1. Introduction
The Southern Ocean plays a central role in the climate system, connecting the major ocean basins and allowing a global overturning circulation to exist (Rintoul 2000; Rintoul et al. 2001). The Antarctic Circumpolar Current (ACC) flows from west to east around the Antarctic continent and consists of three major fronts, from north to south—the Subantarctic Front (SAF), the Polar Front (PF), and the Southern ACC Front (SACCF) (Klinck and Nowlin 2001; Olbers et al. 2004). The fronts are often aligned with particular circumpolar sea surface height (SSH) streamlines (Sokolov and Rintoul 2009) and are associated with strong zonal jets that regulate poleward heat transport (Naveira Garabato et al. 2011). Large-scale upwelling of deep waters from the ocean interior is facilitated by the steeply sloping isopycnals across the ACC (Rintoul et al. 2001; Olbers et al. 2004; Thompson 2008; Tamsitt et al. 2017). This large-scale upwelling and ventilation of the deep ocean is fundamentally important for heat and carbon uptake in the Southern Ocean, where ∼70% of global anthropogenic heat (Frölicher et al. 2015) and ∼40% of anthropogenic carbon (Caldeira and Duffy 2000; Le Quéré et al. 2007; Khatiwala et al. 2009) is absorbed into the ocean.
Southern Ocean circulation and dynamics vary around the Antarctic continent, particularly in regions where the ACC interacts with major topographic features. Flow–topography interactions generate standing meanders and rich eddy fields that are hotspots of poleward heat transport (Phillips and Rintoul 2000; Foppert et al. 2017), large-scale upwelling (Tamsitt et al. 2017; Rintoul 2018), biological activity (Siegelman et al. 2019), carbon sequestration (Sallée et al. 2012; Langlais et al. 2017), and ventilation of the ocean interior (Dove et al. 2022). Standing meanders form due to the westward propagation of topographically generated Rossby waves that balance the eastward mean flow of the ACC (Hughes 2005; Thompson and Naveira Garabato 2014; Meijer et al. 2022; Zhang et al. 2023a). Complex eddy–mean flow interactions within the meander generate barotropic and baroclinic instabilities, resulting in rapid eddy kinetic energy (EKE) generation (MacCready and Rhines 2001; Youngs et al. 2017; Foppert 2019). Vertical motion at the mesoscale is tied to the phase of the meander (Phillips and Bindoff 2014), driven by the ageostrophic component of the gradient-wind flow field resulting from the meander curvature (Meijer et al. 2022). At smaller temporal and spatial scales, submesoscale dynamics in energetic meander regions can drive localized subduction and tracer exchange between the mixed layer and the ocean interior (Rosso et al. 2014, 2015; Llort et al. 2018; Balwada et al. 2018; Taylor et al. 2018; Freilich and Mahadevan 2021; Dove et al. 2021; Morrison et al. 2022).
Submesoscale motions of O(10) km are distinct from mesoscale motions of O(100) km in that they are not as strongly constrained by Earth’s rotation—with a Rossby number close to one, they have a strong ageostrophic component and can develop much stronger vertical velocities in comparison to mesoscale motions (Thomas et al. 2008; Gula et al. 2022; Taylor and Thompson 2023). In addition to their relatively small spatial scales, submesoscale features evolve on time scales of hours to days, making them extremely challenging to observe (Siegelman et al. 2020a; Archer et al. 2020). Mesoscale fronts provide an ideal environment for submesoscale processes to occur due to the strong horizontal density gradients that provide a source of potential and mean kinetic energy to fuel submesoscale instabilities (Yu et al. 2019). The rich eddy fields in energetic meander regions act to stir the large-scale density gradient and generate finescale fronts and filaments (Mahadevan and Tandon 2006; Capet et al. 2008; Gula et al. 2022). The strain field between two counterrotating eddies in an eddy dipole structure can also generate submesoscale fronts and instabilities that drive localized subduction (Klein and Lapeyre 2009; Archer et al. 2020).
Despite their relatively small scale, submesoscale processes can influence components of the large-scale ocean circulation and Earth’s climate (Su et al. 2018; Taylor and Thompson 2023; Hewitt et al. 2022; Swart et al. 2023). They provide an important pathway for the cascade of energy and tracer variance from large scales to dissipative scales (D’Asaro et al. 2011; Gula et al. 2022) and are therefore essential to understanding water mass variability and ocean mixing, as well as localized subduction and ventilation. Submesoscale processes are particularly prevalent in the upper ocean and play a key role in the vertical exchange of heat, nutrients, and carbon across the base of the mixed layer—influencing primary productivity and the carbon cycle (Lapeyre and Klein 2006; Thomas et al. 2008; Mahadevan 2016; Lévy et al. 2018; Archer et al. 2020; Siegelman et al. 2020a; Su et al. 2020). Submesoscale dynamics can alter the depth of the surface mixed layer (Lapeyre et al. 2006; du Pleiss et al. 2017; Bachman et al. 2017; du Pleiss et al. 2019; Bachman and Klocker 2020), controlling the connectivity between the atmosphere and the ocean interior, and thus influencing the strength of the global overturning circulation (Fox-Kemper et al. 2011; Swart et al. 2023).
A key process that alters the upper-ocean density structure and connectivity to the ocean interior is frontogenesis. Frontogenesis is well described in the atmospheric context (Hoskins 1982) and refers to the sharpening of horizontal density gradients, driven by the mesoscale strain field, generating an along-front acceleration (McWilliams et al. 2009; McWilliams 2021). During frontogenesis (Fig. 1a), the front can transition from a state of thermal wind balance to turbulent thermal wind balance, where turbulent mixing becomes a key component in the momentum balance (Hoskins and Bretherton 1972; Gula et al. 2014; McWilliams et al. 2015). Turbulent mixing reduces the along-front vertical shear and leads to an unbalanced cross-front pressure gradient. This drives a cross-front ageostrophic secondary circulation (ASC) that acts to flatten isopycnals and restore the flow back to hydrostatic and thermal wind balance (McWilliams et al. 2009; Archer et al. 2020; Gula et al. 2022). This single-cell circulation results in downwelling on the dense side of the front and upwelling on the light side (Fig. 1a). The ASC tends to be strongest where the magnitude of the horizontal density gradient peaks (Gula et al. 2022). Vertical motion associated with an ASC can influence vertical nutrient fluxes and the residence time of phytoplankton in the mixed layer (Lévy et al. 2018), with consequences for primary production and marine ecosystems.
Schematic of Southern Hemisphere (a) frontogenesis and (b),(c) filamentogenesis associated with cold and warm filaments, respectively. Blue arrows indicate the ageostrophic secondary circulation (ASC). Adapted from McWilliams et al. (2009) and Gula et al. (2022).
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
Filamentogenesis is similar to frontogenesis but, instead of a single cross-front circulation cell, there are two circulation cells—one on each side of the filament (Figs. 1b,c). In the case of a cold (i.e., dense) filament (Fig. 1b), ageostrophic convergence occurs at the surface in the core of the filament, aligned with the confluence axis associated with the deformation flow (McWilliams et al. 2009). The filament narrows and there is an intensification of downward velocities in the filament, with weaker upwelling motions on either side. This process is known as cold filamentary intensification (McWilliams et al. 2009; Gula et al. 2014). While warm filamentary intensification can also occur (Fig. 1c) (Lapeyre and Klein 2006), the surface divergence associated with the ASC acts to broaden the filament and opposes the confluent flow—leading to upwelling that has a slower intensification rate than the downwelling associated with a cold filament (McWilliams et al. 2009).
During frontogenesis and filamentogenesis, submesoscale instabilities can develop as they feed off the local potential and kinetic energy of the intensifying front (Boccaletti et al. 2007; Capet et al. 2008; Fox-Kemper and Ferrari 2008; Gula et al. 2014; Archer et al. 2020). Submesoscale instabilities can prevent further frontal sharpening (i.e., frontal arrest) (McWilliams and Molemaker 2011), enhance vertical motion (Thomas et al. 2013; Callies et al. 2016; Yu et al. 2019), restratify the upper ocean (Boccaletti et al. 2007; Fox-Kemper and Ferrari 2008; Yu et al. 2019), drive localized subduction (Thomas et al. 2013; Freilich and Mahadevan 2021), and facilitate a dynamical route to molecular diffusion by extracting energy from the geostrophically balanced flow (D’Asaro et al. 2011; Gula et al. 2022). Different types of submesoscale instabilities can be distinguished by their energetics, using indicators such as potential vorticity (PV), the buoyancy frequency (N2), and the Richardson number (Ri) (Taylor and Thompson 2023; Gula et al. 2022).
While submesoscale fronts and instabilities are often confined to the mixed layer, ageostrophic motion associated with frontogenesis can extend down to 400–900 m in the weakly stratified Southern Ocean (Siegelman 2020; Siegelman et al. 2020a). Submesoscale processes can also extend deeper in the water column in the presence of steep topography (Rosso et al. 2014; Gula et al. 2022), in strain regions on the periphery of mesoscale eddies (Yu et al. 2019; Siegelman et al. 2020b; Dove et al. 2021), and during wintertime when stratification at the base of the mixed layer is weak (Thompson et al. 2016; Erickson and Thompson 2018; Yu et al. 2019).
Many studies over the past decade discuss the importance of submesoscale dynamics in the climate system; however, there is a lack of finescale observations to test current theory and understanding, particularly in the Southern Ocean. We present a unique set of finescale observations from an energetic meander region of the ACC, south of Tasmania, that supports the theory of cold filamentary intensification. The meander forms between two topographic features—downstream of the Southeast Indian Ridge (SEIR) and just upstream of Macquarie Ridge (MR) (Fig. 2). Zhang et al. (2023a) provide a detailed analysis of the dynamics governing the meander formation in this region. This particular meander has been identified in previous studies as a hotspot of EKE (Thompson and Naveira Garabato 2014), cross-frontal exchange (Thompson and Sallée 2012), poleward eddy heat flux (Foppert et al. 2017) and turbulent mixing (Cyriac et al. 2022). The observations we present here were collected on a voyage on the R/V Investigator in October 2018, and highlight the role of finescale filaments in generating strong vertical motions, intensifying vertical exchange between the mixed layer and the ocean interior, and transporting water mass properties across mesoscale fronts. We find enhanced vertical velocities and evidence of ageostrophic motion extending as deep as 1600 dbar—well below the base of even the wintertime mixed layer. We also find evidence of localized subduction and ventilation associated with a finescale cold filament. The vertical extent of these dynamics has important consequences for ventilation of the ocean interior, vertical heat and nutrient fluxes, interior stratification, and carbon cycling in the Southern Ocean.
Map of the study domain. Trajectories of three EM-APEX floats during rapid sampling are indicated by the colored circles. Triaxus tows are shown as black solid lines. Background color indicates ocean depth (m), highlighting Macquarie Ridge (MR) near 160°E, Campbell Plateau (CP) to the northeast, and part of the Southeast Indian Ridge (SEIR) to the west. White contours show SSH in 0.1-m intervals from −0.8 to 0.2 m, averaged over the float rapid sampling period (21 Oct–5 Dec 2018). The ACC standing meander forms between the SEIR and MR, with the crest (trough) being the poleward (equatorward) excursion of the meander. Black contours show 20-yr (1998–2018) average SSH contours corresponding to the mean position of the PF (−0.65 m; solid), SAF south (−0.4 m; dashed), and SAF north (0.2 m; dash–dot), as in Cyriac et al. (2022).
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
2. Data and methods
We draw together in situ observations from EM-APEX floats, a towed Triaxus and the shipboard acoustic Doppler current profiler (ADCP), as well as satellite observations of SSH and high-resolution sea surface temperature (SST). Details about these data sources, processing steps and derived variables are detailed in their respective subsections.
a. EM-APEX floats
Electromagnetic autonomous profiling explorer (EM-APEX) floats are autonomous profiling floats that follow the current while profiling up and down through the water column, gathering high-resolution vertical profiles of temperature, salinity, pressure, and horizontal velocity. The floats behave similar to an Argo float but profile more frequently, gathering data at higher temporal and spatial resolution. In addition to a CTD (conductivity–temperature–depth) sensor, the EM-APEX floats are fitted with an electromagnetic (EM) subsystem, containing two pairs of electrodes that measure the electric current generated by the motion of seawater through Earth’s magnetic field (Sanford et al. 2005). External fins rotate the float as it ascends and descends, to allow processing that removes the electric self-potential between electrode pairs and leaves the potential difference induced by the ocean currents (Sanford 1971; Phillips and Bindoff 2014). The electric current voltages are converted into velocity components, relative to a depth-independent reference velocity. Relative velocities are then converted to absolute velocities by adding a velocity offset, so that the theoretical resurface position after each up-profile matches the actual surface position from the GPS (see appendix).
During the R/V Investigator voyage (16 October–15 November 2018), six EM-APEX floats were deployed in the standing meander of the ACC upstream of Macquarie Ridge. Two of the six floats failed to collect any data and one failed to retrieve velocity information. In this study, we use data from the three floats that were fully functional—containing temperature, salinity, pressure, and velocity information. The trajectories of these floats are shown in Fig. 2.
The floats were programmed to profile to 1600 dbar, with a vertical resolution of 3–4 dbar, and complete one profiling cycle per day that consists of four vertical profiles, and a drift at 1000 dbar (Fig. 3a). In this analysis, we only use data during the floats’ ascent or descent; data collected during the floats’ drift has been removed. This pattern was designed to resolve velocities close to the inertial frequency (Phillips and Bindoff 2014)—with consecutive down-profiles and consecutive up-profiles separated by approximately half an inertial period (π/f = 7.29 h) at the latitude of float deployment (see appendix for further information). In this study, we only use data from the floats’ rapid sampling period (21 October–5 December 2018), providing the highest temporal and spatial resolution. The lateral distance between consecutive profiles (based on surface GPS positions) during rapid sampling has a mean of ∼3.5 km, with individual separation distances ranging from <1 to ∼15 km depending on the speed of the background current (Fig. 3b).
(a) EM-APEX float profiling pattern showing two daily cycles that each produce four vertical profiles from 0 to 1600 dbar. The drift at 1000 dbar has been removed for this analysis. (b) Histogram of separation distance between consecutive profiles in 1-km bins. Depth-averaged speed for the profiles in each 1-km bin is indicated by the black line and corresponds to the RHS y axis.
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
The data from each float went through a routine quality control procedure (Phillips and Bindoff 2014; Cyriac et al. 2021). This involves a pressure drift correction, where the surface pressure value is subtracted from all pressure values in that profile, to reset the surface pressure to zero. Temperature and salinity profiles below the seasonal thermocline (200 dbar) are compared with the CSIRO Atlas of Regional Seas (CARS) (Ridgway et al. 2002) and satellite gravest empirical mode (SatGEM) (Meijers et al. 2011) climatologies to identify erroneous data and spikes that are then manually removed. Velocity spikes are identified using a depth-dependent cutoff based on the RMS error of velocity, ranging from 0.5 cm s−1 below 900 dbar to 1.5 cm s−1 between 100 and 220 dbar.1 Temperature, salinity, and velocity profiles are vertically interpolated to an even grid of 2 dbar. Following the initial QC, 3 (out of 1072) profiles with discontinuities in the temperature and salinity were identified and manually removed. Similarly, several profiles of relative velocities were manually removed after the initial QC due to large inconsistencies with surrounding profiles. The white vertical lines in the float sections are the erroneous profiles with data removed.
With consecutive down-profiles and consecutive up-profiles being approximately half an inertial period apart, we remove inertial oscillations from the float data by averaging half-inertial pairs and interpolating back onto the original time grid. Absolute velocities are rotated to along and cross trajectory as a proxy for along and cross streamflow. Further details on the EM-APEX velocity processing pipeline and half-inertial pair averaging can be found in the appendix.
Derived quantities
Potential density referenced to the surface (ρ), gravitational acceleration g, and the square of the buoyancy frequency [
The Richardson number, Ri, gives an indication of the presence of ageostrophic dynamics and the susceptibility of the flow to shear-driven overturning (Miles 1961; Thomas et al. 2008). It can be calculated as the ratio of the vertical stratification to the vertical shear,
b. Triaxus and ADCP
In addition to the EM-APEX float deployment, eight high-resolution Triaxus tows to a depth of 300 dbar were conducted across the ACC meander during the R/V Investigator voyage (Fig. 2). The Triaxus is an undulating CTD system that is towed behind the ship. It was equipped with a dual-sensor Seabird SBE911 CTD unit as well as sensors to measure dissolved oxygen (SBE43), nitrate, and chlorophyll (ECO Triplet). Quality control and processing was undertaken by CSIRO Marine National Facility (MNF). This included spike removal, identification of water entry and exit times, conductivity sensor lag corrections and determination of pressure offsets. Data were gridded in 1-dbar vertical bins, using a linear least squares fit as a function of pressure to interpolate the value for the bin midpoint. Vertical casts were created from the vertically gridded data, using linear interpolation to a maximum of two casts’ distance. An estimate of the Triaxus’ average position for each vertical cast was calculated using the wire out, pressure, and ship location. Further information can be found in the processing reports on the MNF website (marine.csiro.au/data/trawler/survey_details.cfm?survey=IN2018_V05). We plot the Triaxus data against along-track distance, using the estimated Triaxus positions for each vertical cast. The average lateral distance between vertical casts is 1.16 ± 0.18 km.
In this study, we use the fourth Triaxus tow that crosses the ACC meander at 54.5°S, between the crest (poleward excursion of the meander) and trough (equatorward excursion). This transect, occupied from west to east, is roughly perpendicular to SSH streamlines and crosses a finescale cold filament that flows northward near 152.2°E, in between an eddy dipole structure (Fig. 4a). The cold filament measured by the Triaxus appears in almost the same location as the larger cold filament sampled by the floats a few days later. With the floats roughly traveling along-stream and the Triaxus being almost perpendicular to the filament, we are able to extract both an along-stream and cross-stream perspective of the cold filaments generated in this region.
(a) Sea surface temperature (SST) map on 29 Oct 2018—the day of the Triaxus tow (black line). The locations of the EM-APEX floats on this day are shown as blue circles. (b) Triaxus tow in relation to the EM-APEX float trajectories (blue lines). Thin gray contours show SSH on the same day (29 Oct 2018), in intervals of 0.1 m from −0.8 to 0.4 m. Anticyclonic and cyclonic eddies associated with the eddy dipole are marked as AE and CE, respectively. The thick gray line underneath the Triaxus transect in each panel is the ship track associated with the ADCP measurements.
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
Vertical profiles of absolute velocity from near-surface (16 dbar) to ∼400 dbar were obtained from the 150-kHz shipboard ADCP mounted on the ship’s hull. The velocities are gridded in 8-dbar vertical bins and 5-min time intervals. Further details on the ADCP processing can be found in the voyage reports on the MNF website. We use the ADCP velocities corresponding to the start and end times of the Triaxus transect, minus 10 min to account for the spatial offset between the ship and the towed Triaxus. A 10-min lag was found to closely align the GPS positions of the ship (and ADCP measurements) with the estimated positions of the towed Triaxus. For consistency with the Triaxus measurements, we only show velocities for the upper 300 dbar of the water column.
Zonal (u) and meridional (υ) velocity components from the ADCP are rotated to along and cross track to avoid the ambiguity of a cross-front/along-front framework in a region rich in finescale features.3 The edges of the Triaxus transect are associated with the anticyclonic and cyclonic eddies of an eddy dipole structure (Fig. 4b), with flow to the northeast on the western side, and to the northwest on the eastern side. At the center of the transect, SSH contours and the SST filament intersect the track roughly perpendicularly. Therefore, in the context of the filament, the rotation to along track is an appropriate proxy for cross front.
Derived quantities
Potential density, N2, and MLD along the Triaxus transect are calculated as described above for the float profiles. The Richardson number is calculated using the gridded ADCP velocities, interpolated onto the N2 pressure and distance grid. Apparent oxygen utilization (AOU) is calculated as the difference between the oxygen solubility of a water parcel (dependent on temperature and salinity) and its measured oxygen concentration (
c. Satellite products
Daily estimates of SSH, provided as absolute dynamic topography, and derived surface geostrophic velocities in Cartesian coordinates were obtained from the 0.25° delayed-time L4-gridded satellite altimeter product, provided by Copernicus. To gain insight into the spatial distribution of submesoscale fronts and filaments in the meander region, we use the 0.02° daily L3-gridded multisensor sea surface temperature (SST) product provided by Australia’s Integrated Marine Observing System (IMOS). We remove data with a quality level flag of less than 3. Further information on the quality control of the SST dataset is described in Griffin et al.’s (2017) appendix A (imos.org.au/facilities/srs/sstproducts/sstdata0/sstdata-references/). Due to the presence of clouds, we take a temporal average over 2–3 days to capture the approximate location of the filaments and improve spatial coverage.
Surface EKE is calculated as
Daily finite-sized Lyapunov exponents (FSLEs), computed backward-in-time from Copernicus L4-gridded geostrophic velocities, were obtained from the 0.04° gridded Aviso+ data product. FSLEs are calculated by measuring the separation rate of pairs of particles in a given three-dimensional velocity field, with fixed initial and final separation distances (Aurell et al. 1996; Artale et al. 1997; Boffetta et al. 2001). They provide a direct measure of local stirring and can be used to characterize the strain field and identify finescale structures of the flow, otherwise called Lagrangian coherent structures (D’Ovidio et al. 2004; Hernández-Carrasco et al. 2012; Cotté et al. 2015). FSLEs computed backward-in-time yield negative FSLE values that provide a measure of horizontal confluence—strongly negative values indicate regions where particles that were initially far apart have converged rapidly. The opposite is true for FSLEs computed forward-in-time, where strongly positive values indicate a high separation rate of particles (i.e., strong stretching) (D’Ovidio et al. 2004). Containing both spatial and temporal information from satellite altimetry, FSLEs can provide information about the growth rate and orientation of submesoscale fronts (Siegelman et al. 2020a). Large absolute FSLE values are often observed around the periphery of mesoscale eddies, associated with submesoscale fronts (Siegelman et al. 2020a), and in high EKE regions in the Southern Ocean (Dove et al. 2022).
3. Results
a. Surface characteristics of the study region
During the float sampling, there was a highly energetic eddy field within the meander, indicated by strong surface EKE just upstream of MR (Fig. 5a), as well as the highly contorted float trajectories (Fig. 2). EKE weakens where the flow encounters MR, as the time-mean current travels through two main gaps in the ridge at 53° and 56°S. The rich mesoscale eddy field in the meander resulted in strong surface FSLE values in the region (Fig. 5b), implying high submesoscale activity. The strongest EKE and FSLE signals within the red box (where the observations were taken) are located within the meander trough at 152.5°E and 53.5°S, where there was a persistent cyclonic eddy with a strong negative Okubo–Weiss parameter (Fig. 5c). The OW parameter (Fig. 5c) also highlights the eddy dipole present in the meander during the float sampling—with the two vorticity-dominated eddies (negative OW), at 150° and 152.5°E, separated by a strain-dominated region (positive OW).
(a) Eddy kinetic energy (EKE), (b) FSLE, and (c) Okubo–Weiss parameter, calculated from satellite geostrophic velocities and averaged over the float rapid sampling period (21 Oct–5 Dec 2018). The thick gray contour in each panel shows the 1500-m depth level of the bathymetry, highlighting Macquarie Ridge and Campbell Plateau. SSH streamlines averaged during the float sampling period are shown as thin contours; black contours show the approximate location of the PF (solid) and southern (dashed) and northern (dash–dot) branches of the SAF, averaged over the same period. The red box marks the map area in Figs. 4 and 6, where the observations were taken.
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
In the high-resolution SST images over the float sampling period, a ∼20-km-wide (Fig. 6b) cold filament was observed (Fig. 6a) and sampled by the floats in various locations (Figs. 6c–e). The Rossby radius of deformation at this latitude is ∼10–20 km (Chelton et al. 1998), indicating that the filament is on the upper end of the submesoscale range. Due to the presence of clouds, the snapshots in Fig. 6 have been averaged over 3 days, introducing some smoothing, and show the approximate location of the filament in relation to the EM-APEX floats.
(a) Satellite sea surface temperature (SST) snapshot in the meander region, averaged over three days (4–6 Nov 2018). The 4.4°C isotherm (black line) highlights the cold filament and two cold-core eddies east of the filament. (b) SST along the dashed line in (a). Shading represents the bounds marked by the 4.4°C isotherm. (c)–(e) The same 3-day SST snapshot overlaid with each float trajectory (white circles), the location of the float during the 3 days (black dots) and the 3-day mean SSH contours in 0.15-m intervals from −0.6 to 0.3 m. Along-trajectory distance at the float locations corresponding to the black dots are displayed at the top left of each panel.
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
The filament originates from south of the front, travels northward at 151.5°E, then turns eastward and appears to feed into the cyclonic eddy at 154.5°E (Fig. 6a). The filament forms in between the eddy dipole structure, where you would expect the mesoscale strain field to drive filamentogenesis (Lapeyre and Klein 2006; McWilliams et al. 2009). Float EM-8492 travels northward in the filament and then crosses over on the western side (Fig. 6d), while EM-8493 follows the filament into the cyclonic eddy (Fig. 6e) and EM-8489 travels out of the filament to the east (Fig. 6c)—making a clockwise loop before traveling northward. The filament is completely eroded a few days later when float EM-8489 travels north (not shown). It is important to note that the EM-APEX floats are not fully Lagrangian because of their profiling—they occasionally cross SSH streamlines and do not necessarily follow the filament either.
b. Subsurface structure along the EM-APEX float trajectories
The EM-APEX floats were deployed upstream of the meander crest, close to the mean position of the PF based on the SSH gradient and the position of the −0.65-m SSH contour on the day of float deployment (Cyriac et al. 2022). For the first ∼200 km, the floats sample a temperature minimum (Tmin) layer of <2°C between 200- and 300-m depth (Fig. 7a) —a characteristic signature of Antarctic Winter Water (WW) at (or just south of) the PF (Orsi et al. 1995; Belkin and Gordon 1996; Klinck and Nowlin 2001; Naveira Garabato et al. 2001). The floats then diverge into different trajectories, steered by the local currents in the rich EKE and FSLE field. Floats EM-8489 and EM-8492 transition into warmer waters north of the PF (Fig. 7a), with a deepening of low-salinity waters (Fig. 7b) indicative of the floats approaching the SAF (Kim and Orsi 2014). Float EM-8493 travels into a cyclonic (cold-core) eddy within the meander trough (Fig. 6e)—where surface waters are ∼3°C warmer than at the beginning of the float trajectory but the WW signature is retained between 200 and 300 dbar (Fig. 7a).
Subsurface sections of (a) Conservative Temperature and (b) Absolute Salinity in the upper 1000 dbar along each of the float tracks. Inertial oscillations have been removed by half-inertial pair averaging, and the data are interpolated onto a regular 3-km distance grid. The thick white contour on (a) is the 2°C isotherm and the black bars at 0 dbar indicate the locations of the floats in Figs. 6c–e. The thin black line marks the mixed layer depth (MLD) and thin white contours show potential density in 0.2 kg m−3 increments.
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
There are a range of scales (from submesoscale to mesoscale) captured in the float data. Between 200 and 300 km in all floats, there is a deepening of the WW layer (to 300–500 dbar) as the water mass subducts along isopycnals (Fig. 7a). This deepening of Tmin waters occurs as the floats travel northward (Figs. 6c–e) and may be a signal of the floats crossing the PF. However, the floats are also crossing submesoscale fronts and sampling the cold filament that flows from the meander crest to the trough (Figs. 6c–e). Float EM-8492 crosses the western side of the cold filament (Fig. 6d) at the location of WW subduction between 200 and 300 km (Fig. 7a), while float EM-8493 continues to follow the cold filament into the cyclonic eddy (Fig. 6e). Float EM-8489 does not travel northward in the cold filament like the other floats, but crosses the eastern side of the filament at its base near 54.8°S.
The subsurface velocity structure along the float tracks helps to distinguish between mesoscale and submesoscale dynamics. A striking feature in the subsurface velocities is the reversal in sign of the cross-track velocities with depth (Fig. 8b). This overturning signal is particularly prominent where the along-track velocities are strongest (Fig. 8a)—between 200 and 450 km in EM-8492 and between 200 and 350 km in EM-8493—and is associated with intense downward velocities of O(10−3) m s−1 extending from the surface to 1600 dbar (Fig. 8c). The inferred vertical velocities are greater than 100 m day−1—an order of magnitude larger than expected from mesoscale motions. This pattern in the velocity structure is coincident with the descent of Tmin waters (Fig. 7a) and is localized where floats EM-8489 and EM-8492 sample the cold filament traveling northward, in between the eddy dipole structure (Figs. 6d,e). We interpret these observations as evidence of cold filamentary intensification. The overturning signal in floats EM-8489 and EM-8492 is consistent with an ASC acting to flatten isopycnals on the western side of the cold filament.5 A reversal in sign of the cross-trajectory velocities is also visible in the first 200 km of the float trajectories (Fig. 8b), upstream of the filament, but the strength of the velocities is weaker. The spatial pattern and magnitude of vertical velocity in Fig. 8c is dominated by wrot [Eq. (2)], associated with the rotation of horizontal velocity with depth. Without information in both the along-stream and cross-stream directions, it remains difficult to unambiguously distinguish features related to the filament from those related to the mesoscale meander. However, the observations reveal an intensification of downward velocities associated with a cross-track overturning circulation, localized within the cold filament, that cannot be explained by mesoscale motions alone.
Subsurface (a) along-trajectory, (b) cross-trajectory, and (c) vertical velocity along each of the float tracks. Inertial oscillations have been removed by half-inertial pair averaging after the rotation and prior to the w calculation (see appendix). The data are interpolated onto a regular 3-km distance grid. Gray contours show potential density in 0.2 kg m−3 increments. The MLD is marked by the black solid line.
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
Submesoscale instabilities
Along the float tracks, we identify regions of the water column that are susceptible to submesoscale instabilities. Negative values of N2 indicate the potential for gravitational instability. Richardson numbers of O(1) and lower indicate an ageostrophic regime associated with intense vertical currents (Thomas et al. 2008; Siegelman et al. 2020b).
The WW layer at the beginning of the float tracks is associated with strong vertical stratification (Fig. 9a), which weakens as the WW descends along isopycnals to deeper depths. Clusters of negative N2 values emerge below the mixed layer (between 100 and 500 dbar), particularly in EM-8489 and EM-8492 from 300 km onward, indicating the potential for gravitational instabilities and convective overturning that could facilitate vertical mixing and subduction (Freilich and Mahadevan 2021). Although these negative values are observed well below the mixed layer, they appear to coincide with a reduction in MLD—notably at 400 km in EM-8489, and 300 and 600 km in EM-8492. While this may reflect another buoyancy-driven or mechanical process, or the movement of the float across a finescale front, submesoscale instabilities can reduce the MLD (Boccaletti et al. 2007; Fox-Kemper and Ferrari 2008; Yu et al. 2019) and may be partially responsible for this restratification. There are no negative N2 values along the trajectory of float EM-8493, implying a lack of gravitational instabilities. This float travels within the cold filament into the core of the cyclonic eddy, while the other two floats get ejected from the filament (Figs. 6c–e) and travel through the strain-dominated region surrounding the cyclonic eddy. This suggests that generation of instabilities and mixing tends to occur on the periphery of the filament and the eddy, in the strain-dominated regions.
Subsurface sections of (a) buoyancy frequency (N2) (upper 700 dbar), (b) vertical velocity shear (
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
Richardson numbers of O(1) and lower are observed in the surface mixed layer in all floats, where stratification is weak, but also appear in vertically coherent bands extending from below the mixed layer down to ∼1500 dbar (Fig. 9c)—indicating the potential for ageostrophic motion in the ocean interior. The spatial distribution of low Ri below the mixed layer is dominated by the velocity shear (Fig. 9b), which may be associated with the development of an ASC during filamentogenesis. The regions where Ri ∼ 1 are similar to those with elevated vertical velocities shown in Fig. 8c. The magnitude of the inferred vertical velocities (>100 m day−1), in addition to the low Richardson number, provide evidence of submesoscale-driven ageostrophic motion in the vicinity of the cold filament.
c. Cross-filament perspective
While the Triaxus does not sample the same filament that the floats do, or at least not at the same stage in its development, it crosses a finescale cold filament (∼5 km wide) that forms in almost the same location several days before the floats arrive. The Triaxus and shipboard ADCP observations provide further evidence of enhanced ageostrophic motion, as well as localized subduction and ventilation, associated with finescale cold filaments in this region.
From the SST map (Fig. 10a), the filament again appears to be drawn up from the southern side of the front and travels northward in between the eddy dipole structure. The filament thins as it travels north and is sampled by the Triaxus at its northernmost point (cyan dot on Fig. 10a) before it is mixed away. The cold filament is visible in the temperature anomaly section (Fig. 10e) approximately halfway along the transect (∼40 km) and extends from the surface to 300 dbar—the maximum depth of the Triaxus measurements. The cold temperature anomaly is less pronounced in the mixed layer and stronger at depth Fig. 10e). The filament is also associated with a fresh anomaly (Fig. 10f) and elevated oxygen (Fig. 10g) and chlorophyll (Fig. 10h) signatures that extend down to 300–200 dbar below the base of the mixed layer. The low AOU values (Fig. 10g), indicative of oxygen-saturated waters, in addition to the elevated chlorophyll observed below the mixed layer (Fig. 10h), provide strong evidence of recent subduction and ventilation associated with the cold filament. There are several weaker cold intrusions with relatively fresh, oxygenated and chlorophyll-rich signatures between 5 and 30 km that extend from the base of the mixed layer down to 200–250 dbar. These features are most pronounced in the interior, below the mixed layer, and do not have as strong an anomalous signature at the surface.
(a) Enlarged SST map on the day of the Triaxus tow (black line) along the ship track (gray line underneath). The thin black contour marks the 4°C SST contour and the cyan dot marks the location of the cold filament in the Triaxus transect. (b),(c) Shipboard ADCP velocities for the upper 300 dbar of the water column, rotated to cross track and along track, respectively. (d) Magnitude of horizontal divergence of along-track velocity, normalized by the Coriolis parameter f. From the Triaxus: (e) temperature anomaly, referenced to the transect mean on each pressure surface, (f) salinity, (g) apparent oxygen utilization (AOU), and (h) chlorophyll concentration. The black bar at 0 dbar in (b)–(h) indicates the location of the cold filament. The black contour in (b)–(h) marks the MLD, and thin contours show potential density in 0.04 kg m−3 intervals.
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
The shipboard ADCP measurements along the Triaxus transect provide a cross-filament perspective of the circulation, which we do not get from the float data. The along-track velocities (Fig. 10c) reveal a large-scale confluent flow—eastward (positive) on the western side of the cold filament and westward (negative) on the eastern side—that is coherent down to the maximum depth of the ADCP measurements (∼400 dbar, upper 300 dbar shown in Fig. 10c). The confluence axis, where the along-track velocities change sign in the horizontal, is aligned with the location of the cold filament near 40 km. This is consistent with the theory of cold filamentary intensification, described by McWilliams et al. (2009). The alignment between the filament and the confluence axis is slightly offset (by 5–10 km) in the mixed layer (Figs. 10c,e), where wind-driven Ekman currents could influence the velocities on short time scales. Winds on the day of the Triaxus tow were north-northwesterly, which would support an Ekman transport to the east-northeast. Based on eddy viscosity estimates (Aυ = 3.25–44.7 × 10−2 m2 s−1) from a standing meander of the ACC north of Kerguelen Island (Roach et al. 2015), the Ekman depth (
The magnitude of the convergence in the along-track direction (Fig. 10d) is mostly smaller than the Coriolis parameter (0.1–0.5f), implying a mesoscale-driven convergence (D’Asaro et al. 2018). However, there are vertical bands approaching f between 40 and 55 km and on the periphery of the cyclonic eddy at 60–65 km that imply submesoscale-driven convergence. The confluent flow and mesoscale convergence likely result from the strain field associated with the eddy dipole structure. The horizontal length scale of the strain field that drives filamentogenesis is expected to be much larger than the width of the filament, with smaller-scale convergent flows developing in response to the mesoscale deformation flow (Lapeyre and Klein 2006; Gula et al. 2014). We do not see a localized ageostrophic convergence at the surface in the location of the cold filament (Fig. 10d); however, the upper 16 dbar is unresolved by the ADCP. The cross-track velocities from the shipboard ADCP (Fig. 10b) reveal a weaker northward velocity on the eastern side (40–50 km) of the filament compared to the western side (30–40 km), consistent with the along-filament flow configuration in Fig. 1b.
At the eastern end of the Triaxus tow (>70 km), there is a strong vertically coherent cold anomaly associated with the cyclonic eddy (Fig. 10e). This feature, like the cold filament near 40 km, has a low salinity, high oxygen, and deep chlorophyll signature down to 250 dbar—potentially resulting from frontogenesis at the eddy periphery, resulting in downwelling on the dense side of the front. Alternatively, the similar properties of the cold filament near 40 km and the cold-core eddy at the end of the transect could suggest a common origin, as opposed to separate ventilation mechanisms.
For the first 40–50 km, the Triaxus is in a strain-dominated region (OW > 0) on the periphery of the cyclonic eddy (Fig. 11b). The strongest strain-dominated region occurs to the southwest of the eddy core, where the cold filament thins as it travels northward (Fig. 10a). This thinning of the filament is characteristic of filamentogenesis, associated with strong convergence and mesoscale strain, as depicted in Fig. 1b. From 50 km onward, the Triaxus enters a strong vorticity-dominated region (OW < 0) associated with the core of the cyclonic eddy. While the Okubo–Weiss parameter, calculated from satellite-derived velocities, gives an indication of the flow regime and where one might expect to find submesoscale filaments, the coarse spatial and temporal resolution of currently available satellite data limits its usage in understanding submesoscale features.
Surface maps of (a) backward-in-time FSLE and (b) the Okubo–Weiss parameter on the day of the Triaxus tow (black line). SSH streamlines in 0.1-m intervals are shown as gray contours. Interpolated values of (c) FSLE and (d) Okubo–Weiss along the Triaxus transect, with shading representing the temporal error (±1 day). The black bar in (c) and (d) indicates the filament location.
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
A strong negative peak in FSLE is present in the middle of the Triaxus transect near 35 km (Fig. 11c), representing an attracting Lagrangian coherent structure associated with a rapid confluence of particles (Boffetta et al. 2001; D’Ovidio et al. 2004; Meunier and LaCasce 2021). This is likely associated with the submesoscale cold filament sampled by the Triaxus and may provide the strong localized surface convergence required for cold filamentary intensification. The FSLE line appears to spiral into the cyclonic eddy (Fig. 11a), supporting a transport of water properties from the filament into the eddy. The Triaxus crosses the FSLE feature again at 50–55 km, but the signal is weaker—this may be associated with the front at the eddy periphery (∼65 km; Fig. 10). The negative FSLE peak in the middle of the transect is offset from the filament (Fig. 10e) and the confluence axis (Fig. 10c) by about 5–10 km; however, because FSLEs are computed from daily satellite-derived geostrophic velocities with a horizontal resolution of 0.25°, the locations are not expected to match up perfectly. In addition, interpolation onto the surface locations and times of the Triaxus is likely to introduce temporal and spatial errors given the transient nature of the filaments and the high-resolution sampling of the Triaxus.
Submesoscale instabilities
In close proximity to the cold filament between 40 and 50 km, there is an increase in the occurrence of negative N2 values, positive PV values, and Ri ≤ 1 below the mixed layer (Fig. 12d)—a signal 3–4 times larger than the background noise. This implies enhanced susceptibility of the water column to submesoscale instabilities and ageostrophic vertical motion, potentially associated with filamentogenesis. The weak stratification suggests a role of turbulent mixing in preconditioning the flow to gravitational instabilities. Positive PV values (opposite sign to f) are dominated by unstable stratification (N2 < 0) in the first term of the PV approximation, rather than the buoyancy gradient term that is often associated with symmetric instability (Thomas et al. 2013; Gula et al. 2022). While we do not account for along-front gradients of velocity and buoyancy in this PV approximation (see section 2), these terms are expected to be small in the location of the filament. The Ri criteria used in Fig. 12d does not include negative values associated with N2 < 0, and thus implies an increase in the susceptibility of the flow to shear instabilities on the eastern side of the filament. The surface mixed layer along the Triaxus transect, ranging from 50 to 100 dbar, is characterized by weak stratification (Fig. 12a) and a low Ri number of <0.25 (Fig. 12c), as expected in a turbulent, well-mixed region. Stratification is enhanced at the base of the mixed layer and also at depth, associated with subsurface thermohaline structures.
(a) Buoyancy frequency N2, (b) Ertel PV approximation, (f + υx)N2 − υzbx, and (c) Richardson number,
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
4. Discussion
a. Evidence of cold filamentary intensification
Cold filamentary intensification is well described in theory (McWilliams et al. 2009, 2015) and models (Gula et al. 2014), but there is a lack of observational evidence due to the short temporal and spatial scales of submesoscale features (D’Asaro et al. 2011; Archer et al. 2020). While there are limitations associated with the observations used in this study (detailed at the end of the discussion), the evidence supports the process of cold filamentary intensification occurring in the meander region. An idealized schematic is shown in Fig. 13, drawing together some of the features observed in this particular meander of the ACC, adapted from the theory of cold filamentary intensification described by McWilliams et al. (2009), Gula et al. (2014).
Idealized schematic of a standing meander of the ACC, illustrating the process of cold filamentary intensification associated with an eddy dipole, adapted from McWilliams et al. (2009). The meandering flow draws water from the cold side of the front up in between the eddy dipole structure, where it is filamented into a submesoscale feature by the mesoscale strain field (plan view map). Below the surface, along the cross section of the filament (lower section of schematic), two ageostrophic secondary circulation cells (ASC) result in surface convergence at the filament location, intensifying downward velocities within the filament.
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
The velocity structure in the EM-APEX observations is consistent with the circulation expected during filamentogenesis on the western side of a cold filament in the Southern Hemisphere (Figs. 1b and 13). When the floats travel north within the filament, the cross-track velocities flow eastward (negative) in the upper ocean (0–500 dbar) and westward underneath (500–1600 dbar) (Fig. 8b), forming an overturning circulation acting in the direction to restratify the water column—consistent with an ASC associated with filamentogenesis. This pattern in the cross-track velocities is amplified in the location of the cold filament, where there is an along-track acceleration (Fig. 8a)—implying steeper cross-front isopycnals through the thermal wind relation. The intense downward velocities (>100 m day−1; Fig. 8c), associated with the vertical velocity shear, provide strong evidence for the contribution of submesoscale/ageostrophic motion associated with cold filamentary intensification.
While the floats capture the overturning circulation in the cross-track velocities, when they travel within the filament, the ADCP velocities capture a cross-filament perspective of the circulation in the upper 300 dbar. These observations were not collected simultaneously, but they were taken in almost the same location a few days apart, and thus offer some complimentary perspectives of the circulation associated with the cold filaments. The ADCP velocities reveal a confluent flow directed toward the cold filament (Fig. 10c) that has a horizontal scale of O(10) km—much larger than the width of the filament. This large-scale confluence, associated with the mesoscale strain field, is a key driver of filamentogenesis and is expected to occur on a much larger scale than the filament itself (Lapeyre and Klein 2006; McWilliams et al. 2009; Gula et al. 2014). A range of scales of convergence (from submesoscale to mesoscale) are captured in the along-track ADCP velocities (Fig. 10d); however, we do not observe a strong localized convergence at the filament location, which is key to cold filamentary intensification (McWilliams et al. 2009). However, the negative FSLE line at 35 km (Figs. 11a,c) could support localized convergence (if the flow is three dimensional) and also implies rapid frontal sharpening—a process that occurs during cold filamentary intensification that we are unable to estimate from the daily SST maps.
Separating the contributions from submesoscale and mesoscale processes is often unattainable, particularly in observations that capture a range of scales. The filament sampled by the floats is ∼20 km wide (Figs. 6a,b). With a characteristic velocity U of ∼1 m s−1 and a length scale L of 20 km, the Rossby number (U/fL) is ∼0.5 at the latitude of the filament—representing a scale somewhere between the submesoscale and mesoscale. However, the SST gradients on the edges of the filament occur over shorter length scales of <5–10 km (Fig. 6b), and would therefore have a Rossby number closer to 1. The magnitude of the vertical velocities (>100 m day−1) exceeds that of mesoscale motions by an order of magnitude (Thomas et al. 2008), and the low Richardson numbers extending into the ocean interior (Fig. 9c) provide evidence for ageostrophic motion associated with the cold filament, rather than a mesoscale process occurring at the PF. There may also be vertical velocity contributions from internal waves, but we have minimized their direct effect on the EM-APEX observations by averaging near-inertial profile pairs. Strain rate is also related to eddy energy (Lapeyre and Klein 2006), so the high EKE in the meander region (Fig. 5a) can increase the vertical motions associated with the filaments. Separating the contributions from mesoscale and submesoscale processes is complicated further by interactions between scales, with mesoscale dynamics influencing the submesoscale field and submesoscale dynamics influencing the mesoscale (Boccaletti et al. 2007; Schubert et al. 2020; Zhang et al. 2023b).
b. Vertical extent of submesoscale motions
A surprising feature of the observations is the depth extent of vertical motion and ageostrophic behavior. Submesoscale processes such as cold filamentary intensification are often described as an upper-ocean process (McWilliams et al. 2009; Gula et al. 2014; Rosso et al. 2015; Lévy et al. 2018), while our observations reveal a larger-scale process—with enhanced vertical velocities and evidence of ageostrophic motion extending as deep as 1600 dbar in a ∼20-km-wide filament (Figs. 8 and 9).
Recent studies have also found evidence of deep-reaching submesoscale activity down to 500–900 m in the energetic regions of the Southern Ocean (Siegelman et al. 2020a; Siegelman 2020), associated with standing meanders and hotspots of EKE downstream of topography (Siegelman et al. 2019; Dove et al. 2021). The weak vertical stratification in the Southern Ocean may explain the much deeper extent of ageostrophic motion and vertical velocities present in these observations, compared to previous literature on cold filamentary intensification (McWilliams et al. 2009; Gula et al. 2014). As described by Lapeyre and Klein (2006), in regions of strong stratification the vertical velocity will be concentrated in the upper tens of meters. This suggests that the energetic and deep-reaching nature of ACC fronts and jets may be an important feature that extends the process of cold filamentary intensification and the influence of submesoscale motions deeper into the ocean interior. The depth of submesoscale instabilities and vertical motion has important consequences for heat, nutrient, and carbon fluxes (Lapeyre and Klein 2006; Mahadevan 2016)—influencing the structure of marine ecosystems (Lévy et al. 2018) and the ability for the Southern Ocean to regulate global climate.
c. Localized subduction and ventilation
Localized subduction and ventilation associated with submesoscale structures has been explored in models (Balwada et al. 2018; Freilich and Mahadevan 2021) and in observations of energetic regions of the Southern Ocean (Llort et al. 2018; Dove et al. 2021). However, there is an ongoing debate about the route along which tracers are carried into the interior, and the relative contributions of localized subduction and diapycnal processes versus remote subduction and subsequent lateral advection or stirring along isopycnals. As discussed by Balwada et al. (2018), Dove et al. (2021), and Morrison et al. (2022), while the initial subduction from the mixed layer is likely to be heavily influenced by localized submesoscale processes, mesoscale motions and isopycnal stirring likely play a key role in transporting tracers within the ocean interior.
In the observations presented here, while we cannot eliminate the potential contribution from remote subduction, we find evidence of localized subduction and ventilation both in the float and Triaxus data. The EM-APEX floats show evidence of subduction on the periphery of the cold filament—where isopycnals are observed to outcrop in the mixed layer and Tmin waters descend along isopycnals (Fig. 7a), coincident with a weakening of the vertical stratification and strong along-isopycnal vertical velocities (Fig. 8c). The susceptibility of the flow to submesoscale instabilities (Fig. 9) implies potential diapycnal mixing and vertical motion within the ocean interior, in addition to the along-isopycnal component. While we do not have vertical velocities along the Triaxus transect, the low AOU and elevated chlorophyll anomalies, extending ∼200 dbar below the mixed layer within the cold filament (Figs. 10e–h), provide strong evidence of recent subduction and ventilation. The orientation of tracer contours across isopycnals in the cold filament (Figs. 10e–h), and the susceptibility of the water column to submesoscale instabilities (Fig. 12), suggest an enhancement of vertical transport through diapycnal mixing—a process that often occurs during filamentogeneis (Callies et al. 2016; Yu et al. 2019; Freilich and Mahadevan 2021). While FSLE alone only provides a measure of horizontal deformation (Hernández-Carrasco et al. 2018), the close alignment between the negative FSLE peak and the submesoscale cold filament (Figs. 11a,c), along with the in situ evidence of subduction and ventilation, suggests that this feature may be associated with strong surface convergence and intense vertical motion. These results support the notion that filamentogenesis associated with finescale cold filaments can influence vertical exchanges of heat, nutrients, and carbon across the base of the mixed layer—with consequences for primary productivity, oceanic heat uptake, and the carbon cycle (Lapeyre and Klein 2006; Mahadevan 2016; Lévy et al. 2018; Su et al. 2020).
In addition to the strong evidence for localized subduction and ventilation in the cold filaments, there are some indications of isopycnal stirring that may play a role in transporting tracers in the ocean interior. For example, there are several weaker cold and fresh intrusions in the first 30 km of the Triaxus transect that are also associated with deep AOU and chlorophyll anomalies. However, unlike the cold filament at 40 km, these intrusions do not have a clear anomaly in the mixed layer and are slightly tilted in the vertical, with tracer contours aligned along isopycnals below 200 dbar. The intrusions have a weaker chlorophyll signature compared to the cold filament in the middle of the transect and may be a consequence of a lateral advection or isopycnal stirring process, as opposed to (or following) local subduction. Alternatively, these features may be remnant structures associated with previous subduction events, with their surface signals diminished by rapid mixed layer processes.
Frontogenesis at the periphery of the cold-core eddy may also result in subduction and ventilation—elevated oxygen and chlorophyll signatures are observed down to 300 dbar at the eastern end of the Triaxus transect on the edge of the cyclonic eddy (Figs. 10e–h). In this case, one would expect downwelling associated with the secondary circulation on the dense side of the front and upwelling on the light side of the front (McWilliams et al. 2009, 2015), where we observe the warm and salty anomaly with low oxygen between 45 and 65 km. While the subsurface water mass properties measured by the Triaxus support this mechanism, and there is a strengthening of the cross-track velocities between 60 and 70 km indicative of a frontal jet (Fig. 10b), a secondary circulation associated with frontogenesis at the eddy periphery is not visible in the along-track ADCP velocities.
d. Implications for cross-frontal exchange
Standing meanders of the ACC, generated by flow–topography interactions, are hotspots of enhanced cross-frontal exchange (Naveira Garabato et al. 2011; Thompson and Sallée 2012) and poleward eddy heat flux (Foppert et al. 2017). Mesoscale eddies generated by baroclinic instabilities can both carry tracers across the front through lateral advection (Dong et al. 2014; Dufour et al. 2015) and through stirring of the large-scale tracer gradient (Naveira Garabato et al. 2011; Frenger et al. 2015). Stirring generates finescale structures that encourage the cascade of tracer variance toward small scales (Naveira Garabato et al. 2016)—enhancing the mixing potential and the cross-frontal exchange of heat and other properties (Naveira Garabato et al. 2011). Submesoscale structures generated by eddy stirring are also associated with lines of high FSLE (Siegelman 2020; Dove et al. 2022)—forming Lagrangian coherent structures that organize the flow and transport tracers from one place to another (D’Ovidio et al. 2004; Hernández-Carrasco et al. 2012; Denes et al. 2022).
The cross-frontal secondary circulation associated with frontogenesis and filamentogenesis acts to flatten density surfaces (McWilliams et al. 2009; Gula et al. 2014) and encourage cross-frontal exchange through along-isopycnal transport (Sanchez-Rios et al. 2020). Submesoscale instabilities that develop as the front intensifies encourage both lateral and vertical mixing (Taylor and Ferrari 2010; D’Asaro et al. 2011), and can cause frontal arrest—preventing further frontal sharpening (McWilliams and Molemaker 2011) and reducing the ability of the front to act as a transport barrier.
From the SST (Figs. 6 and 4a) and FSLE (Fig. 11a) maps, it appears that the cold filaments sampled by the floats and the Triaxus originated south of the PF and were elongated in the northward direction in between the eddy dipole, subsequently feeding into the cyclonic eddy. We suspect, based on a time evolution of SST maps from October 2018 to March 2019 (not shown), that this is a relatively common process in this region, and perhaps in other energetic meander regions associated with mesoscale fronts. The time scale of persistence or destruction of finescale filaments may determine the extent to which they can transport tracers into the ocean interior and also laterally across mesoscale fronts. The mechanisms for, and contribution of, submesoscale processes in the cross-frontal exchange of tracers in energetic meander regions requires further research.
e. Observational limitations
While in situ observations, like those used in this study, provide an invaluable view of the subsurface ocean, they are challenging to obtain, especially in energetic and remote regions of the Southern Ocean. They are limited in spatial and temporal coverage and are subject to the high-frequency variability of the ocean. The two-dimensional nature of the observations (vertical/slant profiles along a transect or float trajectory) limits the ability to resolve the three-dimensional nature of the flow and fully characterize potential vorticity. In addition to the instruments moving in space, there is a time-varying component that can make it difficult to discern whether an observed change along the instrument track is a result of spatial or temporal variability, or a combination of both. For the EM-APEX floats in particular, the along-trajectory evolution of tracer and velocity fields does not necessarily reflect purely along-stream or along-front variations—the floats are subject to ageostrophic flow and the vertical shear of the current while profiling, allowing them to travel across SSH contours. In this region with high EKE and submesoscale activity, it is not surprising that the floats occasionally cross SSH streamlines that capture only mesoscale/geostrophic features. However, the fact that the floats and submesoscale features do not follow SSH streamlines reinforces the idea that finescale features and ageostrophic velocities play a key role in lateral exchange and transport of water mass properties across mesoscale fronts. Finescale ocean observations that simultaneously capture both along-front and cross-front variability, as well as the vertical structure, will be key in developing a better understanding of these three-dimensional processes.
Satellite SST images are often patchy due to clouds, particularly in the Southern Ocean, and most satellite products are too coarse to resolve finescale features, making it difficult to observe the superexponential growth associated with filamentogenesis. However, the negative FSLE peaks reveal finescale confluence zones that imply rapid frontal sharpening and are often associated with submesoscale fronts (D’Ovidio et al. 2004; Hernández-Carrasco et al. 2012; Siegelman et al. 2020a,b). While we have used gridded satellite products that merge data from multiple satellite missions, offering greater temporal and spatial coverage, individual satellite revisit times can vary from a few days to a few weeks. Errors are inevitably introduced during the interpolation to produce the daily gridded product. With the recent launch of the Surface Water and Ocean Topography (SWOT) satellite altimeter with 0.05° spatial resolution (swot.jpl.nasa.gov/mission/overview/), there will be new capabilities to resolve finescale features in the surface ocean. However, the limited temporal resolution associated with the 21-day repeat orbit will make it difficult to track the evolution of these features in space and time.
5. Conclusions
We present a unique set of observations in an energetic meander region of the ACC that reveal evidence supportive of cold filamentary intensification. While previous studies discuss this process in the context of the upper ocean (McWilliams et al. 2009; Gula et al. 2014; Freilich and Mahadevan 2021), our observations reveal a larger-scale process of the same nature, with enhanced vertical velocities and evidence of ageostrophic motion extending as deep as 1600 dbar—well below the base of the mixed layer. The weak stratification and deep-reaching nature of the ACC may allow the process of cold filamentary intensification to extend deep into the ocean interior. The depth of vertical motions, as well as the frequency and persistence of these filaments in the Southern Ocean, has important consequences for vertical fluxes of heat, nutrients, and carbon.
Submesoscale fronts and filaments are not captured in global climate models (Lévy et al. 2018; Hewitt et al. 2022), which may lead to misrepresentation of the magnitude and locations of subduction pathways and tracer transport across the world’s oceans. This is particularly relevant in regions with energetic fronts, where submesoscale dynamics are likely to be strong, such as the ACC, Gulf Stream, and Kuroshio Current. Standing meanders and rich eddy fields generated by flow–topography interactions provide an ideal environment for the generation of finescale filaments that can influence cross-frontal transport pathways and have deep-reaching consequences on vertical motions and interior stratification.
For velocities above 100 dbar, where the surface waves dominate, we do not apply the cutoff based on RMS error. Instead, we exclude velocities with a magnitude greater than 2 m s−1.
The curvature values are insensitive to the temporal discretization.
We also tested rotating the velocity components with respect to SSH contours and with respect to finite-sized Lyapunov exponent (FSLE) eigenvectors. The coarse resolution of the SSH product and the flow curvature associated with the eddy dipole structure led to unrealistic and rapidly varying frontal orientations.
Different time periods of 1, 3, and 10 years were tested for the mean field; however, the spatial pattern and magnitude of EKE were insensitive to the choice of time period, suggesting low interannual variability of mean kinetic energy in the region.
Positive cross-trajectory velocity is to the left of the float track.
Of the float profiles, 94% and 83% had ≤6 missing velocity measurements (≤12 dbar) at the surface and at the bottom of the profile, respectively. The maximum number of measurements filled was 8 at the surface and 15 at the bottom.
Acknowledgments.
We thank the R/V Investigator officers and crew, CSIRO Marine National Facility technical team and scientists who contributed to the collection of the data on the IN2018_V05 voyage. Their efforts enabled a comprehensive finescale survey of an Antarctic Circumpolar Current meander in the presence of significant winds and waves. We acknowledge support from the Australian Government as part of the Antarctic Science Collaboration Initiative, the Australian Research Council’s Discovery Projects (DP170102162) and Special Research Initiative, Australian Centre for Excellence in Antarctic Science (SR200100008) and from the Australian Government under the National Environmental Science Program (Climate Systems). Our team has benefited enormously from the Graduate Program of the Australian Research Council Centre of Excellence for Climate Extremes (CE170100023). AFT was supported by the Resnick Sustainability Institute and the Ginkgo Foundation. MIJ thanks Helen Phillips for writing the original routine EM-APEX quality control MATLAB scripts, and Ajitha Cyriac and Jan Jaap Meijer for adapting and documenting the scripts. We are grateful for the detailed and constructive comments from three anonymous reviewers that have helped improve the paper.
Data availability statement.
The EM-APEX float data and quality control toolbox is publicly available at the Australian Antarctic Data Centre (data.aad.gov.au/metadata/EMAPEX_Macquarie_2018). Triaxus and ADCP data from the voyage are publicly available for download from the MNF website (marine.csiro.au/data/trawler/survey_details.cfm?survey=IN2018_V05). Daily L3-gridded (0.02°) SST data used in this study can be downloaded via the IMOS data portal (researchdata.edu.au/imos-srs-sst-time-australia/1370442?source=suggested_datasets). Daily L4-gridded (0.25°) SSH and derived geostrophic velocities can be obtained from the Copernicus website (data.marine.copernicus.eu/product/SEALEVEL_GLO_PHY_L4_MY_008_047/). FSLEs (0.04° grid resolution) computed from Copernicus L4-gridded geostrophic velocities can be downloaded from Aviso+ (aviso.altimetry.fr/en/data/products/value-added-products/fsle-finite-size-lyapunov-exponents). Python implementation of the TEOS-10 Gibbs Sea-Water (GSW) Oceanographic Toolbox is available at github.com/TEOS-10/GSW-Python. Python scripts for this analysis will be made publicly available at github.com/mayajakes/phd-public.
APPENDIX
EM-APEX Data Processing
a. Absolute velocities
Missing values of relative velocity at the surface and at the bottom of each profile were replaced with the nearest value from that profile, assuming constant velocity.A1 Several erroneous relative velocity profiles in each float were manually removed after the initial QC—these were identified as spikes in the velocities that were not consistent with the surrounding profiles.
To obtain absolute velocities, a velocity offset is calculated for each profile pair (down- and up-profile) based on the difference between the velocity components according to the GPS surface positions and times, and the depth-integrated relative velocity components measured by the float. Figure A1 shows an example of the estimated subsurface path of the float (using deduced reckoning) before and after adding the velocity offset. The depth-integrated subsurface float positions provide a theoretical resurface position, while the GPS measurements provide the actual resurface position of the float. The offset values are large in the ACC due to the large bottom velocities and strong mean flow—the relative velocities represent only the baroclinic component. We encountered several profile pairs that had almost unchanged GPS positions, leading to unrealistic absolute velocities. An example of this is shown in the bottom panel of Fig. A1. Rather than altering the GPS positions, we instead remove these velocity profiles.
Subsurface path of float EM-8489 during a down- and up-profile according to the (left) measured relative velocities and time recorded by the CTD sensor at each pressure level and (center) with the velocity offset added to the relative velocities (i.e., absolute velocities). Red dots indicate the surface GPS positions related to the two profiles. The bottom panel shows how an erroneous GPS position leads to unrealistic absolute velocities.
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
b. Rotated velocities
For the rotation of velocities from Cartesian coordinates into along trajectory and cross trajectory, we use the absolute velocities before half-inertial pair averaging and find the estimated position of the float at the bottom of each down-profile (1600 dbar) using deduced reckoning. The along-trajectory bearing for the down-profiles is taken from the surface GPS position to the position of the float at 1600 dbar. The bearing for the up-profiles is then taken from the subsurface position at 1600 dbar (bottom of the previous down-profile) to the GPS position when the float resurfaces. This method provides an accurate representation of the depth-integrated stream direction experienced by the float as it descends to 1600 dbar and ascends back to the surface. An example of the along-trajectory (stream) bearings for profiles 100–119 in Em-8492 is shown in Fig. A2.
(a) Trajectory of float EM-8492, with the float positions where each bearing is taken from displayed as gray dots and the along-trajectory (stream) bearings for profiles 100–119 shown in the red arrows. (b) Enlarged quiver plot of the velocity bearings at 200 dbar (cyan arrows) and the stream bearings associated with profiles 100–119.
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
Along-trajectory and cross-trajectory velocities are calculated as
c. Removal of inertial oscillations
To remove the effects of inertial oscillations from the float data we perform half-inertial pair averaging of the temperature, salinity, and velocity profiles, then interpolate the values back onto the original time grid. Consecutive down-profiles and consecutive up-profiles are approximately half an inertial period apart at the latitude of float deployment (π/f = 7.29 h at 55.15°S). The sampling frequency of the floats after deployment was not altered to account for the change in latitude of the floats. The half-inertial period (π/f) at the northernmost position of the floats is 7.63 h.
Example plots for absolute velocity half-inertial pairs are shown in Fig. A3. For the rotated velocities, we perform the half-inertial pair averaging after rotation rather than before, as the rotation method uses the absolute velocity profiles to determine the location of the float at the bottom of each down-profile.
Examples of half-inertial pair averaging of absolute velocity components for a pair of consecutive (top) down- and (bottom) up-profiles in float EM-8489. Note the mirror-imaging of zonal and meridional velocity profiles that are half an inertial period apart.
Citation: Journal of Physical Oceanography 54, 3; 10.1175/JPO-D-23-0085.1
REFERENCES
Archer, M., A. Schaeffer, S. Keating, M. Roughan, R. Holmes, and L. Siegelman, 2020: Observations of submesoscale variability and frontal subduction within the mesoscale eddy field of the Tasman Sea. J. Phys. Oceanogr., 50, 1509–1529, https://doi.org/10.1175/JPO-D-19-0131.1.
Artale, V., G. Boffetta, A. Celani, M. Cencini, and A. Vulpiani, 1997: Dispersion of passive tracers in closed basins: Beyond the diffusion coefficient. Phys. Fluids, 9, 3162–3171, https://doi.org/10.1063/1.869433.
Aurell, E., G. Boffetta, A. Crisanti, G. Paladin, and A. Vulpiani, 1996: Growth of noninfinitesimal perturbations in turbulence. Phys. Rev. Lett., 77, 1262–1265, https://doi.org/10.1103/PhysRevLett.77.1262.
Bachman, S. D., and A. Klocker, 2020: Interaction of jets and submesoscale dynamics leads to rapid ocean ventilation. J. Phys. Oceanogr., 50, 2873–2883, https://doi.org/10.1175/JPO-D-20-0117.1.
Bachman, S. D., J. R. Taylor, K. A. Adams, and P. J. Hosegood, 2017: Mesoscale and submesoscale effects on mixed layer depth in the Southern Ocean. J. Phys. Oceanogr., 47, 2173–2188, https://doi.org/10.1175/JPO-D-17-0034.1.
Balwada, D., K. S. Smith, and R. Abernathey, 2018: Submesoscale vertical velocities enhance tracer subduction in an idealized Antarctic Circumpolar Current. Geophys. Res. Lett., 45, 9790–9802, https://doi.org/10.1029/2018GL079244.
Belkin, I. M., and A. L. Gordon, 1996: Southern Ocean fronts from the Greenwich Meridian to Tasmania. J. Geophys. Res., 101, 3675–3696, https://doi.org/10.1029/95JC02750.
Boccaletti, G., R. Ferrari, and B. Fox-Kemper, 2007: Mixed layer instabilities and restratification. J. Phys. Oceanogr., 37, 2228–2250, https://doi.org/10.1175/JPO3101.1.
Boffetta, G., G. Lacorata, G. Redaelli, and A. Vulpiani, 2001: Detecting barriers to transport: A review of different techniques. Physica D, 159, 58–70, https://doi.org/10.1016/S0167-2789(01)00330-X.
Bower, A., and T. Rossby, 1989: Evidence of cross-frontal exchange processes in the Gulf Stream based on isopycnal RAFOS float data. J. Phys. Oceanogr., 19, 1177–1190, https://doi.org/10.1175/1520-0485(1989)019<1177:EOCFEP>2.0.CO;2.
Caldeira, K., and P. B. Duffy, 2000: The role of the Southern Ocean in uptake and storage of anthropogenic carbon dioxide. Science, 287, 620–622, https://doi.org/10.1126/science.287.5453.620.
Callies, J., G. Flierl, R. Ferrari, and B. Fox-Kemper, 2016: The role of mixed-layer instabilities in submesoscale turbulence. J. Fluid Mech., 788, 5–41, https://doi.org/10.1017/jfm.2015.700.
Capet, X., J. C. McWilliams, M. J. Molemaker, and A. F. Shchepetkin, 2008: Mesoscale to submesoscale transition in the California Current system. Part I: Flow structure, eddy flux, and observational tests. J. Phys. Oceanogr., 38, 29–43, https://doi.org/10.1175/2007JPO3671.1.
Chelton, D. B., R. A. deSzoeke, M. G. Schlax, K. E. Naggar, and N. Siwertz, 1998: Geographical variability of the first baroclinic Rossby radius of deformation. J. Phys. Oceanogr., 28, 433–460, https://doi.org/10.1175/1520-0485(1998)028<0433:GVOTFB>2.0.CO;2.
Chrysagi, E., L. Umlauf, P. Holtermann, K. Klingbeil, and H. Burchard, 2021: High-resolution simulations of submesoscale processes in the Baltic Sea: The role of storm events. J. Geophys. Res. Oceans, 126, e2020JC016411, https://doi.org/10.1029/2020JC016411.
Cotté, C., F. D’Ovidio, A. C. Dragon, C. Guinet, and M. Lévy, 2015: Flexible preference of southern elephant seals for distinct mesoscale features within the Antarctic Circumpolar Current. Prog. Oceanogr., 131, 46–58, https://doi.org/10.1016/j.pocean.2014.11.011.
Cyriac, A., H. E. Phillips, N. L. Bindoff, H. Mao, and M. Feng, 2021: Observational estimates of turbulent mixing in the southeast Indian Ocean. J. Phys. Oceanogr., 51, 2103–2128, https://doi.org/10.1175/JPO-D-20-0036.1.
Cyriac, A., H. E. Phillips, N. L. Bindoff, and K. Polzin, 2022: Turbulent mixing variability in an energetic standing meander of the Southern Ocean. J. Phys. Oceanogr., 52, 1593–1611, https://doi.org/10.1175/JPO-D-21-0180.1.
D’Asaro, E., C. Lee, L. Rainville, R. Harcourt, and L. Thomas, 2011: Enhanced turbulence and energy dissipation at ocean fronts. Science, 332, 318–322, https://doi.org/10.1126/science.1201515.
D’Asaro, E. A., and Coauthors, 2018: Ocean convergence and the dispersion of flotsam. Proc. Natl. Acad. Sci. USA, 115, 1162–1167, https://doi.org/10.1073/pnas.1718453115.
d’Ovidio, F., V. Fernández, E. Hernández-García, and C. López, 2004: Mixing structures in the Mediterranean Sea from finite-size Lyapunov exponents. Geophys. Res. Lett., 31, L17203, https://doi.org/10.1029/2004GL020328.
Denes, M. C., G. Froyland, and S. R. Keating, 2022: Persistence and material coherence of a mesoscale ocean eddy. Phys. Rev. Fluids, 7, 034501, https://doi.org/10.1103/PhysRevFluids.7.034501.
Dong, C., J. C. McWilliams, Y. Liu, and D. Chen, 2014: Global heat and salt transports by eddy movement. Nat. Commun., 5, 3294, https://doi.org/10.1038/ncomms4294.
Dove, L. A., A. F. Thompson, D. Balwada, and A. R. Gray, 2021: Observational evidence of ventilation hotspots in the Southern Ocean. J. Geophys. Res. Oceans, 126, e2021JC017178, https://doi.org/10.1029/2021JC017178.
Dove, L. A., D. Balwada, A. F. Thompson, and A. R. Gray, 2022: Enhanced ventilation in energetic regions of the Antarctic Circumpolar Current. Geophys. Res. Lett., 49, e2021GL097574, https://doi.org/10.1029/2021GL097574.
Dufour, C. O., and Coauthors, 2015: Role of mesoscale eddies in cross-frontal transport of heat and biogeochemical tracers in the Southern Ocean. J. Phys. Oceanogr., 45, 3057–3081, https://doi.org/10.1175/JPO-D-14-0240.1.
du Pleiss, M., S. Swart, I. J. Ansorge, and A. Mahadevan, 2017: Submesoscale processes promote seasonal restratification in the Subantarctic Ocean. J. Geophys. Res. Oceans, 122, 2960–2975, https://doi.org/10.1002/2016JC012494.
du Pleiss, M., S. Swart, I. J. Ansorge, A. Mahadevan, and A. F. Thompson, 2019: Southern Ocean seasonal restratification delayed by submesoscale wind–front interactions. J. Phys. Oceanogr., 49, 1035–1053, https://doi.org/10.1175/JPO-D-18-0136.1.
Erickson, Z. K., and A. F. Thompson, 2018: The seasonality of physically driven export at submesoscales in the northeast Atlantic Ocean. Global Biogeochem. Cycles, 32, 1144–1162, https://doi.org/10.1029/2018GB005927.
Foppert, A., 2019: Observed storm track dynamics in Drake Passage. J. Phys. Oceanogr., 49, 867–884, https://doi.org/10.1175/JPO-D-18-0150.1.
Foppert, A., K. A. Donohue, D. R. Watts, and K. L. Tracey, 2017: Eddy heat flux across the Antarctic Circumpolar Current estimated from sea surface height standard deviation. J. Geophys. Res. Oceans, 122, 6947–6964, https://doi.org/10.1002/2017JC012837.
Fox-Kemper, B., and R. Ferrari, 2008: Parameterization of mixed layer eddies. Part II: Prognosis and impact. J. Phys. Oceanogr., 38, 1166–1179, https://doi.org/10.1175/2007JPO3788.1.
Fox-Kemper, B., and Coauthors, 2011: Parameterization of mixed layer eddies. III: Implementation and impact in global ocean climate simulations. Ocean Modell., 39, 61–78, https://doi.org/10.1016/j.ocemod.2010.09.002.
Freilich, M., and A. Mahadevan, 2021: Coherent pathways for subduction from the surface mixed layer at Ocean fronts. J. Geophys. Res. Oceans, 126, e2020JC017042, https://doi.org/10.1029/2020JC017042.
Frenger, I., M. Münnich, N. Gruber, and R. Knutti, 2015: Southern Ocean eddy phenomenology. J. Geophys. Res. Oceans, 120, 7413–7449, https://doi.org/10.1002/2015JC011047.
Frölicher, T. L., J. L. Sarmiento, D. J. Paynter, J. P. Dunne, J. P. Krasting, and M. Winton, 2015: Dominance of the Southern Ocean in anthropogenic carbon and heat uptake in CMIP5 models. J. Climate, 28, 862–886, https://doi.org/10.1175/JCLI-D-14-00117.1.
Griffin, C., H. Beggs, and L. Majewski, 2017: GHRSST compliant AVHRR SST products over the Australian region – version 1. Bureau of Meteorology Tech. Rep., 151 pp., https://imos.org.au/fileadmin/user_upload/shared/SRS/SST/GHRSST-DOC-basic-v1.0r1.pdf.
Gula, J., M. J. Molemaker, and J. C. McWilliams, 2014: Submesoscale cold filaments in the Gulf Stream. J. Phys. Oceanogr., 44, 2617–2643, https://doi.org/10.1175/JPO-D-14-0029.1.
Gula, J., J. Taylor, A. Shcherbina, and A. Mahadevan, 2022: Submesoscale processes and mixing. Ocean Mixing: Drivers, Mechanisms and Impacts, M. Meredith and A. C. Naveira Garabato, Eds., Elsevier, 181–214, https://doi.org/10.1016/B978-0-12-821512-8.00015-3.
Henson, S. A., and A. C. Thomas, 2008: A census of oceanic anticyclonic eddies in the Gulf of Alaska. Deep-Sea Res. I, 55, 163–176, https://doi.org/10.1016/j.dsr.2007.11.005.
Hernández-Carrasco, I., C. López, E. Hernández-García, and A. Turiel, 2012: Seasonal and regional characterization of horizontal stirring in the global ocean. J. Geophys. Res., 117, C10007, https://doi.org/10.1029/2012JC008222.
Hernández-Carrasco, I., A. Orfila, V. Rossi, and V. Garçon, 2018: Efect of small scale transport processes on phytoplankton distribution in coastal seas. Sci. Rep., 8, 8613, https://doi.org/10.1038/s41598-018-26857-9.
Hewitt, H., B. Fox-Kemper, B. Pearson, M. Roberts, and D. Klocke, 2022: The small scales of the ocean may hold the key to surprises. Nat. Climate Change, 12, 496–499, https://doi.org/10.1038/s41558-022-01386-6.
Hoskins, B. J., 1982: The mathematical theory of frontogenesis. Annu. Rev. Fluid Mech., 14, 131–151, https://doi.org/10.1146/annurev.fl.14.010182.001023.
Hoskins, B. J., and F. P. Bretherton, 1972: Atmospheric frontogenesis models: Mathematical formulation and solution. J. Atmos. Sci., 29, 11–37, https://doi.org/10.1175/1520-0469(1972)029<0011:AFMMFA>2.0.CO;2.
Hughes, C. W., 2005: Nonlinear vorticity balance of the Antarctic Circumpolar Current. J. Geophys. Res., 110, C11008, https://doi.org/10.1029/2004JC002753.
Ito, T., M. J. Follows, and E. A. Boyle, 2004: Is AOU a good measure of respiration in the oceans? Geophys. Res. Lett., 31, L17305, https://doi.org/10.1029/2004GL020900.
Khatiwala, S., F. Primeau, and T. Hall, 2009: Reconstruction of the history of anthropogenic CO2 concentrations in the ocean. Nature, 462, 346–349, https://doi.org/10.1038/nature08526.
Kim, Y. S., and A. H. Orsi, 2014: On the variability of Antarctic Circumpolar Current fronts inferred from 1992–2011 altimetry. J. Phys. Oceanogr., 44, 3054–3071, https://doi.org/10.1175/JPO-D-13-0217.1.
Klein, P., and G. Lapeyre, 2009: The oceanic vertical pump induced by mesoscale and submesoscale turbulence. Annu. Rev. Mar. Sci., 1, 351–375, https://doi.org/10.1146/annurev.marine.010908.163704.
Klinck, J., and W. Nowlin, 2001: Antarctic Circumpolar Current. Encyclopedia of Ocean Sciences, J. H. Steele, Ed., Elsevier, 151–159, https://doi.org/10.1006/rwos.2001.0370.
Langlais, C. E., A. Lenton, R. Matear, D. Monselesan, B. Legresy, E. Cougnon, and S. Rintoul, 2017: Stationary Rossby waves dominate subduction of anthropogenic carbon in the Southern Ocean. Sci. Rep., 7, 17076, https://doi.org/10.1038/s41598-017-17292-3.
Lapeyre, G., and P. Klein, 2006: Impact of the small-scale elongated filaments on the oceanic vertical pump. J. Mar. Res., 64, 835–851, https://doi.org/10.1357/002224006779698369.
Lapeyre, G., P. Klein, and B. L. Hua, 2006: Oceanic restratification forced by surface frontogenesis. J. Phys. Oceanogr., 36, 1577–1590, https://doi.org/10.1175/JPO2923.1.
Le Quéré, C., and Coauthors, 2007: Saturation of the Southern Ocean CO2 sink due to recent climate change. Science, 316, 1735–1738, https://doi.org/10.1126/science.1136188.
Lévy, M., P. J. S. Franks, and K. S. Smith, 2018: The role of submesoscale currents in structuring marine ecosystems. Nat. Commun., 9, 4758, https://doi.org/10.1038/s41467-018-07059-3.
Llort, J., C. Langlais, R. Matear, S. Moreau, A. Lenton, and P. G. Strutton, 2018: Evaluating Southern Ocean carbon eddy-pump from biogeochemical-Argo floats. J. Geophys. Res. Oceans, 123, 971–984, https://doi.org/10.1002/2017JC012861.
MacCready, P., and P. B. Rhines, 2001: Meridional transport across a zonal channel: Topographic localization. J. Phys. Oceanogr., 31, 1427–1439, https://doi.org/10.1175/1520-0485(2001)031<1427:MTAAZC>2.0.CO;2.
Mahadevan, A., 2006: Modeling vertical motion at ocean fronts: Are nonhydrostatic effects relevant at submesoscales? Ocean Modell., 14, 222–240, https://doi.org/10.1016/j.ocemod.2006.05.005.
Mahadevan, A., 2016: The impact of submesoscale physics on primary productivity of plankton. Annu. Rev. Mar. Sci., 8, 161–184, https://doi.org/10.1146/annurev-marine-010814-015912.
Mahadevan, A., and A. Tandon, 2006: An analysis of mechanisms for submesoscale vertical motion at ocean fronts. Ocean Modell., 14, 241–256, https://doi.org/10.1016/j.ocemod.2006.05.006.
McDougall, J. T., and P. M. Barker, 2011: Getting started with TEOS-10 and the Gibbs Seawater (GSW) oceanographic toolbox, version 3.0. Tech. Rep. SCOR/IAPSO WG127, 28 pp., https://www.teos-10.org/software.htm.
McWilliams, J. C., 2021: Oceanic frontogenesis. Annu. Rev. Mar. Sci., 13, 227–253, https://doi.org/10.1146/annurev-marine-032320-120725.
McWilliams, J. C., and M. J. Molemaker, 2011: Baroclinic frontal arrest: A sequel to unstable frontogenesis. J. Phys. Oceanogr., 41, 601–619, https://doi.org/10.1175/2010JPO4493.1.
McWilliams, J. C., F. Colas, and M. J. Molemaker, 2009: Cold filamentary intensification and oceanic surface convergence lines. Geophys. Res. Lett., 36, L18602, https://doi.org/10.1029/2009GL039402.
McWilliams, J. C., J. Gula, M. Jeroen Molemaker, L. Renault, and A. F. Shchepetkin, 2015: Filament frontogenesis by boundary layer turbulence. J. Phys. Oceanogr., 45, 1988–2005, https://doi.org/10.1175/JPO-D-14-0211.1.
Meijer, J. J., H. E. Phillips, N. L. Bindoff, S. R. Rintoul, and A. Foppert, 2022: Dynamics of a standing meander of the subantarctic front diagnosed from satellite altimetry and along-stream anomalies of temperature and salinity. J. Phys. Oceanogr., 52, 1073–1089, https://doi.org/10.1175/JPO-D-21-0049.1.
Meijers, A. J. S., N. L. Bindoff, and S. R. Rintoul, 2011: Estimating the four-dimensional structure of the Southern Ocean using satellite altimetry. J. Atmos. Oceanic Technol., 28, 548–568, https://doi.org/10.1175/2010JTECHO790.1.
Meunier, T., and J. H. LaCasce, 2021: The finite size Lyapunov exponent and the finite amplitude growth rate. Fluids, 6, 348, https://doi.org/10.3390/fluids6100348.
Miles, J. W., 1961: On the stability of heterogeneous shear flows. J. Fluid Mech., 10, 496–508, https://doi.org/10.1017/S0022112061000305.
Morrison, A. K., D. W. Waugh, A. M. C. Hogg, D. C. Jones, and R. P. Abernathey, 2022: Ventilation of the Southern Ocean pycnocline. Annu. Rev. Mar. Sci., 14, 405–430, https://doi.org/10.1146/annurev-marine-010419-011012.
Naveira Garabato, A. C., J. T. Allen, H. Leach, V. H. Strass, and R. T. Pollard, 2001: Mesoscale subduction at the Antarctic polar front driven by baroclinic instability. J. Phys. Oceanogr., 31, 2087–2107, https://doi.org/10.1175/1520-0485(2001)031<2087:MSATAP>2.0.CO;2.
Naveira Garabato, A. C., R. Ferrari, and K. L. Polzin, 2011: Eddy stirring in the Southern Ocean. J. Geophys. Res., 116, C09019, https://doi.org/10.1029/2010JC006818.
Naveira Garabato, A. C., K. L. Polzin, R. Ferrari, J. D. Zika, and A. Forryan, 2016: A microscale view of mixing and overturning across the Antarctic Circumpolar Current. J. Phys. Oceanogr., 46, 233–254, https://doi.org/10.1175/JPO-D-15-0025.1.
Naveira Garabato, A. C., and Coauthors, 2019: Rapid mixing and exchange of deep-ocean waters in an abyssal boundary current. Proc. Natl. Acad. Sci. USA, 116, 13 233–13 238, https://doi.org/10.1073/pnas.1904087116.
Okubo, A., 1970: Horizontal dispersion of floatable particles in the vicinity of velocity singularities such as convergences. Deep-Sea Res. Oceanogr. Abstr., 17, 445–454, https://doi.org/10.1016/0011-7471(70)90059-8.
Olbers, D., D. Borowski, C. Völker, and J.-O. Wölff, 2004: The dynamical balance, transport and circulation of the Antarctic Circumpolar Current. Antarct. Sci., 16, 439–470, https://doi.org/10.1017/S0954102004002251.
Orsi, H. A., T. Whitworth III, and W. D. Nowlin Jr., 1995: On the meridional extent and fronts of the Antarctic Circumpolar Current. Deep-Sea Res. I, 42, 641–673, https://doi.org/10.1016/0967-0637(95)00021-W.
Phillips, H. E., and S. R. Rintoul, 2000: Eddy variability and energetics from direct current measurements in the Antarctic Circumpolar Current south of Australia. J. Phys. Oceanogr., 30, 3050–3076, https://doi.org/10.1175/1520-0485(2000)030<3050:EVAEFD>2.0.CO;2.
Phillips, H. E., and N. L. Bindoff, 2014: On the nonequivalent barotropic structure of the Antarctic Circumpolar Current: An observational perspective. J. Geophys. Res. Oceans, 119, 5221–5243, https://doi.org/10.1002/2013JC009516.
Ridgway, K. R., J. R. Dunn, and J. L. Wilkin, 2002: Ocean interpolation by four-dimensional weighted least squares – Application to the waters around Australasia. J. Atmos. Oceanic Technol., 19, 1357–1375, https://doi.org/10.1175/1520-0426(2002)019<1357:OIBFDW>2.0.CO;2.
Rintoul, S. R., 2000: Southern Ocean currents and climate. Pap. Proc. Roy. Soc. Tasmania, 133, 41–50.
Rintoul, S. R., 2018: The global influence of localized dynamics in the Southern Ocean. Nature, 558, 209–218, https://doi.org/10.1038/s41586-018-0182-3.
Rintoul, S. R., C. W. Hughes, and D. Olbers, 2001: The Antarctic Circumpolar Current system. Ocean Circulation and Climate, G. Siedler, J. Church, and J. Gould, Eds., International Geophysics Series, Vol. 77, Academic Press, 271–302, https://doi.org/10.1016/S0074-6142(01)80124-8.
Roach, C. J., H. E. Phillips, N. L. Bindoff, and S. R. Rintoul, 2015: Detecting and characterizing Ekman currents in the Southern Ocean. J. Phys. Oceanogr., 45, 1205–1223, https://doi.org/10.1175/JPO-D-14-0115.1.
Rosso, I., A. M. Hogg, P. G. Strutton, A. E. Kiss, R. Matear, A. Klocker, and E. van Sebille, 2014: Vertical transport in the ocean due to sub-mesoscale structures: Impacts in the Kerguelen region. Ocean Modell., 80, 10–23, https://doi.org/10.1016/j.ocemod.2014.05.001.
Rosso, I., A. M. Hogg, A. E. Kiss, and B. Gayen, 2015: Topographic influence on submesoscale dynamics in the Southern Ocean. Geophys. Res. Lett., 42, 1139–1147, https://doi.org/10.1002/2014GL062720.
Sallée, J.-B., R. J. Matear, S. R. Rintoul, and A. Leyton, 2012: Localized subduction of anthropogenic carbon dioxide in the Southern Hemisphere oceans. Nat. Geosci., 5, 579–584, https://doi.org/10.1038/ngeo1523.
Sanchez-Rios, A., R. K. Shearman, J. Klymak, E. D’Asaro, and C. Lee, 2020: Observations of cross-frontal exchange associated with submesoscale features along the North Wall of the Gulf Stream. Deep-Sea Res. I, 163, 103342, https://doi.org/10.1016/j.dsr.2020.103342.
Sanford, T. B., 1971: Motionally induced electric and magnetic fields in the sea. J. Geophys. Res., 76, 3476–3492, https://doi.org/10.1029/JC076i015p03476.
Sanford, T. B., J. H. Dunlap, J. A. Carlson, D. C. Webb, and J. B. Girton, 2005: Autonomous velocity and density profiler: EM-APEX. Proc. IEEE/OES Eighth Working Conf. on Current Measurement Technology, Southampton, United Kingdom, IEEE, 152–156, https://doi.org/10.1109/CCM.2005.1506361.
Schubert, R., J. Gula, R. J. Greatbatch, B. Baschek, and A. Biastoch, 2020: The submesoscale kinetic energy cascade: Mesoscale absorption of submesoscale mixed layer eddies and frontal downscale fluxes. J. Phys. Oceanogr., 50, 2573–2589, https://doi.org/10.1175/JPO-D-19-0311.1.
Siegelman, L., 2020: Energetic submesoscale dynamics in the ocean interior. J. Phys. Oceanogr., 50, 727–749, https://doi.org/10.1175/JPO-D-19-0253.1.
Siegelman, L., M. O’Toole, M. Flexas, P. Rivière, and P. Klein, 2019: Submesoscale ocean fronts act as biological hotspot for southern elephant seal. Sci. Rep., 9, 5588, https://doi.org/10.1038/s41598-019-42117-w.
Siegelman, L., P. Klein, P. Rivière, A. F. Thompson, H. S. Torres, M. Flexas, and D. Menemenlis, 2020a: Enhanced upward heat transport at deep submesoscale ocean fronts. Nat. Geosci., 13, 50–55, https://doi.org/10.1038/s41561-019-0489-1.
Siegelman, L., P. Klein, A. F. Thompson, H. S. Torres, and D. Menemenlis, 2020b: Altimetry-based diagnosis of deep-reaching sub-mesoscale ocean fronts. Fluids, 5, 145, https://doi.org/10.3390/fluids5030145.
Sokolov, S., and S. R. Rintoul, 2009: Circumpolar structure and distribution of the Antarctic Circumpolar Current fronts: 2. Variability and relationship to sea surface height. J. Geophys. Res., 114, C11019, https://doi.org/10.1029/2008JC005248.
Su, Z., J. Wang, P. Klein, A. F. Thompson, and D. Menemenlis, 2018: Ocean submesoscales as a key component of the global heat budget. Nat. Commun., 9, 775, https://doi.org/10.1038/s41467-018-02983-w.
Su, Z., H. Torres, P. Klein, A. F. Thompson, L. Siegelman, J. Wang, D. Menemenlis, and C. Hill, 2020: High-frequency submesoscale motions enhance the upward vertical heat transport in the Global Ocean. J. Geophys. Res. Oceans, 125, e2020JC016544, https://doi.org/10.1029/2020JC016544.
Swart, S., and Coauthors, 2023: The Southern Ocean mixed layer and its boundary fluxes: Fine-scale observational progress and future research priorities. Philos. Trans. Roy. Soc., A381, 20220058, https://doi.org/10.1098/rsta.2022.0058.
Tamsitt, V., and Coauthors, 2017: Spiraling pathways of global deep waters to the surface of the Southern Ocean. Nat. Commun., 8, 172, https://doi.org/10.1038/s41467-017-00197-0.
Taylor, J. R., and R. Ferrari, 2010: Buoyancy and wind-driven convection at mixed layer density fronts. J. Phys. Oceanogr., 40, 1222–1242, https://doi.org/10.1175/2010JPO4365.1.
Taylor, J. R., and A. F. Thompson, 2023: Submesoscale dynamics in the upper ocean. Annu. Rev. Fluid Mech., 55, 103–127, https://doi.org/10.1146/annurev-fluid-031422-095147.
Taylor, J. R., S. Bachman, M. Stamper, P. Hosegood, K. Adams, J.-B. Sallee, and R. Torres, 2018: Submesoscale Rossby waves on the Antarctic Circumpolar Current. Sci. Adv., 4, eaao282, https://doi.org/10.1126/sciadv.aao2824.
Thomas, L. N., A. Tandon, and A. Mahadevan, 2008: Submesoscale processes and dynamics. Ocean Modeling in an Eddying Regime, Geophys. Monogr., Vol. 177, Amer. Meteor. Soc., 17–38, https://doi.org/10.1029/177GM04.
Thomas, L. N., J. R. Taylor, R. Ferrari, and T. M. Joyce, 2013: Symmetric instability in the Gulf Stream. Deep-Sea Res. II, 91, 96–110, https://doi.org/10.1016/j.dsr2.2013.02.025.
Thompson, A. F., 2008: The atmospheric ocean: Eddies and jets in the Antarctic Circumpolar Current. Philos. Trans. Roy. Soc., A366, 4529–4541, https://doi.org/10.1098/rsta.2008.0196.
Thompson, A. F., and J.-B. Sallée, 2012: Jets and topography: Jet transitions and the impact on transport in the Antarctic Circumpolar Current. J. Phys. Oceanogr., 42, 956–972, https://doi.org/10.1175/JPO-D-11-0135.1.
Thompson, A. F., and A. C. Naveira Garabato, 2014: Equilibration of the Antarctic Circumpolar Current by standing meanders. J. Phys. Oceanogr., 44, 1811–1828, https://doi.org/10.1175/JPO-D-13-0163.1.
Thompson, A. F., A. Lazar, C. Buckingham, A. C. Naveira Garabato, G. M. Damerell, and K. J. Heywood, 2016: Open-ocean submesoscale motions: A full seasonal cycle of mixed layer instabilities from gliders. J. Phys. Oceanogr., 46, 1285–1307, https://doi.org/10.1175/JPO-D-15-0170.1.
Weiss, J., 1991: The dynamics of enstrophy transfer in two-dimensional hydrodynamics. Physica D, 48, 273–294, https://doi.org/10.1016/0167-2789(91)90088-Q.
Youngs, M. K., A. F. Thompson, A. Lazar, and K. J. Richards, 2017: ACC meanders, energy transfer, and mixed barotropic-baroclinic instability. J. Phys. Oceanogr., 47, 1291–1305, https://doi.org/10.1175/JPO-D-16-0160.1.
Yu, X., A. C. Naveira Garabato, A. P. Martin, C. E. Buckingham, L. Brannigan, and Z. Su, 2019: An annual cycle of submesoscale vertical flow and restratification in the upper ocean. J. Phys. Oceanogr., 49, 1439–1461, https://doi.org/10.1175/JPO-D-18-0253.1.
Zhang, X., M. Nikurashin, B. Peña-Molino, S. R. Rintoul, and E. Doddridge, 2023a: A theory of standing meanders of the Antarctic Circumpolar Current and their response to wind. J. Phys. Oceanogr., 53, 235–251, https://doi.org/10.1175/JPO-D-22-0086.1.
Zhang, Z., and Coauthors, 2023b: Submesoscale inverse energy cascade enhances Southern Ocean eddy heat transport. Nat. Commun., 14, 1335, https://doi.org/10.1038/s41467-023-36991-2.
