The Sensitivity of the Antarctic Ice Sheet to a Changing Climate: Past, Present, and Future

Abstract

The Antarctic Ice Sheet (AIS) is out of equilibrium with the current anthropogenic‐enhanced climate forcing. Paleoenvironmental records and ice sheet models reveal that the AIS has been tightly coupled to the climate system during the past and indicate the potential for accelerated and sustained Antarctic ice mass loss into the future. Modern observations by contrast suggest that the AIS has only just started to respond to climate change in recent decades. The maximum projected sea level contribution from Antarctica to 2100 has increased significantly since the Intergovernmental Panel on Climate Change (IPCC) 5th Assessment Report, although estimates continue to evolve with new observational and theoretical advances. This review brings together recent literature highlighting the progress made on the known processes and feedbacks that influence the stability of the AIS. Reducing the uncertainty in the magnitude and timing of the future sea level response to AIS change requires a multidisciplinary approach that integrates knowledge of the interactions between the ice sheet, solid Earth, atmosphere, and ocean systems and across time scales of days to millennia. We start by reviewing the processes affecting AIS mass change, from atmospheric and oceanic processes acting on short time scales (days to decades), through to ice processes acting on intermediate time scales (decades to centuries) and the response to solid Earth interactions over longer time scales (decades to millennia). We then review the evidence of AIS changes from the Pliocene to the present and consider the projections of global sea level rise and their consequences. We highlight priority research areas required to improve our understanding of the processes and feedbacks governing AIS change.

Full text

The Sensitivity of the Antarctic Ice Sheet to a Changing
Climate: Past, Present, and Future
T. L. Noble
1
, E. J. Rohling
2,3
, A. R. A. Aitken
4
, H. C. Bostock
5
, Z. Chase
1
,
N. Gomez
6
, L. M. Jong
7,8
, M. A. King
9
, A. N. Mackintosh
10
, F. S. McCormack
10
,
R. M. McKay
11
, L. Menviel
12
, S. J. Phipps
1
, M. E. Weber
13
, C. J. Fogwill
14
,
B. Gayen
15
, N. R. Golledge
11
, D. E. Gwyther
1
, A. McC. Hogg
2,16
, Y. M. Martos
17,18
,
B. Pena‐Molino
8,19
, J. Roberts
7,8
, T. van de Flierdt
20
, and T. Williams
21
1
Institute for Marine and Antarctic Studies, University of Tasmania, Hobart, Tasmania, Australia,
2
Research School of
Earth Sciences, Australian National University, Canberra, ACT, Australia,
3
Ocean and Earth Science, University of
Southampton, National Oceanography Centre, Southampton, UK,
4
School of Earth Sciences, University of Western
Australia, Perth, Western Australia, Australia,
5
School of Earth and Environmental Sciences, University of Queensland,
Brisbane, Queensland, Australia,
6
Department of Earth and Planetary Sciences, McGill University, Montreal, Quebec,
Canada,
7
Australian Antarctic Division, Department of Agriculture, Water and Environment, Kingston, Tasmania,
Australia,
8
Australian Antarctic Program Partnership, Institute of Marine and Antarctic Studies, University of Tasmania,
Hobart, Tasmania, Australia,
9
School of Technology, Environments and Design, University of Tasmania, Hobart,
Tasmania, Australia,
10
School of Earth, Atmosphere and Environment, Monash University, Melbourne, Victoria,
Australia,
11
Antarctic Research Centre, Victoria University of Wellington, Wellington, New Zealand,
12
Climate Change
Research Centre, PANGEA, University of New South Wales, Sydney, New South Wales, Australia,
13
Steinmann Institute,
University of Bonn, Bonn, Germany,
14
Geography, Geology and the Environment, Keele University, Staffordshire, UK,
15
Department of Mechanical Engineering, University of Melbourne, Melbourne, Victoria, Australia,
16
ARC Centre of
Excellence for Climate System Science, Sydney, New South Wales, Australia,
17
Department of Astronomy, University of
Maryland, College Park, MD, USA,
18
Planetary Magnetospheres Laboratory, NASA Goddard Space Flight Center,
Greenbelt, MD, USA,
19
Ocean and Atmosphere, Commonwealth Scientific and Industrial Research Organization, Hobart,
Tasmania, Australia,
20
Department of Earth Science and Engineering, Imperial College London, London, UK,
21
International Ocean Discovery Program, Texas A&M University, College Station, TX, USA
Abstract The Antarctic Ice Sheet (AIS) is out of equilibrium with the current anthropogenic‐enhanced
climate forcing. Paleoenvironmental records and ice sheet models reveal that the AIS has been tightly
coupled to the climate system during the past and indicate the potential for accelerated and sustained
Antarctic ice mass loss into the future. Modern observations by contrast suggest that the AIS has only just
started to respond to climate change in recent decades. The maximum projected sea level contribution
from Antarctica to 2100 has increased significantly since the Intergovernmental Panel on Climate Change
(IPCC) 5th Assessment Report, although estimates continue to evolve with new observational and
theoretical advances. This review brings together recent literature highlighting the progress made on the
known processes and feedbacks that influence the stability of the AIS. Reducing the uncertainty in the
magnitude and timing of the future sea level response to AIS change requires a multidisciplinary approach
that integrates knowledge of the interactions between the ice sheet, solid Earth, atmosphere, and ocean
systems and across time scales of days to millennia. We start by reviewing the processes affecting AIS mass
change, from atmospheric and oceanic processes acting on short time scales (days to decades), through to ice
processes acting on intermediate time scales (decades to centuries) and the response to solid Earth
interactions over longer time scales (decades to millennia). We then review the evidence of AIS changes
from the Pliocene to the present and consider the projections of global sea level rise and their consequences.
We highlight priority research areas required to improve our understanding of the processes and feedbacks
governing AIS change.
Plain Language Summary The Antarctic Ice Sheet (AIS) is an important component of the
global climate system. Human activities have caused the atmosphere and especially the oceans to warm.
However, the full effect of human caused climate change on the AIS has not currently been realized because
the ice sheet responds on a range of time scales and to many different Earth processes. Modern
observations show that West Antarctica has been melting at an accelerating rate since the 2000s, while the
data for East Antarctica are less clear. Environmental records preserve the history of the climate
©2020. The Authors.
This is an open access article under the
terms of the Creative Commons
Attribution License, which permits use,
distribution and reproduction in any
medium, provided the original work is
properly cited.
REVIEW ARTICLE
10.1029/2019RG000663
Key Points:
•The AIS is a highly dynamic
component of the Earth system,
evolving on a broad range of
temporal and spatial scales
•Paleoenvironmental evidence
highlights the centennial to
millennial response time scales of
the AIS to atmospheric‐ocean
forcing
•Coupling feedbacks in Earth system
components are required to
reduce the uncertainty in AIS's
contribution to past and future sea
level rise
Correspondence to:
T. L. Noble,
taryn.noble@utas.edu.au
Citation:
Noble, T. L., Rohling, E. J., Aitken,
A. R. A., Bostock, H. C., Chase, Z.,
Gomez, N., et al. (2020). The Sensitivity
of the Antarctic Ice Sheet to a Changing
Climate: Past, Present, and Future.
Reviews of Geophysics,58,
e2019RG000663. https://doi.org/
10.1029/2019RG000663
Received 21 JUN 2019
Accepted 11 AUG 2020
Accepted article online 15 AUG 2020
NOBLE ET AL. 1of89

and AIS, which extend beyond the instrumental record and reveal how the AIS responded to past climate
warming. Estimates of how much the AIS will contribute to sea level rise by the Year 2100 have changed as a
result of new information on how the AIS evolved in the past and research into the interactions between the
ice sheet, solid Earth atmosphere, and ocean systems. This review brings together our knowledge of the
major processes and feedbacks affecting the AIS and the evidence for how the ice sheet changed since the
Pliocene. We consider the future estimates and consequences of global sea level rise from melting of the AIS
and highlight priority research areas.
1. Introduction and Motivation
The Antarctic Ice Sheet (AIS) is the largest potential source of and most uncertain contributor to global sea
level rise (Oppenheimer et al., 2019). The marine‐based sectors of Antarctica, meaning the portion of the AIS
that lies below global mean sea level (GMSL), contain enough fresh water to raise GMSL by approximately
25 m (BEDMAP 2: Fretwell et al., 2013, Figure 1). The response of the AIS to anthropogenic climate warming
in terms of the time scales of ice loss and where the ice loss occurs will depend on the extent of climate warm-
ing and interactions between the ice sheet and the atmosphere, ocean, and the solid Earth.
Observations during the satellite era show the AIS have been losing mass at an accelerating rate (Rignot
et al., 2019). Continent‐wide observations show a net acceleration of grounding line retreat between 2010
and 2016, driven by ice shelf thinning (Konrad et al., 2018). Differences in ice discharge from glaciers have
been observed between West and East Antarctica. Mass loss of −214 ± 51 Gt yr
−1
(2008–2015; Gardner
et al., 2018) in the West Antarctic Ice Sheet (WAIS) has been driven predominantly by ocean‐forced ice shelf
melt in the Amundsen and Bellingshausen Sea sectors (Turner et al., 2017). Ice mass loss from WAIS has
dramatically increased over the past two decades (Rignot et al., 2019; Shepherd et al., 2019). Increased ice
velocity and glacier terminus retreat have been observed for outflows of the East Antarctic Ice Sheet
(EAIS) in Wilkes Land, in response to warming ocean temperatures (Li et al., 2016; Miles et al., 2016;
Shen et al., 2018). However, EAIS mass change has been difficult to quantify and remains uncertain. For
example, previous assessments have suggested that the EAIS has remained close to mass balance through
increased snow accumulation (Boening et al., 2012; King et al., 2012; The IMBIE Team et al., 2018), but
the large uncertainty on these estimates has made it difficult to determine whether the EAIS has lost or
gained mass (Gardner et al., 2018; Rignot et al., 2019; The IMBIE Team et al., 2018).
Uncertainties in the future contribution of the AIS to sea level rise are associated with unknowns in
future climate change and in the processes and feedbacks governing the response of the AIS to warming
of the atmosphere and ocean (Bamber et al., 2019; Nowicki & Seroussi, 2018; Schlegel et al., 2018).
The Intergovernmental Panel on Climate Change (IPCC) Fifth Assessment Report (AR5) estimated that
the AIS is likely to contribute only −0.08 to 0.14 m to sea level rise by 2100, even under the most extreme
emissions scenario (Church et al., 2013), where projected ice mass loss is compensated by surface mass gains
in the EAIS. Further assessments of global and regional sea level projections to 2100 have since been devel-
oped based on new ice dynamic feedbacks not explicitly considered in the IPCC AR5. For example,
atmospheric‐driven melting that causes hydrofracturing and leads to ice cliff failure has been proposed as
a mechanism for triggering major, rapid mass loss from the AIS. A model incorporating ice cliff failure esti-
mated an Antarctic contribution to global sea level rise of 1.05 ± 0.30 m by 2100 under the high‐end
Representative Concentration Pathway 8.5 (RCP8.5) climate scenario, increasing to 15.65 ± 2.00 m by
2500 (DeConto & Pollard, 2016). However, this mechanism has not been directly observed in polar ice sheets
and glaciers. Edwards et al. (2019) revisited the projections of DeConto and Pollard (2016) and explored the
uncertainties in ice sheet model projections of sea level rise to 2100. They suggested that under the RCP8.5
scenario the AIS is most likely to contribute 0.45 m to global sea level rise if the mechanism of DeConto and
Pollard (2016) is true, but only 0.15 m otherwise. The expert judgment assessment of Bamber et al. (2019) for
a high‐temperature scenario, estimated “likely”(17–83%) sea level contributions of 0.03 to 0.46 m from
WAIS and −0.04 to 0.11 m from EAIS by 2100, increasing to 0.07 to 2.28 m (WAIS) and −0.14 to 0.51 m
(EAIS) by 2300.
Observations and modeling evidence indicates that the AIS can respond to disequilibria between the ice
sheet and the atmosphere, ocean, and solid Earth on a range of time scales; from hours to decades (e.g.,
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Padman et al., 2018) and hundreds (MacAyeal, 1992a) to thousands of years (Warrick & Oerlemans, 1990).
The need to develop a better understanding of AIS sensitivity to climate forcing is driving multidisciplinary
research, as noted by the many recent publications on feedbacks with sea level and the solid Earth (e.g., de
Boer et al., 2017; Whitehouse et al., 2019) and evolution of continental shelf bathymetry (Colleoni
et al., 2018). Other recent reviews have focused on the processes governing AIS stability and ice sheet
dynamic processes, including the role of ice shelf buttressing and marine ice sheet instability (MISI) (e.g.,
Pattyn, 2018; Pattyn et al., 2017). Sea ice (see Hobbs et al., 2016) and ice shelves have been shown to
provide stability to the past and present AIS through buttressing (Hughes et al., 2017) and through
dampening ocean swell (Massom et al., 2018). The role of ocean processes on AIS stability has also been
reviewed. This includes ice shelf‐ocean processes (Galton‐Fenzi et al., 2016; Rintoul, 2018) and the
influence of the Antarctic Slope Current on ocean heat flux to the Antarctic margin (Thompson
et al., 2018), processes driving ice retreat in the Amundsen Sea Sector (Turner et al., 2017), and the role of
tides (Padman et al., 2018), in ice sheet mass balance and dynamics.
We focus on new knowledge from the past 5 years, identifying gaps and priorities for future research. The
paper is structured as follows: in section 2 new insights into AIS‐Southern Ocean interactions are discussed,
in particular the mechanism for oceanic heat transport to the Antarctic margin. In section 3, we highlight the
ice dynamical processes governing the AIS response to the climate and solid Earth. In section 4, we review
the current understanding of solid Earth interactions and processes over a range of time scales that affect the
ice sheet directly or indirectly. Evidence for ice sheet changes from the observational record for the present
climate‐AIS state, and times in the past are reviewed in section 5. Finally, in section 6, the consequences of a
more dynamic AIS are discussed in terms of our current understanding of sea level and wider effects on
Southern Ocean and global climate processes.
An interdisciplinary perspective of AIS change is required if we are to make progress in Antarctic research
(Figure 2). The motivation of this review was to provide a resource to engage and bring together the different
fields of Antarctic science and thereby improve our understanding of the science and gaps in our knowledge.
A summary is presented of the current knowledge about processes that govern AIS stability across the range
of disciplines included in this review, and we offer a list of research priorities for understanding AIS behavior
and hence for reducing uncertainty in its response to, and role in, future climate perturbations.
2. Ice‐Ocean‐Atmosphere Interactions
The atmosphere and ocean act together to induce changes in the AIS. Understanding how these systems
interact and the associated feedbacks that arise from the AIS are key to quantifying the nature and time
scales of the ice sheet response to a changing climate. Atmosphere and ocean processes impacting the ice
sheet are expressed differently in East and West Antarctica. Variability in atmospheric circulation has
Figure 1. Map of Antarctica showing (a) the regions, glaciers (Gl), ice shelves (IS), and iceberg tongues (IBT) mentioned in the text and (b) the subglacial
basins (SB) and mountainous features of the Antarctic continent (BEDMAP 2; Fretwell et al., 2013). Note that the Antarctic Peninsula is geographically
and geologically part of West Antarctica, but the Antarctic Peninsula Ice Sheet is glaciologically distinct from the WAIS. Source: Quantarctica and the Norwegian
Polar Institute.
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contributed to an increase in the mass loss of the WAIS (section 2.1). For both EAIS and WAIS, mass loss has
been enhanced by intrusions of relatively warm water onto the Antarctic continental shelf (section 2.2), dri-
ven by changes in ocean currents, tides, and variability in surface winds and by feedbacks associated with
increased freshwater flux around the Antarctic margin. The global oceans have absorbed 90% of the excess
heat resulting from anthropogenic emissions, with 75 ± 22% heat uptake occurring in the Southern Ocean
(Frölicher et al., 2015). Understanding how this heat is redistributed and the resulting changes in ocean cir-
culation will impact on projections of sea level change.
Figure 2. The upper panel shows current available methods for observations on the ice and in the ocean. ApRes:
Autonomous Phase‐Sensitive Radio echosounder can be mounted on aircraft or pulled by a snowmobile; AXCTD:
Airborne eXpendable Conductivity Temperature Depth probe; ADCP: acoustic Doppler current profiler; AXBT: Airborne
EXpendable BathyThermograph. The lower multipanels show an overview of the processes and feedbacks discussed
in this review that affect AIS dynamics. Our understanding of these processes is underpinned by direct observations of
the atmosphere, ocean, ice sheet, and solid Earth and the past reconstructions from ice, sub‐ice, and marine
sediment cores.
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2.1. Atmospheric Processes Driving AIS Mass Changes
The AIS mass balance is driven by atmospheric circulation, via its effects on near‐surface air temperature
and precipitation (and therefore snowfall and surface melt), and particularly through its influence on ocean
forcing. The major climate drivers of atmospheric conditions include teleconnection modes such as the
Southern Annular Mode (SAM), Pacific Decadal Oscillation (PDO), and El Niño–Southern Oscillation
(ENSO) and the presence and strength of low‐pressure systems such as the Amundsen Sea Low. These
low‐frequency coupled ocean‐atmosphere interactions have a range of impacts, including altering rates of
snowfall, air temperatures, and sea ice conditions and extent. Changes to surface wind patterns can also alter
the degree of upwelling of relatively warm deep‐water masses, such as Circumpolar Deep Water (CDW) and,
as a result, melting of the underside of floating glacier ice.
SAM is a measure of the pressure gradient between the middle and high latitudes and describes the
north‐south position and strength of westerly winds circling Antarctica. A positive phase indicates a negative
high‐latitude and positive midlatitude pressure gradient, which leads to a southward shift in the storm tracks
toward the Antarctica and a strengthening of the westerly winds (Thompson et al., 2000). Stronger and pole-
ward shifted westerlies (positive SAM) have caused enhanced precipitation across much of West Antarctica
and in the western Antarctic Peninsula over the past 50 years (Goodwin et al., 2016). However, the enhanced
westerly winds cause an orographic “precipitation shadow”on the eastern side of the Antarctic Peninsula
and reduce the relatively small amounts of precipitation over East Antarctica (Marshall et al., 2017). Ice core
records reveal the complex relationship between SAM and other climate oscillators (e.g., PDO) in multideca-
dal analyses (Goodwin et al., 2016). Over the past millennia ice core records show that SAM has evolved
toward a more positive state since CE 1480 (Abram et al., 2014; Dätwyler et al., 2018). This intensified and
poleward‐shifted atmospheric circulation likely results from a combination of stratospheric ozone depletion
and increased greenhouse gas concentrations (Arblaster et al., 2014). In Dronning Maud Land, increased
accumulation between 2009 and 2011 was driven by anomalous large‐scale atmospheric conditions that
enhanced the poleward transport of moisture via atmospheric rivers, long narrow bands of enhanced water
vapor that stretched from subtropical latitudes to the Antarctic coast (Gorodetskaya et al., 2014). Although ice
core records indicate that such accumulation events have occurred previously, climate models show that they
are increasingly likely in a warming climate (Lenaerts et al., 2013).
There is mounting evidence that ENSO drives Antarctic mass balance changes primarily via its influence
on ocean heat transport. Variability in ENSO has been linked to interannual variations in ice shelf melt
rates at a regional scale (Paolo et al., 2018). Prevailing La Niña conditions in 2011 resulted in reduced rates
of sub ice shelf melting at Pine Island Glacier (Dutrieux et al., 2014). A strong El Niño in the 1940s may
have triggered the twentieth century retreat of Pine Island Glacier (Smith, Andersen, et al., 2017).
Similarly, the warming trend observed over West Antarctica (Schneider & Steig, 2008; Steig et al., 2013)
is likely linked to natural variability associated with high‐latitude teleconnections from the ENSO
(Smith & Polvani, 2017). A surface melt event in 2016 over the Ross Ice Shelf and extending to more than
1,000 m above sea level in West Antarctica was also likely favored by prevailing El Niño conditions
(Nicolas et al., 2017). The linkage between ENSO and mass balance in West Antarctica occurs either
directly via wind forcing of ocean currents and/or indirectly via effects on polynya formation and sea ice
(St‐Laurent et al., 2015; Webber et al., 2017).
Internal variability of the regional atmospheric circulation has a profound influence on Antarctic climate,
primarily by affecting the distribution and strength of cyclones and anticyclones around the continent
(e.g., Raphael, 2004). For example, the warming trend on the Antarctic Peninsula from the mid‐1950s to
the late 1990s (Turner et al., 2005) has now reversed, likely due to stronger cyclonic conditions in the north-
ern Weddell Sea, bringing temperature trends on the Peninsula within bounds of natural variability (Turner
et al., 2016). Given the large magnitude of natural variability, long observational periods are required to
establish true statistical significance in interannual to decadal variability (Stevenson et al., 2010).
However, recent modeling and observations of Holland et al. (2019) have shown how the regional retreat
of glaciers and melting of ice shelves in the Amundsen Sea can be attributed to anthropogenic forcing
(e.g., greenhouse gas increase). Their results demonstrated that increasing twentieth century anthropogenic
greenhouse gases changed the shelf break winds in the Amundsen Sea, resulting in increased ice loss
induced by enhanced ocean forcing.
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Future projection experiments performed with CMIP5 models suggest that increased greenhouse gas con-
tent over the coming century could lead to a positive trend in the SAM (Zheng et al., 2013). Improved esti-
mates of future changes in near‐Antarctic winds, and their intrinsic variability, is a high priority for
predicting heat fluxes onto the continental shelf, especially in the Amundsen‐Bellingshausen sector, as well
as studies on more extreme ENSO events (Cai et al., 2015) on Antarctic shelf subsurface temperatures.
2.2. Oceanic Processes Driving AIS Changes
Oceanic processes that drive changes to the floating parts of the AIS influence its evolution predominantly
through changes in buttressing (De Angelis & Skvarca, 2003; Dupont & Alley, 2005a; Gudmundsson, 2013)
(section 3.1.1). The ice shelves that have grounding lines below sea level are vulnerable to ocean forcing and
MISI (section 3.1.2). Ocean‐driven ice shelf mass loss has long been known to occur primarily through melt-
ing at the underside of ice shelves (~67%), but recent observations of iceberg calving have shown that this
process also contributes significantly to mass loss (Depoorter et al., 2013; Liu et al., 2015; Rignot et al., 2013).
Over the last decade, evidence for a significant role of intrusions of relatively warm CDW in driving sub‐ice
shelf melting of the AIS has increased (e.g., A. J. Cook et al., 2016; Dutrieux et al., 2014; Jacobs et al., 2011;
Pritchard et al., 2012; Rintoul et al., 2016; Silvano et al., 2018). However, as the observational record grows it
has become evident that the processes that allow this warm water to be delivered to the ice shelf vary in space
and time. Additional observations and modeling studies are required to better understand the processes
leading to CDW intrusions on the Antarctic shelf and to quantify their impact on slow and rapid changes
in ice sheet evolution. In this section, CDW is used sensu lato; that is, CDW comprises the CDW sensu
stricto, as well as all the deep‐water masses that derive from it such as modified CDW (mCDW), and
Warm Deep Water in the Weddell Sea (Nicholls et al., 2009; Vernet et al., 2019). Processes driving melt
include those that deliver warmer offshore water onto the continental shelf (section 2.2.1), interaction of
warmer water with ice shelves and icebergs (section 2.2.2), and microscale dynamics of sub‐ice shelf melting
(Jenkins et al., 2016; Schmidtko et al., 2014) (section 2.2.3).
2.2.1. Processes Across the Shelf Break: Winds, Eddies, and Waves
Ocean processes can significantly impact the AIS by transporting heat to the Antarctic coastline and so con-
trol the magnitude and distribution of basal melting under ice shelves. Factors involved in controlling the
oceanic heat supply across the continental shelf to the AIS include (1) the topography of the continental shelf
and slope (St‐Laurent et al., 2013), (2) the presence and location of the Antarctic Slope Front (ASF)
(Schmidtko et al., 2014), (3) the micromesoscale eddy field (Stewart & Thompson, 2015), and (4) atmospheric
processes relating to changes in the strength and position of the westerly winds (Heywood et al., 2016;
Jenkins et al., 2016). In areas where the clockwise Antarctic Circumpolar Current (ACC) takes a southerly
path and comes close to the continental shelf, warmer off‐shelf water such as CDW is more readily available,
transferring heat across the continental slope (Figure 3). However, the combination of the continental slope,
the ASF, and its associated westward current provides a barrier to heat transport around much of the con-
tinent. Where the ASF is weak (e.g., Amundsen Sea, West Antarctic Peninsula; Whitworth II et al.,1998;
Thompson et al., 2018), this barrier is less effective and heat can more easily access the continental shelf
(Schmidtko et al., 2014) (Figure 3b). But even where a strong ASF is present, the barrier is broken in key
locations by processes such as eddy transport, overturning circulation and changes in the position and
strength of the westerly winds. Once on the continental shelf the effect of the heat input by CDW transport
is primarily controlled by the salinity budget. The balance between sea ice formation and melt, precipitation
minus evaporation (E −P), lateral advection and freshwater input from melting ice shelves controls the stra-
tification and the buoyancy‐driven circulation, which in turn regulates the amount of heat that is available to
further promote melting of ice shelves. While the notion that oceanic heat supply across the continental shelf
drives AIS melt is well known (Jacobs et al., 1996), new research has been defining what these processes are,
and their relative contribution to cross‐shelf heat transport.
Foremost among these processes is the role of mesoscale eddies (size of order 4 to 40 km) in transporting
warm and saline CDW from the open Southern Ocean onto the Antarctic continental shelf (Heywood
et al., 2016; Stewart & Thompson, 2013). These eddies are relatively small on the Antarctic continental shelf,
due to the combination of a strong Coriolis parameter and weak stratification, meaning that most regular
ocean models are unable to resolve the eddy fluxes (Stewart & Thompson, 2013). However, the increase
in computing power over the last decade has allowed for eddy resolving simulations to be run
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(e.g., Mack et al., 2019), even over regions as large as the Southern Ocean. Variations in topography are
known to localize eddy fluxes (Couto et al., 2017; Moffat et al., 2009). Recent modeling work suggests that
both eddies and tides are important for cross‐slope heat transport, where eddy heat flux is the dominant
driver of heat transport, and this is modulated by tides and wind‐driven Ekman transport (Stewart
Figure 3. The connection between Antarctic ice mass loss and ocean forcing. (a) The black line shows the southern
boundary of the ACC defined as the southernmost extent of the upper CDW. Gravimetric Mass Balance data for
2002–2016 shows ice mass changes across the continent (Groh et al., 2019), which is highest in West Antarctica where the
Antarctic continental shelf and southern boundary of the ACC are in close proximity. (b) Map of the Southern Ocean
bottom potential temperature highlights the presences of warmer waters along the West Antarctic margin, based on the
WOCE global hydrographic climatology. The black line shows the 1,000 m bathymetry contour.
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et al., 2018). More observations and high‐resolution circum‐Antarctic models are required to constrain
uncertainty in the exact processes that move heat on to the shelf and the magnitude of their impact.
The cross‐shelf transport of heat may also be related to the larger‐scale ocean overturning circulation
through the formation of Antarctic Bottom Water (AABW). AABW is one of the main water masses that
comprise the global meridional overturning circulation (MOC) and is derived from dense shelf water
(DSW) that is formed on the Antarctic continental shelf by brine rejection. During sea ice formation, brine
rejection increases the density of very cold near‐surface shelf waters, which sink and form DSW. This dense
water flows off the continental shelf, then sinks along the continental slope, mixing with ambient waters,
and is subsequently exported northward as AABW at the bottom of the Pacific and Atlantic basins (Orsi
et al., 1999; Van Sebille et al., 2013). In the Weddell Sea, sea ice formation over the continental shelf leads
to the formation of High Salinity Shelf Water (HSSW). In austral winter, the HSSW flows beneath the
Filchner‐Ronne Ice‐Shelf, and because of the greater pressure at depth, HSSW melts the ice beneath the
ice shelf and produces Ice Shelf Water (ISW). The HSSW and ISW then flow down the continental slope
to produce an AABW precursor (Nicholls & Østerhus, 2004). A variety of processes can facilitate or inhibit
relatively warm CDW incursions onto the shelf, which in turn influence the rate of DSW production and
thus AABW formation. A reduction in AABW formation has been shown through modeling experiments
to lead to decreases in the oceanic meridional heat transport to high southern latitudes, increased stratifica-
tion, and ultimately increased warming in the subsurface ocean (Menviel et al., 2010) that can potentially
drive increased sub‐ice shelf melting of the AIS (Golledge et al., 2014). The feedbacks associated with AIS
meltwater‐driven suppression of AABW formation, as a consequence of AIS melt to the ocean, are reviewed
in section 6.2.
A negative feedback mechanism relating to cross‐shelf heat transport has been proposed by modeling work
that resolves DSW formation on the Antarctic shelf (Snow et al., 2016). The negative feedback mechanism is
based on the balance between off‐shelf dense water transport and on‐shelf flux, which is composed of fresh
cold AABW and warm saline CDW, respectively. In the model configured by Snow, Hogg, et al. (2016),
reductions in DSW formation and export were balanced by reduced on‐shelf transport of CDW. However,
improved understanding of the interaction between DSW formation and CDW incursions is required.
This should be addressed in the coming years by gathering observational data and by performing sensitivity
experiments using high‐resolution ocean/sea ice models.
The strength and position of the southern hemispheric westerly and/or easterly winds can also impact heat
transport to the shelf through their influence on the oceanic circulation changes and influence subsurface
temperatures on the Antarctic shelf. Large‐scale wind forcing affects the amount of heat available on the
shelf through changes in thermocline depth that modulate access of CDW onto the shelf (Dutrieux
et al., 2014; Galton‐Fenzi et al., 2016; Jacobs et al., 2013; Schmidtko et al., 2014; Spence et al., 2014). At smal-
ler scales, reductions in easterly wind stress near the Antarctic coastline act to alter the sea surface height,
with internal adjustments transporting heat onto the shelf (Spence et al., 2014). This process is an expression
of wider influences, with East Antarctic wind perturbations focusing their activity onto the western side of
the Antarctic Peninsula via barotropic Kelvin waves (Spence et al., 2017), due to the steep bathymetry and
strong offshore temperature gradients at that location. Overall, the complex interactions between eddies,
topographic variations, winds, and dense water formation each act to modify the balance of heat flux onto
the Antarctic shelf. Given the potential importance of these interactions in the future, such processes require
urgent research to improve projections of AIS variability.
2.2.2. Processes Influencing the Heat Content Beneath Ice Shelves
Ice shelves around Antarctica can be broadly classified into two types (Jacobs et al., 1992): cold cavity ice
shelves, which are characterized by considerable areas with basal freeze‐on of marine ice, such as the
Ross Ice Shelf and Filchner‐Ronne Ice Shelf, and warm cavity ice shelves, where little or no DSW is formed
and CDW is ubiquitous across the shelf, such as Pine Island Glacier Ice Shelf, and the West Antarctic
Peninsula ice shelves (e.g., Bernales et al., 2017) (Figure 4). While sub‐ice shelf melting occurs in both cold
cavity and warm cavity ice shelves, the largest melt rates are observed in warm cavity ice shelves, due to the
much higher heat content of CDW (e.g., Paolo et al., 2015; Pritchard et al., 2012). The particular water mass
that drives melting below each ice shelf results from a complex variety of processes, such as atmospheric for-
cing (sea ice formation and melt, precipitation and evaporation, and variation in wind direction and
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strength), bathymetry, meltwater outflows, and tides. These processes
control stratification and the buoyancy‐driven circulation, which in turn
regulate the amount of heat that is available to melt ice shelves. Ice
shelf‐ocean models are critical for researching these processes, as the
coastal and sub‐ice shelf region is typically very difficult to access.
However, there is a lack of understanding of the boundary conditions,
namely, the seafloor bathymetry. While recent efforts have significantly
improved understanding of the geometry of some ice shelf cavities (e.g.,
Arndt et al., 2013; Fretwell et al., 2013; Timmermann et al., 2010), much
of the Antarctic continental shelf and most ice shelf cavities remain very
poorly constrained. As a result, the ability of models to simulate melting
beneath ice shelves is hampered, particularly where that melting is heav-
ily influenced by bathymetric‐controlled flows. As more observations are
available and models become more realistic, it has become clear that the
relative contributions of the many processes influencing melting, and
the interactions between them, vary around Antarctica. Such variation
is responsible for the observed heterogeneity in the thinning response of
Antarctic ice shelves.
Atmospheric forcing explains a large portion of observed variability in the
temperature and salinity properties of Antarctic shelf waters and sea ice
production (Petty et al., 2013, 2014). At the regional scale, offshore direc-
ted winds, such as katabatic winds advect ice away from the coast, result-
ing in large heat fluxes from the ocean to the atmosphere. This
wind‐driven and buoyancy flux process maintains high sea ice production
rates in many Antarctic coastal polynyas (Tamura et al., 2016) and results
in the production of cold DSW, which can reduce basal melting near the
front of the ice shelf (St‐Laurent et al., 2015). At the Mertz Glacier
Tongue, prior to its break off in 2010 (Giles, 2017), increased basal melting
of floating ice occurred during periods when the nearby polynya was
much weaker (Cougnon et al., 2013; Holland et al., 2015). Changes to
the local icescape, in particular the positioning of grounded icebergs, for
example, after the calving of the Mertz Glacier Tongue (Kusahara
et al., 2011), have been shown to play a role in defining the properties of
the shelf waters and the rate of ice shelf basal melting (Cougnon
et al., 2017).
In other locations where cold shelf waters are observed, upstream precon-
ditioning of shelf waters, in particular freshening due to ocean‐driven
melting, is an important process controlling the properties of Antarctic
shelf waters (Kusahara et al., 2017; Williams et al., 2016), as well as deep
ocean properties through the reduction of AABW export (Hellmer, 2004).
Along the Sabrina Coast, East Antarctica, where DSW does not result in
the formation of AABW, the high glacier flow speed of the Totten
Glacier has been linked to basal melting, driven by a pool of warm water
on the continental shelf (Rintoul et al., 2016; Silvano et al., 2017).
Observations and models suggest that the magnitude of basal melting of
the Totten Glacier is modulated by the properties of DSW formed in the
nearby Dalton Ice Tongue polynya (Gwyther et al., 2014; Khazendar
et al., 2013). Even in very active polynyas, the freshening caused by the
input of meltwater from upstream can partially offset the density
contribution from brine released during sea ice formation, preventing
top‐to‐bottom convection and further promoting access of warm water to the glacier cavity, as found in
the Dalton Ice Tongue and Amundsen polynyas (Silvano et al., 2017).
Figure 4. Cross‐shelf exchange and basal melt. (a) Cold cavity ice shelf: A
sharp ASF prevents the access of CDW onto the shelf. Moderate
(e.g., Ross Sea) to large (e.g., George V‐Adelie Land) buoyancy fluxes at the
surface result in the formation of DSW, which drives low basal melt
rates (after Jenkins et al., 2016; Petty et al., 2014; Schmidtko et al., 2014;
Thompson et al., 2018). (b) Warm cavity ice shelf: Weak or absent
ASF allows CDW to spread onto the continental shelf. DSW formation does
not occur and CDW drives high basal melt rates. (c) Troughs and
depressions at the shelf break can also facilitate access of oceanic heat on to
the continental shelf and further to the coast, such as at Pine Island
Glacier. Here, in the Amundsen Sea Embayment bathymetric troughs and
depressions facilitate the exchange of CDW onto the shelf and toward
the grounding line of the ice shelves (e.g., Thoma et al., 2008). The presence
of CDW on the shelf drives melting of the nearby ice shelves
(e.g., Jacobs et al., 2011), with freshwater release potentially leading to a
shutdown of DSW formation where such dense waters are produced
(e.g., Silvano et al., 2018; Williams et al., 2016). With DSW no longer
exported across the shelf break in former areas DSW production the
ASF weakens further promoting CDW access to the shelf
(e.g., Thompson et al., 2018). Large, recurrent polynyas exist on the
Amundsen Sea Embayment shelf; however, no DSW is produced there
probably as a consequence of CDW presence. Panels (a) and (b) based on
Jacobs et al. (1992).
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There is increasing evidence that the continental shelf bathymetry, in particular submarine troughs and
depressions (Figure 4c), is key in allowing the warm water to reach the ice front. For example, flow iner-
tia has been shown to initiate exchange of along‐slope currents onto the shelf when in the presence of a
trough that crosscuts an offshore curving shelf break (Dinniman & Klinck, 2004; Klinck &
Dinniman, 2010). Cross‐shelf exchange has also been shown to result from interaction between an off
shelf‐sourced Rossby wave and the topography of a shelf breakthrough (St‐Laurent et al., 2013).
Conversely, seabed ridges can also block the deepest warmest waters from reaching the grounding line
(De Rydt et al., 2014; Dutrieux et al., 2014; Jenkins et al., 2010; Muto et al., 2016). The local circulation
on the shelf, whether wind (Dinniman et al., 2015; Stewart & Thompson, 2015) or buoyancy driven
(Snow et al., 2016), also plays a role in delivering heat to the ice shelf front. A cyclonic gyre on the con-
tinental shelf helps to transport warm water from the shelf break to the ice shelf front both at the Amery
Ice Shelf in East Antarctica (Galton‐Fenzi et al., 2012; Herraiz‐Borreguero et al., 2016; Liu et al., 2017),
and in the Bellingshausen Sea (Zhang et al., 2016). However, the role of the gyre located in Pine Island
Bay (Thurnherr et al., 2014) in delivering water to the Pine Island and Thwaites Glaciers is unknown
(Heywood et al., 2016). Along the Western Antarctic Peninsula, a projected increase in westerly wind
strength could increase mixing (Brearley et al., 2017), thus enhancing the heat content in shallower water,
leading to decreased melting below deep ice shelves and increased melting below shallow ice shelves
(Dinniman et al., 2012).
The meltwater outflows from warm cavity ice shelves can impact melting. Fine‐resolution observations
(~0.5 km spacing between conductivity temperature depth (CTD) stations) collected along the Pine Island
Ice shelf front have recently shown that as these buoyant meltwater outflows rise they can mix laterally
via centrifugal and gravitational instabilities (Naveira Garabato et al., 2017) and eventually spread at the
base of the winter mixed layer. Adding freshwater to the surface of the ocean has the potential to increase
sea ice formation by raising the freezing point; however, it is the suppression of convective overturning that
is important. These outflows contribute to the preconditioning of polynyas by increasing the near‐surface
stratification and reducing the penetration of the deep convection that forms DSW (Williams et al., 2016).
Weaker deep convection leads to less heat loss from the ocean to the atmosphere and can lead to warmer
waters accessing the ice shelf cavity, possibly further accelerating the mass loss from the ice sheet.
Sub‐ice shelf melting responds to the amount of heat flux present at the ice‐ocean interface, which is in turn
a function of cavity circulation and turbulent mixing. For higher ocean temperatures, stronger melting
releases more buoyant meltwater, thus increasing the overturning circulation within the ice shelf cavity
(Gwyther et al., 2016; Holland et al., 2008). As stronger circulation increases turbulent heat transfer to the
ice shelf, tides also impact melt rate. This occurs through both tidal rectification, where the tidal current var-
ies as a function of Earth's rotation and the topography above the continental slope, as the periodic oscilla-
tion of tidal currents leads to greater mean circulation strength (Padman et al., 2018), and through increased
turbulence, which results in stronger mixing of heat and salt across the ice‐ocean boundary layer (Mueller
et al., 2012). The net impact of tides on basal melting is to increase the total basal mass loss and to change
the spatial distribution of melting (Gwyther et al., 2016). On a broader scale, the rate at which the cavity
is flushed and reventilated affects the heat budget within the cavity and thus the magnitude of basal melting
(Holland, 2017).
Around Antarctica, basal melting contributes nearly twice as much to the total mass loss as iceberg calving
(Liu et al., 2015). However, ice shelves that are currently thinning have an approximately even mass loss con-
tribution from basal melting and calving, as a result of more frequent calving events. The filling of preexist-
ing crevasses with surface meltwater is suggested to have led to the fragmentation and collapse of the Larsen
A and B ice shelves (Scambos et al., 2004). The loss of these ice shelves has been linked to subsequent accel-
eration of tributary glaciers (Scambos et al., 2004). However, the loss of floating ice shelf area does not neces-
sarily lead to increased glacier velocities. There is an area of the ice shelf known as the “passive ice shelf ”
(Fürst et al., 2016) that can be removed without damaging the integrity of the remaining ice shelf (see
section 3.1.3). However, ice shelves in the Amundsen and Bellingshausen Seas were shown to not be pro-
tected by a band of passive ice. The role of a sea ice buffer and fast ice buttressing in maintaining the struc-
tural integrity of the frontal ice shelf region is relatively underexplored. Recent research suggests that the
seasonal absence of sea ice can remove the buffer from ocean swell and allow increased wave‐induced ice
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shelf flexure, which likely contributed to the disintegration of Larsen A, B, and Wilkins Ice Shelves (Massom
et al., 2018). The mechanical buttressing of landfast sea ice upon the front of ice shelves has been shown to be
the primary contributor for controlling seasonal ice shelf velocity within 50 km of the calving front (Greene
et al., 2018).
There is a strong effort to assess the role of atmospheric‐driven processes, such as hydrofracturing, in driving
grounding line retreat (see section 3.1.3), as demonstrated by the development of models of ice shelf rifting
and calving (Benn et al., 2017) and of iceberg evolution (Stern et al., 2017). The formation of icebergs through
rifting and fracturing processes is still poorly understood and difficult to model, particularly on larger spatial
scales and longer time scales. As a result, this process is often excluded from global climate models, although
it may be a very important process, in particular for achieving high rates of sea level rise not observed in the
present but recorded during past deglaciations (e.g., Meltwater Pulse 1A) and warm climate periods (e.g.,
Last Interglacial (LIG) period) (see sections 4.2.2 and 4.2.3).
Low‐frequency intrinsic ocean variability, generated internal to the ocean and in response to stochastic
forcing, also has the potential to drive complex changes in ocean conditions (Leroux et al., 2018;
O'Kane et al., 2013). Flushing and ventilation of the ice shelf cavity by periodic changes in ocean condi-
tions such as thermocline depth, resulting from climate teleconnection modes (Holland et al., 2019;
Jenkins et al., 2018) or intrinsic ocean variability (Gwyther et al., 2018; Huneke et al., 2019), will influ-
ence ice shelf mass loss. This has been demonstrated in a realistic model of the Totten Ice Shelf
(Gwyther et al., 2018) and idealized models (Holland, 2017; Snow et al., 2017). While assessments of
future mass loss from ice shelves can be conducted (e.g., Hellmer et al., 2012), intrinsic ocean variability
and low‐frequency climate models present a challenge for accurately simulating present day mass balance
trends (e.g., Christianson et al., 2016) and projecting the future response of the AIS to anthropogenic cli-
mate forcing.
2.2.3. Microscale Interactions
The accelerated mass loss around Antarctica (see section 5.1) has been largely attributed to the intrusion of
warm and salty CDW under ice shelf cavities and interaction with ice shelves through the ice‐ocean bound-
ary layer (Jacobs et al., 2012; Jenkins et al., 2016, 2018; Khazendar et al., 2013; Payne et al., 2004; Picard
et al., 2012; Swingedouw et al., 2009). Understanding the processes that control melting at the smaller scale
is important for accurate modeling and prediction of ice shelf basal melting; however, in situ investigation of
the ice/ocean interface is limited (Jenkins et al., 2010; Stanton et al., 2013) due to logistical difficulties
(Galton‐Fenzi et al., 2016). As a result, models and laboratory experiments are critical in investigating the
key processes that control microscale interactions: turbulence and convection.
The underlying microscale dynamics of ice melting involve the transport of both heat and salt in a thin
boundary layer at the ice face (Gayen et al., 2016; Kerr & McConnochie, 2015; Mondal et al., 2019; Wells
& Worster, 2011). Ocean modeling studies can only resolve the flow field at scales in the order of 100 m
and must rely on parameterizations to represent small‐scale boundary layer processes to infer the melt rate.
The conditions inside ice shelf cavities change the melt rate by controlling ocean‐ice interactions. The sea-
water temperatures inside ice shelf cavities vary from the freezing point temperature of seawater at the loca-
tion (i.e., −2.3°C to −1.7°C for a cold cavity) to 0°C (for a warm cavity), with the salinity around 34 practical
salinity unit (PSU) (Dutrieux et al., 2014; Jacobs et al., 2013, 2011, 2012; Jenkins et al., 2010; Nicholls
et al., 2009; Payne et al., 2004). At these low temperatures, the primary transfer of solute to the ice lowers
the melting point of the ice and enhances the heat transfer across the ice‐ocean interface. This process is
more accurately described by dissolution (Wells & Worster, 2011; Woods, 1992), meaning that both heat
and salt fluxes must be considered in estimating the rate of ice shelf mass loss.
Heat and salt transfer occur through multiple ocean layers against the ice face. Diffusive fluxes dominate at
the ice‐ocean interface, with the flux controlled by diffusive boundary layers on the order of 1 mm thick. The
thickness of the salinity layer is an order of magnitude smaller than the thermal boundary width due to the
large differences in molecular diffusivity between heat and salt (Josberger, 1983; McPhee, 1981). Because of
the small spatial extent of the layer dominated by diffusive fluxes, it is not possible to make direct measure-
ments of these fluxes in the field. Therefore, we often rely on model simulations or parameterizations to esti-
mate these fluxes.
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Beyond the diffusive layer (further away from the ice interface), fluxes are usually dominated by turbulent
processes in the outer layer which has a thickness in the order of 10 m. The turbulence is due to a combina-
tion of natural convection from buoyant melt water, strong shear ambient flows, internal waves, and ocean
tides (Cenedese & Gatto, 2016; Gayen et al., 2016; Keitzl et al., 2016). Next to the ice face, meltwater is posi-
tively buoyant compared to the ambient seawater and forms a convective plume. This plume is turbulent and
acts to amplify melting by transferring heat to the ice surface from the surrounding ocean (Figure 5) (Gayen
et al., 2016; Kerr & McConnochie, 2015; McConnochie & Kerr, 2017; Mondal et al., 2019; Slater et al., 2016).
Such convection‐driven melting is not considered in present parameterizations used for ocean modeling
(McConnochie & Kerr, 2017) and represents a key challenge in future models of the ice melting process.
Future improvements in model parameterizations need to include stratification in the cavity water, which
can impact the melt boundary layer in a complex way by forming of double‐diffusive layers and enhancing
the rate of mixing of melt water with the surrounding salty water (Gayen et al., 2015; Huppert et al., 1980;
Huppert & Turner, 1978, 1980). In addition, turbulent heat transport across the ice‐ocean boundary layer
is significantly enhanced by roughness of the ice shelf bottom surface, usually caused by the small ice crystal
accretion (frazil) (Robinson et al., 2017), as well as channels at the underside of an ice shelf formed by ocea-
nic melting (Alley et al., 2016; Bindschadler et al., 2011; Mankoff et al., 2012) and possibly initiated by chan-
nelized subglacial meltwater outflow (Le Brocq et al., 2013), significantly enhances turbulent heat transport
across the ice‐ocean boundary layer (see section 4.2.3 for more on sub ice shelf channels). This can influence
the distribution of melting and freezing (Gwyther et al., 2015; Le Brocq et al., 2013; Stanton et al., 2013;
Vaughan et al., 2012) and even ice shelf stability (Alley et al., 2016). Estimates of turbulent‐driven mixing
and the magnitude of sub‐ice shelf melt are currently limited by our lack of understanding of basal friction.
Improving this understanding principally depends on measurements of basal roughness and the impact on
adjacent flow beneath ice shelves.
The complex interactions inside the boundary layer, including ice surface morphology, strongly influence
heat transport across the ice‐ocean interface and the rate of basal melting. Thus better understanding and
representation of these processes in ocean‐climate models are necessary to increase accuracy in the predic-
tion of future AIS evolution.
3. Ice Dynamic Processes
The AIS was traditionally viewed as a slowly evolving and passive component of the Earth system (Warrick
& Oerlemans, 1990). However, recent evidence indicates that the AIS responds to forcings on time scales of
Figure 5. Schematic of microscale circulation processes on the Antarctic continental shelf. Grounded ice sheets are
buttressed by floating ice shelves. Their basal melting is driven by the supply of salinity and heat from the Southern
Ocean via CDW and its derivatives and influences the formation of deep‐water masses and Antarctic Bottom Water
(AABW). Zoomed view of boundary layer processes by showing snapshot of simulated turbulent dissipation (log scale)
field underneath a sloping ice shelf base (Mondal et al., 2019).
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hours (e.g., tides) to millennia (e.g., glacial cycles) (Davis et al., 2018; Pattyn et al., 2018) and with significant
spatial variations, even within ice sheet drainage sectors (e.g., Konrad et al., 2017). In this section, we review
the shift in perspective of the AIS's dynamic interaction with other components of the Earth system in the
context of processes that drive ice discharge from the AIS. We discuss processes that have been proposed
to reconcile the evidence of past AIS change and sea level rise inferred from paleostudies, as well as
grounding line controls, including ice shelf buttressing and the marine ice sheet and ice cliff instabilities
(Figure 6). Ice flow processes, particularly basal sliding and ice deformation, that influence the ice mass
flux over the grounding line, are also discussed. We highlight advances in observational capabilities and
the development of numerical ice sheet models that improve understanding and prediction of the
dynamic AIS evolution and stability.
3.1. Grounding Line Controls
The grounding line (or zone) is the interface between the grounded ice sheet and floating ice shelves and is a
key control on ice discharge from the AIS (Weertman, 1974).
Grounding lines have been retreating since the last glacial period when they were near their peak extent
(The RAISED Consortium et al., 2014). Satellite evidence points to accelerated retreat in key sectors over
the last few decades. For example, Konrad et al. (2018) mapped grounding line migration for the AIS from
2010 to 2016 by combining surface elevation and mismatches in the surface elevation above mean sea level,
breaks in the ice sheet surface elevation slope, and the effect of tides. While much of the continent's ground-
ing lines showed no discernible change over this period, 3.3% of East Antarctic, 21.7% of West Antarctic, and
9.5% of Antarctic Peninsula grounding lines retreated. Grounding line retreat was found to be largest in the
Amundsen and Bellingshausen Sea sectors (Christie et al., 2016; Parizek et al., 2013; Rignot et al., 2014),
which are regions of significant ice sheet thinning and fast flow. Observations since the 1970s confirm
long‐term retreat in basins of West Antarctica and the Antarctic Peninsula (Mouginot et al., 2014), with
35 km retreat of Smith/Kohler Glaciers, 31 km retreat of Pine Island Glacier, and 14 km retreat of
Thwaites Glacier over the period 1992–2011 (Rignot et al., 2014).
Figure 6. Ice sheet dynamic processes. (a) Friction from topographic pinning points or ice shelves impacts ice speeds.
(b) Marine ice sheet instability (MISI). The geometry of the bed topography plays a critical role in ice flow and ice sheet
stability. Specifically, a retrograde bed sloping downward into the interior of the continent can lead to unstable
grounding line retreat when combined with melting of the base of the ice shelf from relatively warm ocean waters.
(c) Marine ice cliff instability (MICI). Ice cliffs can collapse forming icebergs as a result of rapid ice shelf disintegration
caused by meltwater ponding at the ice shelf surface and hydrofracturing. (d) Basal sliding, in the presence of subglacial
meltwater, is one of the main processes by which ice flows from the interior of the continent to the oceans. (e) Ice
deformation is the movement within and between ice crystals, which is a function of the different stress configurations
within the ice sheet.
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Grounding lines migrate in response to various forcing mechanisms. On long (centennial to millennial) time
scales, grounding lines migrate in response to glacial isostatic adjustment (GIA) due to ice loading
(section 4.2.2), and subglacial hydrology (section 4.2.4), long‐term fluctuations in climate forcings and sea
levels, and internal ice sheet dynamics. On time scales of hours to centuries, the ocean and atmosphere drive
grounding line migration predominantly through ice sheet and shelf melting that modifies ice shelf buttres-
sing, and may induce marine ice sheet and ice cliff instabilities.
3.1.1. Ice Shelf Buttressing
Ice shelves can stabilize the position of the grounding line, and ice mass flux across it, by applying a back
stress on the upstream ice through interaction with topographic pinning points, or interaction with topogra-
phy at basal or lateral boundaries—a process known as ice shelf buttressing (Doake et al., 1998; Dupont &
Alley, 2005a, 2005b; MacAyeal, 1987; Rignot et al., 2004; Scambos et al., 2004). Buttressing can alternately
act to slow ice flow (Gladstone et al., 2012; Humbert & Steinhage, 2011; Jenkins et al., 2010; Schmeltz
et al., 2001) or generate a highly localized flow regime, particularly in the case of basal topographic pinning
points (Favier et al., 2016). Several regional and idealized modeling studies (e.g., Berger et al., 2016; Favier
et al., 2012, 2016; Goldberg et al., 2009) have found that grounding line dynamics are strongly sensitive to
the buttressing provided by topographic pinning points.
Ice shelf buttressing is impacted by changes in ice shelf thickness and extent: With the thinning or calving of
floating ice, the buttressing effect is diminished. Past ice shelf recession and collapse has been linked to a
reduced buttressing effect, leading to increased mass flux. For example, the collapse of Larsen B ice shelf
in the Antarctic Peninsula in 2002 led to increased ice shelf flow and ice front retreat in the subsequent dec-
ade, as well as acceleration of the tributary Leppard and Flask Glaciers (Khazendar et al., 2015). Ice shelf
thinning and the reduction in buttressing effect that occurred in Larsen B was potentially sparked by the ear-
lier collapse of the neighboring Larsen A ice shelf in 1995 (Albrecht & Levermann, 2014).
Recent work has highlighted the importance of the ice shelf buttressing effect in determining AIS stability
(Fürst et al., 2016). In their analysis, Fürst et al. (2016) delineated regions of ice shelves that control buttres-
sing from regions where the ice shelf could be removed without impacting flow—the latter denoted “pas-
sive”portions of the ice shelf. A total of only 13.4% of Antarctica's total ice shelf area could be removed
without impact to the grounding line position and ice mass flux across it. Glaciers and ice streams in the
Amundsen and Bellingshausen Sea sectors are particularly vulnerable to dynamic changes in the event of
ice shelf thinning or calving, having passive ice shelf areas of only 7.9% and 5.3%, respectively.
A regional modeling study on the Lazarev and Roi Baudouin Ice Shelves in East Antarctica investigated the
potential for future sea level rise as a result of ice discharge from these ice shelves (Favier et al., 2016). The
study found that the buttressing effect of pinning points beneath the ice shelves led to an overall decrease of
10% in sea level rise from ice discharge compared with when the pinning point was removed. The results
highlight the importance of high‐resolution bed topography in numerical modeling, given that ice shelf sta-
bility can be controlled by kilometer‐scale pinning points.
Ocean‐driven melt of ice shelves is responsible for more than half of Antarctic ice shelf mass loss (Depoorter
et al., 2013; Liu et al., 2015; Paolo et al., 2015; Rignot et al., 2013) and is a strong driver of changes in the but-
tressing effect. Reese, Gudmundsson, et al. (2018) diagnosed the dynamic response of grounded ice and par-
ticularly the corresponding mass flux across the grounding line, to specific spatial patterns of ice shelf mass
loss. Their results highlighted a potential for “tele‐buttressing,”whereby relatively small changes to ice shelf
thickness can initiate an ice dynamic response far upstream of the thickness perturbation. The ice shelves
most vulnerable to change have ice‐ocean fronts close to the continental shelf or in regions prone to intru-
sion of warm ocean waters (see section 2.2.1).
Recent observations of the Totten Glacier in East Antarctica suggest that ice shelves work in concert with
ocean forcing to modify the ice flow regime and surface elevation through back stress generated from topo-
graphic pinning points Roberts et al., 2018). Specifically, when the seawater in contact with the ice shelf base
is cool, a reduction in melting leads to ice shelf thickening, increased back stress at the pinning point, and an
overall decrease in ice velocity. The converse is true when the ocean‐forced melting is high. In studies of ice
sheet retreat, topographic pinning points have also been found to play a leading role in stabilizing the
grounding line position (e.g., Seroussi et al., 2017; Schlegel et al., 2018), highlighting the importance of but-
tressing and back‐stress processes to ice shelf stability.
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Numerical model studies estimating the potential future contribution of the AIS to sea level rise will
benefit from continued improvement in the spatial and temporal resolution of topographic data beneath
the ice sheet and shelves (e.g., Morlighem et al., 2020). The advent of high‐resolution ice surface
elevation data, such as the submeter resolution satellite imagery used in the construction of the
Reference Elevation Model of Antarctica (REMA, Howat et al., 2018), may aid in identifying or charac-
terizing pinning points, where they have a surface expression (see e.g., Budd & Carter, 1971; J. Roberts
et al., 2018).
3.1.2. MISI
MISI (Figure 6) is an important process in evolution of the AIS. It describes the susceptibility of a
marine‐based ice sheet with a bed that slopes in an upstream direction (retrograde bed slope) to grounding
line retreat especially under a warming climate (Mercer, 1978; Thomas, 1979; Weertman, 1974). Specifically,
MISI occurs when the grounding line of a glacier flowing over a retrograde bed slope retreats, which acts to
reduce the ice shelf buttressing effect and results in further grounding line retreat and accelerated ice dis-
charge. If the ice thickness at the grounding line increases when the grounding line retreats, then the ice dis-
charge will also increase, and the grounding line has to retreat further upstream, thereby also reducing the
glacier's catchment area and thus the area of snow/ice accumulation. Schoof (2007, 2011) used analytical
methods to verify MISI, demonstrating that for an ice sheet/ice shelf system that is not confined by an
embayment, ice discharge at the grounding line is extremely sensitive to ice thickness at the grounding line,
proportional to approximately its fifth power. For an ice sheet fringed by an embayed ice shelf, modeling
shows this relationship is expected to be less strong due to the buttressing effect from interaction of ice with
topography at the lateral boundaries (Goldberg et al., 2009).
MISI is triggered by an initial retreat of the grounding line, which can be activated by a variety of processes,
for example, ice shelf thinning, sea level rise due to thermal expansion of seawater, or the sea level finger-
print from melting glaciers in the middle to low latitudes and high latitudes of the Northern Hemisphere.
The grounding line retreat leads to a run‐away effect (e.g., see Joughin & Alley, 2011; Vaughan &
Arthern, 2007) associated with accelerated ice discharge (Pollard & DeConto, 2009, 2012). Several WAIS gla-
ciers that satisfy the geometric requirements of MISI are potentially already undergoing MISI (Favier
et al., 2012; Joughin et al., 2014).
In the Amundsen and Bellingshausen Sea sectors, MISI is believed to be driven by incursions of CDW onto
the continental shelf (Holland et al., 2010; Jacobs et al., 2011; Jenkins et al., 2018; Nakayama et al., 2013),
potentially as a result of broader changes in patterns of wind forcing (Thoma et al., 2008). In the event that
CDW incursions decrease in the future, glaciers already undergoing MISI might continue to experience
grounding line retreat even after the forcing is removed, as recently inferred from numerical modeling by
Waibel et al. (2018). However, sustained retreat of MISI‐prone glaciers is not guaranteed and depends on
complex interactions and feedbacks between the ice sheet and other Earth and climate systems. For
instance, the response time scales vary widely depending on the subglacial setting (Levermann &
Feldmann, 2019) and ice‐ocean interactions such as the magnitude of ocean‐driven melting and the supply
of warm ocean waters to the grounding line (Joughin et al., 2014; Seroussi et al., 2017). Furthermore, ground-
ing line retreat is highly sensitive to, and can even be stabilized by, ice‐solid Earth interactions such as the
effect of significant topography, pinning points, or lateral buttressing (Gudmundsson, 2013; Jamieson
et al., 2012; Schlegel et al., 2018; Seroussi et al., 2017; Thomas, 1979) or the bed slope and depth below sea
level (Golledge, 2014), as well as the negative feedbacks from solid Earth rebound and sea level rise
(Larour et al., 2019).
Challenges remain in using numerical models to predict the timing and magnitude of potential collapse of
marine‐based parts of the AIS as a result of MISI. One challenge is the treatment of ocean melt rates that
drive MISI. These are often parameterized in ice sheet models as a function of the ice shelf draft, accounting
for greater melt rates near the grounding line than at the ice‐ocean front (Asay‐Davis et al., 2016; Walker
et al., 2008). Options exist that allow for a more sophisticated treatment of sub‐ice shelf melt rates, such
as that provided by the Potsdam Ice‐shelf Cavity mOdel (PICO) (Reese, Albrecht, et al., 2018), which simu-
lates the vertical overturning circulation in ice shelf cavities, using temperature and salinity fields to calcu-
late sub‐ice shelf melt rates. Plume models of shelf meltwater (e.g., Jenkins, 2016), may also provide realistic
melt rate parameterizations by accurately representing melting in the ice shelf cavity. A combination of
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PICO with a plume model (Pelle et al., 2019) has been shown to capture ocean melt rate patterns and mag-
nitudes that agree well with observations.
Sub‐shelf melt rate estimates from ocean models are continually improving due to the addition of new phy-
sical processes (e.g., Galton‐Fenzi et al., 2012), improved boundary conditions, and model capabilities (e.g.,
Gwyther et al., 2014; Nakayama et al., 2019) and may be directly incorporated as the ocean forcing in ice
sheet models. However, stand‐alone ocean models are limited in their ability to account for changes in
the geometry of the ice shelf cavity and moving grounding lines and thus cannot provide estimates of the
timing and magnitude of collapse of the marine‐based AIS. The effect of these issues is being explored in
efforts such as the Marine Ice Sheet‐Ocean Model Intercomparison Project (MISOMIP) (Asay‐Davis
et al., 2016), which aims to improve understanding of ocean‐ice shelf interaction for the eventual inclusion
in general circulation models, using coordinated modeling experiments across both standalone ocean and
ice sheet models and coupled ice sheet‐ocean modeling frameworks.
An alternative to parameterizing ocean melt rates in ice sheet models is coupling between ice sheet and
ocean models, which enables accurate treatment of the ice‐ocean boundary interactions and feedbacks that
drive MISI. Recent developments have been made in asynchronous ice sheet‐ocean model coupling, where
each component model is run independently for short periods of time, and the outputs (e.g., geometry, melt
rates, temperature, and salinity) are updated between models (Asay‐Davis et al., 2016; De Rydt &
Gudmundsson, 2016; Seroussi et al., 2017). Asynchronous coupling has the advantage of being relatively
computationally inexpensive as it does not require each component model to be solved at the same time step
where the time scales of evolution vary considerably (e.g., Goldberg et al., 2012). Nevertheless, this may
come at the expense of accurately representing the full spectrum of coupling time scales relevant to each
model component (Snow et al., 2017).
Synchronous high‐resolution coupling of an ice‐ocean model has also recently been undertaken (Jordan,
Holland, et al., 2018). Snow et al. (2017) used synchronous coupling to show that the ice sheet responds most
strongly to long ocean forcing time scales, and the ice volume above floatation changes in response to differ-
ent ocean forcing periods that are generally shorter for the coupled model than for the uncoupled parame-
terizations. In addition to modifying the time scales of the ice sheet response, using a synchronous approach
over an asynchronous one has the advantage of conserving heat, salinity, and mass, allowing for tidal deflec-
tions to be resolved.
Ice sheet models using a combination of the shallow ice approximation (SIA) and shallow shelf approxima-
tion (SSA) (De Rydt & Gudmundsson, 2016; Goldberg et al., 2012; Seroussi et al., 2017; Thoma et al., 2015)
coupled to ocean models have been used in experiments to assess the sensitivity of the ice sheet to changing
ocean temperatures and circulations. These studies consistently predict an enhanced ice sheet response (e.g.,
grounding line retreat or volume above floatation) in the absence of coupling between the ice and ocean
model components. This suggests that projections of the AIS response to climate change and susceptibility
to MISI require that spatial and temporal patterns in ocean melt rates as predicted by ocean models, and
coupled feedbacks between the ice and ocean systems, are sufficiently resolved.
A challenge in predicting the impact of MISI on the AIS is the treatment of grounding line dynamics and
migration. Recent developments in ice sheet model physics, parameterization of physical mechanisms,
and mesh resolution have driven improvements in grounding line dynamics (Nowicki & Seroussi, 2018).
For example, accurately resolving fine‐scale migrations of the grounding line in numerical model simula-
tions requires very fine mesh elements near the grounding line, which comes at a large computational cost
for the time scales typical of ice sheet dynamic evolution. New approaches have accounted for this problem
through subgrid parameterizations of friction or ocean melting (Feldmann et al., 2014; Gladstone et al., 2010;
Leguy et al., 2014; Seroussi & Morlighem, 2018; Seroussi et al., 2014) or adaptive mesh refinement (Durand
et al., 2009; Gladstone et al., 2010). High‐resolution full Stokes models of the grounding zone are more accu-
rate than models with simplified stress balance approximations (e.g., the SIA; Hutter, 1982) in yielding large
short‐term variations in the grounding line position (Drouet et al., 2013; Seddik et al., 2017). Models with
sliding relations that incorporate a dependency on the effective pressure have been able to smooth the tran-
sition between grounded and floating ice (Brondex et al., 2017; Gladstone et al., 2017; Jong et al., 2018; Robel
et al., 2016; Tsai et al., 2015), a problem that arises in models with less physically motivated sliding relations
(see section 3.2.1).
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3.1.3. Surface Melt and MICI
Surface meltwater in the form of supraglacial lakes, subsurface lakes, surface streams, and rivers (Bell
et al., 2018) (Figure 7) can potentially impact AIS mass loss directly by accelerating ice shelf disintegration,
which reduces buttressing back stresses and increases ice flow to the ocean (Scambos et al., 2004). Summer
melting in Antarctica is generally confined to the ice sheet edge (Tedesco & Monaghan, 2009) and is most
intense and widespread on the Antarctic Peninsula (Cook & Vaughan, 2010). Indeed, an increasing duration
of surface meltwater on ice shelves on the Antarctic Peninsula has been observed over the past fifty years
(Abram et al., 2013; Barrand et al., 2013; Vaughan, 2006). This is particularly the case in the eastern
Antarctic Peninsula (Pritchard & Vaughan, 2007; Scambos et al., 2004; van den Broeke, 2005) in response
to atmospheric warming (Scambos et al., 2000) and a result most vulnerable to meltwater‐induced
hydrofracture.
Active surface hydrology systems in East Antarctica ice shelves were previously considered to be rare, as
highlighted by Kingslake et al. (2017). However, an inventory from the peak of the 2017 melt season reveals
extensive and active surface lake formation in East Antarctica. Over 80% of the area of these lakes was
located on ice shelves, making the ice shelves vulnerable to meltwater induced fracturing (Stokes et al., 2019).
East Antarctic ice shelves also show evidence of surface hydrology processes (Langley et al., 2016, 2011;
Phillips, 1998), which may be driven by different processes (e.g., a positive feedback on surface melting by
wind‐albedo interaction; Lenaerts et al., 2017) but can affect ice shelf disintegration through similar hydro-
fracturing mechanisms to those observed on the Antarctic Peninsula.
Surface meltwater can trigger ice shelf collapse through meltwater induced fracture propagation, a process
referred to as hydrofracturing. Ice shelf breakup can result from a combination of both flexure stresses, due
to changes in stress load from filling and draining of surface lakes and hydrofracture (i.e., increasing crevasse
depths) (Banwell et al., 2013; Banwell & MacAyeal, 2015; MacAyeal & Sergienko, 2013; MacAyeal
et al., 2015). Direct observations of ice shelf flexure on time scales of weeks from the McMurdo Ice Shelf,
a small portion of the Ross Ice Shelf, due to changes in the surface meltwater load (Banwell et al., 2019) pro-
vide much needed constraints for implementing this process in ice sheet models. Developments in model
parameterization of hydrofracturing by Robel and Banwell et al. (2019) suggest that the time scales
Figure 7. Schematic of Antarctic surface hydrology features that can influence AIS dynamics. Surface melting is
enhanced due to reduced surface albedo, resulting in a positive melt feedback. Basal hydrology is impacted when
the surface meltwater reaches the subglacial environment through fractures. Fracturing develops through surface lake
and firn aquifer drainage. Hydrofracturing of the ice shelf can arise due to the loading and unloading of lakes. Direct
surface runoff into the ocean can change seawater density properties and potentially result in stratification that can trap
warming in the subsurface, enhancing basal melting of the ice shelf. Adapted from Bell et al. (2018).
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(weeks to years) of ice shelf collapse by hydrofracturing are limited by the rate of surface melting, with an
intrinsic limit defined by the flexual length scale of the ice shelf.
The impact of meltwater on the basal conditions of grounded ice is not well known for Antarctica compared
with the Greenland Ice Sheet. Rapid transfer of meltwater to the ice sheet base, via meltwater‐induced frac-
turing of crevasses, provides a mechanism for changing conditions at the ice‐bed interface (van der
Veen, 2007). However, Bell et al. (2017) suggest that efficient export of surface meltwater off the ice shelf,
through an extensive drainage network, can act to stabilize the ice shelf. Limitations on the depth of melt-
water ponds likely play an important role in preventing ice shelf collapse (Robel & Banwell, 2019).
However, the first evidence for meltwater‐induced speed up of glacier velocity on the Antarctic Peninsula
suggests that a direct coupling and positive feedback between atmospheric forcing and ice flow dynamics
can occur (Tuckett et al., 2019).
To address the difference between simulations of ice sheet retreat and paleoevidence for, and sea level
requirements of, major WAIS collapse, as well as retreat in portions of marine‐based sectors of the EAIS,
Pollard et al. (2015) proposed the MICI. This mechanism describes the vulnerability of ice shelves to both
atmospheric and oceanic warming. MICI includes the physical processes of cliff failure and hydrofracturing,
which can cause ice sheet instability with climatic warming: (DeConto & Pollard, 2016; Pattyn, 2018; Pollard
et al., 2015). Under MICI scenarios, atmospheric warming generates meltwater ponding at the ice shelf sur-
face, leading to extensive crevassing. Crevassed and damaged ice undergoes mechanical failure at lower cliff
heights than noncrevassed ice. Reduction in buttressing from this mechanical failure, in combination with
ocean‐driven ice shelf thinning from below, can excite grounding line retreat (Figure 8) and rapid ice sheet
collapse if sustained. By including the hydrofracture and cliff failure mechanisms, MICI has been used to
reconcile ice sheet simulations of past climate conditions within the geological record, especially for
warmer‐than‐present and Pliocene climates (Cook et al., 2013; Patterson et al., 2014; Williams et al., 2010;
Young et al., 2011) (section 5.2.3).
Under conditions representative of a warm Pliocene period, Pollard et al. (2015) simulated a MICI‐driven
grounding line retreat of large regions of the EAIS into the major subglacial basins. DeConto and
Pollard (2016) developed the representation of MICI further, coupling the ice sheet model to a regional cli-
mate model. To explore uncertainty in their parameterization of MICI, they perform a large ensemble ana-
lysis, fitting their model to LIG and Pliocene sea level targets of 3.6–7.4 and 10–20 m, respectively. Under an
LIG climate, and without the MICI mechanism, subsurface ocean warming greater than 4°C is required to
simulate WAIS retreat. However, by accounting for the additional ice‐ocean‐atmosphere feedback mechan-
ism described by MICI, the model predicted WAIS collapse with ocean forcing of 3°C.
Under Pliocene conditions, the model with the MICI mechanism simulated an Antarctic contribution of
<12 m to GMSL. This estimate is consistent with a Pliocene‐maximum contribution of no more than 13 m
from the AIS during the mid‐Pliocene inferred from Pliocene ice sheet simulations constrained by the oxy-
gen isotope composition of benthic foraminifera (Gasson et al., 2016). DeConto and Pollard (2016) used the
tuned MICI parameterization in simulations of future AIS sea level contributions under the IPCC RCP8.5
scenario, finding a GMSL contribution of 15.65 ± 2.00 m by 2500. The rates of sea level rise corresponding
to the DeConto and Pollard (2016) assessment are at the higher end of estimates based on natural climate
forcing of sea level rise from paleoclimate data, which reveal an average of 1–2 m per century across past
deglaciations, and 4–5 m per century during rapid pulses of sea level rise (Deschamps et al., 2012b; Grant
et al., 2014; Rohling et al., 2013, 2019).
Observational evidence to support MICI is scarce and is currently limited to one study of the maximum
water depths and shapes of iceberg‐keel plow marks eroded during the post‐Last Glacial Maximum
(LGM) deglaciation on the West Antarctic continental shelf (Wise et al., 2017) and another based on glacio-
logical evidence of rapid ice cliff calving in Greenland (Parizek et al., 2019). Wise et al. (2017) found
iceberg‐keel plow marks on the floor of the Pine Island Trough in West Antarctica that may indicate rapid
deglacial ice sheet retreat into Pine Island Bay. Radiocarbon dating of foraminifera used to constrain the tim-
ing of the formation of grounding zone wedges present in this paleo ice stream trough indicates that the
retreat commenced around 12.3 ka and terminated before 11.2 ka (Kirshner et al., 2012; Larter et al., 2014).
The keel plow marks in the study of Wise et al. (2017) indicate smaller icebergs with pinnacle‐shaped keels
calved from an ice cliff front, rather than large tabular icebergs calved from the regions of the current ice
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shelf configuration. The inferred calving‐margin thicknesses of the bergs during this period of retreat are
consistent with the threshold predicted to trigger ice shelf collapse as a result of MICI. Observed calving
of ice cliffs ~100 m above sea level in Helheim Glacier, East Greenland (Parizek et al., 2019), provides
evidence for rapid ice loss by a rotational slumping mechanism that was suggested in a modeling study
by Bassis and Walker (2012). The observations of Parizek et al. (2019) show how surface melting at
the glacier front, driven by atmospheric forcing, results in basal crevassing that promotes rapid calving
and ice loss (Figure 8). This slumping mechanism is relevant to smaller cliffs of 100 m in height above sea
level, as opposed to the 200 m high ice cliffs collapsing under brittle failure in the model of
Figure 8. Schematic of marine ice cliff instability. (a) Warm ocean water access drives basal melting, while surface
melting causes hydrofracturing and surface crevasse formation. (b) Loss of buttressing due to enhanced basal melting
or calving results in grounding line retreat, which on a retrograde sloping bed leads to a feedback loop of increased
melt, calving and increased in ice flow across the grounding line resulting in rapid further retreat (i.e., MISI). Basal
crevasses, possibly enhanced by basal melting, connect with surface crevasses, leading to enhanced iceberg calving.
(c) Melting and calving has advanced to the stage that an ice shelf no longer exists. An unstable ice cliff face of >90 m
leads to loss of structural integrity due to stress buildup exceeding yield strength. (d) The unstable cliff‐face can fail
by retrogressive slumping, leading to buoyancy‐driven, full‐thickness calving (i.e., MICI). (e) An ice melange forms in
front of the newly calved, marine‐terminating glacier, possibly providing some cliff stabilizing buttressing. Figure
redrawn from Parizek et al. (2019) and Pollard et al. (2016).
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Pollard et al. (2016), and provides a mechanism to support rapid ice loss and sea level rise observed from a
paleo sea level and perspectives for future ice sheet stability (e.g., Rohling et al., 2013). However, this MICI
mechanism has not (yet) been observed in Antarctic glaciers.
MICI was proposed in an attempt to reconcile historical estimates of sea level with ice sheet model simula-
tions. The study by Edwards et al. (2019) revisited the analysis of DeConto and Pollard (2016). Edwards
et al. (2019) attempted to take into account model errors and the uncertainties in reconstructions of past
changes in sea level. They were able to reconcile the model with proxy‐based estimates of the Antarctic con-
tribution to global sea level during the Pliocene Epoch and the LIG, even when using a version of the mod-
eling framework that did not include MICI. However, the large and undefined nature of the uncertainties in
the reconstructions mean that past changes in sea level have limited utility for constraining the physics of ice
sheet models. It therefore remains a parameterized mechanism that requires further testing and validation
involving field studies, model development, and the comparison of ice sheet models with paleoenvironmen-
tal and paleoclimate records with improved uncertainty constraints.
3.2. Ice Sheet Processes Influencing Ice Flow
Understanding the processes governing the flow of ice and its spatial and temporal variability is crucial for
determining ice discharge from the AIS. Different processes influence ice flow in different regions of the ice
sheet and at different spatial and temporal scales: viscous ice deformation (creep) operates throughout the
ice sheet, dominating in regions where the ice is frozen to the bedrock below and on ice shelves, while basal
processes (including basal sliding over a wet bed and bed deformation) dominate in fast‐flowing areas, such
as ice streams. The relative contributions of basal sliding and bed deformation to ice flow can vary both spa-
tially (Bueler & Brown, 2009) and temporally (Stearns et al., 2008). The following section primarily focuses
on recent developments in ice deformation relations and basal sliding laws. The development of these
numerical descriptions relies on synthesis between observations, laboratory experiments, and numerical
model simulations.
3.2.1. Ice Deformation
Ice deformation is a key process contributing to the large‐scale flow of the AIS. Ice is a polycrystalline mate-
rial that deforms in response to an applied stress. Deformation is dependent on the microstructural proper-
ties of ice crystals and the type of stresses applied (Budd & Jacka, 1989). Under prolonged exposure to the
same type of stress, patterns of preferred crystal orientations (fabric) develop that are related to the nature
of the stress configuration. The ice strain rates and fabric are both a function of the stress configuration, that
is, the relative proportions of simple shear and normal stresses (Figure 6). Other factors, such as temperature
and impurities, also influence flow rates. Laboratory experiments show that the type of stress applied to the
ice mass is important to the overall ice flow regime, with the ice flow response to simple shear stresses
greater than that under compression stresses (Treverrow et al., 2012). Developing and validating flow rela-
tions that incorporate these features of ice deformation and are appropriate for investigating large‐scale
AIS evolution requires a combined observational and modeling approach.
Studies of ice deformation have been widely used to investigate the relationship between applied stresses
and corresponding strain rates. The pioneering laboratory work of Glen (Glen, 1952, 1953, 1955, 1958) and
Nye (1953) on isotropic ice samples led to the development of an empirical power law flow relation for
isotropic ice—the Glen flow relation—which is arguably the most widely used flow relation in numerical
ice sheet models. Subsequent laboratory work (Steinemann, 1954, 1958) highlighted the importance of
anisotropy (i.e., ice having a preferred crystal orientation) and microstructural properties of ice, following
which numerous flow relations describing the steady‐state creep of polar, anisotropic ice have been pro-
posed. Broadly speaking, there are two general categories of flow relations for anisotropic ice: (1) those
that directly account for the anisotropy through parameterization of specific properties of individual ice
crystals driving deformation (e.g., Azuma & Goto‐Azuma, 1996; Gagliardini et al., 2009; Gillet‐Chaulet
et al., 2005; Pettit et al., 2007; Placidi et al., 2010; Thorsteinsson, 2002) and (2) those that use empirical
methods, describing the deformation of anisotropic ice as being consistent with the broad‐scale stress con-
figuration (Breuer et al., 2006; Budd & Jacka, 1989; Budd et al., 2013; Graham et al., 2018; Li et al., 1996;
Treverrow et al., 2012; W. Wang et al., 2012). Flow relations in the former category rely on knowledge of
the ice fabric at the crystal scale, often with an explicit evolution equation for the ice crystal fabric.
However, flow relations of this type may be too computationally costly to be feasible for
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continental‐scale modeling applications over long time scales. As an alternative, most continental‐scale ice
sheet models use the Glen flow relation for isotropic ice, adding an ad hoc enhancement factor to account
for the tertiary creep mode of deformation prevalent in polar ice sheets relative to the empirical relation-
ships based on secondary creep laboratory data (Greve & Blatter, 2009). A recent idealized modeling study
(Graham et al., 2018) showed that the Glen flow relation can distort ice shelf geometry and overestimate
shelf flow compared with a flow relation for anisotropic ice developed from empirical studies (Budd
et al., 2013; Graham et al., 2018), which has implications for its reliability in modeling AIS evolution
and its contribution to sea level change.
Geophysical studies using polarimetric radar (Fujita et al., 2006; Matsuoka et al., 2003, 2012; B. Wang
et al., 2018) and seismic (Diez & Eisen, 2015; Diez et al., 2014, 2015) systems have demonstrated that ice crys-
tal fabric and microstructural properties that influence deformation play an important role in ice sheet and
glacier evolution. For example, Matsuoka et al. (2012) used radar‐derived estimates of ice crystal fabric and
anisotropy to infer the migration history of the WAIS Divide between the Ross and Amundsen Seas, arguing
that the strain history, and hence evolution, in this region has been consistent for the past five to eight thou-
sand years. More recently, B. Wang et al. (2018) used ice crystal fabric and age information at Dome A, East
Antarctica, to infer characteristics about the dynamics that controlled deformation in the past. Their results
suggest that accounting for processes governing ice deformation is important in accurately dating englacial
ice layers in the AIS and that ice deformation processes can provide alternative explanations for the forma-
tion and evolution of internal structures in the ice, with implications for estimates of past accumulation
rates. Other studies have used seismic methods to demonstrate the importance of the ice fabric in influen-
cing the dynamics of fast‐flowing ice streams (Harland et al., 2013; Picotti et al., 2015; Smith, Baird,
et al., 2017) and highlight the importance of accounting for the deformation of anisotropic ice in numerical
simulations to accurately predict ice mass loss into the future.
3.2.2. Basal Motion
Basal motion arises from the processes of till deformation and from ice sliding over a wet bed. While these
processes may occur simultaneously, they are mechanistically distinct, although both are influenced by con-
ditions at the ice base: temperature, bed elevation and roughness, and meltwater pressure, volume, and evo-
lution (section 4.2.3). Following Cuffey and Paterson (2010), we refer to the processes of basal sliding and till
deformation collectively as “basal slip”and the relation used to describe these processes, that is, relating
basal shear stress to basal velocity, water pressure, and basal conditions, as a “friction law.”
A variety of friction laws has been developed to capture the relationship between basal shear stresses and
sliding over a rigid bed. One of the most commonly used formulae (Weertman, 1957) is a power law relation-
ship between basal shear stress and sliding velocities that includes a basal friction coefficient to represent
bed properties. However, there are a number of limitations of the power law. First, laboratory studies show
that the basal shear stresses are also dependent on the effective pressure at the base (Budd et al., 1979), which
the Weertman‐type friction law neglects. Second, basal shear stress magnitudes in power law relations are
often unconstrained, despite observational evidence to indicate an upper limit. Other friction laws address
these issues by explicitly including effective pressure to represent flow over a rigid bed with water filled cav-
ities or an upper bound on basal shear stresses (Gagliardini et al., 2007; Lliboutry, 1968; Schoof, 2005). For
example, Schoof (2005) proposed a regularized Coulomb friction law with cavitation, that is, the effect of
pockets of liquid water between the ice and basal substrate, with an effective limit on the magnitude of
the basal shear stresses, even for increasing basal sliding.
The friction laws described above account for basal slip over a rigid bed and do not include basal motion as a
result of till deformation—this process generally being described in separate friction laws (e.g., Boulton &
Hindmarsh, 1987; Bueler & Brown, 2009; Hindmarsh, 1997). Recently, Zoet and Iverson (2020) used labora-
tory experiments of ice motion over both hard and soft beds to develop a unifying friction law that accounts
for basal motion as a result of sliding over a rigid bed and from till deformation. This friction law describes
the dependence of basal shear stress on water pressure, basal sliding, and a transition velocity—an experi-
mentally determined parameter that captures the ice speed at which till deformation begins. That is, below
the transition velocity, basal motion occurs by sliding; at or above the transition velocity, basal motion
occurs by sliding and bed deformation. There is evidence to suggest that this new friction law agrees well
with observations of glacier surface velocities in Iceland and Greenland (Zoet & Iverson, 2020).
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Being one of the most fundamental components of an ice sheet model, the friction relation can significantly
impact the ability of the model to reproduce observed glacier behavior. Joughin et al. (2019) compared
numerical simulations of Pine Island Glacier from 2002–2017 using a variety of friction laws and showed
that a regularized Coulomb friction law (e.g., Schoof, 2005) produced the best match to observed patterns
in grounding line retreat and glacier acceleration, due to cavitation effects. That is, the effect of cavitation
reduces the sensitivity of the basal friction to speed, redistributing the stresses over a wider area. This result
has implications for the choice of friction laws used in studies assessing future mass balance and sea level
calculations.
Due to difficulties in directly observing the subglacial environment, there remain uncertainties in how to
accurately parameterize friction laws in ice sheet models. This may be addressed using a technique
called the adjoint (or control) method (MacAyeal, 1992b). For example, the adjoint method has been
used to generate spatially and dynamically consistent values of the basal friction coefficients in friction
laws through matching observed and modeled fields—usually the present ice surface flow field—for a
given snapshot in time. The use of the adjoint method in this context has improved numerical model
accuracy in matching observed fields, providing insight into the AIS dynamics and its evolution
(Favier et al., 2014; Gillet‐Chaulet et al., 2016; Morlighem et al., 2010, 2013). Nevertheless, errors and
uncertainties may be introduced into the basal friction coefficient as a result of incomplete model phy-
sics and parameterizations, inaccuracies in fields used to initialize the model, and by way of its stationar-
ity in time.
Technological advances may also provide the capacity to address uncertainties in friction law parameters.
For example, tomographic radar systems (Jezek et al., 2013) that provides off‐nadir swath coverage (i.e., data
coverage beyond the single or series of lines from ice penetrating radar tracks) have been increasingly used to
improve constraints on the cross track variation in bed properties that are useful for calculating basal rough-
ness and water distribution (e.g., Rippin et al., 2011; Siegert et al., 2005). New methods have also been devel-
oped that allow for direct observations of the ice sheet base and subglacial hydrology drainage systems,
through the use of chemical (Chandler et al., 2013) and electronic (Bagshaw et al., 2012) tracers, and the
deployment of untethered sensors beneath the ice sheet utilizing wireless communications (Lishman
et al., 2013).
4. Solid‐Earth Interaction and Ice Sheet Bed Conditions
The solid Earth interacts with the AIS over time scales of hours to millions of years, and from local (i.e., kilo-
meter) to global spatial scales (Figure 9). The long response time scales of solid Earth processes mean that
these processes are currently out of equilibrium, and continue to adjust in the background, on top of the cur-
rent climate forcing. It has long been understood that ice‐solid Earth interactions are two‐way: the solid
Earth influences the ice sheet dynamics and properties through changing large‐scale ocean and atmospheric
circulation, basal geometry, and heat input; and the ice sheet affects the evolution of the solid Earth through
erosion and sedimentation, GIA, and hydrogeology. However, recent research has revealed new spatial com-
plexity of the interactions (e.g., Burton‐Johnson et al., 2017; Martos et al., 2017) and relevance to Antarctica
of mechanisms previously demonstrated elsewhere (e.g., Loose et al., 2018) and, in more than one case, more
rapid interactions than previously thought (e.g., Barletta et al., 2018; Larour et al., 2019). The need to under-
stand these interactions is pressing, especially with regard to the need to produce accurate projections of
future sea levels. The following sections review the various interactions that have improved our understand-
ing of AIS dynamics.
4.1. Template‐Setting Interactions
The characteristics of the ice sheet bed are required to understand past and future AIS dynamics. These char-
acteristics can be thought of as “template setting”interactions. A template‐setting process is one that gener-
ates boundary conditions that are stable over the desired time scale, for example, glacial‐interglacial time
scales. For example, topographic controls include the elevation and morphology of the ice sheet bed
(Fretwell et al., 2013), the geothermal heat flux across the ice‐rock interface (Martos et al., 2017), and how
these change through time due to erosion, sedimentation, and tectonic processes (Paxman et al., 2019).
Accurate models of past ice sheet states demand a good understanding of these processes despite the fact that
they evolve more slowly than other boundary conditions. Several recent studies focused on the
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reconstruction of past periods of change, such as the last interglacial or the Pliocene as indicators for
potential future ice sheet states (Cook et al., 2013; de Boer et al., 2015; Golledge, Levy, et al., 2017;
Patterson et al., 2014). Knowledge of the nature of the ice sheet bed at these times has been inadequate to
constrain such studies within acceptable bounds for them to be used in a predictive capacity. The
following section reviews recent progress in understanding the long‐term evolution of the basal boundary
conditions to the AIS, including tectonically generated topography, geothermal heat flux, and dynamic
topography.
4.1.1. Topography Resulting From Tectonic Processes
Tectonic processes, such as rifting, subduction, and mountain building, are the primary drivers of topogra-
phy within the Antarctic continent and have determined the long‐term evolution of the AIS (Figure 9). Most
fundamentally, significant differences in the structure of the thick cratonic lithosphere and high viscosity
mantle of East Antarctica and the thin lithosphere and low viscosity mantle of West Antarctica have caused
substantial differences in topography (An et al., 2015a; O'Donnell et al., 2017) as well as in the response of
the AIS to external climate forcing. These tectonic differences play a significant role in determining the rates
and importance of feedback processes that affect ice sheet dynamics on medium to long time scales (Pollard
et al., 2017).
Despite the substantial growth in knowledge of present day AIS topography, uncertainties in bedrock eleva-
tion (on the order of 100–1,000 m) can result in large differences in the retreat patterns of glaciers in key
regions, such as the areas drained by the Recovery and Support Force Glaciers in East Antarctica (Gasson
et al., 2015) and Pine Island Glacier in West Antarctica (Sun et al., 2014). Compounding these uncertainties,
past ice sheet reconstructions may be affected by lack of knowledge of past tectonic movements.
Widespread rifting in West Antarctica has been linked to dynamic thinning of the WAIS (Bingham et al., 2012;
Pritchard et al., 2009). The resulting subglacial topographic setting of the WAIS controls the ice sheet's
susceptibility to MISI (Schoof, 2007; section 3.1) and has caused the WAIS, particularly in the Amundsen
Sea sector, to be vulnerable to external climate forcing (Vaughan et al., 2011) (sections 2.1 and 5.2). These
regions of tectonic activity have been preferentially eroded first by river systems and later during continental
glaciation into deep troughs through erosion by ice streams across the continental shelf (Gohl, 2012). The
Ross and Weddell embayments formed due to rift processes associated with Gondwana breakup. The
Weddell Embayment ceased rifting in the Jurassic (Jordan et al., 2017) and has exerted only passive topo-
graphic control on AIS evolution. In contrast, the Ross Embayment experienced episodic rifting due to the
West Antarctic Rift System (WARS) since the late Cretaceous and continues to show active rifting today in
the Terror Rift, which has been active since the Miocene (Fielding, 2018; Fielding et al., 2008; Huerta &
Harry, 2007). Neogene rifting is also suggested for the Amundsen Sea (Bingham et al., 2012; Jordan et al., 2010;
Kalberg & Gohl, 2014).
Figure 9. Earth processes and time scales. Illustrative plot showing typical length and time scale for ice sheet interactive
processes at the base of the ice sheet. Template‐setting (see text) interactions (dashed lines) occur over time scales of
millions of years and direct interactions (solid lines) are short‐term processes that result in a dynamic relationship
between the ice sheet and surface processes of the solid Earth. Under typical conditions, these processes operate on
length and time scales below and left of the line. Note the different scaling of the solid Earth processes relative to the
surface processes. Threshold responses are not represented.
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Present accelerating AIS mass changes are focused in regions of active rifting that are characterized by con-
tinental shelves cross cut by deep troughs (e.g., overdeepened glacial troughs in the Amundsen Sea
Embayment; Morlighem et al., 2020), which can be narrow or broad, with easily eroded sedimentary depos-
its, and elevated geothermal heat flux (Bingham et al., 2012; Buck, 1991). Narrow‐rift processes create deep
valleys, which are modified by fluvial conditions (Sugden & Jamieson, 2018) and are also are repeatedly
exploited by ice streams over multiple glacial cycles. Well‐defined troughs are created that today can route
warm open ocean waters across the continental shelf to the interior of the WAIS (Joughin et al., 2014).
Flanking high ground is also generated, which can further strengthen the topographic focusing (Egholm
et al., 2017). Examples of this include some rapidly changing glacier systems that are currently shown to
be most sensitive to ocean‐atmospheric forcing (e.g., Pine Island Glacier and Ferrigno Glacier; Bingham
et al., 2012; Jordan et al., 2010). In other cases, broader depressions do not show the same degree of topo-
graphic focusing but have high geothermal heat flux and well‐lubricated sedimentary beds (e.g., Siple
Coast and Thwaites Glacier) (Damiani et al., 2014).
Active rifting is also associated with upwelling of warm buoyant mantle and thermal support of the litho-
sphere, which varies over time. The warm rifted mantle of the WARS provides buoyant support to the
Ross Embayment and may also partially support the Transantarctic Mountains (TAM) (Brenn et al., 2017;
Graw et al., 2016) (Figure 10). Recent seismic studies in the Ross Sea suggest a complex seismic velocity
structure in the lithospheric mantle (Accardo et al., 2014; Hansen et al., 2014; O'Donnell et al., 2017), indi-
cating strong thermal heterogeneity. This implies a highly variable topography with different amounts of
thermal uplift and subsidence throughout West Antarctica (see section 4.1.3). In addition, motion will occur
at different rates associated with a variable viscosity in the mantle. Early paleotopography models accounted
for thermal subsidence (Wilson et al., 2012) but did not capture these heterogeneities. Two recent studies
have reconstructed paleotopography since circa 34 Ma, one from an iterative topographic reconstruction
method (Paxman et al., 2019) and another based on subglacial sedimentary processes in an ice sheet model
(Pollard & DeConto, 2019). Although more sophisticated than their predecessors, incorporating new process
and including variable flexural rigidity, these models do not possess variable mantle properties.
Incorporation of thermal support of topography and variable viscosity in the mantle is important to con-
strain past and future changes in subglacial topography associated with changing conditions of the AIS,
and resolving the associated feedbacks into AIS dynamics, and thus future projections of ice sheet evolution
and sea level change (Colleoni et al., 2018; Whitehouse et al., 2019).
East Antarctica is tectonically more stable than West Antarctica but has very substantial topographic varia-
tions (O'Donnell et al., 2017). New geophysical data sets have helped to resolve tectonic influences on topo-
graphy extending back to the Archean in the form of compositionally buoyant Archean lithosphere that
underlies much of East Antarctica (An et al., 2015b; Shapiro & Ritzwoller, 2004). Superimposed on this is
a variety of morphologies associated with younger deformations. These include orogenic belts (e.g., the
Figure 10. The role of tectonic processes on Antarctic topography. Schematic figure showing the effects of rifting,
subduction, and mountain building on the current topographic relief of Antarctica. (WARS: West Antarctic Rift
System; TAM: Transantarctic Mountains; EARS: East Antarctic Rift System).
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Gamburtsev Subglacial Mountains (Ferraccioli et al., 2011; Paxman et al., 2016) and various forms of rift
basin, including narrow rift (e.g., the Lambert Trough; Phillips & Läufer, 2009), wide rifts (e.g., Wilkes
Subglacial Basin; Jordan et al., 2013), and sag basins (e.g., Sabrina Subglacial Basin; Aitken et al., 2016).
Thermochronology studies suggest that most of these features date back to at least the Mesozoic (e.g.,
Lisker et al., 2007), with focused Cenozoic erosion (S. N. Thomson et al., 2013), and therefore have had a pas-
sive influence on AIS development.
4.1.2. Dynamic Topography
Dynamic topography is the uplift or subsidence of the Earth's surface due to heterogeneities in mantle cir-
culation (Figure 11), which cause long‐wavelength topographic changes over million year time scales
(Conrad, 2013; Flament et al., 2013) with typical magnitudes of hundreds of meters. The long time scales
of dynamic topography, and the dependence on global mantle circulation patterns, means that it has a
template‐setting influence on ice sheet dynamics. Although dynamic topography does not possess direct
feedbacks with ice sheet evolution (in contrast with GIA, section 4.2.2), a knowledge of past dynamic topo-
graphy is nevertheless important for accurate reconstruction of the past AIS in two main ways.
The first is the global effect of dynamic topography on estimates of eustatic sea level (Dutton, Carlson,
et al., 2015; Rowley et al., 2013). In this case, corrections to eustatic sea level beyond the last interglacial
are significantly affected by dynamic topography (Dutton, Carlson, et al., 2015). The geoid (and hence regio-
nal sea level) is affected by dynamic topography on the order of tens of meters. Positive (negative) dynamic
topography is associated with geoid highs (lows), although the underlying low‐density mantle responsible
may generate the opposite relation (Hager et al., 1985). With current levels of model uncertainty, a reliable
correction cannot be made for past dynamic topography. Models can, however, provide indications of which
areas may be most and least affected (Rowley et al., 2013) and so aid in avoiding misinterpretations.
Comparison with ice‐volume reconstructions (e.g., based on deep‐sea benthic oxygen isotope ratios, which
have their own uncertainties but are not affected by dynamic topography) may help with removing
dynamic‐topography influences from sea level benchmark data that are affected by crustal movements.
The second relates to dynamic topography in Antarctica itself. Raising or lowering relative topography
impacts directly on the susceptibility of the AIS to marine instabilities (Austermann et al., 2015) and also
to climatic influences. Global models suggest that the current topography of Antarctica may support as
much as ±1 km of dynamic topography. Features in common for the present day include neutral to positive
dynamic topography of the Indo‐African sector of East Antarctica, neutral to negative dynamic topography
in Wilkes Land, central East Antarctica, and the Weddell Sea, and high dynamic topography along the
Pacific Margin (Conrad & Husson, 2009; Flament et al., 2013; Spasojevic & Gurnis, 2012;
Figure 11. Dynamic topography. Schematic figure (not to scale) showing how heterogeneous deep mantle circulation
alters the elevation of the solid Earth surface (gray); regions of hot mantle upwelling cause high dynamic
topography and cold‐mantle downwelling causes a low dynamic topography. CMB; core‐mantle boundary. Figure drawn
based on Kiefer and Kellogg (1998).
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Steinberger, 2007). These variations suggest that some of the topographic‐induced sensitivities of the AIS,
including MISI affecting the Wilkes Subglacial Basin, and the current high elevation of Dronning Maud
Land, may have changed through time due to dynamic topography. However, dynamic topography within
Antarctica has been explicitly addressed by only a few studies in the Scotia Sea and Ross Sea region
(Nerlich et al., 2013; Spasojevic et al., 2010) and several studies on mantle‐plume derived support of the
Marie Byrd Dome (LeMasurier & Landis, 1996; Spasojevic et al., 2010; Wobbe et al., 2014).
The potential influence of dynamic topography on the AIS has been examined for the warm mid‐Pliocene
period (Austermann et al., 2015). The dynamic topography model used in this study predicted that topogra-
phy was 100 to 200 m lower at 3 Ma, and the modeling results suggest that this lowering causes a destabiliz-
ing effect on the AIS. In the ice sheet model, an additional 200 to 560 km of grounding line retreat into the
Wilkes Basin was found, relative to present‐day topography.
The influence of past dynamic topography on the AIS depends on the direction and magnitude of dynamic
topography change and also on the susceptibility of the ice sheet to relatively small changes in bed elevation
(Austermann et al., 2015). Subglacial regions with deep narrow troughs may be little affected by dynamic
topography, whereas regions with broader shallow depressions may be more affected. In the latter, slight
overdeepening may allow greater ocean access to the ice sheet, as seen in the Wilkes Subglacial Basin mod-
eling study (Austermann et al., 2015). Conversely, a similar magnitude change in the opposite direction may
mean that once‐vulnerable catchments are then better protected from MISI than before.
Dynamic topography is now widely recognized as a substantial driver of topographic change, with substan-
tial impact on AIS sensitivity, and in our ability to understand past ocean and ice sheet changes. Despite this,
a better understanding of how Antarctic topography has evolved under the influence of dynamic topography
remains a significant challenge. Model uncertainties remain high, and the data needed to help reduce these
uncertainties are sparse, especially within Antarctica. In the future, sensitivity studies will be important for
constraining the timing and spatial scale of dynamic topography. In particular, quantitative models have
substantial diversity at regional scales, which means that catchment‐scale understanding remains very lim-
ited. An important goal is to define the potential sensitivity of the AIS to dynamic topography.
4.1.3. Geothermal Heat Flux
The range and spatial distribution of geothermal heat flux at the base of the AIS influences ice dynamics
through changes in ice deformation from changing ice temperature, and via the production of subglacial
meltwater, both of which affect the ice base, and thus flow of grounded ice (section 3.2.1). Except in rare cir-
cumstances, such as subglacial volcanism (Behrendt, 2013; Blankenship et al., 1993; van Wyk de Vries
et al., 2017), geothermal heat flux is a stable feature of the geology on glacial to interglacial cycles, and
changes little over millions of years. It is, however, affected by erosion or sedimentation and the topography,
so paleoreconstructions may need to consider those influences. Despite being a stable condition over glacial
to interglacial cycles, its interactions with the ice sheet may change rapidly as the thermodynamic setting of
the interface changes (Loose et al., 2018). Under certain conditions, the thermodynamics of ice are very sen-
sitive to heat flux, and a change in the overall thermodynamic system may lead to a changing association
with the heat flux, despite the heat flux itself has remained stable. Therefore, a more accurate and precise
knowledge of geothermal heat flux is required for greater confidence in understanding past and present
basal conditions and the pattern of AIS flow.
There are no direct measurements of subglacial heat flux in East Antarctica and only a few from West
Antarctica (Begeman et al., 2017; Fisher et al., 2015; Siple Coast). Measurements of thermal gradients within
the ice support studies of the AIS, particularly in search for basal conditions that enable preservation of the
oldest ice for climatic records modeled for ice core drill sites (Price et al., 2002, South Pole; Salamatin
et al., 1998, Vostok). Temperature gradients within the ice have been used to suggest the contribution of heat
from beneath but can only be trusted if the thermal conditions at the base are well known. In the case of off-
shore temperature gradients, these can be collected in a simpler and quicker manner than over the conti-
nent. These values can be useful to understand tectonic settings and lithospheric response of the
continent if the measurements are acquired in close proximity to the coastal regions (e.g., measurements
in the Amundsen Sea Embayment; Dziadek et al., 2017, 2019).
New drilling programs to access the base of the ice sheet (Goodge & Severinghaus, 2016) will increase the
coverage of these data. However, point measurements of geothermal heat flux remain highly localized,
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with high spatial variability due to the underlying geology and may not be representative of regional
conditions.
Regional geothermal heat flux studies provide insight in the local thermal properties at the base of the AIS.
Lateral variability of crustal heat production is used to refine models that assume constant values (Antarctic
Peninsula; Burton‐Johnson et al., 2017). Studies of Antarctic heat production include those from central East
Antarctica (Goodge, 2018) and around Prydz Bay (Carson et al., 2014). Heat production data have been com-
piled in a recent study by Gard et al. (2019). Estimates of geothermal heat flux have also been derived from
ice mass balance calculations in ice streams, heat flux derived by ice structure from radar data in Dome C
(Parrenin et al., 2017), and analyses of hydrological conditions beneath Thwaites Glacier (Schroeder
et al., 2014). These regional maps provide insight into key areas of change, but upscaling of these methods
is difficult due to the highly localized nature of the information and limited access to continental bedrock.
Previous work has sought to improve continental‐scale estimates of geothermal heat flux. Antarctic heat
flow maps derived from temperature gradients based on seismic wave speed (An et al., 2015b; Shapiro &
Ritzwoller, 2004), magnetic equivalent dipole models (Maule et al., 2005), and ice sheet models (Pollard &
Deconto, 2012) showed continent‐wide and/or regional discrepancies. Recently, however, Martos et al. (2017)
derived a geothermal heat flux model from spectral analysis of airborne magnetic data to determine the
depth to the Curie temperature isotherm. In particular, spectral methods were used to determine the mag-
netic structure of the lithosphere (in particular, depth of the Curie isotherm), from which the geothermal
heat flux was estimated. This geothermal heat flux map (Figure 12) provides higher spatial variability com-
pared with previous estimates based on other techniques (e.g., seismic, An et al., 2015a) and characterizes
uncertainty in geothermal heat flux estimates. Higher heat flux is found in West Antarctica, where the
majority of the volcanoes lie, while East Antarctica is mostly characterized by low heat flux values, in agree-
ment with the cratonic nature of this part of the continent.
The large difference in geothermal heat flux estimates between East and West Antarctica are related to the
tectonic evolution of these sectors. Subglacial volcanism can lead to rapid changes in the local geothermal
heat flux, and the corresponding feedbacks with the ice sheet may lead to a more direct interaction
(Behrendt, 2013; Blankenship et al., 1993; Loose et al., 2018; van Wyk de Vries et al., 2017). However, anom-
alously high geothermal heat flux has recently been reported for the South Pole region, where high heat pro-
ducing Precambrian basement rocks and hydrothermal circulation along a major fault system have been
reported (Jordan, Martin, et al., 2018). Alongside geophysical modeling, efforts have been made to estimate
basal temperatures and basal melt rates by using ice sheet velocity models (Pattyn, 2010; Van Liefferinge &
Pattyn, 2013), which suggest that the interior of East Antarctica represents a frozen bed while coastal East
Antarctica and regions in West Antarctica are characterized by basal melting reaching up to 30 mm yr
−1
,
generally coinciding with higher geothermal heat flux locations (Figure 12).
Despite this recent progress in proxy‐based methods and their application, resolving the geothermal heat
flux still remains challenging. These methods complement each other, as each represents a different physical
phenomenon, and the methods have different sensitivities in depth‐sensitivity and spatial resolution. Each
method also depends on a set of assumptions, and the impact of these on the modeled outputs is difficult to
assess using the resulting geothermal heat flux maps. Not all studies and local estimates have provided
uncertainties, which limits their usefulness. Notwithstanding potential errors, the differences in results
may reflect different but complimentary facets of the complex system required understand Antarctica's heat
flux. Higher‐quality and more consistent geothermal heat flux models are needed to enable a more systema-
tic knowledge of the geothermal heat flux on the generation of subglacial meltwater, the dynamics of sub-
glacial hydrological systems, and their role in AIS dynamics. In addition, advanced sensitivity studies of
geothermal heat flux with respect to basal temperatures and melt rates are needed to understand the influ-
ence of geothermal heat flux on ice sheet behavior, which depends strongly on the thermodynamic setting.
4.2. Direct Interactions
4.2.1. Erosion and Sedimentation
The erosion of rocks and the subsequent redistribution of sediments at the base of the AIS influences the sen-
sitivity of the ice sheet to marine forcing, through changes in ice dynamics and sedimentary processes
beneath ice shelves at the coastal margin, and on the continental shelf, over tens of millions of years down
to subdecadal time scales (e.g., Bart et al., 2016). Broadly, we may separate erosion and sedimentation into
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longer‐term systems that reconfigure the landscape over millions of years, setting the template for individual
glacial cycles, and shorter‐term systems that interact more directly with glacial and oceanic processes.
In general, local sedimentation shallows the bathymetry, while erosion deepens the bathymetry. Both pro-
cesses may generate landforms that act to modify ocean circulation, and both may impinge on the ice shelf.
The regional response to mass loading changes is dictated by flexural‐isostatic processes, and it is common
that focused erosion or deposition will generate a regional uplift or subsidence, respectively. In certain cir-
cumstances, the changing morphology of the seafloor is important in dictating access of CDW to ice shelf
cavities, thus promoting or inhibiting MISI (Bart et al., 2016; Millan et al., 2017; Nitsche et al., 2017).
Similarly, ridges and troughs formed through these processes dictate the likelihood of ice shelf grounding,
and have a particular influence on calving (Arndt et al., 2018).
Relict eroded landscapes are evident in maps of subglacial topography throughout Antarctica (Fretwell
et al., 2013; Morlighem et al., 2020). The degree to which this erosion is associated with fluvial or glacial
action remains unclear (Paxman et al., 2019; Pollard & DeConto, 2019; Sugden & Jamieson, 2018).
Geomorphological features on both the continental margin and buried under the ice can potentially be used
to constrain past ice sheet models (Jamieson et al., 2014). Recent studies both onshore and offshore have
tackled the mapping of these eroded landscapes, and, in some cases, their relationships with offshore deposi-
tional systems (Aitken et al., 2016; Gohl et al., 2013; Lindeque et al., 2013; Paxman et al., 2016; Young
et al., 2011). In general, the landscapes that exist have the potential to be long‐lived or more transient; how-
ever, limited exposures of these features (Sugden & Denton, 2004) prevent precise dating to establish when
they were formed. Over the long term, sediment flux to the ocean has modified the topography of the con-
tinental shelf and shelf break. In particular, the continental shelf has been extended by up to 220 km over the
past 34 Ma, with an overall increase in area of 7% (Hochmuth & Gohl, 2019). In the rest of this section we will
focus on the more short‐lived processes that interact most closely with the AIS.
Sedimentation during phases of grounding line retreat can form stabilizing grounding‐zone wedges, which
are subaqueous moraines at the terminus of ice streams (Alley et al., 1989; Batchelor & Dowdeswell, 2015).
These form at the grounding line and can lead to a temporary pause in retreat (e.g., Alley et al., 2007;
Jamieson et al., 2012; Ó Cofaigh et al., 2008) and may provide stable pinning points (Christianson
Figure 12. The upper map shows a model of the basal melt rate (Pattyn, 2010) with boundary conditions set by averaging
three geothermal heat flux data sets from Fox Maule et al. (2005) and Shapiro and Ritzwoller (2004). The lower map
shows the geothermal heat flux distribution of Antarctica modified from Martos et al. (2017). Blue triangles show the
location of ice cores where rare vertical temperature profiles of the ice sheet exist.
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et al., 2016; Klages et al., 2017). Deposition of grounding zone wedges influence ice stream systems over time
scales from a few hundred to a few thousand years (e.g., Bart et al., 2017; Graham et al., 2010; Jakobsson
et al., 2012) and locally significant sediment erosion and redeposition can occur on time scales as short as
years (Smith et al., 2007) to decades (Smith et al., 2012).
Early on, landward extent of bathymetric cross‐shelf troughs was unknown (Hughes, 1977), but further
research since then has shown that cross‐shelf troughs, which can exceed 1,000 m water depth but are gen-
erally between 200 and 500 m deep (Patton et al., 2016), are indeed paleo‐ice stream troughs extending
beneath the modern AIS. These cross‐shelf bathymetric troughs have been eroded repeatedly by the advance
and retreat of paleo‐ice streams and are found on the West Antarctic shelves in the Ross, Amundsen,
Bellingshausen, and Weddell Seas (e.g., Belgica Trough, Ó Cofaigh et al., 2005; Ferrigno Trough, Bingham
et al., 2012; Filchner Trough, Larter et al., 2012; Rosier et al., 2018; Thiel Trough, Fretwell et al., 2013;
Pine Island‐Thwaites paleo‐ice stream Trough, Graham et al., 2010; Larter et al., 2014; Lowe &
Anderson, 2002; and Thwaites, Holt et al., 2006, subglacial troughs). Overdeepened troughs that we know
about from East Antarctica are less well defined and include along the Totten (Nitsche et al., 2017;
Rintoul et al., 2016) as well as the Academy and Foundation ice streams from East Antarctica (Fretwell
et al., 2013).
Cross‐shelf troughs may increase the vulnerability of the ice sheet to melting at the grounding line by allow-
ing CDW to penetrate directly to the WAIS margin (Bingham et al., 2012; Heywood et al., 2016; Jenkins
et al., 2016; Nitsche et al., 2007; Turner et al., 2017). Historic erosion rates within these troughs may be as
much as 1 m yr
−1
where the bed is weak (Smith et al., 2012). Repeated selective erosion over glacial to inter-
glacial cycles within these troughs may total kilometers in magnitude since the Eocene (Wilson et al., 2012).
An important factor is the width of the basin in which erosion is focused.
Narrow basins exert strong topographic focusing on ice flow, which efficiently focuses erosion within the
trough (Egholm et al., 2017; Jamieson et al., 2005; Maritati et al., 2016; S. N. Thomson et al., 2013). Over time,
selective erosion is associated with an isostatic‐flexural uplift that will be accommodated over a broader
region. This character leads to increasingly focused ice sheet flow through time as the elevated flanks
become cold based, but the lows become more susceptible to either MISI, if marine terminating
(Golledge, Levy, et al., 2017; S. N. Thomson et al., 2013), or to the development of connected subglacial
hydrological systems (section 4.2.4).
In contrast, relatively broad basins, in particular the Wilkes Basin and Ross Embayment, may allow
retreat to occur along broad fronts, with more poorly focused ice streams, resulting in erosion across
broad regions, and relatively subdued trough development (Aitken et al., 2016; Cook et al., 2013)
and/or the development of regional unconformities (Bart, 2003; Bart et al., 2000; Bart & de
Santis, 2012; Brancolini et al., 1995, 1997; De Santis et al., 1999). In such cases, erosion over time may
act to enhance future vulnerability to change as the eroded region forms a broad subsided region increas-
ingly susceptible to MISI processes. Concurrently, the deposition of sediments in a prograding shelf edge
will extend the shelf and increase the distance to the source of CDW and may contribute to the shallow-
ing of the shelf break, inhibiting CDW incursion. With respect to AIS stability, these regional negative
feedbacks may counteract the increased ocean access provided by the deeper bathymetry. With a degree
of selective erosion, and the associated topographic feedbacks, a broad erosional system may evolve into a
more selective system, with uplifted regions developing adjacent to troughs (Egholm et al., 2017; Kessler
et al., 2008).
4.2.2. GIA
GIA is the solid Earth response to past ice‐ocean surface loading changes that resulted in deformation of the
solid Earth and geoid and in changes to the location of Earth's rotation pole and spin rate (Mitrovica
et al., 2009). Due to the viscous properties of the mantle, changes in surface loading at a given point in time
continue to produce effects for thousands of years (Figure 13). For example, the Earth is continuing to
respond to the rearrangement of surface mass loading following the LGM (27–19 ka); this response is mea-
surable and important to consider when interpreting data that give insight into past and present geophysical
processes (Peltier, 2004). Past and present‐day observations of change in geophysical parameters can also
help to inform our understanding of Earth's rheology and ice‐ocean load history (Argus, 1996; Barletta
et al., 2018; Chen et al., 2013; Milne et al., 2001; Mitrovica et al., 2006).
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Accurate sea level projections and measurements of present‐day sea level change cannot be obtained with-
out knowledge of GIA. However, Antarctic GIA is poorly constrained by data. Limitations in GIA models for
Antarctica stem from large uncertainties in AIS evolution during past glacial to interglacial cycles
(section 5.2) and the response of the solid Earth determined by the structure of the lithosphere and mantle
(section 4.1). Advances have been made in understanding the interaction between the solid Earth response
that results in grounding line migration and near‐field sea level change; improved seismic data coverage and
more Global Positioning System (GPS) and Global Navigation Satellite System (GNSS)‐derived surface defor-
mation measurements; and the inclusion of three‐dimensional (3‐D) Antarctic rheology, which determines
the deformation of the bedrock topography beneath the ice sheet and varies according to the viscoelastic
properties of the Earth (Gomez et al., 2018), into GIA models.
Study of Antarctic GIA has intensified over the past two decades as a result of the launch of the Gravity
Recovery and Climate Experiment (GRACE) and GRACE Follow‐on (GRACE‐FO) satellite missions that
have a primary aim of measuring ice sheet contributions to sea level change (Tapley et al., 2004).
However, the accuracy of GRACE‐derived estimates of AIS mass change rely on the accuracy of background
models of other mass‐change processes, with GIA the dominant “contaminating”signal, which is of a simi-
lar magnitude to the estimated ice mass change (King et al., 2012; Velicogna & Wahr, 2006). More recently,
the potential for GIA processes to alter ice sheet dynamics has been suggested, leading to the development of
a new class of models that couple ice sheet and GIA processes (Adhikari et al., 2014; De Boer et al., 2014;
Gomez, Mitrovia, Huybers, et al., 2010, Gomez et al., 2012, 2013, 2015, 2018; Konrad et al., 2015, 2016).
Incorporating GIA processes into ice sheet and sea level models drives a significant advance in the under-
standing of AIS dynamics. GIA feedbacks on ice sheet evolution were first investigated by Gomez,
Mitrovia, Huybers, et al. (2010), who coupled a canonical MISI model to a model of the sea level equation
(Farrell & Clark, 1976) that considers both the solid Earth response to ice mass change and the consequent
changes to the geoid and hence regional sea level. Ice mass loss associated with ice sheet retreat leads to
regional bedrock uplift and a drop in the sea surface gravitational equipotential, producing a combined effect
of sea level lowering at the ice sheet grounding line that impacts ice flow (Figure 13). Gomez, Mitrovia,
Figure 13. Local glacial isostatic adjustment (GIA) schematic. Following ice loss, the loss of gravitational attraction
immediately draws down the local sea surface, and over slower response time scales the solid Earth uplifts elastically
in the vicinity of the region of ice loss, leading to a sea level fall near the ice sheet grounding line. On longer time
scales, the Earth viscously deforms toward isostatic equilibrium, leading to further uplift and sea level fall under the
region of ice loss (upward pointing yellow arrows) and a subsidence of peripheral bulges in surrounding regions
(downward pointing yellow arrows). If the grounding line were to retreat beyond the bedrock high it currently rests on, it
would be in an unstable configuration due to MISI processes. The grounding line may then be subject to irreversible
retreat. The aforementioned sea level fall acts against the MISI, acting to slow or stop grounding line retreat. The
importance of this sea level feedback depends on the shape and depth of the bed beneath the ice, the nature of the climate
(oceanic and atmospheric) forcing on the ice sheet, and the Earth viscosity structure. The viscosity structure
determines the timing and spatial pattern of deformation in response to ice loss and varies by several orders of magnitude
in three dimensions across the Antarctic continent.
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Huybers, et al. (2010) and Gomez et al. (2012) adopted a simplified ice sheet stability model and in turn a
flowline model of ice dynamics to show that this sea level feedback could slow or even stabilize ice sheet
grounding line retreat on retrograde bed slopes, a configuration regarded as fundamentally unstable (e.g.,
Schoof, 2007; Weertman, 1972; section 3.1.2). Building on this, realistic 3‐D ice sheet models have been
coupled with GIA models that consider 1‐D Earth rheologies to explore the potential effect on ice retreat
(De Boer et al., 2014; Gomez et al., 2015, 2013; Konrad et al., 2014). These coupled models demonstrate that
GIA effects are important in modeling past and future ice sheet changes, notably acting to slow or limit
future and Pliocene ice sheet contribution to sea level, with strong sensitivity to the Earth model adopted
in the GIA models (Gomez et al., 2015; Pollard et al., 2017).
Coupling between ice sheet models and solid Earth models with a gravitationally consistent description of
sea level is an active area of research. Challenges lie in correctly translating between the different model grid
types. For example, ice sheet models use regional Cartesian grids, and solid Earth models operate on global
Gauss‐Legendre grids. Many solid Earth models operate in the Laplace domain (see glossary). However,
some now use explicit time stepping to evolve the solid Earth dynamics and are therefore less computation-
ally expensive (e.g., Konrad et al., 2014). GIA changes have been incorporated into ice sheet models to
resolve grounding line dynamics so that the ice sheet and sea level varies consistently in space and time
(De Boer et al., 2014; Gomez et al., 2018, 2015, 2013; Konrad et al., 2015). These studies highlight the negative
feedback associated with a more realistic viscoelastic deformation of the solid Earth and stabilization of MISI
in West Antarctica with a coupled ice sheet‐sea‐level model. de Boer et al. (2017) provide an overview of the
recent developments in coupled ice sheet‐sea‐level modeling and emphasize how synthesis between two dis-
tinct disciplines—GIA and ice sheet dynamics—can greatly enhance our understanding of the interactions
between sea level change and ice sheet dynamics.
Modeling GIA processes requires knowledge of the temporal evolution of past ice‐ocean load changes and
Earth's rheological properties. For Antarctica, both are sparsely observed and are highly uncertain; as a con-
sequence, models of Antarctic GIA are also highly uncertain with large intermodel differences. In particular,
the post‐LGM ice loading history for Antarctica is sparsely observed in space and time (Bentley et al., 2014;
Whitehouse, Bentley, & Le Brocq, 2012), as demonstrated by the wide range (~5–23 m) of recent estimates
of Antarctica's contribution to post‐LGM sea levels (Argus et al., 2014; Lambeck et al., 2014; Pollard
et al., 2016) (see section 5.2). Nonetheless, substantial advances have been made in modeling GIA processes
in recent years, and these have contributed to improved understanding of ice sheet mass balance (Shepherd
et al., 2012; 2019).
A range of observations is useful in improving and/or validating GIA models, including proxy observations
of past ice extent and ice retreat history, proxy, and tide gauge measurements of local and far‐field post‐LGM
relative sea level (RSL), ancient and geodetic observations of Earth's rotation pole, and geodetic observations
of solid Earth deformation and gravity changes (Mitrovica et al., 2009; Tamisiea et al., 2014). These may be
separated into local and global‐scale observations: relevant to Antarctica are paleo RSL data from less than
20 sites across Antarctica (Whitehouse, Bentley, Milne, et al., 2012), spanning the last 12–15 ka, and
GPS‐derived surface deformation measurements at approximately 80 sites with varying time series quality
and length (Martín‐Español, Zammit‐Mangion, et al., 2016). Sediment records of RSL for the period before
15 ka to the LGM are lacking but can be preserved within isolated basins, where lake sediment cores record
the transition from marine outlets or basins during the glaciation to freshwater lakes following deglaciation,
as isostatic rebound outpaced eustatic sea level rise (Hodgson et al., 2016; Roberts et al., 2011).
GPS and GNSS deployments on bedrock outcrops in Antarctica provide estimates of bedrock uplift rates that
are used to validate GIA models and understand missing processes through identification of differences in
the magnitudes of the predicted uplifts and their spatial patterns (Bevis et al., 2009; Groh et al., 2012;
Thomas et al., 2011). Ice sheet reconstructions have been improved by new evidence from ice sheet history
proxy data that the LGM volume of the WAIS was smaller than previous estimates (Whitehouse, Bentley, &
Le Brocq, 2012). Results from dynamic ice sheet modeling driven by climatic changes also support a smaller
LGM AIS volume (e.g., Pollard et al., 2016). These new AIS history proxy data are not uniformly accepted as
reliable constraints on past ice history. Consideration of AIS volume required to explain LGM sea levels that
account for global ice sheet contributions indirectly suggests a larger (23–30 m) contribution from
Antarctica, based on a global sea level budget approach (Lambeck et al., 2014). Nonetheless, based on these
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ice proximal data, a new generation of GIA models has been developed (Ivins et al., 2013; Whitehouse,
Bentley, Milne, et al., 2012) that predict present‐day GIA uplift velocities substantially closer to observed
GPS/GNSS velocities than the previous generation. In combination with GRACE measurements, these mod-
els yield a systematically smaller estimate of AIS mass loss during the last deglaciation (Argus et al., 2014;
Ivins et al., 2013; King et al., 2012) compared to Lambeck et al. (2014). However, these 1‐D GIA models
are unrealistic based on seismic tomographic constraints on Earth rheological structure, which show sub-
stantial rheological variation, notably between East and West Antarctica but also within the regions
(Kaufmann et al., 2005).
Seismology is unable to determine the absolute mantle viscosity. Since 2014, geodetic studies of surface dis-
placement in Antarctica have provided constraints on this missing information. Thomas et al. (2011) demon-
strated nonlinear deformation of the northern Antarctic Peninsula associated with ice mass loss that was
initiated by the breakup of the Larsen B Ice Shelf. This nonlinearity could be explained by a model of viscoe-
lastic deformation, with local upper mantle viscosities of <3 × 10
18
Pa s (Nield et al., 2014). This is around 2
orders of magnitude lower than conventionally adopted viscosities in global GIA models. Meanwhile,
500 km further south, upper mantle viscosities were estimated to be at least 1–2 orders of magnitude higher
(Zhao et al., 2017). Higher viscosities have also been suggested in the southwest Weddell Sea region
(Wolstencroft et al., 2015) and the Siple Coast (Nield et al., 2016). Considering the spatial pattern of change
and high bedrock uplift rates measured by GPS/GNSS, the Amundsen Sea Coast region has also been sug-
gested to be underlain by low‐viscosity mantle (Groh et al., 2012), and this has been confirmed using a more
dense network of 3‐D displacement measurements (Barletta et al. 2018). In contrast to West Antarctica, no
geodetic constraints are available for East Antarctica although there is a consensus supported by seismic stu-
dies and geological history of the region, which suggests that, aside from some localized regions where the
viscosity could be as low as 10
18
Pa s, East Antarctica is underlain by high viscosity mantle and is responding
slowly to LGM‐time scale changes in ice loading.
Global and regional seismic studies indicate substantial 3‐D variations in Antarctic rheology, both between
and within East and West Antarctica. Regional seismic tomography studies have improved our knowledge of
upper mantle structure and spatial variation (An et al., 2015b; Chaput et al., 2014; Heeszel et al., 2016;
O'Donnell et al., 2017). These seismic data indicate that effective upper mantle viscosity may vary by 3–4
orders of magnitude across Antarctica, with values as low as 10
18
Pa s in parts of West Antarctica
(O'Donnell et al., 2017). Such low viscosities, if realistic, suggest relaxation times in the upper mantle on
the order of decades compared to the conventional understanding of thousands of years, as well as more
smaller spatial scale patterns of deformation. However, LGM‐scale ice loading changes will have perturbed
the lower mantle, producing deformation patterns on millennial time scales and large spatial scales in addi-
tion to the deformation occurring in the upper mantle over smaller spatial scales and shorter time scales.
Including lateral variation in GIA models produces distinctly different spatiotemporal patterns of deforma-
tion and associated gravity field change (Kaufmann et al., 2005; King, Whitehouse, & van der Wal, 2016; van
der Wal et al., 2015). For example, regions of lower than average upper mantle viscosity and a thinner litho-
sphere such as the in the Amundsen Sea Embayment (Barletta et al., 2018) result in faster viscous deforma-
tion toward isostatic equilibrium that is more localized to the region of ice loading changes. Various
approaches have been used to convert seismic tomography models into mantle viscosity, but uncertainties
in these approaches result in a wide range of model predictions (e.g., van der Wal et al., 2015), and indepen-
dent data are required to validate models. One characteristic of 3‐D models is that the spatial gradients of
deformation are larger, and peak uplift and subsidence are more smoothed than in 1‐D models (Nield
et al., 2018).
The presence of low upper‐mantle viscosity in parts of West Antarctica, including the Antarctic Peninsula,
implies a strong dependence of present‐day GIA on late Holocene or centennial‐scale ice loading changes;
this is particularly problematic for modeling studies because almost no Antarctic paleoproxy data exist for
this period. Certain regions may have experienced substantial changes in ice cover during the Holocene with
retreat beyond modern grounding lines followed by readvance (Bradley et al., 2015; Kingslake et al., 2018;
Siegert et al., 2013). For example, field data and modeling have been combined to suggest that ice streams
in the Siple Coast region and in the Weddell Sea Embayment have readvanced in the last few thousand years
and that this readvance may have been due to a GIA feedback (Kingslake et al., 2018; Siegert et al., 2019).
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However, these simulations used an upper mantle viscosity of 5 × 10
20
Pa s, which is unrealistically high for
West Antarctica—by perhaps as much as two orders of magnitude. Kingslake et al. (2018) find that even
reducing the viscosity to 1 × 10
20
Pa s prevents this response because, “the bed responds too quickly to ice
unloading.”The most widely adopted global and regional GIA models assume no ice loading change over
the last 2 to 4 ka; therefore, predictions of present‐day deformation and GIA in West Antarctica, including
the Antarctic Peninsula, remain highly uncertain. This has been highlighted by inverse solutions for GIA
based on satellite altimetry, and GRACE and spatially continuous GPS/GNSS measurements (Gunter
et al., 2014; Martín‐Español, Zammit‐Mangion, et al., 2016; Riva et al., 2009; Sasgen et al., 2017). These mod-
els show many of the features of the forward models, although without a clear convergence between the
inverse solutions or with a particular forward model (Martín‐Español, Zammit‐Mangion, et al., 2016).
However, inverse models consistently demonstrate rapid uplift, for instance, in the Amundsen Sea region
that is compatible with the GPS estimates, and have been demonstrated through modeling to originate
purely from unloading due to ice mass loss since 1900 CE (Barletta et al., 2018).
The negative feedback between ice dynamics and bedrock uplift has the potential to stabilize dynamic
retreat of vulnerable WAIS glaciers such as the Thwaites and Pine Island Glaciers over a multicentennial
time scales (Gomez et al., 2015; Larour et al., 2019). A high spatial resolution (1 km) modeling approach
for the ice sheet grounding line migration in response to rapid elastic uplift demonstrated a significant
increase in elastic uplift rates (10 cm year
−1
) beyond 2300 CE (Larour et al., 2019).
The implications of 3‐D variations in Earth structure on ice dynamics are just beginning to be explored.
Given the substantial lateral structural variability beneath the Antarctic, 1‐D Earth models are unable to
accurately capture the response of the Earth to, and feedbacks on, changes in ice loading across the whole
continent. In regions of lower viscosity and thinned lithosphere, uplift in response to ice unloading will take
place faster and be more localized to the ice sheet grounding line, where it can more effectively slow ice sheet
retreat. Building on earlier work, Gomez et al. (2018) developed a coupled model incorporating 3‐D varia-
tions in Earth structure and applied it to the last deglaciation, showing substantial local differences in pre-
dicted ice cover, sea level, and Earth deformation between simulations adopting a 3‐D Earth model and
those using average 1‐D Earth models. They also found that it will be important to consider the impact of
3‐D Earth viscosity structure in future ice sheet projections, particularly given the low viscosities in regions
of expected future ice sheet retreat. Hay et al. (2017) adopted a sea level model with 3‐D Earth structure to
show that viscous Earth deformation, which is typically neglected in sea level projections on century time
scales, will increase sea level fall in regions of ice loss by up to 50% in 100 years, thereby contributing to
the negative feedbacks acting on grounding line dynamics.
4.2.3. Basal Hydrology and Fluid‐Interchange Processes
The presence or absence of liquid water at the ice sheet bed, as well as its distribution, influences ice sheet
flow and discharge and thus affects mass balance (e.g., Ashmore & Bingham, 2014; Bell, 2008, and references
therein; Siegert et al., 2017). Understanding the nature of the ice sheet bed is essential to understanding AIS
dynamics and in particular to accurately model flow over a variably rough and partially lubricated substrate
(section 3.2). Factors influencing the subglacial hydrology include the ice overburden pressure, subglacial
topography, and production and removal of meltwater. New research shows that hydrological systems con-
sisting of active lakes and channel networks are a near‐ubiquitous feature of Antarctica, and a significant
driver of AIS dynamics (Ashmore & Bingham, 2014; Bell, 2008; Fricker et al., 2016; Kirkham et al., 2019;
Siegert et al., 2016; Smith, Gourmelen, et al., 2017; Stearns et al., 2008).
Subglacial fluid transport networks can form continuous or semicontinuous flow paths from the interior of
the ice sheet to the grounding line (Ashmore & Bingham, 2014). These can provide a direct connection
between the inland ice sheet and sub‐ice shelf cavities (Figure 14). The discharge of freshwater to ice shelf
cavities can affect ice shelf dynamics by increasing basal sub‐ice shelf melting rates (e.g., Le Brocq et al., 2013)
and changing the buoyancy‐driven ocean circulation (section 2.2.2). Recent work to map, interpret, and
model subglacial hydrological systems has resulted in new insights into (1) the role of basal accretion
processes in modifying ice sheet thickness, (2) flow network topologies and their influence on ice flow,
(3) the role of active lake systems in providing time‐variable fluid flux, (4) the nature of the interaction of
freshwater melts with ice shelves and their cavities, and (5) the development and validation of new subgla-
cial hydrology modeling approaches.
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A wide range of subglacial flow topologies has been observed beneath the AIS, which may have a role in sta-
bilization and destabilization of glaciers. These flow network topologies include distributed flow, where
water is transported diffusely across a broad region of the bed; and channelized flow, where the flow is con-
centrated into discrete channels incised either into the ice (Röthlisberger, 1972) or into the substrate below,
forming channels (Nye, 1976) or tunnel valleys (Kehew et al., 2012) (Figure 14). The flow network topology
beneath the ice sheet is affected by variations in subglacial environment, including the geothermal heat flux,
subglacial topography and substrate composition, and presence of meltwater. The downstream switch from
distributed to concentrated channels, for example, observed in the upstream and downstream sections of the
Thwaites Glacier (Schroeder et al., 2013), results from characteristics of the substrate and changes in surface
slope, meltwater flux, and basal shear stress. It is not clear the extent to which the distinct anisotropy imaged
at Thwaites Glacier is a general rule; however, the tendency for channelized flow increases from the ice sheet
interior to the margin, as meltwater flux increases (Ashmore & Bingham, 2014). A more distributed flow net-
work topology is associated with more variable behavior in ice stream dynamics, as seen in the Siple Coast
(Bougamont et al., 2015). Subglacial hydrology models have been shown to be able to capture the transition
from distributed to channelized flow, agreeing well with specular reflection estimates derived from radar
echo sounding used to detect the transition in hydrological flow (Dow et al., 2020).
Basal accretion of ice by refreezing of subglacial water influences the thickness and flow characteristics of
the EAIS. A key example of these basal accretion mechanisms is found in the Gamburtsev Subglacial
Mountains, where alpine‐like valleys are well preserved by current ice sheet processes (Bo et al., 2009;
Creyts et al., 2014; Rose et al., 2013). In this system subglacial meltwater is formed in the deeper valley bot-
toms and is driven through pressure gradients up the ancestral drainage paths toward the valley heads,
where it encounters thin ice and freezing conditions (Creyts et al., 2014). The accretion of ice occurs through
two main processes. Conduction‐dominated freezing is driven by the increasingly cold conditions toward the
ice sheet surface, which cause ponded water to freeze; glaciohydraulic supercooling is driven by the reduc-
tion in pressure associated with flow up a steep gradient (Alley et al., 1998; Creyts et al., 2013). The role of
basal accretion in the overall evolution of the ice sheet remains poorly defined. However, observations sug-
gest that basal accretion is an important process in ice dynamics in some areas, and that it has operated over
long time scales (Bell et al., 2011).
Active subglacial lake systems beneath the AIS can modify the flow of ice streams and outlet glaciers
depending on the mechanism of lake drainage. Drainage events from active lake systems have been observed
to influence ice sheet dynamics at several locations including Byrd Glacier (Victoria Land) (Stearns
Figure 14. Basal hydrology and fluid‐interchange processes. Schematic showing channelized subglacial drainage
network (after Ashmore & Bingham, 2014) beneath a major ice stream. The cross section shows subglacial lakes and
channels, and active subglacial hydrogeology in sedimentary sequences. A basal channel in the ice shelf cavity reflects a
localized input of fresh meltwater to the cavity that will impact on the sub‐ice‐shelf melt rates, particularly at the
grounding line, which will in turn affect the ice stream behavior (e.g., Le Brocq et al., 2013). Red arrows represent
variable geothermal heat flux (GHF) interacting with the base of the ice sheet to produce meltwater.
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et al., 2008), Whillans Ice Plain (Siple Coast) (Siegfried et al., 2014), Crane Glacier (Antarctic Peninsula)
(Scambos et al., 2004), and Cook Glacier (McMillan et al., 2013; Miles et al., 2018). The incorporation of lake
drainage processes in subglacial hydrology models is still under development. Understanding of lake drai-
nage processes in Antarctica is based largely on theories for the drainage of lakes in alpine environments
(e.g., G. K. C. Clarke, 2003; Nye, 1976; Werder et al., 2013), with knowledge that the spatial and temporal
scales over which the subglacial network evolves are greater in Antarctica (Fricker et al., 2007). However,
the validation of lake drainage mechanisms in Antarctica is difficult in the absence of comprehensive obser-
vations of the subglacial environment. A new model by Carter et al. (2017) for the AIS suggests a lake filling
phase is followed by distributed flow once the lake level reaches a maximum, after which channelized flow
develops during peak outflow conditions. The filling phase is likely associated with slowed ice sheet flow
downstream, through the withdrawal of basal water from the hydrological system, and an associated
increase in basal shear stress (B. E. Smith et al., 2009). Conversely, the drainage phase is associated with
accelerated flow downstream, with reduced basal shear stress (Fricker et al., 2016). Although these concepts
have been defined and modeled, much uncertainty remains regarding their spatial and temporal frequency,
and the intensity of active lake drainage events on large‐scale AIS dynamics. The same applies to Antarctic
subglacial lake environments and processes in general, which are still mainly inferred from theoretical con-
siderations, combined with findings from ice accreted at the roof of subglacial Lake Vostok (Karl et al., 1999;
Shtarkman et al., 2013; Siegert et al., 2001), ice penetrating geophysical data (e.g., Smith et al., 2018), and or
the paleorecord (Kuhn et al., 2017). However, recent drilling projects, such as drilling of the Subglacial Lake
Whillans as part of the WISSARD project (Christner et al., 2014; Hodson et al., 2016; Michaud et al., 2016),
and Subglacial Lake Mercer on behalf of the SALSA project (https://salsa‐antarctica.org/), may soon have
the potential to answer many open questions.
The transport of freshwater to the ice shelf cavity can have profound impacts on ice shelf basal melt rates
(Jenkins, 2011; Le Brocq et al., 2013), ice streaming (Stearns et al., 2008), ocean stratification and mixing,
and nutrient fluxes (Corbett et al., 2017). For example, Le Brocq et al. (2013) used satellite and airborne
remote sensing to identify large channels under the Filchner‐Ronne Ice Shelf. These channels likely formed
due to the discharge of cool, fresh subglacial hydrological meltwater entering the sub‐ice shelf cavity at the
grounding line. Being fresh relative to the ocean water, this fresh subglacial meltwater rises to the ice shelf
boundary layer, entraining warm ocean water that enhances ice shelf melt rates. However, another assess-
ment by Alley et al. (2016), which mapped the basal channels in the ice shelves surrounding Antarctica,
found that the majority of channels in regions undergoing rapid change and ice mass loss, for example,
the Amundsen Sea sector, are more likely generated by ocean‐driven mechanisms alone rather than by sub-
glacial hydrology processes.
Concerted efforts over recent years have led to the development of numerical models that simulate the spa-
tial and temporal variability in Antarctic subglacial hydrological networks (Carter & Fricker, 2012; Dow
et al., 2016; Le Brocq et al., 2009). Nevertheless, due to a paucity of observations, validation of these models
is difficult, in particular with respect to the dynamics governing their development, evolution, and feedbacks
to ice sheet dynamics. Efforts to improve model physics and represent the processes governing drainage sys-
tems are ongoing, through initiatives such as the Subglacial Hydrology Model Intercomparison Project
(SHMIP; De Fleurian et al., 2018). Furthermore, the application of coupled subglacial hydrology/ice sheet
models will enable sensitivity studies on relatively poorly known hydrological parameters and boundary
conditions (Werder et al., 2013). Combining numerical modeling efforts with field measurements into the
future is an essential step toward a systematic understanding of subglacial hydrology and ice sheet dynamics
in Antarctica.
4.2.4. Hydrogeology Within the Subglacial Bed
Fluid exchanges across the ice sheet bed, and interactions with the groundwater beneath, are likely to be
common in Antarctica and may have a substantial impact on regional ice sheet dynamics (Gooch et al., 2016;
Person et al., 2012). Currently, this aspect of ice sheet behavior is not captured in large‐scale ice sheet mod-
els. Subglacial water at the base of the ice sheet is subject to several forms of exchange with the subglacially
eroded, transported, and deposited sediments (till) and permeable bedrock that underlies the till and the ice
sheet. Fluid exchange occurs across the bed on a variety of depth and time scales. Here we make a distinction
between hydrogeology of the till that is characterized by a thin layer of very high permeability and
low‐rigidity sediments operating on short time scales and bedrock‐aquifer exchange that is characterized
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by deeper storage, lower permeability, and higher rigidity strata, operating on longer time scales. The pre-
sence of relatively weak sedimentary strata is an important precondition for the development of a till‐rich
bed (Anandakrishnan et al., 1998; Bell et al., 1998), so that these hydrogeological systems are likely con-
nected over longer time scales.
The existence of water‐saturated and easily deformed till at the base of the AIS was identified in key studies
of the Siple Coast region (Alley et al., 1986; Blankenship et al., 1987; MacAyeal, 1989). Recent studies there
have continued to define the hydrogeology of till beds, with particular developments in the understanding of
hydrogeological processes in the stagnation and reactivation of ice streams, including interactions between
the ice sheet, water, and till. Important factors include the mechanical behavior of till, which can absorb and
release water under ice sheet loading and influence basal friction (Christoffersen et al., 2014). Till loading
time scales include weeks to months as a result of tidal loading, causing till compaction with the accumula-
tion of tidal effects (Christianson et al., 2013; Leeman et al., 2016). The response of till to tidal loading is sug-
gested as a key factor in stabilizing the grounding zone as the till becomes dewatered and stiffer under
repeated loading (Christianson et al., 2013). Till loading can also be impacted on decadal time scales as a
result of the distribution and routing of water, which can influence ice sheet flow due to changes in the dis-
tribution of highly consolidated and unconsolidated till units associated with hydrological changes. This
flow reorganization will be most substantial where ice streams are defined by the dynamics of flow rather
than by topography, and water is more easily transferred from one ice stream to another. Overall, AIS
dynamics are highly sensitive to the hydrogeology and mechanics of till layers, with processes that operate
rapidly but are associated with substantial reorganizations of ice flow at a regional scale.
Evidence for the role of deformable till layers in ice dynamics has also been observed for other West
Antarctic glaciers, such as Thwaites Glacier (Parizek et al., 2013), Pine Island Glacier (Brisbourne et al., 2017)
and Rutford ice Stream (King et al., 2009; King, Pritchard, & Smith, 2016; Smith & Murray, 2009). East
Antarctica is likely to possess similar characteristic in places where fast‐flowing ice streams flow over
sedimentary‐substrate and generate a till layer on top. However, systematic studies are yet to be conducted
in many places. Further investigation of the existence and characteristics of till layers beneath the AIS, and a
fuller understanding of how these systems evolve over longer time scales is required to integrate these beha-
viors more fully into models of AIS evolution.
In comparison with the highly active hydrogeology of till layers, aquifers in sedimentary strata and crystal-
line bedrock beneath the AIS have received relatively little attention over recent decades (Siegert et al., 2016).
It is very likely, however, that these systems are linked (Anandakrishnan et al., 1998; Bell et al., 1998, 2007)
and so longer‐term cycling of water through bedrock aquifers may be significant in releasing or storing sub-
glacial water at centennial to multimillennial time scales and transporting heat from within the crust to the
base of the AIS.
Ice sheets of the Northern Hemisphere, including the Laurentide and Fennoscandian paleo ice sheets, are
known to have generated substantial changes in bedrock hydrogeology (Person et al., 2012, 2007). Key pro-
cesses include recharge of aquifers from glacial meltwater, with up to 40% of meltwater infiltrating into the
bedrock (Lemieux et al., 2008), and the establishment of excess hydraulic head by glacial loading and water
transport over hundreds of meters to kilometers vertically, and tens to hundreds of kilometers laterally
(McIntosh et al., 2011). Deglaciation is associated with opposite tendencies, that is, discharge and a reduc-
tion in head. These processes generate a positive feedback that may enhance the supply of water to the ice
sheet bed during deglaciation and increase the potential for dynamic ice mass loss. Advance and retreat
effects are not equal because sediment compressibility is partially irreversible (Fowler & Yang, 2002).
Because of this, “fossil”hydraulic head can be preserved from one glacial cycle to the next (Person
et al., 2012). This means that a knowledge of loading history during previous cycles is needed to understand
present conditions, which may lead to longer‐term feedbacks that are not accounted for in short model runs.
]?twb.]Few studies exist of deep‐penetrating groundwater within Antarctica. However, recent evidence sug-
gests that processes operating under the AIS are similar to those understood for the Northern Hemisphere
ice sheets. Gooch et al. (2016) studied the potential for a hydrogeological and hydrothermal system in the
Wilkes Land region. This region is characterized by widespread sedimentary basins (Aitken et al., 2014,
2016; Frederick et al., 2016; Jordan et al., 2013), with the likelihood of active subglacial hydrology, including
many lakes (Wright et al., 2012), and suggestions of moderate to high geothermal heat flux (Carson
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et al., 2014), which are key factors associated with hydrogeological feedbacks. Gooch et al. (2016) model a
single cycle of glacial loading and unloading and find that subsurface pressure effects in the subglacial aqui-
fer are of the order of 10–30 MPa and extend to depths of a kilometer. Following unloading, overpressures of
up to 10 MPa remain in the aquifer. Under the conditions modeled, discharge of water is associated with
increased heat flux, with an additional ~40 mW m
−2
for a flux of 1 mm yr
−1
(Gooch et al., 2016). While rela-
tively simple, this study demonstrates that glacial activity is linked to groundwater processes which are very
poorly characterized, but to which the AIS may be highly sensitive.
5. Evidence for Ice Sheet Change
Direct observations of AIS mass balance changes, with sufficiently low uncertainty to reveal meaningful
changes over time, have revealed accelerating mass loss over the past twenty years, and spatially highly vari-
able AIS change. However, large internal variability of the AIS, predominantly driven by changes in atmo-
spheric and ocean dynamics (section 2), means that attributing the drivers of ice sheet evolution to natural or
human‐induced causes is difficult. One approach to understanding the AIS response to natural climate for-
cing is to look at its past history, when the ice sheet went through phases of disequilibrium with the ambient
climate. Ice and sediment core records reveal the natural long‐term ice sheet evolution over time scales of
decades to multiple millennia. The following section reviews the advances in observations and understand-
ing of AIS dynamics over the past 5 years based on modern measurements and paleoreconstructions, based
mostly on marine sediment cores.
5.1. Modern Observations
All approaches to estimating the overall AIS mass balance suffer from systematic and random errors that
affect estimates in different ways. Three methods are commonly used: (1) The mass budget method where
surface mass balance (snowfall minus surface ablation) is calculated using a regional atmospheric model,
while perimeter loss (flow across the grounding line) is calculated using satellite radar interferometry
(Gardner et al., 2018; Rignot et al., 2008, 2019); (2) satellite laser or radar altimetry (e.g., ICESAT‐1/2 and
Cryosat‐2) where surface elevation changes are repeatedly mapped over distinct tracks and then averaged
to provide estimates of volumetric change (McMillan et al., 2014; Paolo et al., 2015; Pritchard et al., 2009,
2012; The IMBIE Team et al., 2018) that are converted to mass change; and (3) gravimetry (e.g., GRACE
and GRACE‐Follow On) where changes in the strength of the gravitational field over the Earth are used
to estimate changing ice sheet mass (e.g., King et al., 2012; Velicogna et al., 2014). Each method has uncer-
tainties and limitations, and we gain confidence in estimates of mass balance when results from different
methods converge (e.g., Shepherd et al., 2012; The IMBIE Team et al., 2018). Overall, the remote sensing data
reveal that Antarctica is losing mass and contributing to sea level rise (Table 1 and Figure 15).
Assessment of the AIS mass over the past four decades reveals an accelerating rate of ice mass loss, from
40 ± 9 Gt yr
−1
in 1979–1990 to 50 ± 14 Gt yr
−1
in 1989–2000, 166 ± 18 Gt yr
−1
in 1999–2009, and
252 ± 26 Gt yr
−1
in 2009–2017 (Rignot et al., 2019). The accelerating decline in AIS mass has been driven
by the dramatic increase in ice loss of the WAIS during the period 2009–2017 (163 Gt yr
−1
), compared to
1999–2009 (73 Gt yr
−1
) (Rignot et al., 2019). Major outlet glaciers of Pine Island and Thwaites draining
the WAIS into the Amundsen Sea exhibit the most striking changes (Shepherd et al., 2019), including surface
lowering, ice flow acceleration, ice shelf thinning, and grounding line retreat (Christie et al., 2016; Konrad
et al., 2018; Mouginot et al., 2014; Rignot et al., 2014). Significant ice volume reduction is also recorded in
the Antarctic Peninsula, where approximately 30 Gt yr
−1
of ice has been lost to the oceans over the last
10–15 years (Table 1). One of the most significant and concerning findings since publication of the IPCC
AR5 in 2013 is that MISI (see section 3.1.2) may have initiated at Thwaites and Pine Island glaciers in the
Amundsen Sea sector of the WAIS (Favier et al., 2014; Joughin et al., 2014). Although a recent study suggests
that grounding line retreat at Pine Island Glacier may have stabilized (Konrad et al., 2018) several ice sheet
models show that projected retreat of glaciers in the Amundsen Sea embayment due to MISI can potentially
cause hundreds of kilometers of grounding line retreat and high rates of ice loss on centennial time scales
(Cornford et al., 2015; Ritz et al., 2015). In some projections (RCP 4.5 and above) this leads to deglaciation
of much of the WAIS (DeConto & Pollard, 2016; Golledge et al., 2015), and multimeter sea level rise on cen-
tennial to millennial time scales (see section 6.1).
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Table 1
Recent Antarctic Mass Balance Estimates in Gigatons per Year (Gt yr
−1
) for Different Sectors and the Continent as a Whole (Total)
Method East Antarctica West Antarctica Antarctic Peninsula Total
Mass Budget
(Gardner et al., 2018)
+61 ± 73 Gt yr
−1
(2008.0–2016.0) −214 ± 51 Gt yr
−1
(2008.0–2016.0) −31 ± 29 Gt yr
−1
(2008.0–2016.0) −183 ± 94 Gt yr
−1
(2008.0–2016.0)
Cryosat2 (McMillan
et al., 2014)
−3 ± 36 Gt yr
−1
(2010.8–2013.8) −134 ± 27 Gt yr
−1
(2010.8–2013.8) −23 ± 18 Gt yr
−1
(2010.8–2013.8) −159 ± 48 Gt yr
−1
(2010.8–2013.8)
GRACE (Harig &
Simons, 2015)
+56 ± 5 Gt yr
−1
(2003.0–2014.5) −121 ± 8 Gt yr
−1
(2003.0–2014.5) −27 ± 2 Gt yr
−1
(2003.0–2014.5) −92 ± 10 Gt yr
−1
(2010.8–2013.8)
GRACE (Forsberg
et al., 2017)
−92 ± 50 Gt yr
−1
(2002.3–2016.0)
GRACE+SLR (Talpe
et al., 2017)
−56 ± 28 Gt yr
−1
(1993.0–2000.0)
Combined method
a
(Martín‐Español,
King, et al., 2016)
+56 ± 18 Gt yr
−1
(2003.0–2014.0) −112 ± 10 Gt yr
−1
(2003.0–2014.0) −28 ± 7 Gt yr
−1
(2003.0–2014.0) −84 ± 22 Gt yr
−1
(2003.0–2014.0)
IMBIE 2 Reconciled
mass balance
b
(The
IMBIE Team
et al., 2018)
+5 ± 46 Gt yr
−1
(1992–2017) −94 ± 27 Gt yr
−1
(1992–2017) –20 ± 15 Gt yr
−1
(1992–2017) −109 ± 56 Gt yr
−1
(1992–2017)
Rignot et al. (2019)
c
−31 Gt yr
−1
(1979–2017) −68 Gt yr
−1
(1979–2017) −23 Gt yr
−1
(1979–2017) −123 Gt yr
−1
(1979–2017)
Note. Mass balance estimates are from the three remote sensing approaches mentioned in the text, one combined method and the IMBIE reconciled estimate. Observational periods are listed in
brackets. Uncertainties represent 1σvalues. Note that ~363 Gt yr
−1
of mass loss will cause ~1 mm yr
−1
of sea level rise.
a
A Bayesian method combining altimetry and gravimetry data, incorporating glacio‐isostatic adjustment information (Martín‐Español, King, et al., 2016).
b
Reconciled estimate includes all
three methods of assessing mass balance over a common time scale,updated to IMBIE 2 (The IMBIETeam et al., 2018).
c
Values from Table 1, Cumulative Balance (Gt) between 1979 and 2017.
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Prior to the assessment of Rignot et al. (2019), which estimated an EAIS mass loss of −57.0 ± 2 Gt yr
−1
since
1979, multiple and independent satellite observations indicated that the EAIS was in near balance, with esti-
mates ranging from −3 to +36 to +61 ± 73 Gt yr
−1
(Table 1). One set of estimates from radar and laser alti-
metry suggests substantially greater growth, but these estimates have been controversial (Richter et al., 2016;
Scambos & Shuman, 2016; Zwally et al., 2015). EAIS mass balance is regionally heterogeneous, and there are
two notable regions that show anomalies. The Wilkes Land margin exhibits a large amount of interannual
variability in mass balance. In the period from 2010–2013, the Wilkes Land sector of the EAIS lost up to
30 Gt yr
−1
(Martín‐Español, King, et al., 2016), with ice shelf thinning and inland ice flow acceleration,
which is believed to be a result of increased CDW incursion (Flament & Rémy, 2012; Li et al., 2016;
Rignot et al., 2019). In East Antarctica, this mass loss was balanced or exceeded by mass gain in the
Dronning Maud Land sector, where increased snowfall between 2009 and 2011 resulted in a mass gain of
~350 Gt, equivalent to an ~1 mm sea level drop (Boening et al., 2012). This sector of the ice sheet has also
experienced recent grounding line advance (Konrad et al., 2018), likely indicating rapid coupling between
surface accumulation and ice sheet dynamics at a regional scale.
Attributing sea level rise to changes in AIS mass balance requires closure of the sea level budget, where the
observed changes in sea level equal the sum of changes in ocean properties (temperature and salinity) and
exchange between various water reservoirs (groundwater, glaciers, and major ice sheets) (Leuliette &
Willis, 2011). Corrections based on geophysical responses to changes in sea level and the response of the
ocean and solid Earth to past ice loading history remain significant sources of uncertainty. The contribution
of Antarctica to sea level change calculated from GRACE data, available since 2003, shows an accumulated
sea level rise of 5 mm (Forsberg et al., 2017) and a clear acceleration in Antarctic ice melt between 2002 and
2016. Systematic effects are introduced into estimates of the rate of Antarctic ice mass change derived from
GRACE, due to mismodeled mantle mass change associated with GIA (King et al., 2012). Models of GIA
do not fully represent lateral variations in Earth properties and are poorly constrained by a lack of under-
standing of post‐LGM ice mass changes, which has resulted in large intermodel differences (section 4.2.2)
(Martín‐Español, King, et al., 2016). The acceleration in ice mass loss is not debated, but differences arise
in temporal and regional estimates depending on the method (Table 1). Despite the uncertainties, attempts
have been made to improve the attribution of the observed sea level rise into its constituent parts (Chen
et al., 2017; Dieng et al., 2017). These attempts have been aided (Clark et al., 2015) by GRACE observations
of ocean mass and Argo float measurements of thermal expansion in the upper 2,000 m of the ocean (Durack
et al., 2014; Roemmich et al., 2015), at least since the early to middle 2000s. Improved constraints on the his-
tory of the AIS (e.g., LGM to late Holocene) and on heterogeneities in the solid Earth response will improve
Figure 15. A selection of recent estimates of Antarctic ice mass change (left axis) and its sea level equivalent (right axis)
as given in Table 1. Estimates based on the combination of multiple data sets are labeled “Hybrid.”Updated from Allison
et al. (2018).
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estimates of global sea level change. As mentioned above (section 5.1),
Antarctica has been, overall, losing mass into the oceans since the 1990s
and is estimated to have contributed ~0.4 mm yr
−1
to GMSL over
2005–2015 (WCRP Global Sea Level Budget Group, 2018, and references
therein), and has been shown to be accelerating, to an Antarctic contribu-
tion over 2012–2016 of ~0.55 mm yr
−1
(Oppenheimer et al., 2019).
5.2. Paleoenvironmental Observations
The long‐term relationship between sea level and atmospheric carbon
dioxide concentrations in Earth's geological past can provide insight into
future sea level responses and reveal the vulnerability of the polar ice
sheets to relatively small changes in temperature. Foster and
Rohling (2013) demonstrated a sigmoidal relationship between atmo-
spheric CO
2
levels and global sea level throughout the past 40 million
years, which holds despite significant changes in boundary conditions
related to plate tectonic changes (Figure 16). Their assessment indicates
that climate warming of 2°C relative to preindustrial (CO
2
between 400
and 450 ppm) results in sea level rise of more than 9 m above present
(68% confidence) in the long term (see also Rohling et al., 2013). This eva-
luation is within the range of estimated sea level rise of 6–9 m for the LIG
(Dutton & Lambeck, 2012; Dutton, Carlson, et al., 2015; Hay et al., 2014;
Kopp et al., 2009; Rohling et al., 2008) when global temperatures were
≤1°C warmer (i.e., similar to today; Hansen et al., 2017; Hoffman
et al., 2017; Otto‐Bliesner et al., 2013; Turney & Jones, 2010). However,
reconsideration of the ice mass distributions of the preceding glaciation
(Marine Isotope Stage 6), and thus of the glacio‐isostatic changes through
the previous deglaciation (Termination II), suggests that existing LIG
highstand estimates may need revising upward by roughly 2 m (Rohling
et al., 2017).
Four periods in Earth history are especially important for gaining a better understanding of the sensitivity of
the AIS to changes in the climate system: The Holocene (11 ka [thousand years ago] to present), the last
deglaciation (19–11 ka) when temperatures rose by 4–5°C and GMSL rose by 120–140 m; the LIG period
(130–115 ka) when global temperatures were similar to modern values but GMSL stood potentially up to
10 m higher; and the Pliocene (5.33–2.58 million years ago) when global temperatures were 2–3°C higher
and GMSL was 10–30 m above the present. Constraining the spatial and temporal behavior of the AIS
and the corresponding sea level change during these time intervals requires not only AIS‐specific informa-
tion (e.g., its geographical extent and inferred mass) but also the rheologi-
cal properties of the solid Earth and their interaction through time. These
complexities are highlighted in the following sections, in addition to the
chronological dating issues that arise when working close to Antarctica
(e.g., see Livingstone et al., 2012; Siegert et al., 2019) and the complex
and regionally variable patterns of deglaciation (The RAISED
Consortium et al., 2014).
5.2.1. Holocene Ice Dynamics
The Holocene is defined as the period from about 11.6 ka to the onset of
industrial‐era warming (about 1850 CE; see Abram et al., 2016) and follows
the transition out of the last glacial period. Sea level rise continued from
the last glacial termination into the Holocene, but most of the 60 m of glo-
bal sea level rise in the Holocene occurred over a broad multimillennial
interval between 11 and 7 ka (Lambeck et al., 2014; Stanford et al., 2011)
(Figure 17). The AIS contribution to Holocene sea level rise remains poorly
constrained. Geological studies on and off the continent (e.g., Anderson
et al., 2014; Hillenbrand et al., 2014; Leventer et al., 2006; Mackintosh
Figure 16. Sea level versus CO
2
concentrations (and the logarithmic
radiative forcing influence of CO
2
changes expressed by ln (CO
2
/C
0
),
where C
0
represents the preindustrial CO
2
level of 278 ppmv) (after Foster
& Rohling, 2013). Symbols represent reconstructions with uncertainties
for different intervals of the past 40 million years. The black line and orange
envelopes represent a probabilistic assessment that takes into account full
propagation of all uncertainties (black line is the probability maximum;
dark orange is the 68% probability interval; light orange is the 95%
probability interval).
Figure 17. (a) The global sea level curve (thick black line) based on paleo
sea level records (black dots with depth uncertainties shown in vertical
blue lines) (Lambeck et al., 2014) and (b) rate of sea level change in meters
per century for the past 20 ka, modified from Clark et al. (2016).
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et al., 2014; Ó Cofaigh et al., 2014) and GIA modeling (Argus et al., 2014) are complex and show regional het-
erogeneity in the timing and rates of ice sheet retreat since the LGM. Determining the nature of this hetero-
geneity is important in the context of constraining AIS contribution to global sea level rise during the
Holocene. The aforementioned studies do clearly suggest a limited contribution from either the WAIS or
the EAIS, relative to Northern Hemisphere sources, but data and modeling comparison studies indicate that
rapid and millennial‐scale variability of the AIS was a feature of Holocene AIS dynamics (Bakker et al., 2016;
Johnson et al., 2019; Spector et al., 2017). In particular, many studies on AIS deglaciation, based on ice thin-
ning histories, point to rapid deflation events near the present‐day ice sheet margin in the Weddell and Ross
Sea sectors during the middle to late Holocene, as the main phase of ice sheet retreat (Hein et al., 2016;
Johnson et al., 2019; Jones et al., 2015; Spector et al., 2017, and references therein). Conversely, finite radio-
carbon ages obtained from sediments sampled from beneath the grounded WAIS suggest retreat upstream of
its present location, and then readvance back to its modern position during the Holocene (Kingslake
et al., 2018). This readvance scenario is proposed to have resulted from continuing isostatic rebound of the
seafloor, and although the radiocarbon ages cannot pinpoint the timing of retreat, model experiments con-
ducted to test the feasibility of this scenario suggest that the readvance likely happened after the grounding
line had reached its minimum extent by the early Holocene (>9.6 ka) (Kingslake et al., 2018).
Along the margins of West Antarctica, internal processes relating to the overdeepened continental shelf
(Patton et al., 2016) and MISI are thought to have led to ongoing retreat into the Holocene (Anderson
et al., 2014; Hillenbrand et al., 2014; Larter et al., 2014; McKay, Golledge, et al., 2016; Ó Cofaigh et al., 2014;
Smith et al., 2014). Whether this retreat was relatively gradual and sustained through to the present day, or
was more rapid with periods of greatly accelerated retreat, remains open to debate and may have varied
between sectors (Bart et al., 2017; Hillenbrand et al., 2013). However, the pattern and exact timing of retreat
are ambiguous and vary from region to region, in part due to issues defining the chronology and differences
relating to geometry of the ice sheet, as well as dynamic feedback processes, such as basal topography (e.g.,
Kingslake et al., 2018) and sediment deposition (e.g., Bart et al., 2018) that control ice sheet retreat. Despite
uncertainties in the Holocene WAIS configuration, mismatches between grounding lines and isostatic
rebound have revealed a more complex Late Holocene retreat pattern than previously accepted and are
heavily influenced by geological processes (Bradley et al., 2015; Kingslake et al., 2018; Lowry et al., 2020;
Siegert et al., 2013), while it is also increasingly apparent that GIA variations throughout the Holocene need
to be considered in the context of AIS dynamics (Simms et al., 2018) and the impact of GIA corrections and
previous ice history models on ice sheet thinning histories determined by cosmogenic exposure dating (Jones
et al., 2019).
Evidence exists for abrupt mid‐Holocene ice sheet loss events (e.g., Fogwill et al., 2014; Johnson et al., 2019),
which have been linked in some cases to atmosphere‐ocean forcing and ice shelf collapse (Hillenbrand
et al., 2017). Rapid ice sheet thinning in the central Ross embayment during the early to mid‐Holocene
has been shown by comparisons of foraminifera‐based radiocarbon dates from a marine sediment core near
Ross Island and cosmogenic (
10
Be) exposure ages from Transantarctic Mountain outlet glaciers.
Accordingly, to the south of the modern Ross Ice Shelf calving line, model comparisons with geological data
indicate that most of this embayment retreated between 9 and 8 ka (Lowry et al., 2019; McKay, Golledge,
et al., 2016; Spector et al., 2017). However, to the north of the Ross Ice Shelf at Mackay Glacier, rapid thin-
ning occurred at around 7 ka, due to MISI as this EAIS outlet glacier retreated into an overdeepened marine
basin. The rate of AIS outlet glacier deflation events during the early to mid‐Holocene in the Ross and
Amundsen Sea sectors of WAIS was similar to that observed in other rapidly changing parts of Antarctica
today (Johnson et al., 2014; Jones et al., 2015). The contradiction in timing between the cosmogenic data
to the north and south of Ross Island could potentially be resolved by high‐resolution multibeam studies
which indicate that residual EAIS outlet glaciers readvanced into the Western Ross Sea, following wide-
spread grounding line retreat in the wider Ross Sea embayment (Greenwood et al., 2018; Lee et al., 2017).
Abrupt mid‐Holocene changes have also been observed in the WAIS section that drains into the Weddell
Sea. Cosmogenic data in the Ellsworth Mountains suggest that the catchment feeding into the Weddell
Sea remained at close to maximum ice thickness at 10 ka, but a rapid thinning event occurred in the eastern
Weddell Sea embayment between 8 and 6 ka (Johnson et al., 2019), while thinning near the modern ground-
ing line occurred rapidly between 6.5 and 3.5 ka (Hein et al., 2016). However, determining the significance of
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these ice sheet thinning events to the offshore extent of grounded ice is more ambiguous in this region, than
in the Ross Sea, due to a sparse offshore data set and large disagreements in the interpretation relating to
terrestrial‐and marine‐based data sets in this region (Hillenbrand et al., 2014). Model comparisons with ter-
restrial geological data from cosmogenic studies in mountain ranges at the margins of the Filchner‐Ronne
Ice Shelves suggest that an ice sheet was grounded on bathymetric highs toward the continental shelf edge,
with floating ice shelves over the deep bathymetric troughs. One model‐based interpretation, constrained by
cosmogenic deflation data, proposes that the grounding lines had retreated near to their present‐day loca-
tions by 10 ka (Whitehouse, Bentley, Milne, et al., 2012). Conversely, interpretations of reworked foramini-
fera from marine sediment cores and multibeam analysis imply an ice sheet that was grounded near the
continental shelf break until at least 20 kyr BP and that most grounding line retreat was more gradual
and postdates 10 ka (Arndt et al., 2017; Hillenbrand et al., 2014; Hodgson et al., 2018). Review of these
two different scenarios by Siegert et al. (2019) highlights that geophysical and cosmogenic data remain con-
sistent with either scenario but clearly indicate that there was a major redirection of ice flow in the Weddell
Sea embayment during the mid‐Holocene. It was proposed isostatically driven processes could be invoked to
explain either scenario: either through a redirection of subglacial hydrology leading to a change in dynamic
ice flow in the mid‐Holocene or through rebound of the seafloor following early Holocene retreat, resulting a
the regrounding of the ice shelf and subsequent a readvance of the grounding line in the mid‐Holocene
(Siegert et al., 2019, and references therein).
The role of CDW in driving the deglaciation in the Amundsen Sea region and west of the Antarctic
Peninsula, similar to modern observations of wind‐driven incursions of relatively warm CDW onto conti-
nental shelf regions of drivers of deglaciation (Christianson, Jacobel, et al., 2016; Jacobs et al., 2011, 2012;
Jenkins et al., 2018; Klinck et al., 2004; Nakayama et al., 2013; Smith et al., 1999; Wåhlin et al., 2016) have
been investigated using geochemical data and ecological constraints from foraminifera and diatoms. These
paleoenvironmental studies suggest the presence of CDW during the deglaciation, which ceased by ~7.5 ka
(Hillenbrand et al., 2017; Minzoni et al., 2017; Peck et al., 2015). The decrease in CDW inflow onto the
Antarctic continental shelves during the mid‐Holocene is proposed to be due to reduced upwelling at the
Antarctic margin as the westerly winds migrated northward during the early to mid‐Holocene (Anderson
et al., 2009; McGlone et al., 2010; Toggweiler et al., 2006).
Whether the final Holocene retreat, which happened after the main deglaciation, was relatively gradual and
sustained through to present day, or was more rapid with periods of accelerated retreat, is still open to debate
(Bakker et al., 2016; The RAISED Consortium et al., 2014). Furthermore, finite radiocarbon ages obtained
from sediment cores collected beneath the WAIS have revealed that the grounding line in the Siple Coast
region retreated to the south of its modern position and subsequently readvanced as a consequence of iso-
static rebound although the exact timing of this retreat is unconstrained (Kingslake et al., 2018). Despite
these uncertainties, it is clear that each catchment area in West Antarctica responded in different ways dur-
ing the deglaciation, due to differences in regional oceanographic forcings, bathymetry, GIA, and internal
ice dynamics.
5.2.2. The Last Deglaciation
The most recent interval in Earth history when global ice sheets reached their maximum integrated volume
was the LGM. This period is traditionally defined by a sea level lowstand and lasted from ~27–19 ka (Clark
et al., 2009; Rohling et al., 2009). The transition out of the LGM into the present day, referred to as the last
deglaciation (~19–11 ka), saw 5 ± 2°C global mean temperature rise that resulted in loss of large ice sheets in
the Northern Hemisphere. There is a range of estimates for the resulting deglacial sea level rise from ~120 to
130 m (Austermann et al., 2013; Clark et al., 2009; Denton et al., 2010; Masson‐Delmotte et al., 2013; Peltier
& Fairbanks, 2006) to as much as ~134 to 140 m (Lambeck et al., 2014) (Figure 17).
Quantitative assessments of the AIS contribution to the deglacial sea level rise vary substantially depending
on the applied methodology, from 5–8 m based on regional assessments (Briggs et al., 2014; Gomez
et al., 2013; Pollard et al., 2016; Whitehouse, Bentley, & Le Brocq, 2012), to 14–23 m based on GIA models
(Argus et al., 2014; Lambeck et al., 2014). The geological evidence used for these reconstructions and model
assessments derives from marine and ice core records, cosmogenic dating of glacial erratics and nearby sea
level and geodetic data. The lower estimate of 5 m is based on a large ensemble of model runs (Pollard
et al., 2016) and only considers a deglacial sea level contribution from the WAIS. Other studies have
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Table 2
The Values Show the Contribution in Meters of Equivalent Global Mean Sea Level Rise From Melting of the AIS for Past Climate Periods (See Section 5.2) and
Estimates for Future Time Intervals
Note. Projections from IPCC AR5 (Church et al., 2013) show the medi\an and likely ranges and do not include ice dynamic feedbacks. Church et al. (2013) reported
an estimate of 0.07 (−0.01 to 0.16) for “Antarctic ice‐sheet rapid dynamics.”Note that DeConto and Pollard (2016) estimates are lower when tuned to Pliocene sea
level targets (e.g., 0.64 ± 0.49 to 13.11 ± 3.04 for RCP8.5). Their range for LIG estimates is based on modern initial conditions (c) and glacial initial conditions (d),
which reflects the difference and uncertainty in the ice sheet size. Global sea level projections by Golledge et al. (2019) show modeled contributions with (e) no melt
feedback and (f) with melt feedback, where fresh meltwater drives a positive feedback due to stratification of the water column that drives basal melting (see sec-
tion 6.2). Edwards et al. (2019) estimated sea level contributions with (h) no MICI and (i) with MICI. The expert judgment estimates of Bamber et al. (2019) are based
on a low (+2°C above preindustrial) and high (+5°C above preindustrial) temperature change scenarios for contributions from the WAIS (W) and EAIS (E).
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estimated an AIS contribution of 6.7 m when only cosmogenic surface exposure ages are used as constraints
on ice sheet thickness (Golledge et al., 2012) to 14.5 m, when isotope measurements from Antarctic ice cores
are instead used to guide modeled ice sheet thickness (Golledge et al., 2014) (Table 2). GIA model‐based esti-
mates depend on the viscosity of the upper mantle, which depending on the ice volume can result in the
same observed rates of rebound, but very different deglacial histories (cf. Argus et al., 2014). These large dis-
crepancies in the deglacial AIS contribution to post‐LGM sea level rise result from uncertainties associated
with quantifying the exact extent, thickness and therefore volume of the AIS during the LGM (Golledge
et al., 2014; Golledge et al., 2013; Whitehouse, Bentley, Milne, et al., 2012). This is due to not only the poor
data coverage and general inaccessibility but also the range of different approaches, each with their own lim-
itations and uncertainties that have contributed to this problem.
The maximum thickness and thinning history of the AIS has been reconstructed by terrestrial cosmogenic
exposure dating of glacially transported erratics, or previously ice‐covered bedrock and by geomorphological
bedforms mapped on land (e.g., Hein et al., 2016; Johnson et al., 2014; Mackintosh et al., 2011; Small
et al., 2019; Spector et al., 2017; Stone et al., 2003; Todd et al., 2010), as well as radiocarbon ages from glacial
drift deposits, lake sediments, or raised beaches (e.g., Hall et al., 2015). Maximum ice sheet extent has been
reconstructed from subglacial and proglacial bedforms preserved on the continental shelf and dated using
radiocarbon ages from carbonate material in marine sediment cores (Bart et al., 2018; Mackintosh et al., 2014;
Ó Cofaigh et al., 2016; Smith et al., 2019). High‐resolution mapping of glacial bedforms on the Antarctic con-
tinental shelves provides unequivocal evidence for the expansion of grounded ice sheets, although improved
chronological constraints are required to understand the spatial variability of the timing at which the max-
imum extent was reached in each sector (Arndt et al., 2017; Dowdeswell et al., 2016; Fernandez et al., 2018;
Halberstadt et al., 2016; Hodgson et al., 2018; Klages et al., 2014; Larter et al., 2019; Lee et al., 2017; Lynch
et al., 2014; Simkins et al., 2017; The RAISED Consortium et al., 2014; Wise et al., 2017). Quantifying the
LGM extent, and the timing of subsequent retreat, is required to test hypotheses regarding the climate or
oceanographic mechanisms for initiating widespread Antarctic marine ice sheet retreat. In this context, it
is essential to constrain the role that internal ice sheet dynamics, solid Earth processes (e.g., feedbacks from
the historical ice load), and oceanic/sea ice feedbacks had in either accelerating or slowing retreat during
both the last glacial termination and the Holocene.
A full synthesis of reconstructing Antarctic deglaciation since the LGM based on nearshore and onshore
archives was presented by Livingstone et al. (2012), Wright et al. (2008) and more recently The RAISED
Consortium (Anderson et al., 2014; Hillenbrand et al., 2014; Larter et al., 2014; Mackintosh et al., 2014;
Ó Cofaigh et al., 2014; The RAISED Consortium et al., 2014). A number of previous reconstructions
favored a minor AIS retreat that occurred late (after 14 ka) and continued well into the late Holocene
(Bentley, 2010; Mackintosh et al., 2011, 2014; Peltier, 2004). The majority of Antarctic ice mass loss may
have occurred on the overdeepened continental shelves of West Antarctica. These topographic depressions
(e.g., 300–1,500 m deep) can result in increased sensitivity to ocean forcing (see section 2.2.2) and MISI.
Topographic lows can be tectonically controlled, through extension and subsidence associated with the
West Antarctic Rift, (see section 4.1.1) in the Weddell and Ross Seas, and also formed by prior glacial ero-
sion, such as those observed on the inner shelves in the Amundsen and Bellingshausen Seas, and west of
the Antarctic Peninsula (e.g., Livingstone et al., 2012).
The RAISED Consortium concluded that Antarctica may have only contributed a few meters to global melt-
water pulses during the deglaciation; although no quantitative estimate was provided. The GIA‐based I6G
model supports ~13.6 m AIS contribution to global sea level rise since the LGM, with the most rapid period
of ice loss occurring between 12 and 5 ka (Argus et al., 2014). However, such estimates are highly dependent
on the values used for mantle viscosity in those models, and other estimates are approximately half of this
value (Whitehouse, Bentley, & Le Brocq, 2012). Large uncertainties remain regarding the EAIS contribu-
tions to post‐LGM sea level rise. GIA‐based models constrained by RSL records along the East Antarctic
coastline, and by elevation constraints from ice cores, suggest that there was thinning of several hundred
meters in coastal regions, however inland of the continental shelf edge ice thickness changes were minimal
(Argus et al., 2014). This assessment is consistent with the Mackintosh et al. (2014) data‐model synthesis of a
modest (1 m) EAIS contribution deglacial sea level rise.
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An early hypothesis was that post‐LGM ice sheet retreat in the Antarctic was primarily triggered by global
sea level rise, where orbitally driven melting of the Northern Hemisphere ice sheets drove sea level rise
around Antarctica and caused the destabilization of the grounded portion of the ice (Denton &
Hughes, 1983; Thomas & Bentley, 1978). However, recent modeling studies suggest that Southern Ocean
heat advection was a more significant driver for initiating glacial retreat during the LGM (Golledge
et al., 2012; Pollard & DeConto, 2009). Ocean thermal forcing and sea level control have also been used to
argue that, contrary to earlier assessments, the advance to (at 29–28 ka) and retreat from the maximum
extent of parts of the AIS during the LGM was nearly synchronous with Northern Hemisphere ice sheets
(Weber et al., 2011), with deglaciation commencing between 19 and 20 ka—although the timing does vary
regionally (e.g., Livingstone et al., 2012). A new understanding of deglacial AIS dynamics, as a series of mul-
tiple rapid ice discharge events that lasted from centuries to a millennium, has been informed by iceberg
rafted debris records in the so‐called “Iceberg Alley”of the Scotia Sea (Figure 18) (Weber et al., 2014). The
records provide an integrated signal of iceberg discharge from the Indian and South Atlantic sectors of the
AIS and show that Antarctic deglaciation accelerated between 17 and 9 ka, with multiple ice discharge
events that lasted from centuries to a millennium and that closely coincided with times of global meltwater
pulses (Figure 17). Although these records do not quantify the volume of AIS loss or rule out that the AIS was
destabilized by rapid sea level rise originating elsewhere, they do present evidence of centennial to millen-
nial response time scales of the AIS during the deglaciation, with abrupt discharge events that have compar-
able time scales to ice discharge events in the Northern Hemisphere ice sheets (Weber et al., 2014).
Figure 18. Iceberg‐rafted debris (IBRD) flux in Iceberg Alley of the Scotia Sea and other climate proxies during the last
deglaciation. (a) Stacked IBRD flux record. (b) The 500 year averages of stacked IBRD flux relative to Holocene
average (Weber et al., 2014). (c) Antarctic deglacial δ
13
C
atm
stack (Schmitt et al., 2012). (d) Biogenic opal flux records
from SE Atlantic Site TN057‐14 (dashed) (Anderson et al., 2009) and Scotia Sea Site MD07‐3134 (gray) (Sprenk
et al., 2013). (e) EDML ice core record of sea salt (ss) Na
+
‐flux. f) δ
18
O record from EDML ice core (EPICA Community
Members et al., 2006). Note that Meltwater Pulse 1A (MWP‐1a, red vertical arrow) is coeval with Antarctic Ice‐Sheet
Discharge Event 6 (AID6), within the dating uncertainties. Note further that major changes in Southern Hemisphere
commenced ~17 ka when the Northern Hemsiphere was cold during Heinrich Event 1 (H1). ACR: Antarctic Cold
Reversal, B‐A: Bœlling‐Allerœd, YD: Younger Dryas. Figure modified from Weber et al. (2014).
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An Antarctic contribution to rapid sea level rise over multicentennial periods, such as Meltwater Pulse 1a
(MWP‐1a; 14.6 ka) when GMSL rose by ~20 m in 400 years (Deschamps et al., 2012a; Grant et al., 2014;
Lambeck et al., 2014; Stanford et al., 2011), is supported by the timing of largest iceberg rafted debris flux
in the Scotia Sea record (Figure 17). Other millennial‐scale sea level rise events during the last glacial
(Siddall et al., 2003) have also been inferred to reflect considerable AIS meltwater contribution in addition
to melting of the Greenland Ice Sheet (Rohling et al., 2004). These studies challenge scenarios in which
the AIS made only a relatively small contribution to sea level rise since the LGM lowstand (e.g.,
Whitehouse, Bentley, & Le Brocq, 2012). Iceberg‐rafted debris records from Weber et al. (2014) are supported
by an ice sheet model that predicts an AIS contribution of 6 mm yr
−1
during the largest peak in deglacial
Antarctic ice loss centered on the MWP‐1a interval, compared to a deglacial average of 1 mm yr
−1
(Golledge et al., 2014). Overall this equates to a 2–3 m sea level equivalent (SLE) AIS contribution to the
MWP‐1a event, with the remainder coming from the Northern Hemisphere ice sheets (Golledge et al., 2014).
However, further evidence for an Antarctic contribution to MWP‐1a from data on land and the continental
margin in other regions such as the Weddell Sea (Arndt et al., 2017; Nichols et al., 2019) is absent, and this
issue remains a conundrum (Goehring et al., 2019; Hall et al., 2015; Prothro et al., 2020). The large uncertain-
ties in the underpinning sea level constraints (Hibbert et al., 2016, 2018; Stanford et al., 2011) and dating of
geological material on land and in the ocean (e.g., see discussion of cosmogenic dating in Siegert et al., 2019,
and radiocarbon dating in Anderson et al., 2014), and the uncertainty around the AIS size, seaward extent,
thickness and volume above flotation at the LGM, mean that currently it remains difficult to quantify the
exact contribution of AIS melting to the sea level rise recorded during MWP‐1a.
Ocean forcing was inferred as the key driver of deglacial AIS dynamics, modulated by global atmospheric
teleconnections, that decoupled ice sheet elevation and air temperatures in a high‐resolution ice core near
the Weddell Sea and resulted in rapid thinning of the AIS during the period coinciding with MWP1‐a
(Fogwill et al., 2017). The coincidence between changes in AIS elevation (Fogwill et al., 2017), enhanced ice-
berg flux (Weber et al., 2014), and atmospheric temperature trends (Pedro et al., 2016) through the deglacia-
tion suggests a tight coupling between the ice‐ocean‐atmosphere system. A positive feedback mechanism
was proposed where reinvigoration of the Atlantic Meridional Overturning Circulation (AMOC) following
Heinrich Event 1 increased Southern Ocean subsurface heat content and triggered initial melting of AIS
margins. Consequent freshening of surface waters led to a weakening of Southern Ocean overturning, result-
ing in reduced AABW formation, surface ocean stratification, and sea ice expansion (Fogwill et al., 2017;
Golledge et al., 2014; Weber et al., 2014). Sea ice expansion would result in ice‐albedo feedbacks that drive
atmospheric cooling (negative feedback), but the strong insulating effect of sea ice would also trap subsur-
face heat, which in addition to ocean stratification from AIS meltwater would trap ocean heat close to the
grounding line of the AIS and enhance thermal erosion to maintain a positive ice‐ocean feedback.
Defining the details of this dynamic feedback during periods of past climate change is critical to understand-
ing the implications of the high‐latitude Southern Hemisphere environmental changes today. This includes
the need for ocean temperature and circulation records close to the Antarctic margin to improve our under-
standing of the role of ocean forcing in driving ice sheet change, which mostly relies on model results and ice
core data.
Simulations of ocean oxygen isotopes (δ
18
O) suggest that potentially significant disruption of global over-
turning circulation may have coincided with AIS melt between 75 and 20 ka BP (Rohling et al., 2004).
During this period of highly variable climate (Bereiter et al., 2012; Dansgaard et al., 1993; EPICA, 2006;
North Greenland Ice Core Project members, 2004) and sea level (Cutler et al., 2003; Siddall et al., 2003)
millennial‐scale Antarctic warming of 2–3°C events (known as Antarctic Isotope Maxima) were associated
with global seal level rise events of 30 m, at rates of about 2 m per century (Grant et al., 2014). The glacial
ocean simulations of Rohling et al. (2004) highlight major changes in Atlantic overturning circulation, with
severely reduced or near‐collapsed North Atlantic Deep Water (NADW) export, associated with a
50:50 ± 20% contribution of meltwater from both Northern and Southern Hemisphere ice sheets that match
global sea level, marine δ
18
O, and ice sheet temperature records. Turney et al. (2017) modeled the impact of
WAIS meltwater from the Weddell and Ross Seas on the Southern Ocean during times when changes in
Atlantic overturning circulation where insufficient to explain smaller amplitude antiphased temperature
relationships between the Northern and Southern Hemispheres during the last glacial. The simulations of
Turney et al. (2017) showed how atmospheric teleconnections were driven by AIS discharge between 30
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and 28 ka (inferred from ice‐rafted detritus [IRD] layer in the South Atlantic south of the Polar Front;
Kanfoush et al., 2000) and the rapid ocean‐atmosphere feedbacks that followed may have contributed
high‐latitude temperature trends in the Northern Hemisphere.
5.2.3. Last Interglacial and Pleistocene Evidence for a Retreated Ice Margin
During the Last Interglacial period (LIG; 130–115 ka, the geological interval also referred to the Eemian),
globally distributed records of local sea levels indicate that GMSL likely stood 6–9 m higher and mean glo-
bal temperatures were 0.7° ± 0.6°C warmer than in preindustrial times, although atmospheric CO
2
con-
centrations were similar to preindustrial values (Dutton & Lambeck, 2012; Dutton, Webster,
et al., 2015; Hoffman et al., 2017; ;Kopp et al., 2009; McKay et al., 2011; Petit et al., 1999). Data from
the North Greenland Eemian (NEEM) ice core record revealed that the Greenland Ice Sheet's sea level
contribution to the LIG highstand was <2 m (NEEM Community Members, 2013). If this estimate is cor-
rect, then combined with a possible contribution of <1 m from thermal expansion of the ocean and melt-
ing of glaciers (McKay et al., 2011), this requires a contribution from the AIS. Even the lower bound (6 m)
of these sea level estimates implies ice mass loss equivalent to the marine‐based sectors of the WAIS
(4.9 m SLE), while the upper bounds (9 m) require some loss of the marine margins of the EAIS
(21.1 m; SLE estimate based on Bedmap2 data; Fretwell et al., 2013), such as the Aurora, Wilkes, and
Recovery subglacial basins.
In spite of the above circumstantial evidence pointing to a substantial contribution to the LIG sea level high-
stand from the marine‐based portions of West and East Antarctica, geological archives close to the Antarctic
margin have only provided limited and contradicting direct evidence for AIS retreat during the LIG. This is
due to both logistical expense and difficulty in working in Antarctica, and factors relating to sedimentation
rates, erosion and preservation, and accurate dating of sediments proximal to the margin. For example, gla-
ciomarine environments are commonly sediment starved on the Antarctic continental shelves (with sedi-
ment mass accumulation rates <10 cm kyr
−1
; Larter et al., 2012). Indirect yet notable, marine sediment
archives off the continental slope of the Adelie Land margin have been used in provenance studies, where
the mineralogical and geochemical characteristics of glacially eroded sediment transported to the
Antarctic margin provide information about past changes in AIS dynamics, as well as information about
the hidden geology beneath the ice (Figure 18) (Cook et al., 2017; Licht & Hemming, 2017). The
Pleistocene provenance record of these sediments suggests substantial retreat in the Wilkes Subglacial
Basin during the LIG, which could have contributed up to a few meters of global sea level rise (Wilson
et al., 2018).
The LIG sea level commitment is further complicated by recent, higher estimates for the Greenland contri-
bution: Yau et al. (2016) suggest a Greenland Ice Sheet contribution of 5.1 m (4.1–6.2 m, 95% probability
interval) toward the end of the Eemian, although the authors acknowledge several unexplained discrepan-
cies between their study and the NEEM ice core‐based reconstructions. Note also that a recent reevaluation
of GIA corrections suggests that LIG GMSL estimates (including the lower bound) may need upward adjust-
ment by ~2 m (Rohling et al., 2017). This is based on revised ice volume constraints for the preceding glacial
maximum that consider GIA corrections of a considerably different ice volume distribution (larger Eurasian
and smaller North America ice sheets) during the glacial period preceding the LIG, relative to the LGM. If
Yau et al. (2016) are right and the Greenland contribution to GMSL was ~5 m, then the ~2 m upward adjust-
ment of LIG GMSL estimates of Rohling et al. (2017) to 8–11 m above the present‐day level would require
Antarctic ice loss during the LIG of ~5 m, even after accounting for an ~1 m contribution from thermal ocean
expansion (Turney et al., 2020). Note that this argument does not include asynchroneity in the contributions
from Northern and Southern Hemisphere ice sheets, which suggest that the estimated ~5 m Antarctic con-
tribution is a low‐end estimate (Rohling et al., 2019).
Centennial‐scale rates of sea level rise during the LIG were quantified through statistical analysis of Red Sea
based sea level records and comparison with global coral/reef, and speleothem data (Rohling et al., 2019), as
discussed in section 5.2.5. That study also deconvolved the relative contributions from Greenland and
Antarctica by subtraction of the Greenland contributions based on two independent reconstructions from
the global signal. This assessment revealed an early AIS‐derived highstand between 129.5 to 125 ka of
5–10 m, the timing of which had previously been inferred qualitatively (Dutton, Webster, et al., 2015;
Marino et al., 2015; Yau et al., 2016). This may have occurred in response to a warm early‐LIG temperature
“overshoot”during the main Termination 2 deglaciation phase (Marino et al., 2015).
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Global mean temperature might not be the critical parameter or driving
process when considering Antarctica's contribution to sea level rise
(Marino et al., 2015). These authors indicated that the LIG highstand
likely occurred from millennial‐scale warming of Antarctica on the order
of 4°C, due to a bipolar see‐saw temperature event initiated by a massive
meltwater injection to the North Atlantic as the Northern Hemisphere
ice sheets retreated. This interval, known as Heinrich Stadial 11
(HS11:135–130 ka) was in fact the main deglaciation event of
Termination II when sea level rose by 70 m in 5,000 years at a rate of
28 m ± 8 m kyr
−1
(Grant et al., 2012, 2014) (Figure 19). The HS11 fresh-
water flux into the North Atlantic coincided with Northern Hemisphere
cooling and a typical see‐saw response of strong Antarctic warming of 9°
C (Jouzel et al., 2007), of which only 5°C can be apportioned to radiative
forcing (Masson‐Delmotte et al., 2010). The remaining warming is likely
to have resulted from the buildup of ocean heat, located in the global inter-
ior ocean north of the ACC rather than in the Southern Ocean itself (Pedro
et al., 2018), and eddy‐driven heat flux across the ACC to Antarctica. Proxy
evidence supports a reduction in sea ice extent (Chadwick et al., 2020;
Crosta et al., 2004; Esper & Gersonde, 2014) and AABW formation coinci-
dent with warmer surface waters, as well as inferred surface stratification
(Ninnemann et al., 1999) and freshwater input that was associated with
the rise in sea level between 135 and 143 ka (Figure 19). This HS11 climate
event was unlike the last glacial termination (Termination 1) when the
highest rate of sea level rise, MWP‐1a, occurred after, not during, the
abrupt North Atlantic cooling and associated reduction in AMOC of
Heinrich Stadial 1 (HS1: 16–17 ka; Lambeck et al., 2014).
Ice and environmental data from the Patriot Hills blue ice region land-
ward of the Weddell Sea Embayment provide evidence for significant ice
mass loss during the LIG (Turney et al., 2020). This study is the first direct
of substantial WAIS loss during the LIG, which is supported by regional
ice sheet modeling results that reinforce the notion of a centennial‐scale
(200 year) response time scale to 2°C of ocean warming relative to today.
Further afield in the Southern Ocean, there are additional indicators of
oceanographic change (Figure 19) associated with ocean stratification
and a reduction of AABW formation when the LIG sea level high sug-
gested an AIS meltwater contribution (see Rohling et al., 2019).
Warming in Southern Ocean sea surface temperature (SST) reconstruc-
tions (Hayes et al., 2014; Ninnemann et al., 1999) corresponds to a peak
in reconstructed air temperature over Antarctica (Petit et al., 1999) at
the time of early LIG sea level rise, when the Northern Hemisphere was
much cooler (Marino et al., 2015). The SST rise based on TEX
86L
shows
a maximum of +7°C during MIS 5e (Hayes et al., 2014), although
diatom‐derived surface temperature reconstructions suggest that this
value may have been closer to +5°C (Bianchi & Gersonde, 2002;
Chadwick et al., 2020; Etourneau et al., 2013; Schneider Mor et al., 2012).
Local atmospheric warming would tend to result in a contraction of sea
ice, as has been suggested by a reconstruction of sea ice extent during
MIS 5e (Wolff et al., 2006). Ice‐ocean simulations (e.g., Hellmer et al., 2012) have indicated how reductions
in sea ice can redirects warm CDW toward the coast and drives basal melting of ice shelves. Taken together,
these observations may be related to AIS change through meltwater release that results in water column
stratification and disrupts the formation of dense AABW, and sets up a positive feedback that traps ocean
heat in the subsurface, and drives further basal melting of the AIS and ice shelf breakup (e.g., Phipps
et al., 2016). Paleoceanographic data to support this mechanism of AIS loss include surface ocean
Figure 19. Paleoceanographic data consistent with reorganization of the
Southern Ocean climate system during the Last Interglacial period from
proxies records: (a) EDC Antarctic air temperature difference (gray dots and
blue line) (Jouzel et al., 2007), (b) sea surface temperature proxy TEX
86L
(Hayes et al., 2014), (c) Antarctic sea ice extent inferred from sea‐salt
sodium flux (ssNa) (orange) (Wolff et al., 2006), (d) AABW formation
inferred from authigenic uranium concentrations which reflect bottom
water oxygenation (black) (Hayes et al., 2014), (e–g) benthic δ
13
C
(Cibicidoides wuellerstorfi)records from Site 1089 (brown) (Hodell
et al., 2003), (f) Site 1063 (gray) (Deaney et al., 2017), and (g) MD03‐2664
(beige) (Galaasen et al., 2014), (h and i) water mass tracer Nd isotopes
measured in (h) fossil fish teeth (Deaney et al., 2017), and (i) sediment
leachate (Böhm et al., 2015), (j) the rate of the LIG sea level change and 95%
confidence limits (Marino et al., 2015), and (k) the relative sea level
curve with 95% confidence limits (Grant et al., 2014). Figure modified based
on (Rohling et al., 2019). HS11: Heinrich Stadial 11.
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stratification changes (differences between surface to deep foraminifera δ
18
O; Ninnemann et al., 1999);
reduced bottom water oxygenation in the deep Southern Ocean following the peak in Antarctic warming
due to reduced AABW formation (Hayes et al., 2014); δ
13
C records from benthic foraminifera, which docu-
ment the transition from low‐δ
13
C deep waters indicative of AABW, to high‐δ
13
C in bottom waters from
increased NADW north (Galaasen et al., 2014; Hodell et al., 2003); and neodymium isotopes in sediment lea-
chates and fossil fish teeth (Böhm et al., 2015; Deaney et al., 2017), that suggest an expansion of NADW at the
end of Heinrich 11 into the South Atlantic.
More paleoenvironmental studies close to the Antarctic margin are required to adequately constrain the
magnitude and pattern of AIS retreat during the LIG. The relatively indirect and far‐field inferences for
LIG AIS retreat (Figure 19) have yet to be supported by proximal records of stronger CDW upwelling and
advection onto the continental shelf. All of the marine sediment proxy data presented in Figure 19 derive
from north of the Southern Boundary of the ACC or ACC/Weddell Gyre Boundary, which forms an impor-
tant oceanographic front that helps to restrict CDW access to the Antarctic margin (e.g., see Vernet
et al., 2019). Furthermore, the southward expansion of NADW does not necessarily correspond to warmer
CDW reaching the Antarctic margin, and a reduction in AABW formation could instead result from a north-
ward shift and/or weakening of Southern Hemisphere westerly winds (Glasscock et al., 2020). Stratification
and freshening due to meltwater inputs south of the ACC might be explained sea ice changes, particularly in
the Atlantic relative to the Pacific sector of the Southern Ocean (Holloway et al., 2017). Reduced LIG sea ice
is consistent with the hypothesis of a weakened AMOC resulting in heat accumulation in the Southern
Hemisphere and subsequent sea ice reduction but without a major collapse of the WAIS (Holloway
et al., 2016). Targeted data collection close to the Antarctic margin is needed to resolve this issue.
There is a broader base of evidence for large‐scale retreat of the WAIS in other periods. For example,
Quaternary diatoms and elevated
10
Be concentrations in subglacial till samples recovered from beneath
the grounded WAIS indicate that they have been exposed to open marine conditions at least once during
the past 1.3 Ma, implying at least one significant grounding line retreat along the Siple Coast (Kerr, 1998;
Scherer et al., 1998). However, the advection of diatoms and adhering
10
Be beneath an ice shelf to the sites
studied by Scherer et al. (1998) may have occurred without complete collapse of the WAIS, due to the stabi-
lizing influence of isostatic rebound following the initial grounding line retreat (see Kingslake et al., 2018).
In the ANDRILL McMurdo Ice Shelf Project drill‐core AND‐1B, recovered beneath the northwestern margin
of the Ross Ice Shelf recovered as part of the ANDRILL project, deposition of diatomaceous sediments bear-
ing Pleistocene calcareous nanofossils was assigned to Marine Isotope Stage 31 (~1 Ma)—a well‐documented
warm period in circum‐Antarctic and the Southern Ocean (Villa et al., 2012). This provides evidence of the
last known time with a smaller‐than‐present ice shelf in the Ross Sea (McKay, Naish, Powell, et al., 2012;
Naish et al., 2009; Scherer et al., 2008; Villa et al., 2012). Deep‐sea cores from the Amundsen Sea are incon-
clusive regarding WAIS retreat during the past 1 Ma (Hillenbrand et al., 2002), but a depositional anomaly
hints at a reduced WAIS in this sector between 621 and 478 ka (Hillenbrand et al., 2009). Further data from
AND‐1B indicated that in the Ross Sea sector, the WAIS has oscillated in the Ross Sea sector between an ice
sheet and ice shelf state at least seven times since 0.8 Ma, and at least once in the past 250 ka, with a Ross Ice
Shelf that was similar to or less extensive than today during past warm periods of Marine isotope Stages
(MIS) 5 and 7 (McKay, Naish, Powell, et al., 2012). Reworked, marine sediments recovered in ice cores on
Ross Island also indicate that ice shelves formed during the LIG in the southern Ross Sea, but evidence that
the grounded ice retreated further south of Ross Island than today remains equivocal.
Marine Isotope Stage 11 is one of the longest interglacial periods of the Pleistocene (374–424 ka) (Lisiecki &
Raymo, 2005), with warmer global temperatures (by 1–2°C) and similar CO
2
concentrations relative to the
Holocene (Lüthi et al., 2008). Sea level estimates for MIS 11 have been debated, with a suggested highstand
of ~20 m higher than today (Hearty et al., 1999; McMurtry et al., 2007; Olson & Hearty, 2009; van Hengstum
et al., 2009), which would require collapse of the Greenland and WAIS, in addition to significant contribu-
tions from the EAIS. This initial estimate was revised downward between 6–13 m after taking into account
the influence of GIA (Raymo & Mitrovica, 2012), and refined further to 8–11.5 m (Chen et al., 2014). Raymo
and Mitrovica (2012) argued that these revised estimates would require meltwater contributions most of
Greenland and West Antarctica but without significant contribution from the EAIS. However, retreat of
the EAIS, specifically the Wilkes Subglacial Basin, is suggested for MIS 11 based on detrital sediment
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provenance in an offshore sediment core (Wilson et al., 2018) and further supported by the very low initial
δ
234
U measured in subglacial precipitates of calcite and opal (Blackburn et al., 2020). The low δ
234
U values
suggest that subglacial fluids recorded by the precipitates were not isolated for millions of years, as would be
expected based on EAIS stability (e.g., Sugden et al., 1993). Instead, these data reveal that the
234
U‐rich sub-
glacial fluids were renewed around 400 ka, by flooding in the low‐lying Wilkes Subglacial Basin, and
groundling line retreat (Blackburn et al., 2020). Paleoceanographic evidence to support this hypothesis is
currently limited. However, (Hodell et al., 2000) observed a strong export of NADW into the South
Atlantic Southern Ocean and upwelling of CDW with a greater proportion of NADW, while Glasscock
et al. (2020) observed an authigenic U peak (similar to that observed during MIS 5e) that could be interpreted
as a reduction in AABW formation during MIS 11 associated with meltwater forcing. Ocean warming and
basal melting of the marine‐based EAIS have been speculated as driving the EAIS retreat, without significant
climate (atmospheric) warming (Blackburn et al., 2020), but further proximal evidence is needed to fully
understand the processes driving MIS 11 AIS change.
5.2.4. Pliocene Evidence for a Retreated Ice Margin
Research on the Pliocene AIS is relevant due to the similarity between the current climate and conditions
of the early‐to‐middle Pliocene, 5.3–2.6 Ma, when atmospheric CO
2
concentrations were between 350 and
400 parts per million (ppm), global surface temperatures were 2–3°C higher and the Southern Ocean was
4–6°C warmer than present (Pagani et al., 2010; Seki et al., 2010; Martínez‐Botí et al., 2015; McKay, Naish,
Carter, et al., 2012). A range of globally distributed sea level records, alongside considerations of the
benthic oxygen isotope record, suggests that GMSL stood 10–30 m above present, which requires a marked
decline of the Greenland Ice Sheet and WAIS, as well as contributions from marine‐based regions of the
EAIS (Foster & Rohling, 2013; Gasson et al., 2016; Miller et al., 2012; Winnick & Caves, 2015)
(Figure 16). However, determining the exact extent and location of ice sheet loss, and thus eustatic sea level
change, from far‐field sea level records has significant caveats. Critically, this requires a robust understand-
ing of tectonic and mantle processes, and GIA that may result in a departure from eustasy in a sea level
record at any given location (Raymo et al., 2011). Therefore, Pliocene sea level records must be considered
in the context of model results that incorporate these processes, alongside more direct evidence of AIS
retreat from the Antarctic margin.
Pliocene sea level and geological evidence for elevated Pliocene shorelines are limited by a dynamic topogra-
phy overprint on the order of 10 m of vertical motion, which makes GIA corrections highly uncertain (see
section 4.1.2 on dynamic topography) (Austermann et al., 2015; Raymo et al., 2011; Rovere et al., 2014;
Rowley et al., 2013). This, superimposed on true sea level variability within the Pliocene, may explain
why global sea level reconstructions for this period span range from 6 to 20 m or even 30 m (Dumitru
et al., 2019; Foster & Rohling, 2013; Grant et al., 2019). A study that attempted to deconvolve the different
processes used phreatic overgrowth in Mallorcan cave deposits and concluded that sea level during the
early‐middle Pliocene reached between 6 and 27 m above present, at 68% confidence (mean values of 16
and 24 m; Dumitru et al., 2019).
There are currently insufficient geological constraints on the size and extent of the Pliocene AIS, resulting in
many synthesis studies on geological data and climate modeling (de Boer et al., 2015; Haywood et al., 2016,
2016). Susceptibility of the marine‐based WAIS to retreat, for example, during warm intervals of the
Pliocene, was suggested by modeling (Pollard & DeConto, 2009) and supported by the deposition of diato-
maceous oozes under surface water temperatures a few degrees warmer than today at the ANDRILL
AND‐1B site in the Ross Sea during Pliocene interglacial periods (McKay, Naish, Carter, et al., 2012;
Naish et al., 2009). Pliocene deglaciation and retreat of marine‐based EAIS portions during Pliocene warm
periods had previously been debated (e.g., see Barrett, 2013), based on observations of
deglaciation‐indicating diatoms in sediments of the Sirius Group (now Sirius Formation) in the TAM
(Webb et al., 1984). This contrasted with field evidence from the Dry Valleys (e.g., Sugden et al., 1995;
Sugden et al., 2017) that indicated no thinning and the presence of >8 Ma old ice, suggesting that the climate
did not warm enough for large‐scale EAIS deglaciation. However, Bertram et al. (2018) and Cook et al. (2013)
provided additional evidence for the (at least partial) collapse of marine‐based EAIS portions during
Pliocene warm periods (next paragraph). Subsequent ice sheet modeling work used alternative ice sheet
model physics (Pollard et al., 2015) and suggested significant EAIS retreat and wind‐blown emplacement
of Pliocene marine diatoms in the TAMs (Scherer et al., 2016).
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Observational evidence for retreat into the marine‐based portions of the EAIS includes geochemical proxy
data in marine sediment records offshore Adelie Land that suggest significant ice retreat inland of the
Wilkes Subglacial Basin, one of the largest subglacial basins of the EAIS, which may have contributed
3–4 m to Pliocene sea levels (Bertram et al., 2018; Cook et al., 2013; Patterson et al., 2014; Valletta et al., 2018)
(Figure 20). Quantification of the subglacial sediment thickness of the northern Wilkes Subglacial Basin by
airborne gravity and geophysical data reveals thin sediment deposits influenced by glacial erosion (Frederick
et al., 2016; Paxman et al., 2018) that may suggest more frequent Pliocene to Pleistocene ice sheet retreat and
advance. Uplifted glacimarine sediments from four separate formations of the Pagodroma Group on the
flanks of the Prince Charles Mountains adjacent to Amery Ice Shelf suggest a reduced glacial extent of the
AIS margin in the Prydz Bay region at some time between the early Miocene (or older) and the Pliocene
or early Pleistocene (Hambrey & McKelvey, 2000; Whitehead et al., 2004).
The Pliocene retreat patterns reconstructed from geochemical sediment provenance (e.g., Cook et al., 2013)
were not reproduced by the concurrent generation of ice sheet models around that time. Since then, new
Figure 20. Reconstructing past AIS dynamics based on sediment archives. The conceptual model shows an Antarctic glacier (a) advanced to the shelf edge over
two different bedrock lithologies, (b) with an ice shelf, similar to the present day, and (c) retreating under warm conditions, for example, the Pliocene. The
provenance of ice‐rafted detritus (IRD) found at a deep‐water site (d) changes with the location of the grounding line for example, in panel (c), proportionately
more IRD is sourced from inland bedrock (a) due to increased basal erosion and entrainment of sediment, which is transported into the Southern Ocean via ice
bergs (figure based on Ehrmann & Grobe, 1991; Hambrey & McKelvey, 2000).
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models have been developed, both including MISI and MICI (DeConto & Pollard, 2016; Pollard et al., 2015)
and excluding MICI (Golledge, Thomas, et al., 2017), which simulate ice sheet retreat in better agreement
with observed Pliocene highstand sea levels.
Pollard et al. (2017) show the effect of different Earth viscoelastic properties on the varying Antarctic contri-
bution to Pliocene sea level estimates. Following on from the sea level feedback model of Gomez et al. (2015),
Gomez et al. (2018) and Pollard et al. (2017) show that estimates of the AIS contribution to the sea level high-
stand during the Pliocene are sensitive to lateral changes in the viscosity and lithospheric thickness of the
Earth beneath the ice sheet. Pliocene ice retreat from Antarctica is limited to 9 m of equivalent sea level rise
using an Earth profile of a weak upper mantle and thin lithosphere; this compares to 15 m in Earth profiles
with slower viscous bedrock rebound (Pollard et al., 2017). The properties of the solid Earth beneath the ice
sheet have been shown to be more important in the Pliocene compared to the solid Earth response to ice
volume changes during the last deglaciation (Gomez et al., 2018) (also see section 4.2.1). These results high-
light the need for more detailed resolution of the Earth's structure in regions of the ice sheet that are cur-
rently experiencing change and/or are susceptible to future change.
Despite the large uncertainty in Pliocene sea level reconstructions and ice volume, Pliocene records have
been used to tune ice sheet models projecting future sea level contributions from Antarctica (see
section 4.1). Results from the Pliocene Ice Sheet Modeling Intercomparison project revealed the high depen-
dency of Pliocene AIS configuration and sea level contribution based on variable climatology and ice sheet
models (Dolan et al., 2018) (Table 2). Recent Pliocene ice sheet simulations suggest an Antarctic contribution
of 3 to 12 m (Gasson et al., 2016); 7.8 to 11.4 m (Golledge, Thomas, et al., 2017), and 11.3 m (DeConto &
Pollard, 2016); all these values imply that the AIS has significant sensitivity under warmer than present
climates.
Ice sheet models that incorporate hydrofracturing of ice shelves and ice cliff collapse processes (see
section 3.1.3) can simulate significant retreat in the East Antarctic Aurora Basin during the Pliocene
(DeConto & Pollard, 2016). In contrast, seismic data suggest minimal retreat of the Totten Glacier by no
more than 150 km into the Sabrina subglacial basin, although these data are equivocal with regards to abso-
lute timing of the last retreat of the margin deep into this basin (Aitken et al., 2016; Gulick et al., 2017). More
robust data of past retreat or stability of this margin are required in order to determine the climate thresholds
or physical processes that govern retreat in this region. Currently, the geological evidence for WAIS and
EAIS retreat during past warm periods of the Pliocene can only provide a qualitative estimate of AIS contri-
bution to higher sea levels during warm periods in the geological past. Direct evidence of glacial retreat com-
bined with geophysical modeling of Antarctica could help to confirm the extent of deglaciation beneath
now‐grounded portions of the AIS, to fully quantify the AIS contribution to sea level rise during warm cli-
mate intervals (McKay, Barrett, et al., 2016).
5.2.5. Paleoperspective on Current Rates of Sea Level Change
For Late Pliocene interglacials, mean rates of sea level rise were reconstructed at typical values of 0.3–0.5 m
per century, with four interglacials that exceeded 0.5 m per century (Grant et al., 2019). Including hydrofrac-
ture and cliff‐collapse mechanisms, Pollard et al. (2015) simulated rates of Antarctic mass loss under a warm
Pliocene‐like climate of approximately 1 m per century SLE. Alternative simulations that did not include
these particular processes yielded rates of 0.45–0.65 m per century (Golledge, Thomas, et al., 2017).
Last interglacial sea level records provide a basis for understanding the main ocean‐climate drivers behind
sea level changes within a more recent warm period, including their rates of change. Likewise, sea level
records through deglaciations may help in identifying the main driving processes behind very fast rates of
sea level change and their relationship to changes in the radiative forcing of climate and interhemispheric
heat redistribution within the ocean‐climate system. Although past changes arose from different (preanthro-
pogenic) rates of climate forcing than those observed today, the identification of critical processes in the past
may still be relevant to the future. For example, the current incursion of warm CDW onto the Antarctic
shelves (section 5.2.2) bears similarity to the consequences of bipolar temperature see‐saw events of the past,
when the ice sheet‐climate perturbation response time was ~500–700 years (Grant et al., 2012, 2014). This
similarity emphasizes the urgency of including encroachment of warmer water onto the Antarctic shelves
and detailed processes of under‐ice melting, in models for simulating the Antarctic response to future cli-
mate change.
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Sea level fluctuations during the LIG and the timing and magnitude of ice sheet retreat have been widely
debated (Bruggemann et al., 2004; Dutton, Webster, et al., 2015; Kopp et al., 2009; O'Leary et al., 2013;
Orszag‐Sperber et al., 2001; Rohling et al., 2008, 2009; W. G. Thompson et al., 2011). The lack of consensus
on the timing, magnitude of the peak, and variability of sea level (stable, two prominent peaks, several oscil-
lations) relates to difficulties in establishing consistent chronologies and correcting for the effects of tectoni-
cally driven vertical movements and GIA (see section 4.2.2) on paleoshorelines or sill depths. However, the
fossil coral and coralline algae based sea level record from the Seychelles (Dutton, Webster, et al., 2015) seem
to have little sensitivity to different GIA reconstructions that vary according to Earth models and ice sheet
configurations used in the GIA models (see, e.g., the results over a wide range of Earth models and
MIS6‐MIS5e ice histories in Rohling et al., 2017). But a lack of (near)continuity in the Seychelles record
means that inferred rates of sea level change are low‐end estimates.
Other work has used the highly resolved Red Sea sea level records, where relative chronology is tightly con-
strained by stratigraphy, in comparison with coral and speleothem based records (Grant et al., 2012, 2014;
Rohling et al., 2019). A focused study for the LIG suggests that mean rates of sea level rise may have reached
2.3–2.8 m per century (0.9–3.7 m per century at 2σ) (Rohling et al., 2019). However, 122 sea level rise events
in the Red Sea records observed throughout the past 500,000 years, including other interglacial times when
total global ice volume was similar to the present, revealed rates of sea level rise that were more commonly
characterized by rates of up to 1.4 m per century (2σ) (Grant et al., 2012, 2014). These rates were driven by
relatively slow natural climate forcing and feedback processes, and they likely form low‐end estimates of the
rates that may be expected due to the current fast anthropogenic climate forcing. Fitting of theoretical
growth functions for rates of sea level rise, using broad probability distributions for ice volume (sea level)
response times and rates of change obtained from the Red Sea results, suggests potential sea level rise by
2100 to about +1.8 m above the Year 2000 level (upper 95% probability bound) (Rohling et al., 2013).
Again, this considers only naturally precedented rates of change, and addition of historically unprecedented
(new) ice‐dynamical processes may cause exceedance of that value. One exception is the initial fast sea level
rise of the LIG, which may be a case where the ice response was trying to “catch up”with abrupt (regional)
temperature forcing, and hence may provide an example where an ice sheet was responding to a disequili-
brium with climate, not unlike today. This event might therefore portray a potential Antarctic melt scenario
for the future, but it is predicated on large‐scale and long‐lasting invasion of warm waters onto the conti-
nent's shelf areas. Meltwater Pulse 1A may (partially) represent another such event, albeit much smaller
in amplitude, which hinders signal‐to‐noise distinction.
6. Potential Consequences of AIS Melt
Global sea level change is a well‐known consequence of climate change, due to the thermal expansion of
ocean water, and glacier and ice sheet melt. The evidence suggests that most of the AIS loss will come from
the WAIS in the coming century (Shepherd et al., 2019), although the rate and concentration of future green-
house gas emissions will dictate the exact rate of sea level rise and also the wider AIS contributions. Aside
from uncertainty around future emissions, an increasing number of studies has shown a spread in estimates
of the Antarctic sea level contribution to the Year 2100 (Table 2), based on differences in how ice dynamics
are modeled. Another complexity associated with sea level predictions is that the spatial pattern of regional
sea level change is heterogeneous due to the impacts of Earth's gravitation, rotation, and deformation as a
consequence of changes in ice mass. Improvements have been made in assessing the regional sea level
impact or fingerprint of AIS melt, which have been aided by more accurate considerations of Earth's rheol-
ogy beneath the AIS. While most of this paper deals with how climate and solid Earth processes affect the
AIS, the following section also details how AIS sensitivity will in turn affect the climate system. In particular,
this concerns the consequences of increased fresh water supplied to the Southern Ocean in response to
increased ice melt, although we do not consider the impact on ecosystems.
6.1. Future Global Sea Level Change
6.1.1. Projected Sea Level Change
GMSL is projected in IPCC AR5 to rise through the 21st century by 0.26 (RCP2.6) to 0.98 m (RCP8.5), primar-
ily in response to expansion of the warming oceans, and contributions from melting glaciers and ice sheets of
Antarctica and Greenland (Church et al., 2013). The largest uncertainty in these estimates comes from
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quantifying the Antarctic contribution to future sea level rise due to historically unprecedentedly rapid
dynamic ice sheet mass loss (see section 2.1; Oppenheimer et al., 2019). AR5 concluded that collapse of
marine‐based sectors of the AIS is unlikely (<33% probability), but if initiated could add “up to several tenths
of a meter”to the range predicted for 2100. Recent work suggests the possibility of an even larger AIS con-
tribution to global sea level rise of ~1 m by 2100 (DeConto & Pollard, 2016; Pattyn et al., 2017). Goodwin
et al. (2017) considered a very large ensemble of model projections of sea level change from an analytical
model that show a high degree of congruence with CMIP5‐based IPCC AR5 simulations, with a subset cho-
sen based on additional consistency with time series of historical observation. They found a higher ice‐melt
sensitivity than previously recognized, with the consequence that ensemble mean sea level rise would be
13–16 cm higher than projected in IPCC AR5, even without additional consideration of ice dynamics.
Alternative methodologies to those used in AR5 have since been developed to provide an assessment of
events beyond 2100. Slangen et al. (2017) highlighted the shift in sea level projection methodology toward
a probabilistic approach using skewed uncertainty distributions to provide better assessments of low prob-
ability but high risk events in the future. Kopp et al. (2014) presented the first probability distribution for
local sea level change, which can differ significantly from GMSL rise due to nonlinear changes in ocean
dynamics, variations in Earth's gravitational field and crustal height, GIA and local effects of tectonics,
and groundwater/hydrocarbon withdrawal and sedimentation. Sensitivity analyses showed that the nature
of ice sheet mass changes had the largest effect on global and local sea level predictions. Kopp et al. (2014)
highlighted the need for improved ice sheet models that do not depend on expert elicitation, which requires
better constraints on the magnitude of positive and negative feedbacks on ice loss. Since this work, advances
have been made in understanding static‐equilibrium sea level and grounding line retreat processes, and the
sea level feedback (Gomez et al., 2015), as well as the inclusion of new processes in ice sheet models.
Several studies have highlighted the major scientific challenges (e.g., Kennicutt et al., 2019, 2014, 2015, 2016)
of accurately predicting the AIS response, and the consequent global sea level rise, to current and future
ocean atmosphere‐ocean warming (DeConto & Pollard, 2016; Edwards et al., 2019; Golledge et al., 2019,
2015; Ritz et al., 2015). Of these, all except the DeConto and Pollard (2016) study suggest a modest AIS con-
tribution by 2100 within the range of the AR5 suggestion for high temperature forcing—less than 0.30 m
(with 95% confidence; Ritz et al., 2015) or 0.39 m (upper bound; Golledge et al., 2015), with the latter predict-
ing as little as 0.01 m by 2100 in low or moderate emission scenarios. In contrast, the results of DeConto and
Pollard (2016) predicted an Antarctic contribution to GMSL rise of up to 1.05 cm (1σ= 0.30 m) by 2100 using
RCP8.5 (see Table 2). The key difference between that study and the other studies is the inclusion of two MICI
processes not previously considered at the continental ice sheet scale: hydrofracturing of ice due to surface
melt ponding, and ice cliff failure (cf. section 3.1.3). When ice sheet models include these processes, the
sea level contribution from AIS melt to 2100 may be closer to 0.45 m under an RCP 8.5 scenario (Golledge
et al., 2019) versus 0.15 m using traditional ice sheet modeling approaches without MICI (similar to the
AR5 predictions; Golledge et al., 2019). Pliocene sea level estimates are not sufficiently well quantified to
be of use in constraining model projections of the AIS contribution to sea level rise; hence, the most recent
sea level projections are calibrated using the satellite data (The IMBIE Team et al., 2018). The incorporation
of ice‐ocean‐atmosphere feedbacks into simulations of the future AIS response suggests that the process of
hydrofracturing by surface melting may not be important for sea level predictions to 2100 (Golledge, 2020;
Golledge et al., 2019). However, the latest expert judgment assessment of Bamber et al. (2019) highlighted
the increased uncertainty in sea level projections in response to uncertainty in ice processes and feedbacks
and high scenario estimates that are significantly higher than AR5 with a combined contribution of 0.57 m
from WAIS and EAIS (see Table 2). Importantly, all studies show that post‐2100 Antarctica will continue
to lose mass, even in the absence of continued forcing, resulting in many meters of eventual sea level
contribution.
Incorporation of hydrofracturing and ice cliff failure processes into a probabilistic framework for sea level
projections by Kopp et al. (2017) resulted in GMSL rise of 80–150 cm by 2100 for RCP 8.5. Using this
approach, there was little correlation between the current rate of ice mass loss and the contribution to
GMSL by 2100, highlighting the need to better understand the processes driving ice sheet melting. Three
key, but highly uncertain model parameters that are limited by modern analogues and observations include
(1) the rate of sub‐ice shelf melt in response to ocean warming, (2) the sensitivity of crevasse penetration to
meltwater input, and (3) the rate of ice cliff collapse. DeConto and Pollard (2016) and Kopp et al. (2017),
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applied a temperature correction of 3°C to ocean temperatures to 400 m depth, in order to match the
observed sub‐ice shelf melt rates (Mouginot et al., 2014; Rignot et al., 2013). Improvements are required in
these ice sheet models to achieve melt rates that are driven by more realistic ocean forcing, so that regional
melting better matches the current observations. These can be used to understand the downstream feed-
backs associated with freshwater input to the oceans (e.g., Silvano et al., 2018). It is also crucial to develop
deep‐water temperature reconstructions on the AIS margin during past times of sea level highstands (LIG
and Pliocene), in order to test and improve models and their results.
By limiting sea level projections to 2100, the commitment to sea level rise into the future as a consequence of
ocean and atmospheric heating and ice sheet instabilities initiated during the 21st century, is not fully rea-
lized (Bamber et al., 2019; Church et al., 2013; Golledge et al., 2015; Oppenheimer et al., 2019). Projections
beyond 2100 using IPCC RCP‐based warming scenarios highlight both the long‐term lag in ice sheet/ice
shelf responses to current perturbations in the climate and the importance of ocean‐forced warming (e.g.,
Cornford et al., 2015; Golledge et al., 2015; Winkelmann et al., 2015). Anthropogenic warming exceeding
1.5–2°C above present was sufficient in these models to produce a collapse of the major Antarctic ice shelves
within 100–300 years, with a subsequent long‐term commitment to sea level rise due to the centennial to
millennial scale response of the AIS.
6.1.2. Patterns of Sea Level Fingerprints From Antarctica
The sea level changes associated with variations in grounded ice cover are spatially heterogeneous due to
effects associated with gravity, Earth rotation, and viscoelastic deformation of the solid Earth (Clark &
Lingle, 1977; Farrell & Clark, 1976; Mitrovica et al., 2001). The spatial pattern of sea level change is depen-
dent on the source of the ice melt, and hence, these patterns are therefore called “sea level fingerprints.”In
the vicinity of ice loss, sea level falls due to (1) a drawdown of the sea surface associated with changes to the
gravitational field and (2) an uplift of the solid Earth beneath the missing mass (see Figure 13). Conversely,
in the far field away from the ice margin, sea level rises more than the GMSL change associated with the ice
loss, due to first‐order gravitational effects. The sea level fall in the near field can be an order of magnitude,
or more, greater than the sea level rise in the far field (Mitrovica et al., 2001). Earth deformation associated
with adding water to the oceans and a shift of the Earth's rotation axis toward the missing ice mass impart
second order effects to increase spatial and temporal variability of the pattern globally. The modeling and
physics of sea level fingerprints is relatively well understood and reviewed in detail in for example
Mitrovica et al. (2011) and Vermeersen and Schotman (2009).
Several recent studies have considered the sea level fingerprint that would follow collapse of sectors of the
AIS in a warming climate (e.g., Bamber et al., 2009; Gomez, Mitrovia, Tamisiea, et al., 2010; Hay et al., 2017;
Mitrovica et al., 2009; Mitrovica et al., 2011). Most of these studies have focused on the sea level change fol-
lowing a rapid retreat of the WAIS over time scales of hundreds of years or less, for which the solid Earth
deformation has (until recently, see below) been considered purely elastic. Far‐field peaks in the sea level
fingerprint of 1.3 times the global average sea level change are predicted along the U.S. East Coast and in
the Indian Ocean, and a sea level fall is predicted in the near field of WAIS reaching up to 15 times that
of the global average sea level change (Gomez, Mitrovia, Tamisiea, et al., 2010; Mitrovica et al., 2009).
Peaks from melting of the marine‐based sectors of the EAIS are shifted relative to the fingerprint for the
WAIS due in large part to Earth rotational effects and occur in the North Pacific and South Atlantic oceans
(Gomez, Mitrovia, Tamisiea, et al., 2010). Bamber et al. (2009) found lower‐amplitude peaks and minima in
the WAIS fingerprint pattern than Mitrovica et al. (2009), but Mitrovica et al. (2011) later showed that these
differences arose from the adoption of an elastically incompressible Earth model used in Bamber
et al. (2009). In contrast, Mitrovica et al. (2011) and other studies (Golledge et al., 2019; Gomez, Mitrovia,
Tamisiea, et al, 2010; Hay et al., 2017) employed a more realistic, elastically compressible Earth model
and treatment of the inundation of water into marine sections of the WAIS as they were freed of ice.
Sea level fingerprints for melting from different AIS sectors may be combined to estimate the total contribu-
tion from ice loss to sea level change globally. However, until recently, modern sea level records (tide gauges
and satellite records) have been too short or spatially sparse to distinguish the signal from individual ice
sheets over and above the natural variability of the oceans (Hay et al., 2015). The first fully global data assess-
ment of sea level fingerprint contributions to sea level change for 2002–2014 was published by Hsu and
Velicogna (2017), based on GRACE data and ocean pressure measurements from Argo floats. This study
used an updated time series of mass loss in Antarctica, which included mass loss from the WAIS and the
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Antarctic Peninsula Ice Sheet as well as the Aurora basin sector of the EAIS, and mass gain in Queen Maud
Land (Velicogna et al., 2014). The assessment for the period 2002–2014 showed, for example, an Antarctic
sea level fingerprint along the North American coastline that is larger than those presented in previous stu-
dies that considered collapse of the WAIS (e.g., Mitrovica et al., 2009). Examining tide gauges from the
Antarctic Peninsula and West Antarctica, Galassi and Spada (2017) suggested that they could observe the
first evidence of detectable sea level fingerprints from the current melting of Antarctica. Along with these
recent data‐analysis advances, new sea level modeling tools have been developed that apply inverse model-
ing to determine the sensitivity of site‐specific sea level changes to evolving ice mass loss geometries (Larour
et al., 2017; Mitrovica et al., 2017). These new tools along with improvements in sea level and ice sheet mass
change data sets hold a strong potential for future work using sea level fingerprinting to constrain the sea
level hazards associated with regional ice mass changes around Antarctica in the future.
Fingerprint studies have generally assumed that viscous Earth deformation is negligible on time scales of a
few centuries and so have adopted a purely elastic Earth model in their calculations. However, seismic tomo-
graphy studies and geological evidence suggest that upper mantle viscosities across the Antarctic are highly
laterally variable and may be several orders of magnitude lower than the average (see section 4.2.1). Hay
et al. (2017) have adopted a realistic, 3‐D Earth viscosity model that includes variability in lithospheric thick-
ness and mantle viscosity and a more realistic ice loss geometry to reassess the fingerprint of WAIS collapse
and its contribution to GNSS/GPS and gravity data. This and other studies (e.g., Barletta, et al., 2018; Powell
et al., 2019) suggest that viscous effects may in fact be nonnegligible on decadal to centennial time scales in
West Antarctica. For example, Hay et al. (2017) assessed a 100 year collapse event, including lateral variabil-
ity in Earth structure. They found that this leads to a slightly reduced peak sea level rise predicted in the far
field and a sea level fall in the vicinity of the WAIS that is up to 50% larger than the aforementioned estimates
from simple 1‐D elastic Earth models. The sea level fall continued to increase for ice sheet retreat scenarios
over longer time scales. The difference between 1‐D and more realistic 3‐D Earth viscosity models on the AIS
contribution to global sea levels (i.e., the pattern of global sea level change) highlights the need for further
work to define the viscosity profile beneath West Antarctica (e.g., in the Amundsen Sea Embayment,
Barletta et al., 2018).
6.2. Ocean State and Circulation
Increased melting of the AIS will have implications for the Southern Ocean and in turn for the wider global
climate through changes in the MOC. Observations around Antarctica are beginning to show that increased
basal melt can significantly change the local stratification, with potential implications for both the local cir-
culation and water mass formation. Freshwater input from melting ice shelves (Jacobs et al., 2011; Kim
et al., 2016; Loose et al., 2009; Nakayama et al., 2013; Naveira Garabato et al., 2017; Silvano et al., 2017)
can partly reduce or completely inhibit sea ice‐driven convection (Petty et al., 2013; Williams et al., 2016;
Silvano et al., 2018). This prevents the formation of dense (cold and/or salty) shelf water (DSW) masses,
which form the precursors for AABW, and thus an important component of the MOC.
The processes governing AABW formation are complex and include the formation and export of sea ice
linked to wind strength and pattern (Santoso & England, 2008), glacial meltwater input from the AIS
and sea level (Paillard & Parrenin, 2004), and the creation of polynyas close to the coast by winds
(Tamura et al., 2016). Modeling studies show that the role of AABW in controlling heat transport across
the shelf varies between models, depending on which of the aforementioned processes are present and
how they interact.
For example, coarse‐resolution models suggest that meltwater addition in the near the surface close to the
Antarctic coast reduces surface density and thus the potential for AABW formation (Fogwill et al., 2015;
Menviel et al., 2010; Phipps et al., 2016). As a result, the rate of AABW formation can be reduced by as much
as ~50% (Fogwill et al., 2015). Weakened AABW formation decreases the oceanic meridional heat transport
to high southern latitudes and leads to lower SSTs over the Southern Ocean and a larger sea ice extent
(Menviel et al., 2015). Moreover, the increased surface stratification and reduced convection were found
to lead to a subsurface warming. This warming occurs below the mixed layer south of 60°S and reaches
an amplitude of 0.5°C to 1°C (Fogwill et al., 2015; Menviel et al., 2010; Phipps et al., 2016; Swingedouw
et al., 2009). The strongest warming can occur adjacent to sectors of the WAIS (e.g., along the Antarctic
Peninsula, Amundsen Sea, Marie Byrd Land, and Ross Sea) that are particularly vulnerable to grounding
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line retreat and MISI (Fogwill et al., 2015). A high‐resolution modeling study suggests that the magnitude of
the subsurface warming could exceed 3.5°C and 2.5°C along the coast in the Ross and Weddell Seas,
respectively (Bronselaer et al., 2018). If this warming extends to other areas of Antarctic continental shelf,
then it could lead to thermal erosion of the WAIS and sectors the EAIS (e.g., Recovery, Aurora, Wilkes
Basins, and Marie Byrd Land; Fogwill et al., 2015; Golledge, Levy, et al., 2017) and induce a positive feed-
back, accelerating disintegration of sectors of the ice sheet that are highly sensitive to ocean temperature
changes (Golledge et al., 2014; Menviel et al., 2010). Moreover, given that coastal currents around the
Antarctic continent can carry fresher surface waters westward (Phipps et al., 2016; Rignot &
Jacobs, 2002), initial melting of a single sector of the AIS could impact upon the wider ice sheet.
Under this scenario, meltwater‐induced changes in the Southern Ocean could act as a strong positive
feedback mechanism, amplifying the response to climate forcings. However, this theory has not yet been
confirmed by observations.
The process of DSW suppression by meltwater is well understood (e.g., Lago & England, 2019; Snow, Sloyan,
et al., 2016). However, it remains poorly represented in coarse resolution models, in which dense waters are
primarily formed through open ocean convection outside the shelves. The effect of meltwater is likely to be
amplified in coastal polynyas, relative to open ocean convection sites, as polynyas in close proximity to the
meltwater sources are particularly sensitive to freshwater input (Silvano et al., 2018; Williams et al., 2016).
Models suggest that the strength of the deep MOC cell, which involves AABW, decreases under warming
scenarios. This is due to both a reduction in sea ice formation associated with the change in heat flux
(Cougnon et al., 2013; Kusahara & Hasumi, 2013; Morrison et al., 2015; St‐Laurent et al., 2015) and the
increased stratification associated with enhanced glacial meltwater (Pauling et al., 2016). Models without
an ice shelf component underestimate the response of the deep MOC cell in a warming climate (Kusahara
& Hasumi, 2013). Oceanographic data collected along the Sabrina Coast and in the Amundsen Sea, com-
bined with a mixed‐layer ocean model, also demonstrate that freshwater input from the basal melt of ice
shelves can inhibit the formation of DSW and allow CDW to reach ice shelf cavities (Silvano et al., 2018).
At a regional scale, an increase in freshwater flux due to basal melting is likely to enhance heat exchange
across different sectors of Antarctic shelves due to a more vigorous buoyancy‐driven coastal current
(Nakayama et al., 2014).
AIS meltwater input to the surface of the Southern Ocean has been shown in models to result in cooling of
the stratified surface ocean, sea ice expansion, and warming of the subsurface water (500–1,500 m) (Menviel
et al., 2010; Richardson et al., 2005; Stouffer et al., 2007), which reduces the formation of AABW in the areas
where it forms today. This feedback mechanism was demonstrated to encourage further melting at the base
of the ice shelves (Golledge et al., 2014) and has since been supported by other modeling studies (e.g.,
Bronselaer et al., 2018; Fogwill et al., 2015; Phipps et al., 2016). Note, however, that this is different from
the mechanism proposed for the inferred AIS sea level contribution to the LIG highstand, where proxy
records show a decrease (rather than increase) in sea ice and AABW formation following Southern
Hemisphere warming during HS11 (see section 5.2.3). We tentatively suggest that a possible explanation
for the difference between LIG proxy data and modeling results may be the unprecedented scale and dura-
tion of the LIG changes, which related to an extremely pronounced bipolar temperature see‐saw effect that
lasted some 5,000 years. Initial AIS retreat may have driven sea ice expansion as models suggest, but main-
tained Southern Ocean warming over 5,000 years may have overwhelmed the initial response, resulting in
an overall sea ice retreat as suggested by the proxy data. This hypothesis will require testing with runs of
high‐resolution models over >5,000 years. The LIG response to climate warming may be more similar to
the future anthropogenic climate change scenarios, where reduced sea ice would be the result of warming
of the climate system. Under such scenarios, future freshwater input into the Southern Ocean due to melting
of the AIS would be expected to result in stratification of the ocean, a reduction in AABW formation and
warming of the ocean at depth (Phipps et al., 2016).
Models also suggest that AABW reduction could also have significant far‐field effects, such as a reduction in
the formation of NADW within half a century of a meltwater pulse (Swingedouw et al., 2009). Further work
has suggested that meltwater input into the Southern Ocean can have global‐scale impacts upon the climate
system, including an increase in Southern Hemisphere sea ice area, which increases global albedo; cooling of
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the atmosphere throughout the Southern Hemisphere and much of the Northern Hemisphere; and a north-
ward shift in the position of the Intertropical Convergence Zone (Bronselaer et al., 2018).
It has also been suggested that freshwater may not enter the ocean as a simple meltwater pulse but instead as
an armada of melting icebergs in the Southern Ocean. Such an iceberg scenario may enhance the formation
of sea ice and models suggest a minor increase in open ocean deep water formation (Jongma et al., 2009).
Modern observations suggest that the most likely future scenario consists of a combination of both large ice-
berg calving, as seen with the Larsen ice shelves (Hogg & Gudmundsson, 2017; Rack & Rott, 2004; Rott
et al., 1996) and significant meltwater input in areas like the Amundsen and Bellingshausen Sea where cur-
rently no AABW is formed (Rignot et al., 2013). Thus, it is likely that the formation of AABW may decrease
in some locations but may be less affected in other regions around the Antarctic margin. An underestimated
and/or unknown component relating to freshwater loss from the AIS concerns subglacial‐fluvial transport
from deep inland under Antarctica (e.g., from Recovery Lakes, Vostok, and Prince Elizabeth Land) (Le
Brocq et al., 2013).
Increased melting of the AIS may furthermore have biogeochemical implications relating to oceanic carbon
storage and cycling. A 2010–2013 satellite survey of giant icebergs in the Southern Ocean observed increased
chlorophyll concentrations in their wake, which was inferred to result in an increased carbon export, as a
result of iron fertilization (Duprat et al., 2016). Increased iron supply from melting of the AIS (Death
et al., 2014; Gerringa et al., 2012) may act to fertilize the Southern Ocean and increase productivity due to
the supply of iron‐rich glacial meltwater. In the biogeochemical model of Death et al. (2014), Southern
Ocean productivity was increased by 40% in runs that included iron release from the AIS, relative to those
without such an iron source.
Experiments using the intermediate complexity LOVECLIM model suggest that increased AIS meltwater
input leads to an expansion of sea ice, an intensification of the Southern Hemisphere westerly winds,
and a reduction of biological productivity in the Southern Ocean south of 40
o
S (Menviel et al., 2010).
Yet this model also simulates an increased subduction of Subantarctic Mode Water, leading to greater
export of nutrients to the equator. This drives an increase in tropical productivity, thus resulting in no
overall change in the global biological export. Note that the results of this simulation depend on the
strength and position of the Southern Hemisphere westerly winds and that changes in iron fertilization
were not taken into account. While an expansion of sea ice may act to reduce Southern Ocean productivity
by decreasing light availability, enhanced productivity in the high‐latitude Southern Ocean has been sug-
gested to play a significant role in lowering atmospheric CO
2
through sea ice feedbacks (Fogwill
et al., 2020), during a period of cooling known as the Antarctic Cold Reversal (14.6–12.7 ka), which inter-
rupted the climate warming of the last deglaciation (Pedro et al., 2016). High seasonal variability in sea ice
extent during this period are suggested to enhance productivity by increasing nutrient availability through
deepening of the mixed layer, and iron supply during sea ice melt, and through the absorption of CO
2
within the melting sea ice (Delille et al., 2014)
Together, the examples from paleoclimate records and mechanistic insights from coupled climate models
highlight the existence of strong links between changes in the AIS and global ocean circulation, and hint
at important implications of a dynamic AIS. They also highlight the uncertainty around the climate feed-
backs resulting from anthropogenically induced melting of the AIS. Ocean freshening associated with glacial
meltwater input in model simulations shows positive feedbacks regarding both the expansion of sea ice and
subsurface warming (e.g., Fogwill et al., 2015; Phipps et al., 2016; Menviel et al., 2010) and reduction of sea
ice and subsurface warming (Bronselaer et al., 2018), as well as a number of negative feedbacks associated
increased Southern Ocean carbon storage through sea ice feedbacks (Fogwill et al., 2020), the reduction of
MISI due to sea level lowering in the region of AIS mass loss, and the stabilizing effect of rapid bedrock uplift
in regions of West Antarctica (see section 4.2.2). This demonstrates a compelling need to conduct further
research into ice‐ocean and solid Earth feedbacks and incorporate such feedbacks into future climate projec-
tions through the development of coupled climate‐ice sheet models. In doing so, due care is needed in dis-
tinguishing impacts over different time scales and for different amplitudes of ice melt and warming.
While sea ice responses, for example, might lean initially toward expansion, they might reverse toward
reduction during long‐term sustained, large‐amplitude warming.
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7. Summary and Future Research Priorities
The societal impact of sea level rise calls for an understanding of how the AIS will evolve in a warming
world, in terms of which portions of the marine‐based sectors will retreat, including the time scales and rates
of change. This review has focused on research that covers many different processes and feedbacks govern-
ing AIS dynamics with an emphasis on the period since IPCC AR5 and on evidence concerning the potential
for rapid AIS evolution. Understanding the processes, feedbacks, and tipping points that govern the AIS is
key to our ability to accurately predict its future behavior. We have highlighted the developments across
multiple disciplines of Earth Science that underpin such insights. The pace of research to address uncertain-
ties in the AIS response to anthropogenic warming and global sea level rise has been, and continues to be,
rapid across all the disciplines. Exciting new interdisciplinary efforts have been facilitated by the
community‐based identification of the major science priorities (Antarctic Science Horizon Scan (“the
Scan”); Kennicutt et al., 2014, 2015, 2019) and the essential technology and infrastructure requirements
(The Antarctic Roadmap Challenges; Kennicutt et al., 2016). The Scan identified the 80 most important
scientific questions clustered into seven groups. This review has touched on 19 of the questions across three
of the groups, including all but one of the 12 questions in the group “Antarctic ice sheet and sea level”
(Table 3). In this section, we highlight some of the unknowns across the different research areas that will
help in improve our understanding of future AIS change and summarize some of the key points from this
multidisciplinary review.
Table 3
The Key Scientific Questions Discussed in This Review From Three of Seven Clusters Developed by the Antarctic and Southern Ocean Science Horizon Scan
(Kennicutt et al., 2015)
Southern Ocean and sea ice
in a warming world
Antarctic ice sheet
and sea level
The dynamic Earth: probing beneath
Antarctic ice
12. Will changes in the Southern Ocean result in
feedbacks that accelerate or slow the pace
of climate change?
13. Why are the properties and volume of Antarctic
Bottom Water changing, and what are the
consequences for global ocean circulation
and climate?
14. How does Southern Ocean circulation,
including exchange with lower latitudes,
respond to climate forcing?
15. What processes and feedbacks drive changes in
the mass, properties, and distribution of
Antarctic sea ice?
20. How do extreme events
affect the Antarctic cryosphere and Southern
Ocean? (Crosscuts “Antarctic ice sheet”)
21. How did the Antarctic cryosphere and the
Southern Ocean contribute to glacial/interglacial
cycles? (Crosscuts “Antarctic ice sheet”)
23. How will changes in freshwater inputs affect
ocean circulation and ecosystem processes?
(Crosscuts “Antarctic life”)
24. How does small‐scale morphology in subglacial
and continental shelf bathymetry affect Antarctic
ice sheet response to changing environmental
conditions? (Crosscuts “Dynamic Earth”)
25. What are the processes and properties that
control the form and flow of the Antarctic
ice sheet?
26. How does subglacial hydrology affect ice sheet
dynamics, and how important is it? (Crosscuts
“Dynamic Earth”)
27. How do the characteristics of the ice sheet bed,
such as geothermal heat flux
and sedimentdistribution, affect ice
flow and ice sheet
stability? (“Crosscuts Dynamic Earth”)
28. What are the thresholds that lead to irreversible
loss of all or part of the Antarctic ice sheet?
29. How will changes in surface melt over the ice
shelves and ice sheet evolve, and what will be
the impact of these changes?
30. How do oceanic processes beneath ice shelves
vary in space and time, how are they modified by
sea ice, and do they affect ice loss and ice sheet
mass balance? (Crosscuts “Southern Ocean”)
32. How fast has the Antarctic ice sheet changed in
the past and what does that tell us about
the future?
33. How did marine‐based Antarctic ice sheets
change during previous interglacial periods?
34. How will the sedimentary record beneaththe ice
sheet inform our knowledge of the presence or
absence of continental ice? (Crosscuts
(“Dynamic Earth”)
37. What is the crust and mantle structure of
Antarctica and the Southern Ocean, and how
do they affect surface motions due to glacial
isostatic adjustment?
38. How does volcanism affect the evolution
of the Antarctic lithosphere, ice sheet dynamics,
and global climate? (Crosscuts
“Antarctic ice sheet”)
40. How do tectonics, dynamic topography, ice
loading, and isostatic adjustment affect
thespatial pattern
of sea level change on all timescales?
(Crosscuts “Antarctic ice sheet”)
Note. The numbering has been maintained as originally published and does not indicate the relative importance of the questions.
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7.1. AIS Interactions With the Southern Ocean
Observations extending back to 1979 show that the AIS has been losing mass. Later oceanographic observa-
tions linked high glacier melt rates to the presence of warm CDW on the Antarctic shelves of West
Antarctica (Jacobs et al., 1996). The rate of ice mass loss has accelerated since 2000s and resulted in an
increasing contribution of AIS melt to GMSL rise, mainly from thinning and retreat of the WAIS. The mass
balance of the EAIS is currently still unclear and associated with larges uncertainty in estimates, which
requires improved observations of precipitation and constraints on GIA. The rate of global ocean warming
has continued to increase in response to anthropogenic emissions, with the majority of the heat being
absorbed by the Southern Ocean and transported north away from the Antarctic margin (Swart et al., 2018).
More observations are required to attribute and quantify the observed slowdown in AMOC to anthropogenic
forcing (Frajka‐Williams et al., 2019). However, the link between AMOC collapse and buildup of heat in the
Southern Ocean resulting in subsequent AIS mass loss is supported by paleoenvironmental studies, particu-
larly for the Last Interglacial (Marino et al., 2015; Rohling et al., 2019). The temporal resolution of proximal
sediment records around the Antarctic margin currently limits the assessment of abrupt (centennial‐scale)
climatic events in the paleorecord, but modeling work (e.g., Pedro et al., 2018) reveals a physical mechanism
for the centennial‐scale response to AMOC collapse, with the accumulation of heat in the Southern Ocean,
and eddy‐driven heat flux across the ACC resulting in surface warming and sea ice retreat.
The observed increase in Southern Ocean stratification and increased surface buoyancy of the Southern
Ocean in response to accelerated ice mass loss and retreat in Antarctica will have impacts on both sea ice
and overturning circulation. Improvements are required in the representation of DSW and AABW in ice
sheet‐climate models. Since a large part of the AIS mass loss is driven by incursions of warm CDW onto
the shelf including a more accurate quantification of sub‐ice shelf melt rates, well‐resolved measurements
and model representations of subsurface temperature changes on the Antarctic shelf are required for accu-
rate future projections of AIS change. Long‐term ice and ocean monitoring programs are all needed to
improve understanding of the feedbacks and interactions leading to warm water incursions on the
Antarctic shelf, and how these respond to a changing climate.
7.2. Ice Sheet Dynamics, the Solid Earth, and Sea Level
The paleoclimate record can provide insight into the behavior of the natural system over longer time scales
and can be used to assess the validity of coupled Earth system models that are widely used to predict future
changes. These longer millennial‐scale records also reveal the importance of understanding the solid Earth
feedbacks on AIS dynamics. Temperature increases of 1° to 3° warming could result in a fourfold difference
of 6 to 27 cm in sea level rise to 2070 (Rintoul et al., 2018). Predicting the long‐term equilibrated AIS response
and subsequent sea level rise requires insight from past warm interglacial periods of the Pliocene (400 ppm;
2–4°C warmer) and Pleistocene warm times (~280 ppm; 1–2°C). Geological records from close to Antarctica
and ice sheet modeling experiments indicate that the AIS, including the EAIS, was highly sensitive to past
changes in atmospheric CO
2
concentrations. Indeed, a strong correlation exists between GMSL and atmo-
spheric CO
2
concentrations (indicator for global temperature) for these geological times, which cannot be
explained considering the Northern Hemisphere ice sheets alone (Foster & Rohling, 2013).
Ice sheet models incorporating dynamic ice processes forced by ocean‐driven melting indicate up to 0.45 (no
MICI; Edwards et al., 2019; Golledge et al., 2019) to 1 m (MICI; DeConto & Pollard, 2016) of GMSL rise from
the AIS alone by 2100, and up to ~11 m (RPC8.5; Pollard et al., 2017) by the Year 2500. However, the inclu-
sion of hydrofracturing and ice cliff instability processes in ice sheet models compared to traditional ice flow
physics remains uncertain. There is a definite role for improved constraints on past sea level estimates from
the geological record, for robust calibration of ice sheet models where the suitability of dynamic ice mechan-
isms is currently unknown.
Improved projections of sea level rise from the AIS based on ice sheet modeling requires reduced uncertainty
in reconstructions of past global sea level and its rates of change, for the interglacial periods of the
Pleistocene (e.g., LIG; Marine Isotope Stage 11), and the warm Pliocene (e.g., Mid Pliocene Warm Period
or early Pliocene). Paleo sea level estimates for the AIS contributions to global sea level change are hampered
by our limited knowledge of past ice sheet configuration and volumes for key glacial periods, such as the
penultimate glacial maximum (MIS 6) and those preceding the warm interglacials of interest. Such recon-
structions require increased spatial sampling of marine sediment deposits around Antarctica with improved
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dating techniques and novel proxies to constrain AIS retreat and meltwater input. These data can be used to
inform models of GIA, grounding line migration and ice flow, and help reduce uncertainties in future global
sea level predictions. Related to this, more studies of dynamic topography are needed to improve the accu-
racy of sea level estimates in the geological past. Ice core evidence from the subglacial basins of the WAIS
and Wilkes Subglacial Basin (East Antarctica) could be drilled to establish the presence or absence of ice dat-
ing back to the Last Interglacial Period. In addition, ice cores from the AIS periphery could be used to estab-
lish the prevalence of surface melt layers, as indicators of conditions where hydrofracturing may have
occurred at the coast, during past warm climates.
New paleoenvironmental archives have been drilled from different sectors of Antarctica by the International
Ocean Drilling Program (e.g., Ross Sea, IODP 374; McKay et al., 2019); Amundsen Sea, IODP 379; Gohl
et al., 2019); Scotia Sea, IODP 382;Weber et al., 2019) and will facilitate reconstructions of both AIS dynamics
and Southern Ocean feedbacks. Observations are needed over a range of spatial scales (catchment to sub-
catchment), and with well‐defined uncertainty bounds, for parameters such as ocean temperature changes
around the Antarctic margin. In addition, proxies that capture ice mass loss as well as interaction between
the Southern Ocean and the AIS (e.g., changes in stratification and sea ice cover) would be valuable for cali-
bration of coupled ice‐climate modeling studies.
Quantifying the extent and timing of deglacial retreat in multiple sectors of Antarctica is required to test
hypotheses regarding the climatic mechanisms for initiating widespread marine AIS retreat. In this context,
it is essential to constrain the roles that internal ice sheet dynamics, solid Earth processes and
atmosphere‐oceanic feedbacks had in either accelerating or slowing retreat during the last glacial termina-
tion and the Holocene. Isostatic feedbacks are also critical to consider when comparing model and
data‐based reconstructions, with mantle viscosity being a fundamental control on the glacio‐isostatic adjust-
ment rate and therefore ice mass loss. The sensitivity of this feedback between the solid Earth rebound and ice
mass change is important for regions of low Earth viscosity, such as beneath the WAIS, in order to accurately
predict the future AIS contribution to, and rate of, sea level rise. In a similar way, the interaction of glacial
retreat and advance events with other basal conditions (e.g., till mechanics, hydrology, and hydrogeology)
is likely significant, although the overall role of these in AIS dynamics remains a key knowledge gap.
Understanding the ice flow processes that could lead to rapid AIS grounding line retreat is essential in
improving model projections of sea level rise. Key areas for improvement include basal friction laws and
their impact on grounding line dynamics and migration; ice flow relations and the role of deformation in
ice shelves, the slow‐flowing interior of the AIS, and at transitions between deformation‐dominated and
sliding‐dominated flow; and regional investigations of interactions between ice dynamic processes, pinning
points, ocean circulation, and ice shelf response. Improving ice sheet model representation of flow processes
will be facilitated by increased accuracy in boundary conditions, including bed topography and ice shelf melt
rates, and synthesis of observations, for example, in time‐dependent adjoint modeling frameworks.
Collaborative initiatives incorporating both observational and modeling efforts to ascertain the current state
and evolution of entire glacial basins, such as currently being undertaken for the Thwaites Basin as part of
the International Thwaites Glacier Collaboration, will be instrumental in improving our understanding of
grounding line controls and in predicting future mass flux changes to the AIS. Similar programs are required
to understand the time scales of risk in the key marine‐grounded basins of East Antarctica (i.e., in the Aurora
and Wilkes Subglacial Basins).
For progress across disciplines, high‐resolution bed topography and bathymetry data sets are paramount in
development of accurate numerical model simulations of the AIS and Southern Ocean interactions, as well
as solid Earth and paleoenvironmental reconstructions. High‐resolution information of the Earth structure
beneath the ice sheet is needed for incorporation of these 3‐D Earth properties in ice sheet models to advance
predictions of the AIS response to future climate change. The solid Earth has a long memory for past ice
loading changes, so accurate knowledge of Earth's rheology is vital for understanding the interactions and
feedbacks relating to changes between the solid Earth and ice mass, as well as accurate GIA corrections
for past sea level reconstructions.
It is clear that our best chance of making major progress in understanding how the AIS will evolve in the
future will come from a cross‐disciplinary approach. Advances will require integrated modeling and obser-
vational studies across the atmosphere, ocean, ice (ice sheet, ice shelves, and sea ice), and solid Earth
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including subglacial hydrology and high‐resolution bathymetry, across a wide range of temporal and spa-
tial scales. This review demonstrates the value of integrating knowledge from paleoenvironmental obser-
vations and modeling with modern observations of change in the AIS and climate system, for
developing a sufficiently detailed understanding of the processes required to accurately predict the future
AIS evolution.
Data Availability Statement
No new data were used in the review article, which is based on existing data from previously published
sources. The sources of the public or published data used in the figures are detailed in the legend.
Acronyms and Glossary
ACR Antarctic Cold Reversal
AIS Antarctic Ice Sheet
ASF Antarctic Slope Front
CDW Circumpolar Deep Water
DSW Dense shelf water
EAIS East Antarctic Ice Sheet
ENSO El Niño–Southern Oscillation
GIA Glacial isostatic adjustment
GMSL Global mean sea level
GRACE Gravity Recovery and Climate Experiment
LGM Last Glacial Maximum
LIG Last Interglacial Period
MIS Marine Isotope Stage
MICI Marine ice cliff instability
MISI Marine ice sheet instability
MWP‐1a Meltwater Pulse 1a
MOC Meridional Ocean Circulation
NEEM North Greenland Eemian
RCPs Representative Concentration Pathways
SAM Southern Annual Mode
WAIS West Antarctic Ice Sheet
Anisotropic ice fabric patterns of preferred ice crystal orientations that is consistent
with the underlying stress regime
Buttressing tendency of an ice shelf to slow or halt the flow of upstream
grounded ice through interaction with topography, for
example, basal or lateral pinning points
Cartesian grid a regular grid composed of unit square or cube elements with
vertices defined on the integer lattice.
Diamict poorly sorted terrigenous material (e.g., cobbles, sand, silt, and
clay) a wide range of particle sizes associated with ice
transport and deposition.
Geothermal heat flux the amount of heat through unit area and unit time coming
from the Earth's interior. This contains the primordial heat
trapped in the interior of the planet and the heat produced by
radiogenic elements within the crust. The main sources of
geothermal heat flux variation are heat flux from the mantle,
linked to tectonic events such as rifting and collision; heat
production within the crust, derived from high concentrations
of radiogenic isotopes, for example, uranium, thorium, and
potassium; and more local effects of volcanic and shallow
intrusive magmatism.
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Glacial erratics rocks carried by glaciers over large distances away from the
source of origin, ranging in size from pebbles to car‐sized
boulders
Grounding line/zone transition line/zone between grounded ice sheet and floating ice
shelf
Holocene the last ~11,600 years of climate history
Ice shelf draft elevation of the base of the ice shelf
Laplace domain in solid Earth modeling mathematical tool used to work in a spatial‐frequency domain
when dealing with complex signals.
Last deglaciation the period where the climate transitioned from cold to warm
conditions (~19,000–11,600 years ago)
Last Glacial Maximum the most recent glacial period approximately 27,000–
19,000 years ago
Last Interglacial Period the penultimate warm period in Earth's climate history around
130,000–115,000 years ago
Pleistocene Geological epoch spanning the last 2.5 million years of Earth
history until the beginning of the Holocene.
Overdeepened continental shelves topographically confined depressions formed by glacier erosion
across the continental shelves, which can be 200–500 m and up
to 1,000 m below sea level.
Rheology The response of a material to deformation when stress or strain
is applied. Relevant to both the solid Earth, which can deform
elastically, viscously, and plastically; and the ice sheet.
Termination II the transition from glacial climate to interglacial climate
following the end of Marine Isotope Stage 6 (the penultimate
glaciation).
Shallow ice approximation approximation to the full Stokes equations; assumes that ice
flows as a result of gravitational driving stresses, balanced by
basal shear stresses.
Stokes equations equations of motion for ice flow
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Acknowledgments
This review paper is the result of a
workshop sponsored by the Australian
Academy of Sciences Frederick and
Elizabeth White Award, the Australian
Research Council's Special Research
Initiative for Antarctic Gateway
Partnership (Project ID SR140300001),
and the University of Tasmania. This
work contributes to The Australian
Centre for Excellence in Antarctic
Science funded by Australian Research
Council Project ID SR200100008.
T. L. N. was supported by the Science
Industry Endowment Fund John
Stocker Postdoctoral Fellowship. The
authors are grateful to C.‐D.
Hillenbrand and S. S. R. Jamieson for
extensive and thorough reviews and for
sharing their knowledge of all things
Antarctic. Thanks to A. Durand for
initial reference support, S. Cook and T.
Staal for editing text, and L. Bell for
graphic design.

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