastern Gulf of Mexico so that the shelfbreak region is resolved. Realistic shelf bathymetry from the National Center for 

Atmospheric Research is used with 1500 m being the maximum depth beyond 

the shelf-break and 5 m being the minimum depth at the coastline. The 

horizontal grids range in size from 6.5 km to 15 km, with an average of 9 km. 16 

sigma levels are used in the vertical, distributed logarithmically about the middle 

to achieve finer resolution of the surface and bottom boundary layers relative to 

the interior. Other WFS applications of this model grid, along with discussions of 

the open boundary conditions, are given by Li (1998) and Li and Weisberg 

(1999a,b). 

The model is forced from mid-March through April 1998 by a spatially 

uniform, time-dependent wind stress. The wind stress components are 

computed from the in-situ wind velocity vectors observed at the NOAA buoy 

using a wind speed dependent drag coefficient (Wu, 1980). Forcings by surface 

heat and coastal buoyancy fluxes are excluded due to a lack of data. 

To initialize the density field we use hydrographic sections from March and 

May 1998; April 1998 hydrography is not available. Each section shows a welldefined pycnocline. On this basis we initialize the model density field with a 4 

sigma-t density change that spans 20 m in the vertical, centered on 20 m depth. 

To avoid specious baroclinic adjustments, this initial density field is input without 

horizontal gradient. Subsequent density field evolution occurs in dynamical 

balance with the applied forcing. A twin experiment with density set constant at 

1023 kgm-3 is also performed to compare the responses under stratified and 

unstratified conditions. Clarke and Brink (1985), through scale analysis of the 

shallow water equations, show that the response of a wide shelf [with small 

Burger number, (LR/L)2, where LR is the Rossby radius of deformation at the 

shelf-break and L is the shelf width] to synoptic wind fluctuations should be 

barotropic. Mitchum and Clarke (1986) suggest that (LR/L)2

 <10-2 for the WFS, 

and this magnitude is consistent with the stratification used here. The stratified 

and constant density experiments provide an opportunity to examine this 
101

barotropic response concept with a primitive equation model. 15 days of model 

start up is performed using observed winds prior to sampling the model results 

for April 1998. Since we find the pressure field response to be largely barotropic, 

consistent with the Burger number argument, the spin up time is on the order of a 

pendulum day and much less than the 15 days that we used. 

4. Comparison Between the in-situ Data and the Model Results 

4.1 Sea level 

Comparisons between sea level observed and modeled at St. Petersburg, 

FL for the month of April 1998 are shown in Figure 3. The upper and middle 

panels are for the constant density and stratified cases, respectively, and the 

lower panel shows the wind stress vector time series used to force the model. 

Agreement is very good for the stratified case; it is not as good for the constant 

density case. While the general patterns of sea level responses to upwelling and 

downwelling favorable winds are replicated in both cases, considerable 

quantitative mismatch exists between in the model responses and the 

observations for the constant density case. In particular, the overestimate for the 

upwelling response around day 23 implies an excessive across-shelf volume 

transport over the inner-shelf region during this event. Since a vertically 

integrated across-shelf volume transport in excess of the volumetric rate of sea 

level change can only occur in a three-dimensional flow, such disparity on 

agreements with in-situ data between the stratified and constant density model 

simulations implies that stratification significantly affects the three-dimensional 

flow field. 

4.2 Currents 

 Modeled along-shelf and across-shelf components of velocity from 

the 20 m isobath off Sarasota, FL are shown in Figure 4a. As is the case for sea 

level under stratified conditions, the model reproduces the general pattern 

evolution of the observed velocity component variations (Figure 2). All three 

upwelling favorable wind events drive down-coast jets with vertical structures 

similar to those observed, some with subsurface and others with near surface 

maxima. These down-coast jets are accompanied by onshore (offshore) flow 

over the lower (upper) portion of the water column. Similar statements apply for 

the responses to the downwelling favorable wind events, i.e., up-coast jets and 

across-shelf flow reversals with depth. While the pattern evolutions for the 

modeled and observed currents are generally similar, differences in details exist 

for several reasons. First, the model simulations are highly dependent upon 

stratification, and our initial stratification is merely a guess based upon limited 

measurements. Second, no provision is made for (unknown) surface heat fluxes 

or for correcting inevitable departures of the model stratification from nature over 

the course of the 45-day integration. Third, we use spatially uniform winds when 

the spatial structure of each synoptic weather system may be different. Despite 

these shortcomings, the quantitative agreements are surprisingly good, with the 

best agreements occurring during the downwelling/upwelling sequence of April 7-

13. Here the model underestimates the magnitude of the along- and across-shelf 
102

components by about 20%. For other sequences this magnitude mismatch is 

either larger or smaller, and for the reasons stated above there is little gained by 

more detailed comparisons. One clear deficiency comparing the model with the 

data lies within the upper 3 m of the water column where there is no data. The 

model purports large vertical shear over the upper few meters because its 

turbulence closure routine yields small eddy viscosity (turbulence length scale is 

zero at the surface). Whether or not such near surface shear is realistic remains 

to be determined. 

 Also depicted in Figure 4a are the modeled isopycnal variations. 

Relative maxima in isopcynal displacements occur when the across-shelf 

component is zero suggesting that advective changes through horizontal 

divergence and vertical velocity are the primary contributors. Across-shelf 

advection over the lower portion of the water column occurs as a boundary layer 

response to the coastal jet. Since the coastal jet tends to be maximum when the 

sea level response is maximum, across-shelf and vertical advection are also 

maxima at that time. Thus, upwelling and downwelling isopycnal displacement 

maxima occur at the transitions between upwelling and downwelling favorable 

winds, accounting for the lag between the isopycnal displacements and sea level. 

This phase relationship constrasts that of an inviscid, standing baroclinic mode 

for which the isopycnal and sea level responses occur in an anti-phase 

relationship, demonstrating the importance of the frictional boundary layers. By 

not imposing a surface heat flux the isopycnals also tend to slope up with time 

reducing the stratification. 

 As with sea level, the twin experiment without stratification yields 

results for the currents (Figure 4b) that are both qualitatively and quantitatively 

different from the observations. For the along-shelf component the vertical shear 

is much more uniform with depth, but the more interesting contrast is in the 

across-shelf component. Without stratification there is very little turning in the 

surface and bottom boundary layers, and hence there is very little across-shelf 

circulation relative to the results with stratification. This behavior for the acrossshelf component mimics the longer term observational record that shows 

pronounced velocity vector turning during times of anticipated stratification 

versus little turning during times when the water column is expected to be well 

mixed. 

 Based on the general agreements between the in-situ data and the 

stratified model results we now use the model to discuss the flow field kinematics 

and dynamics with emphasis on the downwelling/upwelling response sequence 

from April 7-14. 

5. Model Flow Field Kinematics 

 The order of presentation includes planar views of sea level, crosssections of velocity and density at the Sarasota transect, and planar views of 

currents. We begin with daily snapshots of sea level from day 7 through day 14 

(Figure 5a) for the stratified experiment. Peak downwelling and upwelling 

responses in the sea level fields (as with all other fields) appear on days 9 and 

12, respectively. During these peaks the set-down at the coast for the upwelling 
103

response in general is more than a factor of two larger than the set up at the 

coast for the downwelling response. Moreover, the offshore scale for the 

upwelling response largely exceeds that for the downwelling response. The 

transitions to these peak responses show along- and across-shelf surface height 

gradients of comparable magnitude. While no clear sense of propagation 

emerges, it does appear that the along-shelf variations in geometry, particularly 

the relatively narrow shelf in the north and the partial closure by the Florida Keys 

in the south, play roles in determining the sea level adjustments. The end results 

are monotonic sea level slopes with coastal maxima, but the transitions to these 

end states are more convoluted, suggestive of complicated flow patterns. 

 Similar plots for the constant density experiment are given in Figure 

5b. Contrasted with the asymmetric responses of Figure5a (in both magnitude 

and offshore scale), the constant density experiment yields sea level 

displacement magnitudes that are nearly proportionate with the wind stress and 

that have comparable offshore scales for either upwelling or downwelling. 

 Cross-sections for the along-shelf, across-shelf, and vertical 

velocity components, and density offshore of Sarasota, FL. are shown in Figure 

6a. Our focus is again on the day 9 and day 12 peak responses. On day 9 we 

see an up-coast jet in the along-shelf component, with onshore flow over the 

upper half and offshore flow over the lower half of the water column in the 

across-shelf component, and with downwelling occurring synchronously with the 

offshore flow. In a parallel, but opposite sense, on day 12 we see a down-coast 

jet in the along-shelf component, with offshore flow over the upper half and 

onshore flow over the lower half of the water column in the across-shelf 

component, and with upwelling occurring synchronously with the onshore flow. 

Due to the frictional effects of the boundary layers the coastal jet has a vertical 

shear structure with isotachs bending offshore with depth. This accounts for the 

subsurface maxima that occur in both the in-situ data and the model simulation. 

In summing the geostrophic interior with the frictional boundary layer flows the 

velocity vector turns counterclockwise over most of the water column. This is 

shown for the WFS by Weisberg et al. (2000) and for the New Jersey shelf by 

Munchow and Chant (2000). As is the case for sea level, the magnitudes and 

offshore scale for the peak upwelling response are substantially larger than for 

the peak of the downwelling response. Since the bottom Ekman layer is a 

response to the coastal jet, across-shelf transports over the lower portion of 

water occur over the entire offshore extent of the coastal jet for either upwelling 

or downwelling situations. Since the upwelling response exists over a larger 

offshore scale than the downwelling response, the across-shelf transports of 

water for upwelling occur over larger distances than for downwelling. 

 Similar plots for the constant density experiment are given in Figure 

6b. As found for sea level, the asymmetry in the currents under stratified 

conditions is replaced by a more proportionate response under constant density. 

With the offshore scales of the coastal jet the surface displacement being equal, 

the difference in scales for the upwelling and downwelling responses under 

stratified and constant density conditions parallels that in Figures 5a,b. 
104

 Evidence for this model-based assertion of asymmetry exists in 

simultaneous in-situ data that were collected on the 50 m isobath (Figure 7). The 

same coastal jets as seen in shallower water (Figure 2) appear here, but with 

less vertical structure, as in the model. For each wind-driven flow reversal we 

see that the near surface Ekman layer response (extending down to about 20 m 

depth) is larger than the near bottom Ekman layer response, also consistent with 

the model. Moreover, for the peak downwelling response around day 9 we see 

nearly zero across-shelf flow in the bottom Ekman layer versus about 0.05 msec1

 onshore flow in the bottom Ekman layer for the peak upwelling response 

around day 12. These observations are consistent with larger offshore extents 

for the surface height field and the coastal jet under upwelling favorable winds. 

 The horizontal structure of the velocity field is complex. Horizontal 

velocity vector maps for the depth average, near surface, mid-depth, and near 

bottom fields are shown in the Figure 8 panels a-d, respectively, for the day 9 

peak downwelling response. The depth average flow is the simplest of the fields. 

An up-coast directed jet, confined to the proximity of the shore, is seen with 

maximum speeds occurring where the sea surface height gradient is maximum. 

The jet is fed by surface Ekman layer transports that are only partially offset by 

bottom Ekman layer transports since the bottom Ekman layer is effected only 

where the surface height gradient and coastal jet are established. Thus the 

coastal jet accelerates downstream. The character of the near bottom flow also 

changes from the inner-shelf region, where it is related to the coastal jet, to the 

shelf break region where it arises by topographic wave variability. The origin of 

these near bottom flow variations will be evident in the section 6 dynamical 

analyses. The mid-depth vectors look very similar to the depth average vectors 

since there the across-shelf flows due to the opposing surface and bottom 

Ekman layer flows are minimal. 

 A parallel set of velocity maps are given in Figure 9 for the day 12 

peak upwelling response. Similarities exist with Figure 8 (albeit with opposite 

sign). Contrasting these two figures, however, is the larger offshore scale for the 

upwelling fields. Strong onshore flows in the bottom Ekman layer exist out to 

about the 100 m isobath, particularly in the northern part of the domain where the 

magnitude of the coastal jet is largest. Bottom Ekman layer flows in the southern 

portion of the domain are reduced by the adverse pressure gradient due to the 

partial closure by the Florida Keys. Stratification also tends to promote regions of 

flow field confluence as seen at mid-depth. 

 A result of the flow field kinematics in switching from downwelling to 

upwelling favorable winds is manifest in sea surface temperature (SST). Figures 

10a,b show satellite AVHRR SST images on April 10 and 13 when the 

downwelling and upwelling circulations, respectively, culminate in isopycnal 

displacement maxima. The month of April, in general, shows a cold tongue that 

extends down-coast at mid-shelf from a region of seasonally coldest SST in the 

northeastern Gulf of Mexico. Associated with this boreal spring cold tongue is a 

chlorophyll plume, referred to as the "green river" by Gilbes et al. (1995). This 

cold tongue is evident on April 10 with warmer waters both inshore and offshore. 

Transitioning from downwelling to upwelling winds, the cold tongue on April 13 is 
105

greatly amplified and the inshore waters are several degrees colder than on April 

10. A region of locally coldest SST is also seen south of Tampa Bay on about 

the 25 m isobath consistent both with the region of maximum vertical velocity in 

Figure 6a and with the upwelling case study results of Weisberg et al. (2000). 

Farther south is a diminution of cold SST consistent with the bottom Ekman layer 

inhibition by adverse pressure gradient shown in Figure 9. Right at the shoreline 

in the Florida Big Bend region and along the Florida Panhandle west of 

Apalachicola Bay we also see eruptions of cold water to the surface. While we 

do not imply that surface heat flux is not a contributing factor, these SST changes 

are consistent with the kinematical features of the numerical model simulation 

described in this section. 

6. Model Flow Field Dynamics 

6.1 Momentum Balances 

 We begin with vertically integrated momentum and mass balances 

(Figure 11) calculated at the 20 m isobath offshore from Sarasota, FL. The wind 

stress is the lead term in the along-shelf direction (panel a). This causes a local 

acceleration as the currents tend to adjust to the time varying winds, followed by 

an along-shelf pressure gradient, a bottom stress, and a Coriolis acceleration, 

with all of these response terms being of comparable magnitude (although the 

along-shelf pressure gradient is generally the largest of these). In a purely two 

dimensional flow the bottom stress would be comparable to the surface stress, 

whereas in this three-dimensional result the bottom stress is no larger than the 

other secondary terms. The vertically integrated across-shelf momentum 

balance (panel b) is much simpler (note the scale change between panels a and 

b). Here the vertically integrated flow is essentially geostrophic since the 

ageostrophic contributions from the surface and bottom Ekman layers tend to 

cancel. The magnitudes of the terms in the across-shelf momentum balance are 

about a factor of four larger than those in the along-shelf momentum balance. 

 The three-dimensionality of the flow field is also evident in the 

vertically integrated across-shelf mass transport (panel c). The four time series 

are: 1) the Ekman transport associated with the wind stress, 2) the vertically 

integrated upper and 3) lower layer transports directed either onshore or offshore 

depending on wind stress, and 4) the total vertically integrated transport. The 

Ekman transport is computed as /( ) 0 U f w

e = τ ρ , where w τ is the along-shelf 

component of the wind stress. The upper and lower layer transports are defined 

as U u dz

r d

upper upper ∫ =

η

, U u dz

dr

H

lower lower ∫

−

= , where u is the velocity component in the 

across-shelf direction, r d is the depth at which the across-shelf flow reverses 

direction, η is the free surface elevation, and H is the mean water depth. The 

net across-shelf transport is U udz

H

net ∫

−

=

η

. The upper layer transport is always 

less than the Ekman transport because the 20 m isobath is within the inner-shelf 

where there occurs an upper layer across-shelf transport divergence along with 
106

the surface slope set-up. Seaward of the inner-shelf the upper layer transport is 

comparable to the Ekman transport. The upper layer transport at this particular 

location is generally over compensated for by the lower layer transport. A 

contributing factor is the along-shelf pressure gradient that supports an acrossshelf geostrophic transport. The along-shelf pressure gradient appears to be 

larger for upwelling favorable winds than for downwelling favorable winds (panel 

a). With mass convergence/divergence inferred in the across-shelf direction 

there must be a mass divergence/convergence in the along-shelf direction to 

maintain continuity. The coastal jets offshore of Sarasota, FL are therefore 

spatially accelerating jets, and a consequence of this for upwelling favorable 

winds is an increase in the intensity of upwelling over what would occur without 

the spatial acceleration. This is consistent with the region of maximum surface 

cooling shown in Figure 10b. 

 An asymmetry in across-shelf transport between upwelling and 

downwelling is also evident in panel c. This asymmetry arises from the bottom 

Ekman layer as demonstrated by the ageostrophic portion of the across-shelf 

momentum balance (panel d). The ageostrophic part is defined as the residual 

after adding the Coriolis and pressure gradient terms. Both the magnitude and 

vertical extent of the ageostrophic residual that comprises the bottom Ekman 

layer at this location are larger for the three upwelling events than for the 

downwelling events (in particular, compare the April 9 and 12 patterns). This 

finding is opposite to what is expected from stability arguments for up and down 

slope flows in a stratified fluid (Trowbridge and Lentz, 1991). Here the bottom 

boundary layer is consistently larger for upwelling than for downwelling 

conditions! 

 To further explore these issues we analyze the across-shelf 

momentum balance over the Sarasota, FL. transect. Figure 12 shows the results 

for the peak downwelling event of April 9. Eight panels provide the sequential 

order of terms that comprise the momentum balance in this coordinate direction. 

Panels a and b are the Coriolis and pressure gradient terms. Their sum (the 

ageostrophic residual) is given in panel c for comparison with the frictional terms 

of panel d (where friction is almost entirely of vertical origin). The sum of panels 

c and d, in panel e, is for comparison with the advective acceleration terms in 

panel f. Finally, the sum of panels e and f, in panel g, is for comparison with the 

only remaining term, the local acceleration, in panel h. Closure is observed 

showing that the calculation is performed correctly. The lead terms are the 

Coriolis and pressure gradient terms. The ageostrophic residual is accounted for 

primarily by the surface and bottom Ekman layers, and the only other term of 

substance is the local acceleration. The pressure gradient is maximum near the 

coast and it decreases to zero by about the 50 m isobath. Modified by surface 

and bottom friction, the coastal jet seen in the Coriolis term follows the pressure 

gradient. The pressure gradient is relatively depth independent, i.e., it is 

primarily barotropic consistent with the Burger number arguments of Clarke and 

Brink (1985). The pressure gradient, however, does decrease across the bottom 

Ekman layer due to the sloping isopycnals there. 
107

 The comparable analysis for the peak upwelling event of day 12 is 

shown in Figure 13. The ordering of terms is the same as in Figure 12. The 

difference is in the magnitude and the offshore scale of the Coriolis and pressure 

gradient terms. Here the pressure gradient nodal line extends out to the 100 m 

isobath, the coastal jet is much stronger, and hence the bottom Ekman layer is 

much more developed. Opposite to the downwelling case, the pressure gradient 

increases across the bottom Ekman layer due to the slope of the isopycnals 

there. We will return to this point in section 7. 

6.2 Vorticity Balance 

 Our analysis of the vertically integrated, vertical component of 

vorticity is similar to that of Ezer and Mellor (1994). Taking the curl of the 

vertically integrated momentum equations, we discuss the vorticity balance with 

respect to four terms: 1) stress torque, 2) bottom pressure torque, 3) stretching of 

planetary vorticity by free surface deformation (plus the planetary beta effect), 

and 4) the material rate of change of relative vorticity (plus the stretching and 

tilting of relative vorticity by the flow field). Since we employed a spatially uniform 

wind stress, the stress torque is entirely due to bottom stress. The bottom 

pressure torque, through the bottom kinematic boundary condition, is the 

stretching of planetary vorticity by geostrophic flow across the sloping bottom. 

This is generally much larger than the stretching by free surface deformation (or 

by the planetary beta effect that we lumped together with free surface term). The 

residual of these tendencies is the variation in relative vorticity (that we lumped 

together with the kinematical terms associated with the flow field deformations of 

relative vorticity and with the relatively insignificant horizontal diffusion terms). 

 These four terms calculated at the 20 m isobath offshore of 

Sarasota, FL. for April 1998 are given in Figure 14. The bottom stress and 

bottom pressure torques are the lead terms, although the other two terms may be 

of comparable magnitude during the transitions when the sea surface slope is 

rapidly changing and the coastal jets are weak. Since these time series are 

spatially dependent their horizontal fields help to define the inner-shelf. 

 Figure 15 shows the four fields for the day 9 peak downwelling 

response. Once the pressure field and coastal jets are established, the innershelf is the region where the primary vorticity balance is between the bottom 

stress and bottom pressure torques. At this time the stretching of planetary 

vorticity by free surface deformation is small everywhere except for regions of 

strong eddies that form along the shelf break. The bottom pressure torque is 

large everywhere, but this tendency to induce relative vorticity can only be offset 

where large near bottom currents are capable of producing large bottom 

stresses, i.e., only in the inner-shelf region for this local wind forced experiment. 

Elsewhere the imbalance in these vorticity tendency terms leads to a large 

material rate of change of relative vorticity. Thus, the primary balance for the 

inner-shelf is between the bottom stress and bottom pressure torques, whereas 

the primary balance for the shelf-break is between bottom pressure torque and 

material rate of change of relative vorticity. Similar field representations for the 

day 12 peak upwelling response are shown in Figure 16. The conclusions are 
108

the same as with Figure15 with the added point of asymmetry. The offshore 

scale of the inner-shelf is larger for the day 12 upwelling response than it is for 

the day 9 downwelling response. 

7. Summary and Discussion 

 How water move across the shelf is the underlying physical 

oceanographic question pertinent to multidisciplinary continental shelf studies. 

Boundary layer and eddy transport processes are involved because of the 

vorticity constraint by the sloping bottom. Understanding these processes 

presupposes an understanding of the shelf responses to its various forcing 

agents: momentum input at the sea surface and shelf break and buoyancy input 

at the sea surface, shelf break, and land. Since each continental shelf is 

different, and since so many processes are involved, it is useful to isolate regions 

and forcing conditions to answer smaller subsets of questions. Along this vein 

we focus on the WFS, which is wide enough for its inner-shelf to be distinguished 

from the shelf-break. Using in-situ data and a numerical model simulation we 

ask how well the currents over the inner-shelf are accounted for by local wind 

forcing alone. In doing this we describe the boundary layer effects that account 

for the transports across the inner-shelf, provide a dynamical definition of the 

inner-shelf, and describe an asymmetry in the inner-shelf responses to upwelling 

versus downwelling favorable winds. 

The model and in-situ data comparison shows that a primitive equation 

model (in this case the sigma coordinate POM with Mellor/Yamada turbulence 

closure) is effective at describing the sea level and circulation of the inner-shelf 

when forced by local winds only. We conclude that local winds are a major driver 

of the inner-shelf circulation. However, stratification impacts the nature of the 

circulation and the across-shelf transports, and it does this by affecting 

turbulence and hence the vertical scales of the surface and bottom Ekman 

layers. When these layers are vertically constrained, through the suppression of 

turbulence by stratification, the velocity vector turning within the boundary layers 

increases over that for constant density. Stratification, in this way, largely 

increases across-shelf mass transport in response to local wind forcing. 

Flow field kinematics and dynamics analyses show the WFS circulation to 

be fully three-dimensional (along-shore variations are significant). These 

analyses define the inner-shelf as the region where the surface and bottom 

Ekman layers play an important role in the momentum balance. Kinematically 

the inner-shelf is where the Ekman layers induce a sea surface slope through 

divergence. With respect to vertically integrated vorticity the inner-shelf is the 

region where the primary balance is between bottom stress and bottom pressure 

torques, as contrasted with the shelf break where the primary balance is between 

the bottom pressure torque and the material rate of change of relative vorticity. 

An asymmetry is found in the WFS response to upwelling and 

downwelling favorable winds, wherein the magnitude and offshore extent of the 

responses to upwelling favorable winds exceed those for downwelling favorable 

winds under stratified conditions. While this asymmetry appears in all of the 

dynamical analyses, these analyses alone do not provide an explanation. The 
109

asymmetry is due to stratification as demonstrated by comparing model twin 

experiments, one with and the other without stratification. Asymmetry only 

occurs in the stratified experiment. 

While we are not aware of such downwelling/upwelling asymmetry being 

reported elsewhere, MacCready and Rhines (1991) and Garrett et al. (1993) 

provide a conceptual and analytical basis for it. Their argument is that 

stratification impedes across-shelf transport in the bottom Ekman layer on a 

slope when the buoyancy force tends to balance the Coriolis force of the alongshelf velocity component. We diagnose this concept from both momentum and 

vorticity perspectives. With respect to momentum, the sea level response to 

downwelling winds results in an offshore directed pressure gradient force that 

drives the offshore directed flow in the bottom Ekman layer. Isopycnals bending 

into the bottom result in a buoyancy force that opposes this (Figure 12). In 

contrast to downwelling, the sea level response to upwelling winds results in an 

onshore directed pressure gradient force that drives an onshore directed flow in 

the bottom Ekman layer. Isopycnals bending into the bottom now result in a 

buoyancy force that enhances this (Figure 13). Granted, fewer isopycnals bend 

into the bottom; nevertheless the ones that do act constructively rather than 

destructively. 

The simplest explanation for the asymmetry derives from the streamwise 

component of vorticity. For a stratified Ekman layer we consider the balance 

between the tendencies to induce relative vorticity by: 1) the tilting of planetary 

vorticity by the vertically sheared along-shelf jet, 2) the buoyancy torque by the 

sloping isopycnals, and 3) the dissipation of relative vorticity by the vertically 

sheared across-shelf flow. Equation (1) expresses this balance, with terms 1-3 

on the left hand side and the local rate of change of relative vorticity, plus the 

omitted residual terms on the right hand side: 

R

z

u

z t

u K

x z

g

z

v f +

∂

∂

∂

∂ = ∂∂

∂

∂

+

∂

∂

+

∂

∂ ( ) ( ) 2

2

0

ρ

ρ

 (1) 

where u , v are the across-shelf and along-shelf velocity components, 

respectively, ρ is density, K is the vertical eddy coefficient, and R is the 

residual. For downwelling, planetary vorticity tilting tends to be balanced by 

buoyancy torque, requiring less relative vorticity dissipation. For upwelling, the 

buoyancy torque, while reduced in magnitude from the downwelling case, adds 

to the planetary vorticity tilting, requiring increased relative vorticity dissipation. 

We demonstrate these effects in Figure 17 by comparing the peak downwelling 

and upwelling responses on days 9 and 12, respectively. Destructive 

interference between planetary vorticity tilting and buoyancy torque is observed 

for downwelling, versus constructive interference for upwelling. Similar results 

(not shown) are found for the other two downwelling and upwelling pairs. The 

buoyancy torque (by isopycnals bending into the sloping bottom) has the same 

sign for downwelling and upwelling, whereas the tilting term reverses sign. 
110

A corollary affect, expanded on by Garrett et al. (1993) in relationship to 

the Lentz and Trowbridge (1991) and Trowbridge and Lentz (1991) findings, is an 

asymmetry in the bottom boundary layer vertical scale due to the 

stabilizing/destabilizing influences of up and down slope flows. While related, 

this is a different phenomenon that appears to be contrary to our finding (in 

Figure 11) that the bottom boundary layer scale is actually larger for upwelling. 

We reconcile this discrepancy by noting that our larger boundary layer scale for 

upwelling follows from the larger upwelling response. Had the upwelling and 

downwelling responses on the WFS been of equal magnitude with regard to the 

interior along-shelf flows, as in the numerical experiments reported by Garrett et 

al. (1993), then we would have expected results similar to theirs. Asymmetry in 

the bottom boundary layer is important in nature for the reasons espoused in 

these earlier papers, including Weatherly and Martin (1978). 

The inner-shelf provides added importance to these stratified boundary 

layer concepts on a slope. For the inner-shelf, the surface pressure gradient set 

up by surface Ekman layer divergence, the geostrophic interior flow adjustment 

to the pressure gradient, and the bottom Ekman layer reaction to the interior flow 

all occur nearly in unison. Anything that impedes one of these three steps will 

impede all of them. Thus, by inhibiting (or promoting) the bottom Ekman layer, 

thereby reducing (or increasing) the near bottom divergence that is necessary to 

feed the near surface divergence, stratification produces the asymmetry in the 

magnitude and offshore extent of the responses to downwelling and upwelling 

favorable winds. 

A physical consequence of asymmetry is the rectification of the shelf 

response to oscillatory winds. This provides a possible explanation of why the 

region south of Tampa Bay was found by Yang et al. (1999) to be devoid of 

surface drifter tracks during a year-long set of deployments. Stratificationinduced asymmetry may also have important biological, chemical, and geological 

consequences. In analogy to a flapper valve, with upwelling responses favored 

over downwelling responses, near bottom material properties are more readily 

transferred into, than away from, the near-shore zone. Support for this comes 

from unpublished hydrographic data (G. Vargo, personal communication) that 

shows large chlorophyll fluorescence extending across the shelf in the bottom 

Ekman layer. In the along-shelf direction the rectification of material property 

transports by the coastal jets may also be important; impacting, for instance, 

long-term sediment transport. Further theoretical and observational studies are 

needed to explore the ramifications of stratification induced asymmetry on the 

continental shelf. 

A related conclusion is that the seasonal and synoptic scales can not be 

fully separated. With synoptic scale wind forced responses dependent on 

stratification, it is necessary to either assimilate density data into a model, or 

simulate density changes through a combination of surface buoyancy fluxes and 

active offshore boundary conditions. In our case, a 45 day (15 days of start-up 

and 30 days of analysis) simulation worked well, whereas attempts to carry the 

model integration longer met with increasing deviations from the in-situ data due 

to density field changes. Nowcasting and forecasting of the inner-shelf must be 
111

supported by sufficient in-situ data for density field assimilation and boundary 

conditions. 

 Acknowledgements. Support for this work derived from the Office 

of Naval Research, grant N00014-98-1-0158; the National Oceanic and 

Atmospheric Administration, grant NA76RG0463; and the Minerals Management 

Service, contract MMS 14-35-0001-30804. G. Mitchum offered helpful 

discussions; R. Cole, J. Donovan, R. He, and W Hemme assisted with the data 

collection and analyses. 
112

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115

List of Figures 

Figure 1 West Florida shelf bathymetry and the measurement 

locations (upper panel) and numerical model domain with coordinate system 

description as applied herein (lower panel). The model seaward boundaries are 

open except for the southeast corner where the solid line represents closure by 

the Florida Keys. The solid line offshore from Sarasota denotes the section 

sampled for kinematical and dynamical analyses. 

Figure 2 April 1998 measurements of: a) along-shelf and b) acrossshelf velocity components sampled at the 20 m isobath offshore of Sarasota FL., 

c) sea level at St. Petersburg, d) wind velocity sampled at NOAA buoy 42036, 

and e) wind stress. The contour intervals are 0.05 ms-1 for the velocity 

components, north is up and east is to the right for the wind vectors, and all data 

are low pass filtered to exclude oscillations at time scales shorter than 36 hours. 

Figure 3 Modeled (dashed-doted lines) and observed (solid lines) sea 

level variability for the constant density (panel a) and stratified (panel b) cases, 

along with the April 1998 wind vectors. 

Figure 4a April 1998 stratified model simulations of: a) along-shelf and 

b) across-shelf velocity components, and c) isopycnal variations sampled at the 

20 m isobath offshore of Sarasota FL, along with the wind vectors used to drive 

the model (with north up and east to the right). The contour intervals are 0.05 

ms-1 for the velocity components and 0.5 sigma units for the density. 

Figure 4b April 1998 constant density model simulations of: a) alongshelf and b) across-shelf velocity components sampled at the 20 m isobath 

offshore of Sarasota FL, along with the wind vectors used to drive the model 

(with north up and east to the right). The contour intervals are 0.05 ms-1 for the 

velocity components. 
116

Figure 5a Daily snapshots of the stratified model simulated WFS sea 

surface topography spanning the period of downwelling and upwelling between 

days 7 to 14. The contour interval is 0.02 m, and the ranges for sea level are 

given in parentheses. 

Figure 5b Daily snapshots of the constant density model simulated 

WFS sea surface topography spanning the period of downwelling and upwelling 

between days 7 to 14. The contour interval is 0.02 m, and the ranges for sea 

level are given in parentheses. 

Figure 6a Daily snapshots of the along-shelf, across-shelf, and vertical 

velocity components and density sampled along a cross section offshore of 

Sarasota, FL. spanning the period of downwelling and upwelling between days 7 

to 13. The contour intervals are 0.05 ms-1 for the horizontal velocity components, 

10-5 msec-1 for the vertical velocity component, and 0.5 sigma units for density. 

Figure 6b Daily snapshots of the along-shelf, across-shelf, and vertical 

velocity components for the constant density case, sampled along a cross 

section offshore of Sarasota, FL. spanning the period of downwelling and 

upwelling between days 7 to 13. The contour intervals are 0.05 ms-1 for the 

horizontal velocity components and 10-5 msec-1 for the vertical velocity 

component. 

Figure 7 April 1998 measurements of: a) along-shelf and b) acrossshelf velocity components sampled at the 50 m isobath offshore of Tampa Bay. 

The contour intervals are 0.05 ms-1 for the velocity components, north is up and 

east is to the right for the wind vectors, and all data are low pass filtered to 

exclude oscillations at time scales shorter than 36 hours. 

Figure 8 Horizontal velocity vector fields: a) depth-averaged, b) near 

surface, c) mid-water, and d) near bottom, sampled during the peak downwelling 
117

event on day 9. Vectors are shown for every other grid point with the maximum 

value for each panel given in the panel legend. The near bottom vectors are 

magnified by a factor of two relative to the other panels. 

Figure 9 Horizontal velocity vector fields: a) depth-averaged, b) near 

surface, c) mid-water, and d) near bottom, sampled during the peak upwelling 

event on day 12. Vectors are shown for every other grid point with the maximum 

value for each panel given in the panel legend. The near bottom vectors are 

magnified by a factor of two relative to the other panels. 

Figure 10a Satellite AVHRR temperature image for April 10th, the time of 

warmest SST associated with the peak downwelling of day 9. 

Figure 10b Satellite AVHRR temperature image for April 13th, the time of 

coldest SST associated with the peak upwelling of day 12. 

Figure 11 Time series of the vertically integrated: a) along-shelf and b) 

across-shelf momentum balance terms and the c) across-shelf transports, along 

with d) the across-shelf ageostrophic momentum residual, all sampled at the 20 

m isobath offshore of Sarasota FL. from the stratified model simulation 

Figure 12 The across-shelf momentum balance during the peak 

downwelling event on day 9 sampled offshore from Sarasota FL. from the 

stratified model simulation. The panels are: a) Coriolis; b) pressure gradient; c) 

their sum, or the ageostrophic residual; d) friction (primarily vertical); e) the sum 

of the Coriolis, pressure gradient, and friction terms; f) advective acceleration; g) 

the sum of e and f; and h) the local acceleration. Contour intervals are given in 

each panel and shaded regions are negative. 

Figure 13 The across-shelf momentum balance during the peak 

upwelling event on day 12 sampled offshore from Sarasota FL. from the stratified 
118

model simulation. The panels are: a) Coriolis; b) pressure gradient; c) their sum, 

or the ageostrophic residual; d) friction (primarily vertical); e) the sum of the 

Coriolis, pressure gradient, and friction terms; f) advective acceleration; g) the 

sum of e and f; and h) the local acceleration. Contour intervals are given in each 

panel and shaded regions are negative. 

Figure 14 Time series of the vertically integrated vorticity balance 

terms sampled at the 20 m isobath offshore of Sarasota FL. from the stratified 

model simulation. 

Figure 15 Horizontal distribution fields for the vertically integrated 

vorticity balance terms during the peak downwelling event on day 9: a) the 

material rate of change, plus tilting and stretching of relative vorticity; b) the 

stretching of planetary vorticity by changes in the free surface, plus the planetary 

beta terms; c) the bottom pressure torque; and d) the bottom stress torque. 

Contour intervals and ranges are given with each panel and shaded regions are 

negative. 

Figure 16 Horizontal distribution fields for the vertically integrated 

vorticity balance terms during the peak upwelling event on day 12: a) the material 

rate of change, plus tilting and stretching of relative vorticity; b) the stretching of 

planetary vorticity by changes in the free surface, plus the planetary beta terms; 

c) the bottom pressure torque; and d) the bottom stress torque. Contour intervals 

and ranges are given with each panel and shaded regions are negative. 

Figure 17 An analysis of the along-shelf component of vorticity during 

the peak downwelling and upwelling responses on a) day 9 and b) day12, 

respectively. From left to right in each panel are 1) the tilting of planetary vorticity 

filaments by the sheared along-shelf jet, 2) the buoyancy torque, 3) the 

dissipation of relative vorticity by the sheared across-shelf flow, and 4) the local 

rate of change of relative vorticity, plus the remainder of other omitted terms. 

The contour interval is 10-7 s-2. 
119

Figure 1 
120

Figure 2 
121

Figure 3 
122

Figure 4 
123

Figure 5 
124

Figure 6 
125

Figure 7 
126

Figure 8 
127

Figure 9
128

Figure 10 
129

Figure 11 
130

Figure 12 
131

Figure 13 
132

Figure 14 
133

Figure 15 
134

Figure 16 
135

Figure 17 
136

Figure 18 
137

Figure 19 
138

Figure 20
139

Figure 21 
140

5e. West Florida shelf circulation and temperature budget 

 for the 1999 spring transition 
141

West Florida Shelf Circulation and Temperature Budget for the 1999 Spring 

Transition 

Ruoying He 

and 

Robert H. Weisberg* 

College of Marine Science, University of South Florida 

140 7th Ave. South, St. Petersburg, Florida, 33701 

Continental Shelf Research, in press, 2001

*Corresponding author. Tel: 1-727-553-1568; Fax: 1-727-553-1189 

E-mail address: weisberg@marine.usf.edu
142

Abstract

 Mid-latitude continental shelves undergo a spring transition as the net 

surface heat flux changes from cooling to warming. Using in-situ data and a 

numerical circulation model we investigate the circulation and temperature 

budget on the West Florida Continental Shelf (WFS) for the spring transition of 

1999. The model is a regional adaptation of the primitive equation, Princeton 

Ocean Model forced by NCEP reanalysis wind and heat flux fields and by river 

inflows. Based on agreements between the modeled and observed fields we use 

the model to draw inferences on how the surface momentum and heat fluxes 

affect the seasonal and synoptic scale variability. We account for a strong 

southeastward current at mid-shelf by the baroclinic response to combined wind 

and buoyancy forcing, and we show how this local forcing leads to annually 

occurring cold and low salinity tongues. Through term-by-term analyses of the 

temperature budget we describe the WFS temperature evolution in spring. Heat 

flux largely controls the seasonal transition, whereas ocean circulation largely 

controls the synoptic scale variability. These two processes, however, are 

closely linked. Bottom topography and coastline geometry are important in 

generating regions of convergence and divergence. Rivers contribute to the local 

hydrography and are important ecologically. Along with upwelling, river inflows 

facilitate frontal aggregation of nutrients and the spring formation of a high 

concentration chlorophyll plume near the shelf break (the so-called ‘Green River’) 

coinciding with the cold, low salinity tongues. These features originate by local, 

shelf-wide forcing; the Loop Current is not an essential ingredient. 
143

1. Introduction 

Located at the eastern edge of the Gulf of Mexico, the West Florida 

Continental Shelf (WFS) is one of the broadest continental shelves in North 

America. Between its southern and northern ends, bounded by the Florida Keys 

and the Florida Big Bend, respectively, the WFS isobaths vary smoothly, and 

they generally parallel the coastline. This geometry changes along the Florida 

Panhandle in the north where the coastline undergoes a right angle bend, and 

the shelf width decreases to a minimum at the DeSoto Canyon. 

Long-term observations (Weisberg et al., 1996) show that the WFS 

circulation, forced by tides, winds, buoyancy, and possible interactions with the 

Gulf of Mexico Loop Current, varies on time scales from semi-diurnal to interannual. Monthly mean currents at mid-shelf suggest a seasonal cycle with 

along-shore flows either to the southeast in spring, or to the northwest in late 

summer to early autumn. Weisberg et al. (1996) hypothesized that these 

seasonal currents are of a baroclinic nature based on an observed thermal wind 

shear and the seasonal reversal of the across-shelf density gradient. As a 

consequence of the spring transition in surface heat flux from cooling to warming, 

they argued that spatial differences in heating (from the coast to offshore by 

increasing depth and from the south to north by solar declination) forms a midshelf cold tongue and a seasonally maximum across-shelf density gradient that 

supports a southeastward current. Here we examine this locally forced, seasonal 

circulation hypothesis by focusing on the spring transition for 1999, a year when 

the Loop Current, as evidenced in relatively flat isopycnal topography at the shelf 

break, did not have a strong direct influence the WFS. Our objective is to 

describe the circulation and temperature budget for spring 1999 with respect to 

the shelf-wide winds, surface heat fluxes, and river inflows. 

An argument used to describe the transition from wintertime horizontal 

stratification to summertime vertical stratification for mid-latitude shelves is that 

decreased winds and increased solar heating conspire to form a thermocline. 

The details of this process are not well understood, however. Chapman and 

Gawarkiewicz (1993) reason that nonlinearity in the equation of state can 

account for the elimination of horizontal stratification by spatially uniform heating, 

but Morey (1999) point out that this argument is valid only for certain salinity and 

temperature ranges. Other processes must also be important. Morey (1999), 

using a two-dimensional model with a sloping bottom, argue that the surface heat 

flux divided by the water depth is the critical factor in the seasonal transition, 

essentially the differential heating argument advanced earlier. The degree to 

which their argument is valid in the fully three-dimensional sense and the 

regional partition between ocean dynamical and local heating affects are also 

topics of our paper. 

The observational record [e.g. Niiler, 1976; Mitchum and Sturges, 1982; 

Cragg et al, 1983; Marmorino, 1983; Mitchum and Clark, 1986a] shows that the 

WFS circulation and sea level variations are highly correlated with the synoptic 

scale wind stress variations. The passage of cold fronts also affects the local 

temperature balance (e.g., Price, 1976). Along with these local synoptic scale 
144

variations are baroclinic effects that originate with the Loop Current at the shelf 

break [e.g., Paluszkiewicz et al, 1983]. What remains unclear are the relative 

importance between the momentum and buoyancy that are input either locally, or 

at the shelf break. 

Such questions are of multi-disciplinary interest since, despite its 

oligotrophic description, the WFS supports highly productive ecosystems. These 

include episodic toxic dinoflagellate blooms (red tides) near the coast (Steidinger, 

1981; Vargo et al, 1985), a seasonal chlorophyll plume near the shelf break 

(Gilbes et al, 1996), and important commercial and recreational fisheries 

throughout the WFS. Parallel programs of in situ measurements and numerical 

model experiments are presently in place for an improved understanding of the 

circulation and how it affects seasonally varying water properties and influences 

organism growth and distribution. 

This paper focuses on local wind and buoyancy forcing during the spring 

transition of 1999, independent of the Loop Current. We use the primitive 

equation, Princeton Ocean Model (POM) described by Blumberg and Mellor 

(1987) forced by National Center for Environmental Predication (NCEP) 

reanalysis winds and net surface heat flux and by river inflows. The only role of 

the adjacent Gulf of Mexico is to set the vertical distribution of temperature and 

salinity for initializing the model density field. Once begun, the integration 

proceeds solely on the basis of local forcing. By running twin experiments, one 

with heat flux and the other without, we explore the relative importance of wind 

and buoyancy in effecting the seasonal and synoptic scale variability. 

Section 2 describes the model and forcing fields. Section 3 compares 

model results with in-situ observations. Based upon these comparisons the 

model is used in section 4 to describe the seasonal mean circulation on the WFS 

for spring 1999 and the evolution of the corresponding temperature and salinity 

fields. Section 5 presents a term-by-term analysis of the three-dimensional 

temperature budget. The results are summarized and discussed in section 6. 
145

2. Model and Forcing fields 

2.1. Model 

 We use the POM for the following reasons. First, it has an embedded 

turbulence closure submodel (Mellor and Yamada, 1974, 1982; Galperin et al., 

1988) for parameterizing vertical turbulence mixing. Second, it employs a sigma 

coordinate in the vertical which, with the turbulence closure sub-model, is well 

suited to study the nonlinear dynamics over a shallow, gently sloping continental 

shelf. Third, its orthogonal curvilinear coordinates in the horizontal are 

convenient for resolving the near-shore regions. 

Previous WFS POM applications include Yang et al. (1999a,b), Li and 

Weisberg (1999a,b), Weisberg et al. (2000), and Weisberg et al. (2001). Yang et 

al. (1999a) studied the WFS response to climatological monthly mean wind 

forcing. Qualitative agreements were found between the model and observations 

to some extent, but monthly mean wind stress alone could not account for the 

southeastward current observed at mid-shelf in spring (Weisberg et al., 1996). Li 

and Weisberg (1999a,b) focused on synoptic scale winds, respectively describing 

the kinematics and dynamics of WFS responses to idealized upwelling favorable 

wind forcing under constant density. The inner-shelf length scale was found to 

be a frictional one, consistent with the analytical work of Mitchum and Clarke

(1986b). The same model was also applied to a specific upwelling case study 

with both constant density and stratified conditions (Weisberg et al. 2000). A 

comparison of in-situ data with model results confirmed a simple Ekmangeostrophic route to spin-up and identified regional upwelling centers promoted 

by coastline and isobath geometries. The utility of this model in replicating the 

longer-term synoptic scale variability and the sensitivity of the WFS response to 

its background density state was then shown by Weisberg et al. (2001). An 

asymmetry in the inner-shelf responses to upwelling and downwelling favorable 

winds was found that helps to clarify the scale of the inner-shelf response. 

Consistent with Ekman dynamics, the inner shelf is the region with divergent 

bottom boundary layer. Since thermal wind effects can either enhance or 

decrease the bottom boundary layer development, the inner-shelf under stratified 

conditions can respond asymmetrically to upwelling and downwelling favorable 

winds. 

Many WFS circulation questions remain. For instance, the relative effects 

of buoyancy and wind forcing have not been considered. While previous drifter 

and current meter measurements hint at seasonally varying circulation patterns 

(Tolbert and Salsman,1964; Gaul, 1967; Williams et al., 1967; Weisberg et al., 

1996; Yang et al., 1999b), there are no definitive measures of these. The 

communication of water between the deep ocean and the shelf and the 

communication of water between different regions of the shelf remain poorly 

understood. To address these questions it is necessary to explore a broader 

region and consider more complete forcing functions than in previous studies. 

 The model domain (Fig 1) extends from the Florida Keys in the southeast 

to west of the Mississippi River in the northwest, and it has one open boundary 

arcing between these two locations for which a radiation boundary condition 

(Orlanski, 1976) is used. The model domain includes the major rivers that impact 
146

the WFS and the Desoto Canyon region where the shelf is narrowest, and its 

orthogonal curvilinear grid has horizontal resolution that varies from less than 

2km near the coast to 6 km near the open boundary. Vertically, the sigma 

coordinate has 21 layers with higher resolution near the surface and bottom to 

better resolve the frictional boundary dynamics. In total, the model has 

121×81×21 grid points. Horizontal diffusivities are parameterized using the 

Smagorinsky (1963) formulation with a coefficient of 0.2. Bottom stress, τb, is 

calculated by a quadratic law with variable drag coefficient having a minimum 

value of 0.0025. A mode splitting technique is used for computational efficiency 

(Blumberg and Mellor, 1987). Here we use external and internal time steps of 6 

seconds and 360 seconds, rspectively. 

The model is initialized at rest with horizontally uniform stratification. 

Stratification above 200m is based on temperature and salinity observations 

taken during a March 1999 trans-shelf hydrographic survey [from the Ecology of 

Harmful Algal Blooms (ECOHAB) Program]. Stratification below 200m is based 

on climatology. From this initial zero-baroclinicity state, the model spins up 

rapidly, generating baroclinicity in balance with the wind and buoyancy forcing. 

An alternative is to begin with a baroclinic field and allow the model currents to 

come into balance diagnostically with this field before proceeding with the spring 

simulation. The hydrographic data are not sufficient for this, however, and 

spurious currents due to incorrect density would corrupt the experiment. 

Consistent with our objective of determining the WFS responses to local, shelfwide forcing only, our initial baroclinicity-free state is a sensible choice. 

Tidal forcing is excluded in the present application since we are not 

considering high frequency variability. It is recognized that tidal mixing can affect 

the synoptic and seasonal scales when the tidal currents are large, but here the 

tidal currents are only a few cm s-1. Modeled and observed tidal current analyses 

will be reported on separately. 

2.2 Atmospheric forcing 

Different from previous WFS model studies that considered wind forcing 

only, here we include both wind and thermohaline forcing. The wind and heat 

flux fields are from the NCEP daily reanalysis product for the period February 28, 

1999 to June 1, 1999. These values, with a grid resolution of 2.5o

×2.5o, are 

interpolated onto the model grid. The NCEP winds agree well with in-situ buoy 

winds for the spring 1999 season. Unlike the winds, however, coarse resolution 

renders the NCEP heat flux unrealistic because of smaller scale WFS 

temperature structures. We correct for this using a relaxation method (e.g., Ezer 

et al, 1992; Chu et al, 1999). Thus, the surface heat flux forcing is given by 

0

( )

= ∂

∂

+ − 













 = ∂

∂

z

S K

C

C

Q

z

K

H

obs

P

H

H θ θ

ρ

θ

 (1) 

where QH is the net heat flux, θobs is an interpolation of the monthly-mean 

satellite observed sea surface temperature, and Cp is the specific heat. The 
147

salinity flux in this study is set to be zero. The relaxation coefficient, C, or the 

reciprocal of the restoring time per unit area, is set at 1m/day. Such relaxation 

prevents deviations from observed monthly-mean SST in an attempt to force 

realistic baroclinic flow structures. These structures are facilitated by turbulence 

mixing through the coefficients KM and KH computed with the Mellor and Yamada

(1982) 2.5 level turbulence closure sub-model. 

2.3 Lateral boundary forcing. 

Gulf of Mexico Loop Current forcing is excluded in this study for two 

reasons. First, previous observations and model studies concluded that 

persistent forcing of the middle and inner-shelf by the Loop Current is minimal 

(Marmorino, 1982, 1983 a,b). Second, modeling the effects on the WFS of an 

aperiodically varying Loop Current and its associated eddies (e.g., Sturges and 

Leben, 2000) remains a great challenge (Marmorino, 1982; Cooper, 1987), 

presupposing that the Loop Current itself is being described properly. To better 

assess the role of the Loop Current as a WFS boundary condition it will be 

necessary to nest a regional model with a larger Gulf of 

Mexico/Caribbean/Atlantic Ocean model. This is beyond the scope of the 

present paper that focuses on local, shelf-wide forcing only. We find, however, 

that local forcing is capable of driving much of the observed synoptic and 

seasonal scale variability. 

Seven major rivers are introduced into the model domain for land derived 

buoyancy forcing. These are the Mississippi, Mobile, Apalachicola, Suwannee,

Hillsborough, Peace and Shark rivers. We use the technique of Kourafalou et al

(1996)(also see Pullen 2000), whereby interpolated monthly mean mass flux data 

for these rivers are input to the top sigma level at the grid cells closest to the 

rivers’ locations. 

We define the spring season here as March 1 to May 31, and we focus on 

this period for 1999. As an initial value problem we begin from a state of rest on 

February 28. With no initial baroclinicity the spin-up phase proceeds rapidly over 

the course of a few pendulum days, consistent with the barotropic response 

arguments for a gently sloping shelf of Clarke and Brink (1985). Under the 

conditions of surface cooling that occurs prior to the spring warming transition in 

mid-March, convective mixing very efficiently adjusts the initial density field on 

the shallow shelf. In other words, the “memory” of initial density field for this 

spring transition experiment is short, and sensitivity experiments that we 

performed using longer spin-up times showed very little difference from the 

present model results. 
148

3. Model and Data Comparisons 

3.1 Sea level 

Since the model is forced without tides, all of the model and data 

comparisons are shown after low-pass filtering to exclude tidal and inertial period 

oscillations. Sea surface height comparisons are given in Fig. 2 at four different 

tide-gauge stations from Pensacola in the northwest to Naples in the southeast. 

Agreement is good at all of these with squared correlation coefficients exceeding 

0.80. We conclude that coastal sea level for this three-month period responds 

primarily to local, shelf-wide forcing. 

3.2 Currents 

 Comparisons are made between the modeled and observed velocity 

vector time series at the 50m, 30m, 25m, 20m, and 10m isobaths (moorings 

CM2, EC3, NA2, EC4, EC5, and EC6 in Fig. 1). As examples, we show the 

modeled and observed vector time series at the 50m, 25m, and 10m locations in 

Figs. 3-5, respectively. The observations are from acoustic Doppler current 

profilers (ADCP), and for each location we show comparisons at three different 

depths: near-surface, mid-water column, and near-bottom. These comparisons 

are quantified by a complex correlation analysis (e.g., Kundu 1976). Defining the 

modeled and observed velocity vectors in the Argand plane as 1 1 1 w = u + iv and 

2 2 2 w = u + iv , respectively, the complex squared correlation coefficient is 

(2)

[ ( ) ( ) ( ) ( )]

[ ( ) ( )] ( , )

2

*

1 2

*

1

2

2

*

1

1 2

w t w t w t w t

w t w t

ρ w w =

where the overbar denotes a time average. The complex correlation has 

an amplitude and a phase, the amplitude being the correlation coefficient and the 

phase being the angle (measured counterclockwise) between the modeled and 

the observed currents. Like sea level, the modeled and observed currents also 

compare well as measured by the two sets of numbers provided with each plot. 

The left-hand sets are the seasonal mean east and north velocity components. 

The right-hand sets are the squared correlation coefficient, phase angle, and 

regression coefficient. At all stations and depths the squared correlation 

coefficients range between 0.62 and 0.82, and the orientations agree to within –

10 to +20 degrees. More importantly, as seen directly from the time series, the 

model gets the sense of the velocity rotation correct in both the surface and 

bottom Ekman layers. A deficiency in the modeled currents, however, lies in the 

amplitudes. The regression coefficients show that the model underestimates the 

observed velocity fluctuations by between 20 to 50 percent. 

Notwithstanding the amplitude disparity, the model reproduces the 

patterns of current variability reasonably well. The systematic underestimate of 

the currents may be the result of the low resolution NCEP forcing fields, and 

hence too much smoothing when interpolating these fields onto the model grid. 

Model performance also degrades between the shallowest and deepest 

comparison sites, i.e., the 10m isobath currents agree better than the 50m 

isobath currents. This is expected based on the frictional scale of the inner-shelf 
149

response to wind forcing (e.g., Weisberg et al, 2001). The 50m isobath is at the 

outer half of the inner-shelf so we anticipate a decreased correlation there. 

The modeled and observed velocity comparisons are summarized in Fig. 

6 where we show the mid-depth seasonal mean vectors and hodograph ellipses 

at all of the mooring locations. The mean vectors compare reasonably well, and 

while the ellipse semi-major axes are off by between 20-50%, the orientations 

and eccentricities tend to agree. On the basis of these agreements we now use 

the model to discuss the WFS circulation in spring 1999. 
150

4. Mean Circulation 

4.1 Flow fields 

The seasonal mean circulation, obtained by averaging the model flow 

fields from March 1 to May 31, is presented in Fig 7. Three-dimensional flow 

features arise because of the WFS geometry that includes partial blocking by the 

Florida Keys, the coastline changes of the Big Bend, and the intrusion of Desoto 

Canyon. These features are depicted in horizontal velocity field maps given for 

sigma layers 2 (near-surface), 10 (mid-level), 20 (near-bottom), and for the 

depth-average. 

The mid-level and depth-averaged fields are similar (and with the same 

scale), and they show the general nature of the 1999 spring season currents. A 

jet exists with axis situated between the mid-shelf and the shelf break. This jet 

originates along the northern coast east of the Mississippi River, and it flows 

along the relatively narrow Florida Panhandle shelf as a closely confined coastal 

feature. The coastal jet bifurcates at Cape San Blas into a mid-shelf part that 

heads along-isobath toward the southeast and a coastal part that hugs the Big 

Bend coastline. The mid-shelf part is consistent with the spring season 

southeastward current described by Weisberg et al. (1996). This mid-shelf 

current again bifurcates upon approaching the Florida Keys with a portion turning 

toward Florida Bay and another portion turning farther offshore. Why this occurs 

and how it affects the Florida Bay/Florida Keys region are evident in the nearbottom and near-surface flow fields (note the scale changes). Near the bottom 

we see a convergence of vectors on Florida Bay and the Florida Keys. In 

contrast to this, near the surface we see flow paralleling the Florida Keys. The 

near bottom flow upwells, feeding the near surface flow, and the second 

bifurcation is due to the recirculation associated with the upwelling. Li and 

Weisberg (1999a) discussed a similar recirculation due to the Florida Keys for 

the case of upwelling winds under a constant density setting. From the surface 

current map we can appreciate why surface drifters originating in the north do not 

penetrate the southeast portion of the WFS; the so-called 'Forbidden Zone' 

described by Yang et al. (1999b). We also note that the currents within the Big 

Bend and those that flow southward near-shore between Cedar Key and 

Sarasota are much weaker than the currents at mid-shelf. As will be shown in 

the next section, this is a consequence of a surface heat flux-induced cyclonic 

circulation that adds destructively (constructively) to the wind driven flow nearshore (offshore). 

Three-dimensionality in the flow field is further evident in the vertical 

component of velocity shown for near-bottom, middle, and near-surface sigma 

levels in Fig. 8. Upwelling is prominent north of the Florida Keys with a maximum 

near the Dry Tortugas, as is often evident in satellite SST imagery (e.g., 

Weisberg et al., 2000). Upwelling is also prominent west of DeSoto Canyon, 

around Cape San Blas, and along the west Florida coastline. Other regions 

show downwelling. This vertical circulation is a function of both the bottom 

boundary layer and the interior. The near bottom w can be reconciled based on 

the near bottom horizontal velocity field (Fig. 7) and the bottom kinematic 

boundary condition. Mid-water column divergence then accounts for the 
151

transition of w from its near-bottom to near-surface values. Upwelling occurs 

along the entire near-shore domain near the bottom, whereas upwelling is more 

localized near the surface. 

4.2 Temperature and salinity field 

The modeled surface temperature and salinity fields sampled at the end of 

the model run on May 31 are presented in Fig. 9. Since the initial model 

temperature and salinity fields are horizontally uniform, this figure shows the 

combined effects of the momentum and buoyancy fluxes in changing the surface 

temperature and salinity. The two principal features are the mid-shelf cold 

tongue that extends southeastward from Cape San Blas and the low salinity 

tongue that also extends southeastward, but displaced seaward of the cold 

tongue. Both of these features occur annually on the WFS. The cold tongue is 

imposed to some extent through the surface heat flux relaxation, whereas the low 

salinity tongue is a fully prognostic result of the model. 

The low salinity tongue derives as a river plume, accumulating fresh water 

primarily from the Mississippi River with additions from the Mobile and 

Apalachicola Rivers. Modulated interannually, the low salinity tongue extends 

southward each year, and in some years (1993, for instance – see Dowgiallo et 

al., 1995) it can be traced around the Florida peninsular to the Carolinas. Using 

Coastal Zone Color Scanner data between 1979 and 1986, Gilbes et al. (1995) 

reported a spring chlorophyll plume at mid-shelf (termed the ‘Green River’) that 

also extends southeast from Cape San Blas. Their explanations for the plume 

included: i. nutrient fluxes from the Apalachicola River; ii. nutrient fluxes from the 

Mississippi and Mobile Rivers; iii. seasonal changes in steric height between the 

shelf and deeper Gulf of Mexico; and iv. circulation of water from the Loop 

Current. Our results help to clarify these speculations. Consistent with Weisberg 

et al. (1996) the ‘Green River' is associated with both the cold tongue and the low 

salinity tongue. The low salinity tongue, at least in 1999, appears to originate at 

the Mississippi River. It is advected eastward, on average, by the spring coastal 

jet and then southeastward where the jet bifurcates at Cape San Blas. This 

bifurcation, in part, is related baroclinically to the cold tongue; hence the offset 

between the low salinity and cold tongues. Since the Loop Current is not 

included in our model, and since it did not extend to the northeast Gulf of Mexico 

in spring 1999, we can rule out the Loop Current as a primary conveyance of 

these salinity and temperature features. What remains unclear is the relative 

importance of land-derived versus upwelling-derived nutrients in fueling the 

chlorophyll plume. Aggregation of nutrients and phytoplankton by the frontal 

regions of the cold tongue/low salinity tongue complex can nevertheless account 

for the 'Green River'. 

To explore the roles of momentum and heat flux in the formation of these 

surface temperature and salinity features we ran a model twin experiment forced 

by NCEP winds only. Fig 10 shows the seasonal mean circulation with wind 

forcing only for comparison with Fig.7 with both wind and heat flux forcing. 

Differences occur in all fields. The most obvious is the location of the jet, as 

seen in either the depth-averaged or the middle level fields. With wind forcing 
152

only the jet has a stronger coastal expression over the entire WFS, as contrasted 

with a stronger mid-shelf jet under wind and heat flux forcing. Baroclinicity 

explains the difference. With wind and heat flux forcing a dynamical feedback 

occurs with the cold tongue causing a cyclonic baroclinic circulation that adds 

constructively (destructively) with the wind forced circulation at mid-shelf (nearshore). A baroclinic circulation also arises without heat flux, but this adds 

constructively with the wind forced circulation near-shore. This latter effect for 

wind forcing only is seen in the surface temperature field of Fig. 11. Surface 

temperature is colder everywhere since the ocean circulation merely rearranges 

it. By virtue of mean upwelling the near-shore region is colder than the offshore 

region which causes a southward baroclinic circulation that adds constructively 

with the southward wind-driven circulation. 

Given the importance of surface heat flux on the seasonal circulation it is 

instructive to see how the baroclinic contribution evolves. This is shown in Fig. 

12 for the March, April, and May monthly mean, depth-averaged flows relative to 

the seasonal mean. March shows an anticyclonic circulation with southeastward 

flow near-shore. During April, once the effects of positive heat flux set in, we see 

the formation of the mid-shelf jet and a cessation or reversal of the near-shore 

flow. This further develops into a strong mid-shelf/shelf-break current in May as 

the cyclonic baroclinic flow becomes fully developed. This is consistent with the 

Weisberg et al (1996) hypothesis on the origin of the springtime southeastward 

current. 

Temperature cross-sections at various positions along the WFS provide 

further information on the cold tongue evolution. These are shown for wind and 

heat flux forcing and for wind forcing only on March 15, April 15 and May 15, in 

Figs. 13, 14, and 15, respectively, at four transects offshore of DeSoto Canyon, 

Cape San Blas, the Big Bend, and Sarasota. Since the heat flux is initially out of 

the ocean the March 15 transects, with or without heat flux, are similar at depth. 

They differ on the shelf where surface cooling, coupled with efficient convective 

mixing, produces typical wintertime horizontal stratification. By April 15, with a 

reversal in the sign of the heat flux, the two cases depart everywhere. The Big 

Bend transect shows a dome of cold water at mid-shelf in the heat flux case. 

This originates primarily by the burial (by water being heated from above) of the 

cold water previously formed in the Big Bend. Water is warmer in the shallows 

near-shore for the same reason that they were colder in the previous season (a 

similar heat flux over a shallower layer will result in a larger internal energy 

change per unit area). There are also secondary contributions from upwelling at 

the shelf break. Here the shelf break upwelling is from the Big Bend as opposed 

to the DeSoto Canyon, and this is in part a consequence of bottom topography. 

The shelf break in the Big Bend occurs about 20m deeper than in the DeSoto 

Canyon. Hence, less upwelling is required for water to rise onto the shelf in the 

Big Bend than in the DeSoto Canyon. This finding is clear in continuous 

animations where we also see that the advective time scale between the DeSoto 

Canyon and the Big Bend is longer than the synoptic weather time scale. For 

spring 1999, at least, we do not see a direct connection between the DeSoto 

Canyon and Big Bend regions as would be evidenced by continuous isotherms. 
153

Note that while 210C waters occur on the shelf offshore of both the DeSoto 

Canyon and Big Bend regions these waters are not seen on the shelf at the 

intervening Cape San Blas cross section. Local upwelling in the Big Bend region 

(due to a relatively deep shelf break) is therefore important, and it may help to 

explain the relatively high productivity of the Florida Middle Grounds. A similar 

subsurface cold-water core is seen farther south off Sarasota and for the same 

reasons. It is smaller in magnitude because of differential cooling and heating 

from north to south, i.e., when the heat flux reversed from cooling to warming the 

temperatures off Sarasota were already warmer than those in the Big Bend. 

Dynamically, the doming of cold water at mid shelf (and hence the cold tongue) 

induces the cyclonic baroclinic circulation. This is further facilitated by the no 

heat flux bottom boundary condition that bends the isotherms perpendicular to 

the bottom. These combined effects of differential heating (near-shore to 

offshore by bottom slope and southward to northward by solar declination) and 

dynamical feedbacks (baroclinic pressure gradients associated with the 

temperature gradients) conspire to cause the spring season temperature 

transition from well mixed to stratified with a cold tongue at mid shelf. 

The May 15 transects show the continued evolution of these processes. 

By now the mid-shelf isotherm doming has relaxed due to continued heating and 

so has the mid-shelf circulation due to decreased winds and baroclinicity. The 

location of the jet is now farther offshore at the shelf break, and the modeled 

temperature agrees with observed ECOHAB Program hydrography. 

4.3 Lagrangian drifters 

Model-simulated Lagrangian drifter tracks provide additional information 

on the evolution of the cold and low salinity tongues and the regions for possible 

communication between the deep ocean and the shelf. Dynamically passive 

particles released in the model on April 1stare tracked through May 31st. These 

are shown in Figs. 16 and 17 for particles originating either near the surface or 

the bottom, respectively. Most evident for the particles released near the surface 

is advection by the mid-shelf jet and the bifurcations at Cape San Blas and north 

of the Florida Keys. The particle trajectories agree with the Eulerian maps 

previously shown and with the cold and low salinity tongue features. With the 

exception of those in the mid-shelf jet the particle displacements are not very 

large. The near-surface drifters also show a propensity to flow either alongisobath or offshore. The near-bottom released drifters show subtle, but important 

differences, particularly in the Florida Big Bend/Middle Ground region. There we 

see onshore flow between the shelf break and the mid-shelf. This is consistent 

with the changes in hydrography discussed in the previous section, supporting 

our speculation that the Big Bend may be an important region for deep 

ocean/shelf interactions. 
154

5. Temperature Budget

5.1 The temperature equation 

To analyze the temperature budget, the temperature equation is recast 

from its modeled flux divergence, sigma level form to an advective, z-level form. 

Thus, we diagnose 

















∂

∂

∂

∂  +



 





∂

∂

∂

∂  +



 





∂

∂

∂

∂

+

∂

∂ − ∂∂ − ∂∂ = − ∂∂

y

T A

x y

T A

z x

T K

Z z

T

w

y

T

v

x

T

u

t

T

H H H

 (3) 

 a b c d vf hfx hfy 

which equates the local rate of change of temperature (a) to a combination 

of the flow field advective rate of change (b+c+d), and the rates of change by 

vertical diffusion (vf), and horizontal diffusion (hfx+hfy). The temperature balance 

is explored using: i. time series of vertical averages, ii. time series of vertical 

profiles at four different locations, and iii. term by term spatial maps of the sea 

surface temperature rates of change. 

The four analysis locations (see Fig 1) are chosen with respect to the cold 

tongue and the upwelling region north of the Florida Keys. Point A is on the 50m 

isobath at the seaward side of the Big Bend shelf subsurface cold dome. Point B 

is on the 40m isobath west of Tampa Bay where the cold tongue begins to taper 

off. Point C is on the 15m isobath offshore of Sarasota on the inshore side of the 

cold tongue. Point D is north of the Florida Keys. A term-by-term analysis of (3) 

quantifies the contributions by each physical process in changing the 

temperature. 

5.2 Depth-averaged balances. 

A depth-averaged temperature equation is obtained by vertically 

integrating equation (3). Since horizontal temperature diffusion is generally at 

least an order of magnitude less than the other terms, the depth-averaged 

diffusion term is essentially the depth-averaged vertical diffusion, Q /( C H) ρ P , 

where Q is the net surface heat flux, ρ and CP are the seawater density and 

specific heat, and H is the water depth. The temperature variations depend on 

both ocean advection and diffusion. With a two-dimensional model, Morey

(1999) proposed that temperature, in a depth-averaged sense, could be well 

represented without advection, i.e., 

 c H

Q t

t

T

dt

dT

ρ P

( ) = ∂∂ = 

 (4) 

The validity of this assertion for various locations on the WFS can be 

evaluated by a fully three-dimensional analysis. 

Time series of the depth-averaged ocean advection and diffusion terms 

and their sum for locations A-D are shown in Figs. 18a-18d, respectively. While 

not shown in these figures, we note that the sum of these two terms is nearly 

exactly equal to the negative of the local rate of temperature change, which is 

essentially a check on our budget analysis. At station A, the depth-averaged 

rates of change are relatively small, and a clear distinction exists between the 
155

balances occurring on seasonal and synoptic time scales. The seasonal change 

(as given by the three-month mean values) is primarily by surface heat flux, 

whereas the synoptic variability (as given by the standard deviations) is primarily 

by ocean advection. Although a relatively small contribution to the seasonal 

change, the ocean advection does provide a cooling influence that offsets the 

surface heating by about 4% at this location. Thus, of the total change in 

vertically averaged temperature (1.660

C), the contributions by ocean circulation 

and surface heat flux are a cooling of -0.070C and a warming of 1.730C, 

respectively. In contrast with the seasonal change, for synoptic scale variability 

we see that the ocean circulation is twice as effective as the surface heat flux in 

changing the vertically averaged temperature at station A. 

Somewhat different results are found at station B. While surface heat flux 

and ocean circulation again control the seasonal and synoptic scale variability, 

respectively, the relative magnitudes and the signs change. Of the total change 

in vertically averaged temperature (3.050

C), both the ocean circulation and the 

surface heat flux at the southern terminus of the cold tongue contribute to 

warming at 0.680C and 2.370C, respectively. The ocean circulation also plays a 

proportionately larger role in the synoptic scale variability at station B than at 

station A. 

The relative importance of the ocean circulation versus surface heat flux 

continues to change approaching shallower water. At the 15m isobath, station C, 

of the total change in vertically averaged temperature (4.760C), the contributions 

by ocean circulation and surface heat flux are cooling of -1.240

C and a warming 

of 6.000C, respectively. The shallower the water, the larger the total spring 

transition change. However, the role of the ocean circulation in this change 

varies in both magnitude and sign with location. For instance, at station D in the 

upwelling region north of the Florida Keys, the ocean circulation provides a 

cooling influence of 2.790C that partially offsets the surface heat flux warming 

influence of 8.110C. Without the upwelling influence by ocean circulation the 

water temperature north of the Florida Keys would be much warmer. Along with 

these increased effects by the ocean circulation on the seasonal mean 

temperature change, the magnitude of the controlling ocean circulation effect on 

the synoptic scale variability also increases with decreasing water depth. The 

temperature budget on all time scales is a complex one requiring a fully threedimensional description. 

5.3 Vertical profiles of the term by term balances 

The temperature budget three-dimensionality is further explored through 

time series of the depth profiles of the individual terms that comprise the 

temperature balance at these four stations (Figs.19-22). In each of these figures 

the left hand panels show the horizontal and vertical components of the ocean 

advection and their sum, and the right hand panels show the diffusion, the 

diffusion plus the advection (which is nearly exactly equal to the negative of the 

local rate of change of temperature), and the temperature. With the exception of 

the initial portion of the record when the surface heat flux is out of the ocean and 

convective mixing is evident, the immediate impact of the surface heat flux is 
156

primarily in the near surface region (Fig. 19). Without convective mixing, it is 

then only through turbulence mixing, brought about by the ocean circulation 

dynamics, that the surface heat flux effect accumulates downward to make its 

contribution to the depth-averaged spring transition. From the vertical 

distributions of the seasonal means shown to the right of each panel we see that 

turbulence mixing distributes the surface heat flux over about 20m depth. Along 

with its role in turbulence mixing, the direct role of the ocean dynamics through 

advection on the synoptic scale variability is seen in both horizontal and vertical 

directions. Omission of any coordinate direction would compromise the model’s 

ability to describe the temperature evolution. For instance, the strong 

stratification at the end of May at station A is due largely to the combined effects 

of horizontal and vertical advection. Similar conclusions follow from station B 

(Fig. 20). Station C (Fig. 21), in shallow water, shows the additional effect of 

turbulence mixing in the bottom Ekman layer. Here, temperature is elevated 

near the bottom by ocean circulation-induced mixing. The role of the ocean 

circulation in promoting mixing is seen in the covariability between the advection 

and diffusion terms. These interactions are even more evident at station D (Fig. 

22) where warming by near-bottom turbulence mixing is required to partially 

offset the cooling influences by horizontal advection and upwelling. Another 

interesting feature of the shallow water regions is the tendency for large changes 

in stratification by the ocean circulation. Two downwelling events culminating in 

destratification and warming by advection are seen at station C around April 5th

and May 10th. Similarly, but in an opposite sense, stratification caused by 

horizontal advection is seen at this station at the end of May. 

5.4 Term by term contributions to the seasonal change in SST 

The contributions to the seasonal mean rate of change of sea surface 

temperature (SST) by each of the advection and diffusion terms are discussed 

with respect to Fig. 23. The left panels show the advection terms and their sum, 

and the right panels show the horizontal and vertical diffusion terms, and the sum 

of all terms. Of particular interest here is the spatial distribution of the processes 

controlling SST over the model domain. Horizontal diffusion is minimal 

everywhere except for north of the Florida Keys where it is still the smallest of the 

terms. Vertical diffusion controlled by the surface heat flux, warms everywhere 

with maximum warming tendencies stretching along mid-shelf from the Big Bend 

to the Florida Keys. This warming is the seasonal signal associated with spring, 

which displays the spatial inhomogeneities mentioned earlier. The difference 

between this map and the local rate of change of SST is the ocean circulation. 

By advecting relatively cool water southward, advection provides a cooling 

influence over the WFS with largest values north of the Florida Keys. With 

regard to total advection, Cape San Blas divides the entire domain into two 

sections. To the west, and extending along the north Florida coastline to the 

Mississippi River, advection tends to provide a warming influence, whereas to the 

east, along the west Florida coastline, advection tends to provide a cooling 

influence. Without advection, the center of the cold tongue on the WFS would be 

some 3-50

C warmer. 
157

6. Summary 

Mid-latitude continental shelves undergo a spring transition as the net 

surface heat flux changes from cooling to warming. Using in-situ data and a 

numerical circulation model we investigate the circulation and temperature 

budget on the WFS, including the northeast Gulf of Mexico shelf from the 

Mississippi River to the Florida Keys, for the spring transition of 1999. The data 

consist of sea level from coastal stations, velocity profiles from instruments 

moored across the shelf between the 50m and 10m isobaths, and hydrography 

from ship surveys. The model is a regional adaptation of the primitive equation, 

POM forced by NCEP reanalysis wind stress and heat flux fields and by river 

inflows. Based on agreements between the modeled and observed fields we use 

the model to draw inferences on how the surface momentum and heat fluxes 

affect the seasonal and synoptic scale variability. 

Spring season features of the WFS include a mid-shelf southeastward 

current, cold and low salinity tongues, and a high chlorophyll plume. We account 

for the southeastward current in 1999 by the combined responses to local, shelfwide wind and buoyancy forcing. Wind stress drives a circulation that tends to be 

strongest near-shore. Heat flux provides a cyclonic contribution that adds 

constructively (destructively) at mid-shelf (near-shore), thus forming the observed 

mid-shelf jet. This heat flux-induced baroclinic circulation is related to the spring 

season cold tongue. By advecting Mississippi (and other) River water it forms 

the low salinity tongue that is displaced seaward of the cold tongue. 

Convergence of nutrients and associated phytoplankton growth then accounts for 

the high chlorophyll concentrations ('Green River') that are co-located with these 

surface features. These findings support the hypothesis advanced by Weisberg 

et al. (1996) on the origin of the southeastward current and cold tongue through 

differential heating from the coast to offshore (by shoaling topography) and from 

south to north (by solar declination). Since we arrive at these features with a 

model experiment that explicitly omits the Gulf of Mexico Loop Current we argue 

that the Loop Current is not an essential element of these spring transition 

features. 

Through term-by-term analyses of the temperature budget we describe 

the evolution of WFS temperature in spring. Surface heat flux largely controls 

the seasonal transition, whereas ocean circulation largely controls the synoptic 

scale variability. These two processes are closely linked, however. Since the 

ocean circulation controls the turbulence mixing, the effects of the ocean 

circulation are of increasing importance with decreasing water depth. For 

instance, warming by turbulence mixing near the bottom only occurs when the 

bottom Ekman layer is well developed by strong currents. Thus, the water 

column can warm even under upwelling influence if the mixing is large enough. 

The water column is also found to either stratify or destratify in response to 

ocean circulation changes. Examples of these effects are shown. While our 

temperature analyses support the Morey (1999) one dimensional temperature 

balance argument advanced on the basis of a two dimensional model we 

demonstrate that the temperature balance is more complicated in time and space 

and requires fully three-dimensional thermodynamics. 
158

Bottom topography and coastline geometry are important in generating 

regions of convergence and divergence and hence upwelling centers. In 

particular, we show that the region north of the Florida Keys has strong upwelling 

in spring and we speculate on the importance of the Florida Big Bend as a region 

for communication between the deeper Gulf of Mexico and the WFS. The shelf 

break there is about 20m deeper than at the DeSoto Canyon thereby requiring 

less upwelling for deeper waters to broach the shelf. This may be one reason 

why the Florida Middle Grounds are productive. The northeast Gulf of Mexico 

shelf from Cape San Blas to the Mississippi River also shows mean upwelling 

during spring 1999. 

Acknowledgements 

This work was supported by grants from the Office of Naval Research, 

grant # N00014-98-1-0158 and the National Oceanic and Atmospheric 

Administration, grant #NA76RG0463. We also benefited from an evolving 

program with previously support from the United States Geological Survey and 

the Minerals Management Service. Ocean Circulation Group staff (Messrs. R. 

Cole, J. Donovan, C. Merz, and P. Smith) contributed to the field work and 

analyses, as did students W. Hemme and J. Virmani. We thank F. Muller-Karger 

for providing assistance with the monthly mean SST fields. NCEP reanalysis data 

provided by NOAA-CIRES Climate Diagnostics Center, Boulder, CO was 

obtained from their web site at http://www.cdc.noaa.gov/ 
159

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List of Figures 

Figure 1 The regional model grid (upper panel) and bathymetry and 

station locations (lower panel). Sea level comparisons are with Florida tide 

gauges at Pensacola, Apalachicola, St. Petersburg, and Naples. Velocity 

comparisons are with acoustic Doppler current profiles from instruments moored 

at the 50m, 30m, 25m, 20m, and 10m isobaths (1-6). Temperature is described 

along transects I-IV, and the temperature budget is diagnosed at Stations A, B, 

C, and D. 

Figure 2 Comparisons between modeled (bold) and observed (thin) 

sea level at Pensacola, Apalachicola, St. Petersburg, and Naples as quantified 

by a squared correlation coefficient, along with the NCEP wind velocity sampled 

at station A. 

Figure 3 Comparisons between modeled and observed currents at 

the 50m isobath (mooring CM2) sampled at depths of 3m, 25m and 40m, along 

with the NCEP wind velocity sampled at station A. Each vector current time 

series is accompanied by its seasonal mean east and north velocity components 

(left hand couplet), and each model/data comparison is quantified by its squared 

complex correlation coefficient, phase angle (or angular deviation of the model 

vector from the data vector measured counterclockwise), and regression 

coefficient (right hand triplet). 

Figure 4 Comparisons between modeled and observed currents at 

the 25m isobath (mooring NA2) sampled at depths of 3m, 12m and 20m, along 

with the NCEP wind velocity sampled at station A. Quantitative comparisons are 

as in Figure 3. 

Figure 5 Comparisons between modeled and observed currents at 

the 10m isobath (mooring EC5) sampled at depths of 2m, 5m and 8m, along with 

the NCEP wind velocity sampled at station A. Quantitative comparisons are as in 

Figure 3. 

Figure 6 Comparisons between modeled (bold) and observed (thin) 

seasonal mean velocity vectors and hodograph ellipses at mid-depth for all six 

mooring locations on the WFS between the 50m and 10m isobaths. 

Figure 7 Modeled seasonal mean velocity vectors for the depth 

averaged and near-surface, mid-water column, and near-bottom sigma levels, 

k=2, 10, and 20, respectively. 

Figure 8 Modeled seasonal mean vertical velocity component fields 

(converted to the z-plane) sampled at the near-surface, mid-water column, and 

near-bottom sigma layers, k=6, 12, and 16, respectively. Bold lines indicate 

upwelling. Thin lines indicate downwelling. 
164

Figure 9 Sea surface temperature and sea surface salinity fields at 

the end of spring 1999 model simulation (May 31, 1999). 

Figure 10 Modeled seasonal mean velocity vectors for the depth 

averaged and near-surface, mid-water column, and near-bottom sigma levels, 

k=2, 10, and 20, respectively, from the model twin experiment forced with wind 

stress only. 

Figure 11 Sea surface temperature at the end of spring 1999 

simulation (May 31, 1999) for the model twin experiment forced by wind stress 

only. 

Figure 12 Evolution of the monthly mean, depth averaged velocity 

vectors for March, April, and May relative to the spring 1999 seasonal mean. 

Figure 13 Modeled temperature sections sampled on March 15 across 

transects originating at DeSoto Canyon, Cape San Blas, Florida Big Bend, and 

Sarasota. The contour interval is 10

C. 

Figure 14 Same as Figure 13 except sampled on April 15. 

Figure 15 Same as Figure 13 except sampled on May 15. 

Figure 16 Modeled near-surface Lagrangian drifter trajectories for the 

period April 1, 1999 to May 31, 1999. Drifters were started along nine different 

transects with initial positions given by solid dots. 

Figure 17 Same as Figure 16 except for near-bottom Lagrangian drifter 

trajectories. 

Figure 18a The relative contributions to the depth-averaged temperature 

balance by ocean circulation and diffusion at station A. Three time series are 

shown: the advection, the diffusion, and their sum (which equals the local rate of 

change of depth-averaged temperatuere). Accompanying each time series are 

their seasonal means and standard deviations in units of 0

C day-1 as measures of 

the seasonal and synoptic scale variability. 

Figure 18b Same as Figure 18a except for station B. 

Figure 18c Same as Figure 18a except for station C. 

Figure 18d Same as Figure 18a except for station D. 

Figure 19 Time series of the depth profiles of the individual terms that 

comprise the temperature balance at station A. The left hand panels show the 
165

horizontal and vertical components of the ocean advection and their sum, and 

the right hand panels show the diffusion, the diffusion plus the advection, and the 

temperature. To the right of each panel is the seasonal mean profile. The 

contour interval for each of the budget terms is 0.05OC day-1, and the contour 

interval for temperature is 0.5OC. Shading indicates warming and clear indicates 

cooling. 

Figure 20 Same as Figure 19 except for station B. 

Figure 21 Same as Figure 19 except for station C. 

Figure 22 Same as Figure 19 except for station D. 

Figure 23 The contributions made to the seasonal mean rate of change 

of SST by each of the advection and diffusion terms. The left panels show the 

advection terms and their sum, and the right panels show the horizontal and 

vertical diffusion terms, the sum of all terms. The contour interval is 0.020

C day-1. 

Bold lines indicate warming and thin lines indicate cooling. 
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