The bottom boundary layer of the deep ocean

1977 ◽  
Vol 24 (4) ◽  
pp. 182
Keyword(s):  
Boundary Layer ◽  
Deep Ocean ◽  
Bottom Boundary Layer ◽  
Bottom Boundary

scholarly journals Stratified Inertial Subrange Inferred from In Situ Measurements in the Bottom Boundary Layer of the Rockall Channel

2010 ◽  
Vol 40 (11) ◽  
pp. 2401-2417 ◽  
Author(s):  
Pascale Bouruet-Aubertot ◽  
Hans van Haren ◽  
M. Pascale Lelong
Keyword(s):  
Boundary Layer ◽  
Energy Levels ◽  
Deep Ocean ◽  
Energy Dissipation Rate ◽  
Bottom Boundary Layer ◽  
Inertial Subrange ◽  
Buoyancy Frequency ◽  
Bottom Boundary ◽  
Wide Range ◽  
Taylor Hypothesis

Abstract Deep-ocean high-resolution moored temperature data are analyzed with a focus on superbuoyant frequencies. A local Taylor hypothesis based on the horizontal velocity averaged over 2 h is used to infer horizontal wavenumber spectra of temperature variance. The inertial subrange extends over fairly low horizontal wavenumbers, typically within 2 × 10−3 and 2 × 10−1 cycles per minute (cpm). It is therefore interpreted as a stratified inertial subrange for most of this wavenumber interval, whereas in some cases the convective inertial subrange is resolved as well. Kinetic energy dissipation rate ε is inferred using theoretical expressions for the stratified inertial subrange. A wide range of values within 10−9 and 4 × 10−7 m2 s−3 is obtained for time periods either dominated by semidiurnal tides or by significant subinertial variability. A scaling for ε that depends on the potential energy within the inertio-gravity waves (IGW) frequency band PEIGW and the buoyancy frequency N is proposed for these two cases. When semidiurnal tides dominate, ε ≃ (PEIGWN)3/2, whereas ε ≃ PEIGWN in the presence of significant subinertial variability. This result is obtained for energy levels ranging from 1 to 30 times the Garrett–Munk energy level and is in contrast with classical finescale parameterization in which ε ∼ (PEIGW)2 that applies far from energy sources. The specificities of the stratified bottom boundary layer, namely a weak stratification, may account for this difference.

Simulations of coastal upwelling on the Sydney Continental Shelf

2000 ◽  
Vol 51 (6) ◽  
pp. 577 ◽  
Author(s):  
Patrick Marchesiello ◽  
Mark T. Gibbs ◽  
Jason H. Middleton
Keyword(s):  
Boundary Layer ◽  
Coastal Upwelling ◽  
Deep Ocean ◽  
Bottom Boundary Layer ◽  
Return Flow ◽  
Pressure Gradients ◽  
Bottom Boundary ◽  
Eastern Australia ◽  
Flow Through ◽  
Good Agreement

Two-dimensional numerical simulations of the response of the coastal waters of Sydney, south-eastern Australia, to idealized upwelling-favourable winds are presented. The spin up of the upwelling circulation is investigated, in particular the structure of the nearshore circulation. The intensity of the final upwelling state is found to be strongly linked to the activation of the return flow through the bottom boundary layer, which is also related to the strength of imposed alongshore pressure gradients. Results from a simulation of upwelling forced by a deep-ocean alongshore-current jet also show the final upwelling state to be weak in comparison with upwelling states produced by the action of the local wind stress. Bottom boundary layer shut-down in the presence of such a forcing jet is also discussed. A simulation of a real upwelling event was also performed and good agreement was found between the simulation and observations from a field experiment performed during summer 1994 in the Sydney coastal ocean.

scholarly journals Submesoscale Baroclinic Instability in the Bottom Boundary Layer

2018 ◽  
Vol 48 (11) ◽  
pp. 2571-2592 ◽  
Author(s):  
Jacob O. Wenegrat ◽  
Jörn Callies ◽  
Leif N. Thomas
Keyword(s):  
Boundary Layer ◽  
Mixed Layer ◽  
Baroclinic Instability ◽  
Slope Angle ◽  
Deep Ocean ◽  
Bottom Boundary Layer ◽  
Surface Mixed Layer ◽  
Nonlinear Simulation ◽  
Bottom Boundary ◽  
Topographic Slope

AbstractWeakly stratified layers over sloping topography can support a submesoscale baroclinic instability mode, a bottom boundary layer counterpart to surface mixed layer instabilities. The instability results from the release of available potential energy, which can be generated because of the observed bottom intensification of turbulent mixing in the deep ocean, or the Ekman adjustment of a current on a slope. Linear stability analysis suggests that the growth rates of bottom boundary layer baroclinic instabilities can be comparable to those of the surface mixed layer mode and are relatively insensitive to topographic slope angle, implying the instability is robust and potentially active in many areas of the global oceans. The solutions of two separate one-dimensional theories of the bottom boundary layer are both demonstrated to be linearly unstable to baroclinic instability, and results from an example nonlinear simulation are shown. Implications of these findings for understanding bottom boundary layer dynamics and processes are discussed.

The bottom boundary layer of the deep ocean

1976 ◽  
Vol 81 (27) ◽  
pp. 4983-4990 ◽  
Author(s):  
Laurence Armi ◽  
Robert C. Millard
Keyword(s):  
Boundary Layer ◽  
Deep Ocean ◽  
Bottom Boundary Layer ◽  
Bottom Boundary

Update on the Global Energy Dissipation Rate of Deep-Ocean Low-Frequency Flows by Bottom Boundary Layer

2018 ◽  
Vol 48 (6) ◽  
pp. 1243-1255 ◽  
Author(s):  
Chao Huang ◽  
Yongsheng Xu
Keyword(s):  
Boundary Layer ◽  
Kinetic Energy ◽  
Dissipation Rate ◽  
Low Frequency ◽  
Deep Ocean ◽  
Power Input ◽  
Energy Dissipation Rate ◽  
Bottom Boundary Layer ◽  
Geostrophic Currents ◽  
Bottom Boundary

AbstractThe global dissipation caused by bottom boundary layer drag is one of the major pathways for the consumption of kinetic energy in the deep ocean. However, the spatial distribution and global integral of the drag dissipation are still debatable. This paper presents an updated estimate of the dissipation rate, using the barotropic component of surface geostrophic currents and 632 in situ velocity measurements. Also, the seafloor roughness is proposed as a parameter of drag efficiency in the parameterized method. The results provide a map of the drag dissipation rate with a global integral of ~0.26 TW. Approximately 66% of this dissipation occurs in the Southern Ocean, which is consistent with the proportion of wind power input into this region. Building upon the work in previous studies on the bottom boundary layer drag, more long-period observations are used, eliminating the influence of the baroclinic contribution to the surface geostrophic currents in the construction of the bottom velocity, and taking topographic roughness into account. The estimates have implications for the maintenance of density structure in the deep ocean and understanding of the kinetic energy budget.

scholarly journals Accelerating Dense-Water Flow down a Slope

2009 ◽  
Vol 39 (6) ◽  
pp. 1495-1511 ◽  
Author(s):  
John M. Huthnance
Keyword(s):  
Boundary Layer ◽  
Density Gradient ◽  
Slope Angle ◽  
Deep Ocean ◽  
Bottom Boundary Layer ◽  
Bottom Slope ◽  
Initial Value ◽  
Bottom Boundary ◽  
Downslope Flow ◽  
Horizontal Density Gradient

Abstract Where water is denser on a shallow shelf than in the adjacent deep ocean, it tends to flow down the slope from shelf to ocean. The flow can be in a steady bottom boundary layer for moderate combinations of upslope density gradient −ρx∞ and bottom slope (angle θ to horizontal):Here g is acceleration due to gravity, ρ0 is a mean density, and f is twice the component of the earth’s rotation normal to the sloping bottom. For stronger combinations of the horizontal density gradient and bottom slope, the flow accelerates. Analysis of an idealized initial value problem shows that, when b ≥ 1, there is a bottom boundary layer with downslope flow, intensifying exponentially at a rate fb2(1 + b)−1/2/2, and slower-growing flow higher up. For stronger stratification b > 21/2, that is, a relatively weak Coriolis constraint, the idealized problem posed here may not be the most apposite but suggests that the whole water column accelerates, at a rate [ρ0−1|ρx∞|g sinθ]1/2 if f is negligible.

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