Dynamics of circulation off the west coast of Vancouver Island

ContinentalShelfResearch, Vol. 16, No. 12, pp. 1591-1607, 1996

Pergamon PII: S 0 2 7 8 - - 4 3 4 3 ( 9 6 ) ~ )

Copyright t~) 1996 Elsevier Science Ltd Printed in Great Britain. All rights reserved 0278-4343/96 $15.130+ 0.00

D y n a m i c s of circulation off the west coast of V a n c o u v e r Island B. K. PAL* and G. HOLLOWAY* (Received 3 May 1995; in revisedform 29 November 1995; accepted 4 January 1996) Abstract--The Princeton Ocean Model (Blumberg and Mellor, 1987) is used to examine dynamical balances in the summer and winter three-dimensional circulation along the west coast of Vancouver Island. The model is initialized with horizontally averaged temperature and salinity over a domain 445 km alongshore by 145 km cross-shore. Forcing is by uniform summer and winter winds and by discharge from Juan de Fuca Strait. A topographic stress parameterization (Eby and Holloway, 1994) is included. In the absence of topographic stress and Juan de Fuca discharge, wind-forced summer circulation is dominated by equatorward flow over the shelf and slope. With topographic stress included, poleward coastal flow develops over the shelf and shelfbreak with a transition zone between poleward and equatorward flow over the slope. When the buoyancy forcing is introduced, an eddy is developed off the mouth of the Juan de Fuca Strait. Winter circulation tends to be poleward for all combinations of wind, buoyancy and topographic stress forcings. These results agree qualitatively with observations. Copyright © 1996 Elsevier Science Ltd.

INTRODUCTION The west coast of Vancouver Island is an extensively studied, yet poorly understood, zone in the northeast Pacific. This area supports highly productive fisheries with commercial stock of salmon, herring, sabblefish, cod and shell-fish. However, recruitment to these fisheries is variable and possibly linked to fluctuations in oceanic water properties and circulation. An improved understanding of the circulation dynamics is sought. Currents along the coast of Vancouver Island are dominated by three quasi-permanent features (Thomson et al., 1989): persistent poleward-flowing coastal current over the shelf; reversible summer/winter wind-driven current over the slope; and the cyclonic Juan de Fuca eddy. There is also evidence for a quasi-permanent poleward flowing California Undercurrent Extension over the continental slope (Hickey, 1979; Mackas etal., 1987). In summer, upwelling is a common feature of the region (Denman et al., 1981; Freeland and Denman, 1982). The dynamical processes responsible for the generation and evolution of the Vancouver Island Current system (e.g. poleward-flowing coastal jet and the Juan de Fuca eddy) are not clear. Forcing mechanisms include wind, tide, buoyancy flux from the Juan de Fuca *Institute of Ocean Sciences, Sidney, B.C., Canada V8L 4B2. 1591

1592

B.K. Pal and G. Holloway

Strait and remote forcing from the Pacific Ocean. In summer, wind blows from the northwest, driving equatorward circulation counter to a poleward coastal flow. In winter, wind blows more from the southeast, driving poleward flow in both coastal and offshore regions. Hickey et al. (1991) investigated the dynamics of the Vancouver Island coastal flow, using hydrographic surveys and current/temperature information in an attempt to identify the buoyancy driving of the coastal flow. Their study showed that buoyancy forcing does not fully account for the poleward coastal flow. Although coastal wind and buoyancy flux are consistent with the winter alongshore continuity of the coastal current (Thomson et al., 1989), it is difficult to account for the continuity of the poleward coastal current during summer. Other forcing mechanisms may be important. Clarke (1989) reviewed various mechanisms for poleward flow in eastern boundary current systems, falling under two categories: those related to wind forcing (e.g. McCreary, 1981; McCreary and Chao, 1985) and those related to topographic stress (e.g. Haidvogel and Brink, 1986; Holloway, 1987; Holloway et al., 1989). Eby and Holloway (1994) showed that the inclusion of topographic stress in a conventional large-scale ocean model generates poleward eastern boundary undercurrents and strengthens equatorward tendencies in deep western boundary currents. Alvarez et al. (1994) reported better agreement with the observations for the western Mediterranean circulation by incorporating topographic stress in their model. Fyfe and Marinone (1995) and Holloway etal. (1995) also reported improvement in their modelled circulations of central Strait of Georgia and Japan Sea, respectively. Much of the modeling effort in this area has concentrated on simulation of tidally generated circulation (e.g. Flather, 1987; Foreman and Waiters, 1990; Foreman et al., 1992). Foreman and Waiters (1990) and Foreman et al. (1992) showed that residual currents due to tidal rectification are small ( - 4 cm s- 1), with eddies over the shallow banks and at the entrance of the Juan de Fuca Strait. However, the residual currents did not produce poleward coastal flow. Foreman et al. (1991) used a three-dimensional diagnostic finite element model to describe residual circulation over the shelf. However, their calculations were based on the assumption of 'depth of no motion' at 300 m. In the present study, a three-dimensional primitive equation model, the Princeton Ocean Model (Blumberg and Mellor, 1987; Mellor, 1993) is applied, to examine the dynamical balance in the circulation by isolating the effect of different forcings such as wind, topographic stress and buoyancy. Remote forcing from the Pacific Ocean is unknown for this model domain, and no attempt was made to introduce this in the model simulation. 2. THE M O D E L The Princeton Ocean Model used here is described in detail by Blumberg and Mellor (1987) and Mellor (1993). This present paper gives only an outline of this threedimensional primitive equation model with a free-surface. The prognostic variables in the model are potential temperature T, salinity S, velocity components U, V and W, and surface elevation t/given by

ou+ Ox

+ Oy

O._UU+ OU2 + O(UV) + O(UW) Ot Ox Oy Oz

Oz

=0

fV . . . . .

10P~_O[K OU]+ F. OzL Oz]

Po Ox

(1) (2)

Dynamicsof circulationoff Vancouver Island

1593

_OV+O(UV) + OV2+ o(VW) t-fU=-IO--~P+~z[KM--~z]+ o r ov1 Fv Ot Ox Oy Oz Po Oy

(3)

OT+ 3(UT) + O(VT) + O(WT) _ 0 [KHOT_] O---t Ox Oy Oz Oz [ -~zJ + Fr

(4)

as + o(us) + o(vs) + o(ws)

0-7

ox

oy

oz

o c

o5 + uO,1 + vO,7 = w

Ot

Ox

osq

- 5-z[K' J + Fs

(5) (6)

Oy

hydrostatic approximation yields P - g(o - z) + Po

fV p _ Pogdz z

(7)

Po

where p = p(S, T), and P0 is a reference density. The vertical mixing coefficients K M and KH are calculated according to Mellor and Yamada's (1982) turbulent closure scheme. The horizontal diffusion terms F,, Fv, F r a n d Fs in equations (2), (3), (4) and (5) use a constant coefficient of horizontal viscosity. For several experiments, rather a large value (500 m 2 s-t) of horizontal viscosity is employed to suppress grid-scale noise. However, modelling practice sometimes favours smaller viscosity, and results are compared using 50 m 2 s -~ . The boundary conditions at the free surface z = tl(x,y,t) are

KM \--~Z "-~Z]

(fOx' roy)

{as or] Kg\oz, Oz) = 0 Similarly, at the bottom,

z

=

(8)

(9)

-H(x,y) \ OZ -~Z = (Tbx' "/'by)

(10)

(as, or) KH\az -~z = 0

(11)

w = - v al-l- v °H

(12,)

Ox

Oy

where r0 = (r0x, roy) is the wind stress vector and (rbx, Vby) is bottom drag. Wind stress is calculated from the wind speed following Gill (1982). The bottom drag is given by

(Tbx, Tby) = PoCDI Vbl(gb, Vb) where CD is the drag coefficient and Ub and Vu are the bottom velocities (for details, see Blumberg and Mellor, 1987; Mellor, 1993). The model uses a coordinate in the vertical (Mellor, 1993) with 24 a-levels having finer resolutions on the surface.

1594

B.K. Pal and G. Holloway

3O

25 20 '5 n "o

~5

(.9

10 5 0 0

10

20

30 Grid Points

40

50

60

Fig. 1. Bottom topography of the inner model domain (see text for explanation).

Initial T and S are obtained from the Levitus (1982) seasonal climatology data on relatively coarse resolution (1° × 1°) with most observations from the deep ocean. The model was initialized with horizontally averaged temperature and salinity fields at each depth, avoiding geostrophic adjustment to unbalanced initial condition. The model configuration is that of a periodic channel with no slip walls on the coastal and offshore boundaries and with cyclic boundary conditions on the cross-shelf (channel end) boundaries. Bottom topography is extracted from ETOPO5 (1986) on an inner domain (Fig. 1) with 65 grids alongshore and 33 grids cross-shore. The domain is extended to 85 grids alongshore to admit blending of topography to meet cyclic boundary conditions. Maximum depth is set at 2700 m. Cyclic conditions are used to isolate the model domain from the uncertain role of remote forcing from the broader Pacific. This is not to deny remote forcing in reality but to recognize the uncertainty of such forcing, while here addressing roles of local wind, Juan de Fuca discharge and topographic stress forcing. The Coriolis parameter is given a constant value corresponding to 49.5°N. Cyclic conditions were used similarly by Haidvogel et al. (1991) to study filament formation and evolution in the California coastal transition zone. 3. E D D Y - T O P O G R A P H Y INTERACTION Theoretical and modeling studies involving interaction between the eddies and the bottom topography identified a mechanism (topographic stress) which acts as a systematic force on the mean circulation (Haidvogel and Brink, 1986; Holloway, 1987; Treguier, 1989; Holloway and Muller, 1990). A possible way to introduce this effect into models that are not explicitly sufficiently eddy-resolving was suggested by Holloway (1992) in terms of the tendency of eddies to drive the ocean toward higher system entropy.

Dynamicsof circulation off VancouverIsland

1595

From statistical dynamics in ideal quasi-geostrophic system, Salmon et al. (1976) showed that an ocean with no external forcing and filled with random eddies will tend to set up a mean flow satisfying (a//3 - v 2) (~0) = h

(I3)

where ~p is the stream function, 72 is the two-dimensional laplacian, a/B(= L -z) is a ratio of Lagrangian multipliers, h = fOH/H is the depth integrated potential vorticity due to variation OH about the mean depth, and f is the Coriolis parameter. Ocean models that cannot adequately resolve eddies are systematically corrupted by the lack of tendency towards higher system entropy. The real ocean with external forcing and internal dissipation is not a closed system to which entropy solutions apply. Nevertheless, Holloway (1992) suggested a simplified expression for a transport stream function motivated by the equilibrium solution (13)

~.1

=

(14)

-fL2H

Parameter L, which comes from inviscid quasi-geostrophic theory, is not well understood in reality. Previous studies in mid-latitudes (Alvarez etal., 1994; Eby and Holloway, 1994; Holloway et al., 1995) have considered L ranging from 4 to about 8 km. For present purpose, L = 6 km is chosen. The form (14) is proposed to apply on scales larger than either length parameter L or the first internal Rossby radius. For the domain of interest over depth ranging from 200 to 1000 m with representative stratification, first radii range from 8 to 12 km. Although these scales are comparable with the grid spacing (hence strongly viscously damped), they are smaller than the scales of interest. As an exploratory application (see also Fyfe and Marinone, 1995) here in coastal zone setting, formulation (14) is continued with. In terms of velocity, the maximum entropy solution can be written as u*

10W -

H Oy

v*

-

10q t H Ox

(15)

following Eby and Holloway (1994). The horizontal viscosity parameterization in the momentum equation is modified to force the model solution towards the maximum entropy equilibrium velocities according to

A M F 2 ( U - u*)

A M F 2 ( V - v*)

(16)

where A M is the horizontal eddy viscosity. 4. M O D E L L I N G B U O Y A N T D I S C H A R G E FROM THE JUAN DE FUCA STRAIT Besides forcing from the Pacific Ocean (remote forcing), tides and local winds, the southern part of the Vancouver Island shelf is forced by coastal buoyancy flux. There are two principal sources of low density water (Hickey et al., 1991): outflow from rivers along the coast and the Juan de Fuca Strait. Discharge from the Juan de Fuca Strait is the major source of fresh water (Thomson et al., 1989). The strait circulation is an estuarine circulation that consists of a seaward flow of relatively fresh water in the upper layer and a landward flow of more saline oceanic water at depth. The peak discharge is maximum in

1596

B.K. Pal and G. HoUoway

(a)

15 . . . . . . . . . .

10

o =. . . . . . . . . .

777757777777

10

77:

2o

7

30

40

6o

50

~-c'61e = 0 . 2 8 ( m / s ) . Mox. V e l o c i t y = 0.280m/s

(b)

10

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

5 0 ~

0

. . . . . . . . .

I

10

. . . . . . . . .

I

20

. . . . . . . . .

I

30

. . . . . . . . .

~

. . . . . . . . .

40

[

. . . . . . . . .

50

I

i

i

i ~

60

S'~'ale = 0 . 0 6 ( m / s ) . Max. V e l o c i t y = O.060m/s

Fig. 2. Plots of summer horizontal velocity field near the (a) surface and (b) bottom and cross-shore vertical sections of (c) temperature and (d) salinity across from Tofino after 20 days of integration (wind forced). (Continued opposite)

early summer and minimum in winter. The summer time runoff is not steady but undergoes fortnightly to monthly "pulses" due to variations in the tidally induced mixing (Thomson et al., 1989). The net volume transport from the strait is unknown. Labrecque et al. (1994) examined the residual (non-tidal) flow in Juan de Fuca Strait and estimated a net volume transport of 0.05 x 106 m 3 s - 1 from the strait. However, they concluded that their estimate of net residual outflow is "statistically not different from zero". The influence from Juan de Fuca Strait is incorporated into the model in two ways. Near the mouth of Juan de Fuca (SE corner of the domain shown in Fig. 1), model temperature and salinity fields are restored toward observations through the water column to represent baroclinic forcing. Separately, a net volume source near the mouth of Juan de Fuca with assigned values 1.0 x 104 m 3 s -1 was included (see Section 5.3). For the summer (May-September), the observed temperature and salinity fields (Hickey et al., 1991: their Fig. 13) were approximated with simplified analytical functions: S = 31.3 - 1.75 tan-l{0.09(25 - z)}

1597

Dynamics of circulation off Vancouver Island (c)

o

.... --,2.8-I

I ~---~

' '__~'.:__l ' -'.--

'JL__L_L-L--y.~,,,,~__j__ ~ /

'~'

_1oo

-200

~-

16.8 -

300

-400

i

-500

, ~ , I ~ ~ J ~ ] , ~ ~ , I ~ ~ . 5

10

15

.

.

.

I , , , , I ,

20

25

30

GRID NO.

FROM

(d)

4.818

TO

13.247

BY

0.50

0 32.2-6

~ -100

32.76~"" ]

L

-200

vE 3Z Ld d3 -300

p f

-400

-500

~ 0

~ 5

10

15 GRID NO.

FROM

31.762

20

TO

25

35.000

30

BY

0.28

1598

B.K. Pal and G. Holloway

(a) 50

25

20

15

~0

o ;---~- ~ - -. ~ ~ ~ - , - , - , - , . . ] - ~ - , 0 10

-, - - - ~ - - , - - - 20

~-, ~-~~ ~ ~ - , - ~ _ : . 2 ~ ~ ~ ~ ~ ~ ~ Z 3 E Z ~ Z - 30 40 50

ZZZZ---~ 60

= 0.48 ( m / s ) Mox. Velocity= 0 . 4 8 4 m / s

S-c'ole

(b)

.

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.

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.

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.

~

~ "--2"Z-_-S~-_-_-

15

]0

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10

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.

...............................

20

50

40

. .

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.

; i ~ ; -, -, -, -,~ 50

60

S'-E~le = 0.15 ( m / s ) . Mox. Velocity= 0 . 1 4 9 m / s Fig.

3.

S a m e as in

Fig.

2,

but for winter season.

(Continued opposite)

T = 10.8 + 2.64 tan-l{0.05(20 - z)} For the winter (October-March), temperature and salinity fields from the hydrographic record of Crean and Ages (1968; Station 75) averaged over October-March were approximated. The simplified analytical fit was given by S = 3 1 . 8 - 1.21 tan 1 { 0 . 0 3 ( 6 0 - z)} T = 8.75 + 1.59 tan-l{0.005(20 - z)} 5. RESULTS 5.1. Wind forcing

First, only wind forcing is considered. As summarized by Thomson et al. (1989), summer wind blows mainly from the northwest parallel to the shoreline. In winter, the wind direction changes and blows from the southeast. To simulate the summer circulation, the model was initialized with the horizontally

1599

Dynamics of circulation off Vancouver Island

.... , , , ~ ,

(c)

.... , .... ,y,,,

-100

-200

E

I Eb- 3 0 0 E Ld

--5.9~

- 400

-500

0

5

10

15 GRID NO. FROM

20

4.866 TO

25

30

10.000 BY

0.50

(d)

-100

-200

S OLd E3

-300

-400

~

-500 0

, 5

,

10

15 GRID NO. FROM

t

I

20

31.790 TO

_L_J

I

~

,

, ~ _ L

25

.35.000 BY

,30

0.25

1600

B.K. Pal and G. Holloway

averaged summer temperature and salinity fields with flow initially at rest. A uniform wind of typical summer speed 4 m s -1 (M. Foreman, pers. comm., 1994) and direction parallel to the coastline from the northwest was applied. The viscosity AM -- 500 m 2 s -1 was used except where indicated otherwise. The model was first run for 20 days. Running for another 20 days showed slight increase in magnitude of the velocity vectors, but the spatial structure of the flow field remained unchanged. Figure 2 shows the model output of surface and (near) bottom velocity vectors, temperature and salinity fields for the wind-driven summer circulation after 20 days. The circulation is everywhere toward the south, without any poleward-flowing coastal current. The maximum current at the surface is about 29 cm s -a. An upslope bottom current (Fig. 2b), compensating offshore surface drift, is obtained. The wind produced upwelling along the coastal boundary, as shown by the upward slope of isotherms and isohalines on the cross-shore vertical section plots (Fig. 2c,d). For winter simulation, the model was initialized with winter temperature and salinity, while a uniform wind from the southeast of typical winter speed 6 m s -1 (M. Foreman, pers. comm., 1994) was applied. Figure 3 plots the surface and bottom velocity vectors, temperature and salinity fields for the winter simulation. With the wind direction reversed, the current flows toward north-northwest (Fig. 3a,b). The maximum surface current is 48 cm s -1. The wind causes downwelling at the coast (Fig. 3c,d) with downslope bottom flow (Fig. 3b). The wind-forced winter circulation is stronger than the summer circulation, in part because stronger winds are assumed for winter. Mean stratification also is different between winter and summer. Is there also an intrinsic difference such as Haidvogel and Brink (1986) or Holloway (1987) consider? To test this, summer conditions were retained, but only the wind direction was reversed (at 4 m s -1 summer wind speed). The model response showed a small asymmetry favouring poleward flow. The difference field (i.e. the sum of winter and summer fields) is shown in Fig. 4. 5.2. Wind and topographic stress Retaining wind forcing as previously, the inclusion of topographic stress yields a noticeable change in the flow pattern. The surface circulation (not shown) is dominated by wind forcing. Figure 5a shows the velocity vectors at 50 m depth after 20 days of integration for the summer season. A pronounced poleward-flowing coastal current, an equatorward current over the slope and a transition zone between the two are apparent. These features agree qualitatively with Thomson et al. (1989) and Foreman et al. (1991). Some eddying motions are also evident on the shelf as well as at the entrance of the Juan de Fuca Strait. Although the latter eddy is weak, it has the sense of rotation of the Juan de Fuca eddy. The coastal current is well developed from I = 40 (alongshore grid point) northward. Figure 5b shows the current at 50 m depth for the winter circulation. The currents are poleward over the entire shelf and slope. Compared to summer circulation, the winter circulation is relatively smooth, without eddies, as is consistent with observation (Thomson et al., 1989). 5.3. Wind, topographic stress and buoyancy forcing The effect of wind, topographic stress and buoyancy forcings acting together is investigated. First, only temperature and salinity restoring near the Juan de Fuca is

16(11

Dynamics of circulation off Vancouver Island

F

0

10

20

30

40

50

S-~le =

60

o.os ( m / s )

Max. Velocity=

,

0.053m/s

Fig. 4. Difference field in the barotropic velocity of two wind-driven runs: one with summer wind blowing from the northwest and the second with the wind direction reversed.

included. Later, volume discharge will be included. Figure 6a shows currents at 50 m depth for the summer simulation after 20 days of integration. Circulation at the mouth of the Juan de Fuca is greatly influenced by the inclusion of the buoyancy forcing. The Juan de Fuca eddy (Fig. 6a) is a regular summer feature (Freeland and Denman, 1982). Although the topography has been modified to satisfy the periodic boundary condition, the model reproduced the overall circulation including the Juan de Fuca eddy. For the winter simulation (Fig. 6b), the current (after 20 days of integration) is toward the pole at all depths. Strong current appears over the shelf and slope. The shelf circulation did not show formation of eddies either on the shelf or at the mouth of the Juan de Fuca. This is consistent with the observations (Thomson et al., 1989). Observed summer and winter mean currents (generated from the monthly averaged current meter records collected between 1975 and 1991) are also shown in Fig. 6a,b (thick arrows). It appears that the model reproduces the coastal circulation reasonably well, although the vectors near the centre of the domain (Fig. 6a) do not agree well. Figure 7a shows cross-shore vertical sections of alongshore velocity field across from Tofino for wind, topographic stress and buoyancy forcings together. The velocity crosssection shows northward flow (stippled region) over the shelf and shelf-break and a broad southward flow (unstippled region) seaward. The core of the poleward current lies below the surface over the shelf-break. These features agree with Fig. 17b of Thomson et al. (1989). The narrow southward flow observed on the innershelf is observed in drifter tracks (W. R. Crawford, pers. comm., 1995). The vertical section of the alongshore velocity at the same location but without topographic stress is also shown (Fig. 7b). The flow is always southward. Further questions are how the model responds to wind and buoyancy forcings alone (without topographic stress) and the effect of using smaller viscosity. Three cases were considered, all using summer wind: (1) restoring temperature and salinity fields along the Juan de Fuca entrance using the same viscosity as in the previous experiments (i.e. 500 m:" s-~), (2) same as in (1) but viscosity reduced to 50 m 2 s -~ and (3) the same as (2) but including also a net volume discharge. Figure 8a shows the velocity field at 50 m depth for

1602

B.K. Pal and G. Holloway

(a)

30

25i 20i 15 10

i~-,:

,: _ F -- ,--- -- F -- F -- F F F F -- -- -- ? '---- F F -- -- F - - - -

0

10

-- F__-- F F - - F F F

20

30

F F F F F F F F F F

40

FFF

FL,__--__.LF

50

~

60

~-Eole = 0.28 ( m / s ) . Mox. Velocity= 0 . 2 7 8 m / s

(b)

,

30 25 ZZZZ

FFFF

....

ii?

2 0 ." Z. .. . ... . . . 15_~ZZZZZZ10 . . . . . . . 5

--

0

10

20

30

40

50

60

~3-c'ole = 0.55 (m./s)10'55" Max. V e l o c i t y = m

Fig. 5. Plots of horizontal velocity field with wind and topographic stress at days of integration: (a) summer and (b) winter.

50 m

/s

depth after

20

the first case. The model produces the Juan de Fuca eddy off the entrance of the strait, but it did not produce continuing poleward flow. The second and third cases produced very nearly the same result, and only the third case is shown in Fig. 8b. It was noted earlier that there is no reliable estimate of the rate of volume discharge from the Juan de Fuca Strait. An upper estimate can be taken from the Fraser River outflow, which is the major source of freshwater in the area. Frazer River discharges at the rate of 1.0 x 10 4 m 3 s -1 (Hickey et al., 1991) to the Strait of Georgia/Juan de Fuca system. A part follows the Juan de Fuca Strait to the ocean, while the other part continues northward in the Strait of Georgia. The relative portions are not well known. For the purpose of testing the hypothesis of buoyant driving of the poleward coastal flow, 1.0 x 104 m3/s - t was used as the rate of volume discharge from the Juan de Fuca Strait. Figure 8b shows the velocity field at 50 m depth for the third case (summer wind, temperature and salinity restoring, net volume discharge, and reduced viscosity). Again, the model exhibits a Juan de Fuca eddy but no continuing poleward current. The difference between Fig. 8a and b is due to reduced viscosity, as confirmed by case 2 (not shown), for which there was no net volume discharge. It is found

1603

Dynamics of circulation off Vancouver Island

(a) .

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•

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_'_="z

...........

20

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.....

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' '.'vT~..'=_Z,' .........

7~_-." ".-_-Z_-57--ZZS -:-"i " . ~ :- :- : : : : : " 2. L2------~--__22222-- 27- 2 Z

~5

10 i . . . . . . . . . . . . . .

5E o~=r~,r,=~?~?r????77,-T:Trrrrr:?77777777777777,-777777,-7777,--:,-7-5~. 3

10

20

40

30

50 g-~'61e =

6C

0.26 ( m / s ) 0.265rn/s

Mox. Velocity=

(b)

2o!

, .

.

.

.

. .

, .

~' .

7.

,

:0

.

----_--

----x,J

.

.

-

.

J

iI ii!iiiiiiiiii iii!i i> 0

10

20

30

40

50 ~'61e =

0.57 ( m / s )

60 .

Mox. Velocity= 0 . 5 6 5 m / s Fig. 6. ( a , b ) Same as in Fig. 5, but for wind, topographic stress and buoyancy torclngs. I r o c k arrows on the velocity plots (enhanced b y a factor of 2) represent observations. Current meter positions are marked by circles.

that buoyancy forcing at Juan de Fuca and reduced viscosity can generate the Juan de Fuca eddy but not a continuing poleward coastal current against the summer wind. 6. SUMMARY AND DISCUSSION Dynamical balances in the circulation off the west coast of Vancouver Island were examined using the Princeton Ocean Model. The forcings considered were wind stress, topographic stress and buoyancy flux from the Juan de Fuca. Forcings were applied separately and in combinations. For summer simulation, application of wind-forcing produced flow everywhere toward the equator. With topographic stress and summer wind, the model produced a polewardflowing coastal current over the shelf with a transition zone to equatorward flow over the slope. With the further introduction of the buoyancy forcing, a cyclonic eddy developed off the mouth of the Juan de Fuca Strait. These features agree with observations. Buoyancy forcing and summer wind (without topographic stress) generated a Juan de

1604

B.K. Pal and G. Holloway (a)

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Dynamics of circulation off Vancouver Island

(a)

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50 S'~'ale = 0 . 2 6 ( m / s ) Max. V e l o c i t y = 0.265m/s

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Fig. 8. Plots of horizontal velocity field (a) s u m m e r wind and buoyancy forcing and (b) volume flux, s u m m e r wind and buoyancy forcing with viscosity = 50 m 2 s 1.

Fuca eddy but did not produce poleward flow. Adding a volume flux of 1.0 x 104 m 3 sfrom the Juan de Fuca and reducing viscosity from 500 to 50 m 2 s -1 enhanced the Juan de Fuca eddy but did not produce a continuing poleward coastal current. For winter simulation, flow was poleward everywhere. Inclusion of the topographic" stress made little difference in the spatial structure of the flow except to increase the maximum current. Buoyancy flux from the Juan de Fuca Strait did not create an eddy al: the Juan de Fuca entrance or on the shelf. During the winter season, forcings act in the same direction, unlike the summer when strong velocity shear contributes to the formation of eddies. This summer/winter difference is consistent with observations. These experiments suggest a significant role of topographic stress in the summer

Fig. 7, Cross-shore vertical sections of s u m m e r alongshore velocity across from Tofino after 20 days of integration for (a) wind, topographic stress and buoyancy forcings and (b) wind and buoyancy forcings. Stippled areas mark poleward flow. Speeds are in m s-~.

1606

B.K. Pal and G. HoUoway

circulation off Vancouver Island, when other forcings (buoyancy, tidal) do not account for a poleward coastal flow. Some restrictive assumptions made in this investigation should be relaxed in future study. In particular, assumption of a periodic channel is a limitation to the complex topography of the region and precludes remote forcing. Acknowledgements--The authors thank Mike Foreman, Rick Thomson, Howard Freeland and Michael Eby formany useful discussions. Thanks are also due to Lie-Yauw Oey for helping with the model, Tessa Sou, Patricia Kimber and Bon van Hardenberg for helping with graphics and Robin Brown for supplying current meter data. An anonymous reviewer is also thanked for many useful comments. One of the authors (B. K. P.) acknowledges support from the Office of the Naval Research grant N00014-92-J-1775.

REFERENCES Alvarez A., J. Tintore, G. Holloway, M. Eby and J. M. Beckers (1994) Effect of topographic stress on circulation in the western Mediterranean. Journal of Geophysical Research, 99, 16053-16064. Blumberg A. F. and G. L. Mellor (1987) A description of three-dimensional coastal ocean circulation model. In: Three-dimensional coastal ocean models, Volume 4, N. Heaps, editor, American Geophysical Union, Washington D.C., pp. 1-16. Clarke A. J. (1989) Theoretical understanding of eastern ocean boundary undercurrents. In: Poleward flows along eastern boundaries, S. J. Nashyba et al., editors, Springer-Verlag, New York, pp. 26-39. Crean P. B. and A. B. Ages (1968) Oceanographic records from twelve cruises in the Strait of Georgia and Juan de Fuca Strait, Department of Energy, Mines and Resources, Marine Science Branch, Victoria. Denman K. L., H. J. Mackas, H. J. Freeland, M. J. Austin and S. H. Hill (1981) Persistent upwelling and mesoscale zones of high productivity off the west coast of Vancouver Island, Canada. In: Coastal upwelling, F. Richards, editor, American Geophysical Union, Washington D.C., pp. 514-521. Eby M. and G. Holloway (1994) Sensitivity of a large scale ocean model to a parameterization of topographic stress. Journal of Physical Oceanography, 24, 2577-2587. ETOPO5 (1986) Global 5' × 5' depth and elevation, National Geophysical Data Centre, NOAA, US Department of Commerce, Code E/GC3, Boulder, CO 80303. Flather R. (1987) A numerical model investigation of tides and diurnal period continental shelf waves along Vancouver Island. Journal of Physical Oceanography, 18, 115-139. Foreman M. G. G. and R. A. Walters (1990) A finite element tidal model for the southwest coast of Vancouver Island. Atmosphere-Ocean, 28,261-287. Foreman M. G. G., R. E. Thomson, D. R. Lynch and R. A. Waiters (1991) A finite element model for threedimensional flows along the west coast of Vancouver Island. In: Estuarine and coastal modeling, 2nd International conference/WW Division ASCE, Tampa, Florida, ASCE, New York, pp. 13-15. Foreman M. G. G., A. M. B aptista and R. A. Waiters (1992) Tidal model studies of particle trajectories around a shallow coastal bank. Atmosphere-Ocean, 30, 43-69. Freeland H. J. and K. L. Denman (1982) A topographically controlled upwelling center off southern Vancouver Island. Journal of Marine Research, 40, 1069-1093. Fyfe J. and G. Marinone (1995) On the role of unresolved eddies in a model of the residual currents in the central Strait of Georgia, B.C. Atmosphere-Ocean, 33, 6134519. Gill A. E. (1982) Atmosphere-ocean dynamics, Academic Press, New York, p. 662. Haidvogel D. B. and K. H. Brink (1986) Mean currents driven by topographic drag over the continental shelf and slope. Journal of Physical Oceanography, 16, 2159-2171. Haidvogel D. B., A. Beckman and K. S. Hedstrom (1991) Dynamical simulations of filament formation and evolution in the coastal transition zone. Journal of Geophysical Research, 96, 15017-15040. Hickey B. M. (1979) The California Current system--hypotheses and facts. Progress in Oceanography, 8,191279. Hickey B. M., R. E. Thomson, H. Yih and P. H. LeBlond (1991) Velocity and temperature fluctuations in a buoyancy-driven current off Vancouver Island. Journal of Geophysical Research, 96, 10507-10538. Holloway G. (1987) Systematic forcing of large-scale geophysical flows by eddy-topography interaction. Journal of Fluid Mechanics, 184, 463-476.

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Holloway G. (1992) Representing topographic stress for large scale ocean models. Journal of Physical Oceanography, 22, 1033-1046. Holloway G. and P. Muller (1990) Topographic stress in the oceans. Eos, 71,343-344. Holloway G., K. H. Brink and D. Haidvogel (1989) Topographic stress in coastal circulation dynamics. In: Polewardflows along eastern boundaries, S. J. Neshyba etal., editors, Springer-Verlag, New York, pp. 142159. Holloway G., T. Sou and M. Eby (1995) Dynamics of circulation of the Japan Sea. Journal of Marine Research, 33, 539-569. Labrecque A. J. M., R. E. Thomson, M. W. Staeey and J. R. Buckley (1994) Residual currents in Juan de Fuca strait. Atmosphere-Ocean, 32, 375-394. Levitus S. (1982) Climatological atlas of the world ocean, NOAA Prof. Paper 13, U.S. Government Printing Office, Washington, DC, p. 173. Mackas D. L., K. L. Denman and A. F. Bennett (1987) Least squares multiple tracer analysis of water mztss composition. Journal of Geophysical Research, 92, 2907-2918. McCreary J. P. (1981) A linear stratified model of the coastal undercurrent. Philosophical Transactions of Royal Society London Series A, 302,385-413. McCreary J. P. and S. Y. Chao (1985) Three-dimensional shelf circulation along an eastern ocean boundary. Journal of Marine Research, 43, 13-36. Mellor G. L. (1993) User's Guide for a three-dimensional, primitive equation, numericalocean model. Program in Atmospheric and Oceanographic Sciences, Princeton University, N.J., p. 35. Mellor G. L. and T. Yamada (1982) Development of a turbulent closure model. Reviews of Geophysics and Space Physics, 20, 851-879. Salmon R., G. Holloway and M. G. Hendershott (1976) The equilibrium statistical mechanics of simple quasigeostrophic models. Journal of Fluid Mechanics, 75,691-703. Thomson R. E., B. M. Hickey and P. H. LeBlond (1989) The Vancouver Island coastal current: fisheries barrier and conduit. In: Effects of ocean variability on recruitment and an evaluation of parameters used in stock assesment models, R. J. Beamish and G. A. McFarlane, editors, Canadian Special Publication of Fisheries and Aquatic Sciences 108, Department of Fisheries and Oceans, Ottawa, pp. 265-296. Treguier A. M. (1989) Topographically generated steady currents in barotropic turbulence. Geophysical and Astrophysical Fluid Dynamics, 47, 43--68.