1. Introduction
The vertical exchange of surface waters with intra- and subpycnocline waters is an important process for the subduction of water masses modified by air–sea fluxes and the transport of nutrients necessary for sustaining marine ecosystems in the photic zone. Vertical motions are enhanced in regions of high mesoscale and submesoscale variability owing to a variety of processes. Shear and strain associated with such flows disrupt the thermal wind balance and hence generate vertical circulations to restore geostrophy over subinertial time scales (Hoskins et al. 1978). Vertical circulations can also arise due to lateral variations of turbulent fluxes of momentum and buoyancy generated during the spin down of meso/submesoscale flows (Garrett and Loder 1981; Thompson 2000; Nagai et al. 2006). Forcing such flows by winds can drive Ekman pumping/suction owing to the modification of the Ekman transport by the vertical component of the relative vorticity (Stern 1965; Niiler 1969; Thomas and Lee 2005). Additionally, when the wind blows in the direction of the surface geostrophic shear, Ekman advection of denser water over light leads to localized mixing in regions of strong baroclinicity, which drives an ageostrophic secondary circulation (ASC) (Thomas and Lee 2005). Furthermore, if the mesoscale/submesoscale flow is thrown out of balance, time-dependent vertical motions can be triggered as the fluid adjusts to geostrophic equilibrium (Ou 1984; Tandon and Garrett 1994).
Regardless of the process leading to the generation of vertical motions, it is of interest to diagnose the vertical velocity from observed meso/submesoscale flows. Presently, there are no observational techniques capable of measuring the vertical velocity w directly over areas large enough to encompass meso/submesoscale features. Consequently, w must be inferred. Vertical velocities are typically inferred through solutions of the quasigeostrophic (QG) omega equation (e.g., Pollard and Regier 1992; Rudnick 1996; Viúdez et al. 1996; Allen and Smeed 1996; Shearman et al. 1999). If the meso/submesoscale flow is embedded at least partially in the mixed layer, where the stratification is weak and the Richardson number is low, and the currents have significant Rossby numbers, the QG approximation may not be valid. For large Rossby number flows, semigeostrophic (Naveira Garabato et al. 2001) and intermediate model (Shearman et al. 2000) versions of the omega equation have been solved to diagnose the vertical velocity from observations. Comparisons with numerical simulations show that these higher-order formulations of the omega equation yield more accurate estimates of the vertical velocity, although the presence of inertia–gravity waves and diffusive processes can lead to significant errors (Viúdez and Dritschel 2004). These errors arise because diagnostics based on the omega equation implicitly assume that the ageostrophic velocity is dominantly driven by the shear and strain of the geostrophic flow and, hence, neglect effects due to friction, mixing of buoyancy, and/or time dependence. To account for these effects, Giordani et al. (2006) derived the “generalized omega equation” in which the Q vector, the divergence of which drives the omega equation, was modified to include contributions from turbulent fluxes of momentum and buoyancy and higher-order terms involving time dependence and advection of/by ageostrophic flows. Giordani et al. used the generalized omega equation to diagnose the vertical velocity from realistic numerical simulations of the northeastern Atlantic and found that the nontraditional forcing terms of their equation can result in significant vertical velocities. Similar to these simulations, in the real ocean it is likely that friction, mixing of buoyancy, and/or time dependence are important in the dynamics of the vertical velocity of upper-ocean meso/submesoscale flows, especially when such flows are exposed to atmospheric forcing. This implies that solutions to the traditional omega equation may poorly represent the actual vertical velocity of upper-ocean meso/submesoscale currents and motivates the use of a different method to estimate the vertical circulation. In this paper, we develop an inverse method to infer w that assesses, rather than specifies, the driving mechanism for the vertical circulation. The method is more general than omega-equation diagnostics and hence is more appropriate for use with upper-ocean datasets.
In this article, the method is applied to observations of the subpolar front of the Japan/East Sea (JES). This paper is one in a series of three detailing the dynamics of the subpolar front of the JES. The first article (C. M. Lee et al. 2008, unpublished manuscript) describes hydrographic, velocity, and shipboard meteorological measurements made at the front during periods of strong atmospheric forcing associated with cold-air outbreaks in the winter of 2000. The observations reveal thermohaline intrusions, pycnostads with weak stratification and low potential vorticity, and plumes with high chlorophyll fluorescence that evidence subduction and vertical circulation at the front. The inverse method developed and outlined in this article is used to quantify this vertical circulation. The third paper in the series (Y. Yoshikawa et al. 2008, unpublished manuscript), describes numerical experiments of an idealized representation of the subpolar front of the JES designed to study the response of the frontal vertical circulation and subduction to wind stress and heat loss and to characterize the driving forces responsible for the vertical circulation for various combinations of wind and buoyancy forcing.
The outline of the paper is as follows: first, the equation governing the ageostrophic secondary circulation that forms the theoretical basis for the inverse method is derived. Next, the inverse problem is formulated. Following this, the strategy and numerical method for the inverse calculation are outlined. In section 5, the inverse solution for the vertical circulation at the subpolar front of the JES is described. A discussion on the role of wind-driven time-dependent motions in the frontal vertical circulation and the advection of tracers by this circulation is given in section 6. The paper is concluded in section 7.
2. Governing equations for ageostrophic secondary circulations
Similar to the geostrophic forcing, vertically varying frictional forces and laterally varying buoyancy sources/sinks (terms II and III, respectively) will disrupt the thermal wind balance and hence drive ageostrophic circulations (Eliassen 1951). Also note that the equations governing Ekman dynamics are encompassed in (8) when II is balanced by the lhs terms with coefficients F12 and F22. Term III is important at wind-forced fronts where Ekman flow can advect dense water over light, generating convection and turbulent mixing of buoyancy localized to the front. This mixing drives (through term III) frontogenetic ageostrophic secondary circulations (Thomas and Lee 2005).
It should be noted that the generalized omega equation of Giordani et al. (2006) captures the same dynamics as Eq. (8).1 However, the two equations differ in their choice of variable: that is, a vector streamfunction in (8) and the vertical velocity w in the generalized omega equation. The vertical derivative of the vector streamfunction is equal to the horizontal component of the ageostrophic velocity, a quantity that, unlike w, can be calculated from observations. Consequently, Eq. (8) is better suited to an inverse analysis than the generalized omega equation since observations can be used to constrain its solutions.
3. Formulation of inverse problem
4. Strategy and method for calculating inverse solution
a. Issues involving temporal variability
The inverse method relies on the spatial structure of the inferred ageostrophic velocity (13) to select the columns of VG that best match the observations. However, apparent spatial structure could arise in the gridded velocity observations (13) due to temporal variability since typical SeaSoar/ADCP surveys last several days. This is a longer period than the characteristic time scale of near-inertial motions, internal gravity waves, and tidal flows that could be present during a survey. This masking of temporal variability as spatial variability is problematic for the inverse calculation and requires a strategy to minimize its effects. The strategy employed on the winter 2000 dataset from the subpolar front of JES is described below.
The SeaSoar/ADCP surveys at the subpolar front of the JES consisted of 5–6 north–south, cross-frontal sections separated by ∼20 km in the zonal direction (Fig. 1). The SeaSoar profiled repeatedly from the ocean surface to roughly 350-m depth while being towed at 8 kt. This configuration provided ∼3 km along-track profile separation. Horizontal velocities were measured continuously using a 150-kHz ADCP installed on the ship, yielding velocity profiles with a vertical resolution of 8 m and a vertical extent of ∼350 m. The velocity measurements were averaged every 3 min, resulting in an along-track resolution of ∼0.75 km. The velocity and density fields used in the inverse calculation were objectively mapped using the method of Le Traon (1990). Each field minus a quadratic fit was mapped at each depth using 5% noise and an anisotropic cosine modulated Gaussian covariance function with e-folding lengths of 20 and 10 km in the x and y directions and wavelengths for the cosine modulation of 25 and 20 km in x and y directions, respectively.
The region closest to the front is of most interest for the analysis since it is anticipated that the strongest vertical motions are located there. By using only data collected within a certain time period T before and after the time from when the ship crossed the center of the front, the data can be considered approximately synoptic if T is much smaller than the time scale of the ageostrophic flow. This argument makes the case for having T be small. However, if T is too small, then the distance transited by the ship Ls = UshipT (Uship the ship’s velocity) might not encompass the major features responsible for the vertical circulation at the front. Thus, in choosing T, a compromise must be made between ensuring synopticity and having sufficient spatial coverage.
For the analysis of the vertical circulation at the subpolar front, T was chosen to be 1.3 h or 0.07 inertial periods. This length of time was selected so as to minimize aliasing by inertial oscillations, the dominant time-dependent motions in the region (Takematsu et al. 1999). The distance traveled by the ship in this time is Ls ≈ 20 km. This distance is sufficient to cover submesoscale features in the hydrography that evidence vertical circulation at the subpolar front. These features are associated with plumes of low salinity water with high chlorophyll fluorescence extending from the surface mixed layer to the pycnocline that appear to have been subducted on the dense side of the front. An example of such a plume is shown in Fig. 2. Plumes such as this were observed on the majority of sections and were characterized by a horizontal scale of 5–10 km, a distance smaller than Ls. Thus, the strategy for the inverse calculation is to perform the analysis section by section, using data collected within ±1.3 h or ±20 km from the frontal crossing. In using this strategy, however, it is necessary that the inverse solution calculated on each section be treated independently because significant temporal variation in the ageostrophic flow is likely from section to section, given the ∼0.6 inertial period transit time from one frontal crossing to the next.
b. Two-dimensional methodology
For a fully three-dimensional ageostrophic flow, the value of the vector streamfunction (ϕ, ψ) along a meridional section, at a zonal location x = xo, will be influenced by rhs forcing terms of (8) nonlocal to that section. That is, the solution vector of the inverse problem, p, will in general have nonzero elements for grid locations at x ≠ xo. From (23) it can be seen that, depending on the number of elements in p relative to the number of data points in vag, (23) changes from an under- to an overdetermined system of equations. For the strategy described above, where only data taken on each section is used to constrain the solution to minimize temporal aliasing, the inverse problem involves solving an underdetermined system of equations, since the elements of vag only include observations taken along x = xo, whereas those of p include contributions from locations where x = xo and x ≠ xo. If the ageostrophic flow were approximately two dimensional, that is, invariant in the zonal direction, however, then p would be at most a weak function of x and velocity data along a single section could provide enough information to make (23) a square or overdetermined system of equations. To perform the inverse calculation, we will therefore make the assumption that the ageostrophic flow is approximately two dimensional while performing tests to determine the validity of the two-dimensional assumption.
1) Numerical method
2) Validity of the two-dimensional solution
Inspection of (8) reveals that lateral variability in the ageostrophic flow field can be driven by horizontal variations in the forcing terms I–V and/or the coefficients of the differential operator. These coefficients are functions of the geostrophic flow and buoyancy fields, which vary in the zonal direction (e.g., Fig. 1). It is also likely that the forcing terms of (8) vary zonally as well. Given the three-dimensional structure of these fields, the validity of the two-dimensional, zonally invariant solution to (24) must be discussed.
In using (24) instead of (8) we assume that the terms on the lhs of (8) that are not included in (24) are negligibly small. Scaling arguments show that, of these terms, the largest is a factor of Bu(Ly /Lx) times the leading order terms F12∂2ϕ/∂z2 and F22∂2ψ/∂z2 in the equations, where Bu = N 2H2/f 2Ly2 is the Burger number of the ageostrophic flow and Lx and Ly are characteristic length scales of the ageostrophic flow in the x and y directions. The Burger number of the ageostrophic flow varies with the phenomena responsible for that flow. For example, if the ageostrophic flow is associated with an inertia–gravity wave with a frequency ω, the dispersion relation of these waves predicts Bu ≈ (ω/f )2 − 1. For near-inertial waves with ω/f ≈ 1, Bu ≪ 1, suggesting that for such motions it is justified to neglect the terms in (8) that are not found in (24). If, in addition, the near-inertial waves have a longer wavelength in the zonal versus meridional direction, then Bu(Ly /Lx) < Bu ≪ 1, further justifying the use of (24) to represent their motions. We therefore conclude that, if near-inertial waves are present in the observations, the two-dimensional inverse solution will accurately capture their ageostrophic circulation.
The Burger number of submesoscale geostrophic flows is typically O(1) (Thomas et al. 2008). Through geostrophic forcing these flows can drive ageostrophic motions with Burger numbers similar to those of the geostrophic currents that drive them. Therefore, the use of the two-dimensional version of (8) for such flows comes into question. To check the validity of the two-dimensional method for this case, solutions to the semigeostrophic omega equation were calculated. The vertical velocity obtained by solving the three- and two-dimensional versions of the omega equations, that is, (16) and 𝗟2Dq = p, with p equal to the geostrophic forcing calculated from the objectively mapped buoyancy and geostrophic velocity fields were compared. Scatterplots of the two- and three-dimensional solutions (
5. Vertical circulation and forcing terms of the ageostrophic flow at the subpolar front of the JES
a. Inferred vertical velocity
Cross sections of the vertical velocity inferred using the inverse method for all sections occupied during three of the surveys of the subpolar front of the JES are plotted in Fig. 4. As evident in the figure, submesoscale laterally banded structures in w of O(5–10 km) in width are prominent. As described in the appendix, analyses aimed at assessing random and systematic errors in the inverse method arising from noise in the data, aliasing, and boundary conditions indicate that features in the vertical velocity field are robust for z > −150 m. For depths greater than 150 m, the arbitrary w = 0 bottom boundary condition can influence the solution and lead to significant errors. The magnitude of w is large, with amplitudes of O[1–2 mm s−1 (∼100–200 m day−1)]. Comparing the inverse solution to the solution to the omega equation,
b. Slantwise nature of ageostrophic flow
c. Driving forces for the ageostrophic flow
Apart from the vertical velocity, the inverse method yields, in the 2MN element vector p, estimates for the rhs forcing terms of (8). The first (last) MN elements of p correspond to the sum of all the rhs forcing terms in the upper (lower) equation of (8) and is referred to as RHS1 (RHS2). We would like to know what physical process—geostrophic forcing, friction, mixing of buoyancy, time-dependence, and/or nonlinearities in the ageostrophic flow [i.e., the terms I–V in (8)]—is associated with (RHS1, RHS2). As neither microstructure measurements nor repeated observations of the ageostrophic flow capable of sufficiently sampling inertial/superinertial motions were made during the cruise, terms II–V cannot be estimated directly. The geostrophic forcing (9) can be calculated from the observations; therefore, it is the only forcing term that we can state with certainty does or does not resemble (RHS1, RHS2) in structure and magnitude.
1) Geostrophic forcing
The degree to which the spatial structure of the rhs of (8) resembles the geostrophic forcing is quantified by calculating the correlations between RHS1 and
2) Friction and mixing of buoyancy
Although no direct measurements of turbulent mixing of momentum and buoyancy were made, the meteorological observations of wind stress (τwx, τwy) taken from the ship can be used to derive scalings for terms II and III of (8), assuming that these terms are primarily associated with wind-driven frictional forces and mixing of buoyancy caused by the Ekman advection of dense water over light.
3) Time dependence
In the previous two sections, it was shown that the estimates and scalings for the contributions from geostrophic forcing, friction, and mixing of buoyancy to the generation of ageostrophic motions are all a fraction of the magnitude of the inferred rhs forcing term of (8). This suggests that the remaining terms on the rhs of (8), namely time variability in the ageostrophic shear and/or nonlinear advection by the ageostrophic flow, contribute significantly to the dynamics of the frontal ASC.
The rate of change of the thermal wind imbalance following the geostrophic flow can drive ageostrophic flows through term IV. To fully evaluate this term, information is required on the local rate change and spatial variation in the direction of the geostrophic flow of the ageostrophic vertical shear. Owing to the infrequent sampling of the ageostrophic flow during the cruise, such information is not available. However, there are indications from the structure of the rhs forcing terms and the ageostrophic velocity field that time dependence is a critical driving term of (8) at the front.
Two physical phenomena may explain the CW spiraling of ageostrophic flow: Ekman shear or downward propagation of an internal wave (Leaman and Sanford 1975). If friction were the cause of the spiraling, then the vertical length scale associated with the spiral (∼50 m) would scale with the Ekman depth, requiring a vertical viscosity of around 0.1 m2 s−1. A vertical viscosity of this magnitude in the pycnocline (where the spiraling is enhanced) is implausible. Therefore, it is more likely that the turning of the velocity vector is a consequence of downward propagating internal waves rather than frictional effects. If this spiraling was associated with an internal wave, then the wave’s frequency should follow the internal gravity wave (hydrostatic) dispersion relation ω =
Assuming that the turning of the ageostrophic velocity vector is due to internal waves, then the cross-front contrast in the relative strengths of CW versus CCW energy shown in Fig. 7d can be interpreted as evidence of an asymmetry in the amount of energy propagating up versus down on either side of the front. This asymmetry was not an isolated occurrence specific to this particular frontal crossing. Averaging ECW and ECCW over all the sections for meridional locations south of the front yields 0.0032 and 0.0022 m2 s−2 in the clockwise and counterclockwise components of the flow, respectively. On the north side of the front, there is on average 0.0031 and 0.0029 m2 s−2 in the clockwise and counterclockwise components of the flow, suggesting that there is less energy propagating upward south of the front. The region of enhanced CW versus CCW energy is coincident with an area of anticyclonic vorticity (e.g., Fig. 7d). The increase in the energy of downward propagating waves in regions of negative vorticity was also observed at the subpolar front of the JES during a summertime cruise (Shcherbina et al. 2003), indicating that time-dependent ageostrophic motions influenced by the frontal vorticity field may be a common feature to the front during all seasons. In addition, eddy-resolving numerical simulations reveal an enhancement of downward propagating near-inertial energy in regions of anticyclonic vorticity, consistent with these findings (Lee and Niiler 1998; Zhai et al. 2007).
4) Higher-order effects
For strong ageostrophic flows, higher-order effects can contribute to the dynamics of ASCs. In the framework of Eq. (8), these effects are contained within term V [i.e., (11)] and involve advection of the ageostrophic momentum by the ageostrophic flow. Although term V cannot be accurately evaluated from the observations because of errors arising from temporal variability, the Rossby number Roag = Uag/ fL, a measure of this term’s relative importance as a driving mechanism, can be estimated. The Rossby number was calculated using the standard deviation of the ageostrophic velocity on a given section, that is, 
6. Discussion
a. Inertial pumping due to nonlinear Ekman effects
As described in section 3, there are indications that the vertical velocities inferred at the front are associated with near-inertial motions. In this section, we show how time-dependent Ekman flow modified by the vertical vorticity of the front can induce inertial oscillations and inertial pumping, which could be a significant source of vertical motions at the front.
To determine if time-variable nonlinear Ekman transport can drive vertical motions of similar magnitude to those inferred from the inverse method, Mey and we were calculated using observed winds and the vertical vorticity distribution near the front. An example solution derived using the surface vorticity of the geostrophic flow for section 2 of survey 2 is shown in Fig. 8. The Ekman transport has been set to zero at 0000 UTC 23 January. This instance corresponds to a period of weak winds that preceded the onset of a strong cold-air outbreak with wind stress magnitudes exceeding 0.4 N m−2. If the Ekman transport at 0000 UTC 23 January were known, it could have been added as an initial condition. However, as this information was unavailable, it was not included in the solution. Consequently, the solution shown in Fig. 8 should be interpreted as the portion of the Ekman flow driven by the cold-air outbreak that took place after 23 January, not the total wind-driven flow, which could include currents set in motion from wind events that took place prior to 23 January. As evident in the figure, the Ekman transport and pumping tend to oscillate near the effective inertial frequency. Because of the high Rossby number of the frontal jet, the period of the oscillations differs significantly from an inertial period. In particular, in the anticyclonic portion of the frontal jet (−16 km < y < −4 km) the effective inertial period, 2π/ feff, exceeds a day. In this region the Ekman transport and pumping is amplified as well, yielding meridional Ekman flows (estimated as Mey/
b. Vertical advection of tracers
The vertical transport of tracers by submesoscale secondary circulations is of considerable interest to questions involving biogeochemical processes in the upper ocean as well as the subduction of water masses recently modified by atmospheric forcing. At the subpolar front the spatial structure of salinity (which can be considered a passive tracer at the front) and chlorophyll fluorescence (a proxy for phytoplankton biomass) presented evidence of such vertical transport. An example of this can be seen along section 2 of survey 2, Fig. 9. On this section, a strong slantwise frontal downdraft is aligned with a streamer of water with low salinity as well as with a plume of high chlorophyll fluorescence that extends down from the surface. The correlation of the ageostrophic flow, salinity, and chlorophyll fluorescence suggests a causal relation between the inferred vertical velocity and the apparent vertical displacements in the tracer fields. These observations attest to the potential impact of frontal scale vertical motions in the generation of submesoscale bio–optical features and thermohaline intrusions (such as the ones seen at y = 22 km and y = −17 km in Figs. 2 and 9, respectively), both of which are common features of the subpolar front of the Japan/East Sea (e.g., the majority of the sections of the SeaSoar survey contained such features) and ocean fronts in general (e.g., Barth et al. 2001; Fedorov 1983). The vertical extent of the plumes of water with low salinity and high chlorophyll fluorescence suggest that fluid parcels near the front experience vertical displacements of O(50 m). The issue of whether such displacements are consistent with the vertical velocities inferred from the inverse method is discussed below.
In the previous section, it was argued that the frontal vertical circulation is primarily associated with near-inertial motions driven by time-dependent nonlinear Ekman dynamics. An oscillatory vertical velocity of magnitude wo oscillating at a frequency feff will induce a maximum vertical displacement of 2wo / feff in one cycle. The strongest frontal downdrafts are found on the dense side of the front where feff = (1 − 1.5) f and wo ∼ 2 mm s−1. Given these values, the maximum vertical excursion experienced by fluid parcels should not be much greater than 40 m, a distance similar to the vertical extent of the thermohaline intrusions and bio-optical features observed at the front. This suggests that the large magnitude of the inferred vertical velocities and the interpretation that the frontal vertical circulation is associated with near-inertial motions are both plausible.
The inferred vertical velocity could conceivably be used to estimate vertical advective tracer fluxes by correlating tracer and vertical velocity fields from each section. However, given the apparent time variability in the ageostrophic flow, it is likely that such a calculation would be prone to errors. This is because, although the time-dependent vertical motions lead to large vertical displacements in the tracer fields, these displacements are purely reversible to leading order and hence do not lead to a net vertical tracer flux or net subduction. Therefore, without enough cross-front sections sampled at various phases of the oscillatory flow, the reversible portion of the vertical tracer advection may not be adequately averaged out, resulting in errors in the flux calculation. It may also be the case that the oscillatory motions are of sufficient amplitude that they result in a net displacement of fluid parcels via a Stokes drift, the quantification of which will be the topic of future research.2 However, to estimate from observations this Stokes drift and the net tracer flux that it induces, one would again need a sufficient number of temporal realizations of the flow to minimize the errors in the calculation. Given these potential pitfalls and the infrequent sampling of the cross-front sections, a calculation to quantify the flux of the various tracers observed at the front was not attempted.
7. Conclusions
An inverse method was derived for diagnosing vertical velocities from quasi-synoptic, high-resolution surveys of upper-ocean mesoscale/submesoscale flows. The method consists of calculating the Green’s function to an equation governing the vector streamfunction of ASCs and fitting the horizontal component of the velocity of the Green’s function to observationally derived estimates of the ageostrophic flow. In doing so, the method both yields an estimate for the vertical velocity and assesses the driving force for the vertical circulation. The equation that forms the theoretical basis of the method is quite general, with solutions that can describe ageostrophic motions such as nonlinear Ekman flow, mixing-driven ASCs, solutions to the omega equation, and inertia–gravity waves in a geostrophic flow. Owing to its semigeostrophic nature, the equation accounts for higher-order corrections to the quasigeostrophic omega equation, similar to the equations that form the basis of the vertical velocity diagnostics of Shearman et al. (2000), Naveira Garabato et al. (2001), and Pallàs Sanz and Viúdez (2005). Unlike these diagnostics, however, the inverse method is designed to be able to infer vertical velocities driven by mixing, friction, time-dependent inertial/superinertial motions, or nonlinearities in the ageostrophic flow, making it advantageous for application to datasets of upper-ocean flows forced strongly by the atmosphere, such as the high-resolution surveys of the subpolar front of the Japan/East Sea taken during the winter of 2000 in periods of cold-air outbreaks.
At the subpolar front, the inferred vertical velocity was characterized by submesoscale (5–10 km) features with magnitudes of 1–2 mm s−1 (100–200 m day−1). The vertical circulation took on a slantwise nature with strong frontal downdrafts that tended to parallel the sloped isopycnal surfaces of the front. Downdrafts were located on the dense side of the front and were aligned with plumes of water with anomalously low salinity and elevated levels of chlorophyll fluorescence, suggesting that the frontal vertical circulation contributed to the formation of the submesoscale bio-optical features and thermohaline intrusions observed at the front. The strength of the vertical velocity inferred from the inverse method was nearly an order of magnitude larger than that predicted by the omega equation, a consequence of the weakness of the geostrophic forcing relative to the total forcing of the ASC. As evidenced by the clockwise spiraling of the ageostrophic velocity vector with depth, near-inertial waves of relatively small horizontal scale (∼15 km) may have contributed to the lack of applicability of the omega-equation solution. A simple model for time-dependent nonlinear Ekman transport indicates that, given the high Rossby number of the frontal jet and the strength and variability of the wind forcing, nonlinear Ekman transport and inertial pumping are capable of generating inertial oscillations and vertical motions of the amplitude observed and inferred at the front. In addition, scalings for frictional forces and turbulent mixing of buoyancy by Ekman-driven convection, based on shipboard wind measurements, suggest that wind-driven friction is important in the dynamics of the ASC at the subpolar front, whereas mixing of buoyancy is not.
Intensification of the frontal vertical circulation by atmospheric forcing has also been observed in numerical simulations of an idealized representation of the subpolar front forced by winds and/or cooling (Y. Yoshikawa et al. 2008, unpublished manuscript). Comparing experiments with and without surface forcing, Y. Yoshikawa et al. (2008, unpublished manuscript) found that cooling and winds lead to the formation of submesoscale frontal downdrafts, 2–3 km in width, having vertical velocities O(1 mm s−1) that were nearly an order of magnitude stronger than the
A multitude of high-resolution hydrographic and velocity observations as well as state-of-the-art numerical simulations reveal that submesoscale high Rossby number currents are ubiquitous to the upper ocean, especially in frontal zones (Thomas et al. 2008). When forced by winds, such submesoscale features will likely induce time-dependent nonlinear Ekman flows with corresponding inertial oscillations, Ekman pumping, and inertial pumping. The small horizontal spatial scales of the inertial motions generated in this way could trigger downward propagating near-inertial waves that could efficiently transfer inertial energy from the mixed layer to the ocean interior. Indeed, at the subpolar front of the JES there was evidence for enhanced downward-propagating near-inertial energy associated with the submesoscale structure of the vorticity of the frontal jet. It is possible that the interaction of wind-forced inertial motions with the submesoscale surface vorticity field could affect the global budget of inertial energy in the mixed layer and ocean interior, with implications for interior diapycnal mixing—this is a topic that should be explored in the future.
Acknowledgments
We thank Tom Farrar, Terry Joyce, Ken Brink, and Dan Rudnick for their helpful comments and suggestions. We also thank the members of the Woods Hole SeaSoar group, Frank Bahr, Jerry Dean, Paul Fucile, Allan Gordon, and Craig Marquette, whose efforts made the Japan/East Sea observations possible. Captain Christopher Curl and the crew of the R/V Revelle provided professional, enthusiastic assistance through difficult wintertime operating conditions. They were a pleasure to work with, and we are grateful for both their exceptional skill and good nature. L. T. was supported under NSF Grants OCE-03-51191 and OCE-0549699. The Japan/East Sea program was supported by the Office of Naval Research under Grant N00014-98-1-0370 (CML).
REFERENCES
Allen, J. T., and D. A. Smeed, 1996: Potential vorticity and vertical velocity at the Iceland–Faeroes front. J. Phys. Oceanogr., 26 , 2611–2634.
Barth, J. A., T. J. Cowles, and S. D. Pierce, 2001: Mesoscale physical and bio-optical structure of the Antarctic Polar Front near 170°W during austral spring. J. Geophys. Res., 106 , 13879–13902.
D’Asaro, E. A., C. C. Eriksen, M. D. Levine, P. Niiler, C. A. Paulson, and P. V. Meurs, 1995: Upper-ocean inertial currents forced by a strong storm. Part I: Data and comparisons with linear theory. J. Phys. Oceanogr., 25 , 2909–2936.
Dobrindt, U., and J. Schröter, 2003: An adjoint ocean model using finite elements: An application to the South Atlantic. J. Atmos. Oceanic Technol., 20 , 392–407.
Eliassen, A., 1951: Slow thermally or frictionally controlled meridional circulation in a circular vortex. Astrophys. Norv., 5 , 19–60.
Fedorov, K. N., 1983: The Physical Nature and Structure of Oceanic Fronts. Springer-Verlag, 333 pp.
Garrett, C. J. R., and J. W. Loder, 1981: Dynamical aspects of shallow sea fronts. Philos. Trans. Roy. Soc. London, 302A , 563–581.
Giordani, H., L. Prieur, and G. Caniaux, 2006: Advanced insights into sources of vertical velocity in the ocean. Ocean Dyn., 56 , 513–524.
Hoskins, B. J., 1974: The role of potential vorticity in symmetric stability and instability. Quart. J. Roy. Meteor. Soc., 100 , 480–482.
Hoskins, B. J., and I. Draghici, 1977: The forcing of ageostrophic motion according to the semi-geostrophic equations and in an isentropic coordinate model. J. Atmos. Sci., 34 , 1859–1867.
Hoskins, B. J., I. Draghici, and H. C. Davies, 1978: A new look at the ω-equation. Quart. J. Roy. Meteor. Soc., 104 , 31–38.
Hua, B., D. W. Moore, and S. L. Gentil, 1997: Inertial nonlinear equilibration of equatorial flows. J. Fluid Mech., 331 , 345–371.
Kunze, E., 1985: Near-inertial wave propagation in geostrophic shear. J. Phys. Oceanogr., 15 , 544–565.
Leaman, K. D., and T. B. Sanford, 1975: Vertical energy propagation of internal waves: A vector spectral analysis of velocity profiles. J. Geophys. Res., 80 , 1975–1978.
Lee, D. K., and P. P. Niiler, 1998: The inertial chimney: The near-inertial energy drainage from the ocean surface to the deep layer. J. Geophys. Res., 103 , 7579–7591.
Le Traon, P. Y., 1990: A method for optimal analysis of fields with spatially variable mean. J. Geophys. Res., 95 , (C8). 13543–13547.
Nagai, T., A. Tandon, and D. L. Rudnick, 2006: Two-dimensional ageostrophic secondary circulation at ocean fronts due to vertical mixing and large-scale deformation. J. Geophys. Res., 111 , C09038. doi:10.1029/2005JC002964.
Niiler, P., 1969: On the Ekman divergence in an oceanic jet. J. Geophys. Res., 74 , 7048–7052.
Ou, H., 1984: Geostrophic adjustment: A mechanism for frontogenesis? J. Phys. Oceanogr., 14 , 994–1000.
Pallàs Sanz, E., and A. Viúdez, 2005: Diagnosing mesoscale vertical motion from horizontal velocity and density data. J. Phys. Oceanogr., 35 , 1744–1762.
Pollard, R. T., and L. A. Regier, 1992: Vorticity and vertical circulation at an ocean front. J. Phys. Oceanogr., 22 , 609–625.
Press, W. H., B. P. Flannery, S. K. Teukolsky, and W. T. Vetterling, 1993: Numerical Recipes in C. Cambridge University Press, 994 pp.
Rudnick, D. L., 1996: Intensive surveys of the Azores Front. 2. Inferring the geostrophic and vertical velocity fields. J. Geophys. Res., 101 , (C7). 16291–16303.
Rudnick, D. L., and J. R. Luyten, 1996: Intensive surveys of the Azores Front. 1. Tracers and dynamics. J. Geophys. Res., 101 , (C1). 923–939.
Shcherbina, A. Y., L. D. Talley, E. Firing, and P. Hacker, 2003: Near-surface frontal zone trapping and deep upward propagation of internal wave energy in the Japan/East Sea. J. Phys. Oceanogr., 33 , 900–912.
Shearman, R. K., J. A. Barth, and P. M. Kosro, 1999: Diagnosis of the three-dimensional circulation associated with mesoscale motion in the California Current. J. Phys. Oceanogr., 29 , 651–670.
Shearman, R. K., J. A. Barth, J. S. Allen, and R. L. Haney, 2000: Diagnosis of the three-dimensional circulation in mesoscale features with large Rossby number. J. Phys. Oceanogr., 30 , 2687–2709.
Stern, M. E., 1965: Interaction of a uniform wind stress with a geostrophic vortex. Deep-Sea Res., 12 , 355–367.
Takematsu, M., Z. Nagano, A. G. Ostrovski, K. Kim, and Y. Volkov, 1999: Direct measurements of deep currents in the northern Japan Sea. J. Oceanogr., 55 , 207–216.
Tandon, A., and C. Garrett, 1994: Mixed layer restratification due to a horizontal density gradient. J. Phys. Oceanogr., 24 , 1419–1424.
Thomas, L. N., and C. M. Lee, 2005: Intensification of ocean fronts by down-front winds. J. Phys. Oceanogr., 35 , 1086–1102.
Thomas, L. N., A. Tandon, and A. Mahadevan, 2008: Submesoscale processes and dynamics. Ocean Modeling in an Eddying Regime, Geophys. Monogr., Vol. 126, Amer. Geophys. Union, 17–38.
Thompson, L., 2000: Ekman layers and two-dimensional frontogenesis in the upper ocean. J. Geophys. Res., 105 , (C3). 6437–6451.
Viúdez, A., and D. G. Dritschel, 2004: Potential vorticity and the quasigeostrophic and semigeostrophic mesoscale vertical velocity. J. Phys. Oceanogr., 34 , 865–887.
Viúdez, A., J. Tintoré, and R. L. Haney, 1996: Circulation in the Alboran Sea as determined by quasi-synoptic hydrographic observations. Part I: Three-dimensional structure of the two anticyclonic gyres. J. Phys. Oceanogr., 26 , 684–705.
Wimbush, M., and W. Munk, 1970: The benthic boundary layer. New Concepts of Sea Floor Evolution, Part I: General Observations, A. E. Maxwell, E. Bullard, and J. L. Worzel, Eds., The Sea, Vol. 4A, John Wiley and Sons, 731–758.
Wunsch, C., 1996: The Ocean Circulation Inverse Problem. Cambridge University Press, 442 pp.
Zhai, X., R. J. Greatbatch, and C. Eden, 2007: Spreading of near-inertial energy in a 1/12° model of the North Atlantic Ocean. Geophys. Res. Lett., 34 , L10609. doi:10.1029/2007GL029895.
APPENDIX
Errors in the Inverse Solution
Sources of error in the inverse solution can be both systematic and random. In this appendix, an estimate for the magnitude of random errors will be calculated by probing the sensitivity of the inverse solution to noise in the observed fields. In addition, a discussion of the possible sources of systematic errors will be described.
Sensitivity of inverse solution to noise
Following Rudnick (1996) and Naveira Garabato et al. (2001), the sensitivity of the inferred vertical velocity to noise was investigated by adding noise to the density and velocity fields and repeating the inverse calculation. The noise had amplitude 5% of the standard deviation of b, uADCP, and υADCP. The horizontal structure of the noise was prescribed such that the two-dimensional correlation function of the noise field was isotropic, Gaussian, and had a correlation scale of 10 km. For z > −125 m the noise was uniform in the vertical; beneath this level the noise was set to zero. Noise with this spatial structure was meant to represent internal gravity waves that could not be resolved by the ∼20 km zonal spacing of the survey.
The sensitivity test was performed on the section at x = −2 km of survey 4 using 100 realizations of noise and averaging the results. The spatial structure and sense of the vertical velocity and forcing terms p were nearly identical to w and p calculated using the uncorrupted data. The magnitude of the error was at worst 3% for the vertical velocity and 2% for p, indicating that the features of both fields are robust.
Systematic errors
Errors in the estimation of ageostrophic flow
The inverse method relies on the estimate of the horizontal ageostrophic velocity (12) to infer the vertical circulation; therefore, errors in the magnitude and spatial structure of (12) can propagate into the inverse solution. Systematic errors in (12) can arise owing to aliasing of high-frequency variability to lower frequencies in the sampling of velocity and density. This is not so much of an issue for quantities that are derived from along-track, meridional variability, such as ug, uADCP, and υADCP, because, as described in section 4, the data collected in the along-track direction is sampled quickly enough to resolve near-inertial variability. However, zonal density gradients and the meridional geostrophic velocity could be prone to aliasing errors since they are calculated from data taken on separate sections. Hence, it is likely that aliasing errors would be most manifest in the estimate of the meridional component of the ageostrophic flow. Assuming that the ageostrophic flow is nearly isotropicA1 (i.e., |uag| ∼ |υag|) and that the estimate for uag is accurate,A2 then errors in υag can be assessed by comparing the relative magnitudes of the zonal and meridional components of the ageostrophic flow. In Table A1, the standard deviations of uag and υag, σ(uag) and σ(υag), are tabulated for each section as a metric for the magnitudes of the horizontal components of the ageostrophic flow. The ratio σ(υag)/σ(uag) takes values between 0.5 and 1.5 for the majority of the sections suggesting that errors in the estimate of |υag| are no greater than 50% on these sections.
Comparison of the spatial structure of υADCP and υag on these sections reveals that the structure of υag is dominantly determined by υADCP, not υg. The spatial structure of υADCP on each section is unlikely to be affected by aliasing since the velocity data used in the inverse calculation was collected within ∼0.14 of an inertial period (see section 4). We therefore surmise that the spatial structure of υag is accurately estimated and conclude that errors in the inferred vertical velocity owing to errors in υag are less than 50% for most of the sections.
The spatial structure of the ageostrophic circulation and, hence, the inferred vertical circulation can be affected by the along-track resolution of the observations. As stated in section 4, the along-track separation between hydrographic and velocity measurements was ∼3 and ∼0.75 km, respectively. To test the sensitivity of the inverse solution to the resolution of the observations, the measurements were subsampled and the inverse calculation repeated. Reducing the resolution of the velocity measurements to 3 km resulted in only a slight (less than 10%) modification of the inferred vertical velocity. However, decreasing the resolution of both hydrographic and velocity observations to 7.5 km was found to reduce the magnitude of w by ∼50%. Clearly the best observational strategy for the inverse method is to collect measurements at as high a spatial resolution as possible while sampling rapidly so as to avoid aliasing of inertial motions.
Boundary conditions
Setting the vertical velocity at the bottom and side boundaries to a constant value of zero is an arbitrary choice that could lead to errors in the solution. If we consider that the vertical velocity obtained from the inverse calculation w is equal to the actual vertical velocity plus an error w = wact + werr, then at the boundary w = 0 the inverse solution will have errors as large as wact. The errors attributed to the w = 0 boundary condition are not confined to the boundary but are felt over a finite distance into the interior of the domain. To assess how large this area of error is, (8) was solved with the rhs forcing terms set to zero and a nonzero vertical velocity meant to represent werr specified at the lateral and bottom boundaries separately. As demonstrated in Fig. A1, the area of error depends on the spatial structure of the vertical velocity specified on the boundary. When the vertical velocity on the lateral boundaries extends deeper into the fluid, the area of error reaches farther into the interior of the domain (cf. Figs. A1a and A1b). Specifying a vertical velocity on the bottom boundary with a larger horizontal scale yields an area of error that extends higher up into the water column (cf. Figs. A1c,d).
Given the 5–10-km horizontal and ∼100-m vertical length scales of the inferred w, this sensitivity test indicates that errors due to the boundary conditions should be less than 10% outside of a region 10 km from the lateral boundaries and above z = −150 m, indicating that the solutions for the vertical circulation in the proximity of the front and above and within the pycnocline are robust.
Objective map of the potential density at z = −18 m (contours) and the ship track (thick gray) for survey 2 of the subpolar front of the JES taken on 26–29 Jan 2000. Section numbers are to the left of the ship track. The black line denotes the truncated domain used for the inverse calculation along section 2. Contour interval for the density is 0.1 kg m−3.
Citation: Journal of Physical Oceanography 40, 1; 10.1175/2009JPO4018.1
Evidence of frontal downdrafts in the tracer fields: (a) salinity and (b) chlorophyll fluorescence for section 5 of survey 2, reveal 5–10-km-wide plumes with high chlorophyll fluorescence and low salinity that extend ∼50 m beneath the mixed layer on the dense side of the front. Isopycnals are contoured in intervals of 0.1 kg m−3.
Citation: Journal of Physical Oceanography 40, 1; 10.1175/2009JPO4018.1
Scatterplots of the vertical velocity for the three-dimensional 
Citation: Journal of Physical Oceanography 40, 1; 10.1175/2009JPO4018.1
The inferred vertical velocity at the subpolar front of the JES. Isopycnals are contoured with an interval 0.1 kg m−3. The 26.5 kg m−3 isopycnal is highlighted in magenta.
Citation: Journal of Physical Oceanography 40, 1; 10.1175/2009JPO4018.1
Vertical velocity inferred from solving the omega equation. Note the reduced scale for w compared to Fig. 4. Isopycnals are contoured with an contour interval 0.1 kg m−3. The 26.5 kg m−3 isopycnal is highlighted in magenta.
Citation: Journal of Physical Oceanography 40, 1; 10.1175/2009JPO4018.1
Example of the slantwise nature of the ageostrophic flow near the front: the ageostrophic flow (υag, w) (vectors with dots indicating their base), isopycnal surfaces (gray solid contours), and surfaces of constant geostrophic absolute momentum Mg (gray dashed contours), for section 2 of survey 2.
Citation: Journal of Physical Oceanography 40, 1; 10.1175/2009JPO4018.1
Relation between velocities having properties of near-inertial motions and the structure of the rhs forcing terms of (8), (a) RHS2 and (b) RHS1, for section 2 of survey 2. (c) Hodograph of the ageostrophic velocity at y = −9.6 km, with colors indicating the depth. (d) The total energy of the clockwise (red) and counterclockwise (blue) components of the ageostrophic flow as a function of the cross-front distance. (e) The vertical vorticity of the geostrophic flow normalized by f, with the location of the profile used to make the hodograph marked with the dashed line.
Citation: Journal of Physical Oceanography 40, 1; 10.1175/2009JPO4018.1
Solution for the time-dependent Ekman transport and the vertical velocity that it induces for section 2 of survey 2, forced using observed winds. Hovmöller diagrams of (a) Mey normalized by the time-mean Ekman depth 
Citation: Journal of Physical Oceanography 40, 1; 10.1175/2009JPO4018.1
Indications of vertical advection of tracer fields. Vectors of the ageostrophic velocity with the (a) salinity and (b) chlorophyll fluorescence for section 2 of survey 2. Dots indicate the base of the velocity vector; isopycnals contoured in intervals of 0.1 kg m−3.
Citation: Journal of Physical Oceanography 40, 1; 10.1175/2009JPO4018.1
Fig. A1. Area of error due to lateral and bottom boundary conditions. The magnitude of the vertical velocity obtained solving (8) for survey 4 in the domain centered at x = −2 km with the rhs forcing terms set to zero and w specified as w = wbc at the boundaries. (a),(b) |w| with boundary conditions wbc = wo = const for z ≥ −H and wbc = 0 for z < −H with (a) H = 300 m and (b) H = 150 m. (c),(d) |w| with boundary conditions w = wo sin[2π(y − yo)/Ly] for −Ly /2 < y − yo < Ly /2 and w = 0 for y − yo < −Ly /2 and y − yo > Ly /2 with yo = −16.5 km and (c) Ly = 30 km and (d) Ly = 15 km. Magnitude of the vertical velocity has been normalized by wo; contour interval 0.1.
Citation: Journal of Physical Oceanography 40, 1; 10.1175/2009JPO4018.1
Comparison between the rhs forcing term of (8) and the geostrophic forcing, where r(RHS2,
Scalings for the contributions of friction and mixing of buoyancy in the rhs forcing terms of (8). The scalings are defined in (29), (30), and (32) and have been normalized by the standard deviations of the inferred rhs forcing [i.e., σ(RHS2) and σ(RHS1)].
Table A1. Comparison of the relative magnitudes of the estimated zonal and meridional components of the ageostrophic flow. The standard deviations of uag and υag for each section, σ(uag) and σ(υag), are tabulated along with their ratio.
The forcing terms on the right-hand side of (8) are included in the “generalized” Q vector of Giordani et al. (2006). The generalized Q vector also includes the terms on the left-hand side of (8) that would not be there if the quasigeostrophic approximation were used.
It appears that the amplitude of the vertical velocity associated with inertial pumping can result in displacements large enough to induce a significant Stokes drift 
A1 The observations suggest that the ageostrophic flow is primarily attributable to near-inertial motions that are isotropic (see sections 3 and 6).
A2 The accuracy of uag is limited by errors in the estimate for ug. A measure of the accuracy of ug is the closeness to which the thermal wind relation models the observed vertical shear in the zonal flow. Following Rudnick and Luyten (1996), this measure was quantified by calculating the regression coefficient A = −〈f (∂uADCP/∂z)(∂b/∂y)〉/〈(∂b/∂y )2〉 (the brackets denote an average over volume for a given survey, excluding grid points where the ratio of error to signal variance exceeded 0.3), which would be equal to one if the zonal flow satisfied the thermal wind relation exactly. The regression coefficients for surveys 2, 3, and 4 are 0.93, 0.79, and 0.74, respectively. These relatively high values of A suggest that the zonal geostrophic flow was captured accurately in the surveys.