Observations of barotropic and baroclinic exchanges in the lower Chesapeake Bay

Pergamon

ContinentalShelfResearch, Vol. 15, No. 13, pp. 1631-1647, 1995 Copyright© 1995 Elsevier Science Ltd Printed in Great Britain. All rights reserved 0278--4343/95 $9.50 + 0.00

0278--4343(95)00011-9

Observations of barotropic and baroclinic exchanges in the lower Chesapeake Bay A R N O L D O VALLE-LEVINSON* (Received 1 February 1994; in revisedform 14 December 1994; accepted 28 December 1994) Abstract--Current meter observations from four instruments deployed in two moorings in the area of the Chesapeake Bay mouth have been analyzed to relate the flow structure to meteorological and tidal forcing. The current velocities and water temperatures recorded by these instruments have been compared to meteorologial and sea level data from a tide gauge located in the same area. The data were collected in the period July-September 1993, during the passage of hurricane Emily over the coast of North Carolina and off the coast of Virginia. Intratidal velocity fluctuations coincide with near-bottom temperature variations. Near-bottom temperature oscillations are of greater magnitude than those near the surface and are predominantly semidiurnal. During flood periods, bottom temperature typically drops 6°C with respect to its ebb value. Temperature behavior suggests self-adjustment of a longitudinal temperature gradient during neap tides and tidal advection of such gradient in spring tides. Low-pass filtered fluctuations of temperature and flow, and hence water exchange at the bay mouth, appear to be caused by the superposition of wind events and gravitational circulation modulated by the spring-neap tidal cycle. Wind stress produces a barotropic response on the residual flow. Southwestward winds drive coastal ocean water into the bay and northeastward winds drive water out of the bay. The development of gravitational circulation, near-surface outflow and near-bottom inflow, occurs during neap tides.

INTRODUCTION The study of water exchange between estuaries and the adjacent coastal ocean has received increased attention in the past two decades. The exchange between these two systems determines the long-term transport of dissolved and suspended matter into and out of an estuary. While it is recognized that water exchange is carried out through barotropic and baroclinic processes, considerable effort has focused on describing the barotropic exchange produced by meteorological forcing, namely the wind stress. Weisberg (1976) found that wind effects can be of equal or greater importance to the tidal or the gravitational circulation in the Providence River of Narragansett Bay. Smith (1977) pointed out that meteorologically forced exchanges between the Gulf of Mexico and Corpus Christi Bay are due to Ekman transport. Wang and Elliott (1978), and Wang (1979a,b) demonstrated that the lower Chesapeake Bay responds barotropically to local (north-southward) winds and coastal Ekman flux. This coupled bay-shelf system response to meteorological forcing has been shown to occur in Sandsfjord, Norway (Svendsen, 1980); Alberni Inlet, British Columbia (Stucchi and Bell, 1980); the Strait of Juan de Fuca *Center for Coastal Physical Oceanography, Old Dominion University, Norfolk, VA 23529, U.S.A. 1631

1632

A. Valle-Levinson

(Holbrook et al., 1980); San Francisco Bay (Waiters, 1982); Delaware Bay (Wong and Garvine, 1984); and Great South Bay, NY (Wong and Wilson, 1984). The baroclinic exchange produced by the density contrast between the estuary and the coastal ocean has also been described extensively. This exchange is not constant in time but can be modulated by the spring-neap tidal cycle (e.g. Geyer and Cannon, 1982; Nunes and Lennon, 1987; Griffin and Le Blond, 1990). In the Chesapeake Bay, a fortnightly modulation of water column stratification and residual flow has been depicted by Haas (1977) for the estuaries of the James, York and Rappahannock rivers. This modulation has not been described in the area of the Bay entrance. The dominance of meteorologically driven exchanges over gravitational circulation at the entrance to Chesapeake Bay has been proposed by Goodrich (1987, 1988). The possibility of periodic events that increase the importance of gravitational circulation, however, has been less studied. The purpose of this study is two-fold. First, to describe the intratidal and monthly variations of flow in the lower Chesapeake Bay during summer conditions as they relate to meteorological and astronomical forcing. Second, based on such flow variations, to define periods when barotropic and baroclinic exchanges occur. This is done through the analysis of current velocity and temperature data recorded at two stations across the Chesapeake Bay Mouth and wind velocity, atmospheric pressure and sea level data recorded at a nearby station. Summer conditions are of particular interest in the lower Chesapeake Bay because summer is when fish and crab larvae enter the estuary from the adjacent ocean. Also, this is the time when wind forcing is at its annual weakest and when baroclinic exchanges may be important. Study area and observations

The lower part of the Chesapeake Bay has received increased attention in the last few years. Two of the main reasons for this increased interest are: (1) it is the region where the dissolved and suspended materials, as well as fish and crab larvae, enter the Bay from the adjacent ocean, and (2) appropriate boundary conditions for the numerical models of the Chesapeake Bay need to be established in this area. With these ideas in mind, an observation program with moored current meters was designed and conducted during the summer of 1993. The current meter moorings, Stas 1 and 2 (Fig. 1), were located approximatley 200 m seaward of the Chesapeake Bay Bridge-Tunnel (CBBT). A tide gauge and meteorological station maintained by the National Oceanic and Atmospheric Administration (NOAA) on a pier off the CBBT (Fig. 1) provided sea level, wind and barometric pressure data throughout the period of current velocity measurements. The start date of recording for the instruments at both moorings was 21 July 1993. The instruments at mooring 1 recorded through early 13 September 1993 and at mooring 2 through 31 July 1993. Each mooring was deployed in about 10 m of water with instruments placed about 2 m from the surface and 3 m from the bottom (Fig. 2). The current meters employed were Sensor Data model SD-2000, which record water temperature, and water current speed and direction. These instruments record direction with a precision of 15° Tand an accuracy of +7.5 °. The precision of the speed rotor, which is similar to those of the Aanderaa instruments, is 0.01 ms 1 and its accuracy is +0.02 m s -1. The speed rotor of these instruments is prone to gravity wave pumping that biases high the recorded speeds. The wave heights during the summer months of this study

Barotropic and baroclinic exchanges in the lower Chesapeake Bay

1633

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Fig. 1. Lower Chesapeake Bay showing approximate bathymetry and current meter mooring locations, Stas 1 and 2, just seaward of the Chesapeake Bay Bridge-Tunnel (CBBT). The tide gauge and meteorological station is marked with an M. The bathymetry shows two channels which coincide with the location of the tunnels of the CBBT. Bathymetry contour interval is 10 m. m u s t h a v e r e m a i n e d r e l a t i v e l y low such t h a t r o t o r p u m p i n g effects m u s t h a v e b e e n small. T h e p r e c i s i o n o f the t h e r m i s t o r is 0.01°C a n d it is a c c u r a t e to + 0 . 1 ° C . C u r r e n t velocity is r e c o r d e d via in situ p r o c e s s e d v e c t o r a v e r a g i n g . T h e s a m p l i n g interval for t h e c u r r e n t m e t e r s was 20 rain a n d for t h e i n s t r u m e n t s at N O A A ' s s t a t i o n was 6 rain. All r e c o r d s w e r e r e d u c e d to o n e h o u r intervals. DATA

REDUCTION

T h e i n t r a t i d a l b e h a v i o r was d e s c r i b e d f r o m an a r b i t r a r y p e r i o d o f the r e c o r d at Stas 1 a n d 2 d u r i n g t h e first 11 days. T h e v e l o c i t y was d e c o m p o s e d into p r i n c i p a l axis a n d

1634

A. Valle-Levinson ..

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cross-principal axis component using the tidal currents analysis program developed by Foreman (1978), which uses least squares analysis. The orientation of the principal axis was determined from the semi-major axis of the M2 ellipse. At Sta. 1, this angle was 138 ° T near the surface, and 146 ° T near the bottom. At Sta. 2, it was 164 and 159 ° T for nearsurface and near-bottom, respectively. This principal axis distribution suggests flow nearly-parallel to the coast in this area, i.e. it tends to follow the bathymetry of the region. The long term behavior was described with the low-pass filtered records of water temperature, decomposed current velocity, wind velocity, sea level and complex demodulated amplitude of sea level (to denote tidal phase). The filter consisted of a Lanczos filter with a half-power of 35 h to remove diurnal, semidiurnal and other high frequency fluctuations. The wind stress or momentum flux is parameterized with the 'quadratic law'. The eastwest, rsx, and north-south, rsy, components of the wind stress are given by:

= paCDWxlolW, ol Tsy = pctfDWylo[Wlol

(1) (2)

where p~ is the air's density (1.2kgm-3); W is the wind magnitude (m s -1) with components [Wx, Wy]; CD is a non-dimensional drag coefficient and equals (0.49 + 0.065W10) × 10 - 3 for W ~> 1 0 m s - ' , and 1.14 × 10 - 3 for W < 1 0 m s -1 (Large and Pond, 1981). The subscript 10 denotes wind values at an elevation of 10 m above the sea surface.

1635

Barotropic and baroclinic exchanges in the lower Chesapeake Bay

DATA DESCRIPTION Ocean water temperature is lower than bay water temperature during the summer as shallow, enclosed bodies of water gain (lose) heat faster than the adjacent coastal ocean. In the records presented here cool oceanic water is advected landward during flood periods (Fig. 3). The intratidal records at Sta. 1 show that minimum bottom temperatures coincide with the end of flood periods. Ebb period speeds are of greater magnitude ( - 1 m s- l) than flood period speeds (-0.70 m s-l). This ebb-flood asymmetry is evident between 27 July and 29 July. During those three days the landward advection of low temperature water at flood periods is small. The surface water temperature drop on late 26 July is associated with a relatively strong flood speed (approx. 0.80 m s -1) that was reinforced by northwestward (same direction as flood) wind earlier that day. Throughout the record, the nearbottom 7142tidal constituent leads the near-surface by 12 min due to bottom friction. The flow at Sta. 2 was faster (>1 m s -]) and less rotary than at Sta. 1 as shown by the small amplitude of the cross-axis component (Fig. 4). The water column at Sta. 2 was less thermally stratified than at Sta. 1 indicating more active vertical mixing in the area of Sta. 2 and/or reduced influence of buoyant water due to rotational effects that maintain buoyant 28

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Fig. 3. Intratidal record for the first 11 days of deployment at Sta. 1. The four panels from top to bottom are: near surface (T~ bold), and near-bottom ( Tb grey) temperature records (°C); near-surface (u s bold), and near-bottom (u b grey) cross-principal axis component of the current velocity (cm s 1); near-surface (vs bold), and near-bottom (v b grey) principal-axis component of the current velocity (cm s 1); and wind velocity (m s-1) (stick diagram---oceanographic convention), and sea level variations (m) (grey) at Sta. M.

---

1636

A. Valle-Levinson

waters in the area of Sta. 1 (along the fight boundary looking seaward). Surface temperature at Sta. 2 was lower and more responsive to tidal variations than at Sta. 1, where the surface layer appeared isolated from tidal fluctuations. Similarly, bottom temperatures at Sta. 2 were higher than bottom temperatures at Sta. l: a consequence of vertical mixing. At Sta. 2 the phase of the near-bottom Me tidal constituent preceded the near-surface by 8 rain. Although salinity, more than temperature, is the factor that determines density in this area (Miller and Valle-Levinson, in preparation) and the parameter that indicates the degree of water column stratification, the summer intratidal variations of temperature are expected to be related to those of salinity, i.e. large (small) salinity variations are associated with large (small) temperature variations. Such is the case in other estuaries [e.g. Long Island Sound (Valle-Levinson, 1992)], at a mid-Chesapeake Bay station (ValleLevinson, 1988), and at a Chesapeake Bay mouth station (Fig. 5). Large salinity increments correspond with large temperature reductions and vice versa owing to the gradient advection by the tidal currents. Hence, the changes in thermal stratification observed here should correspond with variations of salinity and density stratification. The point to make here is that the temperature records in the present case can be used to describe stratification patterns. Thus, the region near Sta. 1 is expected to be more stratified than that of Sta. 2 and stratification tends to be greater during neap tides than during spring tides as shown in the low-passed records below.

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The dominance of ebb currents over flood flows, the influence of the wind stress, and the spring-neap modulation of the tidal currents translate into the following low-passed or residual fields in temperature and flow. A clear correspondence between wind stress variations and sea level fluctuations is evident (Fig. 6). South-southwestward wind corresponds with rising sea level, and north-northeastward wind coincides with dropping sea level. This behavior is considered in further detail in the next section. The surface net flow at Sta'. 1 was directed seaward during the observation period, except during 6 September. B o t t o m net flow was also dominantly seaward but exhibited periods of net inflow. Neap tide periods corresponded with landward bottom flow, i.e. intrusion of oceanic waters. This was corroborated by the temperature records, which showed, during neap tides, decrease of bottom temperature and little change of surface temperature. This behavior during neap tides suggests increased thermal stratification induced by near-bottom inflow of oceanic waters that are cooler than Chesapeake Bay's waters during the summer. In addition to neap-tides related intrusions, there were three other distinct events of net

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Barotropic and baroclinicexchangesin the lowerChesapeakeBay

1639

landward near-bottom flow during the observation period. The bottom inflow on 1-2 August occurred after southward wind piled water up against Norfolk's coast, thus causing a baroclinic response to the sea level set-up. The temperature record on these two days shows near-bottom decrease but little change near the surface. The other two bottom inflow periods (5-6 August and 6 September) occurred after the two events of strongest northeastward wind caused the greatest sea level drops of the observation period. The loss of water caused by the sea level drop was compensated by net inflows. On 5-6 August, near surface outflow was weakened and on 6 September (strongest wind) it was reversed. Both events show decreased temperature at the two depths measured as a consequence of the net barotropic inflow. Hurricane Emily passed near the study area on 30 and 31 August, causing the greatest sea level increase of the sampling record (-0.30 m). The passage of the hurricane also produced decreased near-surface and near-bottom outflow and a short period of very weak bottom inflow (-0.01 m s -1) with the brief southwestward winds of 30 August. The effect of the passage of the hurricane was not as dramatic as other events in the year (e.g. winter northeasters) due to the transitional character of the winds.

SEA LEVEL VARIATIONS DUE TO METEOROLOGICAL FORCING The relation between sea level changes and meteorological forcing (wind stress and atmospheric pressure) is explored in this section to illustrate the barotropic response of the flow. Wang (1979a) found that in the lower Bay, velocity fluctuations are mainly barotropic and can be related to sea level changes. Wang and Elliott (1978) suggested that water is driven out of (into) the Bay by northward (southward) wind. These findings are explored further through the assessment o f the effects of the surface and bottom stresses, as well as of the atmospheric pressure on producing the observed sea level fluctuations. This is done with the vertically-integrated equations of motion. For a right-handed coordinate system the momentum balance can be expressed in vector notation: 0U + (U. V)U + gV~ + gV

Ot I

II

III

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pdz + 212 x U = VM,

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V

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VI

where U is the fluid velocity vector with components u, v; ~ and p are the fluid's surface elevation and density, respectively, V is the Nabla or Del operator denoting a horizontal gradient; II is the earth's angular frequency; and F M represents turbulent shear stresses and can include horizontal and vertical momentum fluxes. The depth integrated (from the bottom z = - h to the surface z = ~) equations of motion with the inclusion of the atmospheric pressure Pa (Pa) horizontal gradients, and considering only the barotropic contribution to the pressure gradient, are

OU+ 0 ( U 2 ) +

o-?

0__( U V ] _ f V

=

+fv=

(h + ~) OP,~ O~ + rsx - %______2_~ (4) P Ox g(h + ~) O-x p (h + ~) OP~ p Oy

0__~ rsy

g(h + ~) Oy +

-p %>'

(5)

1640

A. Valle-Levinson

O___U+ OV + 0__~= 0 Ox Oy Ot

(6)

where U and V are the cross-principal-axis (x) and principal-axis (y) components of the depth integrated transports: U=

I

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u dx,

V=

I

v dy

--h

rbx and 7Jby are the components of the momentum flux at the bottom (bottom stress); g is the acceleration due to gravity; f is the Coriolis parameter; and p is a reference density (1020 kg m-a).

Effect of wind and bottom stresses Considering the special case of the principal axis component in the Chesapeake Bay (U = 0) at steady state, and neglecting non-linear terms, the effects of the atmospheric pressure and those due to the earth's rotation, then (5) takes the form: O(~) _ (~h) -- (~m) _ (rsy) -- (rby) Oy Ay gp(h + ~) '

(7)

which describes the subinertial variations in sea level slope inside the estuary as a function of surface and bottom stresses, and water column depth. The brackets in (7) denote residual quantities, which are represented here by low-passed values. The terms (~h) and (~m) represent sea levels at the head and mouth of the estuary, respectively, separated by a distance Ay = 200 km. The term (rby) is parameterized through a linear formulation and equals Or@b) where r is a drag coefficient (4 x 10 -4 m s -1) and (Vb) is the observed bottom residual flow (principal axis component) at Sta. 1. The wind stress (rsy) is obtained from the wind velocities at Sta. M. The sea level variations at the Bay mouth, which are the same as the coastal sea level variations and not a result of the influence of local winds (Wong and Garvine, 1984), are produced, among other forcing agents, by the influence of regional winds. The orientation of the Bay mouth (-45°T) makes (~m) particularly sensitive to regional southwestward and northeastward winds. Southwestward (northeastward) winds drive water directly into (out of) the Bay mouth. Southward (northward) winds, which are less frequent than northwestward and southwestward winds in the lower Bay area (Paraso and Valle-Levinson, in preparation), can produce barotropic inflows (outflows) due to coastal Ekman transport (Wang, 1979a,b). Note from (7) that, inside the estuary, a wind stress with a northward (southward) component would cause the sea surface to slope up towards the head (mouth) and a sea level drop (rise) at the mouth. This is illustrated in Fig. 7, where the predicted variations in sea level difference inside the estuary ((~h) -- (~m)) mirror the observed sea level variations at the Bay mouth, i.e. positive sea level slopes coincide with sea level drops at the Bay mouth and vice versa. The lagged correlation between observed along-estuary wind stress and predicted sea level difference inside the estuary ((~h) -- (~m)) is greatest (0.83) at 0 h lag. This indicates that the subinertial changes in sea level in the lower Bay are produced mainly by wind forcing. Other factors that may influence sea level to a lesser extent in this

1641

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Fig. 7. Observed sea level variations at the mouth of the Chesapeake Bay (m) compared to the variations in slope inside the estuary produced by surface and bottom stresses. Only the N-S component of the wind and the principal axis component of the flow are considered in the calculation.

area during the summer months are bottom stress and barometric pressure. The calculated variations of {~h) -- (~m) lead the observed sea level oscillations at the mouth by -10-20 h, which indicates that it takes the lower Bay less than one day to respond to wind forcing. This differs from the studies of Wang (1979a,b) only in the apparent lack of response to forcing with periods greater than 7 days. During the summer months observed, the most energetic events (1--4 August, 24-28 August, 31 August-2 September and 3-5 September) produced responses with periods of 5 days or less. Of course, fall and winter wind events should be expected to be more energetic and to produce longer-period response than those in the summer•

Effect of atmospheric pressure The exclusive effect of the atmospheric pressure on the sea level is [from (5) with similar assumptions as above]

o<~)_ Oy

1 o
If A~ and AP a represent changes over a finite horizontal distance Ay, then A<~) = -- ! A
(8)

This relation indicates that increases in atmospheric pressure are followed by decreases in sea level, and vice versa. This is the "inverted barometer effect". The inverted barometer effect was relatively small during this time of the year in the

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A. Valle-Levinson

lower Chesapeake Bay as compared to the wind stress effect. Some of the surface elevation changes not accounted for by the calculation with (7) were due to the barometric effect. A high pressure system with a peak of 1.026 x 105 Pa (1,026 mb) on 9-10 August was responsible for the observed sea level drop despite the southwestward winds that, according to (7), tend to increase sea level at the mouth. The barometer effect is relatively large during fall and winter, when the atmospheric pressure systems are more extreme than during the summer. For instance, on 1 December 1993 a high pressure system peaked at 1.040 x 105 Pa and kept the sea level near its predicted value despite southward winds of up to 10 m s -1. Four days later, on 5 December, a low pressure system had a low value of 0.987 × 105 Pa, which produced a 0.30 m increase in sea level despite northward winds (up to 14 m s -1) that tend to depress it in this area. As discussed above, the barotropic response of the sea level to meteorological forcing has been described for this part of Chesapeake Bay. The study of Wang (1979a) examined data obtained during late fall when the winds are stronger than during summer. The summer data analyzed here do not suggest influence of long-period (>7 days) forcing on sea level variations. This behavior, however, should be subject to interannual variability. In addition to the barotropic response of the flow to meteorological forcing, these data also suggest a baroclinic response that becomes evident during neap tides. Goodrich (1988) suggested that in addition to meteorological flushing, other processes operate at the estuary mouth. He speculated that an oscillatory wind-driven circulation could be superimposed on a low amplitude gravitational circulation in the lower Chesapeake Bay. The baroclinic response in the form of a gravitational circulation is dealt with in the following section. BAROCLINIC RESPONSE The baroclinic response is investigated using two approaches. The first looks at the residual flow vertical shears. High shears indicate tendencies for gravitational circulation and baroclinicity. Low shears represent unidirectional, barotropic flow. The second approach is through the examination of the residual momentum balance at the two depths sampled. Subtraction of surface momentum equation from bottom momentum equation eliminates barotropic tendencies of the flow and yields the contribution of inertial, rotation, and frictional forces to the observed residual flow. The vertical difference between near-bottom and near-surface low-passed flows ((Vb) -- (V~)), which represents the vertical shear, tends to increase during neap tides and to decrease during spring tides (Fig. 8). This suggests the development of baroclinic (barotropic) exchanges during neap (spring) tides. The high shears on 2 August were produced after a southward wind piled-up water against Norfolk's coast and induced bottom inflow. In general, small differences between near-surface and near-bottom values are indicative of the unidirectional character of the residual flow. Local wind effects are manifested by the 2-3 day oscillations superposed on the longer term variations analogously to the pattern proposed by Goodrich (1988). The influence of the inertial, rotation and frictional forces on the baroclinic pressure gradient in the lower Chesapeake Bay can be studied in terms of their relative contribution to the momentum balance (3). In that equation, the fluid velocity U encompasses a mean or residual part ((u)) and a fluctuating or tidal part (Uo). To understand intratidal and monthly fluctuations in U, two types of momentum balance can be explored. A tidal

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Barotropic and baroclinic exchanges in the lower Chesapeake Bay

momentum balance explains intratidal variations due to Uo, and a residual momentum balance depicts tidal monthly changes due to both Uo and iu). The tidal momentum balance is essentially linear and dominated by terms I, III, V and VI in (3). The residual momentum balance is determined by the average, over one or several tidal cycles, of terms II, III, IV, V and VI. For the principal axis component iv), the Eulerian residual momentum balance is (neglecting a/ax and u for argument's sake) iv)

+ g Oy

Po

(9)

p dz = fy. --Z

A

B

An additional contribution to the principal axis momentum balance arises from the Coriolis termfiv) of the orthogonal component balance, wherefis the Coriolis parameter. Residual velocity data from Sta. 1 are used to determine the relative contribution of each A and B to balance the flow generated by the pressure gradient force. This is assessed through subtraction of the near-surface from the near-bottom residual momentum balance for the principal axis component (Valle-Levinson and Wilson, 1994). This subtraction eliminates barotropic influences on the residual flow and, through dimensional analysis, yields residuals proportional to ((Vb)2 -- (vs)2)/Ln (inertial term), f(ivb) -- (Vs)) (Coriolis term), rsy/pHb) (wind stress) and Cd (v2)/Hb (bottom friction term). The scale LR is associated with the baroclinic pressure gradient (scale of the bottom front) taken as 10 km; f equals 8.8 x 10 -5 s -1 for a latitude of 37°; Ca is a non-dimensional bottom drag coefficient with a value of 2.5 x 10 -3, and Hb is the water depth (10 m). 20

o?/z9

08/08

o8/18

o8/2s

09/07

15

E

a2

io

<:

~

5

-5 210

220

230

Julian Day

240

250

1993

Fig. 8. Vertical difference between near-surface and near-bottom residual flows ((vs)- (vb~) (principal axis component). The complex demodulated sea level amplitude is shown in grey as reference to the tidal phase. The vertical shear of the residual flow increases during neap tides and decreases during spring tides suggesting the development of baroclinic (barotropic) exchanges during neap (spring) tides.

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0.80

.~ 0.70 0.60

~

' 0.50

< 0.40 o

0.30

E 1.0

0.8

0.6

0.4

0.2

0.0 210

220

230 Julian Day I993

240

250

Fig. 9. Absolute value of the normalized terms resulting from the subtraction of the near surface from the near bottom residual m o m e n t u m balance. The upper panel shows the complex demodulated sea level amplitude to indicate the tidal phase. In the bottom panel, the bottom stress is the shaded area. The inertial term is shown as the bold line and the wind stress term is shown as the grey line. This figure suggests the influence of both barotropic and baroclinic mechanisms in determining the residual flow in the area studied.

The vertical momentum exchange terms associated with the tidal and mean vertical shears are proportional to A,((Vb) -- (vs))/H 2, where Av is the eddy viscosity coefficient (m 2 s- 1) and H 2 is the vertical distance between measurements (m). The influence of these terms has been omitted from the present analysis as the coefficients of eddy viscosity cannot be calculated without knowledge of the varying density field (stratification). However, vertical exchange is expected to increase (decrease) during spring (neap) tides (e.g. Griffin and Le Blond, 1990) as a function of eddy viscosity and be an important contributor to the mean momentum. IfAv is assumed constant, then the vertical exchange associated with the vertical shears will vary similarly to the behavior shown in Fig. 8 because of its dependence on (%) - (vs). This is a restrictive assumption and the effects of varying Av should be explored further. If the inertial, surface stress, and bottom friction terms are normalized by the Coriolis term, residuals proportional to ((%)2 _ (Vs)Z)/[fLR((Vb) _ (Vs))] (inertial term), rsy/[fpHb((Vb) -- (Vs))] (surface stress term) and Cd (v2)/[fHv((Vb) -- (Vs))] (bottom friction term) result. The relative magnitude of each term is ascertained through the examination of their absolute values. Although Co varies with the flow speed, the variations of the bottom friction term will be dominated by (v2). The bottom friction term had greater magnitude than the inertial and rotation terms during periods of spring tides (Fig. 9). This indicates hindrance of inertial and rotation accelerations and the pressure gradient force being balanced by friction. During neap

Barotropic and baroclinicexchangesin the lowerChesapeakeBay

1645

tides, inertial accelerations and rotation effects constituted an important contribution to the dynamics of the system as the terms were of similar magnitude as bottom friction. This is similar to the behavior observed in eastern Long Island Sound (Valle-Levinson and Wilson, 1994). During certain neap periods, such as on 13-14 and 27-28 August, inertia and rotation even dominated the dynamics. This suggests that the flow accelerated into the estuary as bottom inflow, probably due to the adjustment of the estuarine longitudinal density gradient. There were four events in which the wind stress dominated the momentum balance (2-3 August, 17-18 August, 3 September and 9 September). The events coincided with relatively high vertical shears and weak bottom stresses. This behavior suggests an exclusive near-surface response to wind forcing during those events. The dynamical behavior of the residual terms suggests that both barotropic and baroclinic mechanisms influence residual flow in the lower Chesapeake Bay. Based on this dynamical behavior, it could be suggested that salt fluxes will show similar variability as the mean flow (Valle-Levinson and Wilson, 1994). During the summer, mean salt fluxes into the estuary will increase during neap periods and after periods of net outflow associated with sea level drops, i.e. the estuary will "rebound" in response to water losses. This is the response observed at one location and is expected to vary across the estuary. Net inflow and robust salt fluxes into the estuary will be more ubiquitous in the deep navigational channels of the lower Chesapeake Bay than in the shallow areas (Valle-Levinson et al., 1994). SUMMARY As part of a study to look at barotropic and baroclinic exchanges in the lower Chesapeake Bay, current meter observations from four instruments deployed in two moorings in the area have been analyzed. The moorings were located approximately 200 m seaward of the Chesapeake Bay Bridge-Tunnel in water depths of about 10 m. The current velocity and water temperature data recorded by these instruments have been compared to meteorological and sea level data from a tide gauge located in the same area. The data were collected in the period July-September 1993, during the passage of hurricane Emily over the coast of North Carolina and off the coast of Virginia. Intratidal velocity fluctuations coincide with near-bottom temperature variations. Nearbottom temperature oscillations are of greater magnitude than those near the surface and are predominantly semidiurnal. Extreme bottom temperature values occur at tidal slack periods: maximum (minimum) temperature at the end of ebb (flood) due to higher water temperatures inside the bay than the coastal ocean during the summer months. During flood periods, bottom temperature typically drops 6°C with respect to its ebb value. In neap tides, semidiurnal temperature variations are not as evident as in spring tides. This behavior suggests self-adjustment of a longitudinal temperature gradient during neap tides and tidal advection of such gradient during spring tides. Low-pass filtered fluctuations of temperature and flow, and hence water exchange at the bay mouth, appear to be caused by a combination of wind events and the spring-neap tidal modulation. Wind stress produced a barotropic response in which southwestward winds drove coastal ocean water into the bay and northeastward winds drove bay water out of the bay in agreement with the studies of Wang (1979a,b) and Wang and Elliott (1978). Extreme events of northward-northeastward winds produced sea level drops that induced

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A. Valle-Levinson

net outflow. After the cessation of these wind events, barotropic net inflows replaced the water lost by the estuary. During neap tides, decreased tidally induced vertical mixing allowed the development of a baroclinic behavior in the residual flow consistent with estuarine circulation: near-surface outflow and near-bottom inflow. Salt fluxes should behave in a similar fashion as net flows. Maximum salt intrusion into the estuary from the adjacent coastal ocean should occur during neap tides and after strong (1>10 m s -1) northeastward winds. The exchange variability described herein applies to one shallow location in the lower Chesapeake Bay. The lateral structure of these exchanges is not uniform across the bay mouth. A pilot study of the flow field in this area during early October 1993 showed that net inflows occurred preferably over the deep navigational channels and that strong lateral shears were associated with abrupt bathymetric changes (Valle-Levinson et al., 1994). In fact, most of the volume inflow recorded during that study appeared over the channels. Hence, future studies on volume exchange in the lower Chesapeake Bay should take into account the influence of the navigational channels. Acknowledgements--Chris Reiss provided invaluable help in the current meter field work and has shown great interest in the progress of this study. R. C. Kidd and Robert Bray contributed important ideas and effort to the design, deployment and maintenance of moorings. Rob Brumbaugh, Tony Colizzi, Margaret Dekshenieks, Jerry Miller, Donnie Padgett and Andry Ratsimandresy helped with the field work at different times. Richard Garvine, Larry Atkinson, Jerry Miller and two anonymous reviewers provided useful revisions to this document. Thanks to the State Council of Higher Education for Virginia for its financial support. Jim Dixon of the NOAA office in Chesapeake, VA, kindly provided the CBBT station data.

REFERENCES Foreman M. G. G. (1978) Manual for tidal currents analysis and prediction. Pacific Marine Science Report 78-6, Institute of Ocean Sciences, Patricia Bay, Sidney, British Columbia, 70 pp. Geyer R. W. and G. A. Cannon (1982) Sill processes related to deep water renewal in a fjord. Journal of Geophysical Research, 87(C10), 7985-7996. Griffin D. A. and P. H. Le Blond (1990) Estuary-ocean exchange controlled by spring-neap tidal mixing. Estuarine Coastal and Shelf Science, 30, 275-305. Goodrich D. M. (1987) Nontidal exchange processes at the Chesapeake Bay entrance. In: Hydraulic Engineering 1987, R. Ragan, editor, American Society of Civil Engineers, New York, pp. 493-498. Goodrich D. M. (1988) On meteorologically induced flushing in three U.S. East Coast estuaries. Estuarine, Coastal and Shelf Science, 26, 111-121. Haas L. W. (1977) The effect of the spring-neap tidal cycle on the vertical salinity structure of the James, York and Rappahannock Rivers, Viginia, U.S.A. Estuarine and Coastal Marine Science, 5,485-496. Large W. G. and S. Pond (1981) Open ocean momentum flux measurements in moderate to strong winds. Journal of Physical Oceanography, 11,324-336. Holbrook J. R., R. D. Muench and G. A. Cannon (1980) Seasonal observations of low frequency atmospheric forcing in the Strait of Juan de Fuca. In: Fjord Oceanography, NATO Conf. Ser. 4, Marine Sciences, H. J. Freeland, D. M. Farmer and C. D. Levings, editors, Plenum, NY, pp. 305-315. Nunes R. A. and G. W. Lennon (1987) Episodic stratification and gravity currents in a marine environment of modulated turbulence. Journal of Geophysical Research, 92(C5), 5465-5480. Smith N. P. (1977) Meteorological and tidal exchanges between Corpus Christi Bay, Texas, and the Northwestern Gulf of Mexico. Eastuarine and Coastal Marine Science, 5,511-520. Stucchi D. J. and W. H. Bell (1980) Shelf-fjord exchange on the west coast of Vancouver Island. Eos Trans. AGU, 61,280. Svendsen H. (1980) Exchange processes above sill level between fjord and coastal water. In: Fjord Oceanography, NATO Conf. Ser. 4, Marine Sciences, H. J. Freeland, D. M. Farmer and C. D. Levings, editors, Plenum, NY, pp. 355-362.

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Valle-Levinson A. (1988) Variability of temperature and salinity at a mid-Chesapeake Bay summer station. M.S. Thesis, Marine Sciences Research Center, State University of New York at Stony Brook, 70 pp. Valle-Levinson A. (1992) Sill processes and barotropic forcing effects on estuary/ocean exchange. Ph.D. Dissertation. Marine Sciences Research Center, State University of New York at Stony Brook, 237 pp. Valle-Levinson A. and R. E. Wilson (1994) Effects of sill processes and tidal forcing on exchange in eastern Long Island Sound. Journal of Geophysical Research, 99(C6), 12,667-12,681. Valle-Levinson A., K. M. M. Lwiza and B. D. Connoly (1994) Flow lateral structure in the lower Chesapeake Bay. EOS, Transactions, American Geophysical Union, 75(16), 198. Walters (1982) Low frequency variations in sea level and currents in South Francisco Bay. Journal of Physical Oceanography, 12, 6584i68. Wang D. P. (1979a) Subtidal sea level variations in Chesapeake Bay and relations to atmospheric forcing. Journal of Physical Oceanography, 9, 413-421. Wang D. P. (1979b) Wind driven circulation in the Chesapeake Bay, Winter 1975. Journal of Physical Oceanography, 9,564-572. Wang D. P. and A. J. Elliott (1978) Non tidal variability in the Chesapeake Bay and Potomac River: evidence for nonlocal forcing. Journal of Physical Oceanography, 8,225-232. Weisberg R. H. (1976) The nontidal flow in the Providence River of Narragansett Bay: A stochastic approach to estuarine circulation. Journal of Physical Oceanography, 6, 345-354. Wong K. C. and R. W. Garvine (1984) Observations of wind-induced, subtidal variability in the Delaware Estuary. Journal of Geophysical Research, 89(C6), 10,589-10,597. Wong K. C. and R. E. Wilson (1984) Observations of low frequency variability in Great South Bay and relations to atmospheric forcing. Journal of Physical Oceanography, 14, 1893-1900.