-plumes created by overflows - -plumes created by overflows shinichiro kida, james f. price, and...

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  • Ogzokmen, T.M. et al. 2001, On the dynamics of b-plumes. J Phys. Oceanogr 31/12, 3569-3680 Stommel, H. 1982, Is the South Pacific Helium dynamically active? Earth Planet. Sci. lett. 61, 63-67 Swaters, G.E. 1991, On the baroclinic instability of cold core coupled den- sity fronts on a sloping continental shelf. J. Fluid Mech. 224, 361-382 This research was supported by NSF Grant # *****


    5. Conclusions and Future Work  The Overflow can force a time mean circulation in the upper oceanic layer that has a magnitude as large as the over- flow itself.  Baroclinic Instability creates a double β-plume in the oce- anic layer through form drag. This effect can enhance the diabatic β-plume.  The β-plumes formed in this study depend on topograph- ic β and are thus locally confined. How these β−plumes can lead to a planetary scale feature, such as Azores Current, needs further investigation.

    4. Diabatic β-plume

    Diabatic β-plume Eddy driven β-plume

    Figure 8: How eddy driven β- plume can enhance the diabatic β-plume.Figure 7 Time averaged flow


     The effect of entrainment on the oceanic layer was studied by imposing a local diapycnal mass flux in the downstream basin. This created a dense layer going down the slope and a diabatic β-plume in the upper layer (Figure 7ab.) (The mass source/sink in the up/downstream basin was turned off.)  The transport of the β-plume was 3Sv which was larger than a linear solution of 2Sv. This was due to the eddy β− plume as shown schematically in Figure 8.

    Diapycnal cooling source (1.5 Sv Mass flux)

    (a)Overflow layer thickness

    (b) Sea surface height3. Eddy Driven β-plume

    Figure 5: The basic mechanism for the formation of eddy driven β-plume.

    400km 0




    800 km

    300m thick

    400km 0

    Bathymetry lines

    5cm 40cm/s 40km

    Figure 3: Snap shot view(birds eye view) of the two layers Eddies formed.

    Figure 4: Time average the oceanic layer shows the double β-plume.

    (a)Overflow layer thickness

    (b) Sea surface height

    (a)Overflow layer thickness

    (b) Sea surface height

    400km 0

    400km 0

    400km 0

    400km 0

    No induced flow!

    Figure 6: No eddy induced flow when there is no instability.

    (a)Overflow layer thickness

    (b) Sea surface height

    2 Sv 5cm/s

    Lower Basin

    Oceanic layer: Double β-plume


    Form Drag



     The adiabatic interaction between the overflow and upper layer was studied by shutting off mixing be- tween the layers.  The overflow (Figure 3a) was baroclinically unstable (Swaters 1991) and the upper oceanic layer flow (Figure 3b, shown by the sea surface height) had strong eddies [40cm/s, 40km wide.] These eddies in- duced an eddy driven mean flow in the upper layer as a double β-plume (Figure 4b.) Its transport was 2Sv, a size comparable to the overflow transport.  Figure 5 shows the mechanism of how the double β-plume was driven. As the overflow exited the strait, it became baroclincally unstable and created eddies. Then part of its momentum is transfered across layers as a form drag. This form drag created the eddy PV flux divergence term in the vorticity equation (Section 1) and thus drove the eddy driven β-plume.

     The momentum transfer was a drag to the overflow layer (Figure 4a) and caused it to go down the slope with a steeper angle than the Ekman number.  Figure 6 shows a case when the upper layer was thicker (+1000m.) Baroclinic instability did not happen and no eddies formed, so there was no eddy induced upper layer flow

    2. Model setup

    Figure 2: Model Setup

    Upstream Basin (Marginal sea)

    Downstream Basin (Open Ocean)



    Strait 40km wide & 1000m deep

    Slope 0.01

    Dense Overflow Layer

     We used a two layer (overflow and upper oceanic layer) isopycnal model (Hallberg Isopycnal Model) on a f-plane. β is introduced by the bottom topography shown below.  1.5Sv of dense water is pumped from the upstream basin to the downstream basin. This creates the overflow layer going through the narrow strait.  Mixing between layers is either shut off or specified.  A linear bottom drag acts on the layer that is in contact with the bottom topography.

    Sea surface

     Overflows vigorously entrain oceanic water. How does the upper ocean respond to this local entrainment?  A small local forcing can induce a strong horizontal circulation(Figure 1, Stommel 1984). This concept is the β− plume and will be used in this study. The equation that ex- presses this concept is the vorticity equation written below.  This study focuses on the overflow ocean interaction in the presence of a sloping bottom topography using a two layer model.

    1. Introduction

    U,q, and w* are Transport, PV, and cross isopycnal velocity.

    Diabatic β-plume: Cross PV gradient flow by a diapycnal mass flux

    Eddy driven β-plume: Cross PV gradient flow by an eddy PV flux divergence

    Local entrainment w*>0 causing a Cross PV gradient flow

    Figure 1: Schematic of a diabatic β-plume in a rectangular basin.


    NUq = -q’U’+qw*+Friction

    -Plumes Created by Overflows Shinichiro Kida, James F. Price, and Jiayan Yang 1. MIT/WHOI Joint Program in Oceanography 2. Woods Hole Oceanographic Institution

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    Contact: [email protected]