Geostrophy, Vorticity, and Sverdrup October 6

Geostrophic Flow: Assume level of no motion

Geostrophic Currents Figure 10.9 in Stewart. Mean geopotential anomaly of the Pacific Ocean relative to the 1,000 dbar surface based on 36,356 observations.

Figure 10.10 in Stewart. Current meter measurements can be used with CTD measurements to determine current as a function of depth avoiding the need for assuming a depth of no motion. Solid line: profile assuming a depth of no motion at 2000 decibars. Dashed line: profile adjusted to agree with currents measured by current meters 1–7 days before the CTD measurements. Plots from Tom Whitworth, Texas A&M University.

Lagrangean Measurements of Currents Figure 10.13 in Stewart. Satellite systems, especially System Argos or GPS with Iridium telemetry, use radio signals transmitted from surface buoys to determine the position of the buoy

Figure 10.14 in Stewart. ARGO floats are widely used by the oceanographic community to measure deeper currents at a depth of 1 km in the ocean, and to profile temperature and salinity as a function of depth between 2 km and the surface.

Figure 10.15 in Stewart. Distribution of tritium along a section through the western basins in the North Atlantic, measured in 1972 (Top) and remeasured in 1981 (Bottom).

Figure 10.16 in Stewart. Ocean temperature and current patterns are combined in this AVHRR analysis. Surface currents were computed by tracking the displacement of small thermal or sediment features between a pair of images.

Figure 10.17 in Stewart. Trajectories that spilled rubber duckies would have followed had they been spilled on January 10 of different years.

Eulerian Measurements of Currents Figure 10.18 in Stewart. Left: An example of a surface mooring of the type deployed by the Woods Hole Oceanographic Institution’s Buoy Group. Right: An example of a subsurface mooring deployed by the same group.

Figure 10.19 in Stewart. An example of a moored current meter, the RCM 9 produced by Aanderaa Instruments

Vorticity Vorticity – tendency to rotate Planetary vorticity - f Like angular momentum – must be conserved Planetary vorticity - f Local z Side view Top view N Pole Rotation about local z relative to fixed space Eq.

Relative Vorticity – ω – current shear x y Positive rotation (RH rule) Negative rotation

Total or Absolute Vorticity Absolute Vorticity = ω +f Conserved, ie : Potential Vorticity - Move water parcel to a greater latitude without other forces, ω must change (if h=const) - h=depth (or height) of fluid column (book uses D)

Flow over a bump (ridge) Nothing to change ω “Topographic Steering”- flow wants to follow contours of f/h=constant

Sverdrup Harald Sverdrup (1947) showed that the circulation in the upper kilometer or so of the ocean is directly related to the curl of the wind stress. He came upon relating the curl of the wind stress to mass transport within the upper ocean. In deriving the relationship, Sverdrup assumed - the flow is stationary - that lateral friction and molecular viscosity are small - turbulence near the sea surface can be described using an eddy viscosity - the flow is baroclinic and that the wind-driven circulation vanishes at some depth of no motion

Sverdrup integrated these equations from the surface to a depth -D equal to or greater than the depth at which the horizontal pressure gradient becomes zero. Mx, My are the mass transports in the wind-driven layer extending down to an assumed depth of no motion. Note – they include the Ekman transport and geostrophic transport!

Tx and Ty are the components of the wind stress The horizontal boundary condition at the sea surface is the wind stress, and the boundary at depth -D is zero stress because the currents go to zero.   Tx and Ty are the components of the wind stress

Using these definitions and boundary conditions these equations become

Sverdrup integrated the continuity equation over the same vertical depth, assuming the vertical velocity at the surface and at depth -D are zero, to obtain: Differentiating the past two equations with respect to y and x, subtracting, and using the above equation we get: where β = ∂f/∂y is the rate of change of Coriolis parameter with latitude, and where curlz (T) is the vertical component of the curl of the wind stress.

where R is Earth's radius and φ is latitude Over much of the open ocean, especially in the tropics, the wind is zonal (E-W) and ∂Ty/∂x is sufficiently small that

Substituting My into: Sverdrup obtained (in spherical coordinates)

Sverdrup integrated this equation from a north-south eastern boundary at x = 0, assuming no flow into the boundary. This requires Mx = 0 at x = 0. where Δx is the distance from the eastern boundary of the ocean basin, and brackets indicate zonal averages of the wind stress

Figure 11.1 in Stewart. Streamlines of mass transport in the eastern Pacific calculated from Sverdrup’s theory using mean annual wind stress.

Figure 11.2 in Stewart. Mass transport in the eastern Pacific calculated from Sverdrup’s theory using observed winds (solid lines) and pressure calculated from hydrographic data from ships (dots). Transport is in tons per second through a section one meter wide extending from the sea surface to a depth of one kilometer. Note the difference in scale between My and Mx.

Figure 11.3 in Stewart. Depth-integrated Sverdrup transport applied globally using the wind stress from Hellerman and Rosenstein (1983). Contour interval is 10 Sverdrups.

Stream lines, Path lines, and Stream functions At each instant in time, we can represent the flow field in a fluid by a vector velocity at each point in space. The instantaneous curves that are everywhere tangent to the direction of the vectors are called the stream lines of the flow. If the flow is unsteady, the pattern of stream lines change with time. The trajectory of a fluid particle, the path followed by a Lagrangean drifter, is called the path line in fluid mechanics. The path line is the same as the stream line for steady flow, and they are different for an unsteady flow.

We can simplify the description of two-dimensional, incompressible flows by using the stream function ψ defined by: The stream function is often used because it is a scalar from which the vector velocity field can be calculated. This leads to simpler equations for some flows.

Stream functions are also useful for visualizing the flow Stream functions are also useful for visualizing the flow. At each instant, the flow is parallel to lines of constant ψ. Thus if the flow is steady, the lines of constant stream function are the paths followed by water parcels. Figure 11.4 in Stewart. Volume transport between stream lines in a two-dimensional, steady flow.

The volume rate of flow between any two stream lines of a steady flow is dψ, and the volume rate of flow between two stream lines ψ1 and ψ2 is equal to ψ1 – ψ2. To see this, consider an arbitrary line dx = (dx, dy) between two stream lines (Figure 11.4). The volume rate of flow between the stream lines is: and the volume rate of flow between the two stream lines is numerically equal to the difference in their values of ψ.

lets apply the concepts to satellite-altimeter maps of the oceanic topography

In addition to the stream function, oceanographers use the mass-transport stream function Ψ defined by This is the function shown in Figures 11.2 and 11.3.