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Continental Shelf Research 29 (2009) 2318–2332
Contents lists available at ScienceDirect
Continental Shelf Research
journal homepage: www.elsevier.com/locate/csr
Subtidal flow and its variability at the entrance to a subtropical lagoon
Patrick L. Murphy a,, Amy F. Waterhouse b, Tyler J. Hesser b, Allison M. Penko b, Arnoldo Valle-Levinson b
a
b
National Marine Fisheries Service, Southeast Science Center, 3500 Delwood Beach Rd., Panama City, FL 32408, USA
University of Florida, Department of Civil and Coastal Engineering, 365 Weil Hall, P.O. Box 116580, Gainesville, FL 32611, USA
a r t i c l e in f o
a b s t r a c t
Article history:
Received 19 May 2009
Received in revised form
12 September 2009
Accepted 21 September 2009
Available online 12 October 2009
Spatial and temporal variability of the subtidal exchange flow at West Pass, an inlet at the entrance
to a subtropical lagoon (St. Andrew Bay, Florida), was studied using moored and towed current
velocity profiles and hydrographic data. Towed and hydrographic measurements were captured
over one diurnal tidal cycle to determine intratidal and spatial changes in flow. Hydrographic profiles
over the tidal cycle showed that tidal straining modified density stratification asymmetrically, thus
setting up the observed mean flow within the inlet. During the towed survey, the inlet’s mean flow had
a two-layer exchange structure that was moderately frictional and weakly influenced by Coriolis
accelerations. Moored current profiles revealed the additional contribution to the dynamics from
centrifugal accelerations. Along channel residual flows changed between unidirectional and exchange
flow, depending on the forcing from the along-estuary wind stress and, to a lesser extent, the spring–
neap tidal cycle. Increases in vertical shear in the along channel subtidal flow coincided with neap tides
and rain pulses. Lateral subtidal flows showed the influence on the dynamics of centrifugal
accelerations through a well-developed two-layer structure modulated in magnitude by the spring–
neap tidal cycle.
Published by Elsevier Ltd.
Keywords:
St. Andrew Bay
Florida
Estuarine dynamics
Tidal dynamics
Residual circulation
1. Introduction
The forcing mechanisms of subtropical coastal lagoons may
differ from those of typical temperate estuaries. Whereas density
gradients normally play a major role in forcing temperate
estuaries, they tend to be of secondary importance in subtropical
lagoons. General characteristics inherent to subtropical lagoons
include microtides, shallow depths, small watersheds, and narrow
inlets. Consequently, the amount of freshwater input (Lee et al.,
1990; Schroeder et al., 1990; Miller and McPherson, 1991; Lipp
et al., 2001; Sylaios and Theocharis, 2002; Poulakis et al., 2004),
meteorological forcing (Smith, 1977, 1979; Swenson and Chuang,
1983; Liu, 1992; Wang, 1998), and the strength of tidal currents
(Ianniello, 1979; Smith, 1993; Valle-Levinson et al., 2000; Weisberg and Zheng, 2003; Sepulveda et al., 2004; Seim et al., 2006)
are crucial in determining whether subtropical estuaries are
driven by density gradients, tides, or winds.
Narrow inlets at the lagoon/ocean boundary are typically
formed as a result of active littoral drift, wave energy, and
microtides (Kjerfve, 1986). Previous studies have examined the
circulation on either side of subtropical lagoons (Blumberg and
Kim, 2000; Moller et al., 2007; Webb et al., 2007; Murphy and
Valle-Levinson, 2008) while the flow strictly within the inlet has
Corresponding author.
E-mail address: [email protected] (P.L. Murphy).
0278-4343/$ - see front matter Published by Elsevier Ltd.
doi:10.1016/j.csr.2009.09.011
been less studied. The flow structure of an inlet determines the
estuary/ocean exchange and is responsible for dispersion of
nutrients and pollutants, for larval transport, and sediment
dynamics (Lyczkowski-Shultz et al., 1990; Hill, 1995; Brown
et al., 2000; Guyondet and Koutitonsky, 2008).
The purpose of this study is to determine the spatial and
temporal variability of flow at the entrance to a subtropical
lagoon, St. Andrew Bay, Florida. This unique subtropical inlet has
marked freshwater influence and more resembles a temperate
density-driven estuary than a tidally driven system, typical of
subtropical areas. The influence of different forcing mechanisms
in causing spatial and temporal variability and in driving the
inlet’s hydrodynamics is examined over different temporal scales.
Survey and moored observations combined with theoretical
results are used to address these objectives.
2. Study area
2.1. St. Andrew Bay system
Northwest Florida’s St. Andrew Bay system is a drowned river
valley, coastal plain estuary consisting of four subestuaries: North
Bay, West Bay, East Bay, and St. Andrew Bay (Fig. 1). The latter is
the only bay with a direct connection to the Gulf of Mexico, while
West Bay and East Bay connect to the Gulf Intracoastal Waterway.
This estuary is ecologically important but surrounded by a region
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Fig. 1. Study area: D0 and B0 are locations of CTD profiles. The thick solid black line connecting points A–D in the lower figure indicates the survey route of the towed ADCP.
The jagged symbol at the survey mid-point denotes the position of the moored ADCP. Isobaths are represented by thin solid lines, with depths indicated. Line A–B is
Transect 1 and Line C–D is Transect 2.
with growing anthropogenic activity. St. Andrew Bay contains the
most diverse marine population of any estuary on the northern
Gulf of Mexico coast (Ogren and Brusher, 1977) due to low
turbidity despite freshwater inflow, extensive sand flats,
widespread submerged aquatic vegetation and a deep basin
containing both coarse and fine sediments (Brusher and Ogren,
1976).
Tides in the entire estuary are diurnal with a mean tidal range
of 0.5 m and a longer ebb tide than flood tide (Ichiye and Jones,
1961; McNulty et al., 1971). The system has a surface area
of 243 km2, draining a watershed of approximately 2800 km2
(USEPA, 1999), which is located entirely in the state of Florida
(NWFWMD, 2001). The largest freshwater source for the system is
spillover from the Deer Point Reservoir located at the head of
North Bay (Fig. 1).
2.2. West pass and East pass
West Pass is a 150 m wide channel, originally dredged in 1934,
that provides a direct link from St. Andrew Bay to the Gulf of
Mexico. This inlet is armored with jetties at the entrance,
maintaining a minimum depth of 9 m from West Pass to the Port
of Panama City (St. Andrew Bay) and to East Bay. West Pass is the
inlet studied in the present work.
East Pass, the original entrance to St. Andrew Bay, was a
natural, shallow inlet southeast of St. Andrew Bay. East Pass closed
naturally in 1998, was re-opened in 2001, and again closed
naturally by 2004. Theoretically East Pass will not naturally
stabilize while the present depth of West Pass is maintained by
dredging (Jain et al., 2004).
3. Data collection and methods
3.1. Meteorological forcing and water levels
Meteorological data for the St. Andrew Bay system were
obtained from Panama City-Bay County International Airport
(PFN), located 7 km north of the study area (Fig. 1). Water height
measurements from NOAA tide station ]8729108 (Panama City,
FL), located 6.5 km from West Pass, were used to identify subtidal
variability and forcing. Mean daily water flow from the Deer Point
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Reservoir was calculated from daily water level measurements at
the Williams Bayou Pumping Station adjacent to the Deer Point
Reservoir Dam Water level heights at the reservoir were
correlated with values of water flow over the dam using a table
lookup method calculated by the Bay County Water Division
(Personal communication, 2008).
3.2. Profiles of physical parameters
Conductivity–temperature–depth (CTD) profiles were collected
using a Sea Bird SBE19 CTD at two locations along the survey
route. Sampling site B0 was located inside West Pass while
sampling site D0 was at the outer end of West Pass, adjacent to
the Gulf of Mexico (Fig. 1). A total of fifteen profiles were obtained
over the observation period at each site to determine the temporal
variability of water column stratification and the effects of tidal
straining on such variations in stratification.
Tidal straining is the result of velocity shear acting on a
horizontal density gradient creating oscillations in the stratification of the water column. In general, tidal straining creates a
stratified water column during ebb tide and a mixed water
column during flood tide (Simpson et al., 1990; de Boer et al.,
2008). To determine the influence of tidal straining on water
column stratification, the potential energy anomaly, j, of the
water column for each CTD cast was calculated. Following the
approach of Simpson et al. (1978), the potential energy anomaly is
the amount of work necessary to completely mix the water
column (J/m3) which can be calculated from
j¼
1
h
Z
ðr^ rÞg z dz;
ð1Þ
Fig. 2. Grid superimposed on February 14–15, 2008 survey route. J denote grid squares used for data analysis. Columns a–b and i–j were removed along with rows 1–2 and
27–30 to ensure accuracy of interpolation.
Fig. 3. Tidal height (m), wind speed (m s 1), sea level pressure (mbar), and rain fall (cm) during the study period.
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where r^ is the vertically averaged water density,
r^ ¼
1
h
Z
0
rðzÞ dz
ð2Þ
h
and h is the total depth, r is the density of water, z is the vertical
depth coordinate and g is the acceleration due to gravity
(9.8 m s 2).
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Temporal variations of water column stratification caused by
tidal straining can be quantified by
Z
djts
@r^ 0
¼g
ðup u^ p Þz dz;
ð3Þ
dt
@x H
where up is the along-estuary vertical velocity and u^ p is the
vertically averaged velocity, both at a point near the CTD sampling
sites (Simpson et al., 1990). The horizontal gradient @r^ =@x is
Fig. 4. Coherence squared (top) and phase (middle) between along-channel velocity and sea surface over the entire observation period of the moored ADCP with the 95%
confidence interval (dashed-dotted line); Observed along channel (dashed) and sea surface elevation (solid) during the towed ADCP survey from February 14–15, 2008
(bottom).
Fig. 5. Surface (1.86 m) K1 tidal ellipses overlaid on contours of K1 tidal phase for February 14–15, 2008.
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determined by calculating the average density between the two
sampling sites (B0 and D0 ) for each time. Comparison of @j=@t as
derived from j in Eq. (1) and @jts =@t from Eq. (3), determines the
relative contribution of tidal straining to estuarine stratification.
An exact match between the two quantities would indicate that
tidal straining is the only mechanism causing stratification
variations.
3.3. Towed ADCP data
In order to describe the spatial structure of the subtidal circulation
in West Pass, a survey was conducted aboard the R/V Harold B
between February 14 and 15, 2008. The vessel towed a downward
pointing 1200 kHz broad band ADCP (RD Instruments, Poway, CA)
mounted on a 1.2 m sled along a closed ‘‘hour glass’’ shaped trajectory
inside West Pass (Fig. 1). The transect lines at the gulf and bay
extremes of the channel were perpendicular to the thalweg.
The ADCP was towed at 2 m/s for 23.98 h, collecting 1 s pings
averaged over 10 ensembles, yielding a horizontal resolution of
20 m and a vertical resolution of 0.5 m. The surface bin was at
1.5 m depth. Data in the lower 15% of the water column were
discarded due to interference from side lobe effects (RD Instruments, 1996). Data were compass-calibrated and corrected by the
method of Joyce (1989) using a Global Positioning System. A total
of thirty transect repetitions were executed.
3.4. Gridded raw measurements
Raw velocities measured by the ADCP (uobs, vobs) were plotted
onto a grid of 10 30 points covering the entire area of West Pass
resulting in 48 m 54 m grid ‘‘squares’’ (Fig. 2). Squares
containing less than 75% of the measured data and containing
outliers (columns a–b and rows 1–2, 27–30) were eliminated,
leaving 90 grid squares for further analysis. Residual components
of the velocity (u0, v0) were obtained from fitting the observed
currents in each grid square to the dominant diurnal (K1 = 23.92 h)
and semidiurnal (M2 = 12.42 h) harmonics using a least squares
method (Lwiza et al., 1991). Given the record length and the
Rayleigh criterion, complete resolution of all diurnal and
semidiurnal constituents was not possible (Emery and Thomson,
2004).
Predicted currents (upred, vpred) calculated from the least
squares fit were compared to the observed currents using a
goodness of fit parameter, R2,
P
ð/uobs S upred Þ2
R2 ¼ P
;
ð/uobs S uobs Þ2
ð4Þ
where angled brackets represent temporal averages. Further
analysis eliminated 18 grid squares where the least squares fit
did not represent the diurnal nature of the tide despite a high R2
value, leaving 72 grid squares for final analysis. Individual
Fig. 6. Observed sigma–t values and tidally averaged sigma–t at each transect through the towed ADCP sampling period at Stations B0 and D0 , shown on the left hand panels.
Corresponding tidal velocity and tidal average near CTD sampling points are shown on the right hand panels in cm s 1. Note the lowest sigma–t for both transects occur
when velocity is approaching zero signifying tidal straining influence.
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Fig. 7. Top: potential energy anomaly ðfÞ calculated from the observed density profiles at stations B0 and D0 . Middle and bottom: comparisons between potential energy
anomaly changes and the forcing from tidal straining at Transect 1 and Transect 2.
Fig. 8. Top panel: contours represent the observed residual mean flow (cm s 1) in Transect 1 rotated to the angle of maximum variance. Arrows denote observed transverse
residual flow. Bottom panel: contours represent the analytical solution of the mean flow (cm s 1) with the bathymetry of Transect 1 (Ke = 0.8 and Ek = 0.01). Arrows denote
transverse residual flow calculated from the analytical solution. Negative contours (shaded) are flow out of the inlet, positive is into the inlet, from the Gulf of Mexico.
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transects were also studied for information on the depth and
spatial pattern of residual flow across the inlet.
3.5. Residual flow and analytical model
Time series of observations along transect A-B (Transect 1) and
transect C-D (Transect 2) were also examined (Fig. 1). Observed
currents were rotated to the angle of maximum variance (Emery
and Thomson, 2004) and fitted to K1 and M2 harmonics also using
the least squares fit method. The spatial contours of residual
velocities across each transect were then compared to an
analytical model (Valle-Levinson, 2008) that considers the nontidal momentum balance among pressure gradient, Earth’s
rotation and bottom friction. Results from both observed and
Fig. 9. Same as Fig. 9 but for Transect 2 (Ke =1.82 and Ek = 0.03).
Fig. 10. Biharmonic spline interpolation of surface (1.36 m) subtidal circulation resulting from removal of diurnal and semidiurnal tidal signals during February 14–15, 2008
survey. Arrows with dots at base denote observation used for interpolation.
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modeled residual flows were cast in terms of the Kelvin
and Ekman numbers. The Kelvin number, Ke ¼ b=Ri , indicates
whether the Earth’s rotation affects the exchange flow in a
channel by comparing the cross-sectional width of the basin, b, to
the internal Rossby radius, Ri. The Ekman number, Ek ¼ Az =ðfH2 Þ,
is a measure of frictional forces to Coriolis forces where Az is
the eddy viscosity, H is the maximum water depth, and f is the
Coriolis parameter.
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3.6. Moored ADCP data
A bottom-mounted ADCP, equipped with a pressure sensor,
was deployed at the entrance to West Pass (3017.3060 N,
85143.7830 W; Fig. 1). The ADCP was moored in the deepest section
of the channel at a depth of 12.1 m from January 10, 2008 to March 7,
2008, measuring a 57-day time series of current velocity profiles. The
ADCP averaged 120 pings distributed over a 6-min interval in 1 m
Fig. 11. Wind (m s 1) and tidal amplitude (m), depth averaged subtidal along channel currents, subtidal along and cross-channel currents of the moored ADCP located at
the entrance of West Pass (cm s 1). Wind speed and direction are plotted in oceanographic convention. Depth is shown in meters (m). Positive (negative) along channel
residual currents indicates flow into (out of) the inlet from (to) the Gulf of Mexico. Positive cross-channel currents indicate flow toward the Northwest.
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bins. Currents were rotated 451 to the principal axis of maximum
variance such that positive along channel velocities represented flow
into the inlet,while negative velocities indicate flow out of the inlet,
toward the Gulf of Mexico. Subtidal variation of the currents was
determined using a low-pass Lanczos filter with half-power of 34 h
(Emery and Thomson, 2004).
4. Results and discussion
4.1. External forcing
Winter winds in St. Andrew Bay during the beginning of
2008 were characterized by frontal systems with a periodicity of
4–6 days (Fig. 3). Twenty-four hours before the towed ADCP
survey, moderate northwest winds (reaching a maximum of
6.5 m s 1) were observed. Two hours after the initiation of the
survey, wind speeds decreased, and ranged from 0 to 3.5 m s 1 for
30 h. Water level at the Deer Point Reservoir was 1.52 m
throughout the towed survey resulting in a higher than normal
freshwater discharge of 22 m3 s 1 into North Bay.
4.2. Tidal wave and ellipses
The phase derived from cross-spectral analysis between surface
elevation and along channel currents from the moored ADCP
indicated that currents in the inlet lead surface elevation by 4.5 h
(671; Fig. 4 (middle)). Consequently, the tidal wave at the West Pass
entrance had characteristics more suggestive of a standing wave
rather than a perfectly progressive wave (which has an expected
phase lag of 01). The phase lag between surface elevation and along
channel velocity corresponded with measured sectionally averaged
velocities and water level observations obtained during the survey
at the local tide gauge (Fig. 4 (bottom)).
Tidal ellipses calculated from the towed ADCP indicated a
diurnal variation (K1) with strongest flows midway between tidal
height extremes (Fig. 5). Orientation of the ellipses was similar at
all depths. Major axes of the K1 tide were 100 cm s 1 and minor
axes were 10 cm s 1 at all depths. Ellipses in West Pass for both
the K1 and M2 tides were rectilinear, oriented in the direction
of the channel. The K1 ellipses became less rectilinear approaching
the Gulf of Mexico from inside West Pass as well as toward the
shallow east side of West Pass outside of the main shipping
channel. This is indicative of bottom friction effects on tidal flow
similar to the results found by Valle-Levinson and Lwiza (1995). In
this system, the flood tide enters St. Andrew Bay as a jet,
encounters the northern boundary, and bifurcates up East or
West Bay (Murphy and Valle-Levinson, 2008). The orientation of
tidal ellipses adjacent to the connection of West Pass with
St. Andrew Bay illustrates the source/sink model of Stommel
and Farmer (1952) working on the opposite side of the inlet,
where ebb tide acts as a sink of water from the bay and the flood
tide acts as a jet entering the bay.
The phase of the K1 tide increased from 01 on the east side of
West Pass to 151 on the west side of West Pass corresponding to a
time difference of 1 h. The surface tidal phases were similar to
those observed by Uncles (1988) in the Bristol Channel and the
Severn Estuary where the flow at the coastline led deeper parts of
the channel. In our survey, co-phase lines at depth exhibited the
same across channel difference, but led the surface phases at the
same location.
4.3. Density field
During the towed survey, dense water from the Gulf of Mexico
(r =1026 kg m 3) entered West Pass during flood tide creating a
well-mixed water column (Fig. 6). The density profile stratified as
Fig. 12. Variation in sub tidal along channel velocities (cm s 1) with tidal amplitude (top) and wind direction (bottom). Included in the bottom panel is the variation in the
subtidal mean sea level height with wind direction. Wind direction is in oceanographic convention. A minimum wind cut off of 2.5 s 1 was used in calculating the observed
variations with wind direction.
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the tide progressed toward ebb (14.9 days). Stratification became
stronger through ebb tide with lowest densities observed between
15.3 and 15.5 days for B, and at 15.5 days for D0 , toward the end
of ebb.
As a result of density variations, j increased at 15.0 days
(Fig. 7) coinciding with the initiation of ebb tide at both stations
(Fig. 6). j at Station B0 reached a maximum after 15.4 days.
Following this peak, stratification decreased as flood tide began.
As less dense water pushed the higher density water out of West
Pass, j gradually increased, reaching a maximum at the end of the
sampling period. Comparisons between the change in j with time
(as observed from the density field) and the change in j predicted
by tidal straining at Station B (Fig. 7) showed a consistent pattern
despite smaller scale fluctuations. Consistencies between the
same parameters at Station D were not as evident, implying that
other forcing mechanisms, such as the advection of a vertical
gradient, were likely affecting stratification of the water column.
Both parameters were consistent, except at 15.1 days when j
peaked. The tidal asymmetries in tidal straining are expected to
influence the observed residual exchange flow within the inlet
(e.g. Stacey et al., 2008).
4.4. Residual flow
The cross-sections of Transects 1 and 2 showed a two-layer
exchange flow with flow entering the bottom and exiting at
the surface. The isotachs associated with the residual flows
across Transects 1 and 2 had a positive lateral slope (looking into
St. Andrew Bay, Figs. 8 and 9). The slope of the interface between outflows and inflows, the zero isotach, was 0.02 for
Transect 1 (5 m depth difference in 250 m across the inlet) and
0.05 for Transect 2. The sign of these slopes was consistent
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with Coriolis accelerations. As a first diagnostic of determining
how much of this isotach slope was produced by a geostrophic
balance across the inlet, the Margules’ slope g was determined as
(Gill, 1982)
g¼ f ðu1 r1 u2 r2 Þ
gðr2 r1 Þ
ð5Þ
where f is 7.3 10 5 s 1, g is 9.8 m s 2, u1, r1 and u2, r2 are the
upper and lower layer flows and densities, respectively. In
Transect 1, taking nominal values for u1 of 0.08 m s 1, r1 of
1025 kg m 3, u2 of 0.05 m s 1, and r2 of 1026 kg m 3 yielded a
value of g of 0.001, which was one order of magnitude smaller
than the observed slope. Similarly, taking the same r2 and r1 of
1025.5 kg m 3 for Transect 2 plus u1 of 0.15 m s 1 and u2 of
0.02 m s 1, resulted in g of 0.003. These estimates of the
geostrophic slope being much smaller than the observed slope
indicated that friction and advection could have contributed to
shaping the exchange flows at the inlet (e.g. Valle-Levinson and
Lwiza, 1995). This will be explored further with analytical
solutions in Section 4.5.
It is noteworthy that the maximum along channel velocities in
Transect 2 were 50% larger than those in Transect 1. This was
not really an increase in volume outflow from Transect 1 to
Transect 2, but rather it was the result of the flows being more
channelized in Transect 2. At Transect 1 the main orientation of
the flows was toward the western side of the channel (Fig. 8) and
between the two transects the mean flow was affected by
recirculation caused by the complicated morphology of the
inlet (e.g. Li et al., 2006). In fact, the analytical model results
presented in the following section required the prescription
of the same volume inflow (river discharge of 40 m3 s 1) for
both transects, which indicated volume conservation from
Fig. 13. Variation in subtidal cross-channel velocities (cm s 1) with tidal amplitude (top) and wind direction (bottom). Wind direction is in oceanographic convention. A
minimum wind cut off of 2.5 m s 1 was used in calculating the observed variations with wind direction.
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Transect 1 to 2 in the observations. The analytical model allowed a
comparison of the combined effects of Coriolis and friction on the
mean flows. The discrepancies between analytical results and
observations could then be attributed to advective effects, as
explained next.
4.5. Friction and Coriolis importance at the inlet
The analytical solutions (as in Valle-Levinson, 2008) were
calculated in terms of the Kelvin and Ekman numbers for the
bathymetries of Transects 1 and 2. For Transect 1, at the northern
end of the inlet, a reasonable match between observed flows and
analytical solutions (Fig. 8) was obtained for an Ekman number of
0.01 and a Kelvin number of 0.8. The value of the Ekman number
corresponded to a vertical kinematic eddy viscosity coefficient of
2 10 4 m2 s 1. These values indicated that at the northern end
of the inlet, both rotation and friction were influential to the
exchange hydrodynamics and the internal radius of deformation
was of the order of the inlet’s width. The bathymetry had mild
lateral variations that were not influential in the overall water
exchange pattern. Not only the shape of the exchange flows but
also the magnitude of the observed mean flows compared well
with the analytical results, including the depth at which the flow
reversed.
For Transect 2, at the southern end of the inlet and the
transition with the Gulf of Mexico, a reasonable match was
obtained between the analytical solution and the observed mean
flows (Fig. 9). This match was attained with an Ekman number of
0.03, corresponding to an eddy viscosity of 1 10 3 m2 s 1, and a
Kelvin number of 1.8. The increased values of Ekman and Kelvin
numbers, relative to the values at Transect 1, produce the
increased slope in the isotachs. It is evident that frictional effects
were stronger at Transect 2 than at Transect 1 as shown by
decreased stratification and stronger tidal currents at the southern end of the inlet (Fig. 6). The analytical solution for Transect 2
compares favorably with the magnitude of the exchange flows and
the lateral slope of the isotachs. The dynamics at West Pass could
then be characterized by weak to moderate frictional and Coriolis
effects. The role of friction becomes more significant at Transect 2
than at Transect 1. Which is consistent with the dynamics
described in the main part of St. Andrew Bay (Murphy and
Valle-Levinson, 2008).
Discrepancies between observations and the analytical
solution can be attributed to curvature effects (or advection),
which are neglected in the analytical model’s momentum
0.03
2
b)
0.02
4
Depth (m)
0.01
0
6
−0.01
Vertical
Structure
1st Mode
8
−0.02
−0.03
01/13
First Mode (93%)
Along shelf wind stress (Pa)
01/20
01/27
02/03
02/10
02/17
02/24
10
0.2
03/02
0.5
0.5
Depth (m)
4
0.3
0.2
0
01/13
0.4
2
0.4
0.1
0.3
6
Vertical
Structure
2nd Mode
8
Second Mode (6%)
Tidal Amplitude (m)
01/20
01/27
02/03
02/10
02/17
02/24
03/02
10
−0.5
0
0.5
Fig. 14. EOF of the along channel subtidal velocity. (a) The first mode of variability (solid) with the along-shelf wind stress (dashed), (b) the vertical structure of the first
mode, (c) the second mode of variability (solid) with the tidal amplitude (dashed), and (d) the vertical structure of the second mode.
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balance. Recent studies (e.g., Chant, 2002; Li, 2006; Webb et al.,
2007) have shown that advection may be significant in complex
estuarine systems. In West Pass, lateral advection enhances
the flow structure and should contribute to the recirculation
observed in the residual flow (Fig. 10). The discrepancies between
the flow observations and the analytical model imply that
advection is a relevant mechanism at West Pass. The residual
flow depicted here over a single diurnal tidal cycle is only a
snapshot within a variety of patterns that will result from wind
forcing and spring–neap tidal forcing. These mechanisms are
further explored with current profiles (moored ADCP) recorded
over 57 days.
4.6. Temporal variations in exchange flow
Moored ADCP measurements in West Pass were used to
determine the vertical structure of the net flow as well as its
temporal variability (Fig. 11). The vertically averaged flow was
well correlated with wind forcing. The correlation coefficient
between along shelf winds and along channel currents was 0.72.
Visually, the depth-averaged currents were most noticeably
correlated with the wind after January 27, 2008 indicating that
the variability of the along channel subtidal flows was dominated
2329
by wind forcing. Subtidal along channel currents had maximum
outflows of 37.3 cm s 1 and inflows of 19.8 cm s 1 whereas
across channel currents ranged between 9.0 cm s 1 and
7.7 cm s 1 (Fig. 11). Throughout the period of observation,
along channel subtidal flows were most often unidirectional
with depth and well represented by the vertical average. However,
this vertical structure changed between unidirectional to less
frequent exchange flow with a periodicity related to both
the fortnightly modulation of the tide and along-shelf wind
stress. During the survey of February 14 to 15, the along
channel subtidal flow basically showed net outflow, consistent
with Fig. 9.
Across channel (secondary) subtidal flows had a distinct twolayer structure with surface flow toward the Northwest (positive)
and deeper flow toward the Southeast (negative), indicative of
curvature induced two-layer lateral flow (Fig. 11). This two-layer
flow was in agreement with the lateral flow from the survey
transects (Figs. 8 and 9). The two-layer structure was prevalent
over nearly the entire observation period except for January 20,
2008 and several events in the last 2 weeks of the deployment.
Along channel winds were negatively correlated (correlation
coefficient of 0.80) to the across channel currents.
To examine the relationship between along channel velocities
and forcing by winds and tides, the variation of the along channel
0.5
2
0.4
Depth (m)
4
0.3
0.2
0.1
0
01/13
6
Vertical
Structure
1st Mode
8
First Mode (65%)
Tidal Amplitude (m)
01/20
01/27
02/03
02/10
02/17
02/24
10
−0.5
03/02
0.06
0
0.5
2
0.04
Vertical
Structure
2nd Mode
4
Depth (m)
0.02
0
6
−0.02
8
−0.04
−0.06
01/13
Second Mode (30%)
Along shelf wind stress (Pa)
01/20
01/27
02/03
02/10
02/17
02/24
03/02
10
0
0.25
0.5
Fig. 15. EOF of the cross channel subtidal velocity. (a) The first mode of variability (solid) with the tidal amplitude (dashed), (b) the vertical structure of the first mode, (c)
the second mode of variability (solid) with the along-shelf wind stress (dashed), and (d) the vertical structure of the second mode.
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P. Murphy et al. / Continental Shelf Research 29 (2009) 2318–2332
(∂ u/∂ z) *10, Tidal Amplitude, Rain
velocities was grouped according to tidal amplitude and along
channel wind direction (Fig. 12). The along channel flow as a
function of wind direction was sorted for low-pass winds
42.5 m s 1, which is the speed at which wind forcing
overwhelms the density-driven forcing. Unidirectional flow
occurred at the weakest neap tides ( o0.125 m) and strongest
spring tides ( 40.375 m) while exchange flow occurred at all other
tidal amplitudes. Along channel currents showed outflow for
winds toward the South-Southeast (901 to 2001) corresponding
with the effects of an offshore Ekman transport. The subtidal sea
level showed a similar decrease for southward to southsoutheastward winds (Fig. 12). The opposite response was
illustrated for winds between northwestward and northeastward.
The magnitude of the lateral flows varied approximately with
the tidal amplitude (Fig. 13(top)). Strongest lateral flows
developed during spring tides (i.e. tides with largest amplitude)
because of strongest centrifugal accelerations, consistent with
Chant (2002). However, the depth of lateral exchange varied with
wind direction (Fig. 13(bottom); calculated using a minimum
wind speed cutoff of 2.5 m s 1). Winds aligned with outflow in the
channel (toward 451 to 2251) resulted in an increase in the depth
of penetration of the northwestward lateral flows. This
penetration was likely due to an increased wind induced
outflow at the surface within the inlet, which correspondingly
increased the lateral structure at the surface. Conversely, winds
heading toward the southwest to northeast (toward 2251 to 451)
resulted in increased bottom layer thickness.
To further examine the vertical structure of the along and cross
channel flow and their temporal variability, an empirical orthogonal function (EOF) analysis was carried out on the time series.
The EOF analysis showed that the first mode portrayed unidirectional net flows. This mode explained 93% of the variability of the
along channel currents within West Pass (Fig. 14b) and was well
characterized by the along-shelf wind stress (Fig. 14a). The second
EOF mode had a two-layer structure with flow into the inlet at the
surface, and outflow below 5 m depth (Fig. 14d). This mode
represented 6% of the variability and was well characterized by
the tidal amplitude (Fig. 14). Similarly, the lateral currents showed
a two-layer structure in the first EOF mode (Fig. 15), which
represented 65% of the variability of the flow, and correlated with
the tidal amplitude (Fig. 15a). The second EOF mode represented
30% of the variability and had a unidirectional vertical structure
with stronger surface flow than at depth (Fig. 15d). The alongshelf wind stress was of secondary importance in driving the
lateral subtidal flow (Fig. 15c).
Further analysis revealed that the temporal evolution of the
along channel subtidal flow structure showed episodic events of
increased vertical shear between the near-surface and nearbottom records (Figs. 11 and 16). The magnitude of vertical
shear showed peaks coinciding with neap tides and had a 0.70
correlation to the spring–neap cycle. Exchange flows within
temperate estuaries, mostly driven by along channel density
gradients, are typically stronger during neap tides (Haas, 1977;
Linden and Simpson, 1988; Griffin and LeBlond, 1990). Unlike
several inlets in the subtropical region of Florida (Valle-Levinson
et al., 2009), West Pass behaved in a similar manner to temperate
estuaries due to the influence of density-driven flows. Several
pluvial precipitation pulses, occurring in conjunction with alongshelf wind stresses, increased the density-driven flows and the
vertical shear (February 18, 19 and 23, Fig. 16). In particular, the
rain pulse on February 23 coincided with neap tides and caused
the largest vertical shears.
The spatial coverage of the survey allowed an evaluation of the
subtidal exchange flow at two transects of West Pass, including an
assessment of the importance of frictional and Coriolis forcing on
the dynamics of the exchange flow. However, the results from the
0.5
0.05
0.25
0.025
0
0
−0.25
Wind stress
2330
−0.025
−1
(∂ u/∂ z) *10 (s )
Tidal Amplitude (m)
Rain (m)
Along shelf wind stress (Pa)
−0.5
01/13
01/20
01/27
02/03
02/10
02/17
02/24
−0.05
03/02
Fig. 16. Temporal evolution of the along channel subtidal vertical shear calculated between surface and bottom bins of the moored ADCP located at the entrance of West
Pass (solid). The tidal amplitude (dotted), the along-shelf wind stress (dashed) and the rain (open circles) over the sampling period are shown.
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P. Murphy et al. / Continental Shelf Research 29 (2009) 2318–2332
survey and from the moored record showed that the observed
exchange flow was infrequent. The subtidal flow at the inlet was
mainly dependent on the along-shelf wind stress and secondly on
the tidal amplitude.
5. Conclusion
West Pass in the St. Andrew Bay system is a unique subtropical
inlet for the study of subtidal flows because of its marked
freshwater influence. The density gradient arising from the
freshwater influence caused tidal straining and tidal asymmetries
in water column stratification. These asymmetries shaped the
spatial structure of the along channel residual flow, which showed
a two-layer exchange with outflow at the surface and inflow at
depth, as in a typical temperate estuary. However, this exchange
pattern was infrequently observed because the net flow was
often unidirectional throughout the water column and mainly
modulated by wind forcing. Also, as found in temperate estuaries,
the shear in the along channel flow was inversely related to
tidal forcing with strongest shears occurring during neap
tides. The lateral flows over the time series showed a distinct
two-layer structure, whose magnitude was modulated predominantly by the spring–neap tidal cycle. The results of this study
indicate that the exchange hydrodynamics at the West Pass are
essentially determined by pressure gradient and friction with
non-negligible influence from Coriolis and advective (centrifugal)
accelerations.
Acknowledgements
Both surveys were accomplished with the collaboration of
University of Florida students and NMFS Panama City Laboratory
personnel. The scientific support of Jung Woo Lee, Tem Fontaine, and
Peng Cheng is appreciated. Boats were safely operated with the
expertise of John Brusher, Andy David, Doug Devries, and Christopher
Gardner. Laura Rear, Karl Cammerer, and Paul Fanelli from NOAA COOPS were generous with their raw ADCP data. AVL acknowledges
support from NSF project OCE-0551923 and OCE-0726697. Alex
Chester provided valuable comments to the original manuscript.
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