Seasonality of the submesoscale dynamics in the Gulf Stream region

Abstract

Frontogenesis and frontal instabilities in the mixed layer are known to be important processes in the formation of submesoscale features. We study the seasonality of such processes in the Gulf Stream (GS) region. To approach this problem, a realistic simulation with the Hybrid Coordinate Ocean Model is integrated for 18 months at two horizontal resolutions: a high-resolution (1/48°) simulation able to resolve part of the submesoscale regime and the full range of mesoscale dynamics, and a coarser resolution (1/12°) case, in which submesoscales are not resolved. Results provide an insight into submesoscale dynamics in the complex GS region. A clear seasonal cycle is observed, with submesoscale features mostly present during winter. The submesoscale field is quantitatively characterized in terms of deviation from geostrophy and 2D dynamics. The limiting and controlling factor in the occurrence of submesoscales appears to be the depth of the mixed layer, which controls the reservoir of available potential energy available at the mesoscale fronts that are present most of the year. Atmospheric forcings are the main energy source behind submesoscale formation, but mostly indirectly through mixed layer deepening. The mixed layer instability scaling suggested in the (Fox-Kemper et al., J Phys Oceanogr 38:1145–1165, 2008) parametrization appears to hold, indicating that the parametrization is appropriate even in this complex and mesoscale dominated area.

Appendix

1.1 Surface fluxes parametrization

Thermal energy flux into the ocean (f _T [in watt per square meter]) is computed by HYCOM from the balance between incident radiation and emitted ocean radiation ($\mathcal {R}$), the latent heat transfer ($\mathcal {H}$), and the sensible heat transfer due to evaporation (𝜖),

$$ f_{T} = \mathcal{R} - \epsilon - \mathcal{H}. $$

(15)

$\mathcal {R}$ (in watt per square meter), the net solar radiation (as well as its short and long wave components), is positive into the ocean and is provided as forcing. The surface radiation budget also includes a black body radiation correction flux, proportional to the difference between the surface temperature and the air temperature of the data used for the forcing (ERA40).

Latent heat transfer due to evaporation, 𝜖, is computed from

$$ \epsilon = L_{\mathrm{E}} \rho_{\mathrm{a}} u^{*} C_{\mathrm{L}} (T_{s} - T_w), $$

(16)

where L _E is the latent heat of evaporation coefficient (2.47·10⁶ j kg⁻¹); ρ _a is the air density computed from air temperature; u ^∗ is the 10-m wind speed; C _L is the latent heat flux coefficient computed from a polynomial expression, function of stability, and temperature difference between atmosphere and ocean (Kara et al. 2000; Fairall et al. 2003); and T _s and T _w are, respectively, the saturation mixing ratio (from a polynomial expression, function of surface temperature) and water vapor mixing ratio which is provided as part of the (ERA40) forcing.

The last term in the expression of f _T is the sensible heat transfer, $\mathcal {H}$, obtained from

$$ \mathcal{H} = C_{P_{\text{air}}}\rho_{a} C_{\mathrm{S}} u^{*}(T_{\text{sur}} - T_{\text{atm}}), $$

(17)

where $C_{P_{\text {air}}}$ is the specific heat of the air at constant pressure (in joule per kilogram per degree); C _S is the sensible heat flux coefficient computed by Kara et al. (2000) as 0.9554·C _L; T _sur is the model sea surface temperature; and T _atm is the air temperature.

The salinity flux into the ocean, f _S ([10⁻³ kg m⁻² s⁻¹]), is quantified in the model as follows:

$$ f_S = (E-P)\cdot (S\cdot 10^3), $$

(18)

where E the is evaporation rate

$$ E = \epsilon\cdot 10^{-3}/L_{\mathrm{E}}, $$

(19)

P is precipitation (given as a forcing), and S is the salinity at the surface. In addition, a relaxation of sea surface salinity to climatology is included.

In Fig. 19a, the seasonal variability of f _T averaged over region A is shown. Values are negative during the winter season and positive during the summer, representing a heat flux from the ocean to atmosphere during the winter and a net heat gain of the ocean during the summer. Analogously, in Fig. 19b, the spatially averaged values of f _S in time are shown with maxima and minima shifted of approximately half period with respect to f _T . Average values of f _S are always positive in region A, evaporation being larger than precipitation; f _S contributes to the mixed layer being saltier during winter and fresher during summer.

Fig. 19

Temporal variability of the bulk fluxes of a heat (in watt per square meter), b salinity [10⁻³ kg m⁻² s⁻¹], and c wind speed (in meter per second) averaged over region A. The gray line represents the first point of the time series considered.

During the winter season, negative values of f _T act to increase density reducing the ocean temperature, while positive values of f _S contribute to increase salinity and thus density. During the summer season on the other side, f _S is nearly zero due to weak evaporation, and density is driven only by f _T which is positive into the ocean. Buoyancy fluxes ultimately affect MLD by stratifying the upper ocean during summer and destratifying the mixed layer in winter.

Wind speed (u ^∗) enters in the equations as a factor in the formulation of sensible heat and evaporation. The wind speed seasonal cycle (Fig. 19c) shows a maximum in winter and a cycle similar to f _S . Wind speed enters in the formulation of both f _S and f _T favoring the latent heat transfer due to evaporation (16) and thus producing a loss of heat (15) and a gain of salinity (19) at the surface.