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Carsten Eden, Richard J. Greatbatch, and Dirk Olbers

Eq. (2) into the instantaneous tracer budget and averaging the result leads to the equation for the mean tracer b given by An immediate difficulty is presented by the eddy tracer flux 1 . These fluxes couple the mean tracer budget to that of the perturbations, such that the evolution of the perturbations has to be known to predict the mean tracer. Of course, the solution to this problem is thought to be given by parameterizing the perturbation quantities in terms of the mean quantities

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Emily Shuckburgh, Guillaume Maze, David Ferreira, John Marshall, Helen Jones, and Chris Hill

intimately connected to irreversible processes such as lateral small-scale mixing and damping processes associated with air–sea fluxes ( Zhai and Greatbatch 2006a , b ; Greatbatch et al. 2007 ). It is a quantification of the latter process that is the focus of attention here. Figure 1 shows a wintertime instantaneous ( Fig. 1a ) and monthly-mean ( Fig. 1b ) net air–sea heat flux obtained from a global ⅛° eddy-resolving model driven by observed atmospheric fields through bulk formulas that allow the

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G. L. Geernaert

1. Introduction For more than a half century, air–sea interaction experimentalists have employed the bulk aerodynamic method in compiling estimates of momentum, heat, and water vapor fluxes. Flux estimates are applied in turn to numerical models of oceanic and/or atmospheric circulations, wave state, waveguide prediction, and loads on engineered infrastructures. More recently, flux estimates have served to support major remote sensing programs: for example, where microwave signatures of the

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Jonathan P. Fram, Maureen A. Martin, and Mark T. Stacey

during the ebb–flood transition ( Fram 2005 ). Scalars at the ocean–estuary interface, such as salts, are transported by these flow processes. Here, we characterize the effect of each process on salinity exchange under different oceanic conditions, tidal forcing, and freshwater input. a. Estuarine flux decomposition Cross-sectionally integrated salt transport can be described quantitatively using the advection–diffusion equation ( Kay et al. 1996 ; Geyer and Nepf 1996 ), with rivers advecting salt

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Bruce A. Warren

1. Introduction In stationary conditions at a level sea surface, the vertical velocity ( w ) that is induced by evaporation ( E ) and precipitation ( P ) is nearly ( E − P ). In recognition that the mass flux into (or out of) the atmosphere is of freshwater alone, attempts have been made to improve the representation: w = ( E − P )/(1 − S ), where S is the mass-fraction salinity, at the sea surface (e.g., Schmitt et al. 1989 ); or w = ρ F ( E − P )/[ ρ (1 − S )], where ρ is

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B. Curry, C. M. Lee, and B. Petrie

created using further low-passed data (10-day cutoff to reduce tidal and meteorological variability) and spatially averaged into domains defined by depth (0–150, 200–250, and 500 m) and location (e.g., the shelves and WGSC–BIC frontal zone). The mean and variable fields were mapped onto a regular, two-dimensional grid with 4-m cells at depths <150 m and 10-m cells at depths >150 m at a horizontal resolution of 5 km (see appendix B of online supplement). c. Flux calculations Daily volume, freshwater

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Edgar L. Andreas

, thus, reclaim some of the sensible heat they have lost. Quantifying the net heating of the atmosphere that is mediated by spray has therefore been illusive. The rate of this net heating is usually termed the enthalpy flux and is the sum of the total air–sea sensible and latent heat fluxes ( Businger 1982 ). I use the adjective “total” here to recognize the possibility that the relevant fluxes comprise contributions from both the usual interfacial sensible and latent heat fluxes (molecular

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María Aristizábal and Robert Chant

1. Introduction The along-channel salinity structure in estuaries is the result of two competing processes: the advection of freshwater oceanward as a result of river output and the flux of salt into the estuary because of processes such as steady shear dispersion and tidal oscillatory salt flux ( Hansen and Rattray 1965 ; Zimmerman 1986 ; Lerczak et al. 2006 ). The vertical salinity structure is maintained by a competition between the straining of the horizontal salinity gradient, which is

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Jeffrey Coogan and Brian Dzwonkowski

for Eq. (1) may no longer be satisfied, and these changes in the salt budget can be examined by where overbars denote depth-averaged salinity s and velocity u ; primes are vertical deviations from the depth-averaged velocity and salinity; K H is the horizontal diffusion; t is time; x is along-estuary distance; and is the along-estuary salinity gradient. On the right-hand side of Eq. (2) , the three terms account for the salinity flux through advection, exchange associated with

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Edgar L. Andreas, P. Ola G. Persson, and Jeffrey E. Hare

. The fluxes via these two routes scale differently ( Andreas 1994 ; Andreas and DeCosmo 2002 ). For example, although the Tropical Ocean-Global Atmosphere Coupled Ocean–Atmosphere Response Experiment (COARE) version 3.0 bulk flux algorithm ( Fairall et al. 2003 ) has been tuned with flux data collected in wind speeds up to 20 m s −1 and is therefore operationally useful in this wind speed range, it is based strictly on interfacial scaling and thus may not be reliable if it is extrapolated to wind

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