Search Results
instability leads to a forward energy cascade. These submesoscale flows are typically restricted to the mixed layer of the ocean because strong forcing from wind and strain by mesoscale features creates fast flows over short length scales [where Ro ~ O (1)]. Convection and wind also make the near-surface stratification very weak (Ri ≲ 1). Since submesoscale flows occur at the upper boundary layer, they coexist with wind and wave forcing. Despite having a partially geostrophically balanced state, these
instability leads to a forward energy cascade. These submesoscale flows are typically restricted to the mixed layer of the ocean because strong forcing from wind and strain by mesoscale features creates fast flows over short length scales [where Ro ~ O (1)]. Convection and wind also make the near-surface stratification very weak (Ri ≲ 1). Since submesoscale flows occur at the upper boundary layer, they coexist with wind and wave forcing. Despite having a partially geostrophically balanced state, these
motions is dissipated for . In this article, we explore a possible pathway for energy removal from fronts in the regime involving internal waves. Previous studies have shown that internal waves can be trapped by ocean fronts (e.g., Kunze and Sanford 1984 ; Kunze et al. 1995 ; Rainville and Pinkel 2004 ; Whitt and Thomas 2013 ; Joyce et al. 2013 ) and that they can exchange energy with them ( Thomas and Taylor 2014 ). Internal waves in the ocean are usually generated by the winds and tides and
motions is dissipated for . In this article, we explore a possible pathway for energy removal from fronts in the regime involving internal waves. Previous studies have shown that internal waves can be trapped by ocean fronts (e.g., Kunze and Sanford 1984 ; Kunze et al. 1995 ; Rainville and Pinkel 2004 ; Whitt and Thomas 2013 ; Joyce et al. 2013 ) and that they can exchange energy with them ( Thomas and Taylor 2014 ). Internal waves in the ocean are usually generated by the winds and tides and
scales are also important in the mixing and dynamics of the upper ocean. These processes are three-dimensional and nonhydrostatic (see the review by Sullivan and McWilliams 2010 ) and include wave-driven Langmuir turbulence, which consists of disordered collections of counterrotating Langmuir cells. These cells create convergence zones at the ocean surface where foam, plankton, and other debris collect in long “windrows” ( Langmuir 1938 ). The length of the cells may extend up to 1 km, while their
scales are also important in the mixing and dynamics of the upper ocean. These processes are three-dimensional and nonhydrostatic (see the review by Sullivan and McWilliams 2010 ) and include wave-driven Langmuir turbulence, which consists of disordered collections of counterrotating Langmuir cells. These cells create convergence zones at the ocean surface where foam, plankton, and other debris collect in long “windrows” ( Langmuir 1938 ). The length of the cells may extend up to 1 km, while their
( Lien et al. 1995 ; McPhaden 2002 ), the Madden–Julian oscillation ( Chi et al. 2014 ), and tropical instability waves (TIWs) ( Menkes et al. 2006 ; Moum et al. 2009 ; Inoue et al. 2012 ) may have an important impact on the mean state of the Pacific Ocean and thus on global climate. Generated in the eastern tropical Pacific and Atlantic Oceans, TIWs propagate westward with wavelengths of 700–1600 km and periods of 15–40 days ( Qiao and Weisberg 1995 ; Kennan and Flament 2000 ; Willett et al
( Lien et al. 1995 ; McPhaden 2002 ), the Madden–Julian oscillation ( Chi et al. 2014 ), and tropical instability waves (TIWs) ( Menkes et al. 2006 ; Moum et al. 2009 ; Inoue et al. 2012 ) may have an important impact on the mean state of the Pacific Ocean and thus on global climate. Generated in the eastern tropical Pacific and Atlantic Oceans, TIWs propagate westward with wavelengths of 700–1600 km and periods of 15–40 days ( Qiao and Weisberg 1995 ; Kennan and Flament 2000 ; Willett et al
allows frontogenesis. b. The internal-wave continuum Starting with the seminal work of Garrett and Munk (1972) , it was realized that internal waves are characterized by a continuum spectrum spanning frequencies from f to N , over which a few spectral peaks due to tides and inertial motions are superimposed. The internal-wave continuum spectrum is understood to be set through weak interactions of linear waves (e.g., Lvov et al. 2004 ). It is remarkably uniform across the ocean and has been shown
allows frontogenesis. b. The internal-wave continuum Starting with the seminal work of Garrett and Munk (1972) , it was realized that internal waves are characterized by a continuum spectrum spanning frequencies from f to N , over which a few spectral peaks due to tides and inertial motions are superimposed. The internal-wave continuum spectrum is understood to be set through weak interactions of linear waves (e.g., Lvov et al. 2004 ). It is remarkably uniform across the ocean and has been shown
case, but also because oceanic eddies are generally too small to bring the classical criticality parameter to unity ( Jansen and Ferrari 2012 ): the zonally and time-averaged baroclinic velocity is of the order of 0.2 m s −1 , much larger than the phase speed of long baroclinic Rossby waves, which, with a deformation radius of about 15 km, is about 100 times smaller than the observed zonal-mean velocities. Many eddy-permitting and eddy-resolving ACC modeling studies have examined the sensitivity of
case, but also because oceanic eddies are generally too small to bring the classical criticality parameter to unity ( Jansen and Ferrari 2012 ): the zonally and time-averaged baroclinic velocity is of the order of 0.2 m s −1 , much larger than the phase speed of long baroclinic Rossby waves, which, with a deformation radius of about 15 km, is about 100 times smaller than the observed zonal-mean velocities. Many eddy-permitting and eddy-resolving ACC modeling studies have examined the sensitivity of
input. Hence, there remains a gap in our understanding of how the interior ocean subinertial circulation dissipates. Fig . 1. Schematic of this study. Wind power input to the ocean general circulation is estimated to be O (1) TW ( Wunsch 1998 ). This power input is balanced by energy dissipation processes such as bottom drag O (0.1) TW ( Wunsch and Ferrari 2004 ) and lee-wave generation O (0.2 TW) ( Nikurashin and Ferrari 2011 ) near the bottom boundary. The study described here suggests that an
input. Hence, there remains a gap in our understanding of how the interior ocean subinertial circulation dissipates. Fig . 1. Schematic of this study. Wind power input to the ocean general circulation is estimated to be O (1) TW ( Wunsch 1998 ). This power input is balanced by energy dissipation processes such as bottom drag O (0.1) TW ( Wunsch and Ferrari 2004 ) and lee-wave generation O (0.2 TW) ( Nikurashin and Ferrari 2011 ) near the bottom boundary. The study described here suggests that an
waves across the ocean domain, after which the dynamics of the circulation change little and the system continues to adjust to a long-term stratification and overturning rate. Kelvin and Rossby wave propagation was found to influence the equilibrating ocean on a range of time scales from months or years ( Kawase 1987 ; Yang 1999 ) up to several hundreds of years ( Huang et al. 2000 ). Johnson and Marshall (2002) considered the response to a change in deep-water formation at high latitudes using
waves across the ocean domain, after which the dynamics of the circulation change little and the system continues to adjust to a long-term stratification and overturning rate. Kelvin and Rossby wave propagation was found to influence the equilibrating ocean on a range of time scales from months or years ( Kawase 1987 ; Yang 1999 ) up to several hundreds of years ( Huang et al. 2000 ). Johnson and Marshall (2002) considered the response to a change in deep-water formation at high latitudes using
. McWilliams , 2006 : Application of the ROMS embedding procedure for the central California upwelling system . Ocean Modell. , 12 , 157 – 187 , doi: 10.1016/j.ocemod.2005.05.002 . Plougonven , R. , and C. Snyder , 2007 : Inertia-gravity waves spontaneously generated by jets and fronts. Part I: Different baroclinic life cycles . J. Atmos. Sci. , 64 , 2502 – 2520 , doi: 10.1175/JAS3953.1 . Ponte , A. , P. Klein , X. Capet , P.-Y. Le Traon , B. Chapron , and P. Lherminier
. McWilliams , 2006 : Application of the ROMS embedding procedure for the central California upwelling system . Ocean Modell. , 12 , 157 – 187 , doi: 10.1016/j.ocemod.2005.05.002 . Plougonven , R. , and C. Snyder , 2007 : Inertia-gravity waves spontaneously generated by jets and fronts. Part I: Different baroclinic life cycles . J. Atmos. Sci. , 64 , 2502 – 2520 , doi: 10.1175/JAS3953.1 . Ponte , A. , P. Klein , X. Capet , P.-Y. Le Traon , B. Chapron , and P. Lherminier
1. Introduction Abyssal turbulence plays an important role in maintaining the observed oceanic stratification ( Munk and Wunsch 1998 ; Wunsch and Ferrari 2004 ). Internal waves in the vicinity of rough topography in the deep ocean are thought to be a major contributor to abyssal turbulence based on observations of bottom-enhanced turbulence and mixing ( Polzin et al. 1997 ; Ledwell et al. 2000 ; Laurent et al. 2001 ; Klymak et al. 2006 ) at these sites. At deep rough topography, there is
1. Introduction Abyssal turbulence plays an important role in maintaining the observed oceanic stratification ( Munk and Wunsch 1998 ; Wunsch and Ferrari 2004 ). Internal waves in the vicinity of rough topography in the deep ocean are thought to be a major contributor to abyssal turbulence based on observations of bottom-enhanced turbulence and mixing ( Polzin et al. 1997 ; Ledwell et al. 2000 ; Laurent et al. 2001 ; Klymak et al. 2006 ) at these sites. At deep rough topography, there is
