1 Late winter oceanography beneath East Antarctic sea
2 ice during SIPEX
a,b,∗ d,b d,c e a,b 3 G.D. Williams , A.J.S. Meijers , A. Poole , P. Mathiot , T. Tamura , b 4 A. Klocker
a 5 Institute for Low Temperature Science, Hokkaido University, Sapporo, Japan. b 6 Antarctic Climate and Ecosystem Cooperative Research Centre, Sandy Bay, Australia. c 7 Department of the Environment, Water, Heritage and the Arts, Australian Antarctic 8 Division, Kingston, Australia. d 9 Centre for Australian Weather and Climate Research, CSIRO, Hobart, Australia. e 10 Laboratory of Geophysical and Industrial Flow, Grenoble, France
11 Abstract
We report on the late winter oceanography observed beneath Antarctic sea ice offshore from the Sabrina and BANZARE coast of Wilkes Land, East Antarctica (115–125◦E) in September–October 2007 during the Sea Ice Physics and Ecosys- tem eXperiment (SIPEX) research voyage. A pilot program using specifically designed ’through-ice’ Conductivity-Temperature-Depth (CTD) and acoustic Doppler current profiling (ADCP) systems was conducted to opportunistically measure water mass properties and ocean currents during major ice stations. Additional water mass properties across the survey region were collected from Ice-Argo floats deployed during the voyage north of the 3000m isobath. The mean drift of the floats between deployment and the first surface download in summertime was along slope to the west with the Antarctic Slope Current. Vertical profiles of potential temperature reveal the deepest winter mixed layer (WML) in the western sector of the survey northwest of the Dalton Iceberg Tongue polynya. The meridional structure of the Antarctic Slope Front, i.e. the offshore shoaling of the WML across the upper continental slope, is found to be the simillar to previous summer observations. South of the continen- tal shelf break a strong bottom intrusion of modified Circumpolar Deep Water (mCDW) as warm as 0◦C was detected beneath the fast ice near 118◦E. This is in the vicinity of a summertime observation of a shoreward intrusion of mCDW
∗Corresponding author: [email protected] Preprint submitted to Elsevier March 10, 2010 of similar magnitude over ten years earlier. We hypothesise that this strengthens the liklihood that there is a persistent supply of ocean heat to the shelf region and that this is the primary cause of recently reported enhanced melting of the nearby Totten Glacier and Moscow University Ice Shelf. Interestingly there was no detection of locally formed dense shelf water capable of forming Antarctic Bottom Water at the shelf break locations sampled despite the number of minor polynyas across this region. Ocean current measurements, limited to a maxi- mum period of 24 hours and 50–100m depth by the relative scarcity of backscat- ter, found increased mean vertical speeds at the offshore stations (6–17 cm s−1) relative to the shelf break (2.3–6.4 cm s−1cm s( − 1)). The diurnal variation in the ADCP range suggests that the diel migration of zooplankton was occur- ring beneath the sea ice in late winter, with greater range/abundance offshore. Analysis of concurrent time series of wind, ocean current and their influence on sea ice drift from Global Positioning System (GPS) compass measurements were computed but the length of data acquisitions limited the applicability of this analysis. Overall the pilot ’through-ice’ program was a success and with improvements in the accuracy of instrumentation, and a strategic survey plan, promises to provide invaluable observations for model comparison outside of the current datasets available.
1 Keywords: SIPEX, Under-ice, Oceanography, East Antarctica
2 1 1. Introduction
2 The seasonal formation of sea ice around Antarctic each year is a primary
3 mechanism for water mass transformations in the polar Southern Ocean. Brine-
4 rejection from sea ice production promotes convection and the formation of a
5 deep winter mixed layer in the Antarctic Surface Water (AASW). In discrete lo-
6 cations on the Antarctic continental shelf, enhanced sea ice formation and brine
7 rejection increases the density of shelf waters, such that if transported north-
8 wards with sufficient negative buoyancy to mix down the continental slope,
9 Antarctic Bottom Water (AABW) forms. Oceanographic observations around
10 Antarctica are sparse due to its remote location and this is accentuated dur-
11 ing winter months when sea ice presents additional logistic challenges. Aus-
12 tralian Antarctic research expeditions conducted in winter are predominantly
13 field-based sea-ice studies, i.e. a series of ’ice-stations’ ex-ship, that do not allow
14 for standard ship-based oceanography due to time and safety contraints. For
15 this reason a pilot field based ’through-ice’ oceanography program was under-
16 taken during the Sea Ice Physics and Ecosystem eXperiment (SIPEX) voyage
17 (Worby et al., 2010, this volume) with the goal of sampling the water mass prop-
18 erties and currents beneath the late winter sea ice around the East Antarctic
19 margin offshore from the Sabrina and BANZARE coast of Wilkes Land between ◦ 20 115–130 E (Figure 1).
21 [Figure 1 about here]
22 1.1. The Antarctic Margin of the Sabrina and BANZARE Coast, Wilkes Land
23 The continental shelf region of the Sabrina and BANZARE coasts of Wilkes
24 Land has a combination of floating glaciers, iceshelves and grounded iceberg
25 tongues, most notably the Totten Glacier and Moscow University Ice Shelf,
26 together with a series of persistent polynya regions. These range from Cape
27 Poinsett in the west to the Dibble Iceberg Tongue in the east and are labelled
28 in Figure 1 following Massom (2003) with estimates of total sea ice produc- −2 29 tion in 2007 (m m ). These sea ice production estimates come from SSM/I
3 1 satellite data with the ERA-Interim reanalysis (ECMWF, 2009) and the heat
2 flux algorithm of Tamura et al. (2007, 2008). These are all examples of recur-
3 ring latent-heat polynyas which form due to ice divergence in the lee of topo-
4 graphic/ice barriers over the continental shelf. The exception is the Voyeykov
5 polynya, which is a ’deep water’ polynya (?) that forms along the fast ice edge ◦ 6 near the shelf break between 122–125 E.
7 The previous observations in this region are limited to the BROKE (Baseline
8 Research on Oceanography Krill and the Enviroment) survey in the austral
9 summer of 1996 (Bindoff et al., 2000), as shown in Figure 1. Vertical sections of
10 potential temperature for the upper 1000m of BROKE meridional transects at ◦ 11 112, 120 and 128 E are shown in Figure 2. These sections demonstrate the key
12 water masses, boundaries and fronts in summertime and their zonal variability
13 across the SIPEX survey region. In this paper we follow definitions used by
14 Orsi et al. (1995); Whitworth et al. (1998) and recently updated by ?. Offshore n 15 of the shelf break, the two major water masses are cold, fresh AASW (γ ¡ n 16 28.00 kg m−3) above warm, saline Circumpolar Deep Water (CDW, 28.00 ¡ γ
17 ¡ 28.27 kg m−3). The CDW properties weaken polewards through mixing with ◦ 18 cold Antarctic water masses and when cooler than 1.5 C (also known as the
19 Southern Boundary) it is referred to as ’modified’ CDW (mCDW).
20 Over the upper continental slope and continental shelf break this boundary
21 between AASW and mCDW is termed the Antarctic Slope Front (ASF, Jacobs
22 (1991); Whitworth et al. (1998)). There is a broad westward flow, termed the
23 Antarctic Slope Current (ASC), associated with the horizontal desnity gradi-
24 ent across the ASF, that often includes a narrow, fast-flowing jet (up to 50 cm −1 25 s ) topographically pinned between the 750–1250 m isobaths. Though some-
26 times co-located, the ASC is not to be confused with Antarcitc Coastal Current
27 (ACoC), which is associated with the East Wind Drift, shallower and has a more
28 coastal domain. In this paper we will concentrate on the ASC as this is typi- ◦ ◦ 29 cal of the circulation around the East Antarctica margin from 30 E to 150 E
30 (Bindoff et al., 2000; Aoki, 2003; Williams et al., 2008b; Meijers et al., 2010;
31 Williams et al., 2010a) and indeed much of the Antarctic margin away from the
4 1 Weddell Sea and Antarctic peninsula (Fahrbach et al., 1992; Whitworth et al.,
2 1998; Heywood et al., 1998; ?).
3 The vertical structure of the AASW varies seasonally and is shown in Figure
4 2 with a warm, fresh seasonal summer mixed layer (SML) at the surface that
5 forms from the fresh water lens post sea ice melt. The SML deepens through the
6 summer by wind mixing into the temperature minimum layer that is the remnant
7 from the previous winter’s permanent mixed layer (referred to hereafter as the
8 winter mixed layer, WML). The WML forms by convection, driven initially by
9 atmospheric cooling of the surface SML and then by the brine rejection that
10 occurs with sea ice formation. Inshore of the shelf break, in regions of sufficient
11 sea ice formation/brine rejection, the WML extends to the bottom, mixing away
12 all remnants of the summer stratification and reconnects with the base of the n 13 previous year’s WML. If denser than γ = 28.27 kg m−3 and colder than θ ◦ 14 = -1.7–-1.85 C, this new winter water mass is referred to as Shelf Water(SW).
15 Offshore of the continental shelf break, brine rejection is relatively weak and the
16 WML/AASW remains fresher and lighter that the mCDW. In specific locations
17 mCDW can upwell and migrate/intrude onto the continental shelf, transporting
18 heat to the Antarctic coastal domain which has major implications for ocean/ice
19 shelf interactions (Smethie and Jacobs, 2005). The mCDW intrusions are also
20 important for the marine ecosystem if they upwell into the surface layer as they
21 are rich in nutrients. These intrusion are also reportedly an important part
22 of the lifecycle of Antarctic krill, transporting eggs from offshore (Sala et al.,
23 2002). However the influence of mCDW on the continental shelf is ultimately
24 dependent on the local stratification. For example, if SW is present, then by
25 definition mCDW can only occupy the layer above.
26 [Figure 2 about here]
27 The striking feature of the BROKE transects in this region is the south- ◦ ◦ 28 ward migration of the warmer mCDW at 120 E, bringing the sub-surface 0 C
29 isotherm onto the shelf break at station 83 . This was reported as the strongest ◦ 30 mCDW signal of the entire BROKE survey between 80–150 E. This is in broad
5 1 agreement with the southward shift in the Southern Boundary of the Antarctic ◦ 2 Circumpolar Current described by Orsi et al. (1995) east of 115 E (see dashed
3 line in Figure 1). However the BROKE results showed there is more zonal
4 variability in the properties of the mCDW layer, as indicated by the southern ◦ ◦ 5 extent of the 1 C isotherm at 120 E in Figure 2. The other major zonal vari-
6 ation across these transects was in the surface layer over the continental shelf ◦ 7 and the weakly developed seasonal mixed layer at 112 E (i.e. it was observed
8 north of station 72). This was not the result of any temporal lag in the sampling
9 as the transects were conducted from east to west and instead implies greater
10 production and/or persistence of sea ice, which has been found to delay sum-
11 mertime mixed layer development in other regions of East Antarctica (Williams
12 et al., 2008b).
13 In this paper we will present unique late winter observations of the water
14 mass properties beneath sea ice on the continental shelf break (0–550m) and up- ◦ 15 per layer (0–1000m) offshore of the East Antarctic margin between 117–128 E.
16 In addition concurrent measurements of near-surface ocean currents, sea ice
17 drift and surface wind speeds collected at SIPEX ice stations are presented.
18 In section 2 we describe the oceanographic survey and introduce the two main
19 observation platforms, namely a ’through-ice’ Conductivity-Temperature-Depth
20 (CTD) and acoustic Doppler current profiling (ADCP) system, together with
21 Ice-Argo floats deployed on behalf of the University of Washington. In Section 3
22 we present our observations, in particular vertical profiles of potential tempera-
23 ture from the CTD system and Ice-Argo in a description of the ’re-sampling’ of ◦ 24 the mCDW intrusion on the shelf break near 120 E and the spatial variation in
25 the winter mixed layer depth across the survey. This is followed by the vertical
26 profiles of current speed and direction immediately beneath the sea ice at both
27 offshore and fast-ice stations. We analyse the concurrent time series of ocean
28 current and surface wind speed and direction to investigate how each component
29 influences sea ice drift. In Section 4 we discuss the impact of local polynyas and
30 ice shelves on the water mass properties observed in this study and the overall
6 1 success and future of the ’through-ice’ oceanography program.
2 2. Data and Methods
3 This project involved two independent sub-ice observation platforms: A
4 winch-driven Conductivity-Temperature-Depth system for measuring basic wa-
5 ter mass properties and an acoustic Doppler current profiling (ADCP)/GPS
6 system for measuring ocean currents and ice drift. Hereafter these are referred
7 to as the CTD and ADCP systems respectively (Figure 3). Additional water
8 mass data were collected by Ice Argo floats deployed during SIPEX for the
9 University of Washington.
10 [Figure 3 about here]
11 2.1. Ice Argo floats from the University of Washington
12 The Ice-Argo floats deployed during SIPEX are the new generation of Argo
13 float that are capable of autonomously collecting CTD and dissolved oxygen
14 profiles beneath sea-ice. These floats are designed to abort their ascent 10-30m
15 from the surface when sea ice cover is detected by a cold near-surface layer (Klatt
16 et al., 2007). Deployed within the sea-ice field north of the 3000-m isobath, the
17 floats profile the water column (every 2m to 2000m depth) every 7 days without
18 resurfacing, using on-board data storage until the data can be uploaded after
19 the sea-ice melts. Depending on voyage logistics, the floats were either deployed
20 from the back of the ship, through a hole in the ice during ice stations, or from
21 a helicopter into a lead when the ship was south of the 3000-m isobath. The
22 Ice-Argo data were extracted from the publicly-available data portal operated
23 by the National Oceanographic Data Centre (NODC, 2009). We use an abbre-
24 viated World Meteorological Organisation (WMO) naming format, i.e. WMO
25 float number 2900130 is referred to in this paper as a130. Data is expected ◦ 26 to have accuracy of 0.001 C, 0.01 and 5m for temperature, salinity and depth,
27 respectively (Office, 2010). Further information about the Ice-Argo instruments
7 1 and data processing can be accessed through the University of Washington (UW,
2 2009).
3 [Figure 4 about here]
4 Figure 4 shows the deployment locations of Ice-Argo floats during SIPEX,
5 and their first ’download’ position, with the mean speed necessary to cover this
6 distance. The broad pattern for the time mean drift of the floats is one of along
7 slope to the west with the Antarctic Slope Current. The floats deployed in the
8 western sector of the survey were transported west, but with a greater northward
9 component, again following the bathymetry. The one exception to this was a
10 float that surfaced near Cape Poinsett. Of course, the exact trajectories and
11 drift speeds of these floats whilst under sea-ice is unknown.
12 2.2. Conductivity-Temperature-Depth System
13 The CTD system comprised of an Falmouth Scientific Institute (FSI) CTD
14 instrument, a tripod and over 1000m of polyethylene rope on a winch/drum
15 attached to a metal sled (see Figure 3). The winch system consisted of a control
16 box and was powered by a portable generator. The FSI was itself housed in
17 a protective fibreglass case to shield the sensors from rough treatment in the
18 sub-ice environment. Initial deployments demonstrated the sensitivity of the
19 instrument to the cold water and so the deployment procedure thereafter began
20 by taking the instrument down to approximately 50 m and leaving it there for
21 3–5 minutes before returning to near the surface and beginning the cast. The −1 22 rate of descent was kept below 1 m s on the way down. Post-deployment the
23 FSI was returned to the ship and the data downloaded and assessed to ensure
24 the data were copacetic. An interesting side-effect of the CTD system was that
25 the ’noise’ from the speed controller box effectively blocked out the ships HF
26 radio communications. As a result, the CTD system could not be operated
27 when helicopter operations were being conducted within 20 km of the ship.
28 Overall, the conductivity/salinity data collected by the FSI were poor (for
29 further details, please consult Rosenberg (????)). In the later stages of SIPEX,
8 1 a shipboard CTD was deployed (Seabird SBE911+ CTD) with the FSI attached
2 for calibration. The vertical profiles from this cast are shown in the top row
3 of Figure 5 and clearly show a large offset in salinity (> 0.1). The first profile
4 from the nearest Ice-Argo float (a120) show that the float performed better
5 than the FSI, closely matching the Seabird profile. An attempt to calibrate
6 all FSI profiles using the combined FSI/Seabird result was conducted with the
7 results again compared to nearby Ice-Argo profiles. This appeared successful
8 for Station 2 (float a130), in conjunction with a nearby profile from BROKE
9 (Stn 109), but not thereafter, with Stations 10 and 11 showing the ’calibrated’
10 salinity data as still much fresher than at the nearest Ice-Argo data (floats a114
11 and a125 respectively). Similar residual offsets in salinity post-calibration were
12 found when the calibrated salinity profiles at Station 8 were compared at with
13 nearby salinity data from BROKE (Stn 82). These offsets in salinity are simply
14 too large, in particular at depth below the seasonally variable surface layer, for
15 this to be accounted for by spatial or temporal variability.
16 [Figure 5 about here]
17 Upon completion of the SIPEX voyage, further testing of the FSI instrument
18 revealed that the conductivity sensor was highly sensitive to the protective fi-
19 breglass casing it was deployed in, and that minor movements (1–2mm) in the
20 orientation of the instrument relative to the casing during battery changes were
21 likely responsible for the offset and variable behavior during the voyage. For
22 this reason we dismiss the absolute value of salinity from this instrument, but
23 present it here nonetheless for its vertical structure, in particular for the stations
24 on the shelf break (lower panel, Figure 5) that were not sampled by Ice-Argo.
25 2.3. Acoustic Doppler Current Profiling System
26 The acoustic Doppler current profiling (ADCP) system was designed and
27 constructed by Marine Science Support at the Australian Antarctic Division
28 with the goal of observing the ocean current regime beneath the sea ice. To de-
29 termine the absolute ocean current, a GPS compass was included in the system
9 1 to account for sea ice drift. The system consisted of a three-legged support-
2 ing platform that included a central, downward facing retractable pylon for the
3 RDI Workhorse ADCP operating at 307.2 kHz, a flat tray for storing the power
4 supply/control system, mounting points for the radio modem antenna and an
5 upward facing extendable pylon for the Furuno SC50 GPS compass (see Figure
6 3). The observation site was preferentially in the zone perpendicular to the port
7 bow of the ship, between 50–100 m away to minimise any contamination of the
8 flow field from the ships bow thrusters.
9 At each station a hole through the sea ice was prepared and the ADCP
10 lowered up to 1.5 m, being fully submerged and clear of the base of the ice.
11 The ADCP, GPS and radio modem were connected to the 24 volt recharge-
12 able portable battery through a control box. The system was designed so the
13 Furuno GPS compass and ADCP remained in fixed position relative to each
14 other, and the entire system was positioned parallel to the ships heading. The
15 basic procedure began with the initialisation of the GPS compass and establish-
16 ing of communication with the ADCP computer on board the ship via the radio
17 modem. Thereafter the ADCP was initialised from the ship using the RDI Win-
18 River software. After initial testing determined a lack of available backscatter in
19 the water column, attributed to the scarcity of biological matter in late winter,
20 the sampling bin size was set to 8 m to achieve a maximum range of 50–100 m
21 with a 1.76 m blanking window below the transducer head. This subsequently
22 ruled out any small-scale sampling of the immediate boundary layer below the
23 ice. The ADCP collected single ping ensembles every 0.5 seconds, operating
24 using the RDI default settings.
25 The ADCP data were post-processed during the voyage using the RDI Win-
26 River software. This synchronised the GPS heading data with the ADCP ve-
27 locity data streams, producing velocity headings relative to true north and the
28 ADCP position, and removed bad data points where data were below the beam
29 correlation threshold or the water velocity exceeded the threshold of 107 cm −1 30 s . The ice mounted ADCP frame ensured that ADCP tilt was fixed near 0
31 degrees and all four ADCP beams were required to produce solutions. Ensemble
10 1 averaging over three minutes (360 pings) reduced the rms error to below 0.5 cm −1 −1 2 s for all stations except station 3, where it is below 1 cm s . The water
3 velocities relative to the ADCP were converted to absolute u and v velocities
4 by adding the ADCP (ice) drift velocity components calculated based on the 90
5 second boxcar filter smoothed GPS compass position. Due to the high clarity
6 of the water column from the relative absence of biological backscatter in late
7 winter, the ADCP maximum range was limited to between 50-100 dbar, with
8 a strong diurnal variation. All stations exhibited strong vertical coherence in −1 9 velocity with standard deviations of the shear of less than 1–5 cm s absolute
10 velocity (Table 1). Because of the relatively narrow range of depths sampled
11 and the strong vertical coherence observed, we use the calculated vertical mean
12 when referring to horizontal components of velocity.
13 [Table 1 about here]
14 2.3.1. Ocean Tides
15 As the ADCP time series were too short (maximum of 24 hours) to meaning-
16 fully estimate the tidal contribution to currents through frequency analysis, tidal
17 velocities were estimated using the Circum-Antarctic Tidal Simulation (CATS -
18 vers. 2008b) (Padman et al., 2002). Subtracting the tidal predictions from the
19 ADCP data only successfully reduced the variance of the u and v components
20 of velocity at stations 2, 8, 13 and 14, suggesting that the phase of the tide
21 model is inaccurate in the SIPEX region. This inaccuracy may also be caused
22 by non-tidal currents being in phase with the tidal signal, but without a longer
23 time series of observations this ambiguity cannot be resolved. We therefore only
24 use the tide model as a gauge of the tidal velocity magnitudes in this analysis.
25 Table 1 shows that north of the shelf break tidal velocities are approximately an
26 order of magnitude smaller than the ADCP mean current absolute velocities,
27 while over the shelf at the shallower stations, the tidal currents are of a similar
28 magnitude to the mean ADCP velocity and may significantly influence it. The
29 meridional component of station 8 in particular exhibits a in particular exhibits
30 a fluctuation over its duration (11.52 hours) that would be consistent with tidal
11 1 influence.. However, without significantly longer observations or an accurate
2 tide model we are unable to unambiguously separate out the tidal component
3 from the ADCP velocities and the velocities are left unmodified in this study.
4 3. Results
5 3.1. Water Mass Properties from ’through-ice’ CTD and Ice-Argo
6 We begin with the water mass properties observed during SIPEX and focus
7 on two main results: the detection of Modified Circumpolar Deep Water on the
8 continental shelf across fast-ice stations 5, 6 and 8, and the large-scale vertical
9 stratification, i.e. the winter mixed layer depth, across the offshore stations and
10 Ice-Argo profiles.
11 3.1.1. Modified Circumpolar Deep Water on the continental shelf north of the
12 Totten Glacier
13 The most significant oceanographic result from SIPEX is the detection of
14 mCDW on the shelf break at station 8. Figures 5 and 6 show that the bottom ◦ 15 layer of Station 8 had a 50-70m layer of mCDW that is warmer (up to 2 C)
16 and more saline (at least 0.2) relative to the water column above, and that this
17 is consistent with the mCDW layer previously recorded in summertime during
18 BROKE at station 82. The strength of the warming signal implies that the
19 Antarctic Slope Front is on or south of the shelf break, as in the western Ross
20 Sea (Whitworth III and Orsi, 2006; Muench et al., 2009). That this feature was
21 observed during SIPEX in late winter, more than ten years after the summertime
22 observation during BROKE, implies this intrusion is likely to be a persistent
23 feature.
24 [Figure 6 about here]
25 In other regions of East Antarcrtica, most notably the Ad´elieand George V
26 Land continental shelf break, mCDW intrusions are much weaker in terms of ◦ 27 their warming signal relative to the 0 C observed here (?), in particular during
12 1 winter when the mCDW is almost unidentifiable due to strong convective mixing
2 during dense shelf water formation (Williams and Bindoff, 2003). Additionally
3 the mCDW intrusions for the AGV region are lighter than the local shelf water
4 and penetrate the continental shelf water column at mid-depths (300–400m).
5 This is important because this means the heat flux associated with the mCDW
6 cannot directly influence ocean/ice shelf interactions at the grounding lines (>
7 1000m, Rignot and Jacobs (2002)) of the Mertz and Ninnis Glacier in that
8 region. We note that for the Sabrina coast the mCDW intrusions observed
9 during both BROKE and SIPEX occupy the bottom layer of the observed water
10 mass profiles. This is significant for two reasons. It suggests firstly that there
11 was no dense shelf water present in this location at the time of sampling and
12 secondly that when mCDW is the densest water mass, it can influence a greater
13 depth range of the nearby Totten Glacier which has an even deeper grounding
14 line than the Mertz and Ninnis glaciers Rignot and Jacobs (2002).
15 3.1.2. Winter Mixed Layer Depth over the Continental Slope and Rise
16 As introduced earlier, the winter mixed layer depth (WML) defining the
17 base of the AASW layer is the product of the current season’s sea ice produc-
18 tion and the local mesoscale circulation on the continental shelf. It is observed
19 to exhibit a meridional structure from summertime surveys, shoaling north of
20 upper continental slope to form the horizontal density gradient of the Antarctic
21 Slope Front and generate the baroclinic Antarctic Slope Current. One of the
22 goals of SIPEX was to firstly determine if this meridional structure exists im-
23 mediately after the sea ice production season, and secondly to examine whether
24 any large-scale variability in the winter mixed layer depth could be attributed
25 to the spatial variability of sea-ice production across the survey region.
26 [Figure 7 about here]
27 We separate the sampling stations into eastern and western sectors of the
28 SIPEX survey and start by presenting vertical profiles of potential temperature
29 for the eastern sector in Figure 7. The WML is relatively shallow (∼100m)
13 1 at the northern stations, i.e. a127, 3 and 2 and then there is a transition to
2 the deeper WML (from ∼100 to sim200m) across the Antarctic Slope Front
3 between the initial profiles for Ice-Argo float a130 and a115 to a117, northeast
4 of the Voyeykov polynya. The profile from Station 2 is unique in that there is a
5 double-step across the permanent thermocline separating the base of the WML
6 and the mCDW below, indicative of advection.
7 [Figure 8 about here]
8 [Figure 9 about here]
9 The vertical profiles from the western sector of the survey (see Figure 8)
10 show a deepening of the WML depth, relative to the eastern sector, to more than
11 300m. The deepest WML are at FSI-CTD stations 11, 13 and 14. The spatial
12 variability in both WML depth and sub-surface Tmax below the WML during
13 late winter and summer from SIPEX CTD and Ice-Argo respectively is shown
14 in Figure 9. There is a shift to shallower WML depths from south to north,
15 similar to that observed in summer (Figure 2), in relation to the meridional
16 barotropic fields associated with the Antarctic Slope Current. The region of
17 deepest WML, extending furthest offshore from the shelf break is found between ◦ 18 115–120 E. This is northwest of the Dalton Iceberg Tongue polynya region and
19 suggests a link to the ice production therein. However the northern extent of
20 the ’deep’ WML is most likely due to the northward shift in the bathymetry
21 in that region deflecting the Antarctic Slope Current to the north-west. The
22 spatial distribution of sub-surface Tmax below the WML also clearly indicates
23 the southern migration of the warmest mCDW/CDW east of this region at ◦ 24 120 E. The lack of true repeat sampling between BROKE and SIPEX disallows
25 definitive comparison of seasonal and decadal variation.
26 3.2. Ocean Currents from the ’Through-Ice’ ADCP System
27 3.2.1. Offshore stations
28 Of the observations made north of the shelf break, over bathymetry of 1500-
29 4000 m, three stations (2, 10 and 14) were conducted for periods close to 24
14 1 hours, while the remaining three (1, 3 and 13) all lasted for less than seven hours.
2 All stations are highly coherent in the vertical with RMS differences between
3 the mean heading and the heading at individual bin depths being between 5–
4 12 degrees for all stations. The influence of tides is small north of the shelf −1 5 break (mean magnitude of order 1 cm s , Table 1) and their removal does not
6 influence the heading or magnitude of the currents observed at these stations
7 significantly. All statistics and figures show the observations with the tidal
8 component included.
9 [Table 1 about here]
10 [Figure 10 about here]
11 The column mean speed and heading of the water beneath the sea ice at
12 the long period stations are shown in Figure 10. The top 100m below the sea −1 13 ice tends to be relatively fast (mean magnitude 6–17 cm s ), with maximum −1 14 speeds of over 30 cm s at stations 2, 3 and 14. At the three stations occupied
15 for longer than 7 hours the currents vary strongly on time scales of a few hours
16 in both magnitude and direction. This is particularly apparent at stations 10
17 and 14 where the flow reverses direction over a period of less than five hours.
18 Similar sharp changes in flow direction occur at station 3, although the change
19 in heading is not as dramatic. In almost all cases these changes occur over
20 relatively short times and are separated by longer periods where the heading and
21 velocity is relatively steady. This non-continuous nature of the ocean transport
22 heading, in combination with the fact that changes are almost universally in a
23 clockwise sense argues against these being the result of inertial oscillations of
24 the water column. −1 −1 25 A mean westward velocity of 3-33 cm s (variance 1.3–9.7 cm s ) is ob-
26 served at all of the offshelf stations, which is consistent with other hydrographic
27 observations (Bindoff et al., 2000) and the presence of the ASC north of the
28 shelf break. Periods of eastward transport do occur at stations 3, 10 and 14,
29 but with the exception of station 14, all are relatively brief. Less consistency is
15 −1 1 apparent in the meridional component (variance 2.6–18.1 cm s ) and the long
2 period stations all exhibit periods of both northward and southward transport −1 3 with magnitudes greater than 10 cm s . However, the northward transport
4 tends to dominate and the net flux over the observed periods is northwards at
5 all stations except station 10, where it is southward (mean meridional velocity −1 6 -4 cm s ).
7 3.2.2. Fast ice stations near the continental shelf break
8 The stations occupied over the shelf (<800 m) were distinctly different from
9 those north of the shelf break. Smaller currents were encountered (2.3–6.4 cm −1 10 s ) and tidal velocities were stronger with mean magnitudes of up to 3 cm −1 −1 11 s and peak velocities of over 6 cm s . The mean current direction was also
12 more variable. While offshelf currents were broadly northwestward, there was
13 no consistent mean heading between the three onshelf stations, although rela- −1 14 tively consistent eastward currents of 1.6–5.8 cm s were observed at stations
15 6 and 8. Additionally, the onshelf current meters are significantly less vertically
16 coherent than those offshelf, with a RMS heading difference between the mean
17 water column heading and individual bin headings of 22–40 degrees. This re-
18 duction of vertical coherence may be a result of the generally weaker currents
19 and reduced signal to noise ratios. The presence of tidal Ekman spirals may
20 also reduce vertical coherence but the short period of observations (< 18 hours)
21 and relatively large ADCP bin sizes (8 m) means there is insufficient data to
22 resolve such vertical structures reliably.
23 [Figure 11 about here]
24 3.2.3. ADCP range as proxy for the diel migration and relative abundance of
25 zooplankton during SIPEX
26 Previous studies have used ADCP backscatter to detect zooplankton(Cresswell
27 et al., 2009; Kaneda et al., 2002). For the SIPEX results, the clear diurnal
28 variation in ADCP range invites a qualitative assessment, albeit extremely sim-
29 plified, of the diel migration of zooplankton beneath sea-ice in this region in
16 1 late-winter. Figure 12 shows the diel variation in ADCP backscatter and rela-
2 tive distribution of maximum ADCP range across the SIPEX survey. Given the
3 basic assumption that ADCP range is directly proportional to the presence of
4 zooplankton, in particular in late-winter before the sea ice melts and contributes
5 detritus to the backscatter of the water columns, then we find that diel migra-
6 tion of zooplankton was occurring beneath sea ice in late winter. The presence
7 of zooplankton was confirmed by direct observations/sampling through the ice
8 holes and camera surveys conducted after each ’through-ice’ CTD. If we make
9 the further assumption that the species of zooplankton and associated acoustic
10 backscatter was constant, then we also find that there was greater abundance of
11 zooplankton at the offshore stations relative to the stations on the shelf break
12 beneath the fast-ice.
13 [Figure 12 about here]
14 3.3. Wind, Ocean Current and Sea Ice Drift
15 Sea ice drift is the result of the summation of the forces acting upon it,
16 most notably the wind, ocean currents and coriolis force, and to a less extent
17 internal ice dynamics (particularly in regions of weak sea ice concentration) and
18 sea surface tilt. The GPS compass used in the ADCP system to determine the
19 absolute ocean current provides highly accurate measurements of sea ice drift.
20 Simultaneous measurements of wind from the underway system on board the
21 RV Aurora Australis present an opportunity to assess the influence of wind
22 and ocean currents on the sea ice drift measured during SIPEX ice stations.
23 The wind data were taken as a maximum of ten minute averages collected on
24 both the port and starboard side of the ship, at ∼ 10m height.
25 In Figure 13 we present the speed and adjusted heading for ocean current,
26 wind and ice drift at the longer period stations (2, 10 and 14). The heading ◦ 27 is adjusted to ensure there are no step across the 359⇒0 transition. There
28 is a characteristic increase in speed at the end of each station when the winds
29 reached 30 knots and the ice stations deemed unsafe. There are clearly strong
17 1 correlations between the components of the ice/ocean/wind system, with the
2 speeds of all three varying together closely. Additionally there is also an obvious
3 dependence between the heading of the ice and the headings of the wind and
4 ocean. However, it is not always clear whether the wind or ocean is dominant
5 in driving the ice heading, or that the wind dictates the surface ocean heading,
6 although the correlated increase in speed does suggest this may be the case. For
7 example, the ice and ocean headings are almost exactly matched at station 2,
8 while the wind heading is almost 100 degrees counter clockwise of these, much
9 greater and opposite to what might be expected if the wind was forcing an
10 Ekman rotation. In a counter example, the ice and wind heading are closely
11 matched at the beginning of the station 10 time series, but suddenly diverge
12 and the ocean heading appears to drive the ice between day 273.6–273.7. The
13 ocean and wind rotate in opposite directions at station 10 at day 273.6, as well
14 as at station 14 between days 280.6–280.7. Otherwise, however, the ocean and
15 wind generally have very similar headings and trends.
16 [Figure 13 about here]
17 Qualitatively this unclear relationship is continued in Figure 14, where we
18 see the progressive vector diagrams for ice, ocean and wind tracers at SIPEX
19 stations 1, 2, 3, 10, 13 and 14. For the wind tracer we use 2% of the wind
20 speed. Of the longer period stations, stations 2 and 14 show the ocean as more
21 dominant, and the wind more dominant on station 10. The resultant of the
22 combined ocean and wind vectors agreed best with ice drift at Station 13. In
23 Figure 15 we present the principal component and correlation analysis for the
24 ice drift, ocean current and wind time series shown in Figure 13. At station 2,
25 the percentage of variation explained by the 1st principal component of speed
26 for ice/ocean and ice/wind pairs was 89 and 81% with similarly large overlaps
27 of correlation at other stations. This demonstrates that there is a significant
28 degree of coupling between the wind-ocean system, suggesting that the wind
29 acts to substantially influence the surface ocean flow on the time and spatial
30 scales observed here.
18 1 [Figure 14 about here]
2 [Figure 15 about here]
3 4. Discussion
4 The most interesting and important water mass result from SIPEX was the ◦ ◦ 5 observation of modified Circumpolar Deep Water at 117 E warmer than 0 C.
6 Bindoff et al. (2000) reported similar intrusions of mCDW near this location
7 in summer (Jan. 1996) and noted that this was the primary location sampled ◦ 8 during the BROKE survey between 80–150 E to have this feature. Williams and
9 Bindoff (2003) showed ’highly’ modified circumpolar deep water in the Ad´elie
10 Depression in winter that crossed the shelf break at mid-depths above the denser
11 shelf water below. Williams et al. (2008a) showed that these intrusions increased
12 dramatically in summer after the sea ice production season. However our result
13 here finds mCDW that is much warmer and bears more similarity to shelf break
14 exchanges in the western Ross Sea and Amundsen sea where the Antarctic Slope
15 Front extends onto the continental shelf. Clearly two observations of mCDW
16 in the vicinity of each other over a 10 year period does not prove that these
17 warm, saline intrusions have been constant over this time period. Nonetheless
18 this additional observation of mCDW in late winter suggests that these mCDW
19 intrusions are not limited to the austral summer as in other locations and we
20 speculate that this strengthens the case for for these mCDW intrusions being
21 persistent intra- and inter annually.
22 This is very relevant to the melting of continental ice in this region as mCDW
23 provides the necessary heat flux. In addition to the strength of the mCDW
24 intrusion in this region, the density of these intrusions relative to the local shelf
25 water is also critical. To truly influence the melting of ice shelves, the mCDW
26 must be dense enough to reach the grounding line. For the Mertz Glacier in
27 the Ad´elieDepression, the mCDW intrusions occur at mid-depths and occupy
28 the upper layer of the water column, providing some heat flux to the base of
29 the floating glacier (300–400m), but have little influence on the groundling line
19 1 (>1200m). Here, the mCDW is the densest water mass occupying the bottom
2 layer, implying that should it continue its path south, it could reach nearby
3 grounding lines. Though the East Antarctic ice sheet is the most stable part
4 of Antarctica, the nearby Totten and Moscow University IceShelf were reported
5 by Rignot and Thomas (2002) to be in a state of negative mass balance. Most
6 recently Pritchard et al. (2009) report that the glacial thinning in this region
7 could be three times more than previously estimated. Enhanced ocean heat
8 flux from a persistent, bottom-trapped inflow of mDCW would certainly play a
9 major role in such changes.
10 There are two mechanisms that could contribute to mCDW being the densest
11 water mass on the shelf in this location, despite its warm temperature. The
12 influence of cold, fresh Ice Shelf Water (ISW), from ocean/ice shelf interactions
13 beneath the nearby Totten Glacier and Moscow Univsersity Ice Shelf, is likely
14 to provide a freshening feedback to the local shelf water properties. In addition
15 the sea-ice production in the nearby polynya regions may simply be insufficient
16 to produce truly dense shelf water. To examine this second point further we
17 present the sea ice production for the polynyas identified in this region in Figure 3 18 16. The Cape Poinsett and Dibble Iceberg Tongue polynyas produce ∼60 km
19 of sea ice each year. These values are relatively low in comparison with the
20 total sea ice production in the Mertz Glacier and Cape Darnley polynya/dense 3 21 shelf water formation regions (∼ 180 km ), though the region east of the Mertz 3 22 Glacier has similar production to the Dibble (∼ 60 km ) and this has been
23 shown to be sufficient to produce low salinity shelf water cascades resulting in
24 Antarctic Bottom Water (Williams et al., 2010b). The Dalton Iceberg Tongue 3 25 polynya is the next largest at 40 km , with the Voyeykov and Blodgett averaging 3 26 20 km per year, and although are persistent with relatively small interannual
27 variability, are most likely too small to generate sufficient dense shelf water for
28 Antarctic Bottom Water, let alone in conjunction with negative feedback from
29 the fresh water flux from nearby ice shelves.
30 [Figure 16 about here]
20 1 However we note that the it has been shown in modelling studies (?) that
2 dense shelf water overflows and mCDW intrusions can be intrinsically linked.
3 That is, cascades of dense shelf water over the shelf break can induce a compen-
4 sating upwelling of mCDW. Following this idea, there may yet be dense shelf ◦ 5 water formation, upstream or downstream of the mDCW intrusion at 117 E.
6 The BROKE survey detected evidence of dense shelf water overflows and AABW
7 production downstream from the Dibble Iceberg Tongue polynya in the elevated
8 bottom layer concentrations of offshore CFC-11 (Williams et al., 2010b). We
9 demonstrate this in Figure 17, which shows the bottom layer in meridional
10 vertical sections of potential temperature across the BANZARE and Sabrina ◦ 11 coasts. The Antarctic Bottom Water layer, in particular the -0.3 C isotherm ◦ 12 shoals to 1500m at 128 E to the west of the Dibble and then there is a sub- ◦ 13 sequent decrease in temperature below 3000m at 120 E. Bindoff et al. (2000)
14 reported evidence of modified Shelf Water on the upper continental slope west
15 of the Cape Poinsett polynya region. It is likely that given its modest amount of
16 sea-ice production and the significant amount of melting occurring locally that
17 there is very little dense shelf water being exported from the Dalton Iceberg
18 Tongue Polynya.
19 [Figure 17 about here]
20 As enticing as the concurrent measurements of sea ice drift, ocean currents
21 and wind were for examining the comparative influence of wind and ocean on
22 sea ice drift, the mixed results in Figures 13–15 reflected the fact that our time
23 series were simply too short to address this conclusively. At times the ocean
24 sometimes appears to drive the ice, and sometimes the wind (and sometimes
25 both). Similarities in the ocean and wind headings/velocities and overlapping
26 correlation coefficients strongly indicate that the wind is important in driving
27 the ocean, and hence directly or indirectly the sea ice too. However as men-
28 tioned, this is only very short space/time scales, i.e. 24 hours. Over longer
29 timescales and in the presence of stronger ocean currents such as the Antarctic
30 Slope Current, the ocean is clearly dominating, as indicated by the Ice Argo
21 1 drift pattern in Figure 4 and reported by Heil et al. (2010, this volume) from
2 drifting buoys. South of the shelf break, in the absence of strong currents, the
3 synoptic wind patterns, tidal signals and clear inertial oscillations shown by Heil
4 et al. (2010, this volume) show that sea ice ice drift is much more of a random
5 walk than in locations where there are strong ocean currents. Ideally future
6 deployments of this ’through-ice’ ADCP system can be run for much longer
7 periods and perhaps even modified to run autonomously over several months.
8 4.1. Future work and challenges for observing winter oceanography in the SIZ
9 The pilot ’through-ice’ oceanography program initiated on SIPEX was suc-
10 cessful but can be improved in two key areas. Firstly the accuracy of the salinity
11 measurements needs to be improved to facilitate greater comparison of results
12 with more precise ship-based CTD observations. Secondly, the length of ADCP
13 sampling must be increased to better understand the processes driving the intra-
14 diurnal variability found during SIPEX. Perhaps the best result during SIPEX
15 came from the Ice Argo floats deployed on behalf of the University of Wash-
16 ington. The majority of these floats have continued to work over two years
17 capturing two seasonal cycles. This dataset will greatly improve the under-
18 standing of the seasonal formation and partial decay of the winter mixed layer
19 each year and will provide a comprehensive benchmark against which the cur-
20 rent and future generations of climate models can compare with. However the
21 Ice-Argo floats remain limited to sampling the region north of the continental
22 slope, that is north of the 3000-m isobath, and for this reason the ’through-ice’
23 CTD system used during SIPEX is a relatively inexpensive and easy way to
24 sample the regions on the upper slope and on the continental shelf break during
25 sea-ice research voyages.
26 The biggest challenge that remains after not being met on SIPEX is the
27 effective wintertime sampling of the Antarctic Slope Front and Antarctic Slope
28 Current region beneath the sea ice between the shelf break and the 1500-m iso-
29 bath. We can report that the Antarctic Slope Current was ’experienced’ by the
30 RV Aurora Australis when heading south towards the continental shelf break
22 1 when it began drifting within a ’river of ice’ that was several tens of metres wide.
2 The ship was in fact powerless to manuever efficiently due to the loose nature
3 of the crushed sea ice within it and for several minutes found itself drifting
4 helplessly west, less than a kilometre from numerous large icebergs grounded on
5 the shelf break, until it eventually regained thrust and escaped, with all further
6 requests to conduct CTDs in this region declined. Some modelling studies have
7 found that the Antarctic Slope Current increases in winter beneath the sea-ice
8 (Mathiot et al., 2009), perhaps in response to the increased winds and the in-
9 creased air/ice drag coefficient. More work is needed to understand the physical
10 mechanisms behind this phenomena and its seasonal variability. However in
11 situ observations of the ASC in winter remain a challenge.
12 Acknowledgements
13 We would like to thank the officers and crew of the RV Aurora Australis
14 and the Marine Science Support staff from the Australian Antarctic Division for
15 their professionalism and support during the SIPEX voyage. The majority of
16 this equipment used in the CTD system was previously used during the Amery
17 Ice Shelf drilling project and Russell Brand constructed the overall system. The
18 ADCP system was designed and built by the Marine Science Support group at
19 the Australian Antarctic Division, with the loan of the ADCP kindly arranged
20 by Mike Craven. Mark Rosenberg completed the post-processing of the FSI
21 CTD data. Bathymetry data provided by ETOPO1 (2009). Ice Argo floats
22 deployed during SIPEX and presented in this paper were from the University
23 of Washington and we would like to thank Steve Riser, Annie Wong, Dana
24 Swift and Rick Rupin for their efforts in providing these data. These data were
25 collected and made freely available by the International Argo Project and the na-
26 tional programs that contribute to it.(http://www.argo.ucsd.edu,http://argo.jcommops.org).
27 Argo is a pilot program of the Global Ocean Observing System.
28
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