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Kenneth G. Hughes, James N. Moum, and Emily L. Shroyer

descent of the DWL. Using a large-eddy simulation under 6 m s −1 winds, Sarkar and Pham (2019) show that the lower half of the mixed layer warmed over their week-long simulation (net mean surface heat input to the ocean was 95 W m −2 ). Yet the turbulent heat transport causing this warming occurred over only four hours each day during the late afternoon to evening. The timing and magnitude of subsurface turbulent heat flux governs SST and the consequent latent, sensible, and longwave radiative

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Adam V. Rydbeck, Tommy G. Jensen, and Matthew R. Igel

et al. (2010 , 2012a , b) first documented the relationship between downwelling equatorial Rossby waves and ISO convective onset, and their work inspired much of this investigation. Webber et al. (2010) suggested that surface latent heat flux anomalies manifest in response to warm SST anomalies associated with oceanic downwelling equatorial Rossby waves and are responsible for ISO convective onset. However, using reanalysis data, Rydbeck and Jensen (2017) did not observe notable increases

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Corinne B. Trott, Bulusu Subrahmanyam, Heather L. Roman-Stork, V. S. N. Murty, and C. Gnanaseelan

eddy variability ( Schott et al. 2009 ; Dandapat and Chakraborty 2016 ; Mahadevan et al. 2016a , b ). Highly dynamic heat and moisture fluxes drive the ISOs in the BoB and bring in seasonal and complex subseasonal variability ( Goswami et al. 2016 ; Weller et al. 2016 ; Sanchez-Franks et al. 2018 ). The ISOs of the BoB can be categorized into three major components of atmospherically driven coupled air–sea oscillations: the 30–90-day signal associated with the monsoon ISO (MISO) and Madden

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Kenneth G. Hughes, James N. Moum, and Emily L. Shroyer

). The buoyancy input in the same units as Eq. (7) is (9) J b t = g α ρ w c p   J q t , where J b is a buoyancy flux (m 2 s −3 ), g is gravitational acceleration, α is the thermal expansion coefficient of seawater (3 × 10 −4 K −1 ), ρ w and c p are the density and specific heat capacity of seawater, and J q is the heat flux through the air–sea interface minus the radiative flux that penetrates deeper than h . Although J q is sinusoidal in time, for simplicity we will treat it as a

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Wei-Ting Chen, Chien-Ming Wu, and Hsi-Yen Ma

:// . 10.1175/2007JCLI1457.1 Iacono , M. J. , J. S. Delamere , E. J. Mlawer , M. W. Shephard , S. A. Clough , and W. D. Collins , 2008 : Radiative forcing by long-lived greenhouse gases: Calculations with the AER radiative transfer models. J. Geophys. Res. , 113 , D13103, 10.1029/2008JD009944 . 10.1029/2008JD009944 Jakob , C. , and A. P. Siebesma , 2003 : A new subcloud model for mass-flux convection schemes: Influence on triggering

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Benjamin A. Toms, Susan C. van den Heever, Emily M. Riley Dellaripa, Stephen M. Saleeby, and Eric D. Maloney

troposphere, which has been found to be an important factor in the overall evolution of the MJO and other convectively coupled equatorial waves ( Peters and Bretherton 2006 ; Benedict and Randall 2007 ; Adames and Kim 2016 ). It would therefore be interesting to expand upon the existing literature citing the importance of convective features in the upscale redistribution of water vapor using methods such as spectral flux, as detailed in Arbic et al. (2012) and Hayashi (1980) , to identify the scales

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Emily M. Riley Dellaripa, Eric D. Maloney, Benjamin A. Toms, Stephen M. Saleeby, and Susan C. van den Heever

cm −3 ). These concentrations were used successfully in a radiative convective equilibrium simulation with RAMS ( Igel et al. 2017 ). For all simulations the model domain is 1000 km × 1000 km in extent, has horizontal grid spacing of 2 km, and is centered over Luzon ( Fig. 1 ). There are 42 stretched vertical levels up to 25 km. The lowest levels are 125 m apart with seven model levels occurring under 1 km, while the top 13 levels are spaced 1 km apart. A vertical stretch ratio of 1.08 is

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