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1. Introduction Traditionally, the effects of nonlinearities of the equation of state on the transport of buoyancy have often been assumed to be negligible. For example Tziperman (1986) investigates the density budget for a fixed volume with a linear equation of state, and Munk and Wunch (1998) investigate the oceanic energy budget using the same equation of state. However, there are exceptions. Davis (1994) looked into density production by the nonlinearities of the equation of state and
1. Introduction Traditionally, the effects of nonlinearities of the equation of state on the transport of buoyancy have often been assumed to be negligible. For example Tziperman (1986) investigates the density budget for a fixed volume with a linear equation of state, and Munk and Wunch (1998) investigate the oceanic energy budget using the same equation of state. However, there are exceptions. Davis (1994) looked into density production by the nonlinearities of the equation of state and
modeling the fair-weather regime is the simultaneous representation of the wind-induced mixing and the maintenance of the upper-oceanic stratification. We pose the problem in a one-dimensional (1D) vertical column with surface momentum and buoyancy fluxes (including a mean wind stress, mean heating, and solar absorption), vertical turbulent mixing, and idealized 1D representations of the heat flux from the interior 3D circulation that acts to cool the interior and to provide a stably stratified balance
modeling the fair-weather regime is the simultaneous representation of the wind-induced mixing and the maintenance of the upper-oceanic stratification. We pose the problem in a one-dimensional (1D) vertical column with surface momentum and buoyancy fluxes (including a mean wind stress, mean heating, and solar absorption), vertical turbulent mixing, and idealized 1D representations of the heat flux from the interior 3D circulation that acts to cool the interior and to provide a stably stratified balance
that model results depend strongly on how a particular turbulence parameterization balances these competing effects. We revisit the buoyancy arrest initial-value problem here to present two classes of results. First, we obtain general expressions for the Ekman transport and buoyancy shutdown time scales. Second, we derive new scales for boundary layer thickness (hence adjustment time) that account for finite boundary layer static stability in both the upwelling and downwelling cases. These results
that model results depend strongly on how a particular turbulence parameterization balances these competing effects. We revisit the buoyancy arrest initial-value problem here to present two classes of results. First, we obtain general expressions for the Ekman transport and buoyancy shutdown time scales. Second, we derive new scales for boundary layer thickness (hence adjustment time) that account for finite boundary layer static stability in both the upwelling and downwelling cases. These results
) and Chapman (2000b) showed how physics related to the buoyancy arrest mechanism determine the cross-shelf locations of surface-to-bottom fronts in the coastal ocean. Further, Chapman and Lentz (1994) and Chapman (2000a) show how the absence of bottom stress in an adjusted current greatly extends the alongshore scale of an along-isobath current. It is not obvious how these results, derived for a nearly steady flow, might be affected by the presence of a fluctuating flow. In addition, there
) and Chapman (2000b) showed how physics related to the buoyancy arrest mechanism determine the cross-shelf locations of surface-to-bottom fronts in the coastal ocean. Further, Chapman and Lentz (1994) and Chapman (2000a) show how the absence of bottom stress in an adjusted current greatly extends the alongshore scale of an along-isobath current. It is not obvious how these results, derived for a nearly steady flow, might be affected by the presence of a fluctuating flow. In addition, there
convective systems contributing more to the rapid precipitation increases. The proposed physical argument for the precipitation onset is buoyancy centric. Holloway and Neelin (2009 , hereafter HN09 ) showed, using a steady-state entraining plume model, that for environmental moisture values at and beyond precipitation onset, the entraining plumes are positively buoyant near the freezing level. The implication is that if a convective entity—often represented by a bulk-entraining plume—can survive mixing
convective systems contributing more to the rapid precipitation increases. The proposed physical argument for the precipitation onset is buoyancy centric. Holloway and Neelin (2009 , hereafter HN09 ) showed, using a steady-state entraining plume model, that for environmental moisture values at and beyond precipitation onset, the entraining plumes are positively buoyant near the freezing level. The implication is that if a convective entity—often represented by a bulk-entraining plume—can survive mixing
vorticity dynamics, and this part of the dynamics has the same action-at-a-distance features as quasigeostrophic dynamics. This paper presents a vorticity–buoyancy view of gravity waves, examines how this interplay between vorticity and buoyancy affects the evolution of stratified shear flow anomalies, and explores its use as a basis for a kernel view. Our general motivation in developing a vorticity-based kernel view of the dynamics is quite basic: in a similar way in which the KRW perspective has
vorticity dynamics, and this part of the dynamics has the same action-at-a-distance features as quasigeostrophic dynamics. This paper presents a vorticity–buoyancy view of gravity waves, examines how this interplay between vorticity and buoyancy affects the evolution of stratified shear flow anomalies, and explores its use as a basis for a kernel view. Our general motivation in developing a vorticity-based kernel view of the dynamics is quite basic: in a similar way in which the KRW perspective has
the waters sank to deeper depths were the same as the regions where the waters became more dense. Although large buoyancy loss clearly drives water mass transformation in the interior of many basins, observations, modeling, and theoretical studies all suggest that the net downwelling in regions of buoyancy loss in the interior of the ocean is negligible ( Schott et al. 1993 ; Send and Marshall 1995 ; Marotzke and Scott 1999 ; Marshall and Schott 1999 ; Spall 2003 , 2004 ). There is intense
the waters sank to deeper depths were the same as the regions where the waters became more dense. Although large buoyancy loss clearly drives water mass transformation in the interior of many basins, observations, modeling, and theoretical studies all suggest that the net downwelling in regions of buoyancy loss in the interior of the ocean is negligible ( Schott et al. 1993 ; Send and Marshall 1995 ; Marotzke and Scott 1999 ; Marshall and Schott 1999 ; Spall 2003 , 2004 ). There is intense
versus that of the free troposphere in different climate regimes has remained unclear. In addition, the influence of surface conditions versus lateral entrainment in/above the ABL on temperature mixing, humidity dilution, and ice formation of a convective parcel in different oceanic and continental climate regimes have not yet been investigated systematically. For convective instability, parcel theory is commonly used to calculate parcel buoyancy and some derived indices, such as convective available
versus that of the free troposphere in different climate regimes has remained unclear. In addition, the influence of surface conditions versus lateral entrainment in/above the ABL on temperature mixing, humidity dilution, and ice formation of a convective parcel in different oceanic and continental climate regimes have not yet been investigated systematically. For convective instability, parcel theory is commonly used to calculate parcel buoyancy and some derived indices, such as convective available
) in ways that can enhance or suppress deep convection ( Mapes 2000 ; Raymond et al. 2006 ; Kuang 2008 ). Furthermore, observations suggest that water vapor fluctuations in the LFT play a central role in the organization of deep convection ( Raymond 2000 ; Grabowski and Moncrieff 2004 ; Sahany et al. 2012 ). This regulating role is achieved through dry air entrainment and dilution reducing the buoyancy of rising cumulus clouds ( Lucas et al. 1994 ; Hannah 2017 ; Kuo et al. 2017 ). To further
) in ways that can enhance or suppress deep convection ( Mapes 2000 ; Raymond et al. 2006 ; Kuang 2008 ). Furthermore, observations suggest that water vapor fluctuations in the LFT play a central role in the organization of deep convection ( Raymond 2000 ; Grabowski and Moncrieff 2004 ; Sahany et al. 2012 ). This regulating role is achieved through dry air entrainment and dilution reducing the buoyancy of rising cumulus clouds ( Lucas et al. 1994 ; Hannah 2017 ; Kuo et al. 2017 ). To further
exported below the mixed layer must be associated with buoyancy fluxes, whether they come from surface fluxes, diffusive fluxes, advection, or mixing, before taking part of the global overturning circulation ( Walin 1982 ; Speer et al. 2000 ; Kuhlbrodt et al. 2007 ; Cessi 2019 ). The production of AAIW and SAMW results from both horizontal advection induced by wind stress and eddies and from the surface buoyancy fluxes ( Sloyan and Rintoul 2001 ). The surface buoyancy loss during winter produces
exported below the mixed layer must be associated with buoyancy fluxes, whether they come from surface fluxes, diffusive fluxes, advection, or mixing, before taking part of the global overturning circulation ( Walin 1982 ; Speer et al. 2000 ; Kuhlbrodt et al. 2007 ; Cessi 2019 ). The production of AAIW and SAMW results from both horizontal advection induced by wind stress and eddies and from the surface buoyancy fluxes ( Sloyan and Rintoul 2001 ). The surface buoyancy loss during winter produces
