3.1 Melting or dissolving?

The ocean-driven, solid-to-liquid phase transition at the base of Antarctic ice shelves is dominated by either one of two different thermodynamic ablation processes: dissolving or melting (Woods, Reference Woods1992). At ocean temperatures higher than the in situ sea-water freezing temperature (~−2.5°C at 800 m) and lower than the in situ freshwater ice melting point, ice is dissolved by the saline solute. The rate of ice loss is controlled by the supply of salt and heat to the ice–ocean interface, a process which is dominant at the deep grounding line of cold-cavity ice shelves. When the ocean is warmer than ~−0.5°C at the surface (Wells and Worster, Reference Wells and Worster2011), melting takes place; the ablation rate is controlled by heat transport to the ice–ocean interface only (e.g. at ice-shelf fronts in summer surface water). As the molecular diffusivity of salt is much smaller than the diffusivity of heat (mass diffusivity, D ~ 5 × 10⁻¹⁰ m² s⁻¹, while heat diffusivity, κ ~ 10⁻⁷ m² s⁻¹), ablation by melting is of order (κ/D)^3/4 ~ 50 times faster than by dissolving in environments of low turbulent kinetic energy (Woods, Reference Woods1992; Kerr, Reference Kerr1994; Notz and others, Reference Notz2003; Wells and Worster, Reference Wells and Worster2011). However, throughout this paper we will omit the distinction between melting and dissolving. Instead, we use the terms ‘melting’ and ‘ablation’ interchangeably when referring to the phase change of ice to water.

3.2 How does Ice Shelf Water form?

Melting at the base of ice shelves can form Ice Shelf Water (ISW) (Jacobs and others, Reference Jacobs, Fairbanks and Horibe1985), a water mass characterised by a potential temperature below the surface freezing point (called potential supercooling; Jacobs, Reference Jacobs2004). ISW is part of a thermohaline convection mechanism under and in front of ice shelves (Foldvik and Kvinge, Reference Foldvik, Kvinge and Dunbar1977; MacAyeal, Reference MacAyeal1984), that is also referred to as ‘ice pump’ (Section 3.5, Lewis and Perkin, Reference Lewis and Perkin1986). Its presence is a necessary but not sufficient prerequisite for the formation of platelet ice (Mahoney and others, Reference Mahoney2011).

3.3 How does High Salinity Shelf Water form?

Growing sea ice produces cold and saline surface water due to salt rejection during freezing. This process is enhanced in coastal polynyas, where an extensive air–water interface is maintained by seaward katabatic winds continuously driving the new sea ice offshore (Massom and others, Reference Massom, Harris, Michael and Potter1998; Maqueda and others, Reference Maqueda, Willmott and Biggs2004; Tamura and others, Reference Tamura, Ohshima, Fraser and Williams2016). Polynyas are the main sources of such dense High Salinity Shelf Water (HSSW; Grumbine, Reference Grumbine1991). Sufficient negative buoyancy by salt rejection at the ocean surface triggers deep convection that conveys HSSW to the ocean floor, where it either flows towards the continental shelf slope or into the cavities below ice shelves. Beneath many ice shelves, HSSW may penetrate as far as the grounding line several hundred metres below sea level (Nicholls and others, Reference Nicholls, Østerhus, Makinson and Johnson2001). Since the convective and advective transfers are mainly adiabatic (without heat exchange), the potential temperature of HSSW remains close to the surface freezing point.

3.4 The three modes of ice-shelf melting

The freezing temperature of sea water decreases with increasing pressure and with increasing salinity (UNESCO, 1978; IOC and others, 2010). The HSSW temperature exceeds the in situ freezing point at depth by ~1°C at 1300 m (Jacobs and others, Reference Jacobs, Hellmer, Doake, Jenkins and Frolich1992) and constitutes a heat source available for latent heat of melting. In places where local turbulence or molecular diffusion transports sufficient heat and salt to the ice–water interface, the ice shelf melts. Submarine melting of ice shelves can be grouped into three categories, or ‘modes’ (see Fig. 1 in Jacobs and others, Reference Jacobs, Hellmer, Doake, Jenkins and Frolich1992). The HSSW driven mechanism is termed Mode 1 ice-shelf melting, which dominates the deep basal melting near the grounding line (Jacobs and others, Reference Jacobs, Hellmer, Doake, Jenkins and Frolich1992). Mode 2 ice-shelf melting is associated with the inflow of relatively warm Circumpolar Deep Water (CDW; Jacobs and others, Reference Jacobs, Hellmer, Doake, Jenkins and Frolich1992). Mode 3 ice-shelf melting is associated with Summer Surface Water near the ice-shelf front (Jacobs and others, Reference Jacobs, Hellmer, Doake, Jenkins and Frolich1992) which may be transported under the ice shelf by tidal pumping (e.g. Makinson and Nicholls, Reference Makinson and Nicholls1999; Joughin and Padman, Reference Joughin and Padman2003; Malyarenko and others, Reference Malyarenko, Robinson, Williams and Langhorne2019). Only Modes 1 and 3 ice-shelf melting are thought to lead to ISW formation and are thereby important for platelet-ice formation. This is because the CDW associated with Mode 2 melting is well above the surface freezing-point temperature (Jacobs and others, Reference Jacobs, Hellmer, Doake, Jenkins and Frolich1992).

3.5 ISW plumes and marine ice

As detailed above, basal melting can lead to ISW formation, characterised by cooling and freshening of the ambient water. As the water density at temperatures close to the freezing point mainly depends on salinity (UNESCO, 1981), freshening increases buoyancy, and the ISW plume rises along the basal slope of the ice shelf towards its front. The subsequent pressure release during the adiabatic ascent to shallower depths, together with the initial melt-induced freshening, raises the in situ freezing point of the plume. Therefore, when sufficient ascent and/or freshening occurs, the plume's freezing point can exceed its actual temperature, a condition termed in situ supercooling (Foldvik and Kvinge, Reference Foldvik and Kvinge1974, Reference Foldvik, Kvinge and Dunbar1977; Jacobs and others, Reference Jacobs, Fairbanks and Horibe1985).

Thermodynamic re-adjustment leads to the release of latent heat by the generation of individual ice crystals in the plume (Jenkins and Bombosch, Reference Jenkins and Bombosch1995; Smedsrud and Jenkins, Reference Smedsrud and Jenkins2004). However, the conditions under which nucleation occurs and frazil crystals are initially formed are still unclear. The growing ice crystals mixed with the sea water further increase the plume's buoyancy. This accelerates the ascent of water parcels along the ice slope, speeding up the pressure release (Jordan and others, Reference Jordan, Kimura, Holland, Jenkins and Piggott2015) and enhancing small-scale turbulent and diffusive processes. Theoretical considerations suggest a range of conceivable plume behaviours depending on the sensitive balance between crystal nucleation and precipitation, which itself is a function of growth rate (Rees Jones and Wells, Reference Rees Jones and Wells2018, Reference Rees Jones and Wells2015):

1. (1) Under conditions of near neutral crystal suspension and low growth rates, ascending water parcels sustain the increased buoyancy and may exit the cavity at shallower depths.

2. (2) Plumes under conditions where precipitation dominates experience an increase in bulk density. Arriving at the state of neutral buoyancy inevitably suppresses the generation of supercooling and associated frazil formation. Plumes could detach from the ice base and leave the cavity at depth, as observed near the front of the Ross Ice Shelf (Jacobs and others, Reference Jacobs, Gordon and Ardai1979, Reference Jacobs, Fairbanks and Horibe1985; Jacobs and Giulivi, Reference Jacobs, Giulivi, Spezie and Manzella1999).

3. (3) At faster growth rates, precipitation is more likely and crystals accumulate at the base of ice shelves (Rees Jones and Wells, Reference Rees Jones and Wells2018) as evident in observations (e.g. Fricker and others, Reference Fricker, Popov, Allison and Young2001; Craven and others, Reference Craven, Allison, Fricker and Warner2009; Kulessa and others, Reference Kulessa, Jansen, Luckman, King and Sammonds2014; Koch and others, Reference Koch, Fitzsimons, Samyn and Tison2015).

Such marine ice layers have been observed beneath several Antarctic ice shelves (e.g. Gow and Epstein, Reference Gow and Epstein1972; Morgan, Reference Morgan1972; Zotikov and others, Reference Zotikov, Zagardnov and Raikovsky1980; Oerter and others, Reference Oerter1992; Eicken and others, Reference Eicken, Oerter, Miller, Graf and Kipfstuhl1994; Tison and others, Reference Tison1998; Craven and others, Reference Craven, Allison, Fricker and Warner2009), and contribute to the stability of ice shelves by counteracting basal melting and closing basal crevasses, thus retarding calving events (Kulessa and others, Reference Kulessa, Jansen, Luckman, King and Sammonds2014). The overall process of ice-shelf melt and marine ice formation was termed ‘ice pump’ by Lewis and Perkin (Reference Lewis and Perkin1986), which in general terms describes a heat engine melting ice at some depth in the ocean and depositing it at a shallower location (Robin, Reference Robin1979; Lewis and Perkin, Reference Lewis and Perkin1983). Substantial evidence has been found that compacted marine ice is also the origin of occasionally observed green icebergs around Antarctica (Dieckmann and others, Reference Dieckmann, Hemleben and Spindler1987; Kipfstuhl and others, Reference Kipfstuhl, Dieckmann, Oerter, Hellmer and Graf1992, Fig. 2l), whose green colour is attributed to marine-derived organic matter included in the ice (Warren and others, Reference Warren1993), or, according to a more recent study, to iron-oxide minerals (Warren and others, Reference Warren, Roesler, Brandt and Curran2019).

3.6 Frazil crystals outside of the cavity

Generally, an ISW plume will continue to rise until it either loses its characteristic properties through mixing with other water masses, reaches a depth of neutral buoyancy or arrives at the ocean surface in front of an ice shelf. Clouds of frazil crystals suspended in supercooled ISW plumes at depth near ice-shelf cavities have been detected by acoustic methods (Penrose and others, Reference Penrose, Conde and Pauly1994) and retrieved by net hauls (Dieckmann and others, Reference Dieckmann, Rohardt, Hellmer and Kipfstuhl1986). The dependence of frazil concentration on supercooling is well known (Tsang and Hanley, Reference Tsang and Hanley1985), while the properties of frazil crystals and its occurrence are extensively reviewed in Daly (Reference Daly1984, Reference Daly1994). However, the only published studies we are aware of that successfully measured the size distribution, shape and concentration of frazil crystals were conducted in fresh water, primarily in rivers (e.g. Ghobrial and others, Reference Ghobrial, Loewen and Hicks2013; Marko and others, Reference Marko, Jasek and Topham2015; McFarlane and others, Reference McFarlane, Loewen and Hicks2017) or in the lab (e.g. Marko and Topham, Reference Marko and Topham2015).

3.7 Supercooling, platelet ice and sea ice

Frazil crystals that are advected out of ice-shelf cavities may eventually rise to the surface by their own individual buoyancy if they are not melted by overlying warmer water masses. Depending on where frazil crystals aggregate and grow, they interact differently with their surroundings, and have therefore been given distinct names. When accumulating underneath a solid sea-ice cover, crystals may form a semi-consolidated sub-ice platelet layer (Dayton and others, Reference Dayton, Robilliard and Devries1969; Gow and others, Reference Gow, Ackley, Weeks and Govoni1982; Crocker and Wadhams, Reference Crocker and Wadhams1989; Gough and others, Reference Gough2012b). If an in situ supercooled plume of ISW reaches the ocean surface (Hughes and others, Reference Hughes, Langhorne, Leonard and Stevens2014), crystals in such accumulations grow further (Leonard and others, Reference Leonard2006, Reference Leonard2011; Dempsey and others, Reference Dempsey2010; Smith and others, Reference Smith, Langhorne, Frew, Vennell and Haskell2012). Atmosphere driven heat extraction through the solid sea ice above freezes the water in the interstices between the individual crystals, leading to a downward progressing consolidation in the upper portions of the platelet layer, and ultimately to the platelets being incorporated into the growing sea-ice fabric (Dayton and others, Reference Dayton, Robilliard and Devries1969; Jeffries and others, Reference Jeffries, Weeks, Shaw and Morris1993; Gow and others, Reference Gow, Ackley, Govoni and Weeks1998; Smith and others, Reference Smith2001). This layer is called ‘incorporated platelet ice’ (Smith and others, Reference Smith2001) and has a distinct sea-ice crystal structure (e.g. Serikov, Reference Serikov1963; Eicken and Lange, Reference Eicken and Lange1989; Jeffries and others, Reference Jeffries, Weeks, Shaw and Morris1993; Tison and others, Reference Tison1998), which reflects the properties of the surface ocean during freezing (Gough and others, Reference Gough2012b; Langhorne and others, Reference Langhorne2015). The presence and persistence of a sub-ice platelet layer depends on the ratio between the flux of platelet crystals from below, and the growth rate of the overlying sea ice (Dempsey and others, Reference Dempsey2010; Mahoney and others, Reference Mahoney2011; Gough and others, Reference Gough2012b). Sea ice, mechanically strengthened by layers of incorporated platelet ice, can buttress and stabilise isolated ice tongues against neighbouring coastlines, and thus allow ice shelves to form outside of embayments (see Section 7; Hulbe and others, Reference Hulbe, Johnston, Joughin and Scambos2005; Massom and others, Reference Massom2010).

3.8 Anchor ice

Another sibling of platelet ice is anchor ice. It describes ice crystals covering the seafloor (Dayton and others, Reference Dayton, Robilliard and Devries1969; Mager and others, Reference Mager, Smith, Kempema, Thomson and Leonard2013). The continued presence of supercooled water allows these initial ice crystals to grow rapidly (Osterkamp and Gosink, Reference Osterkamp, Gosink, Barnes, Schell and Reimnitz1984) and form delicate, leaf-like ice structures that clump together (Kempema and others, Reference Kempema, Reimnitz, Clayton and Payne1993). This results in a patchy covering of ice that provides a lair for some benthic species, though it is also a potentially fatal environment for shallow water benthic communities (Dayton and others, Reference Dayton, Robilliard and Devries1969; Dayton, Reference Dayton1989; Denny and others, Reference Denny2011). These ice clusters can grow large enough to become buoyant and detach from the sea floor, lifting any adhered sediment, flora, fauna or scientific equipment into the overlying ice (Dayton and others, Reference Dayton, Robilliard and Devries1969; Hunt and others, Reference Hunt, Hoefling and Cheng2003). See Mager and others (Reference Mager, Smith, Kempema, Thomson and Leonard2013) for a review of anchor ice. Related to this effect and to platelet-ice formation, any ropes or instrumentation immersed into in situ supercooled water can also have ice form on them (Figs 2c and f), leading to measurement errors or instrument failures (Robinson and others, Reference Robinson, Grant, Stevens, Stewart and Williams2020).

4.1 Ice-based measurements

Ice-based measurements have been used since the turn of the 20th century. They allow in situ observations and their relatively low operation costs (compared to techniques relying on aircraft or vessels) make them popular to this day. Most methods do however suffer from low spatio-temporal resolution, and all rely on ice conditions that are safe for on ice travel. Thus, observations of platelet ice during initial freeze-up or during spring time melt are rare and observations are often restricted to fast ice.

4.2 Ocean-based measurements

4.3 Airborne remote sensing

Airborne remote sensing forms the link between in situ observations and satellite remote sensing and is crucial for verifying satellite mission data. Currently only one method of airborne observations of platelet ice exists, which is EM by plane or helicopter (Rack and others, Reference Rack, Haas and Langhorne2013). Airborne EM functions in the same way as ground-based EM systems described above, but has a much higher spatial coverage and higher operating costs. A sub-ice platelet layer has been observed in EM thickness estimates below the sea ice in front of the McMurdo Ice Shelf and is in evidence beneath the ice shelf to a thickness of 55 m: the maximum ice-shelf thickness detectable by the airborne EM (induction sounding) instrument (Rack and others, Reference Rack, Haas and Langhorne2013). The larger footprint of the measurement means that small-scale features cannot be resolved. In addition to the much wider spatial coverage, an advantage over ground-based EM is that airborne EM does not require a minimum sea-ice thickness to operate safely, allowing measurements during freeze in and melt/break-up. This method has great potential to fill the current observational gaps especially in remote regions far from coastal research stations, but much more development work with respect to suitable inversion algorithms is necessary.

4.4 Satellite remote sensing

Satellite altimetry measures the elevation of the snow or ice upper boundary above mean sea level. This measurement may then be converted to sea-ice thickness using estimates of snow depth and density, and sea-ice density (e.g. Kurtz and Markus, Reference Kurtz and Markus2012). Global Navigation Satellite System (GNSS) surface elevation data have been used to assess the influence that the presence of a sub-ice platelet layer has on estimates of sea-ice thickness, with overestimates in consolidated sea-ice thickness of up to 19% close to an ice-shelf edge (Price and others, Reference Price2014). Indeed there is strong evidence of the influence on ICESat laser altimeter data of a sub-ice platelet layer beneath multi-year sea ice in McMurdo Sound between 2003 and 2009 (Price and others, Reference Price, Rack, Haas, Langhorne and Marsh2013). If developed and used in conjunction with other tools, satellite altimetry has the potential to provide an Antarctic-wide view of the pervasiveness of platelet ice.

4.5 Frazil ice numerical models

Frazil ice formation in the water column influences the buoyancy of the surrounding water parcel (e.g. Jordan and others, Reference Jordan, Kimura, Holland, Jenkins and Piggott2015) and thus its dynamics as well as potentially influencing sea ice by accumulating as platelet ice. However, the processes involved in frazil ice formation are not well understood and its transport in supercooled plumes is difficult to observe. In addition, the range of scales on which these processes occur, from the molecular scale nucleation to ISW plumes extending hundreds of kilometres, naturally excludes them from being explicitly simulated in the same model for computational reasons. Frazil ice and plume models aim to bridge this gap by providing much needed insights into large scale effects of these small-scale physical processes. Their results can be used to parameterise frazil ice processes in large scale numerical models such as coupled ice-shelf–ocean, Earth system or climate models. Early theoretical work on frazil ice dynamics was undertaken by Daly (Reference Daly1984). There are many numerical models dealing with the representation of frazil generated by supercooled water resulting from heat loss to the atmosphere, e.g. in polynyas, but this section will focus on models dealing with frazil precipitating from ISW. Models focus on the propagation of (frazil-laden) ISW plumes and the effects on marine-ice thickness and sea-ice thickness (e.g. Jenkins and Bombosch, Reference Jenkins and Bombosch1995; Smedsrud and Jenkins, Reference Smedsrud and Jenkins2004; Holland and Feltham, Reference Holland and Feltham2005) and have been steadily improved since then (e.g. Hughes and others, Reference Hughes, Langhorne, Leonard and Stevens2014; Jordan and others, Reference Jordan, Kimura, Holland, Jenkins and Piggott2015; Cheng and others, Reference Cheng2019). The newest frazil ice models more closely investigate the formation of the individual crystals (e.g. Rees Jones and Wells, Reference Rees Jones and Wells2018) based on theoretical investigations of nucleation (Rees Jones and Wells, Reference Rees Jones and Wells2015). Unfortunately, the lack of quantitative observations of frazil ice properties (e.g. frazil crystal size distributions and concentrations, and basal drag parameterisations) and the knowledge gaps in crucial physical processes such as nucleation mean that verification of model findings is difficult to impossible.

4.6 Platelet-ice numerical models

Since platelet ice is difficult and costly to measure on large scales, as seen above, many groups have worked on numerical models describing platelet ice (e.g. Ohashi and others, Reference Ohashi, Sasaki and Yoshimura2004; Ohashi and Kawano, Reference Ohashi and Kawano2007; Dempsey, Reference Dempsey2008; Kawano and Ohashi, Reference Kawano and Ohashi2008; Dempsey and others, Reference Dempsey2010; Wongpan and others, Reference Wongpan, Langhorne, Dempsey, Hahn-Woernle and Sun2015; Buffo and others, Reference Buffo, Schmidt and Huber2018). Since the exclusion of platelet-ice effects from sea-ice models leads to large errors in areas affected by platelet-ice accretion (Buffo and others, Reference Buffo, Schmidt and Huber2018), these efforts are not only of relevance to the platelet-ice community, but also to the wider Antarctic sea ice and climate modelling community. Currently, the most comprehensive 3-D model of platelet-ice accretion is the one by Wongpan and others (Reference Wongpan, Langhorne, Dempsey, Hahn-Woernle and Sun2015), which is however limited to modelling small 10 cm³ volumes at the ice–water interface, due to high computational costs. The most comprehensive 1-D model, capable of re-creating ice core characteristics (columnar to incorporated platelet-ice transition, thickness and growth duration) by forcing with realistic ocean and atmosphere parameters, was created by Buffo and others (Reference Buffo, Schmidt and Huber2018). In the presence of platelet ice, the authors intend the model to calculate sea-ice properties accurately from environmental parameters. However, to date, this has only been tested against the single set of observations used to configure the model.

4.7 Ocean–ice shelf numerical models

In addition to the challenges in modelling ISW plumes, frazil formation and platelet-crystal accretion described above, accurately simulating the precursor processes of platelet-ice formation and incorporation is difficult. On the one hand, these processes affect the large scale circulation via the generation and advection of ISW and supercooled ocean water. On the other hand, they depend on molecule- to centimetre scale processes, such as ice-shelf melting, ocean turbulence and double diffusion. The evolution of the ocean boundary layer beneath the ice shelf–ocean interface is not well captured in current ocean and ice-shelf models. Sub-ice-shelf ocean observations are critically needed to constrain this component in model solutions and improve the underlying algorithms.

In the past decade, many numerical ocean model frameworks have adopted combined algorithmic and parameterised representations of thermodynamic ice–ocean interaction (e.g. MITgcm, Losch, Reference Losch2008; FESOM, Timmermann and others, Reference Timmermann, Wang and Hellmer2012; Fluidity, Kimura and others, Reference Kimura, Candy, Holland, Piggott and Jenkins2013; ROMS, Robertson, Reference Robertson2013; NEMO, Mathiot and others, Reference Mathiot, Jenkins, Harris and Madec2017; MOM6, Adcroft and others, Reference Adcroft2019). However, a detailed digest of the many caveats in modelling ice-shelf cavities is beyond the scope of this study. A comprehensive overview on the evolution in theoretical concepts and numerical technologies of ice-shelf ocean simulation, including the main contemporary challenges, is given in Williams and others (Reference Williams, Jenkins, Determann, Jacobs and Weiss1998) and Dinniman and others (Reference Dinniman2016). To date, there have been very few measurements of salinity, temperature, currents or anything else beneath ice shelves with which ocean–ice shelf numerical models can be tested.

5.1 Weddell Sea and Filchner–Ronne Ice Shelf

Sub-ice platelet layers and associated incorporated platelet ice in sea-ice cores have mainly been observed close to the ice-shelf edge or coastal polynyas (Eicken and Lange, Reference Eicken and Lange1989; Lange and others, Reference Lange, Ackley, Wadhams, Dieckmann and Eicken1989; Smetacek and others, Reference Smetacek1992). However, Dai and others (Reference Dai2010) found platelet ice in drifting pack ice at the northwestern rim of the Weddell Sea Gyre, far from the influence of ice shelves. The Weddell Sea Gyre (Deacon, Reference Deacon1979; Fahrbach and others, Reference Fahrbach, Rohardt, Schröder and Strass1994) may have carried the ice floe from its point of formation to the position it was observed. In front of the Filchner–Ronne Ice Shelf, ice crystals with diameters between 10 and 25 mm have been captured by trawl nets at depths of 250 m (Dieckmann and others, Reference Dieckmann, Rohardt, Hellmer and Kipfstuhl1986).

5.2 Southeastern Weddell Sea

In Atka Bay close to the Ekström Ice Shelf, incorporated platelet ice has been observed in all fast-ice cores that have been published so far (e.g. Kipfstuhl, Reference Kipfstuhl1991; Günther and Dieckmann, Reference Günther and Dieckmann1999, Reference Günther and Dieckmann2001; Hoppmann and others, Reference Hoppmann2015a, Reference Hoppmann2015b; Arndt and others, Reference Arndt, Hoppmann, Schmithüsen, Fraser and Nicolaus2020). The contribution of platelet ice to sea-ice thickness is sometimes so high that cores may not contain any columnar ice (Hoppmann and others, Reference Hoppmann2015a, Reference Hoppmann2015b; Arndt and others, Reference Arndt, Hoppmann, Schmithüsen, Fraser and Nicolaus2020), although this mainly depends on the timing of initial sea-ice formation. At the end of a typical sea-ice growth season, the sub-ice platelet layer is on average 4 m thick, with up to 10 m observed locally (Hoppmann and others, Reference Hoppmann2015a, 2015b; Hunkeler and others, Reference Hunkeler, Hoppmann, Hendricks, Kalscheuer and Gerdes2016b; Arndt and others, Reference Arndt, Hoppmann, Schmithüsen, Fraser and Nicolaus2020). Hoppmann and others (Reference Hoppmann2015b) found evidence of episodic bursts of floating frazil crystals in the water column that are most likely tidal-induced, and resulting from a plume of supercooled ISW emerging from the Ekström Ice Shelf cavity (Smith and others, Reference Smith2020). Comparable sub-ice platelet layer thicknesses (3–7 m) were observed 10–15 km from the coast in Drescher Inlet (e.g. Eicken and Lange, Reference Eicken and Lange1989; Günther and others, Reference Günther, Gleitz and Dieckmann1999b). This layer was found to thin offshore and eastwards to ~0.5 m 35 km offshore of Halley Bay. Incorporated platelet ice was found to make up 11% of the total sea-ice thickness in the area with some cores consisting of 50% incorporated platelet ice (Eicken and Lange, Reference Eicken and Lange1989). The sub-ice platelet layer was observed using cameras mounted on an ROV (Marschall, Reference Marschall1988) which suggested that the platelets were mainly advected from deeper in the water column but also grew in situ (Eicken and Lange, Reference Eicken and Lange1989).

5.3 East Antarctica

At the Amery Ice Shelf, Penrose and others (Reference Penrose, Conde and Pauly1994) recovered frazil crystals from depths of 150 m using trawl nets. Along the East Antarctic coast, platelet ice is intensively formed at the Antarctic shore off Mirny Station and is responsible for the anomalously rapid growth of the fast ice there and its friable structure (Serikov, Reference Serikov1963; Moretskii, Reference Moretskii1965). These authors also concluded that Antarctic platelet ice is formed by supercooling of sea water by ice shelves and icebergs. To our knowledge, there are no other studies that found individual icebergs to be a significant source for platelet-ice formation, although this may locally be the case under favourable conditions similar to those described in Section 3. The platelet ice off Mirny station attaches to the underside of the fast ice there and, after break-up, is carried by currents out to a distance of more than 40 km from the shore (Moretskii, Reference Moretskii1965). The great thickness of the fast ice at Mertz Glacier has been hypothesised to be caused, at least in part, by platelet-ice accretion, though no direct evidence for this exists (Massom and others, Reference Massom2010). Platelet ice has been found close to Dumont D'Urville Station (Delille and others, Reference Delille, Jourdain, Borges, Tison and Delille2007), which may have been formed from ISW advected from Mertz Glacier (Williams and Bindoff, Reference Williams and Bindoff2003).

5.4 Ross Sea and McMurdo Sound

Observations of platelet ice in McMurdo Sound, eastern Ross Sea, were compiled by Langhorne and others (Reference Langhorne2015). The first ever observations of platelet ice were made in McMurdo Sound (Hodgson, Reference Hodgson1907) and it remains the region with the longest time series and the largest number of observations, mainly under fast ice (Langhorne and others, Reference Langhorne2015). Observations include: sub-ice platelet layers (e.g. Bunt, Reference Bunt1963; Gow and others, Reference Gow, Ackley, Weeks and Govoni1982; SooHoo and others, Reference SooHoo1987; Crocker, Reference Crocker1988; Rivkin and others, Reference Rivkin, Putt, Alexander, Meritt and Gaudet1989; Smith and others, Reference Smith2001; Ryan and others, Reference Ryan2006; Hughes and others, Reference Hughes, Langhorne, Leonard and Stevens2014; Price and others, Reference Price2014; Brett and others, Reference Brett2020); incorporated platelet ice (e.g. Gow and others, Reference Gow, Ackley, Govoni and Weeks1998; Smith and others, Reference Smith, Langhorne, Trodahl, Haskell and Cole1999, Reference Smith2001; Jones and Hill, Reference Jones and Hill2001; Dempsey and others, Reference Dempsey2010; Smith and others, Reference Smith, Langhorne, Frew, Vennell and Haskell2012; de Jong and others, Reference de Jong2013); the presence of frazil crystals in the water column (e.g. Leonard and others, Reference Leonard2006; Gough and others, Reference Gough2010, Reference Gough2012b); anchor ice (e.g. Dayton and others, Reference Dayton, Robilliard and Devries1969; Mager and others, Reference Mager, Smith, Kempema, Thomson and Leonard2013, and references therein) and ice on ropes (e.g. Hodgson, Reference Hodgson1907; Wright and Priestley, Reference Wright and Priestley1922; Dayton and others, Reference Dayton, Robilliard and Devries1969; Barry, Reference Barry1988; Leonard and others, Reference Leonard2006; Gough and others, Reference Gough2010; Mahoney and others, Reference Mahoney2011). Due to the local ocean circulation (e.g. Leonard and others, Reference Leonard2006; Mahoney and others, Reference Mahoney2011; Robinson and others, Reference Robinson, Williams, Stevens, Langhorne and Haskell2014), in the western McMurdo Sound platelet ice occurs earlier in the season, and makes up a greater percentage of the sea-ice thickness compared to the eastern McMurdo Sound, where it is sometimes absent (Dempsey and others, Reference Dempsey2010; Hughes and others, Reference Hughes, Langhorne, Leonard and Stevens2014). In areas with strong tidal currents the sub-ice platelet layer has been observed to be separated from the overlying sea ice, and moving independently of the stationary fast ice (Robinson and others, Reference Robinson, Stevens and McPhee2017). The long time series of platelet-ice observations in McMurdo Sound (over 100 years) does not show any trends in the surface ocean temperatures or the thickness of the platelet ice (Langhorne and others, Reference Langhorne2015).

5.5 West Antarctica

No platelet ice has been observed in the Amundsen and Bellingshausen Seas, except for the observation of platelet ice in Pine Island Bay in the winter of 1991/92 by Veazey and others (Reference Veazey, Jeffries and Morris1994). That was the first study of sea ice in the region and the temperature profiles from the area, the first taken in 1994 (Jacobs and others, Reference Jacobs, Hellmer and Jenkins1996, Reference Jacobs, Jenkins, Giulivi and Dutrieux2011), suggest that Veazey and others (Reference Veazey, Jeffries and Morris1994) may have observed one of the last occurrences of in situ supercooled water outflow from Pine Island Glacier. This hypothesis remains speculative however, since the beginning of the oceanographic and glaciological observational record at Pine Island Bay possibly coincides with the beginning, or at least acceleration, of the dramatic melt-induced thinning (Wingham and others, Reference Wingham, Wallis and Shepherd2009) of the glacier due to ocean forcing and associated changes in cavity geometry (Jenkins and others, Reference Jenkins2010; Jacobs and others, Reference Jacobs, Jenkins, Giulivi and Dutrieux2011). The warming and shoaling of CDW in the Amundsen and Bellingshausen Seas (Schmidtko and others, Reference Schmidtko, Heywood, Thompson and Aoki2014), which in 1994 was already present at the Pine Island Glacier calving front and was significantly warmer than the freezing point (Jenkins and others, Reference Jenkins, Vaughan, Jacobs, Hellmer and Keys1997), make the presence of ISW and platelet ice highly unlikely in present-day conditions.

5.7 Seasonality of formation

The seasonality of platelet-ice formation is likely to be different for each location and may indeed vary from year to year depending on factors such as icebergs influencing fast-ice persistence and ocean circulation (Brunt and others, Reference Brunt, Sergienko and MacAyeal2006; Dinniman and others, Reference Dinniman, Klinck and Smith2007; Remy and others, Reference Remy, Becquevort, Haskell and Tison2008; Robinson and others, Reference Robinson, Williams, Barrett and Pyne2010; Mahoney and others, Reference Mahoney2011). Platelet-ice formation from July to September was recorded in several locations in the Mirny area (Buynitskiy, Reference Buynitskiy1967), and has also been seasonally observed in other regions (Wright and Priestley, Reference Wright and Priestley1922; Dayton and others, Reference Dayton, Robilliard and Devries1969; Crocker and Wadhams, Reference Crocker and Wadhams1989; Günther and Dieckmann, Reference Günther and Dieckmann1999), suggesting that the water only becomes supercooled during this period or after sea ice has begun to form. However, there are substantial deviations from these dates in other regions. The mode of melting below the Amery Ice Shelf is known to be seasonal, with Mode 1 dominating melting and Mode 2 contributing in winter (Herraiz-Borreguero and others, Reference Herraiz-Borreguero2015). ISW has been found at mid-depth year round and marine ice formation is highest during winter (Herraiz-Borreguero and others, Reference Herraiz-Borreguero, Allison, Craven, Nicholls and Rosenberg2013). In Atka Bay, Hoppmann and others (Reference Hoppmann2015a, Reference Hoppmann2015b) found a thickening sub-ice platelet-ice layer throughout winter with a decline in thickness starting in December with the arrival of warm surface waters. In McMurdo Sound, the onset of near surface supercooling conditions was observed as early as March, with platelet ice appearing as early as April (Leonard and others, Reference Leonard2006) in a year when the local oceanography was heavily influenced by a grounded iceberg in the Ross Sea (Brunt and others, Reference Brunt, Sergienko and MacAyeal2006; Dinniman and others, Reference Dinniman, Klinck and Smith2007; Remy and others, Reference Remy, Becquevort, Haskell and Tison2008; Robinson and others, Reference Robinson, Williams, Barrett and Pyne2010; Mahoney and others, Reference Mahoney2011; Robinson and Williams, Reference Robinson and Williams2012). Other researchers found ISW arrived as late as mid-July and platelet ice began to form in early August (Mahoney and others, Reference Mahoney2011; Gough and others, Reference Gough2012b). The onset of supercooling conditions in McMurdo Sound is influenced by the location in the Sound, with an outflow of ISW likely present throughout the sea-ice growth season in western McMurdo Sound, as evident in oceanographic observations (e.g. Robinson and others, Reference Robinson, Williams, Stevens, Langhorne and Haskell2014) and the high contribution of platelet ice in sea-ice cores (e.g. Barry, Reference Barry1988; Dempsey and others, Reference Dempsey2010, recorded an absence of columnar ice from sea-ice cores taken in western McMurdo Sound). In the eastern McMurdo Sound, a summertime inflow of warmer surface waters prevents the appearance of surface supercooling at the start of the sea-ice growth season (Barry, Reference Barry1988; Barry and Dayton, Reference Barry and Dayton1988; Robinson and others, Reference Robinson, Williams, Stevens, Langhorne and Haskell2014). This warm inflow subsides during winter as the ISW plume exiting in western McMurdo Sound spreads eastwards (Barry and others, Reference Barry, Dayton, Dunbar and Leventer-Reed1990; Leonard and others, Reference Leonard2006; Mahoney and others, Reference Mahoney2011; Robinson and others, Reference Robinson, Williams, Stevens, Langhorne and Haskell2014) resulting in an onset of platelet-ice occurrence deeper in the core (later in the growth season) for cores taken successively further eastwards (e.g. Hughes and others, Reference Hughes, Langhorne, Leonard and Stevens2014).

The picture of platelet-ice presence painted above should be viewed with some caution. Both the geographical and temporal occurrence suffer from significant sampling bias, with a higher density of observations close to research stations and during conditions with well-established fast ice. In addition to this, the picture is most likely incomplete, with many incidental observations of platelet ice never published. To expand the base of observations beyond platelet-ice observations published in the literature, it would be essential to undertake a search of expedition reports and other such sources. However, this was far beyond the scope of this review.

6.1 Size of individual platelets

Platelet crystals attached to a slab of sea ice removed from McMurdo Sound in 1999 were reported by Smith and others (Reference Smith2001) to be ~2 mm thick and up to 200 mm in diameter. Systematic measurements of the size distribution of individual platelet crystals were published by Stevens and others (Reference Stevens, O'Connor and Robinson2019), where over 100 individual platelet crystals were measured. The crystals were predominantly within the 50–100 mm diameter class, with some up to ~200 mm in diameter. Thicknesses ranged from ~0.5 to 4 mm.

6.2 Ice-volume fraction

The correct interpretation of signals from radar and EM induction measurements of sea ice are dependent on a knowledge of the electrical properties of sea ice. They are discussed in Morey and others (Reference Morey, Kovacs and Cox1984) and a field approach using a multi-frequency EM instrument is described in Hunkeler and others (Reference Hunkeler, Hendricks, Hoppmann, Paul and Gerdes2015, Reference Hunkeler, Hoppmann, Hendricks, Kalscheuer and Gerdes2016b). The electrical conductivity of the sub-ice platelet layer can be used to calculate its porosity ϕ, and subsequently its ice-volume fraction β = 1 – ϕ, via a modified Archie's law (Archie, Reference Archie1942; Hunkeler and others, Reference Hunkeler, Hoppmann, Hendricks, Kalscheuer and Gerdes2016b) given by $$\phi = \left({\displaystyle{{ \sigma_{\rm p} } / {\sigma_{\rm b}} }} \right)^{1/m}$$ The ice-volume fraction β thereby depends on a knowledge of the brine conductivity, σ _b, the conductivity of the sub-ice platelet layer, σ _p and an empirical cementation factor, m. The latter in turn depends on the pore geometry and connectivity, and on the shape of the platelet crystals. All these factors are not well constrained for the sub-ice platelet layer (Hunkeler and others, Reference Hunkeler, Hoppmann, Hendricks, Kalscheuer and Gerdes2016b). A range of other techniques have also been used to try to measure or estimate the ice-volume fraction in the sub-ice platelet layer. Thus it is not surprising that estimates of the ice-volume fraction, collected in Table 2, range from 0.16 (Price and others, Reference Price2014) to >0.50 (Jeffries and others, Reference Jeffries, Weeks, Shaw and Morris1993), where the former used Archimedes' Law and the latter examined horizontal thin sections of incorporated platelet ice and determined what proportion of the ice was platelet crystals when frozen in.

Table 2.

Estimates for the ice-volume fraction β within the sub-ice platelet layer

Table modified from Gough and others (Reference Gough2012b) and Hoppmann and others (Reference Hoppmann2015b).

6.3 Platelet crystal growth rates

Through video camera still analysis, Smith and others (Reference Smith, Langhorne, Frew, Vennell and Haskell2012) reported that individual crystals in the sub-ice platelet layer grew in episodic bursts with growth rates of the order of 10⁻⁶ m s⁻¹ during these bursts. Incorporated platelet ice was reported by Smith and others (Reference Smith, Langhorne, Frew, Vennell and Haskell2012) to have growth rates of the order of 10⁻⁷ m s⁻¹, an order of magnitude slower than the individual platelet crystals. Leonard and others (Reference Leonard2006) reported maximum growth rates for individual platelet crystals in the sub-ice platelet layer that were close to the minimum crystal growth rate observed by Smith and others (Reference Smith, Langhorne, Frew, Vennell and Haskell2012).

6.4 Basal topography and ocean dynamics in the boundary layer

The presence of sub-ice platelet layers forms an enhanced basal topography and increases the drag at the underside of sea ice (McPhee and others, Reference McPhee, Stevens, Smith and Robinson2016; Robinson and others, Reference Robinson, Stevens and McPhee2017) to 6–30 times the value for sea ice unaffected by platelet ice (Robinson and others, Reference Robinson, Stevens and McPhee2017). The growth of platelet ice is hypothesised to itself be limited by turbulent heat transfer, a quantity dependent on drag. Robinson and others (Reference Robinson, Stevens and McPhee2017) described the effect of platelet-ice roughness-induced mixing on limiting (in a stratified ocean with warmer water underlying a supercooled plume) or increasing (in a homogeneous ocean or thick supercooled plume) platelet-ice growth. This also applies to vertical convection due to the latent heat and salt fluxes associated with platelet-ice growth (Robinson and others, Reference Robinson, Williams, Stevens, Langhorne and Haskell2014). From heat flux calculations, Smith and others (Reference Smith2001) inferred a kinematic eddy viscosity for a sub-ice platelet layer to be ε = 2 × 10⁻⁵ m² s⁻¹. Stevens and others (Reference Stevens, Robinson, Williams and Haskell2009) measured average turbulent energy dissipation rates beneath McMurdo Sound fast ice, where platelet ice is commonly present, to be 3 × 10⁻⁸ m² s⁻³.

6.5 Crystal c-axis distribution

c-axis distribution measured from thin sections of ice cores reveal differences between columnar ice and incorporated platelet ice (Weeks and Ackley, Reference Weeks and Ackley1982; Gough and others, Reference Gough2012b; Fig. 5). In columnar ice, c-axes are preferentially oriented in the horizontal, forming a girdle structure when plotting their orientation (e.g. Weeks and Ackley, Reference Weeks and Ackley1982; Gough and others, Reference Gough2010, Reference Gough2012b). In cases where the columnar ice grows in the presence of a current of constant direction, the horizontal c-axes cluster in the direction of the current (Langhorne and Robinson, Reference Langhorne and Robinson1986; Gough and others, Reference Gough2010). Incorporated platelet ice shows random c-axis orientations, with some even aligned in the vertical (e.g. Gough and others, Reference Gough2012b). The orientation of platelet crystals is thought to be dependent on the rotation of the crystal as it settles at the underside of the ice (Dempsey and others, Reference Dempsey2010; Wongpan and others, Reference Wongpan, Langhorne, Dempsey, Hahn-Woernle and Sun2015, Reference Wongpan, Prior, Langhorne, Lilly and Smith2018c). Wongpan and others (Reference Wongpan, Prior, Langhorne, Lilly and Smith2018c) were able to measure the crystallographic a-axis orientations and showed that these too show differences between columnar ice and platelet ice.

Fig. 5.

Selected thin sections along a sea-ice core from McMurdo Sound, Antarctica, viewed under cross-polarised light, showing granular, columnar and incorporated platelet ice. Depth from sea-ice surface in metres. Crystallographic c-axes orientations have been displayed for three horizontal sections at 0.60, 1.38 and 1.56 m. The blue dots correspond to the orientation of the c-axes when projected on to an upper hemisphere Schmidt net, a standard method of visualising c-axes orientation (Langway, Reference Langway1959). Horizontal c-axes will plot along the perimeter of the circle whereas vertical c-axes will lie in the centre. Vertical thin sections are shown for 0.06–1.15, 1.20–1.29, 1.84–1.93 and 2.20–2.29 m.

6.6 Salinity, permeability and brine volume

Through analysis of multiple sea-ice cores taken at one site, Gough and others (Reference Gough, Mahoney, Langhorne, Williams and Haskell2012a) showed that the salinity of incorporated platelet ice is slightly higher than the salinity of columnar ice, after growth rates were taken into account. This is a subtle effect that is not clearly obvious in the C-shaped salinity profiles from single cores (Smith and others, Reference Smith, Langhorne, Frew, Vennell and Haskell2012). Wongpan and others (Reference Wongpan, Hughes, Langhorne and Smith2018a) showed that this is the result of the arithmetic mean permeability for columnar and incorporated platelet ice being indistinguishable beneath thick Antarctic landfast sea ice, implying that porosity is more important than crystal geometry. It should be noted that Wongpan and others (Reference Wongpan, Hughes, Langhorne and Smith2018a) were only able to probe incorporated platelet ice which has a porosity similar to columnar sea ice. Less is known of the sub-ice platelet layer which has been measured to be isothermal. If it is assumed that the system is in thermal equilibrium, then the salinity gradient in the brine in the interstitial space of the sub-ice platelet layer must also be minimal. Diffusive transport is unlikely to dominate in such a system. At the time of writing, the permeability-porosity relationship and fluid transport in the sub-ice platelet layer are poorly understood.

6.7 Mechanical properties

Many authors (e.g. Weeks, Reference Weeks2010) have shown that all mechanical properties of sea ice are strongly dependent on the porosity, and platelet ice is no exception. Thus the mechanical properties of incorporated platelet ice would be expected to be very similar to the properties of columnar sea ice. Measurements of flexural strength and the fatigue properties of the first-year landfast sea ice of McMurdo Sound were made throughout the 1990s by subjecting in situ cantilever beams to repeated bending with zero mean stress (e.g. Haskell and others, Reference Haskell, Robinson and Langhorne1996; Langhorne and others, Reference Langhorne, Squire, Fox and Haskell1998; Langhorne and Haskell, Reference Langhorne and Haskell1999). These beams were comprised mainly of columnar ice, but included between 10 and 35% incorporated platelet ice (Langhorne and others, Reference Langhorne2015). The crystallographic structure was shown to exert much less influence on the mechanical properties of the beams than salinity, temperature or brine volume (e.g. Langhorne and Haskell, Reference Langhorne and Haskell2004). In contrast, in marine ice with low bulk salinities, the anisotropy in crystal fabric controls the rheological behaviour when samples are subject to unconfined uniaxial compression (Dierckx and others, Reference Dierckx, Peternell, Schroeder and Tison2014). Indeed, recently formed marine ice appears to behave as natural isotropic ice, provided the higher temperature of marine ice is taken into account (Dierckx and Tison, Reference Dierckx and Tison2013). To our knowledge, no data on mechanical properties of the high porosity sub-ice platelet layer have been published.

6.8 Oxygen isotopes (δ¹⁸O)

There are two main controls on the δ¹⁸O values of sea ice: (i) the growth rate of the sea ice (higher growth rates lead to lower sea-ice δ¹⁸O values); and (ii) the δ¹⁸O values of the water which the ice formed from (lower parent water values lead to lower sea-ice δ¹⁸O values). As noted above, individual platelet crystals grow in episodic bursts at rates that are an order of magnitude faster than columnar ice. It would therefore be expected that δ¹⁸O values for crystals from the sub-ice platelet layer and for incorporated platelet ice would be lower than for columnar ice; however, this is not the case. Smith and others (Reference Smith2015) reported that there was no discernible difference between the isotopic composition of columnar ice and incorporated platelet ice. Reliable δ¹⁸O measurements for platelet crystals from the sub-ice platelet layer have not been published. Smith and others (Reference Smith2001) reported measurements from such crystals, but noted the possible contamination from the presence of sea water. Attempts to minimise this contamination have been made but not yet published (personal communication from Natalie Robinson). Smith and others (Reference Smith2015) reported a range of derived effective fractionation coefficients (the ratio of the isotopic composition of the sea water to that of the sea ice) to be +1.84‰ to +2.21‰ for McMurdo Sound sea ice in 2009.

7.1 Contribution to sea-ice thickness and volume

Sea-ice thickness in front of an ice shelf is influenced by the stabilisation of the upper water column by glacial meltwater. Using an ocean model with ice-shelf cavity thermodynamics, Hellmer (Reference Hellmer2004) predicted that sea ice is up to 0.2 m thicker for this reason. The overall thickening effect of glacial meltwater emerging from the ice-shelf cavity on sea ice is likely to be higher than 0.2 m since the model of Hellmer (Reference Hellmer2004) does not take into account supercooling, the formation of frazil ice, or platelet ice. Since approximately half of Antarctica's coastline is fringed by ice shelves (Fretwell and others, Reference Fretwell2013), this potentially affects large areas of the Southern Ocean. In the Ross Sea, Jeffries and others (Reference Jeffries, Weeks, Shaw and Morris1993) determined that incorporated platelet ice constituted over half of the thickness of fast ice in some locations in McMurdo Sound. In late winter, in first-year sea ice close to the ice shelf, at least 0.25 m of the 2 m thick cover forms due to heat loss to the ocean (Purdie and others, Reference Purdie, Langhorne, Leonard and Haskell2006; Gough and others, Reference Gough2012b). Similarly, multi-year fast ice attached to the Mertz Glacier tongue was estimated to be between 10 and 55 m thick, and frazil accumulation must have contributed to its thickness (Massom and others, Reference Massom2010).

The percentage contribution of platelets to the total volume of sea ice has been estimated for various regions in Antarctica. Kipfstuhl (Reference Kipfstuhl1991) determined that fast ice in Atka Bay contained up to 50% incorporated platelet ice, while Günther and Dieckmann (Reference Günther and Dieckmann1999) found up to 60% of the fast ice to consist of incorporated platelet ice in the same area. Jeffries and others (Reference Jeffries, Weeks, Shaw and Morris1993) and Smith and others (Reference Smith2001) showed that platelet ice can add up to more than 40% of the total sea-ice mass in McMurdo Sound. Eicken and Lange (Reference Eicken and Lange1989) found their cores from the coastal zones of the southeastern Weddell Sea to contain 11% of incorporated platelet ice, while Lange (Reference Lange1988) and Lange and others (Reference Lange, Ackley, Wadhams, Dieckmann and Eicken1989) estimate that between 20 and 30% of sea ice in coastal regions is comprised of platelet ice. The spatial variability of the contribution of platelet ice to sea ice was studied by Barry (Reference Barry1988), Smith and others (Reference Smith2001) and Dempsey and others (Reference Dempsey2010) based on ice-core analyses. The thickness of sub-ice platelet layers, when present, have been reported to range from consisting of ranging from individual crystals attached to the ice bottom (Mahoney and others, Reference Mahoney2011) to 8 m or thicker (Hughes and others, Reference Hughes, Langhorne, Leonard and Stevens2014; Price and others, Reference Price2014; Hoppmann and others, Reference Hoppmann2015a). In some locations on the eastern side of McMurdo Sound, incorporated platelet ice has been observed in the lower parts of cores only (e.g. Smith and others, Reference Smith2001; the lower 690 mm in a core with a total length of 2150 mm), with some locations not having any incorporated platelet ice present at certain times (e.g. Cape Evans in Smith and others, Reference Smith2001). In contrast, Hoppmann and others (Reference Hoppmann2015a) and Hughes and others (Reference Hughes, Langhorne, Leonard and Stevens2014) took sea-ice cores in Atka Bay and western McMurdo Sound, respectively, which consisted of incorporated platelet ice throughout, i.e. these cores did not contain any pure columnar ice.

7.2 Further effects on sea ice

Sea ice is a critical part of Earth's climate system through its high albedo (especially when snow covered), thermal insulation properties, and through its role in deep water formation. It is generally agreed that the growth of a sub-ice platelet layer causes sea ice to be thicker than it would be with heat loss to the atmosphere only (e.g. Smith and others, Reference Smith, Langhorne, Frew, Vennell and Haskell2012; Hoppmann and others, Reference Hoppmann2015b, Section 7.1). The presence of platelet ice will therefore alter the thermal insulation properties of sea ice and its longevity, with flow-on effects to climate. The overall impact of ice-shelf basal melt on sea-ice formation, and hence on climate, is not well understood. The contribution of ice-shelf melting as a factor in the unexplained discrepancy between global climate models and satellite-derived sea-ice extent around Antarctica has been explored using different climate models, but is controversial (Bintanja and others, Reference Bintanja, van Oldenborgh, Drijfhout, Wouters and Katsman2013, Reference Bintanja, van Oldenborgh and Katsman2015; Swart and Fyfe, Reference Swart and Fyfe2013; Holland and others, Reference Holland2014; Pauling and others, Reference Pauling, Bitz, Smith and Langhorne2016, Reference Pauling, Smith, Langhorne and Bitz2017; Bronselaer and others, Reference Bronselaer2018). As explained in Section 4.7, no ice-shelf–ocean (and thus climate) models currently include the effect of platelet-ice formation on sea-ice extent and thickness, so these papers have only taken into account the direct effects of the freshwater input (e.g. stratification, and sometimes also cooling). As shown earlier in this paper, manifestations of ice shelf–ocean interaction and associated phenomena have been observed in many locations around Antarctica (Fig. 4). Although aspects of these interactions have been investigated in detail for example in McMurdo Sound (Ross Sea) and Atka Bay (Weddell Sea), little is known about the overall role of platelet ice for coastal sea ice in Antarctica. This lack of regional information is a barrier for investigating the effects of platelet ice on sea ice and the climate system.

Although it is relatively simple to accurately predict the growth of undeformed Arctic sea ice by thermodynamic modelling, a thermodynamic model which accurately predicts ice growth in the Arctic will not do the same in the Antarctic when platelet ice is present due to the negative ocean heat flux associated with the presence of platelet ice (Langhorne and others, Reference Langhorne2015). This phenomenon has encumbered efforts to predict growth rates of fast ice in Antarctica ever since Crocker and Wadhams (Reference Crocker and Wadhams1989), as detailed in Smith and others (Reference Smith2015).

7.3 Effects on ice-shelf buttressing

As noted in previous sections, incorporated platelet ice results in thicker sea ice and therefore must contribute considerably to the overall stability of the fast ice around Antarctica and to its prolonged existence, especially in bays and inlets. The stability of the fast ice, with its estimated area of between 5.5 × 10⁵ and 8 × 10⁵  km² (Kozlovsky and others, Reference Kozlovsky, Nazintsev, Fedotov and Cherepanov1977, as cited in Fedotov and others, Reference Fedotov, Cherepanov, Tyshko and Jeffries1998) also plays an important role in the stability of ice shelves, for example by protecting the calving front from ocean swell (Massom and others, Reference Massom2018) or through buttressing (MacAyeal and Holdsworth, Reference MacAyeal and Holdsworth1986). Consequently, the melting of ice shelves resulting in platelet-ice formation forming thicker and more stable fast ice may indirectly lead to ice shelves that are more stable due to protective, more stable fast ice in front of them. However, further research is needed to investigate these connections. The climate effects of platelet ice are therefore complex, involving a series of feedbacks.

The stability of fast ice and ice shelves not only has consequences for the global climate system, but also for the ecology of the coastal regions, as explained in Section 7.5.

7.4 An indicator for ice-shelf processes

More recently, Hoppmann and others (Reference Hoppmann2015b) and in much more detail, Langhorne and others (Reference Langhorne2015), hypothesised that the presence and evolution of platelet ice and its physical properties as a signature of supercooling could be an indicator for the state of an ice shelf with respect to processes in the ice-shelf cavity. Smith and others (Reference Smith2015) suggested using δ¹⁸O samples from sea ice as evidence for ice-shelf melt processes, but noted the challenges of doing this. This is a particular useful and intriguing idea, especially since the sampling beneath Antarctic ice shelves is a huge logistical and financial challenge. Although the exact connection between those two cryospheric elements is not clear, it is certain that the properties of sea ice (or more accurately, fast ice) close to an ice shelf should be an important element in the monitoring of the latter.

7.5 Significance for ecosystems and ecology

The first studies of the biology associated with sub-ice platelet layers date back to the 1960s in McMurdo Sound (Bunt, Reference Bunt1963; Bunt and Wood, Reference Bunt and Wood1963; Bunt and Lee, Reference Bunt and Lee1970). These first explorations described the dense microbial assemblages associated with platelet layers (e.g. Palmisano and Sullivan, Reference Palmisano, SooHoo and Sullivan1985; Dieckmann and others, Reference Dieckmann, Arrigo and Sullivan1992; Arrigo and others, Reference Arrigo, Robinson and Sullivan1993b; Ackley and Sullivan, Reference Ackley and Sullivan1994; Arrigo and Thomas, Reference Arrigo and Thomas2004). The high concentration of biological matter in sub-ice platelet layers is usually immediately evident by the reddish-brown colour of the interstitial waters, and has significant implications for the entire coastal Antarctic food web.