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B. Praveen Kumar, Eric D’Asaro, N. Sureshkumar, E. Pattabhi Rama Rao, and M. Ravichandran

dissipation ε despite the large errors in the forcing. In the range of L T / L MO corresponding to our data, ε increases while ε T decreases. This decrease is consistent with that predicted by (9) within the estimated errors (cloud of small dots) as shown by the red line, implying a prediction of (11) ε T = 1.15 B 0 ⁡ ( H L MO ) 1 / 2 for the wind-forced regime. At the smallest values of L T / L MO , for BoB16, the value of ε T is about 0.8, far below the LG89 scaling and clearly in the

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

4a (500 × 2-m grid cells horizontally and 500 m deep with 0.2-m resolution near the surface), we force with a constant wind speed of 4 m s −1 , and we apply diurnal heating that has zero net heat input on daily average ( Fig. 8a ). Penetrating solar radiation is treated with a nine-band formulation with a solar-angle dependence described by Gentemann et al. (2009) . Because the grid is two-dimensional ( x – z ), we set Coriolis frequency f to zero rather than let the diurnal jet veer

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Dipanjan Chaudhuri, Debasis Sengupta, Eric D’Asaro, R. Venkatesan, and M. Ravichandran

km to the right of the track) and BD10 (near the track) employ temperature and salinity initial conditions constructed from the mooring data interpolated in the vertical. Model vertical resolution is 0.25 m, and the time step is 1 h. Surface forcing is based on observed hourly incoming shortwave and longwave radiation, and turbulent fluxes are estimated from hourly moored measurements of air temperature, surface pressure, sea surface temperature, relative humidity, and wind using the COARE 3

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

surface water was typically advected 3 km farther per day than water at 30 m. The shear that occurs between the diurnal jet and the mixed layer (0.03 s −1 ; Sutherland et al. 2016 ; Bogdanoff 2017 ) is comparable to that found in estuarine flows (0.05 s −1 ; Stacey and Pond 1997 ), at the base of internal solitary waves (0.05 s −1 ; Moum et al. 2003 ), and in the sheared layer above the equatorial undercurrent (0.02 s −1 ; Smyth et al. 2013 ). Under weak forcing (wind < 2 m s −1 ), clear sky, and

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

case, the sun’s heat is spread throughout the mixed layer and warms each parcel of water by O (0.1°C) by midafternoon. In the latter case, warming is concentrated in the top 2 m and, consequently, more of this heat is likely to be transferred from the ocean back to the atmosphere over a short time scale. In between these extremes heat transport is more complicated. Warming of the lower half of the mixed layer, for example, lags the surface solar forcing by several hours because it depends on the

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Jai Sukhatme, Dipanjan Chaudhuri, Jennifer MacKinnon, S. Shivaprasad, and Debasis Sengupta

rotational component. Moreover, from 80 to 10 km, the observed anomalous scaling of velocity increments observed in the ocean data presented here is consistent with three-dimensional stratified turbulence in other geophysical fluids ( Lohse and Xia 2010 ). Specifically, in situ measurements of stratified turbulence, for example, through marine clouds ( Siebert et al. 2010 ) and the atmospheric surface layer ( Chu et al. 1996 ) also show anomalous scaling and non-Gaussian distributions of velocity

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