Local “Mean Field” Theory of Hydraulically Controlled Strait Flow

Melvin E. Stern Oceanography Department, The Florida State University, Tallahassee, Florida

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Abstract

The free discharge of a layer of bottom water in a wide strait (e.g., the Denmark Strait) differs from the classical control problem because of the strong geostrophic turbulence. As a consequence, the cross-stream (x) variation of the time-averaged downstream velocity υ(x) is “underdetermined” and depends on more parameters than available conditions. To resolve this “degeneracy” the classical control condition is generalized; the result requires the discharge Q to be extremized with respect to the degeneracy parameters and with respect to the constraints. One of these constraints is that a branch point or a local stationary wave can be supported at some section of the long channel. By either maximizing Q or by requiring a stationary wave, useful approximations are obtained. Future work should consider the joint variational problem. Specific calculations are made for nonuniform potential vorticity in a rectangular channel and also for variable cross-stream bottom topography. In the latter case, the current has a free streamline whose point of intersection with the bottom is computed. It is suggested that a necessary condition for hydraulic control at any section is that the flow does not separate from either side of the channel. The mean isopycnals in the Denmark Strait satisfy this condition, thereby suggesting that it does in fact (topographically) control the upstream state in the Greenland Sea.

Corresponding author address: Melvin E. Stern, Oceanography Department, The Florida State University, Oceanography/Statistics Bldg., West Call St., Tallahassee, FL 32306-4320. Email: stern@ocean.fsu.edu

Abstract

The free discharge of a layer of bottom water in a wide strait (e.g., the Denmark Strait) differs from the classical control problem because of the strong geostrophic turbulence. As a consequence, the cross-stream (x) variation of the time-averaged downstream velocity υ(x) is “underdetermined” and depends on more parameters than available conditions. To resolve this “degeneracy” the classical control condition is generalized; the result requires the discharge Q to be extremized with respect to the degeneracy parameters and with respect to the constraints. One of these constraints is that a branch point or a local stationary wave can be supported at some section of the long channel. By either maximizing Q or by requiring a stationary wave, useful approximations are obtained. Future work should consider the joint variational problem. Specific calculations are made for nonuniform potential vorticity in a rectangular channel and also for variable cross-stream bottom topography. In the latter case, the current has a free streamline whose point of intersection with the bottom is computed. It is suggested that a necessary condition for hydraulic control at any section is that the flow does not separate from either side of the channel. The mean isopycnals in the Denmark Strait satisfy this condition, thereby suggesting that it does in fact (topographically) control the upstream state in the Greenland Sea.

Corresponding author address: Melvin E. Stern, Oceanography Department, The Florida State University, Oceanography/Statistics Bldg., West Call St., Tallahassee, FL 32306-4320. Email: stern@ocean.fsu.edu

1. Introduction

In the classical hydraulics for the steady inviscid discharge of a single layer of dense fluid from a reservoir into a long channel of transverse width L(y) and bottom elevation M, a steady quasi-laminar downstream (y) velocity υ(x) and cross-stream thickness h(x) are assumed, so that the Bernoulli function on each streamline is equal to its reservoir value; furthermore, the potential vorticity (PV) is assumed to be uniform. Besides the volume discharge Q, there is (at least) one more unknown [say h(0)] than equation, so that some form of the Hugoniot control condition is required to close the problem and to determine Q as a function of the reservoir head (relative to the sill). In one version of the Hugoniot condition, a topographic change (say δL) is made and δh(0)/ δL is computed. If it is finite, the flow is not controlled at this section, because δL merely produces a local perturbation without affecting Q. Otherwise, if a bifurcation point (δh/δL = ∞) exists, the flow Q is controlled1 (although not necessarily at the section in question). Because δh/δL = ∞ implies δh ≠ 0 for δL ≡ 0, the branch-point condition is equivalent to the stationary-wave condition in a straight channel. These alternate conditions are also equivalent to a Q extremal; that is, ∂Q/∂h(0) = 0 (Killworth 1995). It is important to note, however, that some free discharges (see the appendix) are “controlled” in the ordinary sense of the word, even though the stationary-wave condition is irrelevant. Such simple examples point out the need for a wider conception of what constitutes control.

The stationary-wave condition has also been carried over to highly unsteady exchange flows such as exist in the Bosphorus [see Gregg and Özsoy (2002) for comprehensive field observations and numerical calculations pertaining thereto]. This exchange problem is beyond our scope, however, and only one-way flow (nonreversing) over a simple sill or through a narrows is considered.

The hydraulic approach has been extended to the free discharge over a rotating sill (Whitehead et al. 1974; Stern 1974, 1975; Gill 1977; Shen 1981; Borenas and Lundberg 1986; Helfrich and Pratt 2003) with application to the important problem of the overflow of the cold bottom water, such as occurs in an oceanic strait (Fig. 1). In these theoretical studies (typically in a rectangular channel), there is a dynamically active dense layer in the upstream reservoir, and the overlying slightly lighter layer is resting; thus a reduced-gravity (g′) model applies. Upon initiating the discharge, the dense fluid flows over the sill and establishes a steady flow with a uniform PV equal to that assumed in the reservoir. As previously mentioned, the problem is closed by applying the Hugoniot condition. With certain other conditions (Gill 1977) this gives Q as a function of Coriolis parameter f, reduced gravity g′, reservoir head H∗, and the channel dimensions. Mention should be made of the Killworth (1994, 1995) papers, which show that for general PV and M(x) the value of Q has an upper bound; this bound is realized in the PV = 0 calculation of Whitehead et al. (1974). Killworth showed that for a very general class of PV the Hugoniot condition is equivalent to extremizing Q. Killworth (1992) also considered the nonuniform PV problem for a very special Bernoulli function in the upstream reservoir.

For consistency in these long-wave theories, υ(x) ≥ 0 is assumed, and then the negative definite relative vorticity requires the minimum channel width to be less than a certain multiple of the radius of deformation (gH∗)1/2f−1. In the case of wide ocean straits (Fig. 1) there is another limitation because each (synoptic) realization contains large-scale eddies (Spall and Price 1998) and dynamically significant cumulative bottom friction, which modify the fluid in its long passage from the reservoir to the sill. Thus, the time-average Denmark Strait flow (Fig. 1) consists of a jet whose Bernoulli function and PV are not known a priori (Spall and Price 1998). Because of the turbulence, the mean flow poses a statistical mechanical problem in which the mean channel velocity υ(x) is underdetermined. As a consequence, it is necessary to articulate a control principle (section 2a) so that it may be applied to the mean of the turbulent flow. The spirit of this investigation is similar to that in turbulent thermal convection (cf. Howard 1972) and turbulent shear flow (Stern 1980); the goal is to approximate certain integral features of a real and time-dependent channel flow.

In section 2a the generalized control principle is stated, and the implication is that the channel discharge (Q) is to be extremized subject to certain constraints; most notable is the requirement that the mean υ(x) contains a branch point (section 2b). Note that the foregoing two conditions are equivalent in classical hydraulics. When the first (Q) of these two conditions is taken alone, the result is illustrated in section 2c. When the second condition is taken alone, the result is illustrated in sections 3 and 4. When the two conditions are taken together, a problem in the calculus of variation (section 3) is obtained (but the latter is not solved here).

The most interesting, and oceanically relevant, problem should take into account a variable cross-stream bottom topography that is intersected by the interface of the density layer. The simplest version of this free-streamline problem (see Fig. 6 below) is considered in section 4 in which the stationary-wave condition is obtained, but the incorporation of the Q extremization is left for subsequent work. The results obtained are used in section 5 to argue that the observed mean flow in the Denmark Strait is indeed controlled by the topography.

Sections 3 and 4 may be read independently as specific examples of the effect of nonuniform PV [due to variations in h(x), υ(x), or M(x)] on the stationary-wave condition.

2. The control principle

a. Generalized Hugoniot condition

A definition can be given, as follows: if all small local topographic changes (±δL, ±δM) at a given section in a strait merely result in a local υ(x) perturbation without changing the reservoir state (e.g., the discharge Q) then the upstream state is not controlled at this section; otherwise the flow is controlled (either at this section, or some other one).

This definition has operational content; in principle one could make topographic changes and see whether merely local flow changes are produced. The theoretical content of the definition emerges with the assumption that the straits in question do control Q and with the assertion that the underdetermined mean velocity υ(x, λ1, … , λn), which contains more free parameters λ1, … , λn (e.g., Fourier coefficients) than known dynamical constraints, should satisfy the control condition. By a cross-stream integration of υh one relates the reservoir discharge,
QQλ1λnM,L
to the straits. In classical hydraulics for a single layer flow, there is one more unknown (λ) than dynamical equations, and its value is obtained by the Hugoniot branch-point condition; in the turbulence problem, however, there are many more λ. We now assume that all the underdetermined λ solutions have finite a priori probability and that the most probable λ solution satisfies the control condition. The meaning of this is as follows.
At a given topography (M, L) consider a possible λ solution with the specific values (λ01, λ02, … , λ0n)and then examine the effect of changing the local topography by (δM, δL). All concomitant changes δλ1, δλ2, … , δλn are also a priori possible. If the various derivatives ∂Q/∂λ exist (on both sides of λ01, … , λ0n) then the change in transport is
i1520-0485-34-7-1692-e22
Now if ∂Q/∂λi ≠ 0 for any λ0i, say λ01, then there certainly exists a solution with δλ2, … , δλn = 0 and with
i1520-0485-34-7-1692-e23
for which
δQδM,δL
Because this violates our control hypothesis, the solution with λ01, λ02, … , λ0n is not realized. On the other hand, our solution may satisfy the control condition if
Qλii
or if λi is an endpoint at which the derivative does not exist. Both of these contingencies are illustrated in Fig. 2 (see also the Hele-Shaw weir in the appendix for an illustration of an endpoint “extremum”). Thus, the control condition serves to select a discrete number of realizations from the continuum presented by the turbulence problem. Perhaps our theory is best stated by saying that because the solution (2.4) satisfies more requirements it is more probable than the other λi.

b. The branch-point constraint

In addition to the Q change, a second requirement of a topographic change (δL) is necessary for control. If δL merely forces a local velocity perturbation (δυ), then the section in question is not a control; otherwise—that is; if δυ/δL does not exist—a control may exist. The implementation in the oceanic context is as follows. Compute the forced mean perturbation while assuming that the Bernoulli condition is satisfied locally (but not all the way into the reservoir). If δυ/δL is finite then there is no control; otherwise, if there is a bifurcation point (δυ/δL = ∞) then the necessary condition for a control is satisfied. Thus, the most probable solution at a presumed control section is one in which Q is to be extremized with respect to the underdetermined parameters and subject to the stationary-wave constraint. Of course if there are other known constraints on υ, they should be included. The predictions are admittedly incomplete because only interrelationships of parameters describing the real flow can be obtained.

The importance of the stationary-wave condition for nonuniform PV will be indicated in sections 3 and 4, but first we give an example in which Q is extremized without using the stationary-wave constraint. This approach can naturally only give a variational approximation to the full problem.

c. Illustrative example of δQ = 0

Suppose we have a (globally) turbulent flow passing through a channel with triangular bottom topography M(x); the two sides in Fig. 3 slope at angles α0 and α1, respectively. The current has “left-hand” support (h = 0) on α0 > 0. The layer interface h(x) intersects the sides at elevations H0 and H1, respectively, and L0 and L1 are the respective distances of these intersections from the vertex of the triangle. Let us maximize Q by (temporarily) ignoring the stationary-wave constraint and by constraining υ(x) to be (uniformly) equal to the x average of the real geostrophic mean flow. Then
i1520-0485-34-7-1692-e25
If s = tanα1/tanα0 and λ = H0/H1, then a straightforward calculation gives
i1520-0485-34-7-1692-e26
where λ > 0 is a free parameter. For fixed H1, Q is maximized at
i1520-0485-34-7-1692-e27
In particular, for s = 1 or α0 = α1, we obtain
i1520-0485-34-7-1692-e28
Equation (2.9) predicts an approximate value of the mean channel velocity in a controlled flow. In a subsequent variational approximation a uniform shear might be added to υ, and the stationary wave constraint might supplement the δQ = 0 condition.

3. Influence of variable potential vorticity on stationary waves in a rectangular channel

The potentialities of our formalism are best illustrated in the case of a rectangular strait. Let υ(x) be the steady mean downstream (y) velocity in a wide flat-bottomed channel with vertical sidewalls (Fig. 4), the left-hand wall (looking downstream) being straight. If the right-hand vertical wall is slightly curved in the downstream direction, it will force a slowly varying perturbation of the undisturbed upstream current. (This wall curvature may eventually be reduced to zero to obtain the condition for the existence of a stationary wave in a completely uniform channel.) At the upstream end of the channel, the vertical walls are located at x = 0, L, and the assumed semigeostrophic mean velocity is
i1520-0485-34-7-1692-e31
where h is the time-averaged layer thickness. At the downstream end, the vertical walls are located at ξ = 0, L1, and
i1520-0485-34-7-1692-e32
is the mean geostrophic velocity. Consider a streamline at a distance x from the left-hand wall at the upstream end, and at a distance ξ from the wall at the downstream end. The geostrophic volume flux between this streamline and this straight wall is given by
i1520-0485-34-7-1692-eq2
and the equivalent flux at the upstream end is
i1520-0485-34-7-1692-eq3
It follows that the downstream layer thickness is
h1ξh2xa1/2
where
ah21h2
As previously mentioned (section 2b) we assume a local Bernoulli equation υ21(ξ) + 2gh1(ξ) = υ2(x) + 2gh(x) is satisfied, in which case
i1520-0485-34-7-1692-eq4
When this and (3.3a,b) are substituted in (3.2) we obtain the first-order differential equation
i1520-0485-34-7-1692-e34
for the displacement ξ(x) as a function of υ(x) and h(x). The constant of integration a is obtained by integrating (3.4) across the entire channel. Because ξ goes from 0 to L1 as h(x) goes from h(0) to h(L), (3.4) gives
i1520-0485-34-7-1692-e35
for a [= h1(0) − h(0)] as a function of L1L.
From (3.3b), we note that when a = 0 and h1(0) = h(0) the integrand in (3.5) reduces to dh/υ = fdx/g′. Then the right-hand side of (3.5) reduces to fL/g′, and L1 = L is recovered when a = 0. We now compute the value of a for small values of (L1L) by differentiating (3.5) with respect to a and then setting this equal to zero:
i1520-0485-34-7-1692-eq5
Because the derivative of (3.3b) is da = 2h1(0)dh1(0), we obtain
i1520-0485-34-7-1692-e36
This is positive and h1(0) increases with L1L if υ2(x) < gh(x), but if the integral vanishes, that is, if
i1520-0485-34-7-1692-e37
then a branch point [∂h1(0)/∂L1 = ∞] occurs in the forced solution, and no undisturbed upstream solution will exist if the channel width decreases downstream. Thus, (3.7) is the critical Froude number condition for a controlled flow [as obtained previously by Stern (1975) and Pratt and Armi (1987)]. Equation (3.7) can also be interpreted as the condition for a stationary wave [amplitude dh1(0) ≠ 0] in a channel in which both walls are straight (L1L = 0).
If H0 is the minimum layer thickness (at x = 0) and H is the maximum thickness (at x = L), then
i1520-0485-34-7-1692-e38
Equation (3.7) consequently reduces to
i1520-0485-34-7-1692-e311a
According to the two parts of our control condition we extremize Q = (g′/f)(H2H20) subject to (3.11), and the method of Lagrange multipliers implies that this is equivalent to extremizing
i1520-0485-34-7-1692-e311b
with fixed Q, or fixed endpoints:
i1520-0485-34-7-1692-e311c
Equation (3.11b) implies a second-order Euler–Lagrange equation for the structure of ĥ(η), and (3.11c) or ĥ2(L) − ĥ2(0) = 2Q(g′/f)−1 supplies the boundary conditions. Equation (3.11b) serves to indicate the increased amount of information about a globally turbulent flow that is provided by our variational theory (section 2); without it, and with only the assumed Hugoniot condition (3.11a), we have only a single integral for ĥ(η). More important, our formulation of the general problem provides the basis for a variational approximation using suitable form functions for υ(x), as is illustrated below.

a. Constant υ(x)

Consider first the case in which υ is approximated by a uniform current (equal to the x average of the real υ), in which case ĥ varies linearly across the stream so that
i1520-0485-34-7-1692-e312ab
The transport
i1520-0485-34-7-1692-e312d
depends on the maximum layer thickness H and on β. Then (3.11) gives two elementary integrals:
i1520-0485-34-7-1692-eq6
whose evaluation yields
i1520-0485-34-7-1692-e313

The main result [from (3.13)] is that an arbitrarily wide (L) strait can provide a branch-point condition provided that β is large and positive; that is; the current is supported on the left-hand side wall, with H0 = H(1 + β)−1 > 0. We note in passing that the Froude number υ2/gH0 based on the minimum thickness (1 + β)β−1 ln(β + 1) > 1 is “supercritical” whereas the right-hand value based on the maximum thickness Froude number β−1 ln(β + 1) < 1 is “subcritical.” Note that the mean isopycnals in Fig. 1 do intersect the bottom on both sides of the strait, suggesting that the stationary-wave criterion may be relevant here. This conclusion is not foregone, however, because the flow might have separated from the left-hand side, as occurs in the numerical calculation of Spall and Price (1998) in which the overflow banks on the right-hand side of the channel. To further explore this qualitative point (of left-hand support) we will introduce a sloping bottom in section 4. First we want to investigate the effect of relative vorticity in the rectangular channel, however.

b. Variable υ

It is of interest to see how the stationary-wave condition is modified for a jetlike υ(η) with maximum υ at η = 1/2. In (3.11a) let
i1520-0485-34-7-1692-e314a
where λ is a free amplitude parameter. Equation (3.11a) is then written as
βIβ,λI1I2
where
i1520-0485-34-7-1692-e315b
Figure 5 gives the numerical calculation of Ff2L2/gH for any β and for λ ≤ 0. Figure 5 shows that, for any given Q, H, and L, a stationary wave can exist for a jet with prescribed amplitude; for example, if F = 2 and β = 6 then λ ≃ −0.04. Although the simple jet parameterization [in (3.14b)] is insufficient to give a Q (or β) extremum, there is a (well known) upper bound Q = (g′/ 2f)H2 at any F. This bound occurs at β → ∞ and λ → −1/(2π), [i.e., υ(0) = υ(1) = 0], because lines of constant F intersect all λ = constant curves in Fig. 5.

4. Sloping bottom

More interesting (and relevant) than the rectangular channel in Fig. 4 is one with a sloping bottom that is intersected by the interface of the density current, thereby providing a free-streamline condition for the stationary-wave problem; the problem is to predict the width of the current in a free discharge. In the simplest case (Fig. 6), a uniformly sloping (α) bottom M(x) intersects a vertical right-hand boundary, and the velocity is approximated by a uniform geostrophic velocity U equal to the η average of the realized velocity υ. At the (dimensional) transverse coordinate η = L the layer thickness is H(η), with H(0) = 0 and H(L) = H1. This problem differs from the classical trapped coastal wave (LeBlond and Mysak 1978) for which there is no vertical wall and no sloping density interface. Our problem also differs from the Killworth (1992) “global” solution which is based on a small uniform current in a wide upstream reservoir.

The plan view (Fig. 6) of the steady long wave shows a streamline whose distance from the η = 0 origin is η at the upstream section and y(η) at the downstream section. Although such a displacement might be forced by a transverse displacement of the vertical wall to y = LδL, we shall henceforth consider the unforced free wave problem (δL ≡ 0) on a uniform slope α. Therefore the bottom elevation at the downstream streamline is
MyMηαyη
In this Lagrangian formulation, the corresponding geostrophic velocity and layer thickness are respectively given by u(y), and h(y). The mass conservation equation uhdy = UHdη may be written as
ϕudyUdη,
where
ϕhyHη
A second upstream/downstream equation is supplied by the Bernoulli relation
i1520-0485-34-7-1692-e43
The third relation is provided by the geostrophic equations, whose upstream value is
i1520-0485-34-7-1692-e44
and whose downstream value u = (g′/f)dh/dy − (g;ga)/f is expressed as
i1520-0485-34-7-1692-e46
where dH/dy = /dy(dH/) is used.
For infinitesimal perturbations (primed quantities) from the upstream state, the downstream state is given by
i1520-0485-34-7-1692-e47a
The linearization of (4.6), (4.2a), and (4.3) gives a complete set of linear equations for ϕ′, u′, and y′:
i1520-0485-34-7-1692-e48a
with nonconstant coefficients [in H(η)].
In the full nonlinear problem the free boundary condition is h[y(0)] = 0, and in the perturbation problem (4.2b) requires that ϕ′(0) be finite. Because y′(0) must also be finite, we have the free boundary condition
yϕη
and the other boundary condition for the stationary wave is
yL
By subtracting (4.8b) from (4.8a) and then from (4.8c), there results
i1520-0485-34-7-1692-e410
Now the foregoing equations are nondimensionalized using
i1520-0485-34-7-1692-e412a
where M(0) ≡ εH1 denotes the topographic elevation on the left-hand side of the upstream channel. The various coefficients in (4.10)–(4.11) then become
i1520-0485-34-7-1692-e413a
where
i1520-0485-34-7-1692-e413c
Thus the nondimensional differential equations become
i1520-0485-34-7-1692-e414a
By substituting the power series
i1520-0485-34-7-1692-e414c
and collecting the coefficients of z0, we obtain the indicial equations
i1520-0485-34-7-1692-e415
For n > 0, the coefficients of zn satisfy
i1520-0485-34-7-1692-eq7
Thus, the recursion relation is
i1520-0485-34-7-1692-e416
When the amplitude normalization condition b0 = 1 is used, (4.15) gives
i1520-0485-34-7-1692-eq8
and (4.16) then gives the remaining bn. Because Sn < 0 for all n, bn is an alternating series, and (4.14c) converges at all 0 ≤ z ≤ 1. The regular solutions (4.14c) satisfy the boundary condition (4.9a), and (4.9b) will be satisfied if
i1520-0485-34-7-1692-e417
This gives the value of εF as a function of ε. For the interesting case of ε ≪ 1, and εFr = O(1), we get
i1520-0485-34-7-1692-eq9
Equation (4.17) then reduces to
i1520-0485-34-7-1692-e418
where Gm is the sum of the first m terms and m → ∞. For G4 = 0, we get the approximate root r ≅ 1.45, or
i1520-0485-34-7-1692-e419
Because the discharge is Q = ULH1/2 and U = (g′/ f)(H1αL)L−1, we obtain
i1520-0485-34-7-1692-e421
The interesting result (4.20) predicts the width of the current at the control as a function of H1; if α is sufficiently small, then L may be much larger than the radius of deformation. Also note from (4.13b) that gαL/ U2 = r, or
i1520-0485-34-7-1692-eq10

For g′ = 0.1 cm s−2, f = 10−4 s−1, H1 = 5 × 104 cm, and α = (10−2)/4, we have L ≃ 2 × 106 cm and Q = gH21(2f)−1(1 − 0.181). Of course, the cumulative effect of bottom friction will produce a jet with non uniform U(η); perhaps this effect may be taken into account in (4.8). With the increased parameterization of U(η) there may also be a Q extremum that predicts those parameters.

Also interesting in the context of the more general topographic problem is the case of α < 0 (Fig. 6b), which corresponds to a current that separates from the left-hand side of a channel (not shown) and banks on the right-hand side. In this case, ε < 0 and (4.18) cannot be satisfied; this condition means that the flow cannot be controlled at a section where “separation” occurs.

5. Summary and suggestions

The hope of all turbulence theories is to bypass the chaotic realization and predict the mean state; the Denmark Strait overflow provides a good example because of the strong geostrophic eddies observed therein. This paper addresses the qualitative question of whether such a wide current is actually “topographically controlled,” or whether the flow in the straits is merely locally forced without influencing the mean state of the Greenland Sea.

First of all it was necessary to develop a clearer conception of what constitutes “control,” such as is applicable to a wide class of fluid problems. Thus our generalization applies to the viscous Hele-Shaw weir (appendix) as well as to the quasigeostrophic problem that is the main focus of this paper. The distinguishing feature of the latter problem (as compared with classical hydraulics) is that there are many more unknowns (the “λ”) than dynamical conditions. Our generalization (section 2) of the classical Hugoniot idea to the case of a time-average flow requires that Q be extremized with respect to the degeneracy parameters and subject to a bifurcation condition on the mean flow. The scope of the new formulation is revealed by (3.11), which imply an Euler–Lagrange ODE for the structure of h(x). Additional constraints, as become known, may be necessary.

Although it will not be easy to solve for the stationary-wave condition, we may avail ourselves of the idea of a variational approximation, obtained by discarding conditions and/or constraining the form of the mean velocity profile [υ(x)]. Thus in section 2c we discarded the bifurcation condition, assumed that h(x) varies linearly, and maximized Q to predict the mean channel velocity (2.9). In section 3 we discarded the Q condition and evaluated the stationary-wave condition for arbitrary υ(x) (and PV) in a rectangular channel (Fig. 4). The main result (Fig. 5) is that a nonreversing and nonseparating jetlike υ(x) may exist at a hydraulic control section even if the width L is much larger than the radius of deformation. Because the mean hydrography of the Denmark Straits (Fig. 1) also has such left-hand support, it is suggested that the discharge could indeed be controlled.

Further investigation of this physically important point requires explicit consideration of bottom topography, extending the work begun by Killworth (1992). Perhaps the simplest relevant model (Fig. 6) is one with a uniformly sloping (α) bottom that intersects a vertical right-hand wall (at η = L), and which intersects the density interface at the (upstream) free streamline position η = 0. We also discarded the Q condition, constrained the form of υ to be uniform across the stream, and computed the condition for a stationary wave. Perhaps the most significant result of this paper is the prediction (4.20) of the width of the current, but when α < 0 (cf. Fig. 4b) there is no stationary wave. Because Fig. 4b corresponds to a flow that banks on the right-hand side of a strait, we conclude that this cannot be a control section; the latter condition requires the current to be supported on the left-hand side of the channel, that is, no separation.

These conditions may be extendable and testable in a laboratory experiment (see Whitehead et al. 1974) like that in Fig. 7, where the right-hand wall curves slightly into a wide upstream reservoir and the left-hand side of the current is supported on a uniformly sloping bottom. Geostrophic eddies in the reservoir may be produced by injecting a thin jet at middepth. Even if the resulting channel flow is quasi-laminar, it will not be possible to connect it (deterministically) to the source, in which case the experiment may provide a test of the ideas proposed herein.

Note added in proof. Why was U = constant chosen in section 4? For variable U(x) = g′/f∂(h + M)/∂η, the discharge is
i1520-0485-34-7-1692-eq11
For fixed (α, L, H1), the last term increases as ∫ ηU UmaxL2/2, and therefore maximum Q occurs for U(x) = Umax. This “endpoint” condition and the branch-point requirement (4.17) satisfy both control requirements (see section 2a and Fig. 2).

Acknowledgments

We gratefully acknowledge the support of the National Science Foundation (Grants OCE-0092504 and OCE-0236304).

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  • Helfrich, K. R., and L. J. Pratt, 2003: Rotating hydraulics and upstream basin circulation. J. Phys. Oceanogr, 33 , 16511663.

  • Howard, L. N., 1972: Bounds on flow quantities. Annu. Rev. Fluid Mech, 4 , 473494.

  • Killworth, P. D., 1992: Flow properties in rotating stratified hydraulics. J. Phys. Oceanogr, 22 , 9971017.

  • Killworth, P. D., 1994: On reduced gravity flow through sills. Geophys. Astrophys. Fluid Dyn, 75 , 91106.

  • Killworth, P. D., 1995: Hydraulic control and maximal flow in rotating stratified hydraulics. Deep-Sea Res, 42 , 859871.

  • LeBlond, P. H., and L. A. Mysak, 1978: Waves in the Ocean. Elsevier, 602 pp.

  • Pratt, L. J., and L. Armi, 1987: Hydraulic control of flows with nonuniform potential vorticity. J. Phys. Oceanogr, 17 , 20162029.

  • Shen, C. Y., 1981: The rotating hydraulics of the open channel flow between two basins. J. Fluid Mech, 112 , 161188.

  • Spall, M. A., and J. F. Price, 1998: Mesoscale variability in Denmark Strait: The PV outflow hypothesis. J. Phys. Oceanogr, 28 , 15981623.

    • Search Google Scholar
    • Export Citation
  • Stern, M. E., 1974: Comment on rotating hydraulics. Geophys. Fluid Dyn, 6 , 127130.

  • Stern, M. E., 1975: Ocean Circulation Physics. Academic Press, 242 pp.

  • Stern, M. E., 1980: A variational principle for turbulent flow. Phys. Fluids, 23 , 21612163.

  • Whitehead, J. A., A. Leetma, and R. A. Knox, 1974: Rotating hydraulics of strait and sill flows. Geophys. Fluid Dyn, 6 , 101125.

  • Worthington, L. V., and W. R. Wright, 1970: North Atlantic Ocean Atlas of Potential Temperature and Salinity in the Deep Water. Woods Hole Oceanographic Institute, Atlas Series 2, 24 pp. and 58 plates.

    • Search Google Scholar
    • Export Citation

APPENDIX

Supplementary Remarks on the Generality of the Control Principle

Although the main subject of this paper concerns geostrophic turbulence in the free discharge from a reservoir through a strait, it is important to illustrate the generality of the proposed control condition (in section 2) by showing that it also applies to a case for which the control of discharge has nothing to do with “stationary waves.”

Consider a Hele-Shaw weir consisting of two slightly separated vertical walls of downstream extent D, through which there is a free discharge of a very viscous fluid from an upstream reservoir where the given vertical height of the free surface is H at x = 0. Let h(x) denote the height of the free surface at 0 ≤ xD inside the weir and Q denote the volume flux. The well-known governing equation for this low-Reynolds-number flow yields a first-order ODE with given h(0) = H at the upstream end. The problem is undetermined because h(D) and Q are unknown. If we assume that a finite value of h(D)/H = O(1) is realized, then by decreasing/ increasing D, keeping Q constant (at its original value), we could find another solution of the ODE with larger/ smaller h(D) but with the remaining upstream part of h(x) unchanged. This, however, would violate our premise that the discharge is controlled by the weir topography; therefore the realized h(L′)/H cannot be O(1), but it must be h(L′)/H → 0. [This suggests the necessity of a lip boundary layer for which h(L′) = 0 is the leading-order condition.] The latter requirement then suffices to supply the leading-order value of Q. Of course, a little physical insight could have supplied us with the boundary condition [h(L′) ≅ 0] without the control hypothesis, but the point at stake is the general applicability of the control formulation in section 2.

Fig. 1.
Fig. 1.

The average depth (m) of the 2°C potential temperature surface according to Worthington and Wright (1970), along with that of the 1°C surface (inset). Note that the isothermal surfaces in the Denmark Strait (e.g., 500 m) intersect the bottom on both sides of the channel

Citation: Journal of Physical Oceanography 34, 7; 10.1175/1520-0485(2004)034<1692:LMFTOH>2.0.CO;2

Fig. 2.
Fig. 2.

An example of the control principle. Suppose that in addition to Q there is only one more unknown parameter (λ) than equations. The curves Q = Q(M, λ) are plotted for constant Q = Qa > Qb > Qc⋯. If p were an allowed solution point with ∂Q/∂λ ≠ 0, then the control condition would be violated because there is a neighboring point (r/s) for which Q = Qb = constant at either δM > 0 or δM < 0. On the other hand, the point t with ∂Q/∂λ = 0 has no neighbor solution with the same Q = Qb for δM > 0, and therefore this solution point satisfies the control condition. Point u illustrates an “endpoint,” arising from some inequality like G(M, λ) > 0. Although Q = Qb for δM > 0, there is no λ with Q = Qb for δM < 0; therefore point u satisfies the control condition

Citation: Journal of Physical Oceanography 34, 7; 10.1175/1520-0485(2004)034<1692:LMFTOH>2.0.CO;2

Fig. 3.
Fig. 3.

A triangular weir with uniform velocity υ = g′/f(dh/dx) used to compute the maximum discharge Q. For any given H1 the value of H0 is underdetermined (see text)

Citation: Journal of Physical Oceanography 34, 7; 10.1175/1520-0485(2004)034<1692:LMFTOH>2.0.CO;2

Fig. 4.
Fig. 4.

The upstream (y) section of a rectangular channel with arbitrary downstream velocity υ = g′/f(dh/dx). The vertical wall (x = L) curves slightly downstream (not shown), thereby causing the Lagrangian column at x to be displaced to ξ(x) in a steady forced flow. A bifurcation condition is sought (see section 3)

Citation: Journal of Physical Oceanography 34, 7; 10.1175/1520-0485(2004)034<1692:LMFTOH>2.0.CO;2

Fig. 5.
Fig. 5.

The modification of the stationary-wave condition for the jetlike υ(x) in (3.14b). Plot of Eqs. (3.15). The lines of constant λ > −1/(2π) have a contour interval of 0.01. For F = 2 and β = 6, λ = −0.04. For any given F, the value of β increases (and Q increases) as the jet amplitude λ increases, and Q reaches its upper bound at λ = −1/2π

Citation: Journal of Physical Oceanography 34, 7; 10.1175/1520-0485(2004)034<1692:LMFTOH>2.0.CO;2

Fig. 6.
Fig. 6.

A controlled flow over a sloping bottom. (a) (top) Vertical section of the basic upstream state with uniform geostrophic velocity U. The (dimensional) thickness (H) of the density current vanishes at η = 0. (a) (bottom) Plan view of the stationary wave; the streamline at η is displaced to y at the downstream location. (b) Vertical section of a current that has separated from the left-hand side of a channel (not shown) and “banks” on the right-hand side. Such a current cannot support a stationary wave

Citation: Journal of Physical Oceanography 34, 7; 10.1175/1520-0485(2004)034<1692:LMFTOH>2.0.CO;2

Fig. 7.
Fig. 7.

A volume flux (Q) of dense fluid is applied to the center of a relatively wide upstream reservoir with a uniformly sloping bottom. The (desired) geostrophic turbulence in the lower-density layer will build up a head, which causes a free discharge through the narrows of the sloping bottom. The approximate mean current width L at the narrows is given by (4.20)

Citation: Journal of Physical Oceanography 34, 7; 10.1175/1520-0485(2004)034<1692:LMFTOH>2.0.CO;2

1

Pratt (L. J. Pratt 2002, unpublished review manuscript) gives a similar definition.

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  • Borenas, K. M., and P. Lundberg, 1986: Rotating hydraulics of flow in a parabolic channel. J. Fluid Mech, 167 , 309326.

  • Gill, A. E., 1977: The hydraulics of rotating-channel flow. J. Fluid Mech, 80 , 641671.

  • Gregg, M. C., and E. Özsoy, 2002: Flow, water mass changes, and hydraulics in the Bosphorus. J. Geophys. Res.,107, 3016, doi:10.1029/2000JC000485.

    • Search Google Scholar
    • Export Citation
  • Helfrich, K. R., and L. J. Pratt, 2003: Rotating hydraulics and upstream basin circulation. J. Phys. Oceanogr, 33 , 16511663.

  • Howard, L. N., 1972: Bounds on flow quantities. Annu. Rev. Fluid Mech, 4 , 473494.

  • Killworth, P. D., 1992: Flow properties in rotating stratified hydraulics. J. Phys. Oceanogr, 22 , 9971017.

  • Killworth, P. D., 1994: On reduced gravity flow through sills. Geophys. Astrophys. Fluid Dyn, 75 , 91106.

  • Killworth, P. D., 1995: Hydraulic control and maximal flow in rotating stratified hydraulics. Deep-Sea Res, 42 , 859871.

  • LeBlond, P. H., and L. A. Mysak, 1978: Waves in the Ocean. Elsevier, 602 pp.

  • Pratt, L. J., and L. Armi, 1987: Hydraulic control of flows with nonuniform potential vorticity. J. Phys. Oceanogr, 17 , 20162029.

  • Shen, C. Y., 1981: The rotating hydraulics of the open channel flow between two basins. J. Fluid Mech, 112 , 161188.

  • Spall, M. A., and J. F. Price, 1998: Mesoscale variability in Denmark Strait: The PV outflow hypothesis. J. Phys. Oceanogr, 28 , 15981623.

    • Search Google Scholar
    • Export Citation
  • Stern, M. E., 1974: Comment on rotating hydraulics. Geophys. Fluid Dyn, 6 , 127130.

  • Stern, M. E., 1975: Ocean Circulation Physics. Academic Press, 242 pp.

  • Stern, M. E., 1980: A variational principle for turbulent flow. Phys. Fluids, 23 , 21612163.

  • Whitehead, J. A., A. Leetma, and R. A. Knox, 1974: Rotating hydraulics of strait and sill flows. Geophys. Fluid Dyn, 6 , 101125.

  • Worthington, L. V., and W. R. Wright, 1970: North Atlantic Ocean Atlas of Potential Temperature and Salinity in the Deep Water. Woods Hole Oceanographic Institute, Atlas Series 2, 24 pp. and 58 plates.

    • Search Google Scholar
    • Export Citation
  • Fig. 1.

    The average depth (m) of the 2°C potential temperature surface according to Worthington and Wright (1970), along with that of the 1°C surface (inset). Note that the isothermal surfaces in the Denmark Strait (e.g., 500 m) intersect the bottom on both sides of the channel

  • Fig. 2.

    An example of the control principle. Suppose that in addition to Q there is only one more unknown parameter (λ) than equations. The curves Q = Q(M, λ) are plotted for constant Q = Qa > Qb > Qc⋯. If p were an allowed solution point with ∂Q/∂λ ≠ 0, then the control condition would be violated because there is a neighboring point (r/s) for which Q = Qb = constant at either δM > 0 or δM < 0. On the other hand, the point t with ∂Q/∂λ = 0 has no neighbor solution with the same Q = Qb for δM > 0, and therefore this solution point satisfies the control condition. Point u illustrates an “endpoint,” arising from some inequality like G(M, λ) > 0. Although Q = Qb for δM > 0, there is no λ with Q = Qb for δM < 0; therefore point u satisfies the control condition

  • Fig. 3.

    A triangular weir with uniform velocity υ = g′/f(dh/dx) used to compute the maximum discharge Q. For any given H1 the value of H0 is underdetermined (see text)

  • Fig. 4.

    The upstream (y) section of a rectangular channel with arbitrary downstream velocity υ = g′/f(dh/dx). The vertical wall (x = L) curves slightly downstream (not shown), thereby causing the Lagrangian column at x to be displaced to ξ(x) in a steady forced flow. A bifurcation condition is sought (see section 3)

  • Fig. 5.

    The modification of the stationary-wave condition for the jetlike υ(x) in (3.14b). Plot of Eqs. (3.15). The lines of constant λ > −1/(2π) have a contour interval of 0.01. For F = 2 and β = 6, λ = −0.04. For any given F, the value of β increases (and Q increases) as the jet amplitude λ increases, and Q reaches its upper bound at λ = −1/2π

  • Fig. 6.

    A controlled flow over a sloping bottom. (a) (top) Vertical section of the basic upstream state with uniform geostrophic velocity U. The (dimensional) thickness (H) of the density current vanishes at η = 0. (a) (bottom) Plan view of the stationary wave; the streamline at η is displaced to y at the downstream location. (b) Vertical section of a current that has separated from the left-hand side of a channel (not shown) and “banks” on the right-hand side. Such a current cannot support a stationary wave

  • Fig. 7.

    A volume flux (Q) of dense fluid is applied to the center of a relatively wide upstream reservoir with a uniformly sloping bottom. The (desired) geostrophic turbulence in the lower-density layer will build up a head, which causes a free discharge through the narrows of the sloping bottom. The approximate mean current width L at the narrows is given by (4.20)

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