OCS Study
MMS 93/0031
A Review
of the
Physical Oceanography
of the
Cape Hatteras, North Carolina Region
Volume I Literature Synthesis
Science Applications International Corporation
MMS Contract 14-35-0001-30594
Prepared for
U .S . Department of the Interior
Minerals Management Service
Atlantic OCS Region
OCS Study
MMS 93/0031
A Review
of the
Physical Oceanography
of the
Cape Hatteras, North Carolina Region
Volume I Literature Synthesis
Charles E . Adams Jr., Louisiana State University
Thomas J . Berger, SAIC
Willilam C . Boicourt, University of Maryland
James H. Churchill, Woods Hole Oceanographic Institution
Marshall D . Earle, Neptune Sciences Inc.
Peter Hamilton, SAIC
Fred M . Vukovich, SAIC
Robert J . Wayland, SAIC
D. Randolph Watts, University of Rhode Island
October 1993
Science Applications International Corporation
MMS Contract 14-35-0001-30594
Prepared for :
U .S . Department of the Interior
Minerals Management Service
Atlantic OCS Region
DISCIAIMER
This report has been reviewed by the Minerals Management Service 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 .
iii
TABLE OF CONTENTS
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viii
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. 1
1 .1
1 .2
1 .3
Ob j ectives . . . . . .
Scope of the Study . .
Methods and Approach .
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1 .4
Report Organization
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2 .1 .1 B a thyme t ry . . . . . . .
2 .1 .2 Basic Current Structure
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2 .1 .3 Basic Stratification . . . . . . .
2 .1 .4 Counterflow, Deep Western Boundary
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List of Figures .
List of Tables
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List of Acronyms
Glossary
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Acknowledgements
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INTRODUCTI ON
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THE GULF STREAM NEAR CAPE HATTERAS
2 .1
Introduc t ion .
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2 . 2 .1 Ove rvi ew . . . . . . . . . . . . . . . . .
2 .2 .2 Path Envelope/Statistical Summary
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2 .2 .3 Seasonal and Interannual Path Variability
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Meanders, Frontal Eddies, and Filaments .
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2 .3 .1 Meanders, Eddies and Filaments Upstream of
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Current
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Gulf Stream Path Variability
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2 .3 .2 Meanders, Eddies, and Filaments at and
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2 .3 .3 Gulf Stream Related Shelf Features from
Satel lite Observations . . . . . . . . .
2 .4
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Gulf Stream Rings and their Interaction with the
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2 .4 .1 Warm Core Rings
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2 .4 .2 Cold Core Rings
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THE SLOPE SEA .
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3 .1
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3 .2
Near-Surface Slope Sea Waters
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3 .2 .1 Influence of Gulf Stream Warm-Core Rings . . . .
3 .2 .2 Effect of Locally Discharged Gulf Stream Water .
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TABLE OF CONTENTS
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3 .3 .1 Gulf Stream Entrainment of Middle
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Deep Circulation off Cape Hatteras
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3 .4 .1 Overview of the Deep Currents and Processes
3 .4 .2 The Deep Western Boundary Current near Cape
Hatteras . . . . . . . . . . . . . . . . . .
3 .4 .3 Topographic Rossby Waves . . . . . . . .
3 .4 .4 Other Processes of Deep Variability near
Cape Hatteras
IV . THE CONTINENTAL SHELF .
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4 .1 .1 Meteorological Setting .
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4 .1 .1 .2 Regional Patterns . . . . . .
4 .1 .1 .3 Synoptic Scale Disturbances .
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Introduction
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4 .1 .1 .1 Basin Scale Patterns
4 .1 .2 Oceanographic Setting
4 .1 .3 Mean Circulation . . .
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4 .2 .3 Atmospheric and Boundary-Current Forcing .
4 .2 .4 Seasonal Variability .
Shelf Variability and Forcing Mechanisms
4 . 2 .1 Tides
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4 .2 .2 Buoyancy Forcing .
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4 .3
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4 .4
Surface Wave Climatology
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4 .4 .1 Introduction . . . . . .
4 .4 .2 Typical Wave Conditions
4 .4 .3 Extreme Wave Conditions
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4 . 5 .1 Introduc t ion . . . . . . . . . . . . . . . .
4 .5 .2 Synthesis and Interpretation of Observations
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Sediment Transport
4 .5
4 .6
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Regional Bottom Boundary Layer Processes
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4 .6 .1 Bathymetric Setting
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TABLE OF CONTENTS
Section
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4 .6 .2 Sedimentological Setting .
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4 .6 .3 Processes Affecting Sediment Movement
4 .6 .3
4 .6 .3
4 .6 .3
4 .6 .3
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S torms
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Internal Waves
Tidal Currents
The Gulf Stream
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Bottom Fishing
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4 .6 .4 Offshelf vs Onshelf Transport
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4 .6 .5 The Influence of Physiographic Features
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4 .6 .3 .5
V . NEARSHORE PROCESSES
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5 .1 Introduction . . . . .
5 .2 Shelf-Estuary Exchange
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5 .3 Sediment Transport
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5 .3 .1 Basic Concepts . . . . . . . . . . . . . .
5 .3 .2 Research Within the Cape Hatteras Region .
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V I . S UMMARY .
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LIST OF FIGURES
Figure No .
Page No .
Ca t on
Figure 1 .1-1 North Carolina - Virginia Coast
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Figure 2 .1-2 The Gulf Stream current velocity structure on a
vertical section off Cape Fear, NC (from Richardson
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et al . 1969) .
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Figure 1 .1-2 U .S . East Coast Oceanographic Areas
Figure 2 .1-1 The bathymetry of the continental slope near Cape
Hatteras
Figure 2 .1-3
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The Gulf Stream (a) velocity structure, and
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Figure 2 .1-4
Figure 2 .1-5
Figure 2 .2-1
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inset in ( b) (from Watts 1983) .
( b) The mean
bathymetric depth over which the Gulf Stream flows
as a function of distance ( km) along its mean path,
as indicated in inset . . . . . . . . . . . . . . .
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The Gulf Stream (a) velocity structure, and
(b) temperature structure on a vertical section
off Cape Hatteras, NC (from Joyce et al . 1986)
The Gulf Stream ( a) temperature structure, and
(b) velocity structure on a vertical section
about 150 km downstream of Cape Hatteras (from
Halkin and Rossby 1985) . . . . . . . . . . . .
(a) The standard deviation ( km) of lateral
displacements of the Gulf Stream as a function
of distance along the mean path as shown in the
Figure 2 .2-2 (a) The mean path and +/-1 and 2 standard deviation
envelopes of the Gulf Stream path, from Florida to
Cape Hatteras (from Tracey and Watts 1986)
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Figure 2 .3-1 A schematic view of a Gulf Stream Meander (from
Bane et al . 1981) . . . . . . . . . . . . . . .
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Figure 2 .3-2 Types of perturbations observed on the western
boundary in vicinity of and downstream from the
Charleston Bump (from Legeckis 1979)
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Figure 2 .3-3 Structure of meanders, eddies and filaments along
the Carolina Capes (a) thermal structure (from
Bane et al . 1981), (b) schematic of thermal and
velocity structure (from Lee et al . 1981)
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LIST OF FIGURES
Figure No .
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Figure 2 .4-1 Idealized views of the formation of a cold core
ring (top) and a warm core ring (bottom) (from
SAIC 1991)
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Figure 2 .4-2 AVHRR images of CCR event in January 1990 (from
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Figure 3 .1-1 Schematic Slope Sea Circulation (from Csanady and
Hamilton 1988)
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Figure 3 .4-2 Mean equatorward alongshore current speed classified
according to mean potential temperature
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Figure 3 .2-1 MARMAP study ( a) hydrographic stations, and
(b) stations at which Gulf Stream water was
detected and the percentage of casts taken at
these stations which intercepted Gulf Stream
water
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Figure 3 .4-1 Mean current vectors 50-300 m above the ocean
bottom ( from Watts 1991)
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Figure 3 .4-3 Structure of the Deep Western Boundary Current,
constructed by projecting mean equatorward speeds
onto a single section
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Figure 4 .1-1 North Atlantic atmospheric pressure for January
(a) mean ( in mb relative to 1000 mb), (b) standard
deviation .
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Figure 4 .1-2 North Atlantic atmospheric pressure for July
(a) mean (in mb relative to 1000 mb), (b) standard
deviation . . . . . . . . . . . . . . . . . . . . .
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Figure 3 .4-4
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Mean deep current vectors 100 m above the bottom on
a line off Cape Hatteras ( from Pickart and Watts
1990)
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Figure 3 .4-5 Current variance ellipses associated with Topographic
Rossby Waves (TRWs) on a line of deep current meters
100 m above the bottom off Cape Hatteras
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Figure 4 .1-3 North Atlantic temperature for January (a) SST
(°C), (b) air temperature-SST difference (°C)
Figure 4 .1-3c North Atlantic wind speed for January ( m s-1)
Figure 4 .1-4 North Atlantic temperature for April (a) SST
(°C), (b) air temperature-SST difference (°C)
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LIST OF FIGURES
Figure No .
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Figure 4 .1-4c
North Atlantic wind speed for April ( m s-1) .
Figure 4 .1-5
North Atlantic temperature fo r July (a) SST
Figure 4 .1-5c
North Atlantic wind speed for July ( m s-1) .
Figure 4 .1-6
Mean Eulerian currents in the Middle Atlantic
Figure 4 .1-7
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Map of minimum temperature in the water column
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(°C),
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(b) air temperature-SST difference (°C)
Bight (from Be ardsley et al . 1976)
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Figure 4 .1-8 Map of minimum temperature in the water column
below 20 m for August 1971 (from Boicourt 1973)
Figure 4 .1-9 Inferred surface drift from drift bottle returns
(a) July and (b) August 1960-1970 (from Bumpus
1973)
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Figure 4 .1-10 Inferred surface drift from drift bottle returns
(a) September and (b) October 1960-1970
(from Bumpus 1973)
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Figure 4 .2-1 Chapman and Beardsley's (1989) schematic
circulation diagram for the coastal water from
the West Greenland Current to the Middle Atlantic
Bight
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Figure 4 .2-2 Schematic summer cross-shelf velocity profile
for surface currents in the Middle Atlantic
Bight (from Boicourt 1982)
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shoreward intrusion of warm Gulf Stream water during
winter on the South Atlantic Bight, and the
development of the continental shelf front
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temperature (minimum
the cold band, from Nantucket
Point (MP), Hudson Canyon (HC),
Charles (CC) (from
. . .
. . . . . . . . . . . .
.
. 91
Figure 4 .2-3 Schematic diagram of the small and large meander
modes observed during the FRED experiment
.
.
.
Figure 4 .2-4 Oey et al .'s (1987) schematic diagram of the
Figure 4 .2-5 Progression of cold-band
temperature measured in
Shoals (NS), to Montauk
Cape May (CM), and Cape
Houghton et al . 1982)
.
Figure 4 .6-1 Bathymetry of Cape Lookout Shoals showing ridge
and swale features transverse to the shoal's axis
(from Swift et al . 1972)
. . . . . . . . . . . .
x
.
.
.
. 106
LIST OF FIGURES
Figure No . Caption
Page No .
Figure 4 .6-2 Locations of bottle samples taken offshore of Cape
Lookout during September 1969 and the sediment
concentrations found in the samples against bottom
depth (adapted from Rodolfo et al . 1971)
.
.
.
.
Figure 4 .6-3 Total variance of velocities measured within 100 m
of the bottom at locations in the Middle Atlantic
Bight ( from Csanady et al . 1988)
Figure 4 .6-4
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
distance seaward of the Middle Atlantic Bight
.
.
FRED expe r iment
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
Same as Figure 4 .6-6 except showing tidal
current ellipses from deeper current meter
records
. 109
.
.
.
. 111
.
.
.
.
.
. 112
.
.
.
.
.
.
.
. 113
.
.
.
.
.
.
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.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
. 115
.
.
.
.
.
.
.
. 116
Figure 4 .6-8 Low-passed ( 40 h half-power period) filtered
records of currents measured at the 75 m isobath
east of North Carolina as part of the FRED
experiment
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
. 117
Same as Figure 4 .6-8 except showing currents
measured at the 400 m isobath . . . . . . . .
.
.
.
.
.
. 118
.
.
.
.
.
. 120
Sea-bed drifter release locations in Onslow Bay
(dots) and two regions of low drifter return at the
beaches ( shaded) ( from Schumacher 1974) . . . . . .
.
.
. 122
.
.
. 142
Figure 4 .6-10 Low-passed (32 h half-power period) filtered
records of current speed measured over the
continental slope northeast of Cape Hatteras
"Up" is north
Figure 4 .6-11
.
.
Figure 4 .6-6 Semidiurnal M2 tidal current ellipses from the
records of the shallower current meters of the
Figure 4 .6-9
.
.
Near-bottom current speeds measured a short
shelf - edge
Figure 4 .6-7
.
The top plot is a record of beam attenuation
(roughly proportional to suspended solids
concentration) measured by a transmissometer
located 3 m above bottom at a site on the 131 m
isobath off the Delmarva Peninsula (from Churchill
et al . 1992)
Figure 4 .6-5
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
Figure 5 .3-1 Conceptual diagram illustrating the notion of
sediment resuspension and seaward transport by
the combination of gravity and bound infragravity
waves (taken from Wright et al . 1991 and based on
the theory of Shi and Larsen 1984)
xi
.
.
.
.
.
.
.
.
LIST OF FIGURES
Figure No .
Caption
No .
Page
Figure 5 .3-2 Sites of shoreface sediment dynamics studies within
the Carolina Capes region . . . . . . . . . . . . .
.
.
. 143
LIST OF TABLES
Table No .
Caption
Page
4 .4-1 Mean wave height (m) by month and year (after
Hubertz et al . 1992)
. . . . . . . . . . . .
4 .4-2 Occurrences of wave height (m) by month for
all years (after Hubertz et al . 1992)
.
.
No .
.
.
.
.
.
.
. 94
.
.
.
.
.
.
.
.
. 95
.
.
.
.
.
.
.
.
. 96
.
.
.
.
.
.
.
.
. 97
.
.
.
.
.
.
.
. 99
4 .4-6 Wave height (m) as a function of return period
(after Hubertz et al . 1992) . . . . . . . . . .
.
.
.
.
. 100
4 .4-3 Occurrences of peak wave period(s) by month
for all years (after Hubertz et al . 1982)
4 .4-4 Occurrences of peak wave direction (deg) by
month for all years (after Hubertz et al .
1982)
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
4 .4-5 Maximum wave heights (m*10) with associated
peak period(s) and directions (deg/10) by
month and year (after Hubertz et al . 1992)
xii
LIST OF ACRONYMS
AABW AntArctic Bottom Water
AMS American Meteorological Society
AVHRR Advanced Very High Resolution Radiometer
AXBT Airborne ELtpendable Bathythermograph
CAO
CCR
DIALOG
_Cold Air Outbreak
C_old C_ore Ring
Dialog Information Services, Inc . (Reg . Servicemark)
FRED
GABEX
Frontal Eddy Dynamics (Experiment)
_Georgip Bight ZAperiment
genesis of Atlantic Lows Experiment
DWBC
EPA
GALE
GEOREF
GOES
INSPEC
IR
K,
LSW
M=
_Deep Western B_oundary _Current
Environmental Protection Agency
Geovhysical et~erences
_Geostationary Qperational Environmental Satellite
Tnformation Services for the _Physics and _Engineering
C ommunities Database
Tnfra_red
Lunisolar Diurnal Tide Component - 23 .93 hr period
Labrador Sea Water
Principal Lunar Tide Component - 12 .42 hr period
MABL
M_arine Atmospheric Boundary Itayer
MASAR
MECCAS
Middle Atlantic Slope #nd Rise (Experiment)
Microbial Exchange and Coupling in Coastal Atlantic Systems
MARMAP
MMS
NADW
NCDC
NDBC
Mar ine t~,pa ping (Program)
Minerals Management Service
North Atlantic Deep Water
N_ational Climatic Data Center
National Data Buoy Center
NHC
National Blurricane Center
NODC
NTIS
NWS
0,
OCS
P,
psu
RAFOS
_National Qceanographic Data _Center
National Technical Tnformation Service
N ational Weather S ervice
Principal Lunar Diurnal Tide Component - 25 .82 hr period
_uter Continental Shelf
Principal Solar Diurnal Tide Component - 24 .07 hr period
Practical Salinity Unit
SOFAR spelled backwards - refers to an isopycnal following
NOAA
SAIC
National _Oceanic and Atmospheric Agency
float developed after the SOFAR float
Science Applications Tnternational Corporation
SAIW
Sub-A_rctic Tntermediate Water
SPMW
SST
about 700m depth)
Sub-Polar Mode Water
Sea S_urface _Temperature
TRW
USC
VHRR
WCR
Topographic Rossby Wave
Llnited States C_ode
Very l~igh _Resolution Radiometer
Warm _Core _Ring
SEEP
SOFAR
STACS
Shelf _Edge Exchange Program
S,_o und Pixing &nd Ranging - refers to an isobaric float
tracked acoustically in the 'SOFAR' channel (nominally at
Subtropical Atlantic Climate Studies
xiii
GLOSSARY
B-induced motion - Planetary waves in which the restoring force results from
the variation of the Coriolis parameter with latitude .
Barotropic - Motions which are uniform (no variation) with depth .
Baroclinic - Motions which are not baratropic (vary with depth) usually
resulting from stratification and geostrophy .
Cold Dome - A propagating meander trough causes the uplifting of deeper colder
water into near surface water of the trough .
Front - A narrow, horizontal transition zone between two water masses of
differing density .
Geostroyhy - Horizontal motions in which the Coriolis force balances the
horizontal pressure gradient that results from horizontal and vertical
variations in sea water density .
Gulf Stream Western Wall or Northern Wall - front separating lighter Gulf
Stream and Sargasso Sea surface layer waters from denser shelf and slope
surface layer waters . Western Wall is usually used for locations south of Cape
Hatteras and Northern Wall is used north of Cape Hatteras .
Gulf Stream Meander - A large scale (-100 km or longer) horizontal wave on the
GS front propagating in the direction of the current .
Gulf Stream Crest - That part of the meander that displaces GS water closest
to the shelf .
Overwash - The extension of GS/Sargasso Sea water into the surface waters of
the MAB Slope Sea .
Planetary Wave - A wave in which motions occur in the horizontal plane on the
scales of l00s-1000's km . The restoring force results from the conservation
of planetary vorticity or the stretching of Taylor columns .
Trough - The part of the meander that displaces GS water farthest away from
the shelf .
Zone of Maximum Baroclinity - The part of the water column that has maximum
shear and the strongest horizontal density gradients .
xiv
ACKNOWLEDGEMENTS
The continuing cooperative interaction between the Program Manger and Dr .
Robert Miller (MMS/COTR) contributed to the project's success . His support
and suggestions throughout the study are gratefully acknowledged .
xv
I . INTRODUCTION
1 .1
Objectives
Since 1973, the Minerals Management Service (MMS) of the Department of the
Interior has funded projects which provide data and information necessary to
support the MMS's statutory requirement to understand and describe the
"environmental impacts on the human, marine, and coastal environments of the
outer continental shelf (OCS) and the coastal areas which may be affected by oil
and gas development" (43 U .S .C . 1346) . It is in this context that the hIlKS
contracted with Science Applications International Corporation (SAIC) to conduct
"A Review of the Physical Oceanography of the Cape Hatteras, North Carolina
Region ."
During the early 1980's, the MMS leased the oil and gas rights for various blocks
offshore of Virginia and North Carolina . Mobil Oil and their partners acquired
some of these lease blocks and submitted a Plan of Exploration for an area
located 44 .8 miles northeast of Cape Hatteras, North Carolina . This area of
proposed drilling is near one of the most oceanographically complex coastal
regions around the US .
The present study is part of a sequence of programs designed to provide the MMS
with a basis for evaluating the potential environmental impacts of oil and gas
production off of the Cape Hatteras region . Mobil has conducted some preliminary
studies, primarily engineering in nature, of the proposed drill site ; however,
these were identified by a review panel of oceanographers to be insufficient for
the more comprehensive characterization required for impact assessment . As a
consequence, and in keeping with the legislative requirements established by
Congress, two studies have been identified as necessary to provide the needed
oceanographic understanding of conditions in the proposed development area .
These two -studies are : (1) a thorough literature and data review, and (2) a
detailed field measurement program . The present project addresses the first of
these .
The primary objective of this review is to summarize and critique the present
state of knowledge of the physical oceanography of the complex region offshore
of Cape Hatteras, North Carolina, within the context of understanding the
regional circulation and its relation to the fate of any discharges resulting
from offshore oil and gas activity . The two other related objectives are to
produce an annotated bibliography of the pertinent literature, primarily from
1970 to the present, and to identify relevant oceanographic data sets which can
provide a basis for an improved understanding of circulation patterns and
physical oceanographic conditions in the study area .
The intended audience for this review includes trained oceanographers and
government decision makers with some knowledge of oceanography gained through
their work . It is hoped that our educated public will also benefit from this
study .
1 .2
Scope of the Study
A review was conducted of all published scientific literature covering physical
oceanographic processes in the region between 34°30'N and 37°N, westward from
1
73'W to the North Carolina-Virginia coast but excluding the Chesapeake Bay . The
topics addressed, and hence the literature surveyed, included the following :
• Atmospheric forcing,
• Warm and cold core ring effects,
• Gulf Stream path variability,
• Regional bottom boundary layer processes,
• Shelf break exchange processes,
• Shelf current variability,
• Shelf/estuary exchange processes,
• Slope Sea circulation, and
• Tides and tidally induced mixing and particle transport .
The
the
and
for
study area, shown in Figure 1 .1-1, is centrally located on the east coast at
convergence of a number of bathymetric provinces . To the north is the shelf
slope region of the Middle Atlantic Bight and to the south the same features
the South Atlantic Bight .
Figure 1 .1-2 shows the various regions which may be discussed during this report .
George's Bank is to the far north and east and southeast of Cape Cod . The shelf
just to the south, offshore of southern New England and Long Island, is broad,
on the order of 120-150 km, and narrows gradually to the south until the region
between the Chesapeake Bay and Cape Hatteras, where the shelf width decreases
relatively rapidly with southward distance . In the southern half of the Middle
Atlantic Bight, Chesapeake Bay and Delaware Bay contribute regionally important
quantities of freshwater to the shelf system . Offshore of the shelf is the Slope
Sea, which overlies the Middle Atlantic slope and rise . These descend fairly
continuously to depths of 3500 to 4500 m . Still further seaward is the Gulf
Stream, which has a fundamental influence on many aspects of the oceanography of
the region .
To the south of Cape Hatteras, the coastline changes to a regional orientation
of northeast to southwest, which is maintained until Georgia . The shelf tends
to widen with increasing distance south of Cape Hatteras until approximately
South Carolina and Georgia . From northern Florida to the Florida Straits, the
shelf narrows, and its orientation is more north-south . Seaward of the shelf
break in the South Atlantic Bight, the local continental slope descends to
approximately 800-1200 m on the Blake Plateau . The relatively gently sloping
bottom feature, which is the Blake Plateau, widens with increasing distance south
of Cape Hatteras . The offshore edge and the upper portion of the Blake
Escarpment is reasonably defined by the 2000 m isobath . On the inner portion of
the Blake Plateau, and just seaward of the shelf break, is the mean position of
the Gulf Stream thermal front (largest gradient in surface temperature typically
measured by satellite infrared sensors) . This front is variously called the Gulf
Stream front, the "North Wall" or the "western wall ."
Several bathymetric conditions converge at or near Cape Hatteras . The southern
extension of the deeper portion (depths greater than approximately 2000 m) of the
Middle Atlantic Bight slope trends offshore to form the Blake Escarpment, which
tends to direct any southward directed deep flows offshore . Also at Cape
Hatteras, the Gulf Stream tends to continue to flow to the northeast, while the
continental slope changes from SW-NE to S-N, as one proceeds generally northward .
As a consequence, the Gulf Stream changes from a western boundary current that
is laterally constrained by the South Atlantic Bight shelf and vertically by the
2
I
78
38 N
W
77
W
76
W
75
W
74
W
73
W
4P
le
1 37 N
72
~
W
38 N
37 N
/
1 36 N
36 N
1 35 N
35 N
1 34 N
33 N
cP"'~
y
~
34 N
C.P.
78
W
77
W
76
W
75
W
74
W
73
W
72
33 N
W
Figure 1 .1-1 North Carolina - Virginia Coast . Study area is outlined by solid
lines . Locations of the Carolina Capes and offshore shoals are
labelled . Long dashed lines with arrowheads indicates approximate
Gulf Stream axis .
3
43 N
83 W 81
W 79 W 77 W 75 W 73 W 71 W 69 W 67 W
43 N
NY
MA
CT
41 N
OH
~ MD , .2
39 N
VA
. . ... .... ....... .•. :::...
200 m . . .. . .. .::::: : . .. .. . . ::: . . . :. . :
\0~zr~r- ~~ Hudson Canyon
106-Miie Site
NJ
Cape May ~ sea
s
\oPeco~~~
~
a Charles
Chesapeake Bay
37 N
Cape Cod
P~~a~
vvv
~
.. .. .......: - `..
VG
PA
Delaware Bay
Gulf~ of Maine '. . . ... ... . .
.
Gu~1\ Stceam
41 N I
39 N I
~~ .
37 N I
NC
e Hatteras
35 N
Cape Lookout
::
Cape Fear F
SC
33 N -
ape Romain
0
c
GA
Sar~sso Sea
oo°
o. ~
o, F
.. .
4P . ti
Qr :
31 N
35 N I
31 N I
Charleston Bump ' . ..
::
:.. ~~--,.. .
29 N
Blake-Bahama Escarpment
FL
Blake Plateau :/
a • . . .. . . ... .
.
... . . .. . ... .. ... ..
27 N
33 N I
29 N I
27 N I
. ..~ . . . : : . . .. . .. . . .. . . ..:.. .... . ;~~:~.~
. • ~~
I
25 N
~
83 W 81
Figure 1 .1-2
25 N
W 79 W 77 W 75 W 73 W 71 W 69 W 67 W
U .S . East Coast Oceanographic Areas . Study area is shaded . Major
features mentioned in the report are labelled .
4
Blake Plateau, to an open ocean current which is comparatively unconstrained by
bathymetry as it passes Cape Hatteras .
The Carolina Capes, which define a series of crescentic bays, extend offshore as
shoals, which tend to isolate the circulation within the bays . Of particular
importance is Diamond Shoals offshore of Cape Hatteras (Figure 1 .1-1) .
1 .3
Methods and Approach
Identifying literature relevant to physical oceanographic conditions and
processes in the study area involved use of existing bibliographies, searches of
a number of literature databases, and reviews of more recent and particularly
relevant journals . It was recognized that incorporation of these data into a
computer-based bibliographic program would facilitate the compilation of
citations in several standard, user-defined formats, integration of differing
sources of information, and identification and elimination of duplicate
citations . Thus, one of the first project tasks was to identify and structure
a commercially available citation software package .
Initial project activities were directed toward using existing standard and
annotated bibliographies (e .g . Imamura 1989) to build the present literature
database . These were supplemented by bibliographies from the various program
principal investigators . The fairly extensive, centralized bibliography
maintained in Raleigh was also incorporated .
Separate electronic searches were made of Dissertation Abstracts, Geophysical
References (GEOREF), and Water Resources using the Information Services for the
Physics and Engineering Communities Database (INSPEC) search facilities at North
Carolina State University . SAIC's on-line Dialogue Information Services, Inc .
(DIALOG) facilities in McLean, Virginia were used to search the National
Technical Information Service (NTIS) database . Finally, a manual search was made
of the Journal of Geophysical Research, Journal of Physical Oceanography, DeepSea Research, and other common oceanographic journals for papers published from
the end of 1991 through July 1992 .
Not all bibliographic sources provided data in a standard form, nor contained all
the elements required by the citation structures being used for this study .
When possible, additional material was added to individual citations so that the
resulting compilation would be as complete as possible . This often involved
locating the original articles and entering an appropriate abstract . For each
entry, it was necessary to establish and enter a series of key words so that
topical and geographic searches could be made .
When most citations had been incorporated in the computer-based bibliography, a
key word search was made to identify those which were particularly germane to the
present study area . While this relevance was a primary basis for inclusion in
the overall literature search, some existing bibliographies contained citations
which were not relevant to the present program objectives . The key word search
identified approximately 1300 entries which were specific to the topical and
geographic area of interest . This listing was provided to the program principal
investigators for review and comment . The net effect of their additions and
deletions resulted in a bibliography of about 1100 entries covering the period
from 1970-July 1992 .
5
During a July 1992 meeting of the program staff and the principal investigators,
a draft report outline covering the topics listed above was discussed, and
writing assignments were made . Each of the principal investigators, who had been
chosen because of their expertise and interest in a particular aspect of the
program, had assignments to write a specific section or to collaborate with
another principal investigator on a section . In the case of collaboration, one
person was assigned to have the lead role .
1 .4
Regort Organization
The remainder of this report provides reviews of processes generally associated
with the Gulf Stream (Chapter 2) ; The Slope Sea and the Deep Western Boundary
Current (Chapter 3) ; The Continental Shelf (Chapter 4) ; and the Nearshore region
(Chapter 5) . A summary of the processes as related to potential oil and gas
activities is presented in Chapter 6 . The annotated Bibliography is in Appendix
A, which is bound separately because of its bulk .
6
Imamura, E . (ed .) 1989 . Summary of the physical oceanographic processes for the
U .S . Atlantic
and Southeastern Gulf of Mexico .
Battelle Memorial
Institute . Ventura . 199 pp .
7
II . THE GULF STREAM NEAR CAPE HATTERAS
2 .1 Introduction
The continental shelf at Cape Hatteras is much narrower than to the north or
south (20-30 km compared to nearly 200 km wide off New Jersey and Georgia), and
the strong currents of the Gulf Stream flow relatively close to shore . The Gulf
Stream plays a leading role in determining the currents and the sources and fates
of water parcels near the outer shelf and the shelf break . Moreover, the Gulf
Stream path shifts on- and offshore from year to year, season to season, and with
meandering periods from months to as short as a few days . The meanders introduce
variations in the strength and structure of the current, and they produce frontal
eddies as well on its inshore edge that can cause upwelling and important crossshelf exchanges . Consequently a large degree of variability is introduced to the
currents and water properties on the shelf itself near Cape Hatteras by the
proximity of the Gulf Stream and by several processes associated with the
stream's energetic eddy variability .
In order to understand the currents and variability near Cape Hatteras, it is
essential not only to understand the basic structure of the Gulf Stream velocity
and water property fields (such as temperature, salinity and density fields), but
also to understand its variability . The basic structure of the Gulf Stream and
its physical setting are summarized in this introduction, while major processes
contributing to its variability are the subjects of later sections .
2 .1 .1 . Bathymetry
From the Gulf Stream's beginning in the Florida Straits up to just south of Cape
Hatteras, it flows offshore of the continental shelf and over the broad Blake
Plateau in bathymetric depths of approximately 800-1200 m . However, the Blake
Plateau narrows northward from about 31°N, to a small remaining wedge at the
latitude of Cape Fear . It narrows further to become non-existent by 34°30'N
(about 50 km south of Cape Hatteras) . Figure 2 .1-1 (from Newton et al . 1971)
illustrates the topography of the continental slope near Cape Hatteras . The
continental slope falls off extremely steeply at Cape Hatteras, such that the
distance from the 100 m or 200 m isobath to the 2000 m isobath is only about 20
km . North of Cape Hatteras the Gulf Stream continues to flow to the northeast,
while the bathymetric contours turn more northward ; the Gulf Stream thus leaves
the continental margin and, within 200 km downstream of Cape Hatteras, flows in
water over 3000 m deep . To the north of Cape Hatteras also, in the wedge of
waters between the Gulf Stream and the continental shelf are the waters of the
Slope Sea, whose circulation and characteristic processes are the subject of
Chapter III .
2 .1 .2 . Basic Current Structure
The mean current strength, width, and structure of the Gulf Stream and its
associated temperature and density structure are conveniently illustrated by
vertical sections crossing approximately normal to the current and front (a
region with a large horizontal gradient in properties, especially density) .
Figures 2 .1-2 through 2 .1-5 show representative vertical sections just upstream,
at, and just downstream of Cape Hatteras .
9
10
a;
~a
d
~
u
~
x
~a
14
U
10
d
r.
N
a
0
r4
In
b
~~
w
r4
Id '.4
41
a ,4
rl~
.a ct
0
41 4
9: d
,
~
dri
,a
0
0
~z
~d
II
a
~
ws+
o a
d
41 "4
~ d
.a ~d
4)
m a
~
~
~
~
N
h0
.vl
k.
km
0
0
.
20 40 60 80 100 120 140
.
400
600
800
~
120
200 . . •r.
100
. . . .. .
. . .:
. •
. •. .. .
.
.
•
~
80
60
:.
. . -•...... :
....
~
4 0
..
, .
Cape Fear, NC
Jun-Ju I 1968
. ..
\
I
20/
/
,
Figure 2 .1-2 The Gulf Stream current velocity structure on a vertical section
off Cape Fear, NC (from Richardson et al . 1969) . Average isotach
contours (in cm s-1) are plotted from a repeated section, where
the velocity profiles were obtained by dropping transport floats
to several depths .
11
(a)
STATION NUMBER
0
12 10 9 8 7 6
5 4
3
2
~'
A
\ \/ r \\\
STATION NUMBER
12 11109 8 7 6 5 4 3 2
_.
0
8
-40
/
-20
1000-
(b)
-/
o/
1000
.
.
-5
E 2000
_
I-
-/O
`-
Q.
~
N
W
0
O
3000
. ,
•
E 2000
5
•
. , .
W
O . . .
.
~
•
NITY ANQMAL~(~4
SALI
.
. ,
.
a.
_5
'
•
~ 3000
.
'
' ,
SILICATE
.
•
4000
4000
VELOCITY
lcm/sec)
TEMPERATURE
(•C)
•
5000
5000
0
100
•
200
DISTANCE (km)
300
0
100
200
300
DISTANCE (km)
Figure 2 .1-3 The Gulf Stream (a) velocity structure, and (b) temperature structure on a vertical section off
Cape Hatteras, NC (from Richardson 1977) . Geostrophic velocities calculated relative to
currents directly measured by near-bcttom current meters . Station 2, furthest offshore, was
thought to be in a Gulf Stream cold-co :'e ring .
(a
)
f40r 45
42
~
PO T ENT/AL TEMPERATURE
49
(b)
42
45
49
0
20
swim
f0 r
f00
~
20
~
1000
/
5~
f0
~
`
2000
5
j
~
w
~
h
~
~
1000
5
/
---f0
~
'
--
`-'------Z 2000 - - - - -
~
_
'------
~ 3000
.
f5
-
3000
,
4000
,-
4000
20km
. ------ ------------
.~.•.
- -__''-; ._
-_
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.
-p
.
20km
Figure 2 .1-4 The Gulf Stream (a) velocity structure, and (b) temperature structure on a vertical section off
Cape Hatteras, NC (from Joyce et al . 1986) . The velocities were calculated geostrophically
relative to shipborne acoustic doppler current profiler measurements averaged at the 60 m level
between stations .
(a)
(b)
6 11 13 151616 16 1413 9 7 6 11 13 15 1616 16 1413 9 7
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2000
T(°C)
b)
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(cm/s)
<0
-80 -40 0 40 80 120 =80 -40 0 40 80 120
CROSS STREAM DISTANCE ( km )
Figure 2 .1-5 The Gulf Stream (a) temperature structure, and (b) velocity
structure on a vertical section about 150 km downstream of Cape
Hatteras (from Halkin and Rossby 1985) . The velocity profiles
were measured directly using an acoustically tracked profiling
float, called Pegasus . The transect was repeated 18 times during
three years, and the results were translated into "stream
coordinates" before averaging .
14
This discussion begins with the upstream section . Richardson et al . (1969) made
direct measurements of velocity at Cape Fear ; their measurement program repeated
a transect during two months in which they dropped transport floats to several
intermediate depth intervals, separated by 100-200 m, at each station . The Cape
Fear section current speed (in cm s-1, where 1 knot - 52 cm s-1) is contoured in
Figure 2 .1-2 . The current has a width of about 120 km near the surface and depth
of 800-1000 m, and the core speed is over 140 cm s-1 . Because these results
present an average of several sections, during which the current is likely to
have meandered, the width shown may be wider and the core speed is likely to be
slower than would be observed in an instantaneous section .
Figures 2 .1-3a,b and 2 .1-4a,b show two instantaneous vertical sections of
velocity and temperature, on a transect directly off Cape Hatteras . The current
width near the surface is about 100 km . The current speed typically increases
rapidly as one moves seaward across the onshore side, "left" of the current
looking downstream, to a core velocity of 140-180 cm s-1 at about 30-40 km beyond
the current's shoreward edge . With increasing depth the center of the velocity
core characteristically tilts offshore 20-40 km from the surface velocity core .
In some sections the subsurface velocity core exceeds the surface velocity
slightly on the offshore side of the Gulf Stream . In each case, the current
exceeds 20 cm s-1 down to about 1000 m depth, and a deep core of current about
5-10 cm s-1 extends deeper . In one of the examples shown (Figure 2 .1-4) the
northeastward flow extends to the ocean bottom, more than 3000 m deep, whereas
in the other example (Figure 2 .1-3) there is counterflow of 5-10 cm s-1 at the
bottom . Further introduction to the deep flow is presented in Section 2 .1 .4 .
Halkin and Rossby (1985) report on the most representative mean section of the
Gulf Stream, taken on a transect about 150 km downstream of Cape Hatteras near
73°W longitude . The transect was repeated at two-month intervals for three
years, and the velocity profiles were measured directly using Pegasus, an
acoustically tracked, free falling profiling instrument . In Figure 2 .1-5 their
measurements are shown averaged in "stream coordinates", i .e ., with the origin
of the transect shifted to follow the "north wall" of the Gulf Stream with each
survey . By averaging the results in "stream coordinates" the main effects of
apparent broadening and slowing of the current due to meandering are avoided .
In this section the current width is about 140-160 km, the velocity core is over
150 cm s-1 at the surface and tilts offshore with increasing depth, the 20 cm s-1
isotach extends to 1000 m, and a 10 cm s-1 core extends below 2000 m .
2 .1 .3 . Basic Stratification .
The Gulf Stream water temperature and density are vertically stratified, with
density increasing and temperature decreasing downward in accord with buoyancy
constraints . The density and temperature surfaces (isopycnals and isotherms) are
strongly inclined (baroclinic) throughout the water column across the Gulf
Stream ; i .e ., they deepen at all levels by about 800 m from the shoreward side
to the Sargasso Sea side .
This stratification and its inclination across the Gulf Stream are illustrated
in the temperature sections accompanying Figures 2 .1-3, -4, and -5 . Temperature
decreases with depth from above 20°C near the surface to below 5°C in roughly the
upper 1000 m, and it continues to decrease more slowly with depth down to below
2°C in abyssal depths around 4000 m . The main thermocline (where the temperature
changes sharply from about 17° to 6°C) rises shoreward across the Gulf Stream,
15
and many of the isotherms intersect the sea surface . Consequently the Gulf
Stream represents a "front" of relatively sharp horizontal changes in
temperature . The surface expression of this front gives rise to the now-familiar
means of viewing the Gulf Stream using satellite infrared (IR) imagery of the
sea-surface temperature (SST) . The sharp "north wall" or surface SST front,
provides a convenient means of tracking the path of the Gulf Stream, as described
in Section 2 .2 . In the upper 100-200 m there is also seasonal modulation in the
temperatures of as much as 3-4°C warmer or cooler than any one of these three
figures . The high velocity core nevertheless always brings from lower latitudes
a core of warmer waters that is also visible on satellite IR imagery .
On the offshore side of the Gulf Stream in depths around 300-400 m is a
relatively thick lens of "Eighteen Degree Water" that is seen in all three
temperature sections (Figures 2 .1-3, -4, and -5) and is common throughout the
Sargasso Sea .
Water properties such as temperature, salinity, oxygen, silicate, and other
chemical tracers tend to be relatively constant along density surfaces and to
change more rapidly with density . In this sense, the Gulf Stream represents a
"front" of sharp horizontal change in many water properties . For a more complete
summary of water mass properties, Worthington's (1976) monograph on the North
Atlantic circulation has summarized watermass properties in the entire Gulf
Stream system, and Watts (1983) has reviewed additional details regarding water
masses in the Gulf Stream from the Florida Straits to south of New England .
Additional details of water masses are given in Stefansson et al . (1971) and in
Stefansson and Atkinson (1971) .
2 .1 .4 . Counterflow, Deep Western Boundary Current
Lastly, we note on one or both lateral edges of the Gulf Stream that counterflowing currents often exist, such as illustrated in an instantaneous transect
(e .g ., Figure 2 .1-3) or in the mean transect (Figure 2 .1-5) . In instantaneous
transects the counterflows astride the core can vary from less than 10 cm s-1 to
over 80 cm s-1, depending upon the existence and distribution of rings, eddies,
and filaments ; these processes are treated in Section 2 .3 .
In the Cape Hatteras region, the Deep Western Boundary Current (DWBC) flows
south-westward along the continental slope under the Gulf Stream . Evidence of
the DWBC can be seen in both of the instantaneous transects shown earlier
(Figures 2 .1-3 and -4) . Its mean flow is typically only 5 cm s-1 to 10 cm s-l ;
however, the deep currents in this region are highly variable, with peak speeds
of 20 to 30 cm s-1 commonly observed . The DWBC is treated in Section 3 .2 .
Pickart and Watts (1990) have shown that the details of the deep velocity
structure are often substantially altered from one instantaneous transect*to
another by the presence of Topographic Rossby Waves (TRWs) . Hence, in different
transects off Cape Hatteras, whether the Gulf Stream appears to reach the bottom
or not, such as exemplified in Figures 2 .1-3 and 2 .1-4, is likely to depend upon
the amplitude and phase of the TRWs propagating at that time through the
transect . TRWs are summarized in Section 3 .2 .
16
2 .2 Gulf Stream Path Variabilit
2 .2 .1 Overview
The Gulf Stream exhibits wave-like lateral displacements that vary with time all
along its path, from the Straits of Florida to the Grand Banks . These lateral
shifts affect the entire current and its baroclinic structure throughout the
water column . Moreover, the predominant wavelengths and the propagation and
growth characteristics of the path fluctuations change from region to region
along the Gulf Stream, with changes being particularly distinct between south and
north of Cape Hatteras due to topographic differences .
Seasonal and interannual changes are summarized in this section, while shorter
period shifts, denoted meanders, are summarized in Section 2 .3 along with their
important associated frontal eddies and filaments . This description begins with
an overall summary of the envelope through which the Gulf Stream path shifts due
to combined long- and short-period processes .
2 .2 .2 Path Envelope/Statistical Summary
A summary of Gulf Stream path variability and meanders is given in Watts (1983) .
In particular, Figure 2 .2-1 draws together information from three previous
studies on the r .m .s . lateral displacement amplitudes of the Gulf Stream path
from Florida to south of New England (Bane and Brooks 1979 ; Watts and Johns 1982 ;
and Halliwell and Mooers 1979) . The inset of the lower panel in Figure 2 .2-1
indicates the mean path of the Gulf Stream . Along this path is shown the mean
depth over which the Gulf Stream flows ; the abrupt dropoff is evident in the
lower panel, as the Blake Plateau ends near Raleigh Bay and Cape Hatteras
(described in Section 2 .1 .1) .
Several features are evident in the standard deviation envelope for the path, as
shown in the upper panel of Figure 2 .2-1 : From the Florida Straits to about
Savannah (32°N) the r .m .s . displacements are only about 5 to 12 km, associated
mainly with shelf waves and small amplitude meanders . From Charleston to about
Cape Fear (34°N), the displacements have grown to 20-25 km r .m .s ., associated
with a topographic feature known as the "Charleston Bump," as discussed further
in Section 2 .3 .1 . The displacement amplitudes decay downstream of Cape Fear to
a minimum of less than 10 km r .m .s . near Cape Hatteras . This broadening of the
path envelope off Charleston and subsequent narrowing along the Carolina Capes
is also well illustrated in Figure 2 .2-2a, from Olson, et al . (1983) .
North of Cape Hatteras, as the current enters deep water, meanders grow rapidly
downstream, and the path contains increasing interannual variability ; Figure
2 .2-1 shows the corresponding rapid downstream growth of the r .m .s . path
displacement . Histograms of the path displacement, for 3-4 year periods both
upstream and downstream of Cape Hatteras, are also shown in Figure 2 .2-2(a), and
2 .2-2(b) from Tracey and Watts (1986) .
2 .2 .3 Seasonal and Interannual Path Variability
The Gulf Stream path varies on seasonal and interannual time scales . Simple
theoretical ideas have suggested that changes in current strength should be
accompanied by north-south shifts of the current with some expectation that the
17
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Bane and Brooks (1979)
: 100-1•200 Km
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Hal liwel I and Mooers (1979) : 1400-2300 Km
(a)
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6 0
500
1000
1500 2000
2500 3000
km
Figure 2 .2-1 (a) The standard deviation (km) of lateral displacements of the
Gulf Stream as a function of distance along the mean path as
shown in the inset in (b) (from Watts 1983) . (b) The mean
bathymetric depth over which the Gulf Stream flows as a function
of distance (km) along its mean path, as indicated in inset .
18
(b)
~-
°~-
.,~
9•
GULF STREAM NORTH WALL
DISPLACEMENT HISTOGRAMS
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Figure 2 .2-2 (a) The mean path and +/-l and 2 standard deviation envelopes of the Gulf Stream path, from
Florida to Cape Hatteras ; at the right are shown histograms of probability distribution of the
Gulf Stream position across the five transects indicated (from Olson et al . 1983) . (b)
Probability histograms of the Gulf Stream position across four transects just downstream of Cape
Hatteras (from Tracey and Watts 1986) .
current might be stronger in spring after stronger wind-forcing and thermohalineforcing in winter, and weaker in fall after correspondingly weaker summertime
forcing . Iselin (1940) noted seasonal changes in the Gulf Stream transport at
68°W with a winter maximum accompanied by a southward shift of the axis of the
Gulf Stream, and with a fall minimum accompanied by a northward shift of the
axis . In contrast, however, the seasonal cycle in Gulf Stream transport in the
Florida Straits (the only place where a seasonal cycle has been consistently and
accurately observed) is shifted several months from this simple conceptualization
regarding the forcing . Multi-year Sub Tropical Atlantic Climate Studies (STACS)
in the Florida Straits have found an average transport of about 30 Sv (1 Sv - 106
m3 s-1) and a peak-to-peak range of about 8 Sv (i .e ., 26-34 Sv) ; however, the
annual cycle has a range of only about 2 Sv, with maximum transport in summer .
Near Cape Hatteras the story becomes confused regarding whether there is a
seasonal cycle in transport at all . Worthington (1976) had indicated an annual
cycle from hydrographic sections and baroclinic transport estimates far
downstream of Cape Hatteras . Blaha (1984), Tracey and Watts (1986), and Kelly
(1991) discuss various indicators of transport and surface-transport (mean
surface current, from sea surface height differences across the Gulf Stream) .
In their three-year set of Pegasus transects of velocity and temperature just
downstream of Cape Hatteras, Halkin and Rossby (1985) found only very weak
evidence of an annual cycle in either baroclinic or total transport ; their
results showed large interannual variability, with a tendency (possibly resulting
from just three years of data) for the transport averaged in February-July to
exceed that averaged in August-January . To some extent, the more thorough the
study, the less certain is the existence of an annual cycle in transport . It may
be that, in the presence of relatively large long-period and interannual changes
in transport, any study covering only several years will simply project some
portion of the variability onto an annual cycle .
Regardless of the weak state of evidence of seasonal cycle in transport near Cape
Hatteras, the evidence for a seasonal cycle in position appears to be more
consistent . Auer (1987), using five years (May 1980-May 1985) of satellite IR
frontal analyses, found an annual cycle for the Gulf Stream's landward edge in
the longitudinal band 70° to 44°W . (In this region, the Gulf Stream's landward
edge is usually called the "north wall" .) The September mean position was
farthest to the north, and the March position was farthest to the south . Similar
studies for an annual cycle were made in smaller longitudinal bands . In the band
74°-70°W, which is just east of Cape Hatteras, Auer (1987) found no significant
annual shift . SAIC (1990) found no clear annual cycle in 62 months of satellite
IR imagery from Cape Hatteras to 73°W . Vukovich (1990) used 10 years of monthly
average north wall positions from satellite IR data and found an annual cycle at
74°W, 70°W, and 65°W . The amplitude of the annual cycle at 74°W was 7 km, at
70°W, 12 km, and at 65°W, 15 km, which are considerably smaller than Auer (1987)
found further east . At 74°W and 65°W, the maximum northward shift occurred*in
October and the maximum southward shift in March and April, and at 70°W, the
maximum northward and southward shifts were in January and June, respectively .
The annual cycle for the Gulf Stream north wall in the 74°W-65°W band shows an
amplitude of 10 km with maximum northward and southward shifts in October and
April, respectively .
Tracey and Watts (1986) used four years of Inverted Echo Sounder data in the Gulf
Stream at 73°-74°W (about 150 km downstream of Cape Hatteras) to calculate
monthly averages of Gulf Stream position and to look for an annual cycle . On
20
average, superimposed upon relatively larger interannual variability, they found
shifts from its mean path to a maximum in September of about 10 km shoreward, and
about 10 km offshore in April . In contrast, the interannual range in position
is over 60 km even this close to Cape Hatteras .
Vukovich (1990) found a linear correlation coefficient between the annual cycle
of the north wall in the band 74°W-65°W and the Straits of Florida transport
(i .e ., cable data) of 0 .99 with a four month phase lag . The Straits of Florida
transport reached maximum and minimum values about four months prior to the
maximum northward and southward shifts of the north wall .
Besides an annual variation, Vukovich (1990) also showed that there were
long-term trends in the north wall position variations and that these were about
twice as large as the annual variations in the longitude band 74°W-65°W . The
maximum northward shift of the north wall occurred in 1985 when the Straits of
Florida transport was a maximum . The maximum southward shift of the north wall
took place in 1981 . It is not known if the Straits of Florida transport was a
minimum at that time because cable transport data were not available before 1982 .
The linear correlation coefficient for the two data sets was 0 .75 with no phase
lag . The amplitude of the interannual variations in position between 74°W - 65°W
was about 30 km .
2 .3 Meanders . Frontal Eddies, and Filaments
This section treats wave-like lateral shifts (meanders) of the Gulf Stream path,
which have time periods shorter than a year, and their associated frontal eddies
and filaments . A schematic of a Gulf Stream meander is shown in Figure 2 .3-1 . The
important distinguishing feature of meanders of shorter periods, and particularly
of shorter wavelengths, is that eddy circulations are intrinsically associated
with them . They perturb the velocity and temperature structure of the main Gulf
Stream front and they may induce eddy-like circulations on either side of the
current within the crests and troughs . Filaments of warmer Gulf Stream waters
often extend back from the crests on the shoreward side of the main front .
The characteristic wavelengths, periods, and propagation speeds of meanders
change considerably from upstream to downstream of Cape Hatteras, due in part to
changes in bottom topography under the Gulf Stream (as summarized in Section
2 .1) . The Blake Plateau ends just south of Cape Hatteras, and the Gulf Stream
leaves the continental margin as it flows to the north over abyssal depths . The
depth and steepness of the bathymetry also affect the growth and decay
characteristics of meanders . Because of these topographic/geographic
distinctions, the properties of meanders south and north of Cape Hatteras are
summarized in the following two subsections .
2 .3 .1 Meanders, Eddies and Filaments Upstream of Cape Hatteras
Upstream (south) of Cape Hatteras, meanders typically have periods between 2 and
10 days, and wavelengths between 50 and 300 km . The propagation speeds can range
from zero for stationary waves, to 60 km day-1 in the downstream (northward)
direction . The r .m .s . amplitudes of path-deflection range from 10-30 km,
accounting for most of the variability that was shown in Figure 2 .2-1 . Brooks
and Bane (1981) report two dominant bands of meandering, near 3 days and 7-8
21
SCHEMATIC MEANDER WITH WARM FILAMENT
.
PHASE
PROPAGATION
TEMPERATURE SECTION
WARM COOL WARM
0-
OOe ©O® ©
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1 LBACK IN MAIN STREAM
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FRONT (_)
IN COLD TONGUE
IN MAIN STREAM,
CROSSING TRAILINI
FRONT(--)
IN WARM FILAM
Figure 2 .3-1 A schematic view of a Gulf Stream meander . The cyclonic flow
of a frontal eddy in the meander trough may be seen . The time
series of velocity components and temperature measured by an
instrument at a fixed point as the meander passes by are
shown . The distance between crests at A and F is the meander
wavelength . Distance (at A or F) from the mean position to
the instantaneous position is the amplitude . Trough is
indicated at C . Arrows indicate direction of flow and suggest
cyclonic circulations in troughs (from Bane et al . 1981) .
22
days, with the most energetic wavelengths between 100-250 km, and a central value
of propagation speeds of 30-40 km day-1 .
Cyclonic frontal eddies (Lee 1975) are commonly found on the Gulf Stream's
western boundary from 27°N to the Charleston Bump, a bathymetric feature at about
31°30'N seen in the 500-800m isobath . These eddies are generally smaller in
amplitude than those found north of the Charleston Bump, having an average
wavelength and amplitude of about 200 km and 20 km, respectively . They have an
average downstream phase speed of 30 km day-1, and period of about nine days
(Vukovich et al . 1979) . These eddies decay and may become very small before they
reach the Charleston Bump . Often only remnants of these perturbations are
observed to reach the Charleston Bump region ; however, studies summarized below
suggest some continuity between eddies up- and downstream of the Charleston Bump .
Moreover, the frontal eddies and filaments along the Carolina Capes appear to
have similar structure to those observed further south .
Brooks and Bane (1978) note a persistent seaward deflection of the Gulf Stream
path at the latitude of the Charleston Bump . A study by Legeckis (1979) focused
on the wavelike perturbations and used National Oceanic and Atmospheric Agency
(NOAA) satellite IR SST data for the period 1974 through 1977 . Legeckis (1979)
suggested the deflection was a product of bottom steering by the ridge and trough
nature of the Charleston Bump . See also Bane and Brooks (1979), Olson et al .
(1983), and Auer (1987) . Legeckis (1979) noted the lateral meanders of the Gulf
Stream increased by a factor of three on a seasonal time scale (see also Bane and
Brooks 1979 and Olson et al . 1983) . Bane and Dewar (1988) have further indicated,
from combined current meter mooring and satellite observations, that the Gulf
Stream path downstream of the Bump has a bimodal nature, switching between strong
and weak deflection in a matter of a few days . They reported that, depending on
the state of deflection, the low-frequency variability from the Charleston Bump
to Cape Hatteras is affected . The strongly deflected state is accompanied by
larger amplitude, slower propagating (20-25 km day-1, as compared to 35-60 km
day-1), longer period (16 days, compared to one week) meanders .
The Geos-3 altimetry data showed a persistent center of negative sea-surface
height-anomalies over and downstream from the Charleston Bump (Huang et al . 1978 ;
Robinson et al . 1983), suggesting that the Gulf Stream deflection results in a
quasi-stationary cyclonic gyre (i .e ., the Charleston gyre) . Upwelling detected
in association with this gyre produced an enhancement of surface layer primary
production of sufficient strength to be detected using the Nimbus-7 Coastal Zone
Color Scanner (McClain and Atkinson, 1985) .
Legeckis (1979) grouped his satellite IR observations into five typical
perturbations found in the region (see Figure 2 .3-2) .
T,ype 1-Pure deflection . The western boundary of the Gulf Stream, which was
defined by a strong SST gradient on the western side of the Gulf Stream, was
deflected seaward at the Bump and then landward downstream from the Bump with no
apparent wavelike activity on the boundary from the Bump to Cape Hatteras .
Type 2-Deflection with apparent cyclonic rotation . The western boundary is
deflected seaward of the Bump and then appears to rotate in a cyclonic sense .
The cyclonic rotation is inferred through the intrusion of Gulf Stream water onto
the shelf . Waves were not detected downstream in this case either .
23
N
~
in the vicinity
Figure 2 .3-2 Types of perturbations observed on the western boundary of the Gulf Stream
of and downstream from the Charleston Bump (from Legeckis 1979) .
Type 3-Similar to case 2 type except the perturbation moved downstream . The
intrusion of warm Gulf Stream water on the shoreward side of the perturbation
extended for a considerable distance over the shelf (i .e ., as much as 200 km)
parallel to the Gulf Stream . This warm feature was later referred to as a warm
filament . The fourth and fifth types of perturbation in the schemes of Legeckis
(1979) are the more common forms of these perturbations found along the western
boundary of the Gulf Stream .
Legeckis
Type 4-Wavetrain of stable waves extending downstream from the Bump .
called these waves stable because they resembled a series of sinusoids and did
not display any apparent cyclonic rotation .
These
Type 5-Wavetrain of unstable waves extending downstream from the Bumy .
waves were called unstable because they had large peak-to-peak amplitudes and had
cyclonic rotation patterns associated with them, i .e, the intrusion of Gulf
Stream water onto the shelf on the shoreward side of the perturbations .
It was subsequently shown that these warm filaments (i .e ., the intrusions of Gulf
Stream water) are trapped at the shelf break, but at times they separate from the
perturbation and are left behind on the shelf (Vukovich and Crissman 1975) . In
either type four or five perturbations, it was often possible to detect a cold
core center in the trough of the wave . Satellite SST analyses have shown that
cold core centers or cold domes can be found in both stable and unstable waves
(Vukovich and Crissman 1980 ; Vukovich and Maul 1983) . Some of the Legeckis
(1979) data indicated that a mixture of stable and unstable waves existed at
times .
For the stable or unstable wavetrains, Legeckis (1979) found that there were from
two to six wave crests present at any one time and these appeared to move down
stream at speeds that varied from 20 to 60 km day-1 with an average speed of 40
km day-1 . These waves had periods from 4 to 5 days . Wavelengths ranged from 90
km to 260 km with an average wavelength of 150 km . The wave amplitudes were as
large as 100 km, though the unstable waves had large amplitudes . Similar
statistics for these perturbations were found by Vukovich and Crissman (1980)
using NOAA-5 Very High Resolution Radiometer (VHRR) data for the period January
through May 1977 .
Maul et al . (1978) used about three years (1976-1978) of Gulf Stream frontal
analyses from Geostationary Operational Environmental Satellite (GOES) IR data
in a randomly spaced time series to study Gulf Stream meanders . Using least
squares spectral analysis, they found that the dominant periods were 30 and 6
days in the offing of Onslow Bay . They gave no explanation for the 30 day
period, but did note that the 6 day period, and a weaker response at periods of
8-9 days and at 4-5 days, were probably associated with the Gulf Stream western
boundary features observed by Legeckis (1975), Rao et al .(1971) , Maul and Hansen
(1972), DeRycke and Rao (1973), Stumpf and Rao (1975), and Vukovich and Crissman
(1975) .
The surface and subsurface temperature and velocity of meanders, eddies and
filaments along the Carolina Cape is well illustrated in Figure 2 .3-3a, from Bane
et al . (1981) . On the basis of several airborne expendable bathythermograph
(AXBT) surveys and simultaneous moored current meter observations, they were able
to produce this extensive, oblique view of the thermal structure . It resembles
25
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Figure 2 .3-3 Structure of meanders, eddies and filaments along the Carolina Capes (a) thermal structure (from
Bane et al . 1981), (b) schematic of thermal and velocity structure (from Lee et al . 1981) .
the structure of Gulf Stream "frontal eddies" that Lee et al . (1981) summarized
off Georgia, as shown schematically in Figure 2 .3-3b .
The Frontal Eddy Dynamics (FRED) study (FRED Group 1989) showed that frontal
eddies often decay near Cape Hatteras to the point where the cold dome is not
visible in the imagery or in drifter tracks . However, a drifter in one of these
cold dome features showed that the eddy can reform north of Cape Hatteras and
distinct cyclonic circulations reappear . The FRED study also showed that the
decay of frontal eddies in Raleigh Bay was closely related to the mean distance
of the Gulf Stream front to the shelf break . With the mean front displaced
seaward of the shelf break in Raleigh Bay, larger meanders and frontal eddies
were present and able to propagate past Hatteras with little decrease in
amplitude . In one case, a large meander crest and filament came within a few
kilometers of the shore at Cape Hatteras (FRED Group 1989) .
The perturbations that develop on the western boundary of the Gulf Stream just
downstream (north) of the Charleston Bump have an average wavelength of about 150
km, amplitude of about 30 km, phase speed of about 30 km day-1, and periods in
the band from 4 to 10 days . However, as the perturbations move further
downstream and approach Cape Hatteras, average wavelength decreases to about 115
km and amplitude decreases to about 20 to 10 km while average phase speed
increases to about 40 km day-1 and period of 2 to 6 days) .
2 .3 .2 Meanders, Eddies, and Filaments at and Downstream of Cape Hatteras
Eddies and filaments of the Gulf Stream front have similar forms in Raleigh Bay
and just northeast of Cape Hatteras . However, Gulf Stream path displacements are
at a minimum at the latitude of Cape Hatteras, and a statistical study (SAIC
1990) of the Gulf Stream front in the vicinity of Cape Hatteras showed little
correlation between the direction of the path and meander characteristics
upstream and downstream of this position . Studies of Gulf Stream current
structure near Cape Hatteras from the Middle Atlantic Slope and Rise Experiment
(MASAR) (SAIC 1987 ; Churchill et al . 1989) and Mobil studies (SAIC 1990) showed
that vertical shear of the current in the cyclonic side of the front was closely
related to degree of eastward deflection of front from Cape Hatteras .
Meanders grow again rapidly to the northeast of Cape Hatteras as the current
leaves the continental margin as a free jet and flows into deep water . These
meanders tend to have considerably longer wavelengths, slower propagation speeds,
longer periods, and can grow to much larger amplitudes than those found upstream
of Cape Hatteras . Early surveys, such as Fuglister and Worthington (1951),
examined large deep troughs that formed to the east of 65°W and discovered that
they could break off into Gulf Stream rings which persisted and propagated
separately from the Gulf Stream itself . While ring formation occurs well east
of the study area, the process is summarized because the rings can translate
westward back into the Cape Hatteras region and cause major perturbations in the
currents and fronts there .
Hansen (1970) reported on a year-long sequence of monthly surveys of the Gulf
Stream path from about 100 km northeast of Cape Hatteras to about 55°W . These
surveys tracked the location of the 15°C isotherm at 200 m depth . This position
was called the "north wall," and chosen for relative convenience of surveying and
because it corresponded to a sharp temperature front at a depth below most of the
seasonal cycle of warming and cooling . Hansen (1970) estimated that the
27
perturbations on the Gulf Stream's boundary downstream from Cape Hatteras had
wave lengths of about 200-400 km and propagated eastward with wave speeds of
about 5-10 km day-1 . It was characteristic of the several path-surveys to find
elongated loops, particularly in the neighborhood of the New England Seamount
Chain, near 60°-65°W .
As was illustrated earlier in Figure 2 .2-1, from a combination of more recent
satellite remote sensing imagery and in situ observations, the lateral
displacement amplitude grows rapidly downstream from Cape Hatteras . Most of the
variance shown there is due to meandering (although long-period shifts in Gulf
Stream position may also play an important role, as treated in Section 2 .2) .
Watts and Johns (1982) used moored inverted echo sounders to track the path of
the Gulf Stream thermal front in the region 100 to 200 km beyond Cape Hatteras,
where the water depth under the Gulf Stream is approximately 3000 m . They found
the variance doubled in each 50 km step downstream, as the r .m .s . amplitude
increased 15 to 20 to 30 km . They found periodicities from 2-60 days,
wavelengths of 150-600 km, and propagation speeds of 18-36 km day"1 to the
northeast .
Halliwell and Mooers (1979) analyzed satellite imagery for three years, mainly
focused on this large-meander region further northeast of Cape Hatteras . Their
observations showed the path lateral-shift variance to continue increasing to an
r .m .s . value of over 80 km near 65°W . The meander envelope (twice the local
amplitude) became 200-300 km wide, and the waves in this downstream region became
very steep, i .e ., the amplitude was frequently greater than the wavelength .
Halliwell and Mooers (1979) also provided a statistical summary of ring motion .
Watts (1983) summarized the results of several earlier studies of meanders, both
upstream and downstream of Cape Hatteras, regarding characteristic wavelengths,
periods, propagation speeds, and amplitudes . The tendency is clear that a new
mode of instability grows in the deep water downstream of Cape Hatteras, with
longer wavelengths, slower phase speeds, and larger amplitudes, as just
described . These meanders eventually either mask or may be triggered by
fluctuations that propagate into the region from upstream of Cape Hatteras .
However, near Cape Hatteras itself, the meanders have not yet had much time or
space to grow, and consequently, the variability there has a wider range of
wavelengths and periods and still has much in common with the variability seen
upstream off the South Atlantic Bight .
Tracey and Watts (1986) refined and extended the information on Gulf Stream
propagation and growth characteristics just downstream of Cape Hatteras . They
used three years of moored inverted echo sounder data to show that downstream
propagation rates increase smoothly from about 14 km day-1 for meanders with
periods and wavelengths of 33 days and 460 km, to speeds exceeding 45 km day-1
for 4 day, 180 km meanders . The most rapid growth of meanders occurred in two
bands, one near 4-5 days and 180-230 km, and the other near 10-33 days and 300500 km wavelengths .
2 .3 .3 Gulf Stream Related Shelf Features from Satellite Observations
Many of the Gulf Stream frontal eddies have warm filaments associated with them
as shown schematically in Figure 2 .3-1 . These warm filaments are usually found
on the shoreward side of the meander near the shelf break . However, very often
28
these filaments penetrate onto the shelf and affect the shelf circulation and
transport shoreward of mass across the shelf . Three years (1984-1987) of NOAA
IR imagery were used to document the depth of penetration of these filaments on
the shelf in Onslow and Raleigh Bays (FRED Group 1989) . The data showed that
warm water associated with the filaments normally penetrated to the 20 m isobath
in both bays . In Onslow Bay, the 20 m isobath is located about 30 km from the
coast . In Raleigh Bay, it is about 10 km from the coast . On a monthly basis,
warm water intrusions into Onslow Bay only occurred in April, September, October,
November, and December ; whereas in Raleigh Bay, significant intrusions occurred
each month . Occasionally these intrusions detached from the Gulf Stream meander
and were left on the shelf, particularly in Raleigh Bay (Vukovich and Crissman
1975 ; FRED Group 1989) . Vukovich and Crissman (1975) found that the circulation
associated with a remnant warm filament enhanced entrainment of cold, low
salinity water from the shelf into the Gulf Stream . These studies suggest that
the life span of these features was from three to six days . However, Tester et
al . (1991) tracked a similar warm water intrusion on the shelf in Raleigh Bay for
a period of more than 19 days (in late October early November 1987) . They
estimated that this feature was responsible for a "red tide" outbreak on the
North Carolina shelf as the result of Loop Current/Florida Current/Gulf Stream
system transport of the toxic donoglagellate (Gymnodinium breve) to the North
Carolina region .
2 .4
Gulf Stream Rings and their Interaction with the Gulf Stream
2 .4 .1 Warm Core Rings
Brown et al . (1986) provide a historical review of ring processes which is not
repeated here . Warm-core rings (WCRs) form in the Gulf Stream system well
downstream of Cape Hatteras when an extended anticyclonic meander of the Gulf
Stream traps a mass of Sargasso Sea water and separates from the Gulf Stream .
This process, shown schematically in Figure 2 .4-1, injects an anticyclonic eddy
of warmer Sargasso Sea water into the Slope Sea waters north of the Gulf Stream .
Gulf Stream ring formation processes are shown schematically in Figure 2 .4-1 .
Auer (1987) found that 94 percent of the WCRs during a five year period were
formed north of the mean position of the Gulf Stream Landward Surface Edge, while
six percent (4 rings) were south of the mean position . WCRs form most often in
the portion of the Gulf Stream between 65°W and 62°W, the area of the New England
Seamounts (Auer 1987 ; Brown et al . 1986) . No rings were observed in a five year
period to form west of 70°W (Auer 1987) .
After formation, WCRs move westward or southwestward (Richardson et al . 1978) in
the Slope Sea at average speeds of 7 km day-' (Auer 1987) to 56 km day-' (Brown
et al . 1986) ; however, their westward motion may be disrupted by interaction
with bottom canyons (Evans et al . 1985), other rings or the Gulf Stream . When
the rings near the southern end of the Slope Sea, they accelerate as their path
is progressively pinched between the 200 m isobath and the north wall of the Gulf
Stream (Evans et al . 1985) . Other factors that will influence the motion of the
rings are B-induced motion (motion induced by change of the Coriolis parameter
with latitude), advection of the rings by the larger-scale Slope Sea circulation
(e .g ., the Slope Sea gyre), interaction with the Gulf Stream, and interaction
with other rings . Brown et al . (1986) found that the range of the speeds for the
WCRs they investigated was from about 2 km day'1 to 14 km day-' with those WCRs
that were long-lived having a range of about 5 km day'1 to 10 km day'1 . The
short-lived WCRs had larger variations for speed of motion than the long-lived
29
Cool
Slope
Water
as-1,
w0
`l
lo~
T i me
Figure 2 .4-1 Idealized views of the formation of a cold core ring (top) and a warm core ring (bottom) .
As the Gulf Stream loops well to the south (north), a region of cyclonic (anticyclonic) flow
is established, which oftentimes detaches from the jet to become a cold core (warm core)
ring (from SAIC 1991) .
WCRs (i .e ., the standard deviation of the speed of motion for the short-lived WCR
was about 2 .5 km day-1, and that for the long-lived WCR, about 1 .2 km day-1) .
Auer (1987) suggested that about 22 WCRs are formed and 22 are absorbed each
year ; whereas Brown et al . (1986) found that the mean number of rings existing
in any year was eight . Vukovich's (1990) data suggest that at least 16 to 18
separate rings were found each year west of 65°W in the period 1980-1988 . No
explanation for the differences during the three surveys can be given at this
time . The diameter of a typical WCR at formation is, on the average, about 150
km . For 71 WCRs investigated by Auer (1987), the diameter of the ring decayed
at a rate of about -0 .0037 day-1 in the period between formation and absorption
by the Gulf Stream . Brown et al . (1986) found that the decay rate was about
-0 .0043 day-1 over the same period . Auer found that the decay rate was much
larger during the first ten weeks of the existence of the ring ; i .e ., -0 .021
day-1 . The mean diameter of the ring at the time of absorption was about 70 km .
The distributions of the WCR lifespans is bi-modal (Brown et al . 1986) with a
mean of about 130 days . Most WCR lifespans are less than 90 days . The highest
frequency of WCR lifespans were from the 7 to 30 day band with a secondary
frequency peak in the 160 to 195 day band . Auer (1987) found that the maximum
lifespan for a WCR was about 450 days, whereas Brown et al . (1986) found a
maximum of about 400 days .
Just over 80 percent of the WCRs in the Slope Sea were absorbed directly by the
Gulf Stream while the remainder were absorbed by other WCRs (Auer 1987) . Thus
ultimately all were absorbed by the Gulf Stream . The highest frequency of WCR
absorptions by the Gulf Stream was found in the longitude band 75°W to 72 .6°W
(i .e ., just northeast of Cape Hatteras), with a secondary maximum in the band
60°W to 57 .6°W . The spatial frequency of ring absorptions in between the primary
and secondary maximum bands (i .e ., the major portion of the Slope Sea) was
uniform except for a minimum in the band 71°W to 67°W . The short-lived WCRs were
absorbed by the Gulf Stream in the band 58°W to 65°W, which is the same region
in which they were most frequently formed . The longer-lived WCRs, on the other
hand, were absorbed primarily in the band 72°W to 75°W . The data showed that the
track of the short-lived WCRs was most often located near the climatological mean
position of the Gulf Stream's north wall east of 66°W, making them more available
for immediate absorption . The longer-lived WCRs tended to be located further
away from the climatological north wall position (Brown et al . 1986) .
Though absorption by the Gulf Stream is the principal means for the removal of
a WCR from the Slope Sea region, some rings merge, producing one larger ring .
Ring absorbing ring occurred in 18 percent of the cases studied by Auer (1987) .
Evans et al . (1985), reported some interesting interactions that occurred in the
Slope Sea region involving rings, Gulf Stream water, slope water, and shelf
water . When a WCR is situated close to another small, usually cyclonic, eddy,
the strong horizontal shear that exists due to the juxtaposition of the two
eddies enhances lateral fluid transport . Streamers of colder water from the
shelf or slope were observed in association with, and sandwiched between, the two
rings . With a cyclonic vortex to the northeast of a WCR, the advection of cold,
fresh shelf water into the slope was enhanced . When a cyclonic vortex was
located in the southwest quadrant, there was intense advection of Gulf Stream
water into the slope . It was also observed that as WCRs moved toward shallower
topography, cyclonic eddies were formed in the lee of the WCR . These cyclonic
eddies very often formed to the northeast of the WCRs as they encountered canyons
31
or approached Cape Hatteras, increasing the likelihood of advection of shelf
water into the slope .
2 .4 .2 Cold Core Rings
Cold-core rings (CCR) also form most often in the vicinity of the New England Sea
Mounts . After formation (see Figure 2 .4-1), CCRs are normally injected into the
Sargasso Sea, south of the Gulf Stream in the region east of 73°W . Richardson
(1980), for example, offers a schematic of ring formation . No CCRs that were
formed as a result of a Gulf Stream meander were observed to form north of the
Gulf Stream in the Slope Sea (Auer 1987) . They consist of a cold central core
of low salinity slope water surrounded by a ring of warm, high salinity Gulf
Stream water . Auer (1987) indicated that about 35 CCRs are found, on the
average, in the Sargasso Sea on a yearly basis . Since the lifespan of some CCRs
can exceed one year (The Ring Group 1981), the 35 CCRs do not necessarily reflect
the number of new rings that form each year . Furthermore, CCRs are difficult to
detect using satellite imagery because the SST gradients associated with CCRs are
weak (Vukovich 1976 ; Vukovich and Crissman 1978) . The cold dense water associated
with the core of the ring, at times, sinks, wiping out the SST anomaly (Auer
1987) . CCRs are best detected in satellite IR imagery in the fall through spring
when the SST contrasts are greater and when an interaction between a CCR and the
Gulf Stream may produce a tongue of warm water that encircles the ring around its
eastern and southern sides (Vukovich 1976 ; Vukovich and Crissman 1978) . Because
CCRs are sometimes hard to detect in the satellite SST imagery, the yearly
average of 35 CCRs in the Sargasso Sea suggested by Auer (1987) may have included
the effect of counting a particular ring more than once in a year during that
ring's life history .
After formation, the rings move in a westward and/or southwestward direction
(Richardson and Knauss 1972, Richardson 1980, Auer 1987) . Richardson (1980)
reported on a CCR which became attached to the Gulf Stream off Florida and was
advected to the northeast . This ring collided just southeast of Cape Hatteras
with another CCR, which subsequently coalesced with the Gulf Stream . Auer (1987)
found that the mean CCR diameter was 105 km, which is smaller than the 250 km
estimate of Lai and Richardson (1977) . Auer attributed the difference to
measurement methods . Vukovich (1976) found that the rings were elliptically
shaped, having a semi-major axis which ranged from 60 km to 90 km and a
semi-minor axis which ranged from 50 km to 60 km . The rings move erratically
with translation speeds of the order of 7 km day-1 . However, over longer periods
(of the order of ring lifetimes) the translation velocity averaged 1 km day-1 in
a southwesterly direction (Auer 1987) .
The Mobil study (SAIC 1991) reported a case of a cold core ring impacting the
eastern side of the Gulf Stream just southeast of Cape Hatteras in January 1990
(Figure 2 .4-2) . The eddy caused a large amplitude meander to form and the CCR
propagated with Gulf Stream toward the northeast past Cape Hatteras, causing the
Gulf Stream to penetrate to greater than 400 m depth over the slope at the Manteo
site . Richardson (1980) provides a schematic of a similar Gulf Stream-CCR
interaction . There is some evidence (Bane et al . 1988 ; SAIC, 1991) that
interaction of a CCR with the Gulf Stream just north of Cape Hatteras can cause
shifts in the mean position of the Gulf Stream front in the Middle Atlantic Bight
that persist for many months .
32
Figure 2 .4-2 AVHRR images of CCR event in January 1990 . (a) January 16,
(b) January 17, (c) January 22, and (d) January 27 (from SAIC
1991) .
-
33
Vukovich (1976) reports satellite observations of cases when a Gulf Stream
frontal eddy was located at the same latitude as a CCR on the eastern side of the
Gulf Stream . In this case the amplitude of the Gulf Stream frontal eddy
increased in the seaward direction . It would appear that the interaction of the
CCR with the Gulf Stream locally deflected the Gulf Stream seaward .
34
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35
Halliwell, G .R ., Jr . and C .N .K . Mooers . 1979 . The space-time structure and
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37
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38
III . THE SLOPE SEA
3 .1
Introduction
The Slope Sea has been defined (Csanady and Hamilton 1988) as the area of the
ocean between the shelf-break of the Middle Atlantic Bight and the Gulf Stream
(Figure 3 .1-1) . These waters overlie the continental slope and upper continental
rise with depths ranging from a few hundred meters to over 4000 m . The southern
Slope Sea boundary is the dynamic Gulf Stream front ; thus the area of the Slope
Sea changes with changes in the Gulf Stream path . When the Gulf Stream is close
to the slope at Cape Hatteras and separates from the slope farther north, Slope
Sea water and circulations in the upper part of the water column, with origins
to the northeast, may not reach the Cape Hatteras region . Also in the southern
part of the Middle Atlantic Bight where the Gulf Stream is close to the shelf
break, Gulf Stream water and Gulf Stream related flows can impinge directly on
the upper slope and outer shelf and thus "short-circuit" the Slope Sea
circulation . This is quite different from the Slope Sea further to the east
where the Gulf Stream is often several hundred kilometers from the shelf break .
In winter, the northern boundary of the Slope Sea at the shelf break is the
location of the shelf-slope front . This front divides the distinctly different
water masses of the shelf and slope . In summer, the shelf-slope front is
restricted to the lower half of the water column . Waters above the seasonal
thermocline (-20 m depth) show little contrast between slope and shelf, unless
a warm core ring is present in the Slope Sea .
The general circulation of the whole water column of much of the Slope Sea is a
southwestward drift . In the upper several hundred meters, this flow returns
along the outer edge of the Gulf Stream front, forming a gyre (Csanady and
Hamilton, 1988) . The return flow includes water entrained from the Middle
Atlantic Bight shelf, the so called "Ford water" (Ford et al . 1952) . In the
lower part of the water column, below about 1000 m, is the southwestward flow of
the various components of the DWBC . The deepest portions descend and cross
beneath the Gulf Stream at Cape Hatteras and then continue southward along the
Blake Escarpment (Pickart and Watts 1990 ; Richardson 1977) .
The upper water column is often disturbed by warm core rings, which also
propagate to the southwest in the Slope Sea before being reabsorbed by the Gulf
Stream, often in the band 72 .6° to 75°W (Auer 1987) . Warm core rings can induce
exchanges with both the Gulf Stream and the shelf as they move southwestward
along slope . In the lower part of the water column, below 1000 m, energetic
disturbances are found with periods of a week or longer . These waves, known as
topographic Rossby waves (TRWs) (Thompson 1977, Rhines 1970), are thought to be
generated by meanderings of the deep Gulf Stream or during the formation of warm
core rings . These waves propagate westward along the isobaths and towards the
slope, where they are refracted by the steep topography (Hogg 1981 ; Shaw and
Csanady 1988) .
The following sections discuss Gulf Stream related exchange processes in the
vicinity of Cape Hatteras, followed by the upper layer Slope Sea circulations,
and conclude with deep circulations .
39
.
/
,~...
/ SHELF EOGE
~ ._ .J
~
/
_.,----_---,-:
/
Figure 3 .1-1 Schematic Slope Sea Circulation (redrawn from Csanady and
Hamilton 1988) .
40
3 .2 Near-Su rface Slove Sea Waters
Drawing from previously published studies and available data, Csanady and
Hamilton (1988) offered the most complete description to date of properties and
motions within the Slope Sea . Their analysis indicated that the surface few
hundred meters of the Slope Sea is occupied mostly by a distinct 'slope water'
mass, which McLellan (1957) found to be a mix of roughly 20% Coastal Labrador Sea
Water and 80% North Atlantic Central Water . They further noted slope water is
often intruded upon by other water masses, the largest being Gulf Stream warmcore rings . Based on current meter and hydrographic data, Csanady and Hamilton
(1988) proposed a circulation scheme of slope water which includes a cyclonic
gyre within the Middle Atlantic Bight (see Figure 3 .1-1) . The essential features
of the gyre were later observed by drifter tracks (Battelle, 1991) . Recent
studies have shown that the portion of the Slope Sea in our region of interest,
south of 37°N, differs in many ways from the view offered by Csanady and Hamilton
(1988) . As discussed below, the upper Slope Sea is unique in this region due to
the influence of water locally discharged from the Gulf Stream and the diminished
effect of warm-core rings .
3 .2 .1 Influence of Gulf Stream Warm-Core Rings
As noted in Chapter 2, warm-core rings tend to migrate along the continental
margin in the direction of Cape Hatteras . Rings have been shown to have a
tremendous influence on the shelf-edge and upper slope environment . They can
produce strong near-bottom currents over the upper slope (Butman 1987), initiate
exchange of shelf and slope water (Morgan and Bishop 1977 ; Bisagni 1983 ;
Churchill et al . 1986 ; Garfield and Evans 1987 ; Joyce et al ., 1992), and generate
wave motions at the shelf-edge (Ramp 1989) . Although rings have not been
extensively surveyed in the study area, available information suggests that their
direct impact here should be much less than over areas to the north . Brown et
al . (1986) used SST images derived from satellite radiometers to examine ring
behavior over a 10-year period . Their findings indicate that a large fraction
of rings become assimilated by the Gulf Stream before reaching the vicinity of
Cape Hatteras . Their analysis also reveals that those rings which do enter this
region tend to be relatively "old" . Auer (1987) noted that 12 of 60 WCRs
absorbed by the Gulf Stream over a five year period were absorbed in the band
72 .6° to 75°W . Case studies of ring 82B (Joyce and Kennelly 1985 ; Olson et al .
1985) indicate that older rings have considerably smaller velocities and
dimensions than their younger counterparts .
3 .2 .2 Effect of Locally Discharged Gulf Stream Water .
Churchill and Cornillon (1991a) found that water properties measured over a two
and a half year-period in the Slope Sea offshore of the Chesapeake Bay mouth were
dominated by fluid locally discharged from the Gulf Stream . From examination of
SST images, they determined that this water was typically ejected from the
trailing edge of a Gulf Stream meander . The discharged water seen in the images
took on two forms, both markedly different from a warm core ring . One form
closely resembled that of the warm filaments associated with Gulf Stream frontal
eddies commonly seen south of Cape Hatteras . It was distinguished by elongated
shape, strong currents (order 60 cm s-1) directed to the southwest and . a
cyclonically curling tail enclosing relatively cool water . The other form was
markedly different in that it covered a relatively broad area (sometimes filling
the entire Slope Sea south of 38°N), and had relatively weak currents (generally
41
<40 cm s-1) . Parcels taking on the latter form were more frequently observed .
These did not appear to propagate northward, as frontal eddies customarily do,
and were remarkably long-lived, sometimes remaining visible in the SST field for
several weeks . Continuity of potential temperature-salinity properties and
nutrient concentrations along density surfaces indicate that these near•surface
parcels originated within deeper, nutrient-rich layers of the Gulf Stream and had
upwelled before crossing the Gulf Stream front . Their appearance tended to
enhance nutrient concentrations over the upper slope, significantly so near the
base of the euphotic zone .
Hydrographic data and moored current meter measurements reveal that a baroclinic
eastward current, with a magnitude of roughly 30 cm s-i, typically flows at the
northern margin of the large-area parcels of discharged Gulf Stream water
described above (Churchill and Cornillon 1991a,b) . When the parcel extends from
the edge of the Gulf Stream to the shelfbreak (as is common), this current
essentially becomes the southern terminus of the cyclonic slope water gyre
described by Csanady and Hamilton (1988), connecting the southwestward slope
water flow over the slope with the return northeastward flow adjacent to the Gulf
Stream . This current also frequently carries shelf water seaward of the shelfslope front, as evidenced by numerous SST distributions showing a band of surface
shelf water at the northern margin of a discharged Gulf Stream water mass (e .g .
Plate 1 of Churchill et al . 1989 and Figures 1 and 6 of Churchill and Cornillon
1991b) . The transport of shelf water in such bands is poorly resolved by the
data examined thus far . A pair of rough estimates, obtained from hydrographic
and current meter data, put it at approximately 0 .1 Sv (Churchill and Cornillon
1991b ; Churchill et al . 1992), comparable with the estimated rate at which warm
core rings entrain shelf water (Morgan and Bishop 1977 ; Bisagni 1983 ; Churchill
et al . 1986 ; Garfield and Evans 1987 ; Joyce et al . 1992) .
The rate at which discharged Gulf Stream water reaches the continental margin and
potentially causes shelf water export was estimated by Churchill and Cornillon
(1991b) using SST images and data from 51 hydrographic surveys carried out over
a 12-year period as part of the U .S . National Marine Fisheries Service Marine
Mapping Program (MARMAP) project . They found that locally expelled Gulf Stream
water impinged on the continental margin between 36° and 37°N roughly 25% of the
time during the MARMAP project (Figure 3 .2-1) . Gulf Stream water incursions onto
the Middle Atlantic Bight shelf were also observed . They were most frequently
seen, 25% of the time, at a MARMAP station east of Pamlico Sound and were
observed 3-9% of the time at outer shelf stations further north .
Intrusions of discharged Gulf Stream water onto the southern Middle Atlantic
Bight shelf have been examined by Gawarkiewicz et al . (1990) and Churchill and
Cornillon (1991b) . Gawarkiewicz et al . (1990) described an intrusion which
appeared over the shelf east of Virginia . It extended as far shoreward as the
35 m isobath and was mostly confined within the seasonal pycnocline . Churchill
and Cornillon (1991b) also observed Gulf Stream water intrusions within the shelf
pycnocline as well as ones which extended over the entire water column and were
confined to surface and bottom layers . They found that the intruding water did
not significantly influence local currents or the alongshelf density field over
the shelf . Its subtidal flow was largely driven by the alongshelf wind stress
component, and its vertical density profile nearly matched the density profile
of adjacent shelf water (indicating negligible adjustment of the alongshelf
density field in response to the intrusion of Gulf Stream water) . In spite of
42
(a)
39°N
+41 aa
4-29
+45
+40
+30
-f• 31 39
+28
+23
J+22
~-21
37°N
I.,
w
20
12 13 ~ .14
+ 11
+10
+3
+6
T+
+32
+27
+33
(b)
3
47
+3
+ 46
3+
+
38
+3
3T
34
2a + 28
33
3+
25
+9
13
6+
+3
+
9+
9
A
3 ~P
MARMAP
Stations
~
76°W
8
8+ 22
17
+4
35°N
8
+
8+
-~9
~18
8
5
74°W
72°W
76°W
+ 27
25
% of Casts
op
~
Gulf Stream Water
Showing
74°W
72°W
Figure 3 .2-1 MARMAP study (a) hydrographic stations, and (b) stations at which Gulf Stream water was
detected and the percentage of casts taken at these stations which intercepted Gulf Stream
water (from Churchill and Cornillon 1991) .
this, it is unclear whether wind-induced circulation played a major role in
transporting the observed Gulf Stream water intrusions onto the shelf .
3 .3
Gulf Stream Related Current Variability
The upper few hundred meters of the Slope Sea have distinct water properties
derived from mixtures of sub-thermocline Sargasso Sea water advecting along
isopycnals from under the Gulf Stream and Labrador Sea Water input west of the
Grand Banks . There is also a strong seasonal signal resulting from intense
cooling in winter, which results in the formation by overturning of a 200 m deep
mixed layer in January (Csanady and Hamilton 1988) . This characteristic, wellmixed pycnostad (-12°C and -35 .4 psu), becomes capped with a seasonal thermocline
and is eroded from below as spring and summer progress . Slope water can also be
overrun by extrusions or overwashes of warm water from the Gulf Stream and
extrusions of shelf water . Some of these types of events can be caused by
interactions of warm core rings with the Gulf Stream and the shelf-slope front .
Csanady and Hamilton (1988) postulated a cyclonic gyre between the Gulf Stream
front and the shelf break in the southern part of the Slope Sea between Nantucket
and Cape Hatteras . The strongest southwestward part of the flows along the slope
occur off New Jersey in the vicinity of the Middle Atlantic Slope and Rise
(MASAR) study northern transect and the 106-Mile Site . The southern extent of
the gyre appears to be highly variable depending on the configuration of the Gulf
Stream front and the presence of overwashes of Gulf Stream water in the narrow
region between the front and the shelf break south of the mouth of the Chesapeake
Bay .
The return flow of the gyre along the Gulf Stream front has been shown to occur
from satellite tracked drifters released at the 106-Mile Site at approximately
weekly intervals over an eighteen month period (Battelle 1991) . The 106-Mile Site
is a 36 .7 kms by 7 .2 km rectangle oriented north-south with its northeast corner
at 39°N, 72°W . The majority of these drifters initially moved southwestward,
turning somewhat north of Cape Hatteras and returning along the Gulf Stream front
but not crossing over the front and becoming part of the Gulf Stream flow proper .
About 21% of the drifters recirculated, indicating that the eastern end of the
gyre is somewhat diffuse and perhaps intermittent .
Bane et al . (1988) showed that the strength of the southwestward flow through the
northern MASAR transect was closely related to the distance of the Gulf Stream
front from the shelf break such that, on a monthly basis, the closer the front
to the shelf the stronger the southwestward flow over the continental slope .
This could be attributed to either an increase in the intensity of the gyre with
a reduction of surface area of the Slope Sea or a shifting of the position of
maximum flow towards the slope . The MASAR study also showed that for periods
with no rings present and characteristic slope water present off Virginia and New
Jersey, the currents at both transects were quite coherent over a distance of
about 200 km (Csanady and Hamilton 1988) . Upper layer flows were also quite
slab-like when warm core rings were absent . This was confirmed later by EPA
measurements higher in the water column (Battelle 1987) . Warm core rings disrupt
this basic gyre flow, destroying the longshore coherence of the flows as well as
inducing exchanges with the Gulf Stream and the shelf .
44
3 .3 .1 Gulf Stream Entrainment of Middle Atlantic Bight Shelf Water
Shelf water in the Middle Atlantic Bight drifts along isobaths toward Cape
Hatteras at roughly 5 cm s-1 (Boicourt 1973 ; Beardsley et al . 1977 ; Beardsley and
Boicourt 1981) . Shelf water which does not fall victim to export while in
transit to Cape Hatteras ultimately becomes entrained into the Gulf Stream
current . Ford et al . (1952) were the first to recognize the presence of shelf
water adjacent to the Gulf Stream . They found filaments of this water stretching
more than 2000 km along the Gulf Stream's northern edge . These were not
continuous, suggesting that the Gulf Stream draws water from the Middle Atlantic
Bight shelf intermittently . Later studies (Kupferman and Garfield 1977 ;
Churchill et al . 1989 ; Lillibridge et al . 1990) have revealed that the
entrainment of shelf water by the Gulf Stream is a highly variable phenomenon .
Using airborne radiometer measurements, Fisher (1972) confirmed that Gulf Stream
entrainment of surface shelf water is intermittent . Images of SST from satellite
radiometers further show that the location of this entrainment varies
significantly, occurring as far north as 38°N (Churchill et al . 1989) .
Hydrographic and moored instrument data reveal appreciable variation in the
dimensions of shelf water filaments at the Gulf Stream's edge (Fisher 1972 ;
Kupferman and Garfield 1977 ; Churchill et al . 1989 ; Lillibridge et al . 1990) .
The velocities measured within these span a wide range, roughly from 20 to 100
cm s'1 (Ford et al . 1952 ; Kupferman and Garfield 1977 ; Churchill et al . 1989 ;
Lillibridge et al . 1990) . Consequently, estimates of shelf water transport
within individual filaments also vary greatly, from 10 to 500 x 103 m3 s'1
(Kupferman and Garfield 1977 ; Churchill et al . 1989 ; Lillibridge et al . 1990) .
3 .4
Deep Circulation off Cape Hatteras
3 .4 .1 Overview of the Deep Currents and Processes
In contrast to the Slope Sea to the north of Cape Hatteras, where the mean
southwestward flow along the continental margin typically extends throughout the
water column, an important crossing of currents occurs at Cape Hatteras . While
the near-surface waters above about 800 m turn northeastward along the northern
edge of the Gulf Stream as part of the Slope Sea circulation (Csanady and
Hamilton 1988), a major component of the deeper flow crosses under the Gulf
Stream as part of the DWBC . The DWBC descends and turns offshore as it
encounters the main baroclinic front of the Gulf Stream ; south of the crossover
it roughly follows the 1400 m isobath along the outer edge of the Blake Plateau
(Pickart and Watts 1990 ; Richardson 1977) .
Offshore of Cape Hatteras, the deep currents on the continental slope and rise
flow nearly along the bathymetric contours, with much weaker and variable
current components crossing up- or downslope . Stratification in the deep currents
tends to constrain strong vertical motion . The mean DWBC is southwestward at
typically 5 cm s'1 (ranging from 0 to about 10 cm s'1) . There are several sources
of variability which can produce currents up to 25 cm s-1 . These can accentuate
the mean currents or can be strong enough to reverse the deep currents, causing
northeastward flow for the periods of several days . The principal processes of
variability are TRWs, and there have been some observations of deep coupling to
upper water column fluctuations associated with Gulf Stream meanders or Warm Core
Rings (Hamilton 1987) . The DWBC transport (-5-14 Sv) has also been observed to
45
vary . The mean deep currents and processes causing their fluctuations are
reviewed in the following subsections .
3 .4 .2 The Deep Western Boundary Current near Cape Hatteras
The mean DWBC consists of several components at different temperatures flowing
equatorward along bathymetric contours . It is typically found within the bottom
300 m to 10000 m of the water column over bathymetric contours ranging from about
800 m to over 4000 m off Cape Hatteras ( Pickart and Watts 1990) . Hence it is
representative to picture the DWBC as a thin ribbon hugging the continental slope
and rise, since its sloping width may be over 200 km whereas its vertical
thickness may be less than 1000 m . The DWBC is formed mainly from northern
waters that are traditionally envisioned to include North Atlantic Deep Water
(NADW) and Labrador Sea Water ( LSW) . The Stommel and Arons ( 1960) model of
thermohaline circulation suggests that these northern source waters flow
equatorward in a continuous ribbon along the western margin of the North
Atlantic .
This traditional idealization of the DWBC constituents and flow must be modified
in several details to conform to recent observational studies (Watts 1991 ;
Pickart and Watts 1990) . A recent evaluation of a large number of deep current
and temperature measurements in the Middle Atlantic Bight by Watts (1991) has
shown that the DWBC carries a number of different water masses, that the current
structure is banded into higher and lower speeds associated with different water
masses and strata of potential temperature, and that the counterflow beneath the
Gulf Stream is as shallow as 800 m(hugging the continental slope) in the water
column .
The mean near-bottom currents within 50-300 m above the bottom along the
continental slope and rise approaching the Cape Hatteras region in the Middle
Atlantic Bight are shown in Figure 3 .4-1 from Watts (1991) . The uncertainty of
the mean component is indicated by the boxes around the arrowheads . It reveals
a very consistent southwestward current pattern along the bottom contours
throughout the bathymetry range 800-4000 m . For reference on this chart, the
surface path of the "north wall" of the Gulf Stream and its meander envelope are
indicated respectively as the solid and two dashed lines arcing eastnortheastward from Cape Hatteras . The mean deep currents at sites that are more
than about 30 km offshore of the Gulf Stream mean path and in water deeper than
about 3600 m appear to be more variable in direction, probably due to eddy
variability driven by the overlying Gulf Stream meanders . The manner of this
coupling is, however, not yet well understood .
The deep currents were also shown in Watts (1991) to be banded in structure when
categorized by potential temperature, as in Figure 3 .4-2, and when categorized
according to their height above bathymetric contours, as in Figure 3 .4-3 . The
alongshore speed structure in each depiction has three peaks around 6-7 cro s-l,
separated by broader low-speed bands around 3 cm s-1 .
In particular, the DWBC transports waters with a range of potential temperatures
that span from 6 .0°C to 1 .8°C . The water types within the 6 .0°C to 4 .0°C range
(found at depths in the Middle Atlantic Bight between about 800 m to 1100 m) are
found to be Sub-Polar Mode Waters (SPMW) or Sub-Arctic Intermediate Water (SAIW),
probably originating in the southern Labrador Basin or Newfoundland Basin, and
they exhibit relatively swift equatorward flow (6 to 10 cm s-1) . Potential
46
770
750
730
71°
69°
200
40°
° a a
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p
3pp
~
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.
.
.
36°
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.
,
.
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r-I c°
670
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r
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L_J
CURRENTS 50-300m
OFF BOTTOM
d
l
0
I
10cm/s
34°
Figure =
a(from Watts 1991) . Speed key is given
ie mean in the u, v directions .
11 ,
I
10
..
:
E
v
0
W
9
8
'
z
'
W
a 6
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ao
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41
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.
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2 .0
2 .0
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2
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1
SAI
2-S PMW/ -
2 .5
3 .0
3 .5
POTENTIAL TEMPERATURE (° C)
2 .5
3.0
3 .5
MEAN IN SITU TEMPERATURE (°C)
4 .0
4 .0
4 .5
Figure 3 .4-2 Mean equatorward alongshore current speed classified according to mean potential temperature, which is approximately adjusted from mean in-situ temperature for each record
according to its depth . Error bars are standard error of the mean (from Watts 1991) .
DWBC
MEAN EQUATORWARD SPEEDS
PROJECTED ONTO ONE SECTION
z
(m)
0
i -.*typical -= i
I Gulf Stream
north wall ~
I
.
~
15 cm/s
i
1 000
5-10 cm /s
.
.
.
~ 2-5 cm/s
.,
7
•
2000
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..
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. ~ highly variable •
I under Gulf Stream
3000
• j >5 ~ 0-4 cm /s
.
~
'
4000
0
100
200
300
X (km)
Figure 3 .4-3 Structure of the Deep Western Boundary Current, constructed by
proj ecting mean equatorward speeds onto a single section,
preserving potential temperature and position relative to
bathymetry (from Watts 1991) .
49
temperatures from 3 .9°C to 3 .1°C (found at depths in the Middle Atlantic Bight
between approximately 1200 m to 2100 m) would commonly be associated with LSW .
However, this stratum exhibited weak equatorward flow (3 cm s-1) ; and, based on
Freon tracer profiles, Pickart (1990) shows that there is a predominance of NADW
even in this temperature stratum, which is at least a decade old . LSW seems to
remain east of the tail of the Grand Banks and mainly flow into the eastern North
Atlantic (McCartney and Talley 1982) . However, small LSW anomalies may also be
observed sporadically in this region (SAIC/Battelle 1988) . Deeper and colder in
the NADW (down to potential temperatures of 2 .0°C) are a swifter stratum (6 cm
s-1 near 2 .9°C to 3 .1°C in depths of 2100 m to 2400 m) and a slower stratum (3
cm s-1 near 2 .0°C to 2 .8°C in depths of about 2400 m to 3100 m) . The coldest
waters (2 .0°C to 1 .8°C) are associated with Antarctic Bottom Water (AABW), which
also flows southwestward, having joined the DWBC somewhere to the northeast .
AABW appears in a relatively swift thin band (5 cm s-1) just at the juncture of
the continental slope and rise in depths around 3300 m . Separate hydrographic
evidence (not shown) suggests that this AABW lens within the DWBC is closed on
the offshore side as shown in Figure 3 .4-3 and separated from other AABW that may
be found in considerably greater depths farther offshore .
Pickart and Watts (1990) show that, as the DWBC crosses under the Gulf Stream at
Cape Hatteras, it shifts offshore to depths about 600 m deeper than it followed
before encountering the Gulf Stream . Figure 3 .4-4 shows the mean current vectors
100 m above the bottom on a line of current meters just off Cape Hatteras ; the
pattern of flow was parallel to bottom contours except for the middle site at
2800 m that was located just under the steepest mean position of the sloping
density front of the Gulf Stream . This result is dynamically consistent with the
conservation of potential vorticity within the lower layer .
To summarize these deep current measurements briefly, there are three depth
ranges in the Middle Atlantic Bight where higher mean speeds (>6 cros-1) are
found : above approximately 1000 m, between about 2000 and 2400 m, and below 3000
m . In the intervening depth ranges, speeds were more typically 2-4 cm s-1
.
Because the depths of the bands are more precisely related to potential
temperature and water mass structure, the alongshore mean speed structure that
was summarized above for the Middle Atlantic Bight may be shifted about 600 m
deeper in the DWBC south of Cape Hatteras .
These arguments also imply that there is little recirculation of the DWBC flow
back along the path of the Gulf Stream . The limited measurements under the Gulf
Stream in Figure 3 .4-1 do not show evidence of any organized flow . A recent
analysis of the high oxygen core of the DWBC by Pickart (1992) does indicate that
a small portion is recirculated to the north by the deep Gulf Stream at Cape
Hatteras . The high oxygen water is around 2 .3°C, and there is no quantitative
estimate of the amount of recirculation other than that it is probably "small" .
Altogether, Watts (1991) estimates the total transport of the DWBC in the Middle
Atlantic Bight to be 5-14 Sv . While the depth ranges and speeds of the DWBC are
relatively well established, the main uncertainty in estimating transports is the
range of estimated widths of the current . Watts (1991) estimates widths from
historical maps of water properties to be between 40 and 120 km .
Numerical models of the Gulf Stream, e .g ., Thompson and Schmitz (1989), have
suggested that the location where the Gulf Stream separates from the western
boundary may be sensitive to the transport of the DWBC .
50
3 8°
~r
3 6°
:, .,
. .;
. ~
~.. . .
:.
••...
.
:
:
.
y
•r .
..
;. s ® X
~
x
.3°C
2 .4° c,
34°
76°
~
2 .8 ° C 5
2.2 °C
740
c
m
/s
72°
Figure 3 .4-4 Mean deep current vectors 100 m above the bottom on a line off Cape Hatteras (from Pickart
and Watts 1990) . The bathymetric contours and the Gulf Stream thermocline depth topography
(as measured by Inverted Echo Sounders deployed during the same interval) are also
indicated .
3 .4 .3 Topographic Rossby Waves
The TRW is another important process that dominates the current fields below 1000
m . These are planetary waves with dominant periods of 10 to 100 days that are
characterized by columnar motions that are bottom intensified . The low frequency
currents have higher speeds by about 5% to 10% near the bottom than at 1000 m
(Luyten 1977 ; Hogg 1981 ; Hamilton 1984) . The waves tend to propagate westward
and southwestward along the isobaths and show evidence of refraction by the steep
continental slope (Hogg 1981 ; Shaw and Peng 1987 ; Pickart and Watts 1990a) . TRWs
are transverse waves with wavelengths of order 100 to 200 km . Since these are
linear waves, they do not generate a flux of solutes ; however, for settling
particles, the currents associated with TRWs will produce an effective dispersion
by displacing particle paths across the isobaths . TRWs are thought to be
generated by meanders of the deep Gulf Stream . Louis et al . (1982) successfully
explained bursts of TRW energy observed on the Scotian Rise as resulting from an
isolated vortex which was a model representation of a newly formed WCR .
Previous studies have documented energetic wave motions at 16 days (Thompson
1977), 32 days (Schultz 1987), 40 days (Pickart and Watts 1990a) and 25-40 days
(Hogg 1981) . Figure 3 .4-5 shows current principal axis variance ellipses for
TRWs of 40 day period off Cape Hatteras, from Pickart and Watts (1990) . There is
a sharp drop in energy at periods shorter than 10 days as a consequence of TRW
motions not being supported at these periods (the Rossby wave cut-off frequency Rhines 1970) . The major axes of the variance ellipses are at more of an
inclination to the isobaths, which is consistent with TRW theory in that current
fluctuations become perpendicular to the isobath direction as the wave frequency
approaches the cut-off (Rhines 1970) .
Analysis of current fluctuations at deep moorings shows that the currents are
highly coherent with depth at periods longer than 10 days . However, in the
cross-slope and along-slope directions, motions are not strongly coherent . This
is typical of TRWs and a consequence of the short length scales of the motions
(Hogg 1981 ; Hamilton 1984) .
3 .4 .4 Other Processes of Deep Variability near Cape Hatteras
Johns and Watts (1985, 1986) observed deep current variability off Cape Hatteras
that is coupled to lateral translations of the Gulf Stream, in addition to the
TRW variability summarized above . This vertically coherent coupling below 2000
m is only apparent at periods shorter than about 16 days (the longer
periodicities having sufficient TRW variability to mask this effect) . At periods
shorter than about 16 days, and just near the northern edge of the Gulf Stream,
the deep temperature front displacements and cross-stream velocity fluctuations
are coherent with the path displacements of the Gulf Stream . This verticallycoherent organized motion is found only under the northern edge of the current,
near the zone of maximum baroclinicity . Pickart and Watts (1990) show this
result to be dynamically consistent with water columns exhibiting rotational
resistance to vertical stretching, and find that the deep current cross-slope
angle of flow is influenced by the angle at which the Gulf Stream crosses the
bathymetry .
Pickart and Watts (1990) have also attempted to spatially filter out TRWs from
an array of deep current measurements across the continental slope off Cape
52
3 8° ~ • .
•
r ., .
.,
. :
.
t
.
.
•
.
. •
»
.
.
Ln
w
0
O
0
.!. .~ .
36° ' ....... . . . .. .: ,•
0 1 .~ .
.. ~ ~ t ,
;t
.
•
. ..
•
.
20cm2/s2
•
.
.
-.
~
.
,
~
34°
J
Figure 3 .4-5 Current variance ellipses associated with Topographic Rossby Waves (TRWs) on a line of deep
current meters 100 m above the bottom off Cape Hatteras . This example is for TRWs of
periodicity centered on about 40 days .
Hatteras . Interestingly, this lateral "average" DWBC velocity still has
significant variability, particularly at periods longer than about 100 days .
There is evidence from Sound Fixing and Ranging (SOFAR) and RAFOS floats as deep
as 1500-2000 m (Shaw and Rossby 1984) that beneath the Gulf Stream there occurs
intermittent strong flow in the direction of the Gulf Stream flow . These
intermittent deep Gulf Stream flows are possibly associated with large meanders
or displacements of the surface Gulf Stream (Hamilton 1987) . Mean near-bottom
flows deeper than 4000 m are observed to be much more variable with slight
northward trend (Watts 1991), which may be a reflection of these deep
northeastward flows . WCR can produce perceptible warming down to 2000 m, but
ring currents are generally negligible below 1000 m except perhaps during
formation or interaction with a Gulf Stream meander (Joyce 1984) .
54
Auer, S .J . 1987 . Five-year climatological survey of the Gulf Stream System and
its associated rings . J . Geophys . Res . 92 (11) : 11709-11726 .
Bane, J .M ., Jr ., O .B . Brown, R .H . Evans and P . Hamilton . 1988 . Gulf Stream Remote
Forcing of Shelf Break Currents in the Mid-Atlantic Bight . Geophys . Res .
Lett . 15(5) :405-407 .
Battelle . 1987 . Final Report on Analytical Results of Samples Collected During
the 1985 North Atlantic Incineration Site (NAIS) Survey . Final Report .
Prepared for the U .S . Environmental Protection Agency under Contract No .
68-03-3319 . Work Assignment 5 . 184 pp .
Battelle . 1991 . Satellite-tracked surface-layer drifters released at the 106-Mile
Site : October 1989 trough December 1990 . A report submitted to the U .S .
Environmental Protection Agency under Contract No . 68-C8-0105 . Battelle
Ocean Sciences Inc . Duxbury, MA . 26 pp .
Beardsley, R .C . and W .C . Boicourt . 1981 . On estuarine and continental shelf
circulation in the Middle Atlantic Bight . pp . 198-234 . In B .A . Warren and
C . Wunsch, eds . Evolution of Physical Oceanography . MIT Press,
Cambridge, MA .
Beardsley, R .C ., H . Mofjeld, M . Wimbush, C .N . Flagg and J .A . Vermersch, Jr . 1977 .
Ocean tides and weather-induced bottom pressure fluctuations in the
Middle-Atlantic Bight . J . Geophys . Res . 82 (21) : 3175-3182 .
Bisagni, J .J . 1983 . Lagrangian current measurements within the eastern margin of
a warm-core Gulf Stream ring . J . Phys . Oceanogr . 13(4) :709-715 .
Boicourt, W .C . 1973 . The circulation of water on the continental shelf from
Chesapeake Bay to Cape Hatteras . Ph .D . thesis . The Johns Hopkins
University . 197pp . (DAI 34/O1B, p .332 ; AAC 7316636) .
Brown, O .B ., P .C . Cornillon, S .R . Emmerson and H .M . Carle . 1986 . Gulf Stream warm
rings : A statistical study of their behavior . Deep-Sea Res .
33(11/12) :1459-1473 .
Butman, B . 1987 . Physical processes causing surficial sediment movement . pp . 147162 . In R .H . Backus, ed . Georges Bank . MIT Press, Cambridge, MA .
Churchill, J .H . and P .C . Cornillon . 1991a . Gulf Stream water on the shelf and
upper slope north of Cape Hatteras . Cont . Shelf Res . 11(5) :409-431 .
Churchill, J .H . and P .C . Cornillon . 1991b . Water Discharged From the Gulf Stream
North of Cape Hatteras . J . Geophys . Res . 96(C12) :22227-22243 .
Churchill, J .H ., P .C . Cornillon and G .W . Milkowski . 1986 . Cyclonic eddy and
shelf-slope water exchange associated with a Gulf Stream warm-core ring .
J . Geophys . Res . 91(C8) :9615-9623 .
Churchill, J .H ., P .D . Cornillon and P . Hamilton . 1989 . Velocity and Hydrographic
Structure of Subsurface Shelf Water at the Gulf Stream's Edge . J . Geophys .
Res . 94(C8) :10791-10800, 11009-11010 .
55
Churchill, J .H ., E .R . Levine, D .N . Connors and P .C . Cornillon . 1992 . Mixing of
shelf, slope and Gulf Stream water over the continental slope of the
Middle Atlantic Bight . Deep-Sea Res . In press
Csanady, G .T . and P . Hamilton . 1988 . Circulation of Slope Water . Cont . Shelf Res .
8(5-7) :565-624 .
Environmental Protection Agency . 1992 . Final report on current-meter measurements
at the 106-Mile Site in support of municipal waste disposal . U .S .
Environmental Protection Agency, Office of Water . Washington, D .C . EPA 82S-92-012 .
Fisher, A ., Jr . 1972 . Entrainment of shelf water by the Gulf Stream northeast of
Cape Hatteras . J . Geophys . Res . 77(18) :3248-3255 .
Ford, W .L ., J .R . Longard and R .E . Banks . 1952 . On the nature, occurrence, and
origin of cold low salinity water along the edge of the Gulf Stream . J .
Mar . Res . 11 :281-293 .
Garfield, N ., III and D .L . Evans . 1987 . Shelf Water Entrainment by Gulf Stream
Warm-Core Rings . J . Geophys . Res . 92(C12) :13003-13012 .
Gawarkiewicz, G ., R .K . McCarthy, K . Barton, A .K . Masse and T .M . Church . 1990 . A
Gulf Stream-derived pycnocline intrusion on the Middle Atlantic Bight
shelf . J . Geophys . Res . 95(C12) :22305-22313 .
Hamilton, P . 1984 . Topographic and inertial waves on the continental rise of the
Mid-Atlantic Bight . J . Geophys . Res . 89(Cl) :695-710 .
Hamilton, P . 1987 . The Structure of Shelf and Gulf Stream Motions in the Georgia
Bight . Prog . Oceanogr . 19 :329-351 .
Hogg, N .G . 1981 . Topographic waves along 70°W on the continental rise . J . Mar .
Res . 39 :627-649 .
Johns, W .E . and D .R . Watts . 1985 . Gulf Stream Meanders : Observations on the Deep
Currents . J . Geophys . Res . 90(C3) :4819-4832 .
Johns, W .E . and D .R . Watts . 1986 . Time Scales and Structure of Topographic Rossby
Waves and Meanders in the Deep Gulf Stream . J . Mar . Res . 44(2) :267-290 .
Joyce, T .M . 1984 . Velocity and hydrographic structure of a Gulf Stream warm-core
ring . J . Phys . Oceanogr . 14(5) :936-947 .
Joyce, T .M . and M .A . Kennelly . 1985 . Upper-ocean velocity structure of Gulf
Stream warm-core ring 82B . J . Geophys . Res . 90(C5) :8839-8844 .
Joyce, T .M ., J .K .B . Bishop and O .B . Brown . 1992 . Observations of offshore shelfwater transport induced by a warm-core ring . Deep-Sea Res . 39 :S97-S113 .
Kupferman, S .L . and N . Garfield . 1977 . Transport of low-salinity water at the
slope water-Gulf Stream boundary . J . Geophys . Res . 82(24) :3481-3486 .
56
Lillibridge, J .L ., III, G . Hitchcock, T . Rossby, E . Lessard, M . Mork and L .
Golmen . 1990 . Entrainment and mixing of shelf/slope waters in the nearsurface Gulf Stream . J . Geophys . Res . 95(C8) :13065-13087, 13559-13560 .
Louis, J .P ., B .D . Petrie and P .C . Smith . 1982 . Observations of topographic
Rossby waves on the continental margin off Nova Scotia . J . Phys .
Oceanogr . 12 :47-55 .
Luyten, J .R . 1977 . Scales of motion in the deep Gulf Stream and across the deep
continental rise . J . Mar . Res . 35(l) :49-74 .
McCartney, M .S . and L .D . Talley . 1982 . The Subpolar Mode Water of the North
Atlantic Ocean . J . Phys . Oceanogr . 12 :1169-1188 .
McLellan, H .J . 1957 . On the distinctness and origin of the slopewater off the
Scotian Shelf and its easterly flow south of the Grand Banks . J . Fish .
Res . Bd . Can . 14(2) :213-239 .
Morgan, C .W . and J .M . Bishop . 1977 . An example of Gulf Stream eddy-induced water
exchange in the Mid-Atlantic Bight . J . Phys . Oceanogr . 7 :472-479 .
Olson, D .B ., R .W . Schmitt, M . Kennelly and T .M . Joyce . 1985 . Two-layer diagnostic
model of the long-term physical evolution of warm-core ring 82B . J .
Geophys . Res . 90(C5) :8813-8822 .
Pickart, R .S . 1990 . Shallow and Deep components of the North Atlantic Deep
Western Boundary Current . Deep-Sea Res . Submitted
Pickart, R .S . 1992 . Space-time variability of the deep western boundary current
oxygen core . J . Phys . Oceanogr . 22(9) :1047-1061 .
Pickart, R .S . and D .R . Watts . 1990a . Deep Western Boundary Current variability
at Cape Hatteras . J . Mar . Res . 48(4) :765-791 .
Ramp, S .R . 1989 . Moored observations of current and temperature on the shelf and
upper slope near ring 82B . J . Geophys . Res . 94(C12) :18071-18087 .
Rhines, P .B . 1970 . Edge, bottom, and Rossby waves in a rotating stratified fluid .
Geophys . Fluid Dyn . 1 :273-302 .
Richardson, P .L . 1977 . On the crossover between the Gulf Stream and the Western
Boundary Undercurrent . Deep-Sea Res . 24 :139-159 .
Schultz, J .R . 1987 . Structure and Propagation of Topography Rossby Waves
Northeast of Cape Hatteras . Master of Science Thesis . Univ . of North
Carolina, Chapel Hill 63 pp .
Science Applications International Corporation . 1987 . Study of Physical Process
on the U .S . Mid-Atlantic Slope and Rise . Final Report, Volume II
Technical . Minerals Management Service, OCS Study MMS 87-0024, Atlantic
OCS Region, Vienna, VA .
57
Science Applications International Corporation/Battelle . 1988 . Draft Report on
Current Meter Measurements at the 106-Mile Site in Support of Municipal
Waste Disposal . Submitted to the U .S . Environmental Protection Agency
under Contract No . 68-03-3319 . Work Assignment 46 . 75p .
Shaw, P .-T . and C .Y . Peng . 1987 . A numerical study of the propagation of
topographic Rossby waves . J . Phys . Oceanogr . 17 :358-366 .
Shaw, P .-T . and H .T . Rossby . 1984 . Towards a Lagrangian description of the Gulf
Stream . J . Phys . Oceanogr . 14(3) :528-540 .
Stommel, H . and A .B . Arons . 1960 . On the abyssal circulation of the world ocean-II . An idealized model of the circulation pattern and amplitude in oceanic
basins . Deep-Sea Res . 6 :217-233 .
Thompson, J .D . and W .J . Schmitz Jr . 1989 . A limited-area model of the Gulf
Stream : design, initial experiments, and model-data intercomparison . J .
Phys . Oceanogr . 19(6) :791-814 .
Thompson, R . 1977 . Observations of Rossby waves near Site D . Prog . Oceanogr . 7 :128 .
Watts, D .R . 1991 . Equatorward currents in Temperatures 1 .8-6 .0 C on the
Continental Slope in the Mid-Atlantic Bight . pp . 183-196 . jn P .C . Chu and
J .C . Cascard, eds . Deep Convection and Deep Water Formation in the Oceans .
Elsevier, Amsterdam .
58
IV . THE CONTINENTAL SHELF
4 .1
Introduction
As a result of varying water depths and location on the east coast of the
continental United States, oceanographic conditions and circulation patterns on
the shelf offshore of North Carolina can be strongly influenced by local
meteorological processes over a broad range of time and space scales . As a
background for the oceanographic presentation to follow, the initial section of
this chapter describes the meteorological patterns which are important to the
observed oceanographical conditions .
This is followed by discussion of the oceanographic setting, which places the
continental shelf features within the context of the preceding chapters, and the
mean circulation on the continental shelf . Subsequent sections cover sources of
shelf variability, Virginia Coastal Water intrusions past Cape Hatteras, bottom
boundary layer processes, and sediment transport .
4 .1 .1 Meteorological Setting
Surface winds are the primary driving force for the currents over the continental
shelf . The sun provides the energy source to drive the atmospheric and oceanic
circulation patterns by differentially heating the earth's surface, so that low
latitudes (i .e . the tropics) receive a larger amount of solar insolation than do
the polar latitudes . It is through complex three-dimensional convective
circulation patterns and energy exchanges that the atmosphere-ocean system
redistributes and transports this solar energy to maintain a stable global
climate .
The oceans, because of their higher specific heat capacity, are an important
climatic element as a reservoir of heat energy . This allows for air-sea
interactions to occur whereby the overlying atmosphere may be cooled or warmed,
while the oceans maintain a relatively constant thermal structure . To maintain
a global energy balance, heat is transferred from low latitudes to polar regions
by ocean currents and the atmospheric winds . Thus, the prevailing general
circulation of the atmosphere provides the primary driving mechanism for the
surface ocean circulation and subsequent transport of heat energy . These
prevailing atmospheric motions contribute to the generation of such oceanic
circulation features as eastern (e .g . Canaries Current) and western (e .g . Gulf
Stream) boundary currents and equatorial currents .
4 .1 .1 .1
Basin Scale Patterns
The oceanic currents associated with the continental shelf offshore of North
Carolina are a portion of the overall North Atlantic basin ocean-atmosphere
circulation regime . Figure 4 .1-la shows the mean climatological sea level
atmospheric pressure field for January . The dominant North Atlantic region
feature this time of year is the Icelandic Low . This low pressure system (996-998
mb surface pressure) just west of Iceland provides a strong counterclockwise gyre
circulation in the higher latitudes of the North Atlantic basin . This system is
most intense during winter, weakens during the summer months and actually splits
into two separate systems . Figure 4 .1-lb shows the standard deviation of the
climatological sea level pressure field, indicating that this region is also
subject to large variations in the intensity and position of these cyclonic
59
^
do~
_
n
~
.
~
(a)
.•_
A
•B
-
.
.
(b)
.•
r
.
.!
Figure 4 .1-1 North Atlantic atmospheric pressure for January (a) mean (in mb
relative to 1000 mb), (b) standard deviation . Azores indicated
by A and Bermuda by B (U .S . Navy 1992) .
60
systems . This area is a region where migratory lows tend to stall and deepen
(AMS 1959) .
The summer season is characterized in the North Atlantic basin by the weakening
of the Icelandic Low and the subsequent growth of the North Atlantic subtropical
anticyclone, more commonly referred to as the Bermuda or Azores High, because of
its geographical proximity to these two regions (Figure 4 .1-2a) . Variability of
the Bermuda/Azores High (Figure 4 .1-2b) is approximately 50% less than the winter
regime, the feature influences a considerably larger geographic area and
generally has weaker pressure gradients and hence less vigorous winds than in
winter . This large pressure system directly influences both the wind and current
systems offshore of North Carolina during summer (AMS 1959) .
The transition seasons (i .e . spring and fall) in the North Atlantic basin are
characterized by intense thermal contrasts as the general atmospheric circulation
pattern shifts from more stable, cold, dry continental air masses traversing the
region (winter) to more unstable warm, moist tropical air masses (summer) . The
transition periods are marked by intense atmospheric storms .
4 .1 .1 .2 Regional Patterns
The winter season for the region offshore of North Carolina generally runs
between November and March, and is characterized by north-northwesterly wind
flows with mean speeds of about 8-10 m s-1) over the continental shelf areas
consistent with the cyclonic circulation associated with the Icelandic Low
(Figure 4 .1-1) . Figure 4 .1-3 shows the climatological mean sea surface
temperature , air temperature-SST difference and wind fields during January for
this region . The Gulf Stream influence during winter is noticeable in the air
temperature-SST difference diagram (Figure 4 .1-3b), which shows the 4-6°C
difference between cold, dry continental air masses flowing out over the warmer
Gulf Stream waters offshore . These conditions have long been recognized as
important for atmospheric cyclogenesis (generation of cyclonic storm systems) and
ocean-atmosphere energy exchanges .
During April, the general circulation of this region begins to shift towards the
summer circulation regime . The influence of the Bermuda/Azores High begins to
dominate the region south of Cape Hatteras, while the Icelandic Low continues to
persist for the areas north of Cape Hatteras . Near Cape Hatteras the transition
is marked by a change from a north-northwesterly winds to a more southerlysouthwesterly pattern, and the wind speeds decrease to about 7-8 m s-l . As the
subtropical, anticyclonic central pressure builds, its areal extent also
increases until it becomes the dominant summer circulation feature for the North
Atlantic basin .
The summer flow regime (between May and August) is characterized by small thermal
contrasts between the ocean and the atmosphere over the continental shelf region
offshore of North Carolina (Figure 4 .1-4a,b) . The regional wind pattern,
dominated by the anticyclonic circulation around the Bermuda High, produces
generally south-southwesterly winds with speeds averaging less than 6 m s-i
(Figure 4 .1-5c) . The warm, tropical waters in the southern North Atlantic Ocean
during summer are favorable to formation and development of tropical cyclone
systems . These systems provide a very efficient transport of latent heat energy
from both the ocean to the atmosphere and from tropical to polar latitudes . The
61
(a)
-40
`
e
.
6
0
~
~
(b)
.•
A
.
.
IB
~
Figure 4 .1-2 North Atlantic atmospheric pressure for July (a) mean (in mb
relative to 1000 mb), (b) standard deviation . Azores indicated
by A and Bermuda by B (U .S . Navy 1992) .
62
\
.
(a)
~
.B
\
1
1
ca :
.. .n:
..cc>
~
(b)
-
~
/
B
~ .
U
~
> :~o.° s . ./, ., <c>
Figure 4 .1-3
North Atlantic temperature for January ( a) SST (°C), (b) air
temperature-SST difference ( °C) . Bermuda indicated by B (U .S .
Navy 1992) .
63
~
.
l\ ~'
)i
q~niYl~
)
Figure 4 .1-3c North Atlantic wind speed for January ( m s'1) . Bermuda
indicated by B (U .S . Navy 1992) .
64
Figure 4 .1-4
North Atlantic temperature for April ( a) SST (°C), (b) air
temperature-SST difference (°C) . Bermuda indicated by B (U .S .
Navy 1992) .
65
-.o
J
~
.~
.,,
.,
.
Figure 4 .1-4c North Atlantic wind speed for April ( m s-1) . Bermuda indicated
by B (U .S . Navy 1992) .
66
~
.
(a)
~
AB
.
~
\
.
Id
'1
, <s, :m.. wwsr °°.c
'r
(b)
B
~
a.~
-.e
Figure 4 .1-5
ao
'
,
,1
•
ao
-oo
North Atlantic temperature for July ( a) SST (°C), (b) air
temperature-SST difference ( °C) . Bermuda indicated by B (U .S .
Navy 1992) .
67
Figure 4 .1-5c North Atlantic wind speed for July ( m s'1) . Bermuda indicated
by B (U .S . Navy 1992) .
68
North Atlantic Tropical Cyclone Season officially runs between June and November
(U .S . Navy 1989) .
During September, the Bermuda/Azores High begins to weaken and the Icelandic Low
begins to once again impart a degree of influence on the general circulation
features of the North Atlantic basin . Fall season (September and October) winds
near Cape Hatteras are generally north-northeasterly with speeds similar to the
spring season (=7-8 m s-1) . This circulation tends to bring cooler air over the
region, and in conjunction with the reduction in solar insolation, the air
temperature-SST difference again becomes an important mechanism for transferring
heat from the ocean surface to the atmosphere through both latent and sensible
heat fluxes . During fall, partially in response to the increasing baroclinicity
of the atmosphere, extratropical cyclone activity increases in the region
offshore of North Carolina .
The ocean-atmosphere system is a complex entity controlled ultimately by the
influx of solar radiation . However, it is the complex, turbulent interactions
between the ocean and the atmosphere which provide the characteristic patterns
associated with ocean currents and atmospheric wind systems . Through these
interactions a stable, but dynamic global climate is maintained .
4 .1 .1 .3 .
SvnoQtic Scale Disturbances
The amount of heat and moisture available to the atmosphere from the open sea
surface is generally large in regions adjacent to eastern continental margins due
to the presence of warm western boundary currents such as the Gulf Stream . The
conditions are optimized geographically along the east coast of the United States
(e .g . offshore of North Carolina) and seasonally during winter months when the
water-land temperature contrast is maximized . The relatively cold continent is
bounded by relatively warmer shelf/slope water and by the consistently warmer
Gulf Stream to the east . The prevailing meteorological conditions in this area
generate a synoptic wind flow from west to east, with winter weather disturbances
propagating from the land eastward out over waters of the shelf and slope .
(Wayland 1991 ; Wayland and Raman 1989) .
During the winter months four distinct SST regions occur offshore of North
Carolina : (1) Inner Shelf (=10-12°C) ; (2) Outer Shelf (=16-18'C) ; (3) Gulf Stream
Core (=22-24°C) ; and (4) Sargasso Sea ( - 18-20'C) . As synoptic scale atmospheric
flow traverses these changing ocean thermal surfaces, the rapidly developing
Marine Atmospheric Boundary Layer (MABL) may become strongly baroclinic . During
Cold Air Outbreaks (CAOs) these baroclinic conditions are maximized, and
extremely large fluxes of momentum, heat and moisture occur with momentum going
from the atmosphere to the ocean and with heat and moisture from the ocean to the
atmosphere . In terms of energy exchange, the cold air outbreak is one of the
most dynamic air-sea interaction events (Wayland and Raman 1989 ; Wayland 1991) .
In strict terminology, a cold air outbreak occurs when a cold, dry continental
air mass (generally of Polar origin) pushes out over the warmer oceanic waters
offshore of eastern continental margins . These events occur approximately 15-20
times a year along the Middle Atlantic coast and generally have a duration of
less than two days . Approximately one third of these systems are classified as
intense or extreme cold air outbreaks, where the air temperature is less than 0'C
and the core Gulf Stream surface temperature is greater than 20°C (Konrad and
Colucci 1989 ; Grossman and Betts 1990) .
69
During the recent Genesis of Atmospheric Lows Experiment (GALE) detailed
observations were made of the evolution and influence of air-sea interaction
during CAOs . Several CAO episodes illustrated that both sensible and latent heat
fluxes increased dramatically as the colder, drier air moved out over the
comparatively warm shelf waters . The lower atmosphere warmed and moistened as
heat and moisture were extracted from the coastal ocean . The presence of the
Gulf Stream just off the North Carolina shelf provides a continually replenished
source of warmer water, which helps maintain the observed cross-shelf SST
gradient seen during winter .
The examination of several CAOs during GALE also showed that the heat flux tended
to increase with distance offshore . This may have been because the MABL was seen
to respond quite rapidly to the underlying pattern of sea surface temperatures,
which increased with distance from the coast (Wayland and Raman 1989 ; Grossman
and Betts 1990 ; Wayland 1991) . No current measurements were made in the study
during GALE so the response of the coastal circulation pattern near Cape Hatteras
could not be determined ; however, water temperatures showed that a single CAO
resulted in a substantial deepening of the surface mixed layer in the Gulf Stream
and a heat loss corresponding to a decrease in the average mixed layer
temperature of approximately 0 .62°C (Bane and Osgood 1989) .
Frontal systems associated with CAOs and migrating cyclonic disturbances have
been identified as important factors influencing the overall winter circulation
patterns in the South Atlantic Bight (Lee 1989) . The periodicity of
meteorological frontal passage is often seen in surface and near surface currents
as the wind stress field drives a corresponding current pattern (Hamilton 1987) .
Cold air outbreaks and the passage of atmospheric frontal systems are recognized
as important contributors to the observed coastal circulation patterns offshore
of North Carolina . However, strong extratropical and tropical cyclones provide
periodic strong and sometimes sustained wind events, which can substantially
affect oceanographic conditions ranging from circulation patterns and wave
climatology to bottom boundary layer processes and sediment transport . The
location of the present study area relative to the jet stream and cyclone
trajectories, combined with the unique SST structure provided by the presence of
the Gulf Stream just offshore, helps insure that episodic storms have an
important role in defining the varied oceanographic patterns which occur in the
present study area .
The eastern United States coast has long been recognized for the occurrence of
strong extratropical and tropical cyclones : Hurricane Hazel, 1954 ; Ash Wednesday
Storm, 1962 ; President's Day Storm, 1978 ; Hurricane Hugo, 1990 ; Halloween Storm,
1991 . Hayden (1981) generated an extratropical cyclone climatology based on the
frequency of occurrence for storms traversing the eastern two-thirds of the
United States for the period 1885-1978 . The mean results delineated an active
cyclone corridor aligned just offshore of the eastern United States . This axis
shifted offshore and eastward when the standard deviations of the climatology
were analyzed, revealing the prominence of these regions for atmospheric
cyclogenesis .
The variance in cyclone frequency was attributed to changes in the east coast
baroclinic zone (e .g . cold air outbreaks, etc .) and to shifts in the North
American long-wave trough/ridge locations and to blocking patterns in the high
latitudes . Additionally, a decline in the frequency of low pressure systems in
70
the Colorado Rockies was linked to an increase in Atlantic cyclone frequency,
while a general downtrend in cyclone activity resulted in a decrease in midAtlantic coastal cyclogenesis .
Results from GALE show the importance of subsynoptic scale features and thermal
advection in explosive cyclogenesis events (Wash et al . 1990) .
Tropical cyclones, which are driven primarily by latent heat input from the
ocean, originate in summertime tropical waters where the SST is on average >
26°C . These systems, at mature stages, can have winds exceeding 90 m s-1,
torrential rains and destructive, deadly tornadic systems during landfall . The
"official" North Atlantic Hurricane Season begins on June 1 and continues through
the summer until November 30 (US Navy, 1989) .
The National Hurricane Center (NHC) has compiled statistics for storms
originating in the North Atlantic Basin for the period 1871-1980 . The average
duration of these systems is eight days, but ranges from two to 30 days for the
period studied . The most frequently occurring duration (e .g . mode) is six days .
Over this period, 21 storms have made a direct hit on North Carolina coastal
regions, while countless others have tracked through this area (Neumann et al .
1981) . Winds from storms tracking further offshore have little direct influence
other than to create larger and more vigorous waves, which can affect beach
erosion and bottom sediment transport . However, decaying storms tracking further
inland can cause heavy amounts of precipitation across the state . September is
the month of highest storm frequency (=338 of all storms), while August (24%) and
October (22%) follow closely . The remaining months of the year show a marked
decrease (typically < 10%) in tropical cyclone probability (Neumann et al . 1981) .
A U .S . Navy climatology atlas (1989) reported only April had no reported tropical
cyclones during a 101-year period .
4 .1 .2 Oceanographic Setting
For the eastern continental shelf of the United States, Cape Hatteras and its
offshore extension, Diamond Shoals, have been traditionally viewed as an
oceanographic "barrier," separating the waters of the Middle Atlantic Bight, with
their distinct flora and fauna, from the South Atlantic Bight waters (see Figures
1 .1-1 and 1 .1-2) . Although wind-driven breaches, movement of Middle Atlantic
Bight shelf waters into Raleigh Bay in response to northeasterly winds, of this
barrier were well documented prior to the advent of moored instrumentation and
satellite perspectives (Bumpus and Pierce 1955 ; Stefannson et al . 1971), mean
hydrographic and faunal differences across Diamond Shoals are clearly sufficient
to warrant consideration of these bodies as two separate oceanographic provinces .
However, local effects of these exchanges, especially in Raleigh Bay, can be
significant .
One primary characteristic that these two provinces share is a strong influence
from the nearby Gulf Stream . In general, the mode of interaction with the Gulf
Stream differs between the Middle Atlantic Bight and the South Atlantic Bight
(see Chapters 2 and 3) . To the south, lateral variations in the Gulf Stream
position are constrained to a considerably narrower domain than in the Slope Sea
north of Cape Hatteras . Thus, in the Middle Atlantic Bight the influence of the
Gulf Stream is often indirect, for example through the detached anticyclonic
rings and through the Gulf Stream's effect on the Slope Sea gyre . In the South
Atlantic Bight, the influence is usually direct, through filaments and frontal
71
eddies along to the western wall of the Gulf Stream . Satellite imagery has shown
that the north-south differences between these modes of interaction decrease in
the approaches to Cape Hatteras . The primary reason for this merging process is
the decreasing width of the Slope Sea in the southern Middle Atlantic Bight .
To the north of Cape Hatteras, the Middle Atlantic Bight stretches more than 800
km to Cape Cod as shown in Figure 1 .1-2 . As the shelf width expands gradually
toward the north, the depths of both the shelf and shelf break decrease to a
minimum along the broad ridge extending seaward off the mouth of Chesapeake Bay .
Typical mid-shelf depths immediately north of Cape Hatteras are 40 m, while off
Chesapeake Bay, mid-shelf depths are 25 m . To the north of Chesapeake Bay, the
shelf maintains its roughly 100 km width while deepening to the more typical 40
m mid-shelf depths of the remainder of the Middle Atlantic Bight .
Southward from the narrow continental shelf off Diamond Shoals at Cape Hatteras,
the South Atlantic Bight does not expand uniformly to the 200 km width off
Georgia, but undergoes three large oscillations marked by the Carolina Capes and
their offshore shoal extensions (Figures 1 .1-1 and 1 .1-2) . The regularity of the
Capes and the embayments they define--from north to south, Raleigh Bay, Onslow
Bay, and Long Bay--has led to speculation as to their origin . Abbe (1895 ; see
Bumpus 1955, 1973 ; Brooks and Bane 1981) inferred a series of counterclockwise
"back eddies" from the Gulf Stream (reported from ship drifts) that coincided
with the Carolina bays . While the giant cusps continue to fascinate geologists
and physical oceanographers, they also provide a regional demarcation between the
circulation processes of the Carolina Bays and the circulation of the smoother
shelf to the south . These regions differ in the amount of circulation control
exerted by the topography and in the degree and mode of Gulf Stream interactions
(Atkinson and Menzel 1985 ; Pietrafesa et al . 1985b) . An unusual aspect of this
difference between northern and southern regions of the South Atlantic Bight is
that offshore topography plays a crucial role in the separation . At about
31°30'N, a submarine ridge, the Charleston Bump, extends seaward from the
continental slope . This feature produces an offshore deflection in the Gulf
Stream and a series of recirculation eddies or frontal deformations in its lee .
Although these larger eddies appear frequently offshore of the Carolina Capes,
similar frontal filaments on the western wall of the Gulf Stream can occur
throughout the South Atlantic Bight (Atkinson and Menzel 1985) . In the southern
portion of the South Atlantic Bight, as the shelf narrows off the east coast of
Florida, the proximity of the Gulf Stream ensures that these cold-core frontal
eddies are an important contributor to the shelf circulation .
4 .1 .3 Mean Circulation
By the time Henry Bryant Bigelow began his pioneering hydrographic work on the
Middle Atlantic Bight, there was a solid body of evidence indicating a southwest
drift of water from Georges Bank to Cape Hatteras (Beardsley and Boicourt 1981) .
Bigelow (1915, 1922, 1933) described the seasonal change in the Middle Atlantic
Bight, from a cold, well-mixed water column in the winter to one of strong
stratification in summer . He noted the cold band of water along the bottom over
the outer shelf and concluded that it was produced locally and not replenished
during the summer from the north . With the insight provided by 10 years of
drift-bottle and seabed drifter releases (Bumpus 1973), and from moored currentmeter arrays (Beardsley et al . 1976 ; Beardsley and Boicourt 1981) on the Middle
Atlantic Bight, the southward mean flow on the order of 5 cm s-1 has been
established for all depths, including the summer cold band of water . Information
72
from the moored arrays also confirmed Iselin's (1939, 1940) inference of a crossshelf shear in the mean flow, with maximum velocities occurring near the shelf
break .
Beardsley et al . (1976) found that the alongshelf volume transport (2 .0 x lOs m3
s'1) through three cross shelf current-meter arrays was remarkably similar
(Figure 4 .1-6), in spite of the spatial and temporal (including seasonal)
differences in measurements . This uniformity encouraged them to postulate little
net loss of water across the shelf-slope front . Their advective shelf-water
residence time of 0 .75 years was less than Ketchum and Keen's (1955) value of 1 .3
years, which was based on an assumption that the primary transport occurred only
in the cross shelf direction . Beardsley et al . (1976) also speculated that most
of the shelf water observed flowing westward south of New England must be part
of a continuous flow originating on the southern flank of Georges Bank and in the
Gulf of Maine .
Bigelow noted the influence of low-salinity water discharged from the large
estuaries on the east coast of the United States--Narragansett Bay, Long Island
Sound, the Hudson River, Delaware Bay, and the largest, Chesapeake Bay . Although
these estuaries discharge on the order of 4000 m3 s-i of fresh water to the
shelf, the primary inflow of water arrives from Georges Bank, south of Nantucket
Shoals (Figure 1 .1-2) . Despite the addition of fresh water along the coast, mean
salinity of a cross-shelf section increases as the Georges Bank water traverses
the 800 km from Nantucket Shoals to Cape Hatteras . The salt to supply this
increase in salinity is transported from the slope water across the sharp front,
which extends continuously along the shelf break from Georges Bank to near Cape
Hatteras .
As southward-flowing shelf water approaches Cape Hatteras, bathymetry steers the
flow offshore . The traditional view of the fate of this water is that, while
there are occasional southward flows over Diamond Shoals into Raleigh Bay, the
majority moves offshore and becomes entrained in the Gulf Stream, producing
narrow, discontinuous filaments of cold, fresher water stretching for hundreds
of kilometers along the north wall of the Gulf Stream . This process was first
described by Ford and Miller (1952) and Ford et al . (1952) and has been examined
more recently by Fisher (1972), Kupferman and Garfield (1977), Csanady and
Hamilton (1988), and Churchill et al . (1989) . The Gulf Stream can, in turn, send
intrusions of warm, salty water onto the shelf in this region (Churchill and
Cornillon 1991b ; Gawarkiewicz et al . 1992) . An unanswered question with regard
to the onshore-offshore exchange in this region is where the cold-water band over
the outer shelf exits the continental shelf and becomes entrained in the Gulf
Stream . This band is a common component of the "Ford Water" observed in
discrete, elongated filaments on the north wall of the Stream . Hydrographic
measurements (Figures 4 .1-7 and 4 .1-8) suggest that the cold band moves offshore
in the vicinity of 35°40'N, but this position is expected to vary substantially
with Gulf Stream position and with changes in wind forcing . The enhanced
temporal and spatial perspective provided by satellite thermal imagery suggests
that both overflow of Diamond Shoals ("breaching the barrier") and Gulf Stream
entrainment will require a significant amount of study in order to provide a
quantitative measure of these complex and highly time-dependent processes .
As is the case for other shelf regions with strong alongshelf flow, accurate
detection of the cross shelf component of flow has proven difficult in the Middle
Atlantic Bight . Iselin's (1939, 1940) inference, from the salinity structure,
73
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76
of mean offshore flow at the surface and onshore flow along the bottom has been
corroborated by both drifter (Bumpus 1973) and moored current measurements
(Beardsley and Boicourt 1981) . Bumpus (1965, 1973) suggested a "line of
divergence" for the near bottom flow, which was located one-half to two-thirds
the distance from the coast to the shelf break . Offshore of this line, the near
bottom flow was directed offshore . Subsequent support for this inference came
from Csanady's (1976) modeling study which showed that, over the outer shelf, the
offshelf drift in the bottom Ekman layer can overcome the density-driven onshore
flow . Further support has come from current-meter observations (Beardsley and
Boicourt 1981 ; Butman et al . 1982 ; Beardsley et al . 1985 ; Butman et al . 1987) .
In contrast to the shelf regions to the north, the South Atlantic Bight does not
have a clearly defined mean circulation, except perhaps for the buoyancy-driven
flow along the inner shelf (Atkinson et al . 1983) . Observed mean currents are
more a climatological average of the highly variable currents driven by the
atmosphere and by offshore fluctuations of the Gulf Stream than they are a
product of a consistent, large-scale pressure gradient . In this sense, the South
Atlantic Bight is more of an event-dominated system than is the Middle Atlantic
Bight to the north .
In the cross-shelf direction, the South Atlantic Bight can be separated into
three circulation provinces, each with a primary driving mechanism . When
freshwater discharge from rivers and estuaries is high, a buoyancy-driven current
flows southward in a narrow nearshore band . The broad, shallow (30 m) middle
shelf is driven primarily by the winds . Over the outer regions of the South
Atlantic Bight, the flow and variations of the shelf waters are dictated by the
nearby Gulf Stream .
In spite of the weakness of the mean flows in the South Atlantic Bight, at least
directional tendencies have been suggested on the basis of information from both
drifters and moored current meters (Atkinson et al . 1983) . Bumpus (1955, 1973)
deduced that the southward flowing coastal current was narrower and more
transient than the coastal current in the Middle Atlantic Bight . Over the middle
shelf, he inferred a slow, broad northward flow . Bumpus' (1973) drifter diagrams
indicate that, except for Raleigh Bay and Onslow Bay in the northern portion of
the South Atlantic Bight, this flow reverses into a comparatively pronounced
southward drift during late summer and early fall (Figures 4 .1-9 and 4 .1-10) .
Although this flow appears persistent during this interval, it is apparently not
sufficient to reverse the weak annual mean flow to the north .
In recent years, research programs sponsored by the Minerals Management Service,
the National Science Foundation, the Office of Naval Research, and the Department
of Energy have substantially improved the coverage and resolution of current
measurements in the South Atlantic Bight . For the most part, the new
measurements are consistent with the earlier drifter studies, with mean
velocities at the detection limits of the temporal sampling and the accuracy of
the measurement . In the Carolina Capes region, Pietrafesa et al . (1985b)
reported northeastward mean flows in Onslow Bay and off Cape Romain, but
southwestward means in Long Bay . As Bumpus' (1955, 1973) drifters indicated,
these direct measurements showed a seasonal reversal in the middle shelf off Cape
Romain, but a persistent northeastward drift in Onslow Bay .
Observations of a persistent southward flow during the Georgia Bight Experiments
(GABEX I and II ; Lee and Atkinson 1983 ; Lee et al . 1985 ; Lee and Pietrafesa,
77
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1973) .
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Bumpus 1973) .
1987) over the outer shelf (especially in the lower layer) at 32°N were
originally thought to be anomalous, yet similar southward flows appeared at depth
in winter 1986 during the Genesis of Atlantic Lows Experiment (GALE) (Lee et al .
1989) . Lee et al . (1985) suggested that this flow was not anomalous, but rather
a manifestation of the semipermanent recirculation eddy downstream from the
Charleston Bump ( Brooks and Bane 1978 ; Pietrafesa et al . 1978 ; Rooney et al .
1978 ; Chao and Janowitz 1979 ; Legeckis 1979 ; Bane 1983 ; Olson et al . 1983 ;
Singer et al . 1983) . This cyclonic feature is consistent with the observed
doming of isotherms downstream from the Charleston Bump (Bane 1983 ; Singer et al .
1983) .
The continental shelf near Cape Hatteras constitutes a boundary region between
two large provinces with distinctly different water masses and mean circulations .
The combined presence of an occasionally leaky physical obstacle- -Diamond Shoals-and a strong, often dominant oceanographic forcing- -the Gulf Stream- -renders the
definition of a mean circulation in this zone difficult . The evidence points to
a large-scale convergence in the mean circulation at Cape Hatteras, although the
southward and eastward mean flow in the region north of Cape Hatteras appears to
be more established than is the northward mean in the region immediately to the
south of the Cape . Establishment of these mean circulations and the major modes
of variability about the mean is important, not only for understanding the
circulations on the Middle Atlantic and South Atlantic continental shelves, but
also for estimating the offshore flux of water-borne materials . To provide an
assessment of the transport and mixing processes with adequate confidence,
significantly more observational information will be necessary from this
spatially complex and highly time-dependent region . The most glaring
uncertainties are the statistical details of the Gulf Stream entrainment
processes and quantitative measures of the frequency and transport of shelf-toshelf exchanges across Diamond Shoals .
4 .2
Shelf Variability and Forcing Mechanisms
4 .2 .1 Tides
Tidal currents at semidiurnal and diurnal periods are important high frequency
motions on the continental shelf around Cape Hatteras . Maximum tidal flows may
be as large as 20 to 30 cm s-1 in the middle of the shelf . Large tidal currents
have important effects on vertical and horizontal mixing processes on the middle
and inner shelf . The width of the shelf is an important influence on the
magnitude of tidal currents, and in this region it varies from about 100 km of
the mouth of the Chesapeake Bay to about 20 km off Cape Hatteras and to about 60
km at the center of Onslow Bay . The coastline makes almost a right-angled turn
between the Middle Atlantic and South Atlantic Bight at Cape Hatteras, which may
be important for continental shelf waves at diurnal tidal frequencies .
Tidal currents have been measured and analyzed in this region by Pietrafesa et
al . (1985a) for Onslow and Long Bays, the FRED group for Raleigh and Onslow Bays
(FRED Group 1989) and Mobil, along a transect just north of Diamond Shoals (SAIC
1991) . Further north, there are extensive measurements around the mouth of the
Chesapeake and offshore of Cape Henry (Beardsley and Boicourt 1981) . The tidal
characteristics are quite well described by the analytical theories of Battisti
and Clarke (1982) which have been recently reviewed and summarized by Clarke
(1991) . The major assumptions in the theory are that the longshore scales of the
topography are much larger than the cross shore scales so the coast can be
80
considered quasi-straight, and the shelf tides are forced by the offshore, deep
ocean tide .
The semidiurnal (MZ) tides are characterized as an inertial-gravity wave
response, where the major axes of the tidal ellipses are approximately
perpendicular to the isobaths and the largest amplitudes occur in the center of
the shelf where the shelf width is largest . The M2 current amplitudes vary from
5 cm s'1 off Cape Hatteras to about 15 cm s-1 in the center of Onslow Bay . The
currents rotate clockwise through the tidal period . The M2 elevations with
amplitudes of about 45 cm are approximately in phase all along the coast between
Duck and Southport (FRED Group 1989) so the inertial gravity wave can be
considered to have the form of a standing wave between the shelf break and the
coast (Clarke 1991) . Thus, M2 tidal current amplitudes tend to be small over the
steep continental slope and near the coast .
Coastal elevation amplitudes for the diurnal tides (K1, P1 and 01) are about one
fifth that of the M2 (K1, which dominates, has an amplitude of =9 cm), but show
distinct north to south phase propagation . Duck leads Southport by about 1 .8
hours . The external, deep ocean equilibrium K1 tide also shows southward
propagation along the U .S . east coast (Daifuku and Beardsley 1983 ; Redfield
1958) . However, diurnal motions are supported by southward propagating
continental shelf waves at these latitudes (w<f) . Daifuku and Beardsley (1983)
attribute part of the Middle Atlantic Bight, K1 tidal current signal to southward
propagating continental shelf waves . A similar kind of response seems to occur
south of Cape Hatteras (FRED Group 1989) .
Diurnal current ellipses are aligned approximately along the isobaths with the
largest amplitudes near the shelf break . Smaller amplitudes occur nearer the
coast on the wider parts of the shelf but not where the shelf is narrow as at
Cape Hatteras (FRED Group 1989 ; SAIC 1991) . Current vectors rotate clockwise as
would be expected (Clarke 1991) . The region near Cape Hatteras, with its narrow
shelf, is the only region of this coast where the diurnal and semidiurnal tidal
current amplitudes have similar magnitudes of about 5-7 cm s-1 . This makes the
tidal current signal more complicated, with large diurnal inequalities and
spring-neap cycles, than regions to the north or south, which are dominated by
M2 tidal currents .
Internal M2 tides are probably generated along some parts of the shelf break in
this region because of the steep continental slope and the generally strong
stratified conditions which prevail throughout the year . However, there has been
little investigation of internal tides for this region . The FRED group (1989)
found evidence of enhanced near bottom M2 tidal currents over the outer shelf
offshore of Frying Pan Shoals (Cape Fear) which is consistent with active
generation of internal tidal waves (Wunsch 1969) . Further north near Baltimore
Canyon, Burrage and Garvine (1988) reported large vertical excursions, at
semidiurnal periods, of order 15 to 30 m of isopycnals in the lower part of the
water column in the region of the summer shelf-break front . This again suggests
active internal wave generation .
4 .2 .2 Buoyancy Forcing
Although a broad understanding of the mean southward flow over the continental
shelf in the Middle Atlantic Bight has been established for decades, the driving
mechanism for such motion has continued to be a matter of speculation . Until
81
recently, the establishment of such a mechanism has proved frustratingly
intractable . Huntsman (1924), Bigelow, (1927) and Iselin (1955) cite "rules" of
coastal circulation, among which is the dictum that, in the northern hemisphere,
flow is parallel to the coast and with the land on the right-hand side of an
observer facing downstream . Although some early investigators inferred a
buoyancy flow driven by the discharge of fresh water along the coast, the "rules"
were only consistent with geostrophy, and did not reveal the driving force .
Beardsley and Boicourt (1981), Csanady (1982), and Butman et al . (1987) reviewed
the attempts to explain this motion, which is in the opposite direction to the
mean wind stress . Although there was a consensus that an alongshore pressure
gradient must exist to drive the flow, its origin was uncertain . External
forcing from the Slope Sea gyre north of the Gulf Stream was a suggestion by
Sverdrup, et al . (1942), Csanady (1978, 1979), and Beardsley and Winant (1979) .
Shaw (1982) and Wang (1982) have cast doubt on this possibility by showing that
vorticity constraints restrict steady cross-slope motion and thereby isolate the
shelf and slope barotropic motions . Csanady (1979) examined the earlier
suggestion that outflow from the St . Lawrence River could provide the alongshore
pressure gradient sufficient to drive the mean flow from the Gulf of Maine to the
Middle Atlantic Bight . He found that the influence of the St . Lawrence could be
detected in stearic heights over the Scotian Shelf and the Gulf of Maine, but not
further south .
Recently, Chapman et al . (1986) established from oxygen-isotope measurements and
modeling analyses that the alongshelf flow was continuous from the Scotian Shelf
through the Gulf of Maine, around Georges Bank, and through the Middle Atlantic
Bight (see Figure 4 .2-1) . In addition, they concluded that the regional largescale circulation did not create the alongshelf pressure gradient in the Middle
Atlantic Bight, but that it helped keep the shelf water on the shelf . Oxygenisotope data indicated that the origin of the Scotian Shelf water was not the St .
Lawrence River, but some unidentified source to the north . Subsequently, Chapman
and Beardsley (1989) suggested that the origin of the Middle Atlantic Bight water
is located in the northern Labrador Sea . The Middle Atlantic Bight would
therefore be the end of a 5000 km, buoyancy-driven, coastal current originating
along the southern coast of Greenland . Although the flow dynamics are as yet not
proven, Chapman and Beardsley's (1989) hypothesis provides a satisfying
explanation for many of the known features of the alongshelf coastal current and
serves as a valuable framework to structure future observational and modeling
studies .
The southward increase in the sectional mean salinity on the shelf in the Middle
Atlantic Bight, from Nantucket Shoals to Cape Hatteras, despite the discharge of
4000 m3 s-1 of fresh water from rivers and estuaries along the coast, might
suggest that the local buoyancy input is not a strong contribution to the shelfscale, buoyancy-driven flow . Whether or not this is the case, discharges from
these estuaries can drive buoyancy flows that, in the absence of strong wind
driving, are rotationally trapped along the coast (Boicourt 1982 ; Chao and
Boicourt 1986 ; Boicourt et al . 1987 ; see section V .) . When upwelling favorable
(southwest) winds force these estuarine plumes offshore, the Ekman circulation
broadens the low-salinity tongue and detaches the plume from the coast in the far
field . The plume can reach mid-shelf during these events . In spite of the
opposing wind stress, the broad plume moves southward, reversing only under
conditions of unusually strong southeasterly winds . The dynamics of buoyancy
driving under these conditions and the degree of coupling between the plume and
the ambient shelf circulation is presently uncertain .
82
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Figure 4 .2-1 Chapman and Beardsley's (1989) schematic circulation diagram for
the coastal water from the West Greenland Current to the Middle
Atlantic Bight .
83
Bumpus (1955) offered reasons for the lack of a "dynamically driven" (buoyancy
driven) coastal current in the South Atlantic Bight . In light of more recent
work (especially Chapman and Beardsley's (1989) findings for the Middle Atlantic
Bight), most of these reasons appear to have substantial merit . The primary
missing ingredient is a strong buoyancy input from the Middle Atlantic Bight,
around Cape Hatteras . In addition, local buoyancy input along the coast is less
than half that of the Middle Atlantic Bight (Bue 1970) . Bumpus argued that the
southwest winds blow parallel to the coastline, countering any southward tendency
for the coastal current . With the observed strong coupling of the mid-shelf
flows to the wind (Lee, et al . 1985) and the mean winds having an alongshore
component from the southwest (Blanton et al . 1985), this process would contribute
to the prevention of a mean southward flow on the shelf . Over the outer portion
of the shelf, the direct influence of the Gulf Stream counters any tendency for
mean southerly flow, although the details of the "frictional drag" (Bumpus 1955)
process are uncertain in light of the myriad forms of Gulf Stream interactions
revealed in satellite SST imagery .
As is the case for the Middle Atlantic Bight, discharges from estuaries and
rivers along the South Atlantic Bight coast create buoyancy-driven coastal
currents in the nearshore region (Blanton 1981 ; Blanton and Atkinson 1983 ; see
section V .) . In the northern region of the South Atlantic Bight, the exchanges
through Hatteras and Ocracoke Inlets into Raleigh and Onslow Bays appear
sufficiently small to create only minor perturbations of the shelf circulation .
4 .2 .3 Atmospheric and Boundary-Current Forcing
In the Middle Atlantic, moored current observations in the 1970's revealed that
atmospheric forcing could easily dominate the 5 cm s-1 mean flow to the south .
Beardsley and Butman (1974) and Boicourt and Hacker (1976) showed that strong
winter storms could produce alongshore currents on the order of 50 cm s-i over
the mid-shelf . A general description of the response of the Middle Atlantic
Bight to atmospheric forcing has been presented by Beardsley and Boicourt (1981),
Csanady (1982), and Allen et al . (1983) . Beardsley et al . '(1985) found that the
efficiency of wind driving varied seasonally, with summer allowing a relatively
strong current response to weak wind fluctuations during that season . Ou et al .
(1981) used current records to demonstrate the existence of free, coastal-trapped
waves, implying that the currents were at least partially responding to remote
forcing . Noble et al . (1983) used statistics and an analytical model to suggest
that most of this remotely forced variability originated no farther away than the
eastern tip of Georges Bank .
Monthly mean velocities from long-term moorings placed on the outer shelf in the
New York Bight and offshore of the Chesapeake Bay suggest a strong low-frequency
variability, but the records are not sufficiently continuous to document a
seasonality (Mayer et al . 1979 ; Beardsley and Boicourt 1981) . The only evident
seasonality is in the deeper currents in the southern Middle Atlantic Bight,
where the currents tend to reverse during one or more winter months while the
near surface currents remain southwestward . Beardsley et al . (1985) were also
unable to confirm any strong seasonality of the volume flux past Nantucket
Shoals . This lack of seasonal change in the volume flux leads to the question
of what happens in the Georges Bank and Gulf of Maine region to filter out the
strong seasonal variation of inflow to the Middle Atlantic Bight . As Brown and
Irish (1992) point out, the reasons for this filtering are poorly understood .
84
Although the inner portions of the Middle Atlantic Bight were known to undergo
summer "reversals" of the alongshore flow in years of persistent southerly winds
and low freshwater runoff (Bumpus 1969), these reversals were thought to be
relatively rare events . Boicourt (1982) suggested that these reversals were more
the rule than the exception . An analysis of Haight's (1942) current measurements
from lightships shows that the July monthly mean flows (taken over a variety of
years) are consistently to the north, counter to the long-term mean, except at
Nantucket Shoals and in estuarine outflow plumes . If near surface flow over the
inner shelf undergoes an annual reversal while the outer shelf reverses only
occasionally, then the summertime upper-layer alongshelf flow should be banded
into two or three counterflows (see Figure 4 .2-2, from Boicourt 1982) . Williams
and Godshall's (1977) analysis of Bumpus' (1973) drifter results support this
idea and suggest that this banded structure may exist northward to New England .
Pettigrew (1981) and Churchill (1985) demonstrated that the inner-shelf current
fluctuations on time scales of several days are strongly influenced by wind
forcing .
While it may differ somewhat in scale and mode, atmospherically forced
variability of the South Atlantic Bight is typical of other shelf regions .
However, the myriad modes and the frequency and strength of interactions with the
adjacent Gulf Stream represent a singularly dominant influence, especially over
the outer shelf . Although the importance of these interactions has been known
for a long time (Bumpus 1955, 1973 ; Von Arx et al . 1955 ; Webster 1961), their
highly episodic and rapidly evolving nature has not lent itself easily to
definitive observation and analysis . The recent combination of moored
instrumentation and remote sensing from satellites has provided an observational
technique with sufficient coverage and resolution to capture these events in
progress .
The FRED Group (1989) examined eddies off North Carolina south of Cape Hatteras
and found that there were two distinct modes associated with the two Gulf Stream
positions (Figure 4 .2-3) as noted previously by Bane and Dewar (1988) . During
the small meander mode, the Gulf Stream front was along the shelf break and
frontal meanders had a smaller amplitude . When the Gulf Stream front moved
offshore, the amplitude of the frontal eddies and associated warm-water filaments
were larger . Although the Gulf Stream front was farther from shore during this
large meander mode, the eddies and filaments reached closer to shore than during
the small meander mode . The cyclonic circulations associated with the largemeander filaments were occasionally observed drawing Middle Atlantic Bight water
around Cape Hatteras and into the South Atlantic Bight (FRED Group 1989) . These
observations of the breach of the "oceanographic barrier" at Cape Hatteras
provide an additional transport mechanism to the earlier observations of winddriven events (Bumpus and Pierce 1955 ; Stefansson et al . 1971) . When these
meander crests move across the shelf south of Cape Hatteras, the wind seems to
have little effect on the shelf-water motions . During the FRED experiment, the
alternation between the large- and small-meander modes occurred on time scales
of the order three months . The six-month record was not sufficient to determine
if these alternations were associated with a seasonal progression . As a Gulf
Stream frontal eddy propagates northward, it appears to produce a front over the
continental shelf .
Although the wind does not appear a dominant force over the mid-shelf in Raleigh
Bay during the large-meander mode of the Gulf Stream, it is the dominant driving
mechanism of the mid-shelf for the remainder of the South Atlantic Bight . The
85
N
~
I
I
1
~
.
Estuary
---~
%%
Shelf
Break
1
,
~
~
;
,
;
,
~
,
1
lnner
Outer
Coastal
Jet
Figure 4 .2-2
She/f
Shelf
I
~
~
Schematic summer cross-shelf velocity profile for surface
currents in the Middle Atlantic Bight ( from Boicourt 1982) .
86
36
::
..
NOR T H
CAROLINA
,
00
v
•
.• :
34°
N
:
'
: BREAK
..• :.
'cD
C HATTERAS ::
'. SHELF
oo4
C. LOOKOUT
C. FEAR
A~
...;.•;;; ..'
.
•. ;,..
• ,.
.~
._ . • .
• . .:•: :,.; . '' D
CD
;;a ;;•
~'i . ; ~• .
.•••
.•:
: :S i
/, .
..
SMALL
MEANDER
MO DE
32° "
78°
D
760
78°
,, .
b) LARGE
MEANDER
MODE
76°W
74°
Figure 4 .2-3 Schematic diagram of the small and large meander modes observed during the FRED experiment .
Shown are the Gulf Stream front and associated frontal eddy and filament structures . F
indicates warm-water filaments ; CD, cold dome (from FRED Group 1989) .
primary response to alongshelf winds is a nearly barotropic arrested topographic
wave (Csanady 1978), as in the Middle Atlantic Bight (Lee and Atkinson 1983 ; Lee
et al . 1985 ; Lee and Pietrafesa 1987) . This barotropic response led Lee et al .
(1984) to apply a two-dimensional model, which produced a simulation of winddriven flow that agreed well with observations . Alongshelf, sea-surface slopes
determined from the momentum balance are of the order 10- 1 for both winter-spring
(Lee et al . 1984) and summer-fall (Lee and Atkinson 1983) . This value matches
the slope determined from stearic leveling by Sturges (1974) . The tendency for
offshore effects to dominate variability at the shelf break, but for the wind to
dominate at midshelf was corroborated statistically by Li and Wimbush (1985) .
Over the inner shelf, wind stress can augment or oppose the development of a
buoyancy driven coastal current (Blanton 1981 ; Schwing et al . 1983 ; Blanton and
Atkinson 1983) . The Ekman drift driven by southward wind events steepens the
nearshore front and accelerates the low-salinity coastal current . Atkinson et
al . (1983) found apparently anomalous low-salinity water off Florida in October .
They attributed this water to southward wind stress, which retained the fresh
water within the coastal band and advected it from the high-runoff region of the
Georgia and South Carolina coast . Under northward stress, the inner shelf front
flattens, and freshwater is forced offshore (Blanton and Atkinson 1983) . During
the spring runoff, sufficient amounts of fresh water are ejected offshore via
this process to be noticeable at the 40 m isobath (Atkinson et al . 1983) .
However, the transient nature of this process has hampered definitive synoptic
description and a quantitative estimate of the transport .
Cross-shelf transport of Gulf Stream water can profoundly affect the temperature,
salinity, and stratification of the middle and outer shelf . The first
recognition of such Gulf Stream influence was Bumpus and Pierce's (1955)
observation of cold, high-salinity water along the bottom of the outer shelf off
North Carolina . This process has been examined further by Webster (1961),
Stefansson et al . (1971), Blanton (1971), Blanton and Pietrafesa (1978), Atkinson
and Pietrafesa (1980), Atkinson, et al . (1980), and Hofmann et al . (1981) .
Hofmann et al . (1980) concluded that these intrusions were two-part processes .
Both an eastward movement of the Gulf Stream (creating upwelling of cold water
on the inshore edge of the Stream) and upwelling-favorable winds were necessary
to drive cold water onto the shelf . These intrusions can transport greater than
20% of the volume of Onslow Bay and have lifetimes that range from 14 to 60 days
(Atkinson et al . 1980) . Blanton et al . (1981) hypothesized that the divergent
isobaths along the shelf north of Cape Canaveral and north of the Carolina Capes
help induce upwelling as the Gulf Stream flows along the outer shelf .
The summer 1981 intrusion processes appeared similar to those observed earlier
off North Carolina in that both Gulf Stream fluctuations and wind-forced
circulations played crucial roles . However, Lee and Pietrafesa (1987) and
Hamilton (1987) show that upwelling can depend not only on Gulf Stream position
and upwelling favorable winds, but also on the presence or absence of Gulf Stream
frontal eddies . The cold, high-salinity water of these summer intrusions rapidly
increased the stratification over the middle shelf region . Wind and tidal mixing
were inadequate to mix these intrusions, so the stratification persisted for
three to six weeks (Atkinson et al . 1987) .
In winter, Gulf Stream intrusions can also rapidly change the outer shelf
stratification . In contrast to the summer case, winter intrusions occur
primarily via the surface layer, as an Ekman drift driven by northerly winds
88
(Atkinson et al . 1989) . Oey (1986) and Oey et al . (1987) presented a model of
this process (Figure 4 .2-4) that incorporated advection, mixing, and
thermodynamic effects . Oey hypothesized that in winter strong southward winds
or Gulf Stream meanders break down the shelf-break front and that these winds
drive a shoreward intrusion of warm water in the upper Ekman layer . These winter
storms then cool the shelf water and vertically mix the intruded warm Gulf Stream
water . With sufficient storm activity, a mid-shelf front appears . As might be
expected from the forcing mechanism, the mid-shelf front is highly timedependent . Atkinson et al . (1989) found occasional thermal fronts over the
middle shelf during the GALE experiment . These fronts were not dynamic, in the
sense that the salinity distribution compensated for the thermal gradient . On
the contrary, they constituted a local minimum in horizontal density gradient .
Atkinson et al . (1989) concluded that, over the outer shelf, the advective supply
of buoyancy from the warm Gulf Stream intrusions was four times greater than the
buoyancy flux from atmospheric cooling and mixing . Over the inner shelf, they
were of the same order .
4 .2 .4 Seasonal Variability
The Middle Atlantic Bight undergoes a marked seasonal change in stratification .
In winter, the water column is often vertically homogeneous over much of the
shelf, with the coldest temperatures and freshest water occurring nearshore in
February and March . The shelf-slope front separates the cooler, fresher shelf
water from the warmer, saltier slope water . In summer, the water column is
strongly vertically stratified over most of the shelf due to vernal warming and
increased freshwater runoff (Bigelow 1933 ; Bigelow and Sears 1935) . During this
season, vertical profiles on the outer shelf show large gradients in both
temperature and salinity . Temperatures decline rapidly with depth, from 25-26°C
at the surface to less than 10°C in the near-bottom cold-water band . Salinity
profiles often show a subsurface maximum (Boicourt and Hacker 1976 ; Beardsley et
al . 1976 ; Gordon and Aikman 1981) above this cold band . Beardsley et al . (1976)
concluded that this cold-water band (called the "cold pool" by Ketchum and Corwin
1964) was not stationary, as suggested by Bigelow (1933), but was moving
southward with the mean flow . Ou and Houghton (1982) provided a model for the
cold band that explained the southward progression (Figure 4 .2-5) of the
temperature minimum described by Houghton et al . (1982) . They found that an
upstream cold-water source was necessary at the onset of the heating season .
This cold water is supplied by the alongshelf flow from the Gulf of Maine, around
Georges Bank (Butman and Beardsley 1987) .
As is the case for the currents, the seasonal progression of stratification on
the South Atlantic Bight is more a variation of event climatology than the
regular progression of the Middle Atlantic Bight (Atkinson et al . 1983) . Over
the inner shelf, runoff is sufficient to maintain a salinity-controlled
stratification for most of the year, with peak stratification in spring (Blanton
1981 ; Blanton and Atkinson 1983) . The middle shelf develops a more predictable
seasonal progression, with vertically homogeneous conditions during winter and
a stratified, two-layer system during July-September (Atkinson 1985) . However,
both the middle and outer portions of the shelf are prone to rapid stratification
changes induced by Gulf Stream interactions . In addition, the stratification of
the mid-shelf can be affected by southerly wind events forcing coastal, low
salinity water offshore (Blanton and Atkinson 1983) .
89
Cooling
Wind mixing
-Mixed-Layer
depth x 100m
;
COOL ~
Shelf break front--4
I
-7
A
i
~
WARM
~
rong Southward Wind
COOL
. P, ..-m Ekman
.
Intrusion---,
.- % '**
~
~
Flux
; WARM
~,
B
DownwelUng
cool water=
COOL
T
Downward
turbulent diffusion
~
~
WARM
~ -i
Continental Shelf
C Front
~
Figure 4 .2-4 Oey et al .'s (1987) schematic diagram of the shoreward intrusion
of warm Gulf Stream water during winter on the South Atlantic
Bight, and the development of the continental shelf front .
90
-1
SEP
AUG
0
10
2
~
~
/
8
JUL
JUN
6
~
MAY
APR
MAR
4
/
~
~
~
~
/
4
/
8
~
4
00 KM
2
Z
8
I
4--
~
Figure 4 .2-5 Progression of cold-band
temperature
( minimum temperature
measured in the cold band, from Nantucket Shoals ( NS) ; to Montauk
Point ( MP), Hudson Canyon ( HC), Cape May (CM), and Cape Charles
(CC) (from Houghton et al . 1982) .
91
4 .3 VirAz in ia Coastal Water Intrusio
Temperature and salinity distributions led Parr (1933), Bigelow (1933), and
Bigelow and Sears (1935) to refer to the oceanographic "barrier" of Diamond
Shoals, restricting exchange between the southern Middle Atlantic Bight and
Raleigh Bay of the South Atlantic Bight . Bumpus and Pierce (1955), employing not
only hydrographic but also planktonic-indicator tracers (chaetognaths), witnessed
a breaching of this barrier during strong northeast winds . Cold, low-salinity
water devoid of chaetognath indicators was driven into Raleigh Bay from the
Middle Atlantic Bight . Stefannson, et al . (1971) also detected wind-driven
intrusions of Virginia Coastal Water into Raleigh Bay . These intrusions tended
to move inshore and form a narrow (20 km) southwestward current along the coast .
Stefansson et al . (1971) concluded that these low-salinity intrusions have a much
greater influence on the circulation of Raleigh Bay than do the rivers entering
south of Cape Hatteras .
The implied magnitudes of the intrusions observed by Bumpus and Pierce (1955) and
Stefannson et al . (1971) were not inconsequential to the circulation of Raleigh
and Onslow Bays . Although a quantitative estimate of the intrusions would be
beneficial for an assessment of their role in the circulation north of the
Carolina Capes, at present even an estimate of frequency of occurrence is
lacking .
Vukovich (1974) noted the occurrence of one of these events in April 1971 . The
data published by the FRED Group (1989) indicated the presence of a similar
intrusion in May 1987 . The 1971 event was detected immediately after an
atmospheric cold front had passed through the coastal region and was located off
the coast . The winds at the coast were from the north at about 3 m s'1, whereas
the winds off the coast were from the northeast with speeds as large as 15 m s'1 .
At this time, the western boundary of the Gulf Stream at Cape Hatteras had moved
seaward, allowing the cold Virginia Coastal Water to flow around Diamond Shoals
into Raleigh Bay . Lenses of cooler water were also noted in Onslow Bay at the
same time, and it was suggested that these cold lenses were relics of earlier
intrusions of Virginia Coastal Water that had previously passed through Raleigh
Bay .
4 .4
Surface Wave Climato
4 .4 .1 Introduction
Ocean surface waves off Cape Hatteras vary seasonally in a manner typical of midlatitude regions and exhibit spatial and temporal variations that are related to
meteorological conditions, local coastal effects, and the effects of the nearby
Gulf Stream . Because of different meteorological conditions which generate winddriven waves in the region, as well as needs for wave information in both
scientific investigations and practical applications such as offshore oil and gas
activities, wave climatology is best described in terms of typical and extreme
wave conditions .
Typical wave conditions are those that usually occur as a result of prevailing
meteorological conditions (see section 4 .1 .1) . Monthly, seasonal, or annual wave
statistics are generally used to quantify typical wave conditions . Additional
information, such as characteristics of waves generated by the comparatively
infrequent severe extratropical storms, tropical storms, and hurricanes, is
92
The severity and
needed to describe extreme wave conditions quantitatively .
recurrence of these strong, wave-forcing meteorological conditions in and
adjacent to the Cape Hatteras area make them important considerations for the
local wave climate .
Climatological summaries of meteorological and oceanographic data, including
waves, from buoy stations with approximately three or more years of data are
available ( U .S . Department of Commerce and National Oceanic and Atmospheric
Administration 1990) . Data are archived at the NOAA National Oceanographic Data
Center ( NODC) . Selected parameters are archived in formats similar to ship
observation formats at the NOAA National Climatic Data Center ( NCDC) .
4 .4 .2 Typical Wave Conditions
Typical or commonly occurring wave fields are generated by prevailing winds
driven by seasonally varying atmospheric pressure systems . Off the coast near
Cape Hatteras, waves with directions from the coast are usually fetch-limited and
are unlikely to have large heights . During summer, waves are most often from the
east through southeast due to prevailing clockwise wind flow around the Bermuda
atmospheric high pressure region . During winter, waves are most often from the
northeastern quadrant due both to counterclockwise wind flow around the Iceland
atmospheric low pressure region and to storms, which are not necessarily extreme,
in the northern Atlantic Ocean . Toward the eastern ( seaward) side of the region,
which is far enough from the coast so that fetch limitations are less important,
waves have substantial probabilities of being from south to southwest during
summer and from west to northwest during winter . As at most mid-latitude
locations, spring and fall are transition seasons for atmospheric conditions
which control the large-scale wind field and hence wave fields . Because of
coastal fetch limitations and shallow water effects, prevailing wave heights and
periods increase with distance offshore .
The recent Corps of Engineers wave hindcast ( Hubertz et al . 1992) provides a good
quantitative summary of typical wave conditions which are consistent with
seasonal wave patterns from ship observations and seasonal climatological
statistics from available measurements, (U . S . Department of Commerce NDBC 1991) .
Mean hindcast wave heights by month and year are listed in Table 4 .4-1 for the
outermost location ( 36°N, 75°15'W) with the deepest water depth ( 37 m) near the
north-south midpoint of the Cape Hatteras offshore region . Except in very
shallow water, wave conditions at this location should be reasonably
representative of those over the continental shelf within the region . The total
number of occurrences equals the total number of output times ( every three hours)
during the twenty year hindcast . Each wave height on which these and other
hindcast results are based was calculated from the total variance in a wave
spectrum and closely approximates the value that would be obtained from the
definition : significant wave height is the average of the highest one-third waves
in a wave record ( Longuet-Higgins 1952) .
Tables 4 .4-2, 4 .4-3, and 4 .4-4 provide monthly occurrences of wave height, peak
wave period, and peak wave direction, respectively, at the same location . Peak
wave period is the period with the most wave variance in a hindcast wave spectrum
and represents the periods of the larger waves that would occur in a wave record .
Similarly, peak wave direction is the direction with the most wave variance in
a hindcast directional wave spectrum and represents the directions of the larger
waves that would occur in a wave record . Expected winter, summer, and transition
93
Table 4 .4-1
Mean Wave Height (m) by Month and Year
(after Hubertz et al ., 1992)
~
Year
Jan
Feb
Mar
Apr
May
Jun
1956
1957
1958
1959
1960
1961
1962
1963
1964
1965
1966
1967
1968
1969
1970
1971
1972
1973
1974
1975
2 .36
1 .44
1 .82
1 .69
1 .58
1 .52
1 .68
1 .67
1 .88
1 .76
2 .06
1 .40
1 .91
1 .93
1 .58
1 .76
1 .39
1 .67
1 .10
1 .56
1 .23
1 .47
1 .93
1 .46
2 .02
1 .59
1 .56
1 .80
1 .65
1 .65
1 .47
1 .75
1 .83
2 .70
1 .62
1 .41
1 .68
1 .75
1 .50
1 .25
1 .41
1 .40
1 .60
1 .46
1 .86
1 .27
2 .06
1 .38
1 .44
1 .17
1 .30
1 .38
1 .48
1 .74
1 .26
1 .69
1 .34
1 .65
1 .28
1 .77
1 .33
1 .22
1 .74
1 .27
1 .21
1 .32
1 .29
1 .21
1 .18
1 .18
1 .00
1 .48
1 .21
1 .28
1 .53
1 .17
1 .19
1 .26
1 .12
1 .49
1 .00
0 .94
1 .13
0 .83
1 .28
1 .17
1 .01
1 .14
1 .29
0 .82
1 .07
1 .24
0 .92
1 .20
1 .12
1 .10
1 .65
1 .08
0 .92
0 .89
`
1 .11
1 .08
0 .99
0 .75
1 .16
1 .07
1 .00
0 .90
0 .86
0 .98
1 .09
1 .06
0 .71
0 .85
1 .05
0 .76
1 .04
0 .89
0 .76
0 .94
0
0
0
1
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
1
Mean
1 .69
1 .67
1 .50
1 .28
1 .09
0 .95
Jul
Aug
Sep
Oct
Nov
Dec
Mean
.93
.82
.86
.01
.91
.75
.92
.77
.97
.90
.94
.95
.56
.79
.90
.97
.84
.76
.59
.13
0 .93
0 .99
1 .04
0 .85
0 .92
0 .94
0 .96
0 .72
0 .95
0 .75
0 .83
0 .85
0 .61
0 .93
0 .94
1 .00
0 .87
0 .72
0 .66
0 .58
1 .26
1 .03
1 .01
1 .10
1 .24
1 .33
1 .12
1 .33
1 .48
1 .07
1 .08
1 .45
0 .83
1 .22
0 .90
1 .28
1 .33
0 .89
1 .04
1 .12
2 .03
1 .46
1 .86
1 .30
1 .22
1 .48
1 .29
1 .68
1 .31
1 .38
1 .34
1 .11
1 .11
1 .67
1 .73
1 .44
1 .56
1 .16
1 .18
1 .25
1 .42
1 .45
1 .18
1 .66
1 .24
1 .48
2 .65
1 .46
1 .22
1 .29
1 .68
1 .31
1 .26
1 .68
1 .61
1 .42
1 .52
1 .23
1 .41
1 .43
1 .21
1 .64
1 .70
1 .60
1 .54
1 .23
2 .18
1 .42
1 .66
1 .19
1 .63
1 .62
1 .76
1 .96
1 .49
1 .76
1 .41
1 .70
1 .53
1 .77
1 .35
1 .25
1 .40
1 .25
1 .35
1 .26
1 .48
1 .29
1 .32
1 .18
1 .29
1 .30
1 .18
1 .49
1 .31
1 .32
1 .32
1 .23
1 .09
1 .26
0 .86
0 .85
1 .16
1 .43
1 .48
1 .60
Table 4 .4-2
Occurrences of Wave Height (m) by Month for all Years
(after Hubertz et al ., 1992)
Height (m)
ko
Ln
Jan
Feb
Mar
Apr
May
Jun
Jul
Aug
Sep
Oct
Nov
Dec
Total
0 .00-0 .49
45
0 .50-0 .99 1070
1 .00-1 .49 1517
1 .50-1 .99
900
486
2 .00-2 .49
2 .50-2 .99
365
3 .00-3 .49
268
3 .50-3 .99
142
4 .00-4 .49
75
4 .50-4 .99
31
5 .00-5 .49
26
5 .50-5 .99
11
6 .00-6 .49
10
6 .50-6 .99
9
5
7 .00-7 .49
7 .50-7 .99
8 .00-8 .49
8 .50-8 .99
9 .00-9 .49
9 .50-9 .99
10 .00-Greater
57
1019
1236
858
573
358
180
93
51
19
24
23
20
5
2
2
103
1313
1462
961
531
284
158
80
31
13
6
5
3
3
2
1
2
2
154
1538
1725
750
325
129
77
55
31
12
3
1
97
2275
1755
513
165
89
39
12
7
3
4
1
170
2831
1357
268
91
31
20
15
12
2
1
222
3195
1341
143
27
13
4
3
4
2
1
2
275
3227
1138
232
71
10
3
2
2
2
3
78
2373
1302
476
300
140
74
33
11
7
4
1
1
93
1495
1631
727
450
247
137
109
34
17
15
4
1
114
1297
1470
901
501
236
133
62
29
14
11
7
3
3
4
4
7
4
91
1177
1425
950
554
302
208
124
58
34
15
13
7
2
1499
22810
17359
7679
4074
2204
1301
730
345
154
110
68
45
27
13
7
9
6
0
0
0
Total 4960 4520 4960 4800 4960 4800 4960 4960 4800 4960 4800 4960 58440
Table 4 .4-3
Occurrences of Peak Wave Period (s) by Month for all Years
(after Hubertz et al ., 1982)
%0
oh
Period (s)
Jan
Feb
Mar
Apr
May
Jun
3 .0- 3 .9
4 .0- 4 .9
5 .0- 5 .9
6 .0- 6 .9
7 .0- 7 .9
8 .0- 8 .9
9 .0- 9 .9
10 .0-10 .9
11 .0-11 .9
12 .0-12 .9
13 .0-13 .9
14 .0-14 .9
15 .0-15 .9
16 .0-16 .9
17 .0-17 .9
18 .0-18 .9
19 .0-19 .9
20 .0-longer
85
654
795
801
775
533
342
313
314
199
70
47
10
12
2
3
2
120
618
745
796
689
419
297
276
211
185
110
34
13
6
169
771
914
834
727
479
335
214
231
129
86
37
15
14
4
1
193
994
859
692
521
411
357
323
198
127
81
41
3
153
966
821
644
660
800
516
210
105
43
20
16
4
2
174
1003
761
906
944
597
244
137
31
3
4960
4520
4960
4800
4960
4800
Total
1
Jul
Aug
Sep
Oct
Nov
Dec
Total
236 188
1330 956
707 737
551 925
973 1101
703 570
340 300
82
85
15
56
4
25
7
10
9
4
3
3
67
501
661
690
1099
703
518
298
201
39
13
118
497
719
817
929
790
496
268
167
97
46
16
101
638
803
749
648
540
315
313
263
208
134
62
17
2
3
2
2
113
705
721
741
666
515
485
396
282
176
93
42
16
7
2
1717
9633
9243
9146
9732
7060
4545
2915
2074
1235
670
308
90
47
14
6
5
0
4960 4960
4800 4960 4800
4960
58440
6
4
Table 4 .4-4
Occurrences of Peak Wave Direction (deg) by Month for all Years
(after Hubertz et al ., 1982)
Direction (deg) Jan
~
348 .75- 11 .24
11 .25- 33 .74
33 .75- 56 .24
56 .25- 78 .74
78 .75-101 .24
101 .25-123 .74
123 .75-146 .24
146 .25-168 .74
168 .75-191 .24
191 .25-213 .74
213 .75-236 .24
236 .25-258 .74
258 .75-281 .24
281 .25-303 .74
303 .75-326 .24
326 .25-348 .74
698
317
283
623
320
563
139
187
190
141
175
129
172
134
293
596
Feb
Mar
Apr
May
Jun
Jul
Aug
Sep
Oct
Nov
Dec
Total
449
268
359
607
306
240
281
267
222
124
138
132
187
153
331
456
436
216
290
676
279
351
290
295
373
243
210
104
185
165
359
488
353
142
262
596
344
369
301
353
430
274
301
213
193
156
265
248
145
196
277
574
593
840
455
452
393
285
267
131
125
63
90
74
88
95
156
296
346
1026
708
515
558
355
289
152
72
50
40
54
108
80
98
144
261
1630
429
236
514
524
462
185
104
55
52
78
136
146
259
280
308
1519
675
286
475
262
284
107
71
41
45
66
239
362
401
887
745
1162
309
194
92
90
90
41
30
29
40
89
393
403
386
877
810
665
347
169
140
73
80
65
96
63
157
236
482
229
331
845
526
400
215
204
221
157
142
112
127
138
293
378
511
253
294
627
541
403
211
216
289
155
149
138
204
163
334
472
4038
2707
3396
7032
5379
9168
4360
3374
3897
2683
2587
1509
1566
1210
2299
3235
Total 4960 4520 4960 4800 4960 4800 4960 4960 4800 4960 4800 4960 58440
season behavior is clearly seen in these tables . Larger occurrences of waves
from the northeast quadrant during October through March and the higher winter
wave heights indicate the importance of extratropical storms which often develop
and intensify off the U .S . east coast particularly from Cape Hatteras northward .
A significant gradient of more severe wave conditions with increasing distance
from the coast occurs between November and March when extratropical storms
produce the higher waves . Wave heights near 75°15'W are roughly 20% higher than
wave heights near 75°45'W, a location about 44 km closer to the coast . Except
very near the coast, offshore variations are smaller during other months,
especially summer months .
Additional typical wave information from this hindcast at several locations in
the region of interest includes joint occurrences of wave height and peak period
by direction intervals and for all directions, occurrences of wind speeds by
month, and occurrences of wind direction by month (Hubertz et al . 1992) .
Hindcast information pertaining to extremes is described in the following
section .
4 .4 .3 Extreme Wave Conditions
Extreme waves are generated by severe extratropical storms as well as by tropical
storms and hurricanes . The region is well-known as an area where rapid
extratropical storm intensification occurs, especially near the Gulf Stream
during the cool season . These storms, which often develop into so-called
northeasters, may generate high waves that persist for relatively long time
periods, sometimes several days . Such persistent high waves may cause
substantial beach erosion along the coast and are usually a primary concern in
the design of offshore structures .
The wave hindcast of Hubertz et al . (1992), which provides quantitative
information about typical wave conditions, also provides useful information about
extratropical, storm-generated extreme waves . Table 4 .4-5 lists the maximum wave
heights, associated peak wave periods, and associated peak wave directions for
each month during the twenty year hindcast at the same location (36°N, 75°15'W)
used for the typical wave condition tables . Highest waves occur during November
through March . The highest single wave height (8 .8 m) during this hindcast was
generated by the very severe and well-studied March 1962 northeaster that caused
major damage along the U .S . east coast .
Return periods, or recurrence intervals, are often used to describe extreme
waves . The return period corresponding to a given wave height is the average
time between wave height occurrences (which are quasi-random) equal to or greater
than this wave height . For all Corps of Engineers hindcast output locations in
the region, Table 4 .4-6 lists wave heights for return periods between two and
fifty years that were obtained by fitting a Fisher-Tippett Type II extremal
probability distribution to hindcast wave heights (Hubertz et al . 1992) . A Type
I distribution provided somewhat lower (less conservative) wave heights at the
longer return periods . Due to coastal fetch limitations for waves from the
north, refraction effects, and further intensification of storms moving offshore,
extreme waves are higher further offshore in the eastern part of the region than
near the coast .
98
Table 4 .4-5
Maximum Wave Heights (m*10) with Associated Peak Periods (s) and Directions (deg/10) by Month and Year
(after Hubertz et al ., 1992)
Year
1956
1957
1958
1959
1960
1961
1962
1963
1964
1965
1966
1967
1968
1969
1970
1971
1972
1973
1974
1975
Jan
7211 2
36 835
40 814
411018
6413 8
38 836
5411 4
45 833
5210 9
48 935
7413 4
34 832
6912 5
5210 .4
46 933
45 831
39 9 2
34 9 3
26 8 1
5310 3
Feb
29 819
4510 3
41 831
29 734
6514 7
441012
5910 3
5410 2
4910 2
461014
44 936
38 8 1
7913 5
6412 .6
49 918
431016
35 815
56 9 1
37 818
33 910
Mar
41 9 1
35 833
37 836
501116
43 832
32 732
8814 5
47 917
39 835
30 735
46 9 4
39 834
48 919
56 935
33 917
7112 3
27 718
421018
28 810
591113
Apr
51 9 2
36 833
5611 8
37 8 5
35 836
34 815
381018
37 818
30 819
30 718
22 631
41 835
33 917
35 818
42 9 1
47 911
42 9 0
35 819
32 819
48 936
May
Jun
Jul
30 8 3
19 7 2
28 8 0
17 719
3212 9
32 732
18 618
27 820
4010 8
17 636
26 7 0
34 818
33 9 8
23 720
23 717
29 819
5711 5
35 818
26 817
17 636
26 8 7
4310 5
22 6 3
24 631
31 8 1
20 718
33 912
39 814
18 5 3
26 9 7
44 9 1
19 6 3
13 422
20 717
22 7 4
13 6 4
391013
16 618
17 616
6912 4
16 611
16 6 1
14 423
29 717
37 819
13 519
28 813
15 618
23 720
20 719
23 635
18 618
12 514
17 618
19 718
28 7 7
19 614
16 910
12 536
6914 5
Aug
21
25
28
22
24
19
25
15
23
18
21
22
19
23
32
44
22
18
19
10
633
813
7 8
633
7 4
617
814
521
718
634
7 8
719
7 3
735
8 7
914
7 3
619
636
5 5
Sep
44 9 9
33 9 9
30 734
29 712
25 8 4
3811 9
6011 3
33 8 1
5111 3
28 8 6
33 8 3
36 9 3
1611 7
36 8 1
26 720
39 9 7
35 8 1
33 816
33 9 3
22 7 4
Oct
46 9 8
5512 9
6212 4
29 716
39 9 1
5410 1
4110 6
4510 7
391011
42 915
36 8 1
23 619
26 732
35 8 1
5212 6
40 912
31 8 1
33 8 1
44 9 3
45 9 2
Nov
31 733
36 918
33 8 1
38 8 1
28 732
36 835
8714 7
42 913
50 9 2
36 834
35 814
33 836
43 831
541212
421010
37 9 2
35 835
31 8 0
43 834
6011 4
Dec
41 831
46 936
6512 6
50 9 2
5310 1
32 732
6612 9
42 9 8
6010 1
42 9 0
75 833
47 918
521018
44 832
41 833
42 833
32 731
411017
491011
5210 1
Max 7413 4 7913 5 8814 5 5611 8 5711 5 6912 4 6914 5 44 914 6011 3 6212 4 8714 7 6612 9
Example :
7211 2 - 7 .2 m, 11 s, direction from 20 deg .
Max
7211 2
5512 9
6512 6
50 9 2
6514 9
5410 1
8814 5
5410 2
6010 1
48 935
7413 4
47 918
7913 5
6412 9
5212 6
7112 3
5711 5
56 9 1
491011
6914 5
Table 4 .4-6
Wave H eight (m) as a Function of Return Per iod'
(after Hubertz et al ., 1992)
hocation
Deuth (m)
Ret urn Period (y ears)
2
34 .50° W,
76 .75° N
34 .50° W, 76 .50° N
r
0
20
9+
5
10
20
25
50
4 .87
5 .33
5 .68
6 .04
6 .16
6 .55
4 .56
5 .13
5 .57
6 .03
6 .19
6 .70
34 .75° W, 76 .25° N
22
4 .93
5 .28
5 .53
5 .80
5 .89
6 .17
34 .75° W, 76 .00° N
35
5 .42
5 .81
6 .10
6 .39
6 .49
6 .79
35 .00° W, 75 .75° N
18
4 .74
5 .35
5 .82
6 .32
6 .49
7 .03
75 .50° N
35
5 .80
6 .50
7 .05
7 .63
7 .83
8 .46
35 .25° W, 75 .25° N
22
5 .81
6 .49
7 .02
7 .56
7 .75
8 .34
35 .50° W, 75 .25° N
27
6 .38
7 .23
7 .90
8 .60
8 .84
9 .61
35 .75° W, 75 .25° N
29
6 .47
7 .34
8 .01
8 .72
8 .96
9 .75
36 .00° W, 75 .25° N
37
6 .69
7 .56
8 .23
8 .93
9 .17
9 .94
36 .25° W, 75 .50° N
27
6 .01
6 .70
7 .23
7 .78
7 .96
8 .55
36 .50° W, 75 .75° N
18
5 .19
5 .85
6 .36
6 .90
7 .08
7 .67
36 .75° W, 75 .75° N
11+
4 .92
5 .71
6 .33
7 .00
7 .23
7 .98
37 .00° W, 75 .75° N
14
4 .88
5 .49
------ -----------+ wave heights possibly depth-limited .
5 .95
6 .44
6 .61
7 .14
35 .00° W,
Few available studies provide extreme wave or wave return period information
based on hurricane-generated waves . Earle (1975) employed (now outdated)
numerical wave directional spectra and wave refraction models to hindcast wave
conditions in 16m of water during fifteen hurricanes and thirty extratropical
storms between 1944 and 1973 for engineering evaluations of the Diamond Shoals
Light Tower off Cape Hatteras . Earle and Burns (1975) used the same techniques
to determine design criteria for similar evaluations of the Chesapeake Light
Tower off Cape Henry, Virginia, in water 12 m deep at mean low water . Extreme
significant wave heights (using the same definition as the Corps of Engineers
wave hindcast) and their corresponding return periods were 9 .8 m (10 years), 11 .4
m (25 years), 12 .5 m (50 years), and 13 .6 m (100 years) at Diamond Shoals Light
Tower . Corresponding results at Chesapeake Light Tower were 10 .7 m (10 years),
11 .4 m (25 years), 11 .6 m (50 years), and 11 .7 m (100 years) . Although windgenerated storm surges were considered in these extreme waves, the results may
be highly location dependent due to substantial shallow water effects . These
results indicate considerably higher extreme waves than the Corps of Engineers
hindcast partly because of the higher wind speeds that occur during severe
hurricanes compared to severe extratropical storms . However, durations of higher
waves during hurricanes are shorter than durations during severe extratropical
storms . Approximately half of the highest hindcast heights at each location were
due to hurricanes and half to severe extratropical storms .
Ward et al . (1978) used a significant wave height model to hindcast waves during
fourteen hurricanes between 1900 and 1975 that affected the Georgia Embayment
south of the study region and Baltimore Canyon north of the study region . While
these results are not directly applicable to the study region, averaging these
results for locations near the seaward edge of the continental shelf provides the
following "rough" estimates of hurricane-generated deep water extreme significant
wave heights for the region : 11 .0 m (25 years), 12 .2 m (50 years), and 13 .4 m
(100 years) .
The . U . S . Army Corps of Engineers performed a twenty year, 1956-1975, hindcast
of hurricane-generated waves for the United States Atlantic and Gulf of Mexico
coasts (Abel et al . 1989) . Extreme waves were estimated at three locations where
waves could be affected by bottom depths (e .g . wave refraction effects) and two
deep water locations within the Cape Hatteras region . Extreme wave results at
several locations where wave heights were not depth-limited, including off Cape
Hatteras, were unrealistically large . A report addendum provides results using
an improved statistical analysis technique . Averaged extreme significant wave
heights and their corresponding return periods from the addendum at the three
non-deep locations (average depth 44m) were 7 .6m (5 years), 8 .9m (10 years),
10 .2m (20 years), and 11 .8m (50 years) . Results at the two deep water locations
were 9 .3m (5 years), 11 .0m (10 years), 12 .6m (20 years), and 14 .9m (50 years) .
There are concerns about the accuracies and implications of the statistical
analysis techniques that have been used with the Corps of Engineers hurricanegenerated wave hindcast . Consideration of the available extreme wave information
shows that a major wave information shortcoming is the lack of information about
hurricane generated waves using state-of-the-art numerical wave models .
Hurricanes clearly generate the very highest waves in the region .
Long-term measurements at buoy station 41001 (located between 34°54' - 35°N and
between 72° - 73°W) since 1976 are approaching a long enough time period to
provide useful extreme wave information (although not for return periods as long
as 100 years) . The highest significant wave height measured through 1988 was
101
10 .0 m during a March, 1983, extratropical storm . Several extreme wave heights
during other storm events have exceeded 9 m illustrating the severity of waves
in the region .
4 .5
Regional Bottom Boundary Layer Processes
4 .5 .1 Introduction
Continental shelf currents are coupled frictionally to the seabed through a
bottom boundary layer that allows the flow to adjust to the underlying stationary
surface . Because the bottom boundary layer is essentially always turbulent, it
is a dynamic region in which flow energy is dissipated and heat, mass, and
momentum are vigorously mixed . Moreover, processes endemic to the bottom
boundary layer result in the exchange of sediments, organisms, and chemical
species across the fluid-sediment interface . Bottom boundary layer structure and
dynamics reflect the complex interactions among the dominant processes .
Sediment transport is a boundary layer process of great practical significance .
It plays an important role in the dispersal of pollutants and in engineering
applications such as water quality maintenance and the design of stable offshore
structures . A fundamental aspect of the sediment transport problem, and one to
which much theoretical and practical effort has been devoted, is the
determitiation of bottom shear stress, a quantity controlled by the intensity of
turbulence within the bottom boundary layer . The turbulence level, of course,
is a function both of unidirectional and wave orbital motions, which combine in
a nonlinear manner to create the near-bottom turbulence field .
On continental shelves, circulation is driven by various mechanisms, two of the
most important of which are surface winds and tides . Bottom boundary layers
associated with tidal currents are unique in that their dynamics and structure
are strongly influenced by tidally induced, rhythmic changes in flow direction
and the accelerations and decelerations associated with those changes . On many
continental shelves, however, mean circulation is relatively independent of tidal
motions, and flow is driven, primarily, by the alongshore stress of the wind at
the sea surface . The resulting bottom boundary layer flow may be unidirectional
and quasi-steady for time periods characteristic of the forcing time scale and
long in comparison to the tidal period . Recent oceanographic studies suggest
that the bottom boundary layer on the Cape Hatteras shelf is of the latter
category during a significant part of the year . Storms with their associated
strong unidirectional currents and large surface waves are especially effective
for transporting sediments and other particulate materials .
This section presents the significant aspects of boundary layer structure and
dynamics as they relate, specifically, to the wind-driven bottom boundary layer
on the Cape Hatteras continental shelf .
4 .5 .2 Synthesis and Interpretation of Observations
Fluid motions with a wide range of frequencies complicate the acquisition and
interpretation of bottom boundary layer dynamical data . Because of that and
other difficulties associated with making measurements in a harsh environment and
in proximity to a boundary, bottom boundary layer field studies have been
undertaken only recently, with the earliest systematic studies of boundary layer
turbulence going back little more than a decade . Unfortunately, none of these
102
earlier boundary layer studies were conducted on the continental shelf off Cape
Hatteras . It is possible, however, to gain a reasonable understanding of the
bottom boundary layer of this region by a judicious use of theory, observations
from bottom boundary layers of other geographic regions, and near bottom
oceanographic measurements on the Cape Hatteras continental shelf .
The oceanography of the continental shelf near Cape Hatteras is complex, being
influenced by intense, seasonably variable atmospheric phenomena and by
excursions of the Gulf Stream onto the continental shelf . In winter,
extratropical cyclones and cold fronts spawn northerly winds that create
energetic shelf currents and surface waves ranging in height from 6 m inshore to
10 m offshore (NOAA, NWS 1990) . Clearly, the effects of the largest waves reach
to the bottom at all shelf depths . Moreover, vigorous wave activity thoroughly
mixes inner shelf waters, thereby allowing wind-generated currents to penetrate
to the bottom . Both onshore (Pilkey and Field 1972) and offshore (Milliman et
al . 1972) motion of sandy shelf sediments and silty continental slope sediments
have been attributed to current and wave-related motions associated with local
and regional weather patterns .
The study by Lee et al . (1989) indicates that subtidal current variability over
the continental shelf in the vicinity of Cape Hatteras and southward is a direct
response to synoptic scale alongshore winds that recur at periods of 2-10 days .
The winds are coherent over an alongshore scale of greater than 800 km, thus the
alongshore wind-forced responses are nearly in-phase over the same scale . A
bottom boundary layer associated with wind forced currents should be coherent at
the same scale as the current .
Lee et al . (1989) have shown that the cross-shelf response to coherent alongshore
forcing is Ekman-like in the surface layer with oppositely directed return flow
in the bottom boundary layer . The nature of this type of flow has been described
by a number of investigators as Ekman frictional equilibrium response to local,
alongshore wind forcing (e .g ., Winant et al . 1987) . It is well-known that the
current vector in a bottom boundary layer in the northern hemisphere turns to the
.left as the bottom is approached . The result of this veering for the problem at
hand is onshore transport during periods of northward mean current flow and
offshore transport during periods of southward mean current flow .
A number of investigators have shown that onshore movement of the Gulf Stream can
enhance northward wind-driven flow . With sufficiently intense Gulf Stream flow,
outer shelf waters may move in opposition to those of the inner and middle shelf
driven by northerly winds . When this situation arises, principally in winter,
bottom boundary layer flow may be onshore near the shelf edge and offshore
farther inshore . The result would be the development of a zone of convergence
in the vicinity of the outer shelf . Conditions for such a situation to develop
apparently are not uncommon during winter .
Based on previous studies (Atkinson et al . 1983 ; Lee and Atkinson 1983 ; Lee et
al . 1984), the shelf south of Cape Hatteras has been partitioned into three
cross-shelf zones--inner, middle and outer--each zone characterized by a
distinctive set of controls governing the subtidal flow . The inner shelf,
occupying the depth range from 0-20 m, exhibits relatively weak flows . The inner
shelf receives fresh water from rivers and surface runoff and supports a general
baroclinic southward flow . The addition of fresh water enhances stratification
in summer and may induce weak stratification in winter .
103
The middle shelf lies between water depths of 21-40 m . The region may be
stratified in summer, but during winter periods of strong wind forcing, the water
column is well mixed from surface to bottom . Alongshore wind stress is the
dominant forcing mechanism for subtidal flow . Mid-shelf currents are somewhat
more energetic than those of the inner shelf .
The outer shelf, represented by water depths >40 m, is distinguished from the
inner and middle shelf by the frequent presence of northward propagating Gulf
Stream meanders and eddies . The water column in winter normally is well mixed ;
however, onshore motion of the Gulf Stream may result in near-surface
stratification . The most energetic currents are found in this region of the
shelf .
Bottom shear stress and boundary layer thickness exert strong controls on near
bottom transport and, as such, are parameters of great interest . Bottom stress
frequently is written as ro - p CD U2, where CD is a bottom drag coefficient and
U is near bottom vector velocity . Using this expression together with values of
U and CD presented by Lee et al . (1989, Table 10), yields bottom shear stress
values of 1 .3, 1 .2, and 1 .8 dynes/cmz for the inner, middle, and outer shelves,
respectively . It is important to note that these estimates do not incorporate
the important effect of waves .
Using the data cited above, shelf steady-state boundary layer thickness for the
shelf as a whole is estimated to be about 13 m . It should be noted that the
velocities upon which this estimate is based were measured at heights of 3 m
above the bed and, thus, may not be representative of flow in the logarithmic
layer whose thickness is only about 1/10 of the bottom boundary layer thickness
(Weatherly, 1972) .
The intensity of bottom boundary layer turbulence, generally, is mirrored in the
bottom sediments over which the boundary layer flows . Early studies of Cape
Hatteras continental shelf sediment distribution reveal the shelf in the region
to be blanketed largely by noncohesive sediments of sand-size and coarser (e .g .,
Milliman et al 1972) . The sediments in this region, like shelf sediments in
general, have been classified as relict, i .e ., sediments deposited during an
earlier and lower sea-level stage . The implication is that modern river-derived
sediments do not remain on the shelf but are returned to estuaries or transported
across the shelf to the continental slope and beyond (Milliman et al . 1972) .
Estimates of bottom shear stress of greater than 1 dyne cm -2 across the shelf
are characteristic of environments where coarse sediment is the dominant
component .
Recent work by Weston (1988, Figure 2) has shown a sediment distribution on the
middle to outer shelf seaward of Cape Hatteras characterized by a diversity of
sediment types ranging from coarse to very fine sands . Cohesive components (silt
and clay) were absent over much of the area sampled and were never more than 27%
at any location . The data of Weston also indicate an alongshelf gradation of
sediment size with particle size generally increasing to the north . Cross-shelf
gradation exists as well, but is much less obvious . The presence of a zone of
silt and clay in an area dominated by much coarser sediments suggests
significantly different depositional environments in close juxtaposition in an
alongshelf direction . The reason for this difference is not obvious but may be
related to the presence of a convergence zone due to aperiodic current reversals
along the outer shelf . Sediment stratification of the boundary layer may be
104
important in that area during energetic resuspension events but is probably
unimportant in those areas where sand is the dominant size component .
4 .6
Sediment Transport
Prior to the mid-1970s most field investigations of marine sediments were
conducted from aboard ship using mechanical devices to collect bottom sediment
and/or water samples . Analysis of these samples typically focussed on the
composition of the bottom sediments and the extent to which they were relic or
recently deposited . Inferences regarding contemporary sediment dynamics were
generally deduced from clues found in the bottom sediment . Advances in
underwater instrument technology during the 1970s greatly expanded the
capabilities of those studying marine sediments, principally by allowing them to
measure continuously both water velocity and turbidity near the sea floor . This
occurred at a time when government agencies took an interest in the motion of
fine-grained marine particles, due largely to their importance as carriers of
contaminants introduced into the coastal ocean . The "funding atmosphere" and
advances in instrumentation combined to shape several intensive field studies of
fine-grain sediment motion over the shelf and slope off the U .S . east coast
during the late 1970s and the 1980s . Unfortunately for the purposes of this
report, sediments of the Cape Hatteras region were largely ignored during this
period . Nevertheless, using the limited number of measurements from this region
(collected mostly during the "strictly shipboard sampling" stage of marine
sediment research), together with findings from studies in adjacent areas, a
reasonably coherent, though somewhat tenuous, picture of marine sediment dynamics
near Cape Hatteras has been constructed .
In this section processes which act to mobilize fine sediments in the Cape
Hatteras region are considered, and how they are influenced by the physiographic
characteristics of the area will be described . Both shelf and slope sediments
will be dealt with . Also to be considered is transport of sediment across the
shelf . As a necessary prerequisite, the bathymetry and bottom sediment
properties of the Cape Hatteras region will first be described briefly .
4 .6 .1 Bathymetric Setting
The most prominent bathymetric features of the North Carolina shelf are
large-scale shoals extending from the barrier island capes (Figure 1 .1-1) . Each
of these shoals stretches across most of the shelf and rises above the
surrounding seafloor by a height of 20 m or more (Swift et al . 1972) . Embedded
in the shoals are numerous ridge and swale features . The largest of these have
a vertical, crest-to-trough, relief of 10-20 m and are oriented transverse to the
shoal's main axis (Figure 4 .6-1) .
Ridge and swale bathymetry is by no means peculiar to the cape shoals . It
dominates the relief of most of the shelf off the U .S . east coast and occurs on
a variety of scales (Duane et al . 1972 ; Swift et al . 1973 ; McKinney et al . 1974 ;
Stubblefield et al . 1975) . The most prominent ridge and swale features of the
North Carolina shelf have crest-to-crest spacings of 1-10 km and crest-to-trough
vertical extents of 1-10 m (Macintyre and Milliman 1970 ; Duane et al . 1972 ; Swift
et al . 1973) .
The character of the bottom relief changes abruptly seaward of the shelf off of
Cape Lookout Shoals . To the north of the shoals, the shelf is connected with a
105
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40
Figure 4 .6-1 Bathymetry of Cape Lookout Shoals showing ridge and swale
features transverse to the shoal's axis (from Swift et al . 1972) .
106
steeply inclined continental slope . This is incised by numerous submarine
canyons and smaller gullies, as revealed in detail by long-range sidescan sonar
surveys (Twichell and Roberts 1982 ; Scanlon 1984) . Extending seaward of the
shelf-edge south of Cape Lookout Shoals is the Blake Plateau . Compared with the
continental slope to the north, the Blake Plateau is gently inclined, and its
depth contours are smooth and noticeably without evidence of canyons . The
plateau extends to a depth of 1500 m before giving way to the steeply inclined
Blake-Bahama Escarpment .
4 .6 .2 Sedimentological Setting
Like many other properties, the chemical composition of shelf sediments undergoes
a transition in the Cape Hatteras region . Shelf sediments south of Cape Hatteras
tend to be rich in calcium carbonate and have a low mineral content, while the
reverse is true to the north (Milliman et al . 1972) . In spite of this trend, the
textural character of shelf sediments shows no noticeable north-to-south
variation in the Cape Hatteras region . Rather, shelf sediments from Florida to
the Hudson Canyon are predominately well sorted sand (Stetson 1938 ; Hollister
1973) . Except in a few isolated areas, the silt-plus-clay (particle diameter
<0 .0625 mm) fraction of these sediments is small, typically less than 10%
(Hathaway 1971) . Nonetheless, significant small-scale variations in this
fraction are prevalent over the entire shelf (Hathaway 1971 ; Stubblefield et al .
1975 ; Stanley and Wear 1978 ; Boesch 1979) . There is a fairly convincing body of
evidence which indicates that these variations are related to the ridge and swale
bathymetry of the shelf . Evidence from the Cape Hatteras region consists of
bottom photographs which show an abundance of rock fragments mixed with very
little fine material on ridges, and a sand-mud bottom within swales (Macintyre
and Milliman 1970) . More quantitative evidence has been obtained further to the
north . In separate studies, Stubblefield et al . (1975) and Boesch (1979)
examined bottom samples collected from densely spaced sites over the shelf
offshore of New Jersey and Virginia . They found sediment silt-plus-clay contents
to be highest (up to 17%) within the swales and lowest (generally less than 2%)
near the ridge crests .
Unlike shelf sediments, the textural character of surface sediments seaward of
the shelf-edge changes dramatically in the Cape Hatteras region (Hathaway 1971 ;
Hollister 1973) . Sediments over the slope north of Cape Lookout Shoals generally
contain very little sand and have a high silt-plus-clay content, exceeding 80%
in many areas .
By contrast, sediments over the northern Blake Plateau closely
resemble shelf sediments in that they are primarily well-sorted sand mixed with
a small fraction of silt-plus clay .
4 .6 .3 Processes Affecting Sediment Movement
Examination of field measurements has identified several mechanisms which
contribute to sediment resuspension off the U .S . east coast . Those likely to be
prominent in the Cape Hatteras region are storm-induced motions, tides, internal
waves, Gulf Stream currents and the action of towed fishing gear . As discussed
below, the relative importance of each of these is expected to vary as a function
of time, water depth and geographic location .
107
4 .6 .3 .1 Storms
Storm-related winds generate both large-scale ocean currents and the oscillating
motions of surface gravity waves . These two types of currents combine on
continental shelves to exert stress on bottom sediments . The potential
importance of wave-current interaction in generating stress at the seafloor was
first recognized by theoretical studies, notably those of Smith (1977) and Grant
and Madsen (1979) . The basic idea put forth by Grant and Madsen (1979) is that
when considering current-induced stress, the benthic zone can be divided into
nested boundary layers . In a layer extending a few centimeters above the bottom,
surface gravity wave and large-scale currents combine nonlinearly to generate
stress . Later observational studies supported this concept and showed that
surface-wave currents often dominate bottom stress production in shallow
continental shelf waters during storms (Cacchione and Drake 1982 ; Clark et al .
1982 ; Young et al . 1982 ; Grant et al . 1984 ; Cacchione et al . 1987 ; Churchill et
al . 1988, 1992 ; Lyne et al . 1990a,b) .
The effect of storm-induced currents on sediments in the Cape Hatteras region was
examined by Rodolfo et al . (1971) using water samples collected at various sites
going across the shelf near Cape Lookout Shoals . The samples were acquired on
three occasions : one during a period of strong winds and rough seas in the wake
of Hurricane Gerda (which passed the area roughly one day before) and the others
during times of relatively light breezes and calm seas roughly one week before
and one week after the hurricane's passage . During the storm, sediment
concentrations found in the samples (Figure 4 .6-2) show an increase in suspended
mass on the middle and inner shelf . Over the outer shelf, suspended sediment
concentrations found during the storm did not differ appreciably from pre- and
post-storm concentrations .
Limits of the seaward extent of storm-induced sediment resuspension off the U .S .
east coast were revealed in two later studies carried out using current meters
and optical turbidity sensors (transmissometers) moored near the seafloor
(Churchill et al . 1988, 1992) . One study was located south of Cape Cod in a
region of fine-grained bottom sediment, known locally as the Mud Patch . The
other was conducted over the shelf east of the Delmarva Peninsula, where the
bathymetric features and sediment grain size characteristics are similar to those
of the Cape Hatteras shelf region . Data from this study showed frequent episodes
of storm-induced sediment resuspension, occurring at a rate of roughly once per
week, at a site on the 42 m isobath . By contrast, only three such episodes were
observed over a seven month period at a site directly seaward on the 90 m
isobath . The suspended solids concentrations detected at the 90 m site during
these episodes, which were particularly vigorous storms, were several times less
than the suspended solids concentrations typically observed at the 42 m site
during storms . The seaward decline in the frequency and intensity of
storm-induced resuspension in this study and at the Mud Patch was found to be
primarily a result of the attenuation in surface gravity wave current magnitude
with increasing water depth .
4 .6 .3 .2 Internal Waves
Displacements of the internal density surfaces of a stratified fluid may
propagate as free waves . These are called internal waves and are prevalent
throughout the world's oceans . They are confined to a well-defined frequency
band . Its lower limit equals or is close to the Coriolis frequency, which varies
108
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wrura .a.
O
.'nstl+
•
'
•
. .
'
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t
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Min:mNn distonce (rom Gerdas. ctnter dm)
too
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un 1 Nww ts .
.k•~
120
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20
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10
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Distance from shore (kn )
N
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s0
~.
Figure 4 .6-2 Locations of bottle samples taken offshore of Cape Lookout during September 1969 and the
sediment concentrations found in the samples versus bottom depth . Approximate times of
sampling phases I, II and III are, respectively, one week before, one day after and one week
after Hurricane Gerda made its closest approach to the sampling area . Note that concentrations during Phase II are exceptionally high over the middle and inner shelf except at
station 750 which is believed to have been in the lee of the cape (adapted from Rodolfo et
al . 1971) .
with latitude and is 0 .048 cycles per hour (cph) at Cape Hatteras . Its upper
limit is the Brunt-Vaisala frequency, which depends upon the vertical density
gradient . Brunt-Vaisala frequencies found over the U .S . east coast continental
shelf during the summer months typically peak at values on the order of 10 cph
(Flagg 1988) . Internal waves may be generated by a host of mechanisms including
variations in surface wind stress and flow over abrupt topography (Briscoe 1975) .
Their propagation is strongly influenced by bottom slope and water column
stratification (Wunsch 1969) . Recent evidence indicates that motions within the
internal wave band contribute to strong near-bottom flows and sediment
resuspension over the shelf-edge and upper slope north of Cape Hatteras .
The potential importance of internal wave currents at the outer shelf was
demonstrated by Csanady et al . (1988) who examined near-bottom velocity records
taken at numerous locations over the Middle Atlantic Bight shelf and slope . As
part of their analysis, total velocity variances were separated into a number of
frequency bands and plotted against bottom depth . The plots (Figure 4 .6-3)
suggest that internal waves may be generated and/or intensified near the Middle
Atlantic shelf edge . Variance in the two highest bands, which contain
frequencies greater than 0 .066 cph and encompass most of the internal wave band,
is greatest over the outer shelf and upper slope . (It should be noted that the
range of these bands includes the frequency of the semidiurnal tide, which may
propagate as a surface or an internal wave) .
Detailed analyses of high frequency currents and their influence on bottom
sediments over the outer shelf were undertaken by Flagg (1988) and Churchill et
al . (1992) using measurements collected south of Long Island and east of
Virginia . Taken together, their findings reveal a factor of two increase in
near-bottom supratidal (periods less than 10 hr) current strength going seaward
across the outer shelf (roughly from the 80 to the 130 m isobath) . The magnitude
of this increase does not appear to vary significantly with season or with
location along the shelf . Flagg (1988) offered convincing arguments that this
cross-shelf change in supratidal kinetic energy results from dissipation of
internal waves propagating onto the shelf . Churchill et al . (1992) found that
vigorous supratidal motions at the outer shelf contributed to several events of
sediment resuspension at a site on the 131 m isobath (Figure 4 .6-4) . They also
noted that these currents were not, by themselves, of sufficient vigor to
initiate sediment motion at this site ; modestly strong subtidal currents were
also required . However, the data examined by Csanady et al . (1988) shows that
the intensity of high frequency motions continues to increase going seaward
across the Middle Atlantic Bight outer shelf and reaches a maximum just beyond
the shelf edge (Figure 4 .6-3) . The possibility that high frequency motions alone
may be sufficient to resuspend sediments just seaward of the Middle Atlantic
Bight shelf edge is suggested by near-bottom speed records taken close to the 200
m isobath at sites off Cape Cod and New Jersey (Figure 4 .6-5) . These exhibit
numerous spikes due to high frequency currents, many of which exceed 30 cm s-1 .
It is thus clear that high-frequency motions, presumably due to internal waves,
can have an important role in mobilizing shelf sediments of the Middle Atlantic
Bight . At this time we can make no statement of this type regarding the South
Atlantic Bight because little is known about its internal wave climate .
110
TOP06RAPHIC WAVE
90
60
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0 SEEP Summer
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V MASAR South Line
+ NASACS
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1000 2000 3000
BOTTOM DEPTH(M)
Figure 4 .6-3 Total variance of velocities measured within 100 m of the bottom
at locations in the Middle Atlantic Bight . The variances are
broken into frequency bands with period ranges as follows :
topographic wave (5 .4 - 29 days), wind driven (30 h - 5 .4 days),
inertial-diurnal (15 - 30 h), semidiurnal (11 - 15 h) and high
frequency (<11 h)(from Csanady et al . 1988) .
111
:::
~ 1 . 00
I
~
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~
d
GO
MOORING 5, 131m ISOBATH
.
.
0 .25
0 . 25
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50
SPEED
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r
r
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BEAM ATTENUATION
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20
20
U
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0
10
15
0
20
Dec 1988
25
30
Figure 4 .6-4 The top plot is a record of beam attenuation (roughly proportional to suspended solids
concentration) measured by a transmissometer located 3 m above bottom at a site on the 131 m
isobath off the Delmarva Peninsula . The bottom plot shows unfiltered (thin curve) and lowpassed filtered (thick curve) versions of a near-bottom speed record measured near the
transmissometer . Note that events of local sediment resuspension, identified by correspondence
of strong current and high beam attenuation, often occur when pulses of high frequency motion
coincide with a maximum of the lower frequency signal (from Churchill et al . 1992) .
CURRENT SPEED 5 mab
so
s0
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U 40
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~ 30
0
w 20
a
N 10
30
0
~
~
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MMS MOORING A
40
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1 20 1 I1 21
~e b 1984 M ar
311
CURRENT SPEED 5 mab
60
11
Apr
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1
11
May
21
311
USGS MOORING SF
0
60
~, s0
\
so
U 40
40
~ 30
~
W 20
30
20
a
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0
10
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20
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Oct 1983
1
11
Nov
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1
ll
Dec
21
311
11
21
Jan 1984
311
0
Figure 4 .6-5 Near-bottom current speeds measured a short distance seaward of the Middle Atlantic Bight shelfedge . The top record is from a current meter deployed 5 m above bottom at a location east of
Virginia with a bottom depth of 225 m . The bottom record is from a current meter set 5 m above
bottom at a location east of New Jersey with a water depth of 202 m .
4 .6 .3 .3 Tidal Currents
Tidal currents, due to both internal and surface waves, strongly influence the
fluid and sediment dynamics of many shallow water regions . To our knowledge, the
only study which has dealt in any significant way with tidal currents in the Cape
Hatteras region is the FRED experiment . As part of this experiment, eight
current-meter moorings were deployed within and offshore of Raleigh and Onslow
Bays . The semidiurnal ( M2) tidal current ellipses derived from the current-meter
records ( Figures 4 .6-6 and 4 .6-7) reveal that vigorous near-bottom tidal
currents, of order 15 cm s-1 in magnitude, occur over the shelf within Onslow
Bay . They also show a significant increase with depth of semidiurnal tidal
current strength at the shelf edge (roughly the 75 m isobath) near Frying Pan
Shoals . The FRED Group ( 1989) took this as evidence that internal tides are
generated and/or intensified in the vicinity of this particular shoals . The
semidiurnal tidal currents measured in Raleigh Bay had relatively small
amplitudes, roughly 5 cm s-1 . It thus appears that the impact of semidiurnal
tidal currents on shelf-sediment motion is potentially significant in Onslow Bay
but likely to be minor in Raleigh Bay . The diurnal tidal signal observed at all
FRED current meters was small, order 3 cm s-1, so that the diurnal tide is not
expected to appreciably influence sediment dynamics over the North Carolina shelf
south of Cape Hatteras .
4 .6 .3 .4 The Gulf Stream and Gulf Stream Frontal Eddies
As discussed in Chapter 2, flow variability at the shelf edge south of Cape
Hatteras is often dominated by Gulf Stream currents and frontal eddies . The
effect of these currents on bottom sediments has not been explicitly studied ;
however, considerable insight as to their influence on sediments in the Cape
Hatteras region can be gained by examining measurements from the FRED current
meter array together with data from moorings placed over the slope near Cape
Hatteras as part of the Middle Atlantic Slope and Rise ( MASAR) study ( SAIC 1987) .
Velocity records from four FRED moorings that were placed at the shelf edge (75
m isobath) show frequent pulses of strong north-westward flow, marking a
shoreward excursion of the Gulf Stream . Additionally, there were numerous
episodes of strong southward flow due to the passages of frontal eddies ( Figure
4 .6-8) . The near-bottom speeds recorded at the shelf-edge moorings during all
such occasions are much lower than the speeds measured directly above . They are,
nonetheless, often in excess of 30 cm s-1 and likely strong enough to mobilize
bottom sediments .
The FRED array included two moorings at the 400 m isobath : one on the Blake
Plateau seaward of Onslow Bay, and the other on the continental slope offshore
of Raleigh Bay . Both had a current meter within 200 m of the surface and another
at 100 m above bottom . Records from the slope mooring (Figure 4 .6-9a,b) contain
numerous pulses of strong northward flow due to the presence of the Gulf Stream .
Pulses of the deeper current meter record peak at roughly 50 cm s-1, while those
of the shallower record are roughly twice as strong . From the Blake Plateau
mooring northward, relatively weak flows due to the Gulf Stream are also seen in
the records, seldom exceeding 30 cm s-1 at 100 m above bottom (Figure 4 .6-9c,d) .
Because these measurements were not taken close to the bottom, they do not lend
themselves to firm conclusions regarding sediment transport . Nevertheless, they
suggest that Gulf Stream currents at the 400 m isobath are unlikely to resuspend
sediments over the Blake Plateau but could be strong enough to mobilize sediments
114
M2 Surface Tidal Current Ellipses
36
° 78°
77°u
76°
75°
- 1 36°
udc
Amplitude : 47.0 cm
Phase : 359 .5'
~•-~~~~_ Hatteras
~ Amplitude : 44 .4 cm ~
Phase : 353 .'
~
1 35°N
Amplitude : 442 cm
Phase : 7.4'
s~
15
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0
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z
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33°
78°
77° u
76°
J ~30
75
Figure 4 .6-6 Semidiurnal M2 tidal current ellipses from the records of the
shallower current meters of the FRED experiment . The bold-faced
number of the ellipse label is the mooring number and the other
number gives the current meter depth . Arrows on ellipses
indicate the Greenwich phase angle and the rotation direction of
the current vector . Velocity scale is given by the arrow at the
lower right ( adapted from the FRED Group 1989) .
115
M2 Bottom Tidal Current Ellipses
78°
36°
77°u
76°
I
35°N
36°
~ I ,
. r :..y.
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75°
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7
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75
Same as Figure 4 .6-6 except showing tidal current ellipses from
deeper current meter records .
116
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o
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ipp ~ A 8 20 m
•1pp
108
_tpp
k) 8 35m
TIME (days)
1987
Figure 4 .6-8 Low-passed (40 h half-power period) filtered records of currents
measured at the 75 m isobath east of North Carolina as part of
the FRED experiment . The records are labeled by mooring number
(see Figure 4 .6-7 for locations) and current meter depth (adapted
from the FRED Group 1989) . "Up" is directed along the general
trend of the isobaths at each station ; 53°T for 8, 45°T for 6,
42°T for 4 and 2 .
117
a) 3 100m
^
y 100
~ -10
`-'
8
F-r
r-r
00
b) 3
300 m
W •00
1
a ~
C) 7 138m
w
d) 7
100
300m
~ -100
U
t 0
140
O
160
2
2
0
240
260
2
0
300
320
TIME (days)
1987
Figure 4 .6-9
Same as Figure 4 .6-8 except showing currents measured at the 400 m isobath . "Up" is
directed along the general trend of the isobaths at each station ; 45°T for 7 and 42'T for 3 .
over the slope east of Raleigh Bay . This must be viewed with suspicion because
it is at odds with the sediment texture observations of the area . As noted in
Section 4 .6 .2, these observations indicate predominately coarse sediment over the
Blake Plateau and mostly fine sediment over the slope to the north .
The MASAR study moorings mentioned above were located over the continental slope
roughly 30 km northeast of Cape Hatteras . They were at water depths of 350 and
710 m . Strong northeastward flows of the Gulf Stream appear in the near-surface
and mid-water velocity records from both moorings (Figure 4 .6-10) . Such flows
are noticeably absent in the moorings' deepest records, which show currents of
no greater than 20 cm s-1 in magnitude . A similar pattern has been seen in other
current meter records taken from the same area (SAIC 1982) . The obvious
implication- -that the Gulf Stream does not appreciably affect sediments over the
slope north of Cape Hatteras--is consistent with the predominance of fine-grains
in these sediments .
4 .6 .3 .5 Bottom Fishing
Over the last few decades the continental shelf off the U .S . east coast has been
subject to intensive bottom fishing with the most commonly used bottom gear being
an otter trawl . Its basic components are a net and two plates, commonly called
the trawl doors, which are towed forward of the net and serve to spread it
horizontally . Underwater observers have noted that turbulence generated in the
wake of trawl doors acts to create plumes of highly turbid water (Main and
Sangster 1981) . Evidence of turbid water behind trawls has also come from the
records of transmissometers moored within Long Island Sound (Bohlen and Winnick
1984) and in the Middle Atlantic Bight (Butman and Noble 1979 ; Churchill et al .
1988) . These observations prompted a study, reported by Churchill (1989),
considering the effect of trawling on sediment transport over the Mud Patch and
Nantucket Shoals . Using records of trawling activity compiled by the U .S .
National Marine Fisheries Service and a simple mathematical model, Churchill
determined that trawl-induced sediment resuspension may account for a large
portion of the time-averaged suspended load over these regions, particularly at
the outer shelf . His calculations also indicated that trawling should initiate
a significant net seaward migration of sediment from these areas . Churchill et
al . (1992) found that sediment movement over the shelf east of the Delmarva
Peninsula is also significantly impacted by trawling . Using probability
analysis, they showed that sediment resuspended by trawls could be responsible
for a number of the turbid water parcels detected during quiescent conditions by
transmissometers moored in this area .
The North Carolina shelf region supports a productive fishery and, like the rest
of the shelf off the U .S . east coast, is heavily trawled . In view of the
findings presented above, it seems certain that bottom fishing should
significantly influence the movement of sediment in this region .
4 .6 .4 Offshelf vs Onshelf Transport
An issue rather hotly debated during the early 1970s was the fate of the solid
material which makes its way into coastal waters . One view was that most of this
material eventually becomes trapped within estuaries and coastal lagoons . An
alternate opinion was that it tends to migrate seaward, "bypass" the shelf and
accumulate in deep water (Meade 1972 ; Drake 1976) . Evidence cited as indicating
onshore transport of sediment in the Cape Hatteras region include the
119
MOORING Q- 350 m WATER DEPTH
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U 100
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300 m
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MOORING R- 700 m WATER DEPTH
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695 m
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20
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Dec 1985
1
11
Jan
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1986
3l1
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21
1
Figure 4 .6-10 Low-passed (32 h half-power period) filtered records of current
speed measured over the continental slope northeast of Cape
Hatteras . "Up" is north .
120
distribution in bottom sediment of garnet (Duane 1962) and phosphorite grains
(Luternauer and Pilkey 1967) . Evidence offered in support of the opposite view
includes turbid plumes seen issuing from the barrier island inlets in satellite
photographs (Mairs 1970) and by suspended sediment surveys (Buss and Rodolfo
1972) . When considered together, these sets of evidence suggest that exchange
of shelf and coastal sediments occurs . Neither set is in any way conclusive with
regard to the long term net transport of sediment from the North Carolina shelf .
A type of data employed by proponents on both sides of the cross-shelf transport
debate is bottom drifter tracks . There have been a number of bottom drifter
studies encompassing the North Carolina shelf (Harrison et al . 1965 ; Bumpus 1973 ;
Schumacher 1974) . Taken together, they indicate that near-bottom flow tends to
be directed offshore at points seaward of the mid-shelf and onshore at locations
shoreward of the mid-shelf . An exception to this occurs in the areas of the
Carolina Capes where offshore, near-bottom flow appears to be the rule across
most of the shelf (Figure 4 .6-11) . While these results are interesting, they
cannot be taken as definite indicators of sediment transport . Because suspended
sediment concentrations undergo significant temporal variations, net water and
sediment transports are unlikely to be the same . Detailed knowledge of
cross-shelf sediment transport in the Cape Hatteras region must therefore await
field investigations in which suspended sediment flux is directly measured .
4 .6 .5 The Influence of Physiographic Features
Though vague on details, the information presently available indicates that
sediment motion over the North Carolina shelf is significantly influenced by the
physiographic character of the region . With regard to sediment transport, the
most important bathymetric features of the North Carolina shelf may be the cape
shoals . As noted above, persistent offshore flow appears to be the rule near the
seafloor in the vicinity of the shoals . In addition, the tidal signal seen in
the FRED current meter data (FRED Group 1989) suggests that strong internal tides
are generated at the shoals (Section 4 .6 .3 .3) .
Sediment transport is also likely to be affected by the ridge and swale
bathymetry of the North Carolina shelf . The tendency for sediment silt-plus-clay
content to be higher in the swales than in the ridges (Section 4 .6 .2) suggests
that the swales may be temporary deposit centers for fine material .
Other bathymetric features of the Cape Hatteras region that may influence
sediment transport are submarine canyons . Major canyons further to the north
have been shown to be robust environments, characterized by strong internal wave
motions and relatively large mean flows directed along the canyon axis (Keller
et al . 1973 ; Hotchkiss and Wunsch 1982 ; Mayer et al . 1982 ; Bothner et al . 1983 ;
Hunkins 1988 ; Gardner 1989a,b) . Evidence also indicates that these canyons serve
as conduits for the seaward transport of fine-grained material (Bothner et al .
1983 ; Gardner 1989b) . Very little is known about the canyons of the Cape
Hatteras region . Field measurements of their flow conditions are limited to
photographs of dye motions (Jenkins 1980) and a few single velocity measurements
taken by a current meter lowered from a drifting ship (Rowe 1971) . Based solely
on their physical characteristics, these canyons seem unlikely to affect shelf
sediment transport to the extent that the major canyons further to the north do .
The latter extend well shoreward of the shelf break, whereas canyons of the Cape
Hatteras region are essentially confined to the continental slope .
121
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~
,
~
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7
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.
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~
.
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T
Figure 4 .6-11 Sea-bed drifter release locations in Onslow Bay (dots) and two
regions of low drifter return at the beaches ( shaded) . Only 2
of the 985 drifters released in the southern low-return region
were recovered on beaches . Corresponding numbers for the
northern region are 7 and 545, respectively ( from Schumacher
1974) .
122
Fluid and sediment dynamics over the inner shelf south of Cape Hatteras are
undoubtedly influenced by the region's cuspate shoreline, peculiar to the North
Carolina coast . An obvious effect of this is the sheltering of near-shore areas
in the lee of the capes . Rodolfo et al . (1971) called upon this to explain a
relatively low suspended sediment concentration found in the lee of Cape Lookout
during hurricane Gerda (see Figure 4 .6-2) .
123
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137
V . NEARSHORE PROCESSES
5 .1
Introduction
In the southern Middle Atlantic Bight, the nearshore circulation comprises five
primary modes of motion . The two wind-driven modes are the coastal jet, parallel
to shore (Scott and Csanady 1976), and upwelling and downwelling, which are shore
normal (Wells and Gray 1960 ; Boicourt 1973 ; Singer and Knowles, 1975) . The
large-scale, annual-mean flow to the south in this region is now thought to be
buoyancy driven (Chapman and Beardsley 1989), and part of a 5000 km long coastal
current originating off the coast of Greenland . This component of the flow is
of the order 3-5 cm s-1 over the inner shelf of the Middle Atlantic Bight (Bumpus
1973 ; Boicourt 1993) . An additional, highly episodic buoyancy-driven current
results from the Chesapeake Bay outflow (Boicourt 1973, 1982 ; Boicourt et al .
1987) . The Bay outflow initially forms a plume with a large anticyclonic turning
region off the Bay entrance . If the plume is not interrupted by strong
upwelling-favorable (southerly) winds, a high-velocity coastally trapped jet is
formed in the far field . Low-salinity water emanating from Pamlico and
Albemarle Sounds via Oregon Inlet can create buoyancy driven plumes that are
often detected in the visible and IR bands of satellite sensors, but the volume
fluxes are inadequate to develop a significant coastal jet .
The fifth primary mode of nearshore motion is the-tides, which drive currents
typically of order 20 cm s-1 or less, except in the vicinity of estuaries or
inlets (see Section 4 .2 .1) . In the vicinity of Chesapeake Bay, the tidal
currents can reach velocities of 120 cm s-l . Ebb tidal currents (which are a
combination of the astronomical tide and the mean gravitational-circulation
outflow) can be substantially greater in magnitude, especially when a quarterwave seiche of Chesapeake Bay produces an outflow surge (Chuang and Boicourt
1989) .
5 .2
Shelf-Estuary Exchange
Bigelow (1915, 1933) and Bigelow and Sears (1935) recognized the influence of
rivers and estuaries on the nearshore currents and salinity distributions in the
Middle Atlantic Bight . Ketchum, et al . (1951) noted the tendency of the Hudson
River outflow to turn and form a plume parallel to the New Jersey coast during
the spring runoff . As river flow diminished, the plume of the Hudson was less
obvious in the surface salinity . During winter when shelf waters were not
strongly vertically stratified, the Hudson plume was confined to a narrow band
along the New Jersey coast . The Chesapeake Bay outflow plume, which was recently
examined by a large multidisciplinary program called MECCAS (Microbial Exchange
and Coupling in Coastal Atlantic Systems ; Boicourt et al . 1987), has a similar
tendency to expand upon discharge onto the continental shelf, and then turn to
the south to form a narrow coastal jet (Boicourt 1973, 1980, 1982) . The
withdrawal of shelf water by estuarine inflows can influence bottom currents in
the Mid Atlantic Bight as far as 30 km seaward of the estuary entrance (Bumpus
1965, 1973 ; Bowman and Wunderlich 1976 ; Beardsley and Hart 1978 ; Boicourt 1980,
1982 ; Chao and Boicourt 1986 ; Masse 1990) .
If the outflow from the Chesapeake Bay is strong and the resulting coastal jet
is undisturbed by opposing (southerly or easterly) winds, low-salinity water
originating from the Bay can be tracked 170 km from the Bay entrance to Diamond
Shoals off Cape Hatteras (Boicourt 1973, 1982 ; Boicourt et al . 1987) . This jet
139
is a narrow, (10 km) relatively steady flow behind the nose of a bore intrusion
propagating down the North Carolina coast (Chao and Boicourt 1986) . Typical mean
velocities in the jet are 50-100 cro s"1 . When this buoyancy-driven jet is
combined with a (northerly) wind-driven coastal jet, the velocities can reach 200
cm s"1 . During large outflow surges from Chesapeake Bay (Chuang and Boicourt
1990), the coastal jet can be detected southward of Oregon Inlet on the North
Carolina coast . Presumably, this water becomes entrained into the Gulf Stream,
or on occasion, passes around Diamond Shoals to Raleigh Bay (Bumpus and Pierce
1955) . Upwelling-favorable winds between Cape Hatteras and the Chesapeake Bay
entrance can force the low-salinity plume water offshore or block the development
of a jet altogether .
Although freshwater discharge along the coast of the South Atlantic Bight may not
produce a consistent southward coastal current in the presence of adverse
northward directed winds (Bumpus 1973), the discharge is sufficient to produce
a narrow (10 km) band of low-salinity water that is often bounded on the offshore
by a front (Blanton 1981) . This discharge occurs, not only via the
distributaries of the larger rivers, but also through the many inlets and sounds
of the "perforated barrier" coastline of the South Atlantic Bight between Cape
Fear, North Carolina, and Jacksonville, Florida (Blanton and Atkinson 1978) .
When downwelling favorable winds (from the north) augment the buoyancy driven
flow, a narrow, southward coastal current develops that appears continuous to
Cape Canaveral (Blanton and Atkinson 1983) . Upwelling favorable winds force the
surface Ekman layer offshore and oppose the development of a southward coastal
current (Blanton 1981) . Under these conditions, parcels of low-salinity water
are found over the middle and outer shelf (Blanton and Atkinson 1983) .
5 ..3 Sediment Transport
5 .3 .1 Basic Concepts
In the study of marine sediments, the nearshore zone is often taken to coincide
with the shoreface, a steeply sloped transitional region connecting the beach
with the gently inclined inner shelf . The shoreface off the U . S . east coast has
maximum slopes in the range of 0 .05-0 .005 . It typically merges with the inner
shelf, which has inclinations as low as 5x10-1, at depths of
10-15 m and at distances of 5-10 km from shore .
From a dynamic viewpoint, the shoreface can be divided into two zones . The more
shoreward of these is the surf zone, where circulation and energy dissipation are
dominated by breaking surface waves . The wave energy to this zone is supplied
primarily by incident gravity waves . However, as these waves propagate into
shallow water, their energy is transferred to lower frequency infra-gravity waves
(e .g . Holman 1981 ; Guza and Thornton 1982, 1985a) . Within the inner surf zone,
the energy level of infragravity waves can exceed that of gravity waves by a
factor of four or more (Guza and Thornton 1982 ; Wright et al . 1982) . The effects
of infragravity waves on sediment movement have only recently come under
scrutiny . Results of field studies presented by Guza and Thornton (1985b) and
Huntley and Hanes (1987) indicate that while incident gravity waves tend to move
sediment shoreward, the interaction of gravity and infragravity waves results in
seaward sediment movement within the surf zone .
Other types of flow likely to affect cross-shore sediment transport in the surf
zone include rip currents and near-bed seaward currents compensating for wave140
induced near-surface flow (Mei and Liu 1977 ; Holman and Bowen 1982 ; Roelvick and
Stive 1989) .
Sediment and fluid dynamics in the shoreface region seaward of the surf zone are
affected both by surface wave currents and more slowly varying flows (as those
due to tides and wind forcing) . Bottom stress production in this region arises
primarily from the nonlinear interaction between these two types of currents
(Smith 1977 ; Grant and Madsen 1979) . Wave related processes affecting the
cross-shore advection of sediment in this region include orbital asymmetries
(Wells 1967) and the interaction between gravity and infragravity waves (Shi and
Larsen 1984 ; Dean and Perlin 1986) . Shi and Larsen (1984) demonstrated that the
latter process can result in seaward sediment transport for the case where
infragravity waves are bound to gravity wave groups (see Figure 5 .3-1) .
Cross-shore sediment movement across the shoreface outside the surf zone is also
affected by wind-induced upwelling and downwelling circulation and by tidal
residual flows (Geyer and Signell 1991) .
5 .3 .2 Research Within the Cape Hatteras Region
The object of much of the research conducted in shoreface areas is the transfer
of sediment to and from the intertidal zone . It is well established that this
zone tends to gain sediment during fairweather conditions, resulting in beach
accretion, and to lose sediment during storms . Over the past decade this
phenomenon has been the focus of a number of field investigations within the
Carolina Capes region, most sited at the U . S . Army Corp of Engineers Field
Research Facility in Duck, North Carolina (Figure 5 .3-2) . Two of the studies
carried out at the Duck site, during 1982 and 1985, have been extensively
documented in the literature (see Mason et al . 1984 and 1987 for overviews of
these studies) .
The 1982 Duck study was the first to document detailed changes of nearshore
morphology during a storm and subsequent fairweather period . Bathymetric profile
measurements of this study, presented by Mason et al . (1984), show the formation
of two sand bars during a storm that passed the Duck site in mid-October 1982 .
Both bars grew in amplitude and migrated offshore during the course of the storm .
Over a four-day period after the storm, the more onshore bar progressed shoreward
and shrank to an indistinguishable amplitude while the offshore bar remained
relatively stable . Mason et al . (1984) noted that both bars took shape well
within the surf zone and thus could not have resulted from breaker-induced
scouring of a trough . Sallenger et al . (1985) put forth the hypothesis that the
formation and movement of the inner bar were caused by a standing infragravity
wave . As noted by Sallenger and Holman (1984) infragravity waves produced r .m .s .
cross-shore currents in excess of 0 .5 m s-1 over the bar crest during the height
of the storm . Other analyses revealed a complex pattern of sediment transport
during and after the storm . Jaffe et al . (1984) found that coupling between
suspended material concentration and onshore wave motion produced a net onshore
sediment flux even though the net fluid transport was offshore . The movement of
different sediment sizes was examined by Richman and Sallenger (1984) who
concluded that fine and coarse material likely moved in opposite directions
during the storm .
The 1985 Duck Study featured a larger suite of instruments and more rapid
sampling than the 1982 experiment and was expanded to include the middle
shoreface . Changes in nearshore morphology observed in this study were similar
141
SHOREWARD
SEAWARD
HIGH WAVES LOW WAVES
sWl
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~
N
FORCEO
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-r
LONG
PERIOD
~
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LOW
A
.
..
.
.
Figure 5 .3-1 Conceptual diagram illustrating the notion of sediment resuspension and seaward transport by
the combination of gravity and bound infragravity waves (taken from Wright et al . 1991 and
based on the theory of Shi and Larsen 1984) .
~
~`
.
<
~
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.
y
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ss:•
46
'J
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_
Sandbndge , ~ ,
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Site'
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.
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Figure 5 .3-2 Sites of shoreface sediment dynamics studies within the Carolina
Capes region .
143
to those seen in 1982 . Howd and Birkemeier (1987) reported that the passage of
a storm in mid-September 1985 coincided with the formation and seaward migration
of a nearshore sand bar . This took on a crescentic shape as the storm abated .
The long-shore transport of surf zone sediment was estimated by Kraus and Dean
(1987) using current meter records and material collected in an array of sediment
traps . They concluded that sediment movement during the storm was primarily
the result of suspended load (and not bedload) transport . Green et al . (1989)
found that suspended load transport also dominated sediment movement over the
middle shoreface during the storm . This transport apparently produced dramatic
changes in shoreface morphology . As reported by Wright et al . (1986), bed level
at a site at the middle shoreface experienced a 5 cm drop during the first phase
of the storm followed by a 15 cm rise as the storm abated .
The most in-depth examination of sediment movement over the middle and lower
shoreface off the Carolina Capes was carried out by Wright et al . (1991) .
Employed in their study were data collected at Duck during 1985-1987 and from a
site 65 km to the north of Duck during 1988 (the Sandbridge site in Figure
5 .3-2) . Using these measurements Wright et al . (1991) estimated sediment fluxes
and attempted to deduce the primary causes of sediment transport during various
sets of conditions . They concluded that wind-driven downwelling circulation
should dominate the offshore flux of sediment seaward of the surf zone during
northeasterly storms . According to their calculations, "a fairly commonplace
northeasterly storm is capable of transporting more sand offshore in an hour than
fairweather processes can move onshore in two or more days" . They found that
cross-shore sediment fluxes during fairweather conditions resulted primarily from
tidal flows . Their analysis also demonstrated that infragravity wave motions
should have a measurable, but not dominant, effect on cross-shore sediment
movement at the middle shoreface . However, they found no compelling evidence in
support of the transport mechanism proposed by Shi and Larsen (1984) . Another
model for which their data lent little support is that in which incident wave
asymmetries are considered to be the principal cause of onshore sediment flux
over the middle shoreface (e .g . Wells 1967 ; Swift et al . 1985) . This was found
to be true only during one experiment conducted during swell-dominated
conditions .
144
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148
VI . SUMMARY
Reviews of the major oceanographic systems and processes that affect circulation
in the study area have been given in the preceding chapters . The study region
encompasses coastal, shelf and offshore waters from Cape Lookout, North Carolina
to the mouth of the Chesapeake Bay on the eastern seaboard of the United States .
This region around Cape Hatteras is one in which the convergence and interaction
of widely different water masses and processes, operating over a range of space
and time scales, makes for one of the most complex oceanographic systems in
coastal U .S . waters .
The organization of the reviews is such that processes are grouped from offshore
to onshore, starting with the Gulf Stream over the deep waters of the continental
slope and rise, and concluding with nearshore waters (excluding the surf zone)
with shallow depths of order 10 m or less . Roughly speaking, the degree of
knowledge and the amount of field data concerning these systems also proceeds in
the same direction, in that the basic structure of the Gulf Stream and the modes
of its variability have been extensively studied in this region and all along the
eastern seaboard, but shelf and nearshore circulation processes have had little
systematic study on the Cape Hatteras shelf .
The Gulf Stream review concentrated on path variability, meanders and frontal
eddies, and interactions with warm and cold-core rings . Since the Gulf Stream
transitions from a slope current, in the South Atlantic Bight, to a free flowing
jet over water depths of 3000 m and greater, in the vicinity of Cape Hatteras,
the behavior of the Gulf Stream path in both large subregions is relevant to
local variability . The complex motions of the Gulf Stream have been monitored
extensively in the last 1-1/2 decades . Much has been revealed from both South
Atlantic Bight and Middle Atlantic Bight studies on meanders, the growth, decay
and propagation characteristics of frontal eddies, seasonal and interannual
shifts in position of the north wall, and changes in transport . However, it
would be fair to say that we do not have a good idea why it behaves as it does .
For example, the factors controlling the position where the Stream separates from
the slope, which can vary from Cape Hatteras to as far north as the eastern
shores of Virginia, are not well understood . The Gulf Stream provides major
impacts on the outer shelf and upper slope including filaments and frontal eddies
over Raleigh Bay, overwashes of Gulf Stream derived water north of Cape Hatteras,
and the entrainment of Middle Atlantic Bight shelf (Ford water) and upper Slope
Sea water along the north wall as it turns eastward away from the slope . These
interactions have been much less studied and thus are less understood than the
structure and variability of the Gulf Stream proper .
Between the shelf-slope front and the north wall of the Gulf Stream in the Middle
Atlantic Bight, there is a wedge-shaped region with distinct water masses and
circulation processes known as the Slope Sea . In the vicinity of Cape Hatteras,
the presence of Slope Sea water in the narrow region between the shelf and the
Gulf Stream is only intermittent depending upon the configuration of the Gulf
Stream . The southward drift of slope waters along the general trend of the
isobaths is present throughout the water column in the Middle Atlantic Bight .
The drift in the upper layers is short-circuited by the Gulf Stream, or by
extrusions of Gulf Stream derived water, with a return flow of slope water along
the north wall of the stream, thereby forming the southern limb of the cyclonic
Slope Sea gyre . The various water masses below 800-1000 m form the different
components of the Deep Western Boundary Current system . The deeper components
149
of the DWBC descend under the Gulf Stream, move offshore in the vicinity of the
Hatteras Corner and continue southward along the Blake Escarpment . This is
another example of interactions between current systems, which are only just
beginning to be explored . The variability in the upper layers is complex and is
influenced by the Gulf Stream, warm core rings, smaller eddies, and shelf-slope
exchanges . The latter may be forced by both wind and eddy induced circulations .
In the lower layers of the DWBC, the variability is primarily from topographic
Rossby waves, probably generated by large Gulf Stream meanders occurring well to
the east . TRWs have characteristic periods of weeks to months with energy
propagating westward and southwestward along the isobaths and with a small
component towards the slope .
These Slope Sea topics were reviewed indicating that considerable current data
are available for the Slope Sea and DWBC systems . In the vicinity of Cape
Hatteras, most of the upper layer measurements over the continental slope are
from the MASAR, Chevron and Mobil studies . However, apart from long term
measurements of the Gulf Stream by current meter and Inverted Echo Sounder arrays
and statistical studies of AVHRR imagery, most of the time series in-situ
measurements are for a year or less . An exception is the two-year long current
measurements of the MASAR study . Thus, our knowledge of long term variability
of circulation processes, particularly related to influences of the interannual
changes in the Gulf Stream path, is quite limited . Similarly, exchanges between
the Gulf Stream and Slope Sea with the Cape Hatteras shelf have not been directly
measured, though much can be derived from the slope and shelf studies in both the
Middle Atlantic Bight (SEEP) and the South Atlantic Bight (FRED and GABEX ; most
notably) .
One of the major characteristics defining Cape Hatteras oceanography is the rapid
changes in water mass characteristics and circulation processes in both the
along-shore and cross-shore directions . Changes in topography, Gulf Stream
characteristics, and the presence of the Slope Sea north, but not south, of the
Hatteras Corner are immediately apparent . Besides these offshore influences on
the shelf, Middle Atlantic Bight and South Atlantic Bight shelf waters differ
considerably, and these two water masses meet and interact at Diamond Shoals .
It is necessary to review shelf circulation for both the Middle Atlantic Bight
and the Carolina Capes to gain an appreciation of their contributions to the
Hatteras shelf . Shelf current measurements are very limited between Diamond
Shoals and the Chesapeake mouth, though there is a reasonable data base of water
mass properties . In the Carolina Capes, the available current measurements
(primarily from FRED and GALE) are of short duration (six months or less) .
On continental shelves, a set of shorter time scale processes become much more
important including synoptic (two days to two weeks) meteorological forcing,
tides and surface waves . The seasonal changes in the basic atmospheric
circulations and synoptic systems are reviewed because of their importance to
wind-forcing and seasonal changes in both stratification and current
characteristics over the shelf . The Carolina Capes region is a primary region
for winter storm cyclogenesis . These ocean-atmospheric interactions were studied
in GALE . To date, synoptic wind events have not been shown to be an important
influence on the Gulf Stream or Slope Sea surface circulations . On the shelf,
however, wind forcing, both direct and indirect (via the mechanism of continental
shelf waves), generate the major source of subtidal variability and are of
primary importance in generating longshore and cross-shore transports . Tides
make large contributions to current variability with the semidiurnal component
150
dominating, except near Cape Hatteras . There is some evidence of internal tide
generation at the shelf break at Frying Plan Shoals and associated with the
shelf-slope front in the Middle Atlantic Bight . Both tides and surface waves are
important for bottom boundary layer dynamics over the shallow shelf and, along
with meteorological forcing, determine the intensity of vertical mixing .
Sediment characteristics are reviewed and an attempt made to determine the
intensity of sediment resuspension and transport from all the above processes
with the addition of trawling activity .
Buoyancy inputs to the shelf also differ north and south of Diamond Shoals . The
Middle Atlantic Bight north of Hatteras is now thought to be the terminus of a
5,000 km coastal current that originates in the Greenland and Labrador Seas and
has contributions from major estuaries along the eastern seaboard from the St .
Lawrence to the Chesapeake . On the Middle Atlantic Bight shelf, this coastal
current has a southward drift of 3 to 5 cm s-1 and includes such notable features
as the ribbon of cold bottom water known as the "cold pool" . The salinity of the
shelf increases from north to south reflecting exchange with Slope Sea waters
across the shelf-slope front . However, the salinity and temperatures of the
shelf just north of Diamond Shoals are still less than are characteristic of
Raleigh and Onslow Bays, which have only small brackish water inputs from the
coast and are frequently partially flushed by warm Gulf Stream filaments . The
mean drift of the South Atlantic Bight is a few cm s-1 to the north, though it
is a much less persistent feature of the circulation than in the Middle Atlantic
Bight, being more influenced apparently by seasonal wind patterns . Since there
is virtually no evidence of South Atlantic Bight water on the Hatteras shelf, it
is speculated that some South Atlantic Bight water is entrained into the Gulf
Stream front through frontal eddy exchange in Raleigh Bay .
Some portion of the southward coastal drift in the Middle Atlantic Bight is also
expected to cross the shelf-slope front north of Hatteras and be entrained along
the north wall, where it is often visible in AVHRR imagery many hundreds of
kilometers downstream . This is known as Ford water, and characteristics and
positions of this major shelf water export are not presently known with any
certainty, though it is likely to be strongly influenced by Gulf Stream behavior .
The shelf-slope front is present along the shelf break from Georges Bank to Cape
Hatteras . It surfaces in winter but is capped by a seasonal thermocline in
summer . Exchange processes across this thermohaline front are complex and
include intrusions of slope water at thermocline depth, frontal instabilities,
wind-forced cross front flows and entrainment by slope eddies . Little is known
about the characteristics of the shelf-slope front and its interactions with the
Gulf Stream south of the mouth of the Chesapeake .
On occasion, part of the cooler, fresher Middle Atlantic Bight shelf water rounds
Cape Hatteras and flows into Raleigh and sometimes Onslow Bays . These events are
known as Virginia Coastal Water intrusions, and their frequency of occurrence,
causes, and importance to the ecology of Raleigh and Onslow Bays are not well
known at the present . Most of our quantitative data comes from imagery and the
FRED and Mobil experiments,
which were restricted to spring and summer
conditions . These studies showed that both strong northerlies and the occurrence
of Gulf Stream frontal eddies could be possible causes of Virginia Coastal Water
intrusions .
151
In the northern part of the study area, the outflow from the Chesapeake Bay (the
Chesapeake plume) is an important local buoyancy input to the shelf . Depending
on the strength of the outflow and the prevailing wind conditions, the plume can
form a nearshore jet along the coast, which may reach as far south as Oregon
Inlet during the peak of spring runoff, or it can diffusively spread across the
width of the shelf, with a consequent influence on the stratification, when winds
are from the south or southwest .
Apart from the Chesapeake plume studies, nearshore circulation studies along the
North Carolina coast have generally been surfzone studies (at Duck, for example)
or inlet studies (such as Oregon Inlet) . The importance of the limited windforced and tidal exchanges with Pamlico Sound, through the inlets of the Outer
Banks, to the nearshore circulation has not yet been established . Indeed most
of the basic information on the coastal boundary layer is unknown for this
region .
In summary, the wide variety of processes, ranging from the Gulf Stream to
nearshore dynamics, coupled with rapidly changing oceanographic characteristics
between Cape Lookout and Cape Henry, make the Hatteras shelf and slope complex
and subject to time and space scales that range from internal waves and frontal
instabilities, through wind and eddy events, to intra- and inter-annual changes
in the Gulf Stream path and its characteristics . The lack of long-term time
series measurements on the shelf indicates that many of the basic shelf
circulation processes, including cross-shelf exchange, have not been quantified .
152